SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES ? NUMBER 33
Great Basin Aquatic Systems History
Robert Her shier, DavidB. Madsen, and Donald R. Currey
EDITORS
fSS'JIH)
DEC 112002
SMRHSONIAN INSTITUTION
Smithsonian Institution Press
Washington, D.C.
2002
ABSTRACT
Hershler, Robert, David B. Madsen, and Donald R. Currey, editors. Great Basin Aquatic Sys-
tems History. Smithsonian Contributions to the Earth Sciences, number 33,405 pages, 190 fig-
ures, 21 tables, 2002.?The 14 papers collected herein treat diverse aspects of the aquatic
history of the Great Basin of the western United States and collectively attempt to summarize
and integrate portions of the vast body of new information on this subject that has been
acquired since the last such compilation was published in 1948. In the first section, four papers
(Lowenstein, Negrini, Reheis et al., Sack) focus on physical aspects of Great Basin paleolake
histories, whereas a fifth paper (Oviatt) summarizes the contributions to the study of Bonne-
ville Basin lacustrine history made by two early giants of the field, Grove Karl Gilbert and
Ernst Antevs. In the second section, four papers synthesize perspectives on Great Basin aquatic
history provided by diatoms and ostracods (Bradbury and Forester), fishes (Smith et al.),
aquatic insects (Polhemus and Polhemus), and aquatic snails (Hershler and Sada), whereas a
fifth (Sada and Vinyard) summarizes the conservation status of the diverse aquatic biota that is
endemic to the region. In the final section, three papers integrate terrestrial biotic evidence per-
taining to Great Basin aquatic history derived from pollen from cores (Davis), floristics
(Wigand and Rhode), and the mammal record (Grayson), whereas a fourth (Madsen) examines
the relationship between Great Basin lakes and human inhabitants of the region. Although
diverse in scope and topic, the papers in this volume are nonetheless linked by an appreciation
that integration of geological, biological, and anthropological evidence is a necessary and fun-
damental key to a mature understanding of Great Basin aquatic systems history.
OFFICIAL PUBLICATION DATE is handstamped in a limited number of initial copies and is
recorded in the Institution's annual report, Annals of the Smithsonian Institution. SERIES COVER
DESIGN: Volcanic eruption at the island of Krakatau in 1883.
Library of Congress Cataloging-in-Publication Data
Great Basin aquatic systems history / Robert Hershler, David B. Madsen, and Donald R. Curry, editors.
p. cm. ? (Smithsonian contributions to the earth sciences ; no. 33)
Includes bibliographical references.
1. Limnology?Great Basin. 2. Paleolimnology?Great Basin. I. Hershler, Robert. II.
Madsen, David B. III. Currey, Donald R. IV. Series.
QE1 .S227 no. 33
[QH104.5.G68]
550 s?dc21
[551.48'0979]
2002026685
? The paper used in this publication meets the minimum requirements of the American
National Standard for Permanence of Paper for Printed Library Materials Z39.48?1984.
Contents
Page
Introduction, by David B. Madsen, Robert Hershler, and Donald R. Currey 1
LAKE HISTORIES
Pluvial Lake Sizes in the Northwestern Great Basin throughout the Quaternary
Period, by Robert M. Negrini 11
Pliocene to Middle Pleistocene Lakes in the Western Great Basin: Ages and Connec-
tions, by Marith C. Reheis, Andrei M. Sarna-Wojcicki, Richard L. Reynolds,
Charles A. Repenning, and Martin D. Mifflin 53
Pleistocene Lakes and Paleoclimates (0 to 200 Ka) in Death Valley, California, by
Tim K. Lowenstein 109
Bonneville Basin Lacustrine History: The Contributions of G.K. Gilbert and Ernst
Antevs, by Charles G. Oviatt 121
Fluvial Linkages in Lake Bonneville Subbasin Integration, by Dorothy Sack 129
AQUATIC BIOTIC PERSPECTIVES
Environment and Paleolimnology of Owens Lake, California: A Record of Climate
and Hydrology for the Last 50,000 Years, by J. Platt Bradbury and R.M.
Forester 145
Biogeography and Timing of Evolutionary Events among Great Basin Fishes, by
G.R. Smith, T.E. Dowling, K.W. Gobalet, T. Lugaski, D.K. Shiozawa, and
R.P. Evans 175
Basins and Ranges: The Biogeography of Aquatic True Bugs (Insecta: Heteroptera)
in the Great Basin, by Dan A. Polhemus and John T. Polhemus 235
Biogeography of Great Basin Aquatic Snails of the Genus Pyrgulopsis, by Robert
Hershler and Donald W. Sada 255
Anthropogenic Changes in Biogeography of Great Basin Aquatic Biota, by Donald
W. Sada and Gary L. Vinyard 277
NON-AQUATIC BIOTIC AND HYDROCLIMATIC PERSPECTIVES
Late Neogene Environmental History of the Northern Bonneville Basin: A Review
of Palynological Studies, by Owen K. Davis 295
Great Basin Vegetation History and Aquatic Systems: The Last 150,000 Years, by
Peter E. Wigand and David Rhode 309
Great Basin Mammals and Late Quaternary Climate History, by Donald K. Gray-
son 369
Great Basin Peoples and Late Quaternary Aquatic History, by David B. Madsen ... 387
m
Introduction
David B. Madsen, Robert Her shier, andDonaldR. Currey
It has long been known that the huge, arid region of internal
drainage in western North America known as the Great Basin
contained a series of large, intermontane lakes during wetter, or
pluvial, periods of the late Cenozoic. During the last 100 years
numerous geologists and biologists, including many preemi-
nent scientists beginning with Grove Karl Gilbert (1843-1918)
and Israel Cook Russell (1852-1906), have studied the extent,
integration, and contemporaneity of these lakes, and the rela-
tion between regional hydrographic history and biogeography.
The first attempt to synthesize this information was in June
1942, when the Western Division of the American Society of
Ichthyologists and Herpetologists met in Salt Lake City, Utah.
The highlight of the meeting was a symposium on the glacial
and postglacial history of the Great Basin, which was jointly
arranged with the Pacific Division of the American Association
for the Advancement of Science (Croker, 1942). Papers pre-
sented at the meeting included summaries of the geological de-
velopment of the Great Basin, pluvial lake distribution and per-
tinent ichthyological evidence, and paleoclimates, which were
later collected in The Great Basin, with Emphasis on Glacial
and Postglacial Times (Antevs, 1948; Blackwelder, 1948;
Hubbs and Miller, 1948). This modest volume has stood as a
landmark in the study of what Grayson (1993:xvi) referred to
as the "natural prehistory" of the Great Basin.
Immense strides have been made in the study of Great Basin
aquatic history during the last 50 years. On the physical side,
our understanding of geological evolution of the Great Basin
has been revitalized in the wake of modern concepts of moun-
tain building, tectonics, and volcanology. Geochronology has
provided many new dating methods to aid interpretation of key
geologic events. New methodologies also have enabled much
more precise delineation and dating of ancient lake beds and
better discernment as to whether such data represent records of
long-lived lakes or ephemeral water bodies (playas). Climatic
David B. Madsen, Environmental Sciences, Utah Geological Survey,
Salt Lake City, Utah 84114, USA. Robert Hershler, Department of Sys-
tematic Biology, National Museum of Natural History, Smithsonian
Institution, Washington, D.C. 20560-0118, USA. DonaldR. Currey,
Department of Geography, University of Utah, Salt Lake City, Utah
84112, USA.
modeling has developed into an important field that further
aids elucidation of pattern and sequence of aquatic-system de-
velopment. On the biological side, evidence derived from
fishes (both living and fossil) has been expanded and refined,
and other aquatic biota now has been similarly studied. The
coupling of biogeographic analysis with phylogenetics has pro-
vided explicit, testable methods of inferring historical relation-
ships among geographic areas (drainages). The blossoming
field of phylogeography (Bermingham and Moritz, 1998) of-
fers additional tools for fruitful analysis at the population level,
although this has been little applied to aquatic biota of the
Great Basin (but see Duvernell and Turner, 1998).
The purpose of this volume is to assemble and integrate this
large body of relevant data from the geological and biological
schools as it pertains to the history of aquatic systems in the
Great Basin. This is hardly a new concept, but instead it reiter-
ates the spirit of Hubbs and Miller (1948:120):
Even greater advances should ensue, when the biological and geological ap-
proaches are more definitely integrated. We eagerly look forward to the time
when we may join with the glacial geologist to journey, helping hand in hand,
across the desert mountains and flats of the Great Basin, together attempting to
unlock some of the doors that are yet closed to us.
This volume is a product of the Great Basin Aquatic System
History (GBASH) Symposium, which was convened in Salt
Lake City at the University of Utah from 17 September to 21
September 1997 to celebrate the 50th anniversary of the publi-
cation of papers from the original 1942 symposium. More than
75 individuals attended the GBASH conference, including in-
vited speakers, poster presenters, representatives of land man-
agement agencies, and the general public. Participants enjoyed
a highly structured program designed to summarize recent ad-
vances in Great Basin natural history in an integrated, interdis-
ciplinary fashion. Presenters included an array of geologists,
climatologists, botanists, zoologists, and anthropologists under
the general direction of conference chair Donald R. Currey.
The presentation format for the 25 invited papers followed the
same tripartite structure we employ herein. Fourteen of those
papers were revised subsequent to the symposium and are col-
lected in this volume. Together these papers complement the
more generalized treatments of Great Basin aquatic history that
are currently available.
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 1.?The limits of the Great Basin, as defined by areas of internal drain-
age (after King, 1986).
Defining the Great Basin
The Great Basin is a vast complex of internal drainage sys-
tems lying in interior western North America (Figure 1). These
systems are composed primarily of high, north-to-south trend-
ing mountains and long, narrow valleys into which mountain
streams and rivers drain. Depending upon how it is defined
(and it can be defined in a number of ways), the Great Basin
extends as far north as 44? north latitude, as far south as 33?
north latitude, as far east as 110? west longitude, and as far
west as 121? west longitude. Although the many definitions of
the Great Basin are appropriate for specific purposes, none ex-
ceeds that of John C. Fremont's (1845:175) original description
in terms of simple utility:
The Great Basin?a term I apply to the intermediate region between the Rocky
mountains and the next range [the Sierra Nevada and the Cascade Range], con-
taining many lakes, with their own system of rivers and creeks, (of which the
Great Salt is the principal,) and which have no connexion with the ocean, or the
great rivers which flow into it.
A number of general accounts reviewing the natural history
of the Great Basin are available (e.g., Houghton, 1976;
McPhee, 1981; Trimble, 1989; and particularly, Grayson,
1993), and there is also a variety of extensive treatments of
specific topics (Lanner, 1983; Ryser, 1985; Fiero, 1986; Moz-
ingo, 1987; and Sigler and Sigler, 1987). As a result, we have
not attempted to produce yet another generalized description of
the Great Basin herein. Furthermore, because a variety of bio-
logical, geomorphological, and hydrological topics are covered
in the papers that follow, it is important to recognize that the
physiographic, hydrographic, and floristic definitions of the
Great Basin employed in these papers are somewhat different.
Physiographically, the Great Basin is part of the Basin and
Range Province (Hunt, 1967) that, although coterminous with
the hydrographic Great Basin in the north, extends into south-
ern Arizona and New Mexico and covers much of northern
Mexico (Figure 2). The fault-block mountains that characterize
the province began forming during the middle Tertiary, approx-
imately 40 million years (My) before present (B.P.) (King,
1977). This system of mountains and valleys was formed as
part of an ongoing extensional process related to plate tectonics
(Zoback et al., 1981), and the Great Basin continues to expand
as the Pacific Plate shears off the western margin of the North
American plate (Stewart, 1971; Fiero, 1986). This east-west
expansion has created a series of north-to-south trending fault-
block mountains separated by long, narrow graben and/or half-
graben valleys filled with Quaternary alluvium. The larger val-
leys exceed 100 miles (161 km) in length and 30 miles (48 km)
in width. The Great Basin is bounded on the west and east by
the higher ranges of the Sierra Nevada and the western Rocky
FIGURE 2.?The limits of the Basin and Range Province (after Stewart, 1971;
Fiero, 1986).
NUMBER 33
FIGURE 3.?The limits of the floristic Great Basin (after Cronquist et al., 1972).
Mountains, respectively, but many of the smaller interior
ranges reach almost as high because the Great Basin is dome-
shaped in cross section. Many ranges reach 12,000-13,000 ft
(3660-3960 m) and extend 5000-6000 ft (1524-1829 m)
above valley floors.
The biotic Great Basin is defined primarily by the distribu-
tion of plant communities rather than by the animals that are
dependent upon them (Gleason and Cronquist, 1964). The most
widely employed floristic definition is that of Cronquist et al.
(1972) (Figure 3). In this interpretation, the floristic Great Ba-
sin extends from the Sierra Nevada on the west to the Wasatch
Front ranges of Utah on the east (Cronquist et al., 1972; Os-
mond et al., 1990). The eastern boundary is not as sharp as that
on the west, however, and many early treatments of Great Ba-
sin flora (e.g., Shreve, 1942) included much of the northern
Colorado Plateau in the floristic Great Basin. Plant communi-
ties are dominated by sagebrush (Artemisia tridentata), shad-
scale (Atriplex confertifolia), and saltbush (Atriplex canescens)
on valley floors, juniper (Juniperous osteosperma) and single-
leaf pinon (Pinus monophyla) on lower mountain slopes in the
central and southern Great Basin, and white fir (Abies con-
color), Engelmann spruce (Picea emglemannii), blue spruce
(Picea pungens), Douglas fir (Pseudotsuga menziesii), and
quaking aspen (Populus tremuloides) at higher elevations.
These communities extend north into the Snake River Plain,
and the floristic Great Basin, as a result, does not conform well
with the Great Basin as defined either physiographically or hy-
drographically. There is similar terminological confusion re-
garding the southern margin of the floristic Great Basin be-
cause this boundary is defined mostly by the appearance of
creosote bush (Larrea tridentata), yucca {Yucca brevifolia),
and other plants of the Mojave Desert south of about 37? north
latitude. When either a hydrographic or physiographic defini-
tion of the Great Basin is employed, the Great Basin actually
contains three of the four biotic North American deserts (Fig-
ure 4).
The Great Basin can be described most definitively in terms
of a series of contiguous interior drainage systems. Although
subject to some debate over whether or not a few individual
drainage basins do indeed have internal drainage (e.g., Mifflin
and Wheat, 1979; Smith and Street-Perrott, 1983; King, 1986),
this hydrographic definition is the most detailed of those avail-
able because drainage basin boundaries can be identified with
precision. Several hydrographic definitions have been pro-
duced, which differ because of the varying number of drainage
NM
FIGURE 4.?The three biotically defined North American deserts with the
hydrographically defined Great Basin (after Shreve, 1942; Spaulding et al.,
1983).
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 5.?Hydrographic basins within the Great Basin (after G.Q. King, unpublished data).
basins included. A hydrographic definition of the Great Basin
is most useful herein because our interest is in Great Basin
aquatic systems, and we employ that of King (1986), who iden-
tified 187 separate drainage basins within the Great Basin as a
whole (Figure 5).
Many of these individual drainages coalesced to form larger
drainage units during cooler and wetter periods when lake lev-
els rose and overflowed into adjoining basins. The largest of
these, by far, are the Lahontan Basin on the west and the
Bonneville Basin on the east, where the relatively more mas-
sive Wasatch Mountains and Sierra Nevada, respectively, cap-
ture substantial amounts of precipitation. During warm and dry
intervals, the rivers of the Lahontan Basin, such as the
Truckee, the Humboldt, the Carson, and the Walker Rivers,
flowed into separate basins and form individual lakes, such as
Pyramid and Walker Lakes. Under higher precipitation and
lower evaporation regimes, however, these basins filled differ-
entially, forming variable combinations of basin systems as
lake waters rose (Davis, 1982). At their maximum, these wa-
ters formed a single lake, covering nearly 8670 square miles
(22,440 km2), during the last glacial cycle. The situation is
similar in the eastern Great Basin where the Sevier and Great
NUMBER 33
V * i
200 km
FIGURE 6.?Large lakes within the Great Basin during the height of the last pluvial period (18-16 Ka) (after Wil-
liams and Bedinger, 1984).
Salt Lake subbasins at times have been either completely sepa-
rate, connected by an overflow channel, or coalesced into a
single lake. At its height, Lake Bonneville extended more than
500 miles (805 km) from north to south, covered more than
19,900 square miles (51,640 km2), and overflowed into the
Snake River and the Columbia River drainage system. At such
times the Bonneville Basin was not, hydrographically speak-
ing, part of the Great Basin at all.
A number of other basins within the Great Basin also over-
flowed during wet periods, but they formed a series of inter-
connected lakes rather than a single, large lake. The most
extensive chain of lakes and basins were formed in the south-
western Great Basin, where Mono, Owens, China, Searles, and
Panamint Lakes (and other smaller lake basins) were connected
at the peak of pluvial periods by a series of overflow channels
that eventually extended into Death Valley (Smith and Street-
Perrott, 1983). A number of shorter, interconnected lake sys-
tems, such as the Malheur Lake-Harney Lake chain in Oregon,
were found in other parts of the Great Basin. Most of the basins
in the central and southern Great Basin contained only isolated
lakes during moist intervals because the mountains that form
their margins are limited both in extent and elevation; however,
several of these isolated lakes, such as Pleistocene Lake Frank-
lin in central Nevada, were relatively large because of the high
elevation of both the basin floor and the surrounding mountain
crests in the central Great Basin. More than 80 percent of the
187 Great Basin drainages contain playas, and most of these
basins likely supported lakes at least seasonally during the peak
of the last pluvial interval (e.g., King, 1986; Williams and Be-
dinger, 1984) (Figure 6).
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 7.?Major subregions of the Great Basin (after Houghton, 1976).
Major subdivisions within the Great Basin have been defined
in a number of ways, much like the basin as a whole. These
subdivisions, also like the basin as a whole, differ markedly de-
pending upon the criteria that are employed in making the divi-
sions. Cronquist et al. (1972) divided the Great Basin floristic
province into 15 separate sections. This number can be reduced
considerably because parts of seven of these sections lie well
outside the Great Basin, as defined either hydrographically or
physiographically, on the Snake River Plain and the Colorado
Plateau. On the other hand, they include none of the plant com-
munities of the Mojave or Sonoran Deserts in the southern hy-
drographic Great Basin. Many of these floristic subdivisions
cross drainage system boundaries. Both Hunt (1974) and
Houghton (1976) divided the Great Basin into five major sub-
regions based upon a combination of hydrographic and physio-
graphic factors (Figure 7). These subdivisions are in relatively
close agreement, differing primarily in the labels that are ap-
plied to them, and they provide a far simpler and more coherent
division of the Great Basin than that of Cronquist et al. (1972).
These include the Bonneville Basin to the east of the central
basin area, the Lahontan Basin to the west of the central area, a
Northwest Lakes (or Lava and Lakes) area to the north of the
Lahontan Basin, and a Southern or Death Valley system to the
south of the central area.
The confusion surrounding various biotic, physiographic,
and hydrographic definitions of the Great Basin is symptomatic
of the difficulties in understanding the history of the Great Ba-
sin aquatic systems. The Great Basin is related in many differ-
ent ways to the areas around it, and it is difficult to tease out
how changes in those relationships through time have differen-
tially affected Great Basin biomes. For example, at times the
Bonneville Basin has been hydrographically part of the Great
Basin, and at other times it has been part of the Columbia River
drainage system. As a result, the aquatic biota of this basin has
a history that is both richer and more complex than that of
many other Great Basin drainages. These different definitions
are also illustrative of the difficulties in making comparisons
across drainages within the Great Basin. Lakes in the south, for
example, are surrounded by vegetation that is so markedly dif-
ferent from that surrounding lakes in the north that floristically
they are not even in the Great Basin. Yet it is these differences
that lie at the heart of the Great Basin and that, by extension, lie
at the heart of the papers in this volume. The Great Basin is at
once both obvious and enigmatic, making attempts to explore
and interpret its natural history both irritating and rewarding.
Chronological Aspects of Great Basin Aquatic Systems
A Great Basin aquatic systems history must include an un-
derstanding of the changing climatic conditions that have
largely structured fluctuations in the size of Great Basin lakes
and the development of biotic communities in the region.
These long-term climatic changes are related most directly to
alterations in the relative distance between the earth and the
sun and, hence, vary with perturbations in the earth's orbit and
axis. These climatic changes operate on a number of different
scales, ranging from about 20,000 to 100,000 or more years,
and, depending upon where the earth is in each of these Mi-
lankovitch cycles, they can reinforce or counteract each other.
For example, the wobble in the earth's axis alternately moder-
ates and exacerbates the change in the amount of energy re-
ceived from the sun as the earth's orbit changes on a longer
scale. For Late Quaternary climates, the 19,000- and 23,000-
year precession cycles and the 41,000-year obliquity cycles are
particularly important (e.g., Grootes and Stuiver, 1997). Cli-
matic change on a millennial scale of 1000-2000 years also has
been recognized in proxy records from ice cores and from
cores taken from the ocean and large continental lakes. As yet,
however, it is not clear how short-term Dansgaard-Oeschger
cyclical variations may relate to the celestial mechanics that
drive the longer cycles. All of these cycles appear to produce
globally linked climatic changes, although the nature of these
changes varies spatially.
The longer-term cycles defined in the oxygen isotope record
of deep-sea cores are also readily recognized in the fluctuation
histories of many Great Basin lakes (e.g., Smith and Bischoff,
1997; Benson et al., 1998). Several of the papers in this volume
explore the effects of these longer-term climatic cycles on
Great Basin lakes, and there is no need to discuss the issue here.
Less well known, however, is the relationship between millen-
nial-scale climatic changes and the fluctuation histories of
NUMBER 33
Great Basin lakes. Most dramatically, it is becoming increas-
ingly clear that the transitions between the steady-state condi-
tions that characterize these millennial-scale cycles are rela-
tively abrupt, often on the order of a decade or less, and
certainly within the lifetime of a human being. Evidence from
Greenland ice cores and from cores taken from the North At-
lantic Ocean indicates that these cycles assumed a characteris-
tic form during the glacial period. They were initiated by a
rapid rise in temperature of 5?-8? C in a matter of decades,
gradually returned to moderate conditions over the course of
about 1000 years, and ended with a rapid return to very cold
temperatures prior to the start of a warming event that initiated
a new cycle. The amplitude of the cycles appears to have been
much smaller during the Holocene (on the order of 1??3? C),
but the pattern seems to have been relatively similar. Twenty-
four cycles have been identified in the last glacial period, and
nine have been recognized in these detailed ice core records
during the 12,000 years of recorded human occupation in the
Great Basin (e.g., Bond et al., 1997). These millennial-scale cy-
cles extend back at least 100,000 years (e.g., Brooks et al.,
1996) and appear to be part of the entire Pleistocene record.
Although it is becoming evident that change in the water
budget of many Great Basin lakes is correlated with millennial-
scale cycles evident in the ice core record (e.g., Benson et al.,
1997; Oviatt, 1997), all of the late glacial and Holocene cycles
have yet to be recognized locally. This is probably due to a
combination of factors, such as the relatively coarse nature of
local environmental records and the presence of other, longer
cycles with periodicities of 6100 and 11,100 years that may ob-
scure local manifestation of millennial-scale fluctuations (e.g.,
Mayewski et al., 1997). Once detailed and sensitive records,
such as those from tree rings and lake varves, are available,
many, but perhaps not all, of these cycles may become evident.
Regardless, it is as important to incorporate a notion of time
into an understanding of Great Basin aquatic systems as it is to
understand the spatial dimensions that define them. At times
Great Basin valleys were filled with water, and at other times
they became completely desiccated; lakes came and went; hy-
drological systems connected and disconnected. These cyclical
changes, together with the basin and range topography, com-
prise the basic structure that underlies Great Basin aquatic sys-
tems.
Organization of the Volume
The first section of this volume includes papers treating the
physical history of Great Basin paleolakes. Negrini's compre-
hensive summary of pluvial lake histories in the northwestern
Great Basin (and evaluation of Antevs' model for Quaternary
climate change) will prove useful not only to geologists but
also to those seeking information relative to aquatic biogeogra-
phy. Millennial-scale oscillations are evident in this sequence,
but it is apparently slightly out of phase with those of the east-
ern Great Basin lakes. Similarly, the paper by Reheis et al.,
which carefully documents new evidence supporting an en-
larged Lake Lahontan drainage system during the late Neo-
gene, also fits well the cross-disciplinary focus of this sympo-
sium. They suggest that lake levels were as much as 70 m
higher during the early to middle Pleistocene than during the
late Pleistocene highstand?a sequence that is also quite differ-
ent from that found in the eastern Great Basin. The highly dy-
namic nature of Quaternary paleoclimates and associated hy-
drological cycles in the Great Basin is emphasized by
Lowenstein, who reviews pertinent high-resolution data from
lake beds of the Death Valley hydrographic system. Lowen-
stein makes it clear that rapid climate-change events evident in
the Greenland ice core and in North Atlantic core records also
are present in Great Basin lake histories. Oviatt summarizes the
lasting contributions of Gilbert and Antevs, two of the contrib-
utors to the 1948 Great Basin volume, to the study of aquatic
history of the Bonneville Basin. Sack marshals geomorphic and
hydrologic evidence to unravel changes in subbasin relation-
ships within a component of the Bonneville Basin, again indi-
cating that complexity is a hallmark of Great Basin aquatic sys-
tems history.
The section on aquatic biology focuses on the latest contribu-
tions that biotic data have made to our understanding of Great
Basin aquatic systems history, and it is intended as a supple-
ment to the earlier papers collected by Harper and Reveal
(1978). Regional aquatic biota includes various taxa that dis-
perse readily, speciate slowly, and whose distribution and oc-
currence more reflects ecological and climatic conditions than
history of drainage interconnections. The fossil record of such
organisms thus can provide a detailed picture of environmental
history. In the paper by Bradbury and Forester, high-resolution
paleontologic data for two such groups from lake bed cores (di-
atoms and ostracods) are integrated with other biotic and phys-
ical evidence to infer the late Pleistocene history of Owens
Lake within the context of Sierra Nevada glaciation and re-
gional climatic history. On the other end of the spectrum, speci-
ose groups, whose members are often locally endemic, can pro-
vide important clues of drainage interconnections, although
this utility is strongly dependent upon the availability of robust
phylogenetic hypotheses as the basis for analysis. Fishes have
long paved the way in this arena, and the paper by Smith et al.
represents a significant advance from the classical treatment of
Hubbs and Miller (1948). In this paper phylogenies of several
groups, based on mitochondrial DNA sequences, together with
a comprehensive review of fossil evidence, are used to refine
previous interpretations of historic drainage relationships and
rates of fish evolution. Fish evidence suggests the importance
of old drainage relationships and vicariant events that preceded
the late Pleistocene pluvial period, as presaged by Minckley et
al. (1986). Papers on aquatic snails (Hershler and Sada) and in-
sects (Polhemus and Polhemus) suggest additional ancient
drainage relationships that predate modern topography, al-
though these hypotheses await testing within the context of
phylogenetic analysis. Great Basin biogeography has changed
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
(extirpation of biota, introduction of exotic taxa, etc.) as a re-
grettable consequence of anthropogenic activities relating to
water development in this arid region. It is appropriate to ask
whether, using the phraseology of Kornfield and Echelle
(1984:251), attention to "tending of the flock" is currently suf-
ficient to ensure long-term persistence of the unique aquatic ec-
osystems in the Great Basin. The chapter by Sada and Vinyard
summarizes the conservation status of aquatic biota endemic to
the Great Basin. One cannot but conclude that conflicting de-
velopment and conservation needs can be better balanced.
The final section on nonaquatic biology and hydroclimatic
perspectives is focused on the climatic mechanisms that drive
changes in Great Basin lake levels and on the biological com-
munities that surround the Great Basin lakes. Because the his-
tory of these biotic systems is, in turn, mostly determined by
fluctuations in the size of the lakes, these changes are necessar-
ily an integral part of Great Basin aquatic systems history. The
papers in this section vary between generalized accounts appli-
cable to the Great Basin as a whole and specific accounts re-
lated to a single lake-basin history, and between consideration
of very long-term changes and of relatively short-term records.
As such, these papers reflect the breadth of studies involving
Great Basin ecosystems that are currently underway. Davis fo-
cuses on the long-term vegetational history of the Bonneville
Basin, as delineated from analysis of the pollen in deep cores
from the Great Salt Lake. He suggests that there was a gradual
evolution of desert flora during the last five million years that
was punctuated by a number of noteworthy events. Pluvial-in-
terpluvial cycles began after about 750,000 years ago, and
Davis suggests that an increase in pollen concentration indi-
cates the Bear River was diverted into the Bonneville Basin
shortly before about 300,000 years ago. Wigand and Rhode
provide a detailed compilation of Great Basin floristic changes
during the last 150,000 years, focusing primarily on the last
15,000-20,000 years. Using an extensive array of data derived
from pollen, midden, and tree-ring records from the northwest-
ern, southern, and northeastern Great Basin, they suggest that
Great Basin ecosystems have not evolved in a holistic fashion.
Rather, Great Basin floristic communities are a continually
changing mosaic of individual organisms, and Wigand and
Rhode suggest that modern and prehistoric communities are
not appropriate analogues for one another. Grayson uses a de-
tailed, small-mammal record derived from the collecting activ-
ity of owls to assess the nature of climatic change during the
Pleistocene-Holocene transition and the early Holocene in the
Great Salt Lake basin. He suggests that the early Holocene was
both cooler and wetter than at present, in contrast to climate
models that indicate otherwise. He too notes that faunal com-
munities are composed of species that react differently to
changing climatic conditions and finds support for species-
richness response models that predict increasing diversity with
increasing precipitation. Finally, Madsen reviews the relation-
ship between Great Basin lakes and Great Basin peoples. He
notes that this relationship is mostly structured by the dichot-
omy between the relatively rich and productive riverine and
lacustrine ecosystems and the relatively dry and less productive
landscapes that surround them. As a result, changes in the size
and distribution of lakes have been critical components in the
development of prehistoric human society in the Great Basin.
Acknowledgments
We thank the authors for their timely submissions and for the
patience they displayed with delays in the editorial process. We
also thank the many colleagues who served as reviewers of
these contributions. The Smithsonian Institution provided
funds to support the publication of this volume. This institu-
tion, together with the University of Utah and the Utah Geolog-
ical Survey, also provided support for the GBASH conference.
Linda Williams and Peter Hovingh assisted in many ways with
the organization of the conference. We thank Guy King for per-
mission to use his unpublished map of Great Basin drainage
units.
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434.
Pluvial Lake Sizes in the Northwestern Great Basin
throughout the Quaternary Period
Robert M. Negrini
ABSTRACT
Sediments deposited by pluvial lakes of the northwestern Great
Basin are good repositories of paleoclimate data, which have been
used to test models of global climate change whose scale ranges
from multidecade to megayear. The principal driving models
include (1) Milankovitch-scale (106-104 yr) jet stream migration
in reponse to changes in the sizes of continental glaciers, (2) mil-
lennial-scale (104-102 yr) climate change as a hemispheric
response to large-scale climate changes in the North Atlantic
Ocean (i.e., Dansgaard-Oeschger cycles, Heinrich Events, and
Bond Cycles), and (3) Antevs' (1948) phenomenological model of
climate change in the Holocene, which primarily predicts, on the
average, a relatively warm and dry interval in the middle part of
the Holocene. Changes in the Earth's orbital geometry are the ulti-
mate causes for Milankovitch-scale climate changes. Changes in
the solar constant might prove to be the ultimate cause for climate
change with periods long enough to overlap the part of the spec-
trum occupied by the Holocene changes and by the millennial-
scale climate changes; however, the magnitude of change expected
from solar processes would be much smaller than those associated
with millennial-scale climate changes, so if there is a causal con-
nection, it would have to be strongly nonlinear.
Climate records from the northwestern Great Basin, particularly
those from pluvial Lakes Modoc, Lahontan, and Chewaucan, sup-
port models related to Milankovitch-scale climate changes by
exhibiting the appropriate periodicities and/or phases of predicted
climate change and by exhibiting the appropriate dynamic spectral
content of climate change throughout time (e.g., the change from a
dominance of periods lasting 400 thousand years (Ky) and 41 Ky
to a dominance of 100 Ky starting at -900 thousand years ago
Robert M. Negrini, Department of Physics and Geology, California
State University, Bakersfield, California 93311, USA.
Reviewers: Andrew S. Cohen, Department of Geosciences, University
of Arizona, Tucson, Arizona 85761, USA; Timothy K. Lowenstein,
Department of Geological Sciences, State University of New York at
Binghamton, Binghamton, New York 13902, USA; James F. Luhr,
Department of Mineral Sciences, National Museum of Natural His-
tory, Smithsonian Institution, Washington, D.C. 20560-0129, USA;
and one anonymous reviewer.
(Ka)). Records of millennial-scale change from the northwestern
Great Basin are beginning to appear in the literature, but at the
time of this review, no published records have been dated precisely
enough for comparison with the archetypical records from North
Atlantic marine and ice cores. A collective record associated with
pluvial lakes throughout the northwestern Great Basin indicates a
weak trend toward overall lower lake levels and warmer, drier cli-
mate during the middle Holocene, thereby supporting Antevs'
(1948) model of Holocene climate change for the Great Basin.
Introduction
Approximately 50 years ago, Ernst Antevs (1948) proposed
a model of Quaternary climate change for the Great Basin. The
inductive reasoning that led to this model was based in part
upon observations made by Allison (1945) and Russell (1895)
of changes in pluvial lake levels through time in the northwest-
ern Great Basin. The pluvial lakes of the internally drained
Great Basin hydrological province have continued to be a valu-
able source of nonmarine paleoclimate data primarily because
the sizes of these lakes are especially sensitive to changes in
precipitation and evaporation (e.g., Morrison and Frye, 1965;
Mifflin and Wheat, 1979; Smith and Street-Perrott, 1983; Ben-
son and Thompson, 1987a; Benson et al., 1990; Morrison,
1991a; Spaulding, 1991; Oviatt et al., 1992; Grayson, 1993;
Benson et al., 1995). Furthermore, these size changes are re-
corded by the presence of a wide range of observable features
in the lake system, including tufa deposits, strandlines, and
biota, and by a wide range of measurable features, including
the geochemistry, physical lithology, and magnetism of lake
sediments (Currey, 1994b). Finally, because most of the basins
have been actively subsiding along marginal faults for much of
the late Cenozoic era, their deposits potentially carry paleocli-
mate records throughout the Quaternary and earlier into the lat-
est Tertiary period.
Records from lakes in the northwestern part of the Great Ba-
sin are set apart from those of the rest of the Great Basin because
11
12 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
this region is located downwind from the volcanoes of the Cas-
cade Range. Consequently, volcanic ashes are ubiquitous in
lake-bottom sediments and are commonly found in context with
other geomorphic features associated with these lacustrine sys-
tems. For example, a 17 m section of lake-bottom sediments ex-
posed near Summer Lake, Oregon, contains more than 50 tephra
layers, and the overlying sand dunes contain fragments of the
Mazama tephra (Allison, 1982; Davis, 1985). The geochemical,
petrographic, and, to a lesser extent, the paleomagnetic signa-
tures of such ashes allow excellent intrabasinal correlations
among outcrops and cores as well as extrabasinal correlations
with other well-dated sequences of lake sediments or volcanic
deposits. The absolute ages of the ashes often can be determined
directly by using a variety of radiometric and luminescence
methods, thus extending the age control of these sediments be-
yond the range of radiocarbon dating (Berger, 1991; Sarna-
Wojcicki and Davis, 1991).
This paper is a review of research on northwestern Great Ba-
sin pluvial lake levels published since Antevs' 1948 paper. An-
te vs' original hypotheses are summarized first, followed by a
review of improvements made on these models during the last
50 years. Next is a brief review of other climate-change models
that potentially impact the northwestern Great Basin. Finally,
advances in the associated observations (i.e., the prehistory of
northwestern Great Basin lakes) made during the last several
decades are summarized, focusing on studies done after previ-
ously published, related reviews (e.g., Davis, 1982; Smith and
Street-Perrot, 1983; Benson and Thompson, 1987a; Morrison,
1991a, 1991b; Grayson, 1993). The spatial coverage includes
pluvial Lake Lahontan and all of the Great Basin pluvial lakes
to the north and west of Lake Lahontan for which published re-
search is available (Figure 1). The temporal scope of this paper
is the Quaternary period, during which the late Tertiary Great
Basin was influenced by the advance and retreat of continental
glaciers. The latest Quaternary receives the most emphasis be-
cause of the prevalence of data from that time interval. Unless
otherwise noted, this paper follows Russell (1895), Hubbs and
Miller (1948), and Morrison and Davis (1984a) in the use of
the same proper name for all successive pluvial lakes that occu-
pied each major basin.
Models of Great Basin Climate Change
ANTEVS' JET STREAM MIGRATION MODEL FOR PLEISTOCENE
CLIMATE CHANGE
Near the end of the nineteenth century, Russell (1895) sur-
mised that the rise and fall of pluvial Lake Lahontan was syn-
chronous with the advance and retreat of continental glaciers in
North America. Antevs (1925, 1938, 1948, 1955) agreed with
this premise and advanced the idea into a more complete model
whereby pluvial episodes in the Great Basin resulted from the
southward deflection of the northern jet stream storm track
(and associated weather regimes) by North American continen-
tal ice sheets. Some of the details in Antevs' model have since
been proven to be wrong. For example, his absolute chronolo-
gies of climate change were in error principally because his
dating methods were limited to long distance correlations of
discontinuous varve sequences and to salt balance calculations.
In addition, he was overly attached to the assumption of pre-
cisely synchronous climate change in Europe and North Amer-
ica (Antevs, 1955:324; 1962; Frye et al., 1965:57). Neverthe-
less, Antevs' model linking ice-sheet extent to pluvial lake
sizes has, to a first order, stood the test of time and, as a result,
has become the prevailing model for research on high-ampli-
tude Quaternary climate change in the Great Basin (e.g., Ben-
son and Thompson, 1987a; Benson et al., 1990; Morrison,
1991a; Spaulding, 1991; Benson etal., 1995).
IMPROVEMENTS TO ANTEVS' PLEISTOCENE MODEL
Estimating the Size of the Ice Sheets through Time
DIRECT EVIDENCE.?The timing of lake-level change pre-
dicted by Antevs' model is, of course, dependent upon the size
of the continental glaciers through time. With the development
of additional absolute dating techniques and the inclusion of
additional data, constraints on the extent of North American
continental glaciers through time have improved considerably
since Antevs' 1948 paper (e.g., Crandell, 1965; Frye et al.,
1965; Goldthwait et al., 1965; Lemke et al., 1965; Muller,
1965; Pewe et al., 1965; Schafer and Hartshorn, 1965; Wayne
and Zumberge, 1965; Wright and Ruhe, 1965; Hamilton and
Thorson, 1983; Mickelson et al., 1983; Waitt and Thorson,
1983; Andrews, 1987; Booth, 1987). This improved knowledge
base has led to more robust testing of Antevs' jet stream migra-
tion model. Unfortunately, even these advances cannot make
up for the obvious fact that most of the direct evidence for the
extent of the North American continental glaciers during the
Quaternary period has been erased by the last glacial advance.
As a result, abundant direct evidence of glacial size is only
available for the last maximum and subsequent retreat posi-
tions.
INDIRECT EVIDENCE.?One of most important advances in
Quaternary climate models during the last 50 years was the dis-
covery that the relative concentration of the oxygen isotopes in
the shells of marine microfossils provides a good first-order
proxy for global ice volume for the entire Quaternary period
(Shackleton, 1967; Imbrie and Kipp, 1971; Imbrie and Imbrie,
1979). This concentration is conventionally expressed as 818O,
the difference between a ratio of 18O and 16O in a given sample
and that same ratio observed in a standard sample. Of particular
importance to Antevs' model and to the Great Basin is the use-
fulness of this proxy to represent the size of the North Ameri-
can ice sheets through time. Indeed, despite low-amplitude,
high-frequency discrepancies and an approximately 1000-year
time lag, the decrease of 518O in marine microfossils has been
shown to occur more or less synchronously with the 14-6 Ka
NUMBER 33 13
Lake
Fort Rock
Lake
Malheur
Lake
Modoc Lake
Buck Lake
Round Lake
Malheur
Lake
Laki
Chewauc
Oregon_
Nevada
N
Lake
Lahontan
Walker
Lake
FIGURE 1.?Pluvial lakes of the northwestern Great Basin. Lake sizes during pluvial maxima are shown in gray.
Modern remnant lakes are shown in black. The names of pluvial lakes discussed in this paper are shown in itali-
cized print with large font size; the names of relevant modern remnants are indicated with smaller font.
retreat of the North American ice sheets (Mix, 1987; Ruddiman,
1987). Considering that many of the radiocarbon ages used to
date both the marine records and the terrestrial limits of the ice
sheets were from entirely different materials exposed to radi-
cally different geochemical environments, the timing of the two
records agrees quite well. Furthermore, there is no reason to be-
lieve that this result is not representative of previous glaciations.
Another advantage in using marine isotope records as a
proxy for ice-sheet size is that their ages can be determined
with an accuracy, precision, and continuity previously un-
known for middle- to late-Pleistocene sedimentary records.
The dating method used assumes that a large component of cli-
mate change is driven by variations in solar insolation incident
on the earth's northern hemisphere, which in turn are driven by
variations in the earth's orbital geometry, the so-called Milank-
ovitch cycles (Hays et al., 1976; Imbrie et al., 1984; Martinson
et al., 1987; Berger et al., 1992). These variations consist of the
precession of the earth's axis of rotation about the ecliptic nor-
mal (an ~21 Ky period), the oscillation of the axial obliquity
angle (an ~41 Ky period), and the oscillation of the eccentricity
of the orbit (an ~100 Ky period). Marine oxygen isotope
records are dated by tuning their phase and amplitude spectra
14 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
to those of the Milankovitch cycles. This analysis includes the
expected lag in the melting of the ice sheets and the associated
lowering of marine 518O. Because the timing of Milankovitch
cycles can be accurately and precisely extrapolated backward
in time for hundreds of thousands of years from modern and
historical astronomical observations, the age control for the ox-
ygen isotope global ice volume proxy is arguably precise to
within a few thousand years to at least as far back as 300 Ka
(Martinson et al., 1987). There is strong support for the accu-
racy of the marine oxygen isotope chronology well back into
the early Quaternary (Johnson, 1982; Baksi et al., 1992; Izett
and Obradovich, 1994; Edwards et al., 1997; Kerr, 1997). It
should be noted that terrestrial temperature proxies from ice
cores and vein calcites, which should respond more quickly to
insolation variations than the marine oxygen isotope records, at
times agree very precisely with the marine records but at other
times are quite discordant (e.g., Winograd et al., 1992; Wino-
grad et al., 1997). This inconsistency calls into question the
constancy of the lag times used in the orbital tuning of the ma-
rine oxygen isotope chronology and, thus, the accuracy of this
chronology.
An abrupt increase in mean 518O, a corresponding decrease
in the mean percentage of CaCO3, and a coeval increase in
variability of both of these indices in marine sediments point to
the onset of ice sheets in the northern hemisphere at about 2.5
Ma (Ruddiman and Wright, 1987, and references therein).
These early ice sheets apparently fluctuated in size at obliquity-
like periods of 41 Ky from -2.5 Ma until -900 Ka (Ruddiman
et al., 1986; Ruddiman and Wright, 1987; Ruddiman and
Raymo, 1988). From 900 Ka to the present, the 23 Ky preces-
sion-like period became as important as the 41 Ky period, and
both became subordinate to the 100 Ky eccentricity-like period
(Imbrie, 1985; Ruddiman and Wright, 1987). The above-de-
scribed character of the marine-sediment proxies for ice-sheet
extent can be used in conjunction with Antevs' model to pre-
dict similar features in records of pluvial lake-level sizes
through time. Furthermore, assuming that the lake-level data
have the requisite amplitude resolution and age control, finer-
scale features in the marine proxies, such as the oxygen isotope
stage (OIS) and the substage boundaries of Hays et al. (1976),
also can conceivably be employed as predictive tools.
Quantifying the Jet Stream Model
ATMOSPHERIC CIRCULATION MODELS.?Spurred by ad-
vances in computer technology during the last 25 years, An-
tevs' model has been expressed quantitatively as part of a more
general effort to model global atmospheric circulation in re-
sponse to a variety of factors (Kutzbach, 1987; COHMAP,
1988; Kutzbach and Ruddiman, 1993). From such models, jet-
stream position versus time has been predicted in response to
changing global climatic conditions for the last 18 Ky. The
models are based upon the fundamental laws of physics and are
constrained by a myriad of boundary conditions, including the
surface area and elevation of continental ice sheets, latitude-
and time-dependent Milankovitch-driven changes in solar inso-
lation, atmospheric chemistry (e.g., the concentration of CO2),
sea-surface temperatures, albedo of the earth's surface, etc.
Such efforts have shown that, to a first order, significant
changes in precipitation, temperature, and weather distribution
patterns should, indeed, occur in response to the advance of the
continental ice sheets as proposed in Antevs' original model
for the Great Basin (Kutzbach and Wright, 1985; Manabe and
Broccoli, 1985; Kutzbach and Guetter, 1986; Kutzbach, 1987;
COHMAP, 1988).
Although the results of the above-described models gener-
ally support Antevs' jet stream migration model, the spatial and
temporal grid spacings in the models are relatively large com-
pared with the spatial and temporal resolution and the accuracy
of the observational database. That is, the general atmospheric
circulation models do not generate specific-enough predictions
regarding the effect of the jet stream on the magnitude and tim-
ing of pluvial lake size in a region as small as the northwestern
Great Basin. For example, such models show a maximum
southward mean position of a strong winter jet stream at
~25?-35? N from before 18 Ka to as late as 15-12 Ka (Kutz-
bach, 1987, fig. 11; Thompson et al., 1993). In contrast, north-
western Great Basin lakes are spaced only tens to hundreds of
kilometers apart, and the lake-level prehistory of at least one of
them (Lake Lahontan) is known throughout this time interval
with a precision as small as a few hundred years (see "Latest
Pleistocene," below). As knowledge of the lake-level prehisto-
ries for the other lakes approaches that for Lake Lahontan,
higher-resolution models will be required.
EFFECT OF JET STREAM MIGRATION ON LAKES.?More-lo-
calized modeling has been conducted to determine if predicted
climate conditions would result in lake-size changes that agree
with observations (Mifflin and Wheat, 1979; Benson, 1986;
Benson and Paillet, 1989; Hostetler and Bartlein, 1990;
Hostetler and Benson, 1990; Hostetler, 1991; Freidel, 1993).
Input for these models includes the general climate conditions
dictated by the output of the larger-scale models, independent
observations of pluvial climate conditions (e.g., temperature
and precipitation estimates from fossil pollen analyses), or ex-
treme values from modern measurements (e.g., river discharge
values from the 1983 wet year). Such efforts have demon-
strated that pluvial lakes can indeed reach their maximum sizes
under the conditions corresponding to a mean annual position
of the jet stream storm track near or south of the latitude of the
lake in question in response to decreased evaporation from
lower temperatures and increased cloud cover and in response
to enhanced precipitation in the drainage basins. The amount of
required increase in precipitation is reduced significantly by
greatly reduced evaporation if the temperature decreases
enough to freeze the lakes during the winter.
NUMBER 33 15
DANSGAARD-OESCHGER CYCLES, HEINRICH EVENTS, AND
BOND CYCLES AS MODELS FOR HIGH-AMPLITUDE
PLEISTOCENE CLIMATE CHANGE
In the last two decades, high-frequency climate events and
cycles not explainable by Milankovitch cycles have been ob-
served in several different types of climate records, especially
high-resolution ice cores and marine sediment cores (Broecker,
1994). In particular, Dansgaard-Oeschger (D-O) cycles are
variations in several climate parameters and were first ob-
served in Greenland ice cores (Dansgaard et al., 1982; Dans-
gaard et al., 1993). They have a distinctive square-wave mor-
phology and periods of a few thousand years. These cycles
usually oscillate about a progressively cooler temperature
mean until a relatively instantaneous Heinrich event occurs.
Heinrich events are marked by distinct marine sedimentary de-
posits that indicate extensive ice rafting from North America
into the North Atlantic Ocean (Heinrich, 1988). These litho-
logic deposits are coeval with the coldest of the cold parts of
the D-0 cycles since the last Heinrich event. A rapid warming
usually follows, marking the beginning of a new sequence of
D-0 cycles that, in turn, end in the next Heinrich event (Bond
et al., 1993). Each set of D-0 cycles and its associated terminal
Heinrich event is collectively referred to as a Bond cycle (Bro-
ecker, 1994). Within Bond cycles, each cold part of the higher-
frequency D-0 cycles also are correlated with high percentages
of lithic detritus presumably rafted on icebergs (Bond and
Lotti, 1995).
The balance of evidence suggests the above phenomena are
global in nature. Heinrich events, in fact, appear to have oc-
curred near major global climate boundaries (e.g., at the transi-
tion from OIS 6 to 5, from 5 to 4, and from 2 to 1 (Broecker,
1994). The most-favored mechanisms for connecting Heinrich
events (and now the cool parts of the D-0 cycles) with global
climate change are (1) the impact of fresh water into the North
Atlantic Ocean on the marine thermohaline circulation and re-
lated atmospheric/oceanic transfer of heat and water vapor,
and (2) the impact on the position of prevailing weather pat-
terns caused by a rapid retreat of the Laurentide ice sheet (Bro-
ecker, 1994; Bond and Lotti, 1995). Although the causes of D-
O cycles, Heinrich events, and Bond cycles have not been iden-
tified unequivocally, their widespread nature and distinct tim-
ing have led to their use as potential phenomenological models
for high-frequency climate change in the Great Basin (e.g.,
Oviatt, 1997; Benson et al., 1998).
ANTEVS' NEOTHERMAL MODEL FOR HOLOCENE
CLIMATE CHANGE
The understanding of Holocene climate change is important
for several reasons, including its bearing on the distribution and
nature of archeological sites (Baumhoff and Heizer, 1965; Aik-
ens, 1983; Mehringer, 1986; Madsen, 2002) and its importance
in establishing a baseline for anthropogenic climate change
(Scuderi, 1993). Unfortunately, for the most part, models of
Holocene climate change are still in the inductive stage. They
are either entirely descriptive or consist of time-varying global
phenomena that have only speculative, generalized links to cli-
mate. Part of the problem for the northwestern Great Basin lies
in the unremarkable nature of the observations relative to those
in the latest Pleistocene. Holocene pluvial lake-level changes in
this region either are very low in magnitude or, if high-magni-
tude changes do exist, are of too high a frequency (< 100-year
periods) to be easily observed in the geological record. Despite
these problems, the existing models of Holocene climate
change are worth reviewing as points of departure.
Antevs (1948, 1955) based his model for Holocene (or, in his
terms, "Neothermal") climate change in the Great Basin on ob-
servational evidence from North America and Sweden and on
the assumption of synchronous climate change for the entire
Northern Hemisphere (Grayson, 1993). Loosely interpreted,
Antevs divided the Holocene into three intervals of roughly
equal duration, as shown in Figure 2a. The early Holocene
Anathermal age (10-7.5 calendar thousand years ago (cal Ka);
8.4-6.6 radiocarbon thousand years before present (14C Ky
B.P.)) was characterized by modern-like temperatures that grad-
ually increased throughout the interval and by aridity levels
that were slightly moister than today. The middle Holocene
Altithermal age (7.5-4.5 cal Ka; 6.6-4.0 14C Ky B.P.) was dis-
tinctly warmer than present, and it was characterized by higher
aridity. In the late Holocene Medithermal age (4.5 cal Ka to
present), conditions returned to cooler temperatures and less
aridity, although they were slightly more arid than the Anather-
mal. Quantitative modeling has not yet been attempted to deter-
mine the extent of lake-level rise and fall in response to Ho-
locene climate change; however, assuming Antevs' model is
correct, it is reasonable to extrapolate from this model that
northwestern Great Basin pluvial lakes would have been at late
Holocene levels or slightly higher during the early Holocene
and at their lowest levels during the middle Holocene.
As with Antevs' jet stream migration model, the particulars
of his Neothermal model have fallen prey to additional data
that have been analyzed with more modern techniques. For ex-
ample, his attribution of moister conditions for the Anathermal
age is based in part upon several meters of lake sediments
found above the Mazama tephra in an outcrop that is 25 m
above the present-day lake level of Summer Lake, Oregon (Al-
lison, 1945; Antevs, 1948). After the advent of geochemical
fingerprinting of tephra layers and better absolute dating tech-
niques, it became clear that Allison's tephra was, instead, the
considerably older (~45-50 Ky old) Mount St. Helens Cy
tephra (Allison, 1966a; Davis, 1985; Berger and Davis, 1992).
Generally speaking, however, when data are averaged tempo-
rally over hundreds of years and spatially over the entire north-
western Great Basin, Antevs' original hunch has proven to be
more or less correct, although a strict adherence to this model
is commonly advised against, mostly because of its limited ap-
plicability to very specific localities and times (Bryan and
16 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Medithermal Altithermal Anathermal
800-
jQQ.
Q.
CO
?JO 400-
o
0J
-34-
early Holocene
cooling event
luminosity changes
and higher overall
solar luminosity
i i i
10 11 12
cal Kyr B.P.
FIGURE 2.?a, Approximate age range of Antevs' Neothermal ages, b, Frequency and severity of Holocene erup-
tions as indicated by residual sulfate anomalies from the GISP2 ice core (after Zielinski et al., 1994, fig. 4). c,
Holocene temperature as indicated by oxygen isotope data from the GISP2 ice core (after Stuiver et al., 1995, fig.
5). d, Solar luminosity throughout the middle and late Holocene as inferred from 14C concentration anomalies
corrected for variations in magnetic field intensity (after Damon et al., 1989, fig. 4; Magny, 1993, fig. 3).
NUMBER 33 17
Gruhn, 1964; Mehringer, 1986; Grayson, 1993; Wigand and
Rhode, 2002).
Aside from the issue of its applicability, Antevs' model is
less than satisfying because it is entirely descriptive in nature.
That is, there is no cause attributed to the transitions between,
and the durations of, Antevs' three Neothermal ages. Toward
this end, Antevs' model appears to be consistent, at least in a
qualitative sense, with his jet stream migration model until the
end of the middle Holocene. That is, the moist, not-quite-hot
Anathermal could be considered to be related to slightly more
frequent storms in the northern Great Basin if the storm track
had not quite retreated to its modern position associated with
ice-free conditions. The Altithermal, which reached its peak
hot, dry conditions 5.5-6.0 cal Ka, corresponds quite well with
the final disappearance of the North American ice sheets. The
return of moderately warm, arid to semiarid conditions in the
Medithermal age after 4.5 cal Ka is not explained by this model
because the North American ice sheets have not returned.
Results from quantitative modeling, however, indicate that
the northern jet stream storm track would have reached its
modern position and intensity (or lack thereof) by 9 Ka or ear-
lier, which is well before the peak of the Altithermal age (Kutz-
bach, 1983, 1987). Thus, subsequent reductions in the size of
an already insignificant ice sheet should not have affected
Great Basin climate. On the other hand, although atmospheric
circulation models also take into account the more subtle shift
in weather patterns caused solely by Milankovitch-driven
changes in insolation in the absence of ice sheets, the low-am-
plitude changes in precipitation and temperature that result di-
rectly from such insolation changes are usually below the sen-
sitivity level of these models (Kutzbach, 1987). On a more
localized scale, Benson (1986) has shown that changes in the
evaporation rate that were induced solely by insolation changes
were an order of magnitude lower than those needed to fill plu-
vial Lake Lahontan in conjunction with increased discharge.
Thus, if insolation variations were nonetheless responsible for
the subtle changes in lake level observed in the northwestern
Great Basin during the Holocene, then they must have acted in
the absence of the other factors that Benson found to be much
more important (e.g., the difference between air and water-sur-
face temperatures, the type of clouds, and the degree of cloudi-
ness). In any case, the peak value of summer insolation at the
latitude of the northwestern Great Basin occurred at ~11 Ka,
well before the onset of the middle Holocene Alithermal age.
Again, timing is a problem with regard to nondescriptive mod-
eling of Holocene climate in the northwestern Great Basin.
OTHER MODELS FOR HOLOCENE CLIMATE CHANGE
Volcanic Eruptions
Large volcanic eruptions significantly increase the concen-
tration of atmospheric aerosols and, as a result, decrease global
temperatures slightly (<1.5? C) for a few years until the aero-
sols settle out (Stuiver et al., 1995). The 1991 eruption of
Mount Pinatubo lowered the earth's mean temperature by as
much as ~0.5? C for a year or two, as determined both by direct
observations and by models that postulated the effects on the
atmosphere caused by the observed increase in aerosol concen-
tration (Hansen et al., 1992; Houghton et al., 1996, fig. 14).
The effects of larger eruptions on climate (e.g., Tambora in
1815), as inferred from increased levels of acidity in Greenland
ice cores, do not seem to last much longer (Hammer et al.,
1980, fig. 2).
Despite the short-term nature of the effects of single erup-
tions on climate, some time intervals containing frequent, large
eruptions may coincide with intervals characterized by lower
temperatures over much of the earth. For example, Hammer et
al. (1980, fig. 5) demonstrated longer-term (decade to century)
correlations between periods of lowered Northern Hemisphere
temperature (by 0.5?-1.0? C) and intensified and/or more fre-
quent volcanism. On the millennial scale, Stuiver et al. (1995)
used 818O-derived temperature estimates from the GISP2 ice
core to argue that volcanism caused no more than a 0.1? C av-
erage decrease in temperature during the last 1000 years.
The last 1000 years, however, may not be representative of
the entire Holocene. Zielinski et al. (1994) presented an in-
triguing record of volcanic eruptions for the past 9000 years,
based upon their analysis of sulfate concentrations in the
GISP2 ice core (Figure 2b). They found that extremely large
eruptions occurred more frequently during the early Holocene
(-9.0-7.0 cal Ka) as opposed to the remainder of the Holocene.
These authors tentatively concluded that this increased volcan-
ism contributed to observed climatic cooling in the Northern
Hemisphere during a time interval that otherwise would have
been warmer because of higher summer insolation values in the
Northern Hemisphere. Thus, Holocene climate change could
conceivably be explained by the combined effects of volcanism
and insolation variations. That is, the highest insolation values
in the early Holocene could have been counteracted by the ef-
fects of anomalous volcanism. The middle Holocene (Antevs'
Altithermal age) would have been a time of higher tempera-
tures driven by still-high insolation values, but the cooling ef-
fect of the anomalous early Holocene volcanism would have
ended. As Northern Hemisphere insolation values decreased
steadily to present values, cooler temperatures would have pre-
vailed in the late Holocene. This combined model is appealing,
but it remains to be shown that the larger, more frequent erup-
tions of the early Holocene would result in a significantly
larger temperature anomaly than the mere 0.1? C effect deter-
mined by Stuiver et al. (1995) for the last 1000 years. It also
must be shown that the increased level of sulfate in the GISP2
ice core during the early Holocene was caused by globally sig-
nificant eruptions rather than by local events in the vicinity of
Greenland (Zielinski et al., 1994).
18 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Solar Variability
Low-amplitude temperature change during the Holocene
also has been attributed to variations in solar insolation caused,
in turn, by inherent variability in the sun's luminosity on the or-
der of 0.1% (Stuiver, 1965, 1980; Denton and Karlen, 1973;
Reid, 1987; Stuiver et al., 1991; Stuiver et al., 1995). This
causal relationship is most directly based upon correlations of
the timing of various global climate estimates during the last
century with both the number of sunspots (Eddy, 1977; Reid,
1987) and the period of the sunspot cycle (Friis-Christensen
and Lassen, 1991). Such variations in solar luminosity can ac-
count for an increase of as much as 0.25? C, which is approxi-
mately 50% of the long-term temperature increase observed
since the middle of the nineteenth century. These values are
based upon recent studies that modeled the effect on global
temperature caused by the relative contributions of heat input
from changes in solar luminosity and from changes in the con-
centrations of greenhouse gasses and aerosols (Kelly and Wig-
ley, 1992; Lacis and Carlson, 1992; Schlesinger and Raman-
kutty, 1992).
Proxy indices have been generated to extend estimates of
changes in solar activity back through the last several tens of
thousands of years. These indices are based upon observed cor-
relations of lower solar luminosity with higher concentrations
of cosmogenic radioisotopes (Stuiver, 1965; Stuiver and Quay,
1980; Damon et al., 1989; Stuiver et al, 1991; Stuiver et al.,
1997). One such proxy index (Figure Id) is based upon A14C,
which is a ratio expressing the difference between measured ra-
diocarbon activity and the expected activity of a sample of a
particular age (Stuiver and Pollach, 1977; Sonett et al., 1990;
Stuiver et al., 1991). The data shown in Figure Id are residual
values of A14C after the removal of variations caused by well-
documented changes in the intensity of the earth's magnetic
field (Damon et al., 1989).
Such proxies for solar luminosity have been compared with
various estimates of climate change throughout the Holocene
(e.g., glacier extent around the world and lake-level fluctua-
tions in Europe), and correlations have been found corre-
sponding to periods of duration of-90, 200, and 2200-2300
years (Damon and Jirikowic, 1992; Magny, 1993; Stuiver and
Braziunas, 1993; O'Brien et al., 1995; Stuiver et al., 1995).
The largest magnitude of temperature change at these peri-
ods, as inferred from Holocene variations of 518O in the
GISP2 ice core (0.4? C; Stuiver et al., 1995), is reasonably
close to the previously mentioned modeled results for solar-
induced temperature change during the last century or so (see
above).
Assuming that changes in the solar constant during the Ho-
locene really do contribute to the A14C variations shown in Fig-
ure Id, one could conceivably use this data to predict the tim-
ing and duration of the relatively subtle Holocene climate
changes observed in the Great Basin. For example, the overall
low values of A14C from 7.0 to 3.5 cal Ka would indicate an
extended period of higher solar luminosity during a time period
that, for the most part, corresponds with Antevs' Altithermal
age. The same time period also was characterized by higher-
frequency variations in luminosity, suggesting that one might
expect to find high-frequency climate changes in the middle
Holocene superimposed on an overall warmer climate. Finally,
relatively high values of A14C shown in Figure 2t/at time inter-
vals of 6.0-5.2, 3.4-1.7, and 0.6-0.2 cal Ka predict low values
of solar luminosity at those times. Magny (1993) noted that
cooler, moister climate in the European Alps indeed occurred
at approximately those times, as evidenced by three dominant
intervals of lake-level transgressions during 6.5-4.5, 3.5-2.0,
and 0.60-0 cal Ka. Furthermore, the solar-luminosity proxy
and an earlier period (8.7-8.0 cal Ka) of transgressions noted
by Magny (1993) correlate reasonably well with the low-am-
plitude, century- to millennial-scale, global cooling events
(8.8-7.8, 6.1-5.0, 3.1-2.4, and 0.6-0 cal Ka) recognized both
in glaciochemical time series from the GISP2 ice core and in
previously published, coeval paleoclimate indicators from
around the globe (O'Brien et al., 1995). This set of observa-
tions suggests that solar variability is worthy of further investi-
gation as a potential cause of low-amplitude climate change in
the Holocene.
A Short-Term, High-Amplitude Cooling Event
in the Early Holocene
A spike-like, early Holocene cooling event centered at
-8.4-8.0 cal Ka (e.g., Figure 2c) is recognized in several cli-
mate records from around the world (Alley et al., 1997) and
may be potentially observable in pluvial lake records from the
northwestern Great Basin. The geographic distribution of the
type of climate change observed at each locality is similar to
that of the terminal Pleistocene, Younger Dryas event (Alley et
al., 1997, and references therein). Furthermore, the amplitudes
of the anomalies in several climate proxies for the early Ho-
locene cooling event are comparable (approximately one-half)
to those of the same proxies for the Younger Dryas event. For
example, a temperature decrease in Greenland of about 6? C
was implied for the early Holocene event by a 2%o decrease in
518O in the GISP2 ice core (Alley et al., 1997). The Younger
Dryas event has often been attributed to diminished thermoha-
line circulation induced by a sudden influx of fresh water into
the North Atlantic Ocean (Broecker et al., 1990; Alley et al.,
1997), suggesting that a similar influx caused the early Ho-
locene event; however, Alley et al. (1997) noted that no such
influx of fresh water is evident in the geologic record during
this time interval. Alternatively, this cooling event is coincident
with the time interval containing four closely spaced, very
large volcanic eruptions, as inferred by the residual sulfate re-
cord from the GISP2 ice core (Zielinski et al., 1994). It is
tempting to draw a causal relationship between the two phe-
nomena, but, again, it has yet to be shown that large eruptions
disrupt climate for periods significantly greater than a couple
of years (see "Volcanic Eruptions," above).
NUMBER 33 19
Pluvial Lake Sizes through Time
in the Northwestern Great Basin
METHODS
During the last 50 years, pluvial lake studies in the north-
western Great Basin have benefited from an increasingly di-
verse array of new or improved methods with which to deter-
mine the timing and magnitude of lake-size changes. These
methods include radiometric and luminescence dating
(Rosholt et al., 1991; Berger, 1988), petrographic and
geochemical fingerprinting of tephra layers (Sarna-Wojcicki
and Davis, 1991), paleoecology of lacustral biota (Forester,
1987; Bradbury, 1991, 1992, 1997), trace element and isotope
geochemistry of lake sediments (Stuiver, 1970; De Dekker and
Forester, 1988; Benson and Peterman, 1995; Benson et al.,
1996), particle-size analysis of sediments (McCave and
Syvitski, 1992), paleomagnetic-polarity stratigraphy and secu-
lar-variation dating as applied to lacustrine sediments and vol-
canic tephra (Reynolds, 1977, 1979; Verosub, 1988, Rosholt et
al., 1991; Negrini and Davis, 1992), environmental magnetism
(Thompson and Oldfield, 1986; Verosub and Roberts, 1995),
improvements in sedimentological and geomorphological
models of lacustrine systems in arid environments (Currey,
1994a, 1994b), and the empirical and theoretical analyses of
post-depositional deformation of basins in this region (Pezzo-
pane, 1993; Bills, unpublished).
The age estimates of several key northwestern Great Basin
tephra layers have been improved since the last review paper
was written on northwestern Great Basin lakes, and previous
reviews (e.g., Davis, 1982; Smith and Street-Perrot, 1983;
Benson and Thompson, 1987a; Morrison, 1991a, 1991b; Gray-
son, 1993) have not covered the use of paleomagnetic secular
variation as a dating tool in the northwestern Great Basin.
These two topics will be addressed below.
Recent Work on the Ages of Key Tephra Beds
in the Northwestern Great Basin
EARLY PLEISTOCENE TEPHRA.?The ages of a few early
Pleistocene tephra layers have been revised recently to slightly
older ages, in part as a result of improved technology of the
laser-fusion 40Ar/39Ar method (Table 1). For example, the
ages of the Bishop and Lava Creek B tephra layers (and, by as-
sociation, the Rye Patch Dam) have recently been revised
downward by an average of a few percent to ~760 ? 2 Ky
(Sarna-Wojcicki and Pringle, 1992; Izett and Obradovich,
1994) and 665 ? 10 Ky (Izett et al., 1992), respectively. The
same technique has been used to revise downward the age of
the Brunhes/Matuyama geomagnetic polarity reversal to agree
better with that predicted by marine 818O stratigraphy (e.g.,
Johnson, 1982; Baksi et al., 1992; Izett and Obradovich, 1994).
MIDDLE PLEISTOCENE TEPHRA.?Several middle Pleis-
tocene tephra layers have ages that are poorly constrained be-
cause they are either too old for the easy acquisition of precise
luminescence or radiocarbon ages, too young, or of the wrong
chemical composition, mineralogy, and/or grain size for the
easy acquisition of precise fission-track or potassium-argon
ages on distal volcanic ashes (Naeser and Naeser, 1987;
Rosholt et al., 1991; Sarna-Wojcicki and Davis, 1991). As a
result, many ashes in this age range have age estimates with
TABLE 1.?Tephra layers, from youngest to oldest, found in the pluvial lakes of the northwestern Great Basin.
Tephra
Tarupah Flat
Mazama
Tsoyawata
Mount St. Helens Mp
Trego Hot Springs
Wono
Mount St. Helens Cy
Olema
Pumice Castle Set
Summer Lake 2
Summer Lake V
Wadsworth
Summer Lake GG
Summer Lake KK
Rockland
Dibekulewe
Lava Creek B
Rye Patch Dam
Glass Mt. D
Glass Mt. G
Rio Dell
Bear Gulch
Pluvial lakes
Lahontan
Lahontan, Chewaucan
Lahontan, Modoc
Lahontan, Chewaucan
Lahontan, Chewaucan, Modoc
Lahontan, Chewaucan
Lahontan, Chewaucan
Chewaucan, Modoc
Chewaucan, Modoc, Malheur
Chewaucan
Chewaucan, Modoc
Lahontan, Modoc
Chewaucan
Chewaucan, Modoc
Lahontan, Modoc
Lahontan, Modoc
Lahontan, Modoc
Lahontan, Modoc
Lahontan
Lahontan
Modoc
Modoc
Age
600-1500 calyrB.P.
6.845 ?0.05 14CKyB.P.
7.015 ?0.045 14CKyB.P.
20.5 ? -0.5 14CKyB.P.
23.2?0.3 14CKyB.P.
27.3 ?0.3 14CKyB.P.
45-50 Ky
50-55 Ky
72?6Ky
67.3 ? 7.2 Ky
155-200 Ky
155-200 Ky
~190?~30Ky
~200?~30Ky
410 Ky
-510 Ky
665?10Ky
-675 Ky
-0.9-1.0 My
-1.0-1.1 My
1.45 My
1.94 My
References
Davis (1978)
Davis (1978), Bacon (1983); Davis (1985)
Davis (1978), Bacon (1983); Reick et al. (1992)
Davis (1978), Davis (1985); Negrini et al. (1988); Negrini and Davis (1992)
Davis (1978), Davis (1985); Reick et al. (1992); Benson et al. (1997)
Davis (1978), Davis (1985); Benson et al. (1997)
Davis (1978), Davis (1985); Berger and Busacca (1995)
Reick et al. (1992); this paper
Bacon (1983); Davis (1985); Reick et al. (1992); Botkin and Carambelas
(1992a, 1992b); this paper
Davis (1985); Berger (1991)
Davis (1985); Negrini et al. (1994)
Davis (1978), Reick et al. (1991)
Davis (1985); Herrero-Bervera et al. (1994); Negrini et al. (1994)
Davis (1985); Berger (1991); Herrero-Bervera et al. (1994)
Sarna-Wojcicki et al. (1985); Sarna-Wojcicki et al. (1991)
Davis (1978); Negrini, Verosub et al. (1987); Sarna-Wojcicki et al. (1997)
Davis (1978); Reick et al. (1991); Izett et al. (1992)
Davis (1978); Reick et al. (1991); this paper
Sarna-Wojcicki et al. (1991); Parker (1993)
Morrison and Davis (1984a); Sarna-Wojcicki et al. (1991)
Izett (1981); Sarna-Wojcicki et al. (1987); Sarna-Wojcicki et al. (1991)
Sarna-Wojcicki et al. (1991)
20 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
0
5-
Q.(D
Q
10-
15-
? I
Mt St Helens Mp(C-14)
Trego Hot Springs (C-14, tl_)
-Wono(C-14)
MtSt Helens Cy(TL)
Pumice Castle (K-Ar)
N(TL) Prominent Unconformity
GG ("excursion" age) GG (Ar-Ar)
0 100 200
Age (Ky)
FIGURE 3.?Chronology of Lake Chewaucan sediments exposed in the Ana River canyon north of Summer
Lake, Oregon. Age control is based upon a variety of methods that are summarized in Negrini et al. (1994) and in
this paper. Tephra names are after Davis (1985). Note the pronounced decrease in precision for dates older than
150 Ky. Shaded region includes all age estimates as allowed by standard deviations of these ages.
NUMBER 33 21
uncertainties of 5%-20%. This problem is exemplified by the
tephra beds from the pluvial Lake Chewaucan sediments ex-
posed in the Ana River canyon near the town of Summer Lake,
Oregon (Figure 1). As shown in Figure 3, age uncertainties of
tephra from lower in the section are much greater than those
from the upper part of the section (see Negrini et al., 1994, for
a review of the age control at this section). The large scatter in
tephra ages is reflected in the age-range (-180-230 Ky) as-
signed to the Pringle Falls/Biwa I geomagnetic excursion.
This is a geomagnetic phenomenon recorded in several sedi-
mentary records, including two that contain Summer Lake
tephra layer GG at the Ana River locality (Negrini et al., 1994)
and two that contain the correlative tephra layer D at the Prin-
gle Falls, Oregon, locality 170 km to the northwest (Herrero-
Bervera et al., 1994). McWilliams (1995) has suggested that
the true age of the excursion is toward the older part of this
range, based upon the similarity of its transitional virtual geo-
magnetic pole (VGP1) path with that of an excursion found in
223 ? 4 Ky old volcanic rocks from New Zealand. Conversely,
a correlation of the Pringle Falls event with an excursion found
in a marine sediment record from the Blake/Bahama Outer
Ridge in the northwestern Atlantic Ocean (Henyey et al.,
1995) suggests that the Pringle Falls excursion correlates with
the 190 ? 10 Ky old Jamaica event found throughout the world
in marine sediments (e.g., Tauxe and Shackleton, 1994;
Yamazaki and Ioka, 1994; Henyey et al., 1995; Weeks et al.,
1 VGPs are magnetic pole positions pointed to by each paleomagnetic vector
recorded at a particular site. Because the earth's magnetic field is not a perfect
dipole, coeval VGPs recorded at different sites will not necessarily coincide.
1995; Guyodo and Valet, 1996; Lehman et al., 1996). The cor-
relation with the New Zealand excursion is problematical be-
cause the VGP path of the Pringle Falls excursion has previ-
ously been shown to have a common morphology with the
VGP paths of several different polarity transitions (Trie et al.,
1991). The correlation with the Jamaica event is more con-
vincing because it is based upon distinct paleomagnetic direc-
tional and intensity features found in both the excursion itself
and in the subsequent secular variation (Negrini et al., 1994;
Henyey et al., 1995).
PUMICE CASTLE TEPHRA SET AND THE OLEMA AND MOUNT
ST. HELENS CY TEPHRA LAYERS.?The three tephra layers
comprising the Pumice Castle set were first identified by
Davis (1985) in lake sediments from a sequence of 54 water-
lain tephra beds in the Ana River canyon near Summer Lake,
Oregon (Figure 4). Based upon their major element composi-
tions, Davis correlated Summer Lake tephra layers 4, 6, and 8
with the Pumice Castle-like 2, the Pumice Castle, and the
Pumice Castle-like 1 tephra layers, respectively, from Crater
Lake, Oregon (Bacon, 1983; Davis, 1985). Negrini et al.
(1988) added support to this geochemical correlation by com-
paring the paleomagnetic direction of Summer Lake tephra 6
with an identical direction for the Pumice Castle tephra mea-
sured near the Crater Lake volcano by D. Champion of the
United States Geological Survey (USGS) at Menlo Park. The
absolute ages of the two older tephra are constrained by brack-
eting lavas that are dated by K-Ar at 117 ? 6 Ka and 72 ? 6 Ka;
the youngest tephra lies above the 72 ? 6 Ky old lava and be-
low a 52 ? 4 Ky old andesite flow, which also was dated with
Lake Chewaucan
sediments
Crater Lake
volcanics
Tephra 2 (67.3 ? 7.5 Ky)
poorly developed
soil on tills and
glaciated surfaces
-50 cm
? lava (52 ? 8.0 Ky)
? Pumice Castle-like 1
? lava (72 ? 6.0 Ky)
? Pumice Castle
? Pumice Castle-like 2
? lava (117 ?6.0 Ky)
FIGURE 4.?Stratigraphic relationships supporting ~70 Ky old age of the Pumice Castle tephra set (after Bacon,
1983; Davis, 1985). These three tephra layers are nearly the same age because they are found within 10 cm of
each other in fine-grained lacustrine sediments at several different localities separated by several kilometers in
the Summer Lake subbasin. Thus, they are approximately as old as the 72 + 6 Ky old lava flow that separates two
of them at Crater Lake, their proximal location. This age is supported by the thermoluminescence date deter-
mined for tephra 2 in the Lake Chewaucan sediments (Berger and Davis, 1992).
22 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
the K-Ar method (Bacon, 1983; C.R. Bacon and M.A. Lan-
phere of the USGS, in litt. to J.O. Davis, 1984; Davis, 1985).
Reick et al. (1992) reported finding the Pumice Castle-like 2
tephra, or perhaps the Pumice Castle tephra, in a core taken
from the sediments of pluvial Lake Modoc. (The core was
named Tulelake for the nearby town of Tulelake, California,
rather than nearby Tule Lake, a remnant of Lake Modoc.) They
determined this ash was deposited -120 Ka, in part because of
its position below two samples showing negative paleomag-
netic inclinations that were provisionally correlated to the 100
-130 Ka Blake Reversed-Polarity Subchron (Reick et al.,
1992:419, 422). This age was considered to be reasonably
close to the above-mentioned 117 ? 6 Ka K-Ar date on the lava
flow below these tephra.
Ongoing work on the bottom sediments of pluvial Lake
Chewaucan indicates that an age of ~70 Ky, near the middle of
OIS 4, is a more reasonable estimate for the Pumice Castle set
of tephra beds. The three ash layers in this set are observed to
be separated by only a few centimeters of lacustrine silty clays
in several exposures 1-2 km apart along the Ana River canyon
and in a core taken several kilometers away at an Ana River
outcrop (Davis, 1985; Negrini et al., 1988; Erbes, 1996). Be-
cause this relationship argues that the three ashes were depos-
ited within no more than a few hundred years of each other, and
because a 72 ? 6 Ky old lava flow is found between the correla-
tives of the upper two tephra found at Crater Lake (Figure 4),
all three tephra layers in the Pumice Castle set must be approx-
imately the same age as the 72 ? 6 Ky old lava flow. This con-
clusion is supported strongly by a thermoluminescence date of
67.3 ? 7.2 Ky on Summer Lake tephra 2 located 35 cm below
the Pumice Castle set (Berger and Davis, 1992).
The ages of the Mount St. Helens Cy tephra and its correla-
tives (e.g., the Marble Bluff bed) also have been the subject of
controversy. The ages of these tephra layers have most often
been quoted in the range of 35-30 14C Ky B.P. (e.g., Davis,
1978, 1985, and references therein) on the basis of radiocarbon
ages of materials found in close proximity to the tephra. In con-
trast, radiocarbon dating of some of the near-vent tephra depos-
its infers an older age for these tephra layers of perhaps greater
than 40 14C Ky B.P., which in turn, infers a calendar date of at
least 45-40 Ka (Mullineaux, 1986; Crandell, 1987; McDonald
and Busacca, 1992). Thermoluminescence dates on the Cy
tephra (46.3 ?-4.8 Ky) in lacustrine sediments at Summer Lake,
Oregon, and on the slightly older Cw tephra (-55 ? 5 Ky) in east-
ern Washington loess sediments support the older ages (Berger,
1991; Berger and Davis, 1992; Berger and Busacca, 1995).
The Olema tephra, originally found in cores from Clear Lake
and Tule Lake, California (Sarna-Wojcicki et al., 1988), also
was found in the Lake Chewaucan sediments of the WL core
(N. Foit, in litt., 1993; Erbes, 1996). Given its position relative
to the Pumice Castle set and Mount St. Helens Cy tephra layers
in the WL core and given the above-described revisions of the
ages of the latter tephra layers, the estimated age of the Olema
bed is probably closer to 55-50 Ky than to the previous estimate
of 75-55 Ky (Sarna-Wojcicki et al., 1988; Reick et al., 1992).
WONO AND TREGO HOT SPRINGS TEPHRA.?New ages for
the Wono tephra (27.3+0.3 14C Ky B.P.) and the Trego Hot
Springs bed (23.2?0.3 14C Ky B.P.) have been published by
Benson et al. (1997). These new ages are based upon 11 precise
accelerator mass spectrometer (AMS) radiocarbon dates from
adjacent organic samples in a core from Pyramid Lake.
Paleomagnetic Secular Variations for Correlating
Fine-Grained, Lacustrine Sediments
Paleomagnetic secular variation (PSV) records from fine-
grained sediments must satisfy strict criteria in order to estab-
lish that they record the ambient geomagnetic field vector (dec-
lination, inclination, and paleointensity) at or near the time of
deposition (Verosub, 1988). This is especially true of the paleo-
intensity component (Tauxe, 1993). If these criteria are satis-
fied, PSV records can be used as a method for high-resolution
(century- to millennial-scale) correlation among fine-grained
sediments of Quaternary age over a spatial range of continental
scale. Anchored by a solid foundation of PSV reference curves
for the last 30 Ky (Lund et al., 1988; Lund, 1996, and refer-
ences therein), this methodology has been used on many sedi-
ments of the Great Basin (Negrini and Davis, 1992; Ewing,
1996; Liddicoat and Coe, 1997; Benson et al., 1998; Benson et
al., unpublished). The potential value of this method is illus-
trated in Figure 5a in which the PSV signal from a North At-
lantic Ocean core is compared with that from the sediments of
Lake Chewaucan. This particular interval of time immediately
follows the 190 Ky Pringle Falls/Jamaica/Biwa I geomagnetic
excursion (Herrero-Bervera et al., 1994; Negrini et al., 1994;
Henyey et al., 1995) discussed in the previous section. Note
that this example demonstrates the utility of the PSV method
for direct correlations between marine and terrestrial sedi-
ments. Clearly, there is an advantage to such an independent
correlation of marine and nonmarine climate records without
relying on either circular reasoning (i.e., the correlation of cli-
mate events to prove that the events occurred simultaneously)
or a comparison of radiocarbon dates from entirely different
geochemical environments. Another advantage of this method
is that, under ideal conditions, it allows for continuous age con-
trol throughout the sedimentary section.
Obviously, as with all relative age-dating methods, this tech-
nique is most valuable when used in conjunction with an inde-
pendent method, such as tephrochronology (Verosub, 1988;
Negrini, 1993; Liddicoat and Coe, 1997). For example, the ac-
curacy of the 14C chronology in the sediments of pluvial Lake
Russell is confirmed using such an approach. Figure 5b is a plot
of the inclination component of the PSV records from pluvial
Lakes Chewaucan and Russell. The Lake Chewaucan record is
from the B&B core taken from the middle of the Summer Lake
play a (Erbes, 1996). The Lake Russell record is after Lund et
al. (1988). The recently redated Wono tephra layer (Benson et
NUMBER 33 23
Inclination Inclination
40 80 40 80
170
180
Declination RelativeDeclination
Inclination
0 30 60 0 30 60 90
20T
projection of Wono
tephra age onto Lake
Russell inclination record
using Lake Chewaucan
inclination record
30
Normalized
Paleointensity
0 2
Normalized
Paleointensity
0 2
1 i i
Lake
Chewaucan Lake Russell
170
180
Lake
Chewaucan
Northwestern
Atlantic Ocean
FIGURE 5.?a, Correlation of paleomagnetic vector components recorded in the sediments of Lake Chewaucan
(Negrini et al., 1994) with those recorded in marine sediments from the northwestern Atlantic Ocean (S.P. Lund,
unpublished data). The bottom of the correlated intervals begins at the end of the -190 Ky old Pringle Falls/
Jamaica/Biwa I geomagnetic excursion. Inclination and declination are in degrees. Declination of the marine
sediments is relative rather than absolute because the core was not azimuthally oriented. The paleointensity data
in both records are normalized to the mean as per Tauxe (1993:340). The dashed correlation lines of all three
components were fixed by the two prominent inclination troughs shaded in black, b, Correlation of inclination
records from pluvial Lakes Russell (Lund et al., 1988) and Chewaucan (Negrini et al., 1993). The correlation
model is supported by the projection of the new age (27.3 + 0.3 14C Ky B.P.) for the Wono tephra layer (Benson
et al., 1997) onto the previously dated PSV record from Lake Russell.
al., 1997) is not found in the Lake Russell record but appears in
the Lake Chewaucan section not too far above the interval of
sediments that contains a record of the Mono Lake excursion
(Negrini, 1994). Using the PSV records, the new age of this
tephra layer (27.3 ?0.3 14C Ky B.P.) can be precisely projected
onto the Lake Russell record (Figure 5b), thereby confirming
the accuracy of this portion of the Lake Russell chronology de-
termined previously by radiocarbon dating of carbonates (e.g.,
Benson et al., 1990; Benson et al., unpublished).
PLUVIAL LAKE RECORDS
Late Pliocene to Middle Pleistocene (3 Ma-30 Kd)
LAKE LAHONTAN.?Our knowledge of the pluvial lacus-
trine history of the northwestern Great Basin is bimodal in na-
ture. On the one hand, no other pluvial lake in the Great Ba-
sin, with the exception of Lake Bonneville, has been studied
as extensively as Lake Lahontan (Figures 1, 6). On the other
hand, although significant advances have been made during
24 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
118?
__OREGON
CALIFORNIA[" NEVADA
J BLACK ROCKPLAYA
BOUNDARY OF
LAKE LAHONTAN
M
EXPLANATION
PRIMARY SILLS
SECONDARY SILLS
PYRAMID LAKE
HUMBOLDT-/'/[ )
PRESENT-DAY LAKES
OR PLAYAS
PRIMARY SILLS
Pronto
Chocolate
Adrian Valley
Darwin Pass
Mud Lake Slough
Astor Pass
Emerson Pass
MAJOR SUBBASINS
Smoke Creek/
Black Rock Desert
Carson Desert
Buena Vista
Walker Lake
Pyramid Lake
Winnemucca Dry Lake
Honey Lake
FIGURE 6.?Detailed map of Lake Lahontan pluvial system (from Benson et al., 1990, fig. 2).
the last 50 years regarding the prehistory of the other pluvial
lakes in the northwestern Great Basin, we still know little
about these other lakes relative to what is known about Lake
Lahontan.
The emphasis on Lake Lahontan is understandable. At high-
stand, this pluvial system covered more than 22,000 square ki-
lometers, approximately 50% more than all of the other pluvial
lakes in the northwestern Great Basin combined (Grayson,
1993). Its sheer size leads to the probability of finding key ex-
posures of critical geomorphic features and lake sediments.
Furthermore, the presence of large remnant lakes, particularly
Pyramid Lake and Walker Lake, provide for modern analogs
and, at times, for sites of continuously deposited sediments, en-
abling the potential retrieval of long cores containing nearly
continuous climate records.
Despite the complexity of the Lake Lahontan system (six
major streams feeding seven major subbasins that are separated
at different elevations by seven sills), the overall lake-level his-
tory has converged toward a coherent story for the latest Pleis-
tocene (see "Latest Pleistocene," below). The prehistory of this
pluvial system prior to 30 Ka is less well developed for the ob-
vious reason that most of the evidence either has been super-
posed by the sediments of the last highstand or has been eroded
or degraded through time. Nevertheless, great exposures of
NUMBER 33 25
sediments as old as late Pliocene in age can be found in can-
yons cut by the Humboldt, Walker, Carson, and Truckee Rivers
(Russell, 1895; Morrison and Davis, 1984a, 1984b; Morrison,
1991b). These exposures and their associated geomorphic fea-
tures (e.g., strandlines) are the foundation for the Quaternary
stratigraphy of the Lake Lahontan area, which was first devel-
oped by Russell (1895) and later refined by Morrison and oth-
ers (Morrison, 1964, 1991b; Morrison and Frye, 1965; Morri-
son and Davis, 1984a, 1984b). The latest stratigraphic
framework for this region consists of a series of alloforma-
tions, as shown in Figure 7. According to Morrison and Davis
(1984a: 255), an alloformation is defined as "a mappable strati-
form body of sedimentary rock that is defined and identified on
the basis of its bounding discontinuities." For example, lacus-
tral alloformations include all deposits of a single lacustrine ep-
isode, such as bars, deep-water sediments, deltas, tufas, etc. In-
terlacustral episodes could even include lake sediments in the
larger remnants of a great lake, e.g., Pyramid Lake (Morrison,
1991b). This usage follows the Stratigraphic Code of the North
American Commission on Stratigraphic Nomenclature (1983).
Restrictions on the use of such terminology, such as the need to
define alloformations as mappable units, limit the ability to di-
rectly correlate the contacts between allounits with climate-
change boundaries (e.g., see later discussion of Mount St.
Helens Cy tephra as a marker for the contact of the Wyemaha
and Sehoo Alloformations).
Given adequate spatial sampling of the Quaternary sediments
of the Lake Lahontan system, perhaps with the help of an exten-
sive coring program, it is this author's opinion that allounits
corresponding to each and every marine OIS cycle will be
found somewhere in the system. Only after such data is col-
lected will Quaternary geologists be able to make final deci-
sions on which groups of cycles best fit into collected allofor-
mations or, indeed, whether a separate alloformation should be
assigned to each cycle. With this end in mind, the present state
of the older Lahontan stratigraphy and its paleoclimatic impli-
cations are reviewed below.
The oldest sediments in the Lake Lahontan area are the
"deep-lake/deltaic sediments" (Figure 7) that lie below the in-
terlacustral Lovelock Alloformation in at least three locations:
(1) 700 m south of Rye Patch Dam on the west side of the
Humboldt River (Morrison and Davis, 1984a:256); (2) possi-
bly in the Truckee River valley, 6.5 km south of Nixon, Nevada
(Morrison, 1991b); and (3) at various locations in the badlands
near Weber Reservoir (Morrison and Davis, 1984a; Morrison,
1991b). The absolute age of these sediments is constrained to
be at least 800 Ky, based upon preliminary magnetostrati-
graphic results and on the identification of the ~1.0-1.1 My
old Glass Mountain G tephra layer in an upper unit of these
sediments near the Weber Reservoir (Morrison and Davis,
1984a; Morrison, 1991b, and references therein). A lower pre-
Lovelock lacustrine unit also appears near Weber Reservoir. Its
relative antiquity is inferred by the fact that it is extensively
faulted and folded and that these structures are truncated by an
TIME
years
x io3
10 -
20 - ?
1OO -
200 -
CO
flC
> 300
Ik
O
COI
400
o
2 500 ?
UJ
2
600 -
700 -
800 -
900 -
J
ALLO3TRATIQRAPHIC AND
PEDOSTRATIQRAPHIC UNITS
FALLON AF(?hallow-la
TOYEH QEOSOL _ _(aubaerlal)
. Upper member.^moderate lalceaL
Middle member (deep lake
Thinollte unit (shallow lake!
Lower member (deep lake!
CHURCHILL QEOSOL AND
WYEMAHA ALLOFORMATION
(subaerial/shallow-lake)
EETZA ALLOFORMATION
(sediments of three deep-
lake cycles)
PAIUTE ALLOFORMATION AND
COON QEOSOL COMPLEX
(alluvium and eolian sand
with many paleosols)
RYE PATCH ALLOFORMATION(deep lake sediments)
?
LOVELOCK ALLOFORMATION
AND PEDOCOMPLEX
alluvium with many paleosols)
DEEP-LAKE/DELTAIC
8EDIMENTS
JLAKE CYCLES(aolld)
ISOIL FORMINGIEPI800E8
I (pattern)
Ro
FIGURE 7.?Stratigraphic units of the Lake Lahontan pluvial system (from
Morrison, 1991b, fig. 6).
angular unconformity that separates it from the unit containing
the Glass Mountain G tephra layer.
Although these older lacustrine units are presumed to predate
the Lovelock Formation, this formation may itself be quite old
at some localities. For example, the ~0.9-1.0 My old Glass
Mountain D tephra has been found in the upper part of the
Lovelock Alloformation approximately 10 m below the Rye
Patch Alloformation at a locality in the Humboldt River can-
yon (Parker, 1993, and references therein). In fact, Parker
(1993) used magnetostratigraphy to show that the 25 m of sedi-
ments found below this tephra were at least ~2.0 My old. No
26 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
lacustrine sediments were observed in this section below the
Rye Patch Alloformation, prompting Parker (1993) to presume
that these 25 m of sediments are within the Lovelock Allofor-
mation. If one interprets the stratigraphy depicted in Figure 7 in
a strict sense, then one must conclude either that the pre-Love-
lock lacustrine sediments were deposited in an areally re-
stricted lake or that an unconformity exists in Parker's section
(i.e., the pre-Lovelock lacustral units in this part of the Hum-
boldt River canyon were eroded, and the lower part of Parker's
section belongs to an older interlacustral alloformation). Addi-
tional tephrochronological and magnetostratigraphical studies
in the Humboldt River canyon and Weber Reservoir badlands
are required to test the continuous nature of older alloforma-
tions in the Lahontan Basin.
The two allolacustral members of the Rye Patch Alloforma-
tion contain two tephra layers in the Humboldt River canyon,
namely the Rye Patch Dam bed and the Lava Creek B ash
(Morrison, 1991b; Davis, 1978). Morrison (1991b) quoted ages
of 630 Ky and 620 Ky for these two ashes, respectively. As dis-
cussed above, the Lava Creek B has since been redated at
665 ? 5 Ka by Izett et al. (1992), which puts the Rye Patch
Dam bed at a slightly older age of ~675 Ky. These tephra lay-
ers imply a correlation of this lacustral episode with OIS 16.
The Dibekulewe tephra of Davis (1978) is found above the
Lava Creek B tephra in the Humboldt River canyon (Morrison
and Davis, 1984a; Negrini, Verosub, et al., 1987). Morrison
(1991b) quoted an age of 610 Ky for the Dibekulewe tephra
and placed it at the top of the Rye Patch Alloformation. This
age estimate was probably derived from the work of Reick et
al. (1992), who found this tephra bed in close proximity to the
Lava Creek B ash in the Tulelake core. Although this was their
preferred age, these authors acknowledged that the age could
be much closer to that of the overlying Rockland ash (410 Ky)
because of the likelihood of disconformities in the interval be-
tween the Lava Creek B and Rockland ashes. A more recent
study on a long core from Owens Lake placed the age of the
Dibekulewe tephra at approximately 510 Ky based upon con-
stant-mass accumulation rates in this core (Sarna-Wojcicki et
al., 1997). This age is more in line with an earlier estimate by
Negrini, Verosub, et al. (1987:10,625) that was based upon the
relative position of these three tephra in a sedimentary section
near Blairsden, California. If the Dibekulewe is ~500 Ky old,
then either the Rye Patch Alloformation is younger or, alterna-
tively, the Dibekulewe ash is within the overlying Paiute Allo-
formation. Indeed, the Rockland ash lies within the Paiute Al-
loformation in the Humboldt River canyon (Parker, 1993);
thus, the upper age limit of the base of this alloformation is 410
Ky.
The Eetza Alloformation corresponds roughly to Russell's
(1895) lower lacustral clay. This alloformation contains at least
three major lacustral allomembers, which consist primarily of
silts and clays separated by thick lenses of alluvium or collu-
vium. All told there are possibly as many as six to eight lake
cycles recorded in exposures along the Truckee River between
Nixon and Wadsworth, Nevada (Morrison and Davis, 1984a;
Morrison, 1991b). The Wadsworth tephra of Davis (1978) is
found toward the top of this alloformation near Pyramid Lake.
The age of this tephra was estimated to be in the range of 155
-200 Ky (Morrison, 1991b; Reick et al., 1992, and references
therein). Because of this age and the observation that the Rock-
land tephra resides below the Eetza in the Paiute Alloformation
in the Humboldt River canyon, Morrison (1991b) correlated
the Eetza lacustral episodes with OIS stages 10, 8, and 6. This
correlation is supported by 230Th ages of 290 ? 90 Ky and 128
? 10 Ky for samples from near the bottom and top of the Eetza,
respectively (Kaufman and Broecker, 1965).
The Wyemaha Alloformation represents a long interlacustral
interval from ~130 Ka to ~35 Ka and is characterized by eolian
and alluvial coarse-grain deposits and well-developed paleo-
sols at higher elevations and by finer-grained sediments depos-
ited in shallow lakes in the depocenters of the major subbasins
(Morrison, 1991a). Morrison proposed that the boundary be-
tween the Wyemaha Alloformation and the overlying lacustral
Sehoo Alloformation should be put at the Mount St. Helens Cy
tephra, now thought to have formed 45-50 Ka (see above).
This is a reasonable suggestion as it would make the Wye-
maha/Sehoo transition relatively easy to map. This tephra layer
is thick, is widespread throughout much of the Lahontan Basin,
is distinctive both petrographically and geochemically, and is
commonly found at the base of lacustrine sequences associated
with the Sehoo Alloformation (Davis, 1978:45). There is a
trade off, however, between the mappability of this contact and
its correlation with global climate transitions. The Sehoo Allo-
formation is commonly considered to be roughly coeval with
OIS 2 and perhaps with the early Holocene (Davis, 1978;
Morrison, 1991b). If the Mount St. Helens Cy tephra layer is
indeed as old as it appears to be based upon recent chronomet-
ric studies (see above), then the base of the Sehoo Alloforma-
tion is pushed back to early OIS 3 time. Perhaps the 27.3 14C
Ky B.P. Wono tephra would be a better choice for the base of
the Sehoo Alloformation. It is as widespread as the Mount St.
Helens Cy tephra layer in the Lahontan Basin (Davis, 1978)
and has a more appropriate age. Furthermore, it was deposited
in a lake of moderate depth preceding the last significant reces-
sion of Lake Lahontan, which, in turn, preceded the prominent
transgressive phase correlated to the OIS 2 ice sheet advance
(Benson et al., 1995, 1997). The 23.2 14C Ky B.P. Trego Hot
Springs tephra layer is, in some sense, still a better choice as it
was probably deposited in the shallow-water lakes of the reces-
sion immediately prior to the OIS 2-related transgression. The
Trego Hot Springs tephra, however, is not as widespread as the
Wono bed (Davis, 1978), probably because it was deposited
under shallower lake conditions.
In addition to its limited application for exact correlations to
global or regional climate change, the Lake Lahontan stratigra-
phy places only very loose constraints on the size of this plu-
vial lake versus time. That is, one can determine whether or not
the lake existed during a given time interval but not the exact
NUMBER 33 27
elevation of the shoreline. To most effectively use the geologic
record to determine the effects of climate change on the lake,
one must determine the surface area of the lake versus time
(Smith and Street-Perrot, 1983; Benson and Thompson, 1987a;
Benson and Paillet, 1989). To do this, one must, of course, de-
termine the precise elevation of the lake surface. Although this
has been done exceedingly well for the time period after 30 Ka
(see next section), very little has been done to determine abso-
lute lake levels for earlier time periods. Morrison (1991b) has
done this in a general way by documenting the maximum ele-
vation occupied by lacustral alloformations. Reheis et al.
(2002) also have obtained estimates of the lake size during plu-
vial maxima in the early and middle Pleistocene by mapping
highly eroded but still discernable strandlines and near-shore
deposits around the Lahontan Basin and nearby basins. In do-
ing so, they have reconstructed an enormous lake correspond-
ing to OIS 16 that would have covered Reno and Carson City,
Nevada, and would have backed the Humboldt River at least to
the town of Battle Mountain, Nevada. Lao and Benson (1988,
and references therein) and Szabo et al. (1996) determined
lake-level minima by dating tufa deposits with the uranium-
series method. The results of the earlier study indicated that the
surface of Lake Lahontan in the Walker subbasin (the first to
fill up if it is fed by the Walker River) would have reached a
level more or less equal to its late Pleistocene maximum of
1330 m at 290 ?90 Ka, -200 Ka, 128 ? 10 Ka, and 49 ?5 Ka.
According to the model of Benson and Thompson (1987b),
Pyramid Lake, Winnemucca Dry Lake, and Smoke Creek/
Black Rock Desert subbasins would have coalesced and risen
to their common elevation of 1207 m at Emerson Pass. It is
conceivable that all of these subbasins could have joined the
Walker subbasin at its 1330 m elevation if they received suffi-
cient spillover from the Walker subbasin through the Carson
River subbasin. Szabo et al. (1996) combined dates from sev-
eral studies in the Pyramid Lake subbasin near the elevation of
the 1207 spillpoint into the Smoke Creek/Black Rock Desert
subbasin. They found that a lake filling at least the local subba-
sins would have been at or above the Emerson Pass sill for
much of the middle Pleistocene, especially in the time interval
from 250 Ky to 170 Ky. Also, these subbasins would have
filled at least twice during the time period from 65 Ka to 30
Ka.
LAKE MODOC AND ENVIRONS.?The longest, most continu-
ous pluvial lake record in the northwestern Great Basin comes
from the sediments of pluvial Lake Modoc (Figure 1) and
nearby Buck Lake and Round Lake in south-central Oregon
(Klamath County) and northernmost California (Modoc and
Siskiyou Counties), respectively. Upper and Lower Klamath
Lakes and Tule Lake are the remnants of pluvial Lake Modoc,
which once filled the subbasins that these lakes occupy. At plu-
vial maximum, the surface of Lake Modoc stood at an eleva-
tion of 1292 m, approximately 63 m above the surface of mod-
ern Tule Lake (Dicken, 1980). The Lake Modoc system is not
considered to be part of the Great Basin hydrologic province as
defined by Grayson (1993) because most of it presently drains
into the Pacific Ocean via surface flow through the Klamath
River and via groundwater flow through permeable volcanic
rocks into the Klamath River and Pit River systems. Neverthe-
less, it is often included as a source of Great Basin paleocli-
mate data because it lies within the botanical Great Basin and
its size has responded dramatically to climate changes in the
past regardless of its connection to the Pacific Ocean (Dicken,
1980; Bradbury, 1991; Grayson, 1993). Also, as with most plu-
vial systems in the hydrographic Great Basin, Lake Modoc has
been a site of lacustrine sedimentation for millions of years as a
consequence of its presence within the Basin and Range struc-
tural province (Dicken, 1980).
Sediment cores were taken from seven sites in this system by
the USGS (Adam et al., 1989; Adam, 1993; Klamath Core
Study Team, 1995). At two of the sites (Tulelake, 334 m; and
Butte Valley, 102 m) coring was completed into sediments
older than 3 My. At three of the other sites (Wocus Marsh, 52
m; Buck Lake, 41m; and Round Lake, 50 m) the oldest cored
sediments were from 1.4 to ~1.0 Ma. The Grass Lake core (29
m) penetrated sediments no older than ~400 Ky, and the Cale-
donia Marsh cores (15.3 m) bottomed out in sediments
younger than 100 Ky. The ages of all of the cores were con-
strained using tephrochronology (Adam et al., 1989; Reick et
al., 1992). Several well-known tephra were identified, includ-
ing the 1.94 My old Bear Gulch ash, the 1.45 My old Rio Dell
ash, the 665 Ky old Lava Creek B ash, the 410 Ky old Rock-
land ash (Sarna-Wojcicki et al., 1985; Meyer et al., 1991), the
50-55 Ky old Olema ash (Sarna-Wojcicki et al., 1988), a pre-
cursor ash (the Tsoyawata bed) to the Mount Mazama eruption
(Davis, 1978; Bacon, 1983), and several ashes that correlate
with the collection of tephra found nearby at Summer Lake,
Oregon, such as one of the Pumice Castle tephra set (Bacon,
1983; Davis, 1985; Reick et al., 1992). The ages of the Tule-
lake, Butte Valley, and Buck Lake cores were also constrained
by magnetostratigraphy (Reick et al., 1992; Roberts et al.,
1996; Rosenbaum et al., 1996). Radiocarbon dating was con-
ducted on samples from the top of the Grass Lake, Caledonia
Marsh, and Wocus Marsh cores with limited success because
of inverted ages, low concentrations of carbon, and a prepon-
derance of anomalously young ages (Bradbury, 1991; Klamath
Core Study Team, 1995). Despite similar difficulties encoun-
tered with the Tulelake core, Bradbury (1991, fig. 6) (Figure 8)
was able to assemble a chronology for the top of this core by
using selected radiocarbon dates and tephra.
In addition to age control, several other studies have begun
on the Klamath Basin cores that have generated data from
which interpretations of pluvial lake conditions can be made.
These include lithologic description, palynology, diatom and
ostracode analyses, environmental magnetism, and geochem-
istry. An initial study on the Tulelake core found that two ma-
jor shifts in the character of pollen type versus depth curves
occurred at -2.5 Ma and -1.65 Ma (Adam et al., 1989). The
2.5 My shift is best expressed by a slight overall increase in
28 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
a
a.o
0 75 0
percent
75 0 500 300 750 500
diatoms/mm
FIGURE 8.?a, Diatom biostratigra-
phy for the entire Tulelake core from
Bradbury (1992, fig. 6). For detailed
descriptions of all genera, groups,
and species the reader is referred to
Bradbury (1991, 1992). b, Fragi-
laria and Aulacoseira diatom pro-
files for the last 180 Ky, after Brad-
bury (1997, fig. 9). Note positions of
key tephra layers.
Fragilaria Aulacoseira
50
age (Ky)
100
150
Tephra T2023
percent
Artemisia count and by the transition toward higher-frequency
variations in both Artemisia and TCT (Taxodiaceae +
Cupressaceae + Taxaceae, which includes Western Juniper).
This older shift was correlated with the onset of continental
glaciers in North America and the beginning of the Praetiglian
cold period in Europe. Some researchers believe that the Plio/
Pleistocene boundary should be moved back to this point in
time, specifically to the Gauss/Matuyama polarity chronozone
boundary, rather than the currently accepted time, the top of
the Olduvai normal-polarity subchronozone (see Morrison,
1991a, and references therein). The latter candidate for the
Plio/Pleistocene boundary is also expressed in the Tulelake
core. After 1.65 Ma there are common interruptions in pollen
preservation, a drop in TCT amplitude, increases in the peak
percentages of Artemisia, Cheno-ams, and Poaceae (grasses),
and high-amplitude oscillations in total algae, all of which
were interpreted to indicate a pattern of alternating lacustrine
and nonlacustrine phases (Adam et al., 1989). Two diagnostic
species of diatoms, Anomoeoneis costata (Kutzing) Husted
(saline water) and Cocconeis placentula Ehrenberg (shallow
water), were most abundant between 90 and 50 m of depth
(800 to <400 Ka), indicating an overall drying of the lake and
a hydrologically closed lake that was rich in solutes. These ob-
servations are consistent with well-preserved ostracodes and
with high values for "factor 3," an association of geochemical
entities (e.g., (carbonate), Mg, and Sr) indicative of dry lake
conditions (Adam et al., 1989).
Bradbury (1991, 1992, 1997) later used a more detailed and
comprehensive diatom stratigraphy to reconstruct lake condi-
tions in the Tule Lake subbasin throughout the entire three mil-
lion years represented by that core. In these reports he made
several intriguing interpretations that bear on global climate-
change models. First, he noted an ~40 Ky periodicity in total
diatom concentration in the two densely sampled time intervals
(3.0-2.5 Ma and 2.05-1.85 Ma) prior to 1.0 My. This observa-
tion is consistent with the dominance of the obliquity period in
the late Pliocene and early Pleistocene observed in marine oxy-
gen isotope records (Ruddiman et al., 1986; Ruddiman and
Raymo, 1988). Second, based upon the gradual disappearance
of Aulacoseira solida (Eulenstein) Krammer and the subse-
NUMBER 33 29
quent takeover by a Fragilaria-Stephanodiscus?Cyclotella as-
semblage, he noted a transition from "more equable and prob-
ably warmer Pliocene climates to more variable, more
seasonal, drier, and probably colder Pleistocene climates"
(Bradbury, 1992:284). Third, three broad peaks in Aulacoseira
solida alternating with broad peaks in Fragilaria species oscil-
late with a period approximately equal to the 413 Ky eccentric-
ity-driven insolation cycle (Ruddiman and Raymo, 1988).
Fourth, from 900 to ~300 Ka, the diatom stratigraphy in the
Tulelake core is characterized by several peaks in Fragilaria
sp. concentration that are more or less in tune with glacial
stages recorded by marine oxygen isotope curves (Bradbury,
1992, fig. 7). Furthermore, these peaks are terminated by in-
creases in diatoms that prefer more saline or alkaline waters
(e.g., Anomoeoneis costata, Cocconeis placentula). Bradbury
interpreted these data to imply that, during glacial periods,
Tule Lake was probably an open, freshwater marsh formed
during cold and somewhat dry climates. During interglacials,
effective moisture decreased but winter precipitation was prob-
ably significant as the presence of Stephanodiscus and Aula-
coseira species in these intervals suggests at least seasonally
high lake levels and dilute waters. Fifth, Bradbury (1992)
noted that the more or less uniform response to glacial cycles
during the 900-300 Ka time period breaks down in the latest
Pleistocene. The interval from 300 to ~180 Ka is represented
by Anomoeoneis costata, a diatom that prefers "exceptionally
shallow and alkaline conditions." Thus, the low effective mois-
ture interval that might correspond to OIS 7 appears to begin
80 Ky too soon in the Tulelake core. The interval from ~ 170 to
~130 Ka contains abundant Fragilaria, which would be ex-
pected for a glacial interval like OIS 6 based upon its occur-
rence patterns in this core from 900 to 300 Ka; however, as
Fragilaria abundance declines, the diatom concentration be-
comes characterized by deeper-water diatoms (e.g., Aula-
coseira ambigua (Grunow) Simonson, Cyclotella meneghini-
ana Kutzing, Cyclotella bodanica Grunow) until A. ambigua
decreases at ~50 Ka. Thus, if the 900-300 Ka diatom assem-
blage is a good analog for the latest Pleistocene, then it appears
as if the entire interval from 170 to 50 Ka is characterized by a
glacial interval; that is, the OIS 5 interglacial is not represented
in the Tulelake core. Bradbury (1991:228) noted that diatoms
such as Fragilaria, which characterized pluvial intervals in the
middle Pleistocene, might also do well during a warm-dry in-
terpluvial period like OIS 5; however, the relatively low abun-
dances of Fragilaria during previous interglacials is left open
for question.
Two cores from the Klamath Basin have proven to be fertile
ground for the nascent field of environmental magnetism. Rob-
erts et al. (1996) conducted sediment magnetic measurements
on bulk samples and reflected light microscopy and x-ray de-
fraction (XRD) analysis on magnetic separates from the Butte
Valley core. They found that most of the magnetism in the 3.0
My of core was carried by titanomagnetite, which was most
likely derived from basalt detritus in the catchment area.
Greigite (Fe3S4), an authigenic magnetic mineral, was found in
several horizons in the age interval from 3.0 to ~1.0 Ma. Be-
cause the formation of greigite depends upon the availability
of dissolved sulfate, Roberts et al. (1996) tentatively con-
cluded that greater residence time of lake water may have
caused the slightly anoxic conditions under which the avail-
ability of dissolved sulfate would have been enhanced. Thus,
lake conditions were significantly different, at least in inter-
vals, for the periods before and after ~1.0 Ma. Notably, a nine-
fold increase in sedimentation rate also occurred at ~ 1.0 Ma,
supporting a major change in the environment of deposition at
this time.
Climate change during the middle Pleistocene (~400 Ka to
-300-200 Ka) also was interpreted with the help of an environ-
mental magnetism record from Buck Lake, 15 km west of Kla-
math Basin (Rosenbaum et al., 1996). Variations in magnetic
properties were found to correlate strongly with changes in cli-
mate, as indicated by pollen type. In particular, zones of high
magnetic susceptibility, hence, high concentrations of magne-
tite, coincide with cold and dry climate zones, and zones char-
acterized by low susceptibilities were identified with warm cli-
mate zones. The magnetic variations preceded the changes in
pollen, an observation that prompted Rosenbaum et al. (1996)
to propose that Buck Lake magnetic properties responded more
quickly to climate change than did vegetation. High hematite
concentrations, inferred from two ratios (HIRM and the S pa-
rameter) calculated from measurements of isothermal remnant
magnetization (Thompson and Oldfield, 1986), were found to
be in phase with high magnetite concentrations. The overall
content of heavy minerals, as indicated by the concentrations of
Ti and Zr (immobile elements mainly held in heavy minerals),
was also found to be in phase with increases in magnetic-min-
eral content; however, because the magnetic-mineral concen-
trations varied by two orders of magnitude between cold and
warm climates and, in contrast, the heavy-mineral concentra-
tion varied only by a factor of three, Rosenbaum et al. (1996)
concluded that the low amount of magnetic minerals during
warm climates relative to other heavy minerals was caused by
sourcing from soil and highly weathered rock in the catchment.
In contrast, during pluvials, greater runoff would have brought
in a greater concentration of relatively fresh detritus derived
from basaltic bedrock.
LAKE CHEWAUCAN.?Four subbasins (Summer Lake, Upper
and Lower Chewaucan Marshes, and Lake Abert) combined
during pluvial maxima to become pluvial Lake Chewaucan in
south-central Oregon (Figure 1). At its maximum, the lake was
113 m deep, and its surface stood at an elevation of 1328 m.
The climate prehistory of this pluvial lake (Allison, 1945) was
made well known when Antevs (1948) used it as a constraint
for his Neothermal climate model. Although this prehistory
proved to be wrong in light of the application of more modern
chronological techniques (see earlier discussion under "Other
Models for Holocene Climate Change"), pluvial Lake Chewau-
can has continued to be the focus of climate-change research
30 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
CO
o
0
50
Mill
lower
?
i i i ii i
higheri
i 111 n i
500
, i ij
more
saline
0
J
fresher
JLi
20
I 1 1
? coarser
f~ i *
0
finer
?
\i i i i i i i i
^
200
outcrop magneticsusceptibility
(SI units)
250
lower higher
WL core magnetic
susceptibility
(SI units)
i i i i i II i i II
1
outcrop
salinity index
(dimensionless)
saZ
WLcore
salinity index
(dimensionless)
WL core silt/clay ratio
(dimensionless)
-
y,
marine 0-18
(norm
III
alized)
III
interpluvial
(Stage 7)
interpluvial
(Stage 3)
pluvial
(Stage 4)
interpluvial
(Stage 5)
pluvial
(Stage 6)
0.001 0.001 500 1.0 -1.0
FIGURE 9.?Lake-size proxies for Summer Lake subbasin of pluvial Lake Chewaucan compared with a marine
oxygen isotope proxy for ice-sheet extent (Martinson et al., 1987, fig. 18). Magnetic susceptibility has been
shown to correlate with other lake-level proxies by Erbes (1996). The salinity index is a weighted ratio of salino-
phile ostracode faunal counts over counts of ostracodes that prefer fresh water (Palacios-Fest et al., 1993). The
outcrop samples are from the Ana River canyon near the northern margin of the subbasin. The WL core samples
are from a locality closer to the depocenter, near the northern edge of the modern playa lake.
owing to the combination of excellent exposures of pluvial lake
sediments, abundant tephra layers, climate-controlled geomor-
phology, and well-preserved records of fossil biota and magne-
tism (e.g., Allison, 1982; Grayson, 1993). Furthermore, in the
last 50 years many studies on archeology, climate modeling,
geochemistry, geomorphology, fossil magnetism, hydrology,
mineralogy, paleontology, sedimentology, structural geology,
and tephrochronology have set the stage for continuing paleo-
climate and related research in this basin (Conrad, 1953;
Brown, 1957; Phillips and Van Denburgh, 1971; Van Den-
burgh, 1975; Allison, 1982; Negrini et al., 1984; Davis, 1985;
Negrini et al., 1988; Simpson, 1990; Banfield et al., 1991a,
1991b; Berger and Davis, 1992; Gobalet and Negrini, 1992;
Negrini and Davis, 1992; Freidel, 1993; Palacios-Fest et al.,
1993; Pezzopanne, 1993; Negrini et al., 1994; Oetting, 1994;
Roberts et al., 1994; Wigand et al., 1995; Erbes, 1996; Wigand
and Rhode, 2002).
The lake-level history for the Chewaucan system during the
last 200-300 Ky has been estimated from several proxy indica-
tors derived from the analysis of bottom sediments exposed in
outcrops and sampled from cores (Figure 9). The correlation of
cores with outcrops and the age control of the resulting com-
posite section (Figure 3) are based upon a variety of methods,
including lithostratigraphy, tephrochronology, sediment mag-
netism, paleomagnetic secular variation, radiometric methods,
and thermoluminescence (Davis, 1985; Berger and Davis,
1992; Negrini et al. 1994; Erbes, 1996). Notably, this chronol-
ogy was constructed completely independent from a priori as-
sumptions regarding correlations to climate-change events.
Variations in lithology type, sediment grain size, and the con-
centration of magnetite grains, as indicated by magnetic sus-
ceptibility, all indicate that the environment of deposition
changed systematically through time more or less in tune with
the marine OIS signal (Figure 9). Complementary variations
NUMBER 33 31
through time in the paleosalinity of the lake, as estimated by
the faunal distribution and geochemistry of fossil ostracodes,
and in the shallowness of the lake, as estimated by the noisi-
ness of the trace-element concentration of the same ostracodes,
strongly suggest that the changes in the depositional environ-
ment were caused by increases in lake size as the ice sheets ad-
vanced and vice-versa (Figure 9; Palacios-Fest et al., 1993; Er-
bes, 1996). Thus, both the nature of the variations of climate
proxies in the Chewaucan sediment and their timing indicate
that, to a first order, the lake-level history of pluvial Lake
Chewaucan supports Antevs' jet stream migration model.
Marine OIS 4 is well resolved in the Summer Lake record
(Figure 9). In fact, the freshest water, hence the deepest lake,
appears to have occurred in OIS 4 rather than in OIS 6 or 2.
There are two explanations for this observation. First, because
the signatures and durations of OIS 4 are either subdued, short-
ened, or nonexistent in lake-level proxies from southern Great
Basin lakes (e.g., Bradbury, 1997; Lowenstein, 2002), it may
be that the southernmost position of the jet stream occupied a
more northerly latitude (i.e., over Summer Lake) than it did
during the OIS 2 glacial maximum (Lowenstein, pers. comm.,
1997). More specifically, if the jet stream remained for a while
at its southernmost point over southern Oregon during OIS 4,
this lesser glacial stage would have had a very profound effect
on Lake Chewaucan. In contrast, if the jet stream moved
quickly through Oregon on its way south and back during the
OIS 2 event, then the pluvial episode seen at Summer Lake
during OIS 2 may have been relatively minor. A second and
equally plausible explanation is the following. Lake Chewau-
can could have been deeper and fresher during OIS 6 and 2
than it was during OIS 4, but the records of the freshest lakes
were erased by deflation during the subsequent interpluvial
maxima. The latter explanation is supported by the gap in sedi-
mentation shown in Figure 9 from -160 to 115 Ka and also by
the fact that the sediments at the surface of the modern Summer
Lake playa are no younger than 23 14C Ky B.P. (Negrini et al.,
1993). Note that without good independent age control one
would never realize that cold/wet and hot/dry extrema corre-
sponding to both sides of glacial terminations had been erased
by deflation. Thus, the Summer Lake example serves as a
warning to those drawing conclusions regarding climate ex-
trema from poorly dated paleoclimate records associated with
deflation-prone desert basins.
In addition to climate variations, drainage diversions also
could have affected the surface level of Summer Lake. In fact,
the Chewaucan River, the main stream entering the Chewaucan
system, is currently diverted from the Summer Lake subbasin
by its own fan-delta complex (Allison, 1982; Davis, 1985). It is
unlikely, however, that such diversions would have resulted in
a lake-level history that matches the OIS signal so well (Figure
9). Thus, climate variation is probably the major influence on
lake level in this subbasin. A similar study of the Lake Abert
subbasin could provide a test for this hypothesis.
In an early study of the paleomagnetism at this locality,
Milankovitch-like periodicities were reported in spectra of the
magnetic vector time series (Negrini and Verosub, 1987; Ne-
grini, Darke, et al., 1987; Negrini et al., 1988). This was cer-
tainly a provocative result but has since proven to be incorrect
by improvements in the chronology (e.g., Negrini et al., 1994).
Even though the PSV time series are no longer considered to be
related to climate change, they still have potential as tools for
high-resolution correlation of other climate records with the
Chewaucan record (see above).
Comparison of Lake Chewaucan and Lake Modoc Re-
cords: The Lake Modoc climate record, as described in a pre-
ceding section, contrasts greatly with the Chewaucan record.
Specifically, OIS 4 is a major feature in the Chewaucan record
but is barely noticed in the Tulelake record. Furthermore, the
OIS 5 interglacial episode at Lake Modoc is characterized by
continuous deposition in an environment similar to that of gla-
cial episodes earlier in the Pleistocene. Most of the discrepancies
between these records can be explained by differing interpreta-
tion of chronologies rather than by radically different climates in
regions that are in close proximity. Figure 10 presents a compar-
ison of the two records after the Lake Modoc chronology has
been reinterpreted by aligning the records using four tephra lay-
ers common to both records. The age scale assumes that the
chronology for the Lake Chewaucan record is accurate, but be-
cause the comparison is based upon the relative position of the
four tephra layers, the comparison would still hold up if the
Chewaucan chronology changes with improved age control.
With the reinterpreted chronology, the two records now have
much in common with each other and with the marine OIS sig-
nal (Figure 10). In particular, the interval of time from ~190 to
-160 Ka is characterized in both records by more pluvial con-
ditions, and a strong OIS 4 signal is now observed in both
records. Note also that the reinterpretation forces a major OIS 5
unconformity on the Lake Modoc record. This is not entirely
unreasonable considering that such an unconformity occurs in
the Chewaucan record.
PLUVIAL LAKE MALHEUR.?A large region (2400 km2) of
very low relief (< 10 m) in eastern Oregon presently contains
Malheur, Mud, Harney, and Silver Lakes (Figure 1). Presently,
all except Malheur are ephemeral lakes. In 1986, after a series
of abnormally wet years, Malheur overflowed through a sill
called The Narrows into Mud Lake and then into Harney Lake
(Hostetler and Bartlein, 1990). At its peak, this combined lake
reached a depth of only 4.4 m, and its surface stood at an eleva-
tion of 1250.5 m. Had the effective moisture level increased so
that the lake rose another one meter or so, then the combined
lake also would have overflowed into the Silver Lake basin. If
the lake deepened yet another three meters to a surface eleva-
tion of-1254 m, it would have overflowed into the Columbia
River system.
One of the tephra layers from the -70 Ky old Pumice Castle set
was identified in lacustrine sediments at an elevation of 1247
m near the base of two trenches associated with archeological
32 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
50
100
15
To
E
T5
HI
150
200 -
Lake Chewaucan (Ana River Outcrop)
magnetic susceptibility
0.001 jsl units) 0.01
Lake Modoc (Tulelake Core) Marine 0-18
0 40 0 40 ? 0
Olema Tephra
Tulelake T2382
Pumice Castle Set
OIS 4 expressed in Tule Lake
core as high percentages of
Fragilaria or Aulacoseira
gap in record corresponding to
unconformity observed in
Chewaucan sediments
gap in record corresponding to
unconformity predicted for Tule
Lake sediments between Pumice
Castle and "V" tephra layers
Tulelake T1193
??????????????????????????I
Summer Lake Tephra V
Tulelake T2023
Summer Lake Tephra
OIS 6 expressed as high
percentage of Fragilaria
50
100
150
200
500 0 0 40 0 40
ostracode salinity index % Fragilaria % Aulacoseira
FIGURE 10.?Comparison of Lake Chewaucan and Lake Modoc/Tulelake core (from Figure 8) paleolake envi-
ronments. Pluvial conditions increase to the right in all diagrams. Relative age-control is anchored by four com-
mon ash layers, as per Reick et al. (1992). The absolute chronology of the Lake Modoc record is reinterpreted in
terms of preferred ages for the tephra layers, as discussed in the text. The positions of the these tephra layers in
the original Tulelake chronology are shown in Figure 86. The new chronology for the Lake Modoc record
resolves the anomalous occurrence of Fragilaria sp. in the OIS 5 interpluvial, which was noted by Bradbury
(1991, 1992), by placing this occurrence in OIS 6 and OIS 4. This chronology also indicates a major gap in the
Tulelake record between the Pumice Castle tephra and the Summer Lake V tephra layers.
investigations near the southern margins of modern Malheur
Lake (Elston, 1992; Botkin and Carembelas, 1992a, 1992b). In
one of the trenches, the ash is a continuous layer found near the
base of at least one meter of stratified lacustrine sandy clays.
Thus, for approximately 70 Ky, during the middle of OIS 4,
pluvial Lake Malheur had enough water in it to deposit more
than one meter of sediments on a lake bottom at 1248 m, which
is two meters above the present-day lake surface. If the OIS 4
lake had a depth of at least 2 m, then pluvial Lake Malheur
would have been at least as large as it was in 1986.
Pluvial Lake Malheur also appears to have included all of the
four major subbasins at -32 14C Ky B.P., as evidenced by a ra-
diocarbon date on mollusk shells collected on the "lakeside
face of a gravel beach ridge" (Gehr, 1980:167). The elevation
of this ridge is 1251 m.
Latest Pleistocene (30 to -10 Ka)
LAKE LAHONTAN.?The late-Pleistocene pluvial history of
Lake Lahontan has been the subject of several reviews published
NUMBER 33 33
in the last 50 years (e.g., Morrison, 1965; Mifflin and Wheat,
1979; Davis, 1982; Smith and Street-Perrot, 1983; Benson and
Thompson, 1987a, 1987b; Dansie et al., 1988; Benson et al.,
1990; Morrison, 1991b; Grayson, 1993). This current review
draws heavily on those earlier reviews and adds supplemental
material from subsequent research.
Extensive radiocarbon dating associated with tufas, organic
materials from packrat middens and other terrestrial animals,
tephra, wood, shells, bulk sediments, fishing cordage, etc.,
have allowed a detailed and coherent picture to emerge for the
past 30,000 years (especially the last 25,000 years) of lake-
level prehistory for the Lahontan Basin. The pioneering radio-
metric studies in the Lake Lahontan area were done on carbon-
ates by Broecker and Orr (1958) and Broecker and Kaufman
(1965). Many of these dates have been contested by Benson
(1978), particularly those indicating an unexpectedly young
age (11.5-9.5 14C Ky B.P.) for the last 1330 m Lahontan high-
stand, because they were found to be inconsistent with the
stratigraphic framework as reviewed in Morrison and Frye
(1965). Since the work of Broecker's group, more than 100 ra-
diocarbon dates have been done on Lahontan Basin materials,
most of them by L. Benson and colleagues (e.g., Benson, 1978,
1994; Benson et al., 1995). This group has demonstrated that
dense, lithoid tufas and aragonitic gastropods yield consistent
dates, whereas massive forms of tufa (e.g., dendritic tufa) yield
significantly younger ages. Through careful petrographic anal-
ysis, Benson (1978) demonstrated that the latter type of tufa
contains abundant evidence for the introduction of secondary
carbon (e.g., shell fragments in pores and precipitated acicular
carbonate) well after the initial formation of the tufa. Thus,
massive tufa is usually not an appropriate material for radiocar-
bon dating. After careful selection of dateable materials from
his petrographic analysis and from previous radiocarbon analy-
ses on deltaic wood from Born (1972), Benson (1978) pub-
lished a new prehistory of the Lake Lahontan surface level
through time and, on its basis, pushed the highstand date back
to 13.6-11.1 14CKyB.P.
The next major advance came with the addition to the data-
base of several radiocarbon dates on organic material from pri-
marily packrat middens, which in contrast to carbonates depos-
ited under water, constrained the maximum depth of the water.
When combined with the previous analyses and the addition of
new tufa dates, these complementary data led to an even clearer
prehistory (Thompson et al., 1986; Benson and Thompson,
1987b; Benson et al., 1990). An exhaustive review on the for-
mation history of the several types of tufas in the Lake Lahon-
tan area, additional tufa dates, and new dates and lake-level
context for two critical chronostratigraphic horizons, namely
the Wono and Trego Hot Springs tephra beds, refined this pre-
history even further (Benson, 1994; Benson et al., 1995; Ben-
son etal., 1997).
The current state of the late Pleistocene Lake Lahontan pre-
history is represented by the surface level of Lake Lahontan in
the Pyramid subbasin, as shown in Figure 11. Although the
lake level of the Pyramid Lake subbasin is not a precise proxy
for the surface area (the parameter most directly affected by
climate) of Lake Lahontan as a whole, modeling by Benson
and Paillet (1989, fig. 8) showed that the overall shapes of both
quantities over time are quite similar.
According to radiocarbon dates on tufas, tufa-coated wood,
and organic remains of a camel (see below), low to moderate
lake levels were present in this subbasin from ~35 to 23-22 14C
Ky B.P., although Lake Lahontan was high enough (1172 m in
elevation) for Pyramid Lake to spill over into Winnemuca Lake
for a few thousand years -30 14C Ky B.P. (Figure 11). The lake
was low enough for a camel to become stuck in the mud near
the modern (post-irrigation) shoreline during the "Wizard's
Beach Recession" (Dansie et al., 1988), which lasted perhaps
as long as 4 Ky from -26 to -22 14C Ky B.P. Further support
for this lowstand is provided by coarse-grained alluvial depos-
its underlying the Trego Hot Springs tephra in the Black Rock
Desert (Ewing, 1996; F. Nials, pers. comm., 1996) and by high
values of 818O and 813C measured in tufa of that approximate
age (Benson et al., 1996).
After the Wizard's Beach recession, abundant data summar-
rized in Benson et al. (1990) indicate that the lake rose very
quickly at -23-22 14C Ky B.P. to a level high enough to con-
nect all of the major subbasins except for Walker Lake (Figure
11). The lake remained stable near this elevation until -16.0
14C Ky B.P. when it underwent a series of high-amplitude oscil-
lations ending with the terminal recession to below the Emer-
son Pass sill (1207 m) at -9.5 14C Ky B.P. The high-amplitude
oscillations included a spike-like, major transgression to the
maximum highstand centered somewhere between 14.0 and
13.5 14C Ky B.P. and a return to levels high enough for Pyramid
Lake to spill into the Smoke Creek and Black Rock Desert ba-
sins at just before 10 14C Ky B.P. This last transgression is sup-
ported by dates on shells, tufa, and marl from the Carson
Desert subbasin at similar elevations (Elston et al., 1988; Cur-
rey, 1990; Benson et al., 1992).
The overall changes in the size of Lake Lahontan, as repre-
sented in Figure 11, also are recorded in the stable isotope sig-
natures of tufas from the Pyramid Lake subbasin (Benson et al.,
1996). Benson et al. (1996) found that when Pyramid Lake is a
closed lake the values of 818O are dependent upon the balance
between input through river discharge, which decreases 818O,
and evaporation, which increases 818O. When Pyramid Lake is
open 818O is dependent mostly upon the fraction of water input
lost to spillover into the next basin; 813C was found to depend
upon lake volume indirectly through its affect on the photosyn-
thesis-respiration balance. Regarding the lake-level prehistory
shown in Figure 11, the major lake-level minima at 26, 15.5,
and 12 14C Ky B.P. were found to be associated with relatively
heavy S18O and 813C values, and the opposite was found to be
true for the major maxima at 14 and 10.5 14C Ky B.P. (Benson
et al., 1996). Buoyed by these results, Benson et al. (unpub-
lished) derived an even higher-resolution record of change in
the level of Pyramid Lake with 818O measurements of the total
34 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
1350
1300
1250
Lake Surface
Elevation (m)
1200
1150
1100
1050
Younger Dryas-
aged transgression
Mud Lake
Slough Sill
o
middle Holocene
regression?
(see Figure 12a)
i i i r i i i i
*Jet stream returns to
the Lahontan
Basin on its way
north. The already
large lake in-
creases size to
maximum.
Darwin Pass Sill
Jet stream
south of
Lahontan
Basin.
Reduced
evaporation
from colder
temperatures
maintains lake
size.
Emerson Pass Sill
Jet stream enters
the Lahontan
Basin on its way
p south. Increased
precipitation causes
the lake to increase
size ~ to level of
Darwin Pass sill.
t i i i I i i i t I i i i i I i i ( i I i i il i I i i i i I i i i i
0 10 20
14CAge(KyB.P.)
30
FIGURE 11.?High-resolution lake-level prehistory of the Pyramid Lake subbasin for the latest Pleistocene and
Holocene after Benson et al. (1995, fig. 11). The various symbols denote the type of subaqueous materials
dated. See Benson et al. (1995) for detailed explanation. Abundant, unplotted terrestrial dates, mostly from
packrat middens, constrain the maximum lake level from 13 I4C Ky B.P. to present. Interpretation of data in
terms of supporting Antevs' jet stream model is after Benson and Thompson (1987a, 1987b) and Benson et al.
(1995).
inorganic carbon (TIC) fraction from a sediment core sampled
at about 200-year intervals. These data, in conjunction with
measurements of total organic carbon (TOC), were found to
vary through time in high-frequency oscillations with square-
wave morphologies that can be correlated in a general way to
D-0 cycles and Heinrich events (Benson et al., unpublished).
Timing of the Last Lake Lahontan Highstand: The only
major point of contention during the last 20 years with the
22-9.5 14C Ky B.P. lake-level record for Lake Lahontan has
been the timing of the maximum highstand (Bradbury et al.,
1989; Benson, 1991; Grayson, 1993:96). Bradbury et al. (1989)
argued that a Lake Lahontan that was high enough to incorpo-
rate the Walker Lake subbasin could not have occurred after 23
14C Ky B.P. because a well-dated sequence of sediments from
Walker Lake shows this lake to have been a saline lake or a
playa for the entire interval from 23 to 5 14C Ky B.P. Further-
NUMBER 33 35
more, evidence for a deep lake during this highstand interval is
missing from the lake-bottom sediments at a coring location in
the Black Rock Desert (Ewing, 1996), although this latter re-
cord does indicate a spike-like interval of freshwater conditions
at the slightly older age of 14.5 I4C Ky B.P.
An attempt to resolve this controversy with a core from the
Carson Sink proved to be inconclusive because of inconsistent
radiocarbon dates (Thompson et al., 1990). Benson (1991), in
an article that focused on replying to Bradbury et al. (1989),
made a strong case for deflation as a means to remove the one
meter or less thickness of lacustrine sediment he estimated to
have been deposited in the Walker Lake subbasin during the
approximately 200-year duration of the Lake Lahontan high-
stand in this location. Although the Black Rock Desert would
have had highstand water in it for a few hundred more years,
deflation also could have removed the record of the highstand
from the sediments studied by Ewing (1996).
The lack of evidence from lake-bottom sediments from only
two locations can be explained by deflation or even, perhaps,
by erosion from underwater currents or slumping; however, if a
highstand lake did indeed exist at this time, then there must be
a lake-bottom sediment record preserved somewhere in the
Lahontan subbasins. Recent studies in the Carson Desert sub-
basin by graduate students from the University of Nevada,
Reno (UNR), appear to have found exactly this type of evi-
dence. One of these students (Cupp, 1998) found evidence for a
surface elevation of the lake at -1280 m and dated it at 12.8 14C
Ky B.P. The dated materials were gastropods sampled from
well-bedded lacustrine sands interpreted to be a near-shore de-
posit. When the Carson Desert subbasin is filled to 1280 m the
effective wetness of the entire Lahontan system is approxi-
mately 80 percent of maximum (Benson and Paillet, 1989). Al-
though this is not high enough for the lake to reach the Adrian
Pass spill point (1302 m) into Walker Lake, the lake only
reached this elevation during the highstand interval. The work
of K. Adams, the other UNR student, provided even more di-
rect evidence for the highstand during the time indicated by
Benson and others. Working in the Jessup embayment at the
north margin of the Carson Desert subbasin, Adams (1998) and
Adams and Wesnousky (1998) dated three shorelines, two with
gastropod shells and the third with a camel bone preserved in
these lakeshore deposits. The older dates, 13.28 ? 0.11 14C Ky
B.P. and 13.11 ? 0.11 14C Ky B.P., were on the shells found in
shoreline deposits at elevations of 1327 m and 1331 m, respec-
tively. The youngest date, 13.07 ? 0.06 14C Ky B.P., was on the
camel bone found in shoreline deposits at an elevation of
1338.5 m. These dates and elevations provide strong, indepen-
dent support for the 14-12.8 14C Ky B.P. highstand of Benson
et al. (1995). Incidentally, further evidence for this highstand is
provided by the low values of 518O found by Benson et al.
(1996; unpublished) in tufas and lake-bottom sediments with
ages of~1414CKy B.P.
Testing Paleoclimate Records with the Late Pleistocene Lake
Lahontan Record: The high-resolution Lahontan lake-level
record from the Pyramid Lake subbasin has been used to test
models of late Pleistocene climate change. Benson et al. (1995:
25-26) demonstrated that, in response to the deflection of the
jet stream during the last glacial advance, "Lake Lahontan
should have experienced two lake-level maxima corresponding
to the pre- and post-maximum glacial passages of the jet
stream core over the Lahontan basin." The timing of these
lake-level maxima are indicated by the dashed vertical lines in
Figure 11. The corresponding glacial maximum was at ~20?18
14C Ky B.P., at which time the jet stream would have been de-
flected to its southernmost position. In this scenario, the lake
would have maintained a high level while the jet stream was
south of the Lahontan area because of reduced evaporation as-
sociated with lower temperatures and the probable occurrence
of ice cover during the winter. This is consistent with the mod-
eling results of Benson (1986), Hostetler and Benson (1990),
and Hostetler (1991). The larger pluvial maximum occurred as
the jet stream returned north after the glacial maximum be-
cause the initial size of the lake was larger than it was when the
jet stream passed through the Lahontan area on its way south
(Benson et al., 1995). Thus, to a first order, the size of Lake
Lahontan through time supports Antevs' jet stream migration
model.
The Lake Lahontan record is sufficiently detailed that it
also has been used to test higher-resolution models of climate
change. For example, Benson et al. (1992) used the abrupt
transgression at -10 14C Ky B.P. as evidence in support of the
Younger Dryas climatic interval as a global climate event.
Furthermore, Benson et al. (unpublished) used the higher-res-
olution 618O record from the Pyramid Lake core to show that
relatively dry periods coincided with the HI, H2, and possi-
bly H3 Heinrich events. This observation supports the de-
scriptive model that Heinrich events represent global climate
change.
LAKE MODOC.?The latest Pleistocene age control in the
Tulelake core is provided by the 23.2 ? 0.3 14C Ky B.P. Trego
Hot Springs tephra (Bradbury, 1991; Reick et al., 1992; Ben-
son et al., 1997) and a complementary date of 25.06 ? 0.32 14C
Ky B.P. from slightly lower in the section. Above the Trego
Hot Springs tephra is a several meter gap and then sediments
containing a series of three precise but inverted radiocarbon
dates in the age range from 14.9 to 9.4 14C Ky B.P. (Bradbury,
1991).
From -40 Ka to a few meters above the Trego Hot Springs
bed, the Tulelake core is characterized by diatoms that are in-
dicative of "somewhat shallower and more alkaline conditions"
(Bradbury, 1991:231). Because the Trego Hot Springs bed is
exposed in littoral sediments in an outcrop 15m higher than the
position of that tephra in the Tulelake core, the lake must have
been -15 m deep at that time, 5 m deeper than the historical
maximum (Bradbury, 1991). After that time, presumably dur-
ing a time interval corresponding to OIS 2, the climate in the
Lake Modoc region was dry overall but was punctuated with
rapid and large fluctuations in moisture.
36 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
1400
o
I 1350
<D
<DO
03
Fort Rock
3
CO
highest shoreline
basin bottom
0
O
a>I
Co
1300
1250
1400
1350 -
1300 -
1250
B
-?-?-(1
Freidel(1993)>
Various (summarized
in Freidel, 1994)
10 15 20
Chewaucan
25 30
" highest shoreline
J. Licciardi
(unpublished)
X"^ im
? X?
? basin bottom
.... i .... i ..
Allison (1982)
" \ '
t& Negrini and
Davis (1992)
/ ;
?
? . 1 ? . . . 1 . . . . 1 . . . . "
10 15 20 25 30
Q.
<DTJ
E
E"s
E
o
c
ctio
LL
1
0.8
0.6
0.4
0.2
0
Composite
10
14
15 20
CAge(KyB.R)
25 30
FIGURE 12.?Lake-level prehistories of pluvial Lakes Fort
Rock and Chewaucan for the latest Pleistocene (after
Negrini and Davis, 1992; Freidel, 1993, 1994). Solid circles
denote ages from archeological sites and from other terres-
trial materials formed above lake level. Open squares repre-
sent materials formed below lake level. In the Chewaucan
record from -30 to 18 14C Ky B.P. at an elevation of 1276 m,
the bold line alternating between black and grey indicates
lake levels above and below, respectively, the Ana River
outcrop sampled by Negrini and Davis (1992). The ages of
these transitions have been changed slightly to conform to
the new ages for the Wono and Trego Hot Springs tephra
layers (Benson et al., 1997). Unpublished radiocarbon shell
dates from near-shore sediments were provided by J. Liccia-
rdi. In the composite record, the question mark (?) at ~24
I4C Ky B.P. indicates a lack of data constraining the maxi-
mum lake depth. A question mark (?) also is placed at ~21
14C Ky B.P. because the gap in the bottom sediment record at
this time interval is represented by a slight angular uncon-
formity. This could be caused by either a lowering of the
lake or a post-depositional deformation.
PLUVIAL LAKES CHEWAUCAN, FORT ROCK, AND ALKALI.?
The latest Pleistocene/early Holocene lake-level prehistory of
pluvial Lake Chewaucan and nearby pluvial Lake Fort Rock
and Lake Alkali (Figure 1) have been summarized by Allison
(1966b, 1979, 1982) and Freidel (1993, 1994). Lake levels
through time are constrained in these basins by the age of bot-
tom sediments (Allison, 1982; Davis, 1985; Negrini and Davis,
1992), fossils from ancient shorelines (Allison, 1982; Freidel,
1993), and dates associated with archeological sites (Ferguson
and Libby, 1962; Bedwell, 1973; Allison, 1979; McDowell and
Benjamin, 1991; Freidel, 1994). The individual data for the
Lake Fort Rock and Lake Chewaucan basins and a composite
lake-level history are presented in Figure 12.
The oldest radiocarbon date shown is from gastropod shells
collected in beach deposits exposed in Fandango Canyon at the
southern margin of Lake Fort Rock (Freidel, 1993). X-ray dif-
fraction analysis demonstrated that the shells were made of
100% aragonite, i.e., no secondary calcite was detected. More
NUMBER 33 37
details of the predating treatment of the shells were given in
Freidel (1993). The shells were collected from two exposures
of beach deposits. The oldest date of-34 14C Ky B.P. corre-
sponds to samples collected at an elevation of 1361 m; the
youngest date of ~30?0.6 14C Ky B.P. was from 1353 m. These
dates and elevations imply that Lake Fort Rock was at -65%
maximum depth twice between -34 and 30 14C Ky B.P.
Paleomagnetic secular variation dating of Lake Chewaucan
bottom sediments showed that the lake had water at least to the
level of the Ana River section (-13% of maximum) for the
time interval containing the Mono Lake geomagnetic excur-
sion. This excursion is dated -28 14C Ky B.P. at Mono Lake,
California (e.g., Lund et al., 1988). Approximately 4 Ky of de-
posits are missing above the excursion level, probably because
the lake level was below the level of the Ana River canyon out-
crop (Figure 12). The lake dropped below the level of the Ana
River section sometime during this time interval, as evidenced
by the incorporation of the 27.3 ?0.3 14C Ky B.P. Wono tephra
layer (Negrini and Davis, 1992; Benson et al., 1997) in a
beach-lag deposit. The Summer Lake subbasin of Lake
Chewaucan held at least some water during this entire time in-
terval because the PSV record from a basin-bottom core is
complete for several thousands of years after the Mono Lake
excursion (Figure 5 b). Lake Chewaucan rose to approximately
50% of its maximum level at -23-22 ?-0.7 14C Ky B.P. This
age is based upon two pieces of evidence. First, Allison (1982)
collected and dated gastropod shells (22.1 ?0.7 14C Ky B.P.) in
a baymouth bar deposit at an elevation of 1312 m. This date
and elevation is supported by the presence of the Trego Hot
Springs tephra (23.2 ?0.3 14C Ky B.P.) in the lacustrine sedi-
ments of the Ana River canyon (Negrini and Davis, 1992; Ben-
son et al., 1997). A 2 Ky hiatus in the PSV record from the Ana
River section and an associated cross-bedded sand deposit are
consistent with the lake again dropping below 13% of its maxi-
mum depth sometime after the deposition of the Trego Hot
Springs tephra and before the deposition of the 20.5 ?-0.5 14C
Ky B.P. Mount St. Helens Mp tephra bed (Negrini and Davis,
1992). After that, the lacustrine sediments at Ana River were
deposited more or less continually until at least 16.7 ?-0.5 14C
Ky B.P. This latter date marks the top of the lake sediments at
an unconformity over which lie eolian sands containing the
Mount Mazama tephra (Davis, 1985). The extent of deflation
or other post-depositional erosion of overlying lacustrine sedi-
ments is unknown. Another shell date of 17.5 ?0.3 14C Ky B.P.
from a basalt pebble conglomerate collected at 1349 m by Alli-
son (1982) indicates that the long-lived lake implied by the in-
terpretations of Negrini and Davis (1992) rose to at least 75%
of its maximum elevation during this time. This long-lived lake
probably represents the pluvial maximum at Lake Chewaucan.
Lake levels between the highstand and its terminal Pleistocene
recession (Figure 12) are constrained by unpublished radiocar-
bon dates on shells from near-shore sediments (J. Licciardi,
pers. comm., 1995).
Evidence for the timing of the termination of the pluvial
maximum in these Oregon lakes also is available from the Lake
Fort Rock basin (Bedwell, 1973; Freidel, 1993, 1994). Sub-
aqueous dates on gastropod and ostracode shells (Bedwell,
1973; McDowell and Benjamin, 1991) constrain the lake level
to no lower than 1335 m and 1312 m at 12.9 ?0.23 14C Ky B.P.
and 11.5 ?0.09 14C Ky B.P., respectively. Thus, pluvial Lake
Fort Rock probably dropped through elevations corresponding
to 36% and ~2%-3% of its maximum elevation in this -1.5 Ky
interval. This timing is well supported by several dates on ar-
cheological materials that must have been above lake level, in-
cluding charcoal, camel bone, and a sandal (Figure 12; Bed-
well, 1973; Freidel, 1993; McDowell and Dugas, 1993).
Further support is lent by similar evidence for extremely shal-
low lakes in the nearby Alkali Lake basin by -11 14C Ky B.P.
(Freidel, 1993, and references therein; McDowell and Dugas,
1993).
LATE PLEISTOCENE PALEOCLIMATE MODELS AND RECORDS
FROM OREGON AND CALIFORNIA LAKES.?Without better age
control it is difficult to use the Tulelake record as a test for late
Pleistocene climate models. One can, however, compare the
overall character of the latest Pleistocene climate data from the
Tulelake core with the well-dated record from Lake Lahontan
during the same time interval. For example, the hydroclimate
data implied by the Tulelake core for the time of and before the
Trego Hot Springs tephra layer (relatively shallow and alka-
line) are consistent with the low to moderate levels of Lake
Lahontan during late OIS 3 times (Bradbury, 1991). Also, the
rapid and large fluctuations in Tulelake-area hydroclimate after
the Trego Hot Springs tephra and before the onset of the Ho-
locene (Bradbury, 1991) are in character with many of the rapid
fluctuations seen at Lake Lahontan.
Lack of data is the main problem with the latest Pleistocene
record from the Oregon lakes. Again, the record compares with
the coeval Lake Lahontan record in a general sense. The pre-
dominance of shallow lakes from -27 to 23 14C Ky B.P. follow-
ing a moderately deep lake is quite similar to Lake Lahontan
prehistory in that same time interval (Figures 11, 12). After
that, both systems were characterized by long time intervals
with very large lakes until the Holocene, perhaps interrupted
only by short-term recessions. The paucity of data in the
records for Oregon lakes, especially in the important time inter-
val between 17 and 12 14C Ky B.P., makes a more detailed
comparison futile at this time.
Holocene
Several previous reviews focussed heavily on Holocene pa-
leoenvironments, principally because of their importance to
the archeological sciences (Baumhoff and Heitzer, 1965; Meh-
ringer, 1977, 1986; Davis, 1982; Aikens, 1983; Dansie et al.,
1988; Grayson, 1993; Hattori and Tuohy, 1993; McDowell
and Dugas, 1993; Aikens and Jenkins, 1994; Madsen, 2002).
38 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
2 calKyB.P. 6
j I 10
1350
1300
c
1 1250
QjI
2
1200
1150
1100
6 14C Ky B.P. 8 10
level of Pleistocene highstand
Pyramid Lake subbasin
KDD
O pack rat middens O tufa
D deltaic wood # human burial sites
KM
0 O
K>H?-O-
archeological index of
prevalent lake level
relative to sill
weighted sum of dates
on archeological sites
in Fort Rock basin
dates on high shorelines
small
very small
standing water in
northern Alkali basin
'C Ky B.P.
FIGURE 13.?a, Holocene lake-level prehistory for the Pyramid Lake subbasin
of the Lahontan pluvial system. Published data are from Born (1972), Benson
and Thompson (1987b), Benson et al. (1990), Hattori and Tuohy (1993), Ben-
son et al. (1995), and Tuohy and Dansie (1997). Evidence for the middle
Holocene 1153 m lowstand is provided by a recent radiocarbon date (5.91 ?
0.125 14C Ky B.P.) on a nearly complete human skeleton with a partial cranium
coated in tufa (A. Dansie and D. Tuohy, in litt, 1998). The archeological index
corresponds to a qualitative estimate of lake level by Hattori and Tuohy (1993),
which was based upon the frequency of occupation in the Winnemucca Caves.
These caves would have been abandoned except when lake levels were above
the mud lake sill and, thus, Winnemuca Lake held water. A = lake levels above
the sill, B = below, and V = variable, b, Holocene lake-level prehistory for Ore-
gon pluvial lakes. C = the date on a shoreline corresponding to a relatively high
early Holocene Lake Chewaucan (S. Robinson, pers. comm. to Gehr, 1980);
M = dates of a relatively high Lake Malheur (Gehr, 1980). Relevant data for
early Holocene standing water in the northern Alkali basin is from Willig
(1988) and Pinson (MS.). An index of relative availability of water on the sur-
face of Lake Fort Rock basin is indicated by the weighted sum of radiocarbon
dates published by Mehringer and Cannon (1994). Diamond Pond interpreta-
tion is from Wigand (1987).
NUMBER 33 39
Selected data for Holocene lake levels in the northwestern
Great Basin are summarized in Figure 13.
LAKE LAHONTAN.?Because Holocene lake-level changes
were too small in magnitude to reach most of the sills between
major subbasins, the case for each basin is presented sepa-
rately. Holocene lake-level minima in the Pyramid Lake subba-
sin are constrained by radiocarbon dates on deltaic wood and
tufas; maxima are constrained by organic materials from pack-
rat middens and archeological sites (Figure 13a). The lake-
level curve shown in Figure 1 la is considered to be smoothed
because it reflects general trends in lake levels as indicated by,
for example, the occupation of archeological sites. The reader
should bear in mind that there are several excursions within
these general trends to higher or lower lakes that probably rep-
resent higher-frequency lake-level changes. For a related dis-
cussion, see Hattori and Tuohy (1993) and Mehringer and Can-
non (1994).
The Holocene record in the Pyramid Lake subbasin begins
with the recession that terminated the relative highstand associ-
ated with the Younger Dryas stade (Benson et al., 1992; Ben-
son et al., 1995). The lowest lake level during this recession
(-1153 m) is indicated by two radiocarbon dates. One of these
(9.66 ? 0.13 14C Ky B.P.) was on a piece of fishing line made
from sagebrush bark found at an archeological site on the
northwestern shoreline of Pyramid Lake (Tuohy, 1988). The
other (9.5-9.1 14C Ky B.P.) was on a partial human skeleton
found at approximately the same elevation at the north end of
Pyramid Lake (Tuohy and Dansie, 1997). Thereafter, the lake
probably rose to a level above 1168 m, as indicated by an age
of 8.80 ? 0.09 14C Ky B.P. on a piece of wood from sediments
interpreted to have been deposited on a delta slope (Born,
1972). The occupation of Winnemuca Caves from -8.4 to 7.8
14C Ky B.P. indicates that water levels high enough for Pyramid
Lake to spill into Winnemuca Lake were probably sustained, at
least intermittently, until ~8 14C Ky B.P. (Hattori, 1982; Hattori
and Tuohy, 1993).
Except for a few radiocarbon dates from the Winnemuca
Caves (Hattori and Tuohy, 1993), no evidence has been found
for a long-lived lake above historical levels for the time inter-
val between -8 and 5 14C Ky B.P. Davis (1982) and Benson
and Thompson (1987b) argued that such lack of radiocarbon
evidence does not indicate a desiccation of Pyramid Lake.
They noted that elevations of the lake far below the historical
(before modern agricultural water diversion) lowstand level of
-1177 m would have killed two species of endemic fish, the
cui-ui (Chasmistes cujus Cope) and the Lahontan cutthroat
trout (Oncorhynchus clarki henshawi Richarson). This argu-
ment is supported by the failure of these fish to survive during
the present century without help from hatcheries after the lake
was lowered artificially to below 1160 m by irrigation-related
diversion of the Truckee River. The lowest possible level for
the survival of these fish is, of course, constrained by the 1153
m regression at 9.66 14C Ky B.P., as discussed in the previous
paragraph. The lake reached this low level at least one other
occasion, in the middle Holocene. Evidence for this 1153 m
lowstand is provided by a recent radiocarbon date (5.91 ?
0.125 14C Ky B.P.) on a nearly complete human skeleton with a
partial cranium coated in tufa found at Wizard's Beach at Pyra-
mid Lake (A. Dansie and D. Tuohy, in litt., 1998). A rise to
levels as high as 1185 m during the late Holocene is indicated
by abundant dates for inorganic carbonates and wood depos-
ited near lake-surface level in a fluvio-deltaic environment
(Born, 1972; Benson et al., 1995).
The only direct information on Holocene lake levels with
quantitative age control in the Carson Sink area is provided by
the tephrochronology of Davis (1978). A Mazama-age lake
level was inferred by the identification of the Mazama tephra
(6.845 ? 0.05 14C Ky B.P.) and its precursor, the Tsoyawata bed
(7.015 ? 0.045 14C Ky B.P.), at two localities in this subbasin.
The two tephra were found in the alluvium in the West Hum-
boldt Range at an elevation of 1204 m. They were also identi-
fied in the shallow lake or marsh sediments described by
Morrison as Section 22a (Morrison, 1964:128) east of Fallon,
Nevada, near the Stillwater Reservoir at an elevation of ?1193
m, which is only one or two meters above the modern low
point in the entire Carson Sink subbasin (Davis, 1978; Elston et
al., 1988). Thus, in the few hundred years of time spanned by
the volcanic eruptions associated with these two tephra, the
lake in the Carson Desert subbasin was less than 10 m deep
(Davis, 1982). In contrast, sometime after the eruption that cre-
ated the Tarupah Flat bed (1500-600 years old), the lake had
risen to a level of 1207 m, as evidenced by the wave-cut plat-
forms crosscutting this ash layer at two locations near Salt
Wells, Nevada (Davis, 1978).
Walker Lake was interpreted to have been an ephemeral
playa lake during the early Holocene (Bradbury et al., 1989;
Benson et al., 1991). This interpretation was based upon the as-
semblage of microfossils (diatoms, ostracodes, aquatic pollen),
lithology, and an abrupt change in sedimentation rates. After
-5.0-4.5 14C Ky B.P. the lake quickly filled to a depth as great
as 90 m. This rapid transgression has been attributed to the di-
version of the Walker River into the Walker Lake subbasin
from its previous destination, the Carson Desert subbasin
(King, 1978, 1993; Davis, 1982; Morrison and Davis, 1984a,
1984b; Bradbury et al., 1989; Benson et al., 1991). Based upon
similar evidence to that mentioned above and also on the stable
isotope analyses of Benson et al. (1991), a coherent lake-level
prehistory of this subbasin has been determined for the latest
Holocene. The coherence of this prehistory with that of nearby
Mono Lake (Stine, 1990) gives confidence that late Holocene
lake levels in the Walker Lake subbasin accurately reflect cli-
mate change (Bradbury et al., 1989; Benson et al., 1991).
There is some controversy as to whether an extreme low-
stand of the lake from -2.5 to 2.0 14C Ky B.P. was caused by
climate change or by another diversion of the Walker River.
Bradbury et al. (1989) contended that climate change is the
more likely cause because the very existence of even a shallow
lake would have required some recharge from the Walker River
40 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
into the basin, where evaporation greatly exceeded precipita-
tion, and because climate records from nearby localities in the
Great Basin also indicate drier climates during this time inter-
val. They further noted that Walker Lake filled more slowly af-
ter 2.0 14C Ky B.P., as evidenced by step-like increases in key
ostracode and diatom taxa, than it did after the previous desic-
cation. Such a slow transgression is more consistent with cli-
mate change than with the sudden diversion of a major river,
like the Walker River, back into this subbasin.
Benson et al. (1991), on the other hand, preferred river diver-
sion over climate change as a cause for this near-desiccation
event because (1) the usually coherent discharge of the Truckee
and Walker Rivers would not have allowed flow in the Walker
River to almost cease while the Truckee was flowing enough to
keep Pyramid Lake above the level required for the spawning
of cui-ui and emerald trout, (2) the Walker River must have
continued to flow somewhere to have allowed the survival of
the cutthroat trout while the lake-dwelling cui-ui population
was wiped out as Walker Lake nearly dried out, and (3) cur-
rently, discharge from local streams is sufficient to maintain
Walker Lake at very shallow levels (< 1 m) even in the absence
of the Walker River.
LAKE MODOC.?Despite the poor age control in the Ho-
locene part of the Tulelake core, Bradbury (1991:233) asserted
convincingly that "the strong fluctuations between planktonic
and benthic diatoms reflect real lake level changes that are im-
portant for resolving paleoclimatic and archaeological ques-
tions." Clearly, this record will become a valuable one with
improved age control. Even as it stands, it provides important
information surrounding the Tsoyawata tephra (7.015 ? 0.045
14C Ky B.P.), which was identified in the Tulelake core (Reick
et al., 1992). Two complementary 14C dates were obtained
within ~1 m below this tephra in the correct stratigraphic or-
der, thereby confirming the ash correlation. The dominance of
Aulacoseira ambigua in the vicinity of the Tsoyawata tephra
bed suggests the presence of a relatively deep lake at this time,
perhaps the 20 m lake suggested by an archeological study at
Nightfire Island (Sampson, 1985; Bradbury, 1991). Above the
Tsoyawata tephra, A. ambigua is replaced by Fragilaria sp.
This indicates a lower lake in the middle Holocene. The return
of A. ambigua at the top of the core marks a deepening lake at
a poorly constrained time, perhaps by ~3 Ka (Bradbury,
1991).
OTHER OREGON LAKES.?Most of the evidence for Ho-
locene lake-level changes in Oregon lakes has been reviewed
by Freidel (1993, 1994), McDowell (1992), Grayson (1993),
McDowell and Dugas (1993), and Aikens and Jenkins (1994).
A summary of these reviews with a minor amount of supple-
mental material follows.
The early Holocene appears to have been characterized by
pervasive shallow lakes throughout the Oregon portion of the
Great Basin (Figure 136). Four very shallow lakes were re-
ported to have come and gone during the early Holocene
within the confines of pluvial Lake Alkali (Figure 1). The first
of these lakes would have been ~0.5 m deep in the Dietz sub-
basin, located in the northern part of the Alkali Lake basin.
The shore of this lake, named Lake Koko, was occupied by a
Clovis-age culture. This conclusion was based upon fluted-
point assemblages that were found distributed along a linear
trend, which was interpreted to be a shoreline (Fagan, 1988;
Willig, 1988; Willig and Aikens, 1988). Another assemblage
of artifacts, which was characterized by large, stemmed, shoul-
dered, and lanceolate points, was found distributed around a
slightly higher shoreline and has been attributed (by Willig,
1988) to the Western Stemmed Point culture. Willig (1988) in-
terpreted this association as a later culture that occupied the
shores of the younger, slightly deeper (~2 m) Sand Ridge
Lake. Thus, she concluded that Lake Koko existed sometime
during the time interval attributed to the Clovis-age culture
(11.5-11.0 14C Ky B.P.) and that Sand Ridge Lake existed dur-
ing the time interval attributed to Western Stemmed Point tra-
dition (10.8-7.0 14C Ky B.P.). Although the overlap in age of
these two cultures may be larger than implied herein (e.g.,
Beck and Jones, 1997:196-197), the observation of Lake
Koko shorelines covered by Sand Ridge Lake sediments
strongly supports an older age for Lake Koko. Charcoal frag-
ments associated with Sand Ridge Lake were radiocarbon-
dated (Willig, 1988) at 9.61 ? 0.1 14C Ky B.P., providing fur-
ther support for the occupation of Sand Ridge Lake by a West-
ern Stemmed Point culture. Willig (1988) postulated the exist-
ence of two younger lakes, an intermediate-sized Lake
Delainie and a deeper (~4 m) Big Cut Lake, based upon the su-
perposition of shorelines and lake sediments. The ages of these
two younger lakes are not constrained.
The Dietz site lake-level history has since been reinterpreted
by Pinson (MS.). In this later study, the stratigraphy was inter-
preted to consist of deep-lake sediments covered by an eolian
deposit, which in turn was covered by a clay-silt layer that was
interpreted to have been deposited in a marsh, based upon its
internal distribution of gastropods and organic material. Over
this last unit is another eolian deposit from which Willig's char-
coal date was obtained. A second charcoal date of 9.43 ?0.08
14C Ky B.P. from this horizon confirmed Willig's date (Pinson
(MS.)). This eolian unit is also the one from which most of the
buried, early Holocene archeological materials were derived.
Notably, the top of this unit appears to have a nascent soil hori-
zon directly over which lies the Mazama pumice. Pinson (MS.)
argued convincingly that this stratigraphy can be interpreted as
the drying up of the terminal, deep Pleistocene lake at -10 14C
Ky B.P. and that thereafter only a shallow marsh existed. This
marsh, in turn, dried completely by ~9 14C Ky B.P., and there
has been no standing water in the Dietz subbasin since that
time.
Fort Rock Basin has been the location of several important
late Pleistocene and early Holocene archeological sites (Bed-
well, 1973; Grayson, 1993; Aikens and Jenkins, 1994; Beck and
NUMBER 33 41
Jones, 1997), some of which reveal a great deal about paleo-
lake levels (e.g., Figure 136). Early interpretations, based upon
evidence from the Connley Caves excavation, pointed to con-
tinuous human occupation for the time from -11 to 3 14C Ky
B.P., with the notable exception of the interval from 7.2 to 4.4
14C Ky B.P. (Bedwell, 1973). The apparent disappearance of
human occupation was initially interpreted, at least in part, to
have been the result of a catastrophic impact on the population
caused by the eruption of Mount Mazama (Bedwell, 1973);
however, it was later attributed to the shrinkage of large,
nearby bodies of water, as evidenced by the near disappearance
of waterfowl remains in the archeological middens of these
caves (Grayson, 1979, 1993).
Since the work at Connley Caves, Aikens and Jenkins
(1994) conducted a new, basin-wide project. They concluded
that when more sites are considered the hypothesis for the mid-
dle Holocene abandonment of the Fort Rock Basin begins to
fall apart. For example, Mehringer and Cannon's (1994) inves-
tigation of "Kelly's Site" in the Fort Rock Basin indicated that
people not only occupied the basin on occasion during that
time interval, but they were eating waterfowl. Analyses of ra-
diocarbon dates from several additional studies done as part of
this basin-wide project, however, revealed shorter-term inter-
vals of relatively lower population density at the bottom of Fort
Rock Basin that were presumably related to shorter intervals of
lower lake levels and dried-up marshes (Figure 136). These pe-
riods are roughly 7-5 I4C Ky B.P. (Jenkins, 1994; Mehringer
and Cannon, 1994) and 3-2 14C Ky B.P. (Jenkins, 1994). Al-
though much more work has to be done to precisely define
these dry/wet intervals, it is interesting to note that the onset of
wetter conditions at ~5 14C Ky B.P. implied by the above chro-
nology coincides almost exactly with the transition from
ephemeral marshes to open water at Diamond Pond/Malheur
Maar in eastern Oregon south of pluvial Lake Malheur (Wi-
gand, 1987).
In addition to dispelling the notion of a 3000-year long mid-
dle Holocene abandonment in the Fort Rock Basin (Bedwell,
1973), new work by P. McDowell and J. Benjamin suggests
that the Neopluvial (post-altithermal) lake postulated by Alli-
son (1979) also did not exist in the Fort Rock basin. Their con-
clusions were based upon observations of "extensive dunes
with elevations low enough and dates early enough that they
would have been destroyed or badly eroded had they been in-
undated by a basin-wide Neopluvial lake" (Greenspan, 1994:
488; P. McDowell, pers. comm. with Greenspan, 1994).
The subbasin remnants of Lake Chewaucan also appear to
have been significantly higher during the early Holocene than
during modern times. The evidence for this includes a personal
communication cited in Gehr (1980) from S. Robinson of the
USGS in which Robinson reported a radiocarbon age of 9.39 ?
0.045 14C Ky B.P. on a tufa-coated beach gravel found 20-22 m
above the modern level of Lake Abert. Summer Lake rose in
the late Holocene, as evidenced by a shoreline 12 m above the
bottom of the Summer Lake basin. This shoreline cuts across
sand dunes consisting of pumice from the Mount Mazama
eruption (Allison, 1982). There is no age control constraining
the minimum age of this feature. There is also evidence that
Lake Abert rose again, probably in the late Holocene. This evi-
dence consists of a higher shoreline (18 m above the modern
lake) that truncates presumably young colluvium (Allison,
1982). This feature has not been dated.
Pluvial Lake Malheur rose to levels high enough to join all of
the subbasins into one lake and perhaps to overflow into the
Columbia River system at least three times during the early
Holocene (Figure 136). Two of these episodes are dated at
9.62 + 0.05 14C Ky B.P. and 8.68 ? 0.55 14C Ky B.P. (Gehr,
1980). No record for the surface levels of this lake exists for
the remainder of the Holocene.
TESTING HOLOCENE PALEOCLIMATE MODELS WITH NORTH-
WESTERN GREAT BASIN PLUVIAL RECORDS.?The balance of
evidence from Figure 13 indicates that pluvial lake levels were
slightly higher in the northwestern Great Basin during the early
and late Holocene than they were in the middle Holocene. This
set of observations, at first glance, appears to support Antevs'
Neothermal climate model (Figure 2a), although the observa-
tions call for a slightly older Altithermal age than was antici-
pated by Antevs (Grayson, 1993). As more evidence has accu-
mulated, however, it has become increasingly apparent that
exceptions to the rule are abundant and that data must be spa-
tially and temporally averaged to conform to the Neothermal
model (Grayson, 1993; Mehringer and Cannon, 1994). In par-
ticular, the Holocene appears to be characterized by short-pe-
riod fluctuations in climate of significant magnitude superim-
posed on the long-term averages (e.g., Benson et al., 1991;
Mehringer and Cannon, 1994). Such variability is consistent
with the high-frequency variations in solar luminosity inferred
by cosmogenic isotope concentrations (Figure 2); however,
there is no obvious correlation between the observed climate
change presented in Figure 13 and those predicted by the lumi-
nosity proxies shown in Figure Id. In addition, there is no evi-
dence in northwestern Great Basin pluvial records, as of yet,
for the early Holocene "cooling event" at 8.4-8.0 cal Ky B.P.
(Figure 2c).
The lack of correlation between northwestern Great Basin
pluvial records and the predictions of Holocene climate models
probably results from deficiencies in the collective data set and
the lack of specific climate changes inferred from the models.
The sampling of Holocene paleoclimate data from individual
pluvial lake basins in the northwestern Great Basin is usually
too sparse for the detection of high-frequency climate change.
Furthermore, little modeling has been done to predict the ef-
fects of low-amplitude changes in Holocene climate on pluvial
lake levels in this region. For example, one recent study on the
geochemistry of Great Barrier Reef corals (Gagan et al., 1998)
has demonstrated that tropical oceans in the Pacific were ~1? C
warmer -5.35 cal Ky B.P. than they are today and that they
42 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
were enriched in 18O, implying increased levels of evapotrans-
piration. This intriguing result is consistent with intensified El
Nino events during the middle Holocene, which in turn could
explain the apparent increased variability in lake levels during
this time (e.g., Grayson, 1993; Mehringer and Cannon, 1994).
Before such a hypothesis can be anything more than specula-
tive, atmospheric circulation models with the ability to predict
the resulting low-amplitude changes in lake levels need to be
constructed, and extremely high-resolution pluvial lake records
need to be available to test the applicability of this idea to the
northwestern Great Basin.
Summary and Future Trends
SUMMARY OF MAJOR RESULTS
Since the time of Antevs, many contributions have been
made toward the understanding of climate change from studies
of pluvial lakes in the northwestern Great Basin. The following
list is a sample of these contributions.
1. Both of the global climate events proposed for the Plio/
Pleistocene boundary at ~2.5 Ma and -1.65 Ma have been ob-
served in several paleoclimate indicators from the Tulelake
core (Adam et al., 1989).
2. Variations in the diatom stratigraphy of the Tulelake core,
which have been interpreted to represent cycles between gla-
cial and interglacial climates, change frequency in accordance
with the evolution through time of the spectral character of ma-
rine oxygen isotope curves and other marine paleoclimate
proxies (Bradbury, 1991, 1992). Specifically, periods of 400
Ky and 41 Ky appear to dominate the climate variations prior
to ~900 Ka; thereafter, 100 Ky variations dominate.
3. Evidence for an extremely large lake in the Lahontan and
surrounding basins during the largest excursion toward glacial
oxygen isotope values in the Quaternary (OIS 16) supports a
connection between the sizes of ice sheets and the sizes of plu-
vial lakes in the northwestern Great Basin (Reheis, 2002).
4. Milankovitch-like periodicities were reported in the spec-
trum of a middle to late Pleistocene PSV record for Lake
Chewaucan (Negrini and Verosub, 1987; Negrini et al., 1988).
Subsequent improvements on the age control of the Chewau-
can sediments have disproved Milankovitch-like periodicities
as the origin of these PSVs (Negrini et al., 1994).
5. Several lake-level proxies in a 250-20 Ka paleoclimate re-
cord from Lake Chewaucan support an in-phase relationship
between increases/decreases in lake levels and advances/re-
treats of the North American ice sheet as inferred from marine
OIS 7-3 (Davis, 1985; Berger and Davis, 1992; Negrini et al.,
1994; Erbes, 1996). This represents a significant test of An-
tevs' jet stream migration model because the age control of the
Lake Chewaucan record is entirely independent of a priori as-
sumptions of correlations with climate change.
6. Bradbury (1991, 1992) explained the unusual occurrence
of a pluvial diatom assemblage during OIS 5 as being the result
of a tectonic or geomorphologic change that affected the local
Tule Lake subbasin. In this review paper, late Pleistocene Tule-
lake stratigraphy has been reinterpreted, resulting in a signifi-
cantly younger age for the Pumice Castle tephra (Figure 10).
Thus, the diatom assemblage is determined to be a likely ex-
pression of OIS 4.
7. OIS 4 is expressed strongly in the northwestern Great Ba-
sin by the presence of Pumice Castle set tephra layers in plu-
vial-lake sediments from three lake basins (Lake Modoc, Lake
Chewaucan, Lake Malheur).
8. An exceptionally precise and accurate lake-level prehis-
tory for Lake Lahontan during the latest Pleistocene is consis-
tent with increases in lake size both when the jet stream passes
by the Lake Lahontan area on its way south during OIS 2 ice
sheet advance and also when it passes by on its way north dur-
ing ice sheet retreat (Benson and Thompson, 1987a, 1987b;
Benson etal., 1995).
9. Higher-frequency climate-change events inferred by stable
isotope and total organic carbon analyses of Pyramid Lake bot-
tom sediments have the same square-wave type morphology
and duration as the D-0 cycles and Heinrich events observed in
paleoclimate records from marine sediments and ice cores
(Benson et al., unpublished).
10. The observation by Allison (1945) of several meters of
Lake Chewaucan sediments above the Mount Mazama ash was
a key piece of data that inspired Antevs' theory of a moist
Anathermal age. Improvements in tephrochronology and abso-
lute-age dating demonstrated that this ash was, instead, the
much older Mount St. Helens Cy tephra (Allison, 1966a;
Davis, 1985).
11. Bedwell (1973) noted the lack of middle Holocene arche-
ological remains in the Connley Caves of the Lake Fort Rock
basin. An analysis of the associated archeofauna revealed the
near disappearance of waterfowl remains in the area, prompt-
ing Grayson (1979, 1993) to hypothesize a middle Holocene
human abandonment of this basin because of lowered water
levels in a nearby lake/marsh complex. Subsequent work on
dune sites in the area demonstrated that Lake Fort Rock basin
was indeed occupied during the middle Holocene and that the
presence offish and waterfowl remains ... challenge notions
of a long uninterrupted mid-Holocene drought" (Mehringer
and Cannon, 1994:321).
12. As more records of Holocene lake-level changes have ac-
cumulated, it is becoming more and more apparent that high-
frequency change characterized this entire period of time, as
suggested by Mehringer and Cannon (1994). Against this back-
ground, some long-term trends have emerged (e.g., lower lake
levels during the middle Holocene) that loosely fit Antevs'
Neothermal age divisions. Neither the long-term changes nor
the high-frequency changes appear to correlate unambiguously
with newer models of Holocene climate change, such as vari-
ability in solar luminosity.
NUMBER 33 43
REMAINING PROBLEMS
Despite the progress that has been made during the last 50
years, the potential for understanding climate change from
northwestern Great Basin pluvial lake records has hardly been
tapped. For example, long-term records from pluvial lakes
other than Lake Modoc need to be obtained to determine if the
climate record from the Tulelake core is representative of the
rest of the northwestern Great Basin, considering its proximity
to the Pacific Ocean and the fact that Lake Modoc drained into
the Pacific after reaching only modest depths.
The source vents and correlative proximal volcanic rocks
have yet to be identified for many of the water-lain tephra beds
in the northwestern Great Basin. Further work of this nature
will improve the overall chronology of the region because it is
inevitably easier to apply radiometric- and radiogenic-age
methods to proximal volcanic rocks than to water-lain ashes. In
addition, promising dating techniques, such as luminescence
methods, should be used more frequently on the water-lain
ashes.
The only well-constrained pluvial lake-size record in the
northwestern Great Basin is that of the Pyramid Lake subbasin
of Lake Lahontan for the time interval from -30 to 9 14C Ky
B.P.; the Walker Lake subbasin may be well-constrained for the
late Holocene. The provocative results from these prehistories
need to be tested with records of similar resolution and accu-
racy from other basins.
The present state of the composite Holocene record does not
allow accurate testing of models for global climate change dur-
ing the Holocene. This problem can only be alleviated with
well-dated, high-resolution, continuous Holocene records from
several basins within the northwestern Great Basin.
THE NEED FOR COMPREHENSIVE CORING PROGRAMS
Most of the problems mentioned above could be solved with
a comprehensive program of coring bottom sediments
throughout the northwestern Great Basin. Although core stud-
ies can never completely replace field-based surface studies, at
this stage in the study of the northwestern Great Basin an ex-
tensive coring program would most efficiently provide a
wealth of hypotheses. These hypotheses could be tested later
by well-focused studies on the extensive exposures of lake de-
posits in this region, which represent much of the Quaternary
period.
There are several indicators that point to the success of such
an extensive coring program, many of which are discussed in
this review. The work of Benson and others (see "Latest Pleis-
tocene," above) has demonstrated that geochemical climate
proxies in cores provide estimates of lake sizes through time
that are consistent with independently derived lake-size prehis-
tories from surface materials and that these proxies can be ac-
quired relatively rapidly with high resolution. The consistency
of paleoclimate results from geochemical measurements with
other bottom-sediment proxies, such as the distribution of fau-
nal and floral taxa, lithostratigraphy, granulometry, and envi-
ronmental magnetism, points to a variety of useful techniques
that may be applicable for each core.
Unprecedented age control for terrestrial records can be pro-
vided by an already rich, tephrochronological database, which
will only improve through iteration as results from additional
cores are added to a composite data set. Under ideal condi-
tions, the promising and complementary application of paleo-
magnetic secular variation often can provide nearly continu-
ous relative age control, which conceivably can be tied
directly to marine paleoclimate records. This method, too, will
grow in usefulness as a reference PSV record continues to de-
velop. Advances in luminescence and radiocarbon methods
have contributed greatly to our understanding of northwestern
Great Basin pluvial lakes and will continue to do so in the fu-
ture.
Coring programs should be preceded by careful consider-
ation of coring sites optimally selected on the basis of sound
sedimentological models and geophysical studies. Several lo-
calities around most basins should be cored in order to sample
and compare the many varieties of depositional environments
observed within each basin. At least two cores should be taken
at each coring site to guard against localized synsedimentary
deformations, loss of core samples during withdrawal, etc.
Cores should be oriented azimuthally whenever possible to en-
able the accurate measurement of magnetic declination for use
in PSV studies. Finally, and perhaps most important, a full ar-
ray of lithological, paleontological, geochemical, tephrochro-
nological, and magnetic studies should be done on these cores
in a timely fashion and with careful management of data and
archival of core splits.
Acknowledgments
Athough the content and interpretations expressed herein are
ultimately the responsibility of the author, this work benefited
greatly from conversations or written correspondences with D.
Adam, K. Adams, L. Benson, P. Bradbury, A. Cohen, A. Dan-
sie, J.O. Davis, R. Elston, D. Erbes, F. Foit, R. Forester, D. Fre-
idel, K. Gobalet, T. Jones, T. Lowenstein, S. Lund, F. Nials, M.
Palacios-Fest, A. Sarna-Wojcicki, R. Thompson, and P.
Wigand. Also, K. Edgett, P. Wigand, and M. Sutton provided
copies of key pieces of literature cited in this paper. I am grate-
ful for the use of unpublished data provided by J. Licciardi and
S. Lund. M. Quilliam assisted with library research and proof-
reading. The final version of this paper benefited greatly from
careful review by A. Cohen, T. Lowenstein, J. Luhr, and an
anonymous reviewer.
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Pliocene to Middle Pleistocene Lakes
in the Western Great Basin:
Ages and Connections
Marith C. Reheis, Andrei M. Sarna-Wojcicki, RichardL. Reynolds,
Charles A. Repenning, and Martin D. Mifflin
ABSTRACT
During the Pliocene to middle Pleistocene, pluvial lakes in the
western Great Basin repeatedly rose to levels much higher than
those well-documented levels of the late Pleistocene. Sedimento-
logic, geomorphic, and chronologic evidence indicates that Lakes
Lahontan and Columbus-Rennie were as much as 70 m higher at
about 650 thousand years ago (Ka) than during their late Pleis-
tocene highstands. At its former altitude of 1400 m, Lake Lahon-
tan submerged present-day Reno, Carson City, and Battle
Mountain in Nevada and flooded other now-dry basins. To the
east, Lakes Jonathan (new name), Diamond, Newark, and Hubbs
reached highstands during the early-middle(?) Pleistocene that
were 25?40 m above their late Pleistocene shorelines; at these very
high levels, the lakes became temporary or permanent tributaries
to the Humboldt River and, hence, to Lake Lahontan. This sce-
nario presents a startling vision of a western Great Basin in which
climatic and tectonic effects conspired to integrate presently iso-
lated basins into one enormous drainage system during at least one
brief period during the Quaternary. Such a temporary connection
could have permitted fish to migrate from the Humboldt River
southward into the presently isolated Newark Valley and from
Lake Lahontan into Fairview Valley. The timing of drainage inte-
gration provides suggested maximum ages for fish to populate the
basins of Lake Diamond and Lake Jonathan.
Marith C. Reheis and Richard L. Reynolds, United States Geological
Survey, MS-980, Federal Center, P.O. Box 25046, Denver, Colorado
80225, USA. Andrei M. Sarna-Wojcicki, United States Geological Sur-
vey, MS-975, 345 Middlefield Road, Menlo Park, California 94025,
USA. Charles A. Repenning, 9010 West 5th Place, Lakewood, Colo-
rado 80226, USA. Martin D. Mifflin, Mifflin and Associates, HCR38,
Box 126, Las Vegas, Nevada 89124, USA.
Reviewers: Kenneth L. Adams, Desert Research Institute, 7010 Dan-
dini Boulevard, Reno, Nevada 89512, USA; M.A. Chan, Department
of Geology and Geophysics, University of Utah, Salt Lake City, Utah
84112, USA; Robert Hershler, Division ofMollusks, National Museum
of Natural History, Smithsonian Institution, Washington, D.C. 20560-
0118, USA; D.E. Wilkins, University of Utah, Salt Lake City, Utah
84112, USA.
The Walker Lake subbasin of Lake Lahontan held deep lakes at
least four times during the early Pleistocene; these were at least
15-25 m above the Sehoo highstand. One site on the Humboldt
River in the northeastern Lahontan Basin displays exposed, faulted
lake sediments of early Pleistocene age that are now 40-55 m
above the Sehoo limit. These sediments suggest that the northern
and Walker Lake subbasins were contiguous parts of Lake Lahon-
tan by about one million years ago (Ma), but fish fossils indicate
that drainage connections existed before 1 Ma. For example, high
beach deposits suggest that Lake Russell, now a tributary to the
Owens River, previously overflowed to the Walker River, poten-
tially providing a migration route for fish southward along the east
side of the Sierra Nevada into the Owens River system. Deposits
of Pliocene lakes suggest connections existed among proto-Lakes
Lahontan, Columbus-Rennie, Clayton, and possibly Tonopah.
Tephrochronologic ages of these Pliocene lakes provide con-
straints on the availability of aquatic migration routes at different
times.
Reconstructing and dating these lake levels also has important
implications for paleoclimate, tectonics, and drainage evolution in
the western Great Basin. For example, shorelines in several basins
form a stair-step sequence downward with time from the highest
levels, thought to have formed at about 650 Ka, to the lowest,
formed during the late Pleistocene. This descending sequence indi-
cates progressive drying of pluvial periods, possibly caused by
uplift of the Sierra Nevada and other western ranges relative to the
western Great Basin. These effects, however, cannot account for
the extremely high lake levels during the early-middle Pleistocene;
rather, these high levels were probably caused by a combination of
increased effective moisture and changes in the size of the Lake
Lahontan drainage basin.
Introduction
Pluvial lakes in the western Great Basin have been the focus
of many classic studies on paleoclimate and on biogeographic
distributions of aquatic species, such as fish and snails. Until
recently, these studies have focused on sedimentary records of
the late Pleistocene and on the distributions of modern species
(especially fish), because few pre-late Pleistocene deposits of
53
54 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
pluvial lakes were known and because fossil localities were
scarce and commonly not well dated. In their seminal work,
Hubbs and Miller (1948) argued for relatively rapid evolution
of fish in now-isolated drainage basins on the assumption that,
in most cases, the last aquatic connections between the fish-
bearing basins were during the last pluvial period, now known
to have ended about 12 thousand years ago (Ka). They recog-
nized that in some cases modern sills between basins were far
too high to have been exceeded by late Pleistocene lake levels;
they called upon transfers by interdivide stream diversions or
"pre-pluvial" connections to explain these fish distributions. As
more fossil localities were identified and dated, research in the
last 20 years has shown that many modern species of fish and
springsnails are very similar to ancestral species that were
widely distributed in the western Great Basin and areas to the
north and south during the Pliocene (Miller and Smith, 1981;
Taylor and Smith, 1981; Taylor, 1985; Minckley et al., 1986;
Gobalet and Negrini, 1992; papers in this volume). As a result,
less credence has been placed in rapid evolution, and modern
aquatic-species distributions are more commonly attributed to
inheritance from widespread Tertiary populations followed by
extinctions and intermittent population-mixing caused by
wet-dry climatic cycles and tectonic disruptions during the
Pleistocene.
The possibility that modern species-distributions could be
produced in part by aquatic migration routes older than late
Pleistocene in age but younger than Tertiary in age has been
given little consideration in biogeographic studies for the very
good reason that such intermediate-age migration routes were
largely unknown. This paper documents connections among
several large lakes in the western Great Basin during part of the
early or middle Pleistocene but that probably have been severed
for the past one-half million years or more. Arguments for the
evolution of modern species from widely distributed ancestral
ones that suggested rapid evolutionary changes took place dur-
ing the last few thousand years (e.g., Hubbs and Miller, 1948)
have been challenged recently (Smith et al., 2002) as being not
representative of the much slower rates observed in the fossil
record during the past several million years. We suggest that if
the observed divergence between some species can be averaged
over a half a million years rather than over a few thousand, the
apparent discrepancy in evolutionary rates may be greatly re-
duced. This new information at least provides alternative hy-
potheses about rates of evolution from independent evidence,
as urged by Smith et al. (2002). In some cases, these older
Pleistocene lakes may require revision of previous interpreta-
tions about ages and relations among fossil fish (e.g., Miller
and Smith, 1981) that were based on incorrect age assignments.
In addition to far-reaching implications for biogeographic
studies, the new lacustrine records hold significance for studies
of paleoclimate, tectonics, and drainage evolution in the west-
ern oreat Basin. The existance of enormously expanded lakes
at about 650 Ka and earlier imply that much wetter conditions
prevailed at that time than are indicated in global climate
records, such as oxygen-isotope data from marine cores; how-
ever, sudden and dramatic integrations of previously discon-
nected basins with the Lahontan Basin may help to account for
the large size of Lake Lahontan in the early-middle Pleis-
tocene. Progressively shrinking lake highstands from the mid-
dle to late Pleistocene indicate different regional climatic con-
ditions. Some Pliocene and younger basins were probably
isolated by tectonic processes, such as compressional folding
and extensional faulting along the Walker Lane tectonic belt
(Stewart, 1988).
This paper focuses on pluvial lakes in the Lahontan Basin
and the basins of Columbus Salt Marsh and Fish Lake Valley,
where most of the field work in the western Great Basin has
been concentrated. Detailed descriptions of the stratigraphy,
chronology, geomorphology, and geology of pre-late Pleis-
tocene lacustrine deposits in the Lahontan Basin are presented
first, followed by a correlation of the deposits from sampled lo-
calities. Next, data from reconnaissance field work for this
study and from previously published studies are discussed
briefly for Lake Russell (Mono Basin) in western Nevada and
for lakes in central Nevada, including those in the basins of
Kobeh (Lake Jonathan)-Diamond, Long (Lake Hubbs), New-
ark, and Pine Valleys. These lakes are described in chronologic
order, beginning with Pliocene localities. A discussion of the
ages of the pre-late Pleistocene lakes in the study area is fol-
lowed by an examination of the possible causes of these very
large lakes. The paper concludes by discussing the implications
of very high lake stands and prior connections among drainage
basins on the migration of aquatic species.
PREVIOUS WORK
During its well-studied late-middle (Eetza) and late Pleis-
tocene (Sehoo) highstands, Lake Lahontan reached a maximum
altitude of about 1330 m and extended more than 350 km, from
the Nevada-Oregon border to beyond the southern end of
Walker Lake (Figure 1; Russell, 1885; Benson et al., 1990;
Morrison, 1991). Many smaller, isolated lakes also occupied
other basins in the western Great Basin (Mifflin and Wheat,
1979), but their late Pleistocene highstands are generally not
well dated, and little information has been available on older
pluvial lakes and their high shorelines.
Previous workers have observed that the Eetza (~180-130
Ka, but possibly in part as old as 300 Ka; Morrison, 1991) and
Sehoo (~35?12 Ka; Benson et al., 1990) shorelines in un-
FIGURE 1 (opposite).?Regional map showing pluvial lakes of northern
Nevada and vicinity, late Pleistocene shorelines, and areas of maximum middle
Pleistocene (where known) and late Pleistocene lakes. See Table 1 for shoreline
altitudes. Triangles show study sites (data for selected sites is in Table 1);
arrows on drainages show possible links between basins during highest lake
stands. Localities: CD = Carson Desert, CLV = Clayton Valley,
CV = Campbell Valley, FLV = Fish Lake Valley, HL = Honey Lake,'
LCR = Lake Columbus-Rennie, LM = Lone Mountain, LT = Lone Tree Hill,'
MW = McGee Wash, PV = Pine Valley, RM = Rhodes Salt Marsh, RP = Rye
Patch Dam, SA = Sunshine Amphitheater, SC = Smoke Creek Desert,
TB = Thorne Bar, WL = Walker Lake. Towns: B = Battle Mountain,'
C = Carson City, E = Elko, R = Reno, W = Wei Is, WI = Winnemucca.
NUMBER 33
121?
55
117?
42
Area of late Pleistocene lakes
(not all are shown)
Maximum known area of middle Pleistocene lakes
f Possible additional area of middle Pleistocene lakes
Early and middle Pleistocene lake deposit locality
& Pliocene lake deposit locality
Late Pleistocene overflow
Inferred earlier overflow and modern sill
height (m) above late Pleistocene shoreline
Late Pleistocene boundary of Lahontan basin
Possible additional area of basin in
middle Pleistocene
Major drainages
56 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
faulted positions throughout the Lahontan Basin lie at about
1330 ? 12 m (Mifflin and Wheat, 1979; Benson, 1978). Iso-
static rebound has caused shoreline altitudes to range from
about 1318 m at the northeastern margin of the lake to as much
as 1341 m in the Carson Desert (Adams, 1996, 1997). The
highest vertical faulting rates of late Pleistocene shorelines are
along the Wassuk fault, which bounds the west side of the
Walker Lake subbasin (Figures 1, 2). Demsey (1987) estimated
a Holocene slip rate of 0.4-0.5 mm/yr and total offset of 6-7 m
of the Sehoo shoreline along this fault.
Morrison (1991) summarized the previously known ages and
altitudes of early-middle and early Pleistocene nearshore and
deltaic deposits of Lake Lahontan. Four sites were in outcrops
below the Sehoo shoreline where incision exposed older depos-
its buried by Sehoo- and Eetza-aged sediment. The fifth site,
Thorne Bar, is described below. The pre-Eetza lake cycles
identified by Morrison (1991) are the Rye Patch Alloformation,
from about 700 to 600 Ka, and two sets of unnamed deposits
containing and underlying the Glass Mountain tephra, which
were erupted about one million years ago (Ma) (only found in
the Walker Lake subbasin). He believed that the 0.76 My old
Bishop ash bed was deposited during a long period of alluvial
119?15' 119?00' 118?45' 118?30'
- Adrian
-4T Valley
EXPLANATION
Area of late Pleistocene
lake with 1330-m shoreline
j 1 Area of middle Pleistocene
lake with 1400-m shoreline
Probable additional area of
middle Pleistocene lake
accounting for sedimentation
? Locality of pre-late Pleistocene
MW lacustrine sediment or shoreline
FIGURE 2.?Map of Walker Lake basin and southernmost Carson Desert (Raw-
hide Flats) showing study sites and lake areas. Area of lake at 1400 m is drawn
using modern topography, acknowledging that this is only an approximation of
the topography in the middle Pleistocene because of subsequent isostatic
rebound, faulting, erosion, and sedimentation. From north to south, letters indi-
cate the following: SA = Sunshine Amphitheater, CV = Campbell Valley,
WR = Weber Reservoir, MW = McGee Wash, TW = Twentymile Beach, TB =
Thorne Bar, SSV = Soda Springs Valley.
and eolian sedimentation (the Lovelock Alloformation) during
which no deep lakes existed in the Lahontan Basin; however,
the Bishop ash bed is contained within lacustrine deposits in
basins to the south, including pluvial Lakes Rennie (Fish Lake
Valley; Reheis et al., 1993), Owens (Smith et al., 1997), Tecopa
(Sarna-Wojcicki et al., 1987), and Manley (Death Valley; Knott
et al., 1996). Also, an ash bed with Bishop-like chemical com-
position was identified in beach gravel exposed in a fault scarp
below the Sehoo shoreline south of Schurz, Nevada (Davis,
1978).
Only in the Walker Lake subbasin of Lake Lahontan are
there previously known outcrops of beach gravel at altitudes
higher than the Sehoo and Eetza shorelines. W.J. McGee ob-
served that the eastern end of the "crest of the gravel embank-
ment separating Walker Lake Valley from the... Walker River
... is now fully 200 feet above its original position" (that is, the
Sehoo shoreline) (Russell, 1885:142, pi. 28). This locality,
which apparently had not been revisited until the time of the
present study, is southeast of Weber Reservoir (Figure 2) and
is informally named McGee Wash (Morrison and Davis,
1984). The Thorne Bar site, at the southeast edge of Walker
Lake (Figures 2, 3) was first reported by King (1978) and Mif-
flin and Wheat (1979) to have pre-Eetza "shore gravel with
layers of lake-deposited carbonate, preserved in a much-
eroded V-bar complex...up to 1443 m altitude" (Morrison,
1991:296). At the McGee Wash site, Russell (1885) thought
the gravels had been folded into an anticline; mapping for this
study (see Figure 7) and that of Morrison and Davis (1984)
confirms the presence of numerous faults and tilted strata. The
beach ridge at Thome Bar has been attributed either to faulting
(King, 1978:67-68) or to northward tilting of the entire
Lahontan Basin (Mifflin and Wheat, 1979; Mifflin, 1984;
Morrison, 1991). Detailed leveling surveys by Demsey (1987),
however, revealed no evidence of Quaternary faults or defor-
mation either at Thorne Bar or northward along the Gillis
Range, in contrast to several localities of superelevated beach
gravel along the footwall block of the Wassuk Range fault
(King, 1978; R. Bucknam, U.S. Geological Survey, in litt. to
M.C.R., 1996).
Little work has been done on other pluvial lakes in western
Nevada (Figure 1). Mifflin and Wheat (1979) mapped late
Pleistocene shorelines for all of the pluvial lake basins of Ne-
vada. These authors also briefly noted higher, older shoreline
features of Lakes Diamond, Newark, and Hubbs; Mifflin
(1984) thought that these features supported his hypothesis of
the northward tilting of northwestern Nevada. Tackman (1993)
obtained ages on the late Pleistocene shorelines of Lake Dia-
mond. He also located many pre-late Pleistocene shorelines
and discussed Lake Diamond's history as controlled by over-
flow from Kobeh Valley (Lake Jonathan) and the incision of
sills. Davis (1981) investigated Tertiary lacustrine deposits in
Clayton Valley. Gilbert et al. (1968) mapped Pliocene lacus-
trine deposits interbedded with basalt and rhyolite flows on the
NUMBER 33 57
118? 40' 118? 39' 118 118? 37'
Qfh
Qfy
Qly
-Qfm-
23 m omitted
North of Thome Bar
Section 1
Pebbles lagged on bedrock slope
up to altitude 1398 m (4585 feet)
i? 5
1?0
Variable
amounts
of tufa
cement
Fishy bones
EL21-WL(Bishop ash?)
Bwk--20 cm
!] Bw?--15cm
Btk-30 cm
Matrix-poor
Btk-50 cm
Btk-90 cm
Btk--30 cm
EXPLANATION
Sandstone and silty
sandstone, locally
conglomerate beds
Sandy pebble con-
glomerate, well
rounded clasts (near-
shore deposits)
Sandy cobble
conglomerate, well
rounded clasts
(beach deposits)
Boulder conglomerate,
poorly bedded
(shore-face deposit)
o o"<
0 0 0
O 0 <
0 0 0
A -4 ''A
XXX
EL21-WL
Sandstone and con-
glomerate, poorly to
moderately sorted
Volcanic rocks
Steep dips
(backset beds)
Calcareous cement,
locally tufa
Tephra layer and
sample number
Buried soil and
diagnostic horizon
thickness
2 KILOMETERS
1 MILE
CONTOUR INTERVAL: 40 FEET;
200 FEET IN MOUNTAINS
MAP EXPLANATION
Fan deposits (Holocene); includes
some Holocene lake deposits
Fan deposits (late Pleistocene)
Deposits of late Pleistocene
(Sehoo) lake (unit QI7T)
Fan deposits (middle Pleistocene)
Fan deposits (middle
and early? Pleistocene)
Qlo
\/ A A '
/TVS. 7
Nearshore deposits and beach gravel of middle
to early Pleistocene lakes (units QI1T to QI6T)
Volcanic and sedimentary rocks
(Pliocene and Miocene)
Late Pleistocene shoreline
Middle Pleistocene shorelines;
dashed where inferred
Tephra locality and measured section
Soil pit (Adams and Wesnousky, 1999)
FIGURE 3.?Geologic map of area around Thorne Bar (Figure 2, Table 1) and measured section north of Thome
Bar (triangle in northern map area).
58 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 1.?Information for representative sites (see Figure 1; within each basin, sites are in order beginning in
the south and proceeding clockwise or northward).
Site General location Latitude (N) Longitude (W)
Altitude range of
nearshore features (m) Type of deposit or shoreline feature
Lake Lahontan (late Pleistocene shoreline, 1330 ? 12 m)
US-95 roadcut
Thorne Bar area
20-Mile Beach
McGee Wash
Campbell Valley
Sunshine Amphi-
theater
Clifton site
1-80 roadcut (?)
Secret Valley cutoff
Burro Mountain
Freds Field
Iron Point
Treaty Hill area
Lone Tree Hill
Drumm Summit
Rhodes Salt Marsh
Southeastern Walker
Lake valley
West-central Walker
Lake valley
Northern Walker Lake
valley
Northern Walker Lake
valley
Northern Walker Lake
valley
Carson River canyon
Truckee River canyon
North of Honey Lake
valley
Smoke Creek Desert
Smoke Creek Desert
Humboldt River
Humboldt River
Humboldt River
Fairview Valley
38?14.7'
38?39'-44'
38?46.6'
38?45.1'
39?01'-03.5'
39?04-O7'
39?08.6'-09.4'
39? 18.1-18.4'
39?30.8'
40?25.0'
40?32.5'
40?44.7'-45'
40?56.6'-59.5'
40?50.3'-58.0'
40?49.4'-50.5'
39? 18-18.5'
118?03.3'
118?37.3'-39'
118?45.4'
118?46.5'
118?47.5'-50.7'
118o59'-119?01.2'
118?59.5'-119?00.7'
119?27.4'-28.0'
119?38.7'
120?16.9'
119?50.6'-51.4'
119?31.4-32.2'
117?17.2-19.5'
117?04'-12'
117?12.0'-12.7'
118?07.9'
1402-1423
1360-1400, 1350-1365
1330-1350
Depends on footwall
location; <1475
1420-1433, 1393-1402
1347-1365
1378-1390,1356-1378
1341-1426
1333-1380
1360
1370-1372
1396-1402, 1373
1378-1400, 1350-1370
1353-1368
1344-1362, 1390-1414
1326-1414
1355-1384
Sinter mound
Borrow pit
Rock Hill
Coaldale Junction
Fish Lake Valley
Southern Columbus
Salt Marsh
Northern Columbus
Salt Marsh
Lake Columbus-Rennie (late Pleistocene shoreline, 1402 m)
1442-1460
38?00'-00.9'
38?08.4'-09'
118?00'
118?01.2'-02.2'
118?56.9-57.8'
Southeastern Columbus 38?00.7'-01.5' 117?52.6'-53.8'
Salt Marsh
1442-1452,
1426-1433, 1417
1390-1417
1422-1430
Lake Russell (Mono Basin) (late Pleistocene shoreline, 2155 m)
Southern outlet area
Rancheria Gulch
Kirkwood Spring
Andesite sill
US-50 site
Slough Creek
Borrow pit on
county road
Eastern Mono Basin
Northern Mono Basin
Northern Mono Basin
Northeastern Mono
Basin
Eastern Kobeh Valley
Southwestern side Lone
Mountain
Northeastern Kobeh
Valley
38?00'-04'
38?05.5'-06'
38?08.7'-09'
38?18.5'
H8?44'_47.5'
119?05'-06'
118?58'-59'
118?46.7'
2207-2277
<2268,2195-2225
2205-2252
2237
Lake Jonathan (Kobeh Valley) (no late Pleistocene lak
39?34.1'-35.3'
39?33.5'-34'
39?42.0'
116? 16.6-17.4'
1871-1920
1856-1890
1914-1925
Nearshore and sublacustrine deltaic deposits
Tufa-cemented beach gravel and sand; 3
shoreline levels, 6 ages of deposits
Tufa-cemented beach gravel and sand; 2
shoreline levels, 2-4 ages of deposits
Lake mud, silt, sand, and gravel; 3 shoreline
levels, 5 ages of deposits
Lake mud, silt, sand, and gravel; shoreline
angle; 2 shoreline levels, 3 ages of deposits
Lake mud, silt, sand, and gravel; 3 shoreline
levels, 5 ages of deposits
Beach or fluvial gravel lag on bedrock
Deltaic(?) and fluvial deposits
Large botryoidal tufa heads, locally attached
to well-rounded clasts; scattered patches
of well-rounded pebbles and cobbles
Beach gravel lag and tufa; 2 different levels
Rounded, water-polished gravel; 2 different
levels
Beach or fluvial gravel lag
Sand and gravel in extensive delta plain to
north; upper level is berm on Treaty Hill
Lake mud, silt, sand, and gravel; 1 shoreline
level, 3 or 4 ages of deposits
Lake mud and silt (Pliocene?) overlain by
beach(?) gravel of uncertain age
Sinter-cemented silt and sand; Bishop ash
(Reheis et al., 1993)
Lake silt, sand, and gravel; shoreline angle; 3
shoreline levels, 4 ages of deposits
Beach sand and gravel in moderately pre-
served berm; pit exposes buried lacustrine
deposits containing Bishop ash
Beach sand and gravel in eroded berms
Lake silt, sand, gravel, and beach rock; at
least 2 ages of deposits, oldest is late
Pliocene; upper limit defined by beach
gravel lagged on bedrock
Beach gravel; 2 shoreline levels, lower
marked by berm
Scattered beach cobbles lagged on bedrock
Bedrock sill bearing scattered rounded clasts
Lake silt, sand, and gravel; 2 ages of deposits
Lake clay, silt, sand, and gravel; 3 ages of de-
posits
Lake sand and gravel; 3 ages of deposits sep-
arated by paleosols; well-preserved berms
in vicinity
NUMBER 33 59
TABLE 1.?Continued.
Site General location Latitude (N) Longitude (W)
Altitude range of
nearshore features (m) Type of deposit or shoreline feature
Lake Newark (late Pleistocene shoreline, 1847 m)
39?27'-27.4' 115?39.6'-40.6' 1856-1859, 1850-1855 Beach sand and gravel; 2 shoreline levels,
West of Beck Pass
Stage Road
Maple Syrup well
Long Valley Canyon
Elko-White Pine
County line
Independence
interchange
Wells
Angel Island
Emigrant Pass
section
The Gap
Redlich Summit
Southern Newark
Valley
Eastern Newark Valley 39?36'
Northern Newark
Valley
115?39.0'
39?52.7'-53.5' 115?39.2'-40.4'
more eroded upward
1850-1855 Beach sand and gravel in moderately pre-
served berm
1887-1890, 1868-1875, Beach gravel in berms and as lag on bedrock;
1856-1859, 1850-1855 3 or 4 shoreline levels (possibly one re-
peated by fault), more eroded upward
Lake Hubbs (Long Valley) (late Pleistocene shoreline, 1920 m)
Southern Long Valley 39?32.2'-34' 115?25'-28' 1935-1940, 1920-1927, Beach sand and gravel; 3 different shoreline
Eastern Long Valley 39?35.3' 115?20.8'
1907-1912
1930-1940?,
1920-1927
levels, more eroded upward
Beach gravel in moderately preserved berm;
lag gravel above
Lake Gale (late Pleistocene shoreline, 1915 m?revised from 1905 m as stated by Mifflin and Wheat, 1979)
Northeastern Butte
Valley
Independence Valley
40?03.5'-07.5' 115?01'-04' 1917-1922
Lake Clover (late Pleistocene shoreline, 1730 m)
41?05.5'-06.3' 114?42.2'-43' 1754-1763
Humboldt River head- 41?06.4'-07.4' 114?58.5'-59.1'
waters
1737-1750
Eastern Clayton Valley 37?40M4'
Pliocene localities
117?32'-35' 1340-1440
Northeastern Fish
Lake Valley
Columbus-Fish
Lake sill
Rhodes-Columbus
divide area
117?55' 37?53.4'
37?54.5'-57' 117?57'-59'
1475-1560
1432-1463
38?12.5'-14.5' 117o59'-118?01.5' 1450-1585
Beach sand and gravel in subdued berms
above well-defined late Pleistocene shore-
line features
Beach (?) gravel exposed in 1-80 borrow pits
North-dipping delta foreset beds of gravel,
sand, and silt; locally tufa-cemented; lo-
cally overlain by marl beds
Indurated lake clay, silt, sand, and gravel; at
least 2 ages of deposits, many tephra lay-
ers; Pliocene (Davis, 1981)
Indurated lake clay, silt, and sand; at least 2
ages of deposits, many tephra layers;
Pliocene (Reheis et al., 1993)
Indurated lake clay, silt, and sand; at least 2
ages of deposits, many tephra layers;
Pliocene
Indurated lake clay, silt, sand, and gravel; at
least 2 ages of deposits, many tephra lay-
ers; Miocene to late Pliocene
east side of Mono Lake. Reheis et al. (1993) documented sev-
eral Pliocene to middle Pleistocene cycles of pluvial Lake Ren-
nie in Fish Lake Valley (Figure 1) and proposed that Lake Ren-
nie merged with Lake Columbus to the north about 760 Ka at a
shoreline altitude of about 1460 m.
METHODS
The vastness of the Lahontan Basin and the large number of
other pluvial lakes in Nevada precluded a systematic search for
shorelines older and higher than those of late Pleistocene age.
This search relied heavily upon (1) extrapolation from known
localities in the Walker Lake subbasin and in Newark, Long,
and Diamond Valleys (Figure 1), (2) Mifflin's recollections of
possible shoreline localities from field work in the 1970s, and
(3) observation from roads. Geologic maps were searched for
outcrops of late Tertiary or Quaternary fine-grained deposits;
several sites in the northern Walker Lake subbasin, for exam-
ple, had been mapped as Miocene to Pliocene sedimentary
rocks (Greene et al., 1991) but in fact are Quaternary lacustrine
deposits. Potential sites were examined on l:24,000-scale topo-
graphic maps and (or) aerial photographs for relatively flat
benches above late Pleistocene shorelines. Dozens of potential
sites were checked in the field, but only those with evidence of
lacustrine deposition are listed in Table 1.
The altitudes and ages of shorelines (Table 1) were estimated
from shoreline features and from sedimentologic evidence.
Classic shoreline morphologies, such as wave-cut cliffs and
60 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
beach berms, become increasingly difficult to discern with age
and are rarely preserved after half a million years or more of
erosion. Thus, the highest altitude of beach-gravel beds or lag
on bedrock slopes was taken as a minimum shoreline altitude.
Where beach morphology was discernible, the preservation of
berm features and the degree of development of relict soils per-
mitted assignment of relative ages. Fine-grained beds and even
some gravels locally contained tephra layers; these were identi-
fied by chemical correlation with tephra of known age by the
United States Geological Survey (USGS) Tephrochronology
Laboratory. Ages of some units also were constrained by mag-
netostratigraphy and by vertebrate biostratigraphy.
Glass shards from tephra layers found within lacustrine se-
quences during mapping were analyzed by electron microprobe
for major-oxide composition (Table 2) using methods de-
scribed by Sarna-Wojcicki et al. (1984) and Sarna-Wojcicki
(2000). Sample compositions were compared with composi-
tions in a data base of previously analyzed tephra layers using
statistical programs (Table 3). Correlations were based on simi-
larity in chemical composition, petrographic characteristics
(such as shard morphology, phenocryst mineralogy), and strati-
graphic data. Values for glass compositions in Tables 2 and 3
were normalized to 100 percent to correct for the variable
amounts of hydration of the volcanic glass, as reflected by the
original totals. In addition to hydration, some postdepositional
alkali exchange had occurred in many of the samples. Variabil-
ity in sodium and potassium contents was particularly apparent
for the older tephra layers.
Samples used for paleomagnetic directions were collected as
large (typically 500-1000 cm3) blocks of sediment that were
oriented in the outcrop using a magnetic compass and a level to
mark the horizontal plane. Many samples were sufficiently in-
durated to carve from the outcrop; some friable samples were
hardened using sodium silicate before extraction. In the labora-
tory, smaller samples were shaped from block samples to fit
into plastic cubes (3.2 cm3 volume). We estimate that the final
orientation of each sample was within ?10? of the magnetic
field declination and within 10? of the original horizontal atti-
tude. Such orientation is adequate for polarity determinations at
these sites. We corrected for tectonic tilting when bedding dips
exceeded 5?. The natural remanent magnetization of the ori-
ented samples was measured using a 90 Hz spinner magnetom-
eter having a sensitivity better than 10~5 amperes per meter
(A/m). Samples were demagnetized by alternating-field (AF)
demagnetization with peak fields of 5, 10, 20, 40, 60, and 80
milliteslas (mT). We evaluated the paleomagnetic reliability of
each sample from an examination of orthogonal demagnetiza-
tion diagrams that show the change in remanent direction and
magnitude in response to each demagnetization step (Zijder-
veld, 1967). Samples that showed unstable behavior (lack of
systematic decay in magnitude and direction toward the origin
of the diagram over most of the demagnetization range) were
not considered further and are represented by a question mark
on the diagrams of stratigraphic sections. Polarity was deter-
mined for stably magnetized samples: normal, when the rema-
nent vector was northward (clockwise between 60? and 300?)
with a moderate to steep, positive inclination (between 20? and
80?); reversed, when the remanent vector was southward
(clockwise between 120? and 240?) with a moderate to steep,
negative inclination. The expected inclination for the latitude
of 39?N through the study region is about 58?. Several samples
exhibited magnetic behavior that reflected the removal of a
normal polarity overprint in response to low AF inductions to
isolate a reversed component of magnetization. Although incli-
nations in these samples may not stabilize in moderately steep
angles, the magnetization clearly is reversed. Such samples are
listed as having a reversed component (R comp; see Figure lie
for sample P10). Samples that showed stable magnetization are
listed as questionably normal ("N?") or questionably reversed
("R?") if they had moderate to steep positive or negative incli-
nation, respectively, with a declination beyond ranges for a
clearly defined polarity.
ACKNOWLEDGMENTS
This project was initially funded as a Gilbert Fellowship by
the USGS; Martin D. Mifflin was the coproposer with Reheis
and was the behind-the-scenes advisor, freely sharing his inspi-
ration and vast store of Lahontan Basin knowledge. Roger B.
Morrison reviewed early versions of the proposal and kindly
led us to several areas with "old" Lake Lahontan deposits. Ken-
neth D. Adams (University of Nevada, Reno) made observa-
tions at the Campbell Valley and Sunshine Amphitheater sites
and subsequently made the sites available to us for study. Scott
Stine (California State University at Hayward) participated in
the initial, startling finds of very high beach gravels in the
Mono Basin. Bruce Braginton (Newmont Gold Corporation)
graciously arranged and shepherded our field work in the Lone
Tree Mine and gave us access to one of the long cores. James
Honey (University of Colorado) and Robert R. Miller (Univer-
sity of Michigan) kindly identified some of the vertebrate fos-
sils. Richard M. Forester and J. Platt Bradbury (USGS) identi-
fied and interpreted ostracode and diatom samples,
respectively. Charles E. Meyer (USGS) performed much of the
tephrochronologic analyses. Joseph G. Rosenbaum and Will-
iam C. Rivers (USGS) instructed Reheis on the initial analyses
of paleomagnetic samples and advised her on their interpreta-
tion. Jonathan Rademaekers and Michael O'Donnell (USGS)
produced the color figure using Geographic Information Sys-
tems (GIS) techniques; George Garcia (USGS) provided some
of the computer drafting. Finally, Reheis thanks her field assis-
tants, who worked very hard in bad weather and rough camping
conditions: Andrea Lonner, Nadine Calis, Cassandra Fenton,
Rebecca Gagliardi, Kathleen Johnson, Marketta McGuire, and
Joanna Redwine.
NUMBER 33 61
Evidence for Old Pluvial Lakes
in the Western Great Basin
We have adopted several conventions for ordering the infor-
mation presented and for naming the lacustrine units so as to
accomodate the complex geographic, stratigraphic, and chro-
nologic relationships among the many sites discussed in this
paper. Because most of the work to date has focused upon the
Lahontan Basin and upon the Walker Lake subbasin in particu-
lar, we begin by describing these primary study sites from
south to north, followed by evidence from supporting locali-
ties, also from south to north. Other pluvial-lake records in
western and central Nevada are described in chronological or-
der by lake basin, working roughly from southwest to north-
east. At each study site, the lacustrine units were numbered in
order from oldest to youngest, followed by a subscript letter
designating the site: for example, the youngest lake unit at
Thome Bar is Q17T. As a result of differential deposition and
preservation, this convention can produce same-numbered
units of different ages at different study sites; however, it
avoids bias in the nomenclature and in the correlation of some
lake units that are undated or whose ages are only relative.
Based upon current evidence and best judgment, we then corre-
late the units from site to site and assign names to deposits of
similar age; names were chosen for localities where a unit is
well exposed. For example, deposits associated with the 1400
m highstand of Lake Lahontan were named "lake deposits of
McGee," for the well-preserved arcuate beach ridge above Mc-
Gee Wash that was deposited during this highstand.
THE LAHONTAN BASIN
The Lahontan Basin consists of a sprawling complex of sub-
basins (Figure 1), many of which have independent hydro-
graphic histories (Davis, 1982; Reheis and Morrison, 1997)
controlled not only by climate but also by interbasin river di-
versions and changes in threshold altitudes. Of the four princi-
pal rivers?the Truckee, Carson, and Walker, draining the Si-
erra Nevada, and the Humboldt, draining northern and eastern
Nevada?all but the Carson River have been diverted among
different subbasins, some more than once. Thus, lake levels re-
corded in some subbasins may not apply to other subbasins, es-
pecially at moderate to low water levels.
EVIDENCE FROM THE WALKER LAKE SUBBASIN.?The
Walker Lake subbasin (Figures 1, 2) contains much of the
stratigraphic evidence of early and middle Pleistocene lakes
with high shorelines in the Lahontan Basin. This evidence must
be interpreted with caution because the area is tectonically ac-
tive. The Walker Lake subbasin is bounded on the west by the
Wassuk Range and the active Wassuk normal fault and is
bounded on the east by the Gillis Range. At the north end of the
Gillis Range, faults of the strike-slip Walker Lane belt (Stew-
art, 1988) intersect the basin. The Thorne Bar area lies in a tec-
tonically quiescent area of the basin (Demsey, 1987), but other
sites discussed below have been affected by faulting. The
Walker River, the only major source of water, arises in the Si-
erra Nevada and presently feeds Walker Lake. At times in the
past, however, the Walker River flowed north into the Carson
Desert via Adrian Valley (Morrison and Davis, 1984; Bradbury
etal., 1989; King, 1993).
Thorne Bar Area: The beach ridge, or bar, at Thorne Bar is
a large, V-shaped barrier complex of shore gravel built at the
mouth of a large canyon that drains the Gillis Range (Figures 2,
3). The part of this complex above the Sehoo shoreline was
previously described as pre-Sehoo lacustrine gravel that was
elevated by northward tilting of the Lahontan Basin (King,
1978; Mifflin and Wheat, 1979; Morrison, 1991). Detailed lev-
eling studies by Demsey (1987) and interpretation (this study)
of aerial photographs revealed no evidence of Quaternary de-
formation on this side of the Walker Lake subbasin. We inter-
pret the altitudes of the bar and its morphologic subdivisions to
represent essentially undeformed shorelines.
The bar consists mainly of pebble- to cobble-sized gravel
that is locally tufa-cemented; the bedding ranges from horizon-
tal to angled at 25? or more. The bar can be divided into four
morphologic units (Figures 3, 4A). The lowest and youngest
unit is the group of historic shorelines below 1250 m in alti-
tude, formed during the last 100 years; this unit is not sepa-
rately numbered. The next higher group of shorelines is distin-
guished by sharp, well-preserved barriers (unit Q17T) that are
marked at the top by the Sehoo highstand shoreline at an alti-
tude of about 1330 m (~4360 ft; Figure 3). A small loop bar
bears a soil profile with a thin vesicular A horizon and a
slightly oxidized Bw horizon with a few thin clay films and lit-
tle pedogenic CaCO3 (Table 3, profile WL-5c of Adams and
Wesnousky, 1999). This degree of development is typical of
Sehoo-aged or younger soils (Chadwick and Davis, 1990).
Between 1330 m and about 1370 m (4360-4500 ft), the bar
complex consists of two nested, eroded V-shaped barriers; the
lower (unit Q16T, marked by vehicles in Figure 4A) is better
preserved than the upper (unit Q15T). Soil pits on the eroded
surfaces of these barriers show increasing thicknesses of Av
and Bw horizons and increasing cementation by pedogenic
CaCO3 and silica of subsurface horizons relative to barrier alti-
tude and to the Sehoo-aged soil, although clay and silt content
remain low (Adams and Wesnousky, 1999). These soil features
indicate that the barriers increase in age upward and are older
than the Sehoo highstand (Chadwick and Davis, 1990), but the
soils are not adequately preserved for a more accurate age as-
sessment. Between 1370 m and 1402 m (4500-4600 ft) the
complex has no preserved morphology, but deep arroyos have
exposed well-bedded, cemented shore gravel of unit Q14T, and
beach pebbles locally occur as lag on basalt outcrops.
In summary, the degree of morphologic preservation and the
degree of soil development in three reconnaissance soil pits (Ad-
ams and Wesnousky, 1999) indicate the presence of at least four
lacustrine units above the historic limit at 1250 m: the Sehoo-
aged bar, unit Q17T, extending upward from the historic limit to
62 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
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asi
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erof
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coin
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IIs
NUMBER 33 63
?6
B
'f-
LE 2.?Co
TAB
ituc
e Long
3
Lat
6
Total
1(O)
o
u
9.
z
o
H
CaO
O
i>
MgO
61
U.
o
<N
n
(W)
Study site
u
"E.
ca
00
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ON
.25
ON
in
ooNO
r-
rs
s
?
ro
?
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?
s
fN
NOOO
NO
00
(~-
30
00
r of Mono Basin
Southeast corne
fN
L-53-ML #UJ
NO
fN?
in
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s
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fN
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fN
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G
tN
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100
.81
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ON
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00
ro
00
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o
inin
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NO
r--fN
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00
r-
30<
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r of Mono Basin
Southeast come
L-54A-MLUJ
ro
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3
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in
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100
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ro
ro
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L-66-FVUJ
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in
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100
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in
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ro
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r?
t--
m
fN
NO
fN
in
ro
ON
ro
ley
e Mountain, Kobeh Va
East side of Lon
L-69-KVUJ
?
fN
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in
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n=16;?s
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ON
ON
ON
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fN
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fN
?
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?
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ro
NO
in
NO
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r-
,91
00
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ro
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Southwest Lone
..-72-KV #
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fN?
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ro
?
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?
ON
?
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NO
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1\
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100
.77
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fN
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in
?
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tN
00
NO
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r-NO
s
NO
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NO
00
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Mountain, Kobeh Vail
Southwest Lone
fN
..-72-KV #w
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fNfN
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?
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in
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99
.00
ON
ON
ON
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?
in
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s
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?
ON
rs
00
r-r-
ro
fN
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ro?
?
fN
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fN
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in
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8
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fN
ro
fN
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t-~
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, Lone Tree Mine
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fN
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in
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100
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ON
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in
ro
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s
?
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?
r-
NO
fN
rofN
ro
r-NO
c~-
NO
tN
NO
inro
ON
ro
lley
e Mountain, Kobeh Va
East sideof Lon
L-80-KV
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?
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fN
?
s
?
ro
??
s
?
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NO
ON
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s?
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NO?
ON
in
fN
ro
NO
fN
r-.
NO
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Southwest Lone
.-81-KV
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NO
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ro
?
s
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fN
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s
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s
?
tN
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tN
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100
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CS
in
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ro
fN
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in
?
ro
??
fN
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NONO
"?
ro
rs
ON00
NO
r-
fN
r-
<
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ON
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Core hole, Lone
..-92-LT
UJ
fN
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s
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fN
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ro
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vO
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100
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fN
ro
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in
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in
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rot--
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rofO
t?
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ro
gs Clayton Valley
Drill hole cuttin
Q
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o
-o
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en+1
in
2
s
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99
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ro
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s
?
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?
rs
fN
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++
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ro
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s
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in
roro
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ro
r-
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<
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S ? CB
a. a. a.E E E
64 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 3.?List of tephra layers analyzed in this study and their compositionally closest matches, based on electron-
microprobe analysis of volcanic glass shards. All samples were from lacustrine deposits unless otherwise noted; lo-
calities are given in Table 2. Analyses and methods are given in Table 2. Closest matches are based on calculations
of similarity coefficients using oxides of the elements Si, Al, Fe, Ca, Na, and K, except as noted below, and data not
used in the calculations are shaded in gray. Closely matched samples that did not provide age constraints or strati-
graphic context were omitted from the table.
Output
no. Field sample no. SiO2 A12O3 Fe2O3 MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer in roadcut south of Rhodes Salt Marsh (1). Closest matches are with the Teels Lake tephra sequence
(2-4, 11), the tephra layers in lake sediments of the Mono Basin (7, 9; 2.3 Ma; R. Fleck, in litt. to A.S.-W., 1997),
and the tephra layer at Willow Wash, Nevada (8; <2.57 Ma, >2.22 Ma; Sarna-Wojcicki et al., In review).
1
2
3
4
7
8
9
11
1
2
3
4
5
6
8
9
10
1
4
5
6
9
EL-2-RM
TLN-1
TLN-3
TLN-4
EL-54A-ML
FLV-36-WW
EL-54B-ML
TLN-2
77.54
77.28
77.20
77.26
77.38
77.47
77.53
77.06
12.72
12.76
12.87
12.83
12.76
12.83
12.64
12.96
0.55
0.56
0.54
0.58
0.55
0.58
0.53
0.58
0.03
0.04
0,04
0.04
0.04
0.05
0.03
0.05
0.06
0.08
0.09
0.08
0.07
0.09
0.07
0.09
0.44
0.44
0.43
0.45
0.40
0.43
0.41
0.42
0,06
0.08
0.09
0.08
0.08
0.11
0.07
0.09
3.94
4.09
4.03
3.95
3.83
3.66
3.82
4.22
4.65
4.69
4.70
4.74
4.91
4.76
4.91
4.54
99.99
100.02
99.99
100.01
100.02
99.98
100.01
100.01
Tephra layer at The Gap (1). Closest matches are with the tuff of Taylor Canyon tephra layers and related units (2-6,
8-10; Sarna-Wojcicki et al.,
EL-3-FL
FLV-150-CS
PC-C
WAC-6
CH88004
FLV-151-CS
TTC-6
TTC-19(2)
DG-4
76.66
77.26
76.76
77.20
76.67
77.34
77.20
77.54
76.96
In review).
12.88
12.71
13.14
12.74
13.20
12.70
12.93
12.57
13.03
0.62
0.61
0.60
0.60
0.64
0.60
0.58
0.59
0.56
0.04
0.03
0.03
0.04
0.05
0.04
0.05
0.05
0.04
0.08
0.09
0.08
0.06
0.09
0.09
0.09
0.07
0.09
0.33
0.34
0.32
0.34
0.34
0.34
0.33
0.33
0.33
0.07
0.05
0.07
0.09
0.08
0.05
0.09
0.06
0.08
4.13
3.96
3.90
3.98
3.80
4.09
4.08
3.91
4.00
5.19
4.96
5.10
4.94
5.13
4.74
4.64
4.89
4.92
100.00
100.01
100.00
99.99
100.00
99.99
99.99
100.01
100.01
Tephra layer at The Gap (1). Closest matches are with the tuff near Mount Jackson (5), the tuff (surface) in Clayton
Valley (4), and the tephra layers in Mud Spring Canyon north of Mono Basin (this study;
EL-4-FL
D77-27B
MJ-1A
EL-37-MS
EL-35B-MS
76.21
77.05
76.95
76.17
76.90
13.14
13.00
13.05
13.64
12.93
0.56
0.58
0.56
0.60
0.58
0.04
0.04
0.04
0.05
0.04
0.07
0.04
0.07
0.05
0.04
0.59
0.53
0.50
0.52
0.48
0.06
0.09
0.08
0.10
0.06
6,9).
3.50
3.45
3.26
3.47
3.34
5.83
5.22
5.49
5.40
5.61
100.00
100.00
100.00
100.00
99.98
1.0000
0.9884
0.9850
0.9821
0.9705
0.9703
0.9676
0.9647
1.0000
0.9747
0.9739
0.9727
0.9706
0.9699
0.9678
0.9675
0.9674
1.0000
0.9539
0.9507
0.9492
0.9451
Tephra layer in alluvial deposits at McGee Wash (1). The only reasonably close match is to the tephra layer at Weber
Reservoir, in Lovelock Formation(?) (2; J.O. Davis, unpublished data); possibly the same sample locality. Note: 8-
element comparison, excludes Mn.
1 EL-5-WD 71.98 14.97 2.31 0.25 0.04 1.49 0.20 5.29 3.47 100.00
2 DR-96 72.64 14.93 2.33 0.24 0.05 1.46 0.24 4.71 3.41 100.01
Tephra layer at Campbell Valley (1). Closest matches are with the Bishop ash bed in Owens Lake core OL-92 (3, 5,
9, 10), at Lake Tecopa, California (7), and in Adobe Valley (6); and with the ~1 Ma Glass Mountain tephra at Mc-
Gee Wash (8).
12.83 0.76 0.03 0.05 0.44 0.06
12.91 0.74 0.04 0.03 0.44 0.07
13.05 0.73 0.03 0.04 0.44 0.02
13.23 0.76 0.03 0.05 0.44 0.07
12.89 0.75 0.04 0.03 0.43 0.06
12.88 0.78 0.03 0.04 0.44 0.06
12.95 0.74 0.03 0.03 0.45 0.07
12.97 0.73 0.04 0.04 0.45 0.06
Tephra layer in alluvial deposits at McGee Wash (1). Closest matches are with the upper Miocene tephra layers
(3-5; 11.81 ?0.03 Ma correlated age; A.M. Sarna-Wojcicki, unpublished correlated age to date in Perkins et al.,
1995). Note: 7-element comparison, includes Ti.
1 EL-19-WD 75.88 12.23 2.02 0.09 0.03 0.70 0.25
3 DSDP173-21-4 (27-28 cm) 76.40 12.16 2.14 0,06 0.03 0.69 0.26
4 ALVES-1 76.33 12.12 2.08 0.05 0.03 0.65 0.23
5 DSDP173-21-4 (32-35 cm) 76.47 12.08 2.08 0.06 0.05 0.69 0.23
1
3
5
6
7
8
9
10
EL-15-CA
OL92-1030
OL-92-2 (97)
EL-38-AV
TECO-28B
EL-20-WD
OL92-1022
OL92-1028 (major)
76.55
76.94
76.88
76.69
77.00
76.78
76.95
76.89
1.0000
0.9532
4.14
3.90
4.12
3.81
3.80
4.31
3.99
3.97
5.14
4.93
4.70
4.93
5.00
4.68
4.79
4.85
100.00
100.00
100.01
100.01
100.00
100.00
100.00
100.00
1.0000
0.9773
0.9748
0.9746
0.9740
0.9731
0.9721
0.9709
3.36
3.12
3.24
2.99
5.42
5.14
5.28
5.36
99.98
100.00
100.01
100.01
1.0000
0.9651
0.9633
0.9622
NUMBER 33 65
TABLE 3.?Continued.
Output
no. Field sample no.
1
7
9
10
11
SiO2 A12O3 Fe2O3 MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer at McGee Wash (1). Closest matches are with the Bishop ash bed (11, 12, 14), with the younger tephra
layers of Glass Mountain (7; 1.2-0.8 Ma), and with the tephra layers near the Thorne Bar (9) and in Campbell Val-
ley (10).
76.78 12.88 0.78 0.03 0.04 0.44 0.06 4.31 4.68 100.00
77.29 12.83 0.79 0.03 0.04 0.45 0.06 4.03 4.48 100.00
76.52 13.08 0.73 0.03 0.03 0.42 0.06 4.44 4.69 100.00
76.55 12.83 0.76 0.03 0.05 0.44 0.06 4.14 5.14 100.00
77.20 12.83 0.76 0.04 0.08 0.42 0.06 4.20 4.42 100.01
76.52 12.90 0.80 0.04 0.05 0.44 0.06 4.10 5.11 100.02
77.40 12.81 0.74 0.03 0.04 0.44 0.06 4.20 4.27 99.99
EL-20-WD
PICO-73
EL-21-WL
EL-15-CA
OL92-1019
12 OL92-1027 (minor)
14 JC-239
Tephra layer at site north of Thorne Bar (1). Closest matches are with the Bishop ash bed (2, 6, 10, 14), with the
younger tephra layers of Glass Mountain (5; 1.2-0.8 Ma), and with the tephra layer at McGee Wash (3).
1 EL-21-WL 76.52 13.08 0.73 0,03 0.03 0.42 0,06 4.44 4.69 100.00
2 OL-92-2(97) 76.88 13.05 0.73 0.O3 0.04 0.44 0.02 4.12 4.70 100.01
3 EL-20-WD 76.78 12.88 0.78 0.03 0.04 0.44 0,06 4.31 4.68 100.00
5 PICO-78 (2) (100-200 cm) 76.33 12.93 0.73 0.03 0.04 0.41 0.06 4.24 4.24 99.01
6 OL92-1019 77.20 12.83 0.76 0.04 0.08 0.42 0.06 4.20 4.42 100.01
10 OL92-1029 76.92 13.06 0.71 0.03 0.04 0.42 0.05 3.92 4.85 100.00
14 OL92-1028 (minor) 76.73 13.16 0.71 0.04 0.01 0.45 0.07 4.08 4.75 100.00
Tephra layer east of Mina (1). No good specific matches. Generally similar to upper Miocene and lower Pliocene
tephra layers of the Snake River Plain eruptive source areas, Idaho (2), and with the tephra layers at Drumm Summit
(3,4).
1.0000
0.9745
0.9734
0.9731
0.9731
0.9729
0.9704
1.0000
0.9789
0.9734
0.9702
0.9702
0.9693
0.9672
1
2
3
4
1
3
5
9
10
11
12
13
14
1
3
11
1
2
3
4
6
1
6
8
12
13
EL-24-M
F8-90-G21 (39-41 cm)
(high Fe, low Na)
EL-66-FV
EL-59-FV (2)
75.58
76.32
76.28
76.52
12.56
12.37
12.22
11.99
1.77
1.67
1.86
1.88
0.06
0,07
0.04
0.04
0.03
0.04
0,02
0.03
0.59
0.59
0.60
0.58
0.17
0,17
0.18
0,17
3.17
2.71
2.94
2.73
6.09
6.06
5.87
6.06
100.02
100.00
100.01
100.00
Tephra layer in roadcut south of Rhodes Salt Marsh (overlies EL-2-RM) (1). Closest matches are with the tuff of the
Taylor Canyon tephra layers (3, 5, 9-14; 2.22-2.14 Ma; Sarna-Wojcicki et al., In review).
EL-26-RM
DG-4
FLV-SP-11
DG-3A
TTC-6
FLV-43-WW
P-169Q
PC-C
TTC-19 (2)
Tephra layer at
76.29
76.96
77.07
76.88
77.20
78.00
77.59
76.76
77.54
12.98
13.03
13.07
12.92
12.93
12.32
12.55
13.14
12.57
0.56
0.56
0.56
0.63
0.58
0.56
0.58
0.60
0.59
0.03
0.04
0.03
0.02
0.05
0.03
0.04
0.03
0,05
: Redlich Summit (1). Closest matches are
-1.95-1.79 Ma; Sarna-Wojcicki et al., In review).
EL-29-RM
CH88014
TTC-23
76.61
76.84
77.58
12.81
12.91
12.62
Tephra layer at Redlich Summit, dated 5
the tephra at Willow Wash, California (6;
EL-34-RM
FLV-83-HT
FLV-80-HT
EL-65-RM
FLV-105-WW
73.84
74.11
74.55
73.50
74.57
13.52
13.89
13.45
14.26
13.81
0.76
0.76
0.71
.9Ma(l).
0,05
0.02
0,03
o.n0.09
0.09
0.09
0.09
0,08
0.10
0.08
0.07
with the
0.08
0.07
0.03
0.32
0.33
0.33
0.31
0.33
0.32
0.33
0.32
0.33
0.06
0.08
0.07
0.05
0.09
0,06
0.06
0.07
0.06
older Glass Mountain
0.33
0.32
0.32
0,09
0.06
0.03
4.82
4.00
3.89
4.45
4.08
3.79
3.87
3.90
3.91
tephra
3.68
3.71
3.58
4.84
4.92
4.89
4.63
4.64
4.84
4.87
5.10
4.89
layers (3,
5.60
5.31
5.10
Closest matches are with the tuffs in Horsethief Canyon (2
100.01
100.01
100.00
99.98
99.99
100.00
99.99
100.00
100.01
11;
100.01
100.00
100.00
,3),
Reheis et al., 1991), and the overlying tephra layer at Redlich Summit (4).
1.24
1.19
1.17
1.11
1.21
0.09
0.10
0.07
0.09
0,09
Tephra layer in alluvial deposits at Mud Spring Canyon (V
Wash (6, 8, 12,
EL-35A-MS
FLV-34-WW
FLV-121-WW
FLV-33-WW
FLV-25-WW
13; <2.83, >2.57 Ma; Sarna-Wojcicki et al.,
77.37
76.85
76.61
76.87
77.60
12.59
12.93
12.95
12.95
12.45
0.63
0.62
0.65
0.64
0.61
0.05
0.08
0.04
0.07
0.07
0.10
0.09
0.10
0.10
0.10
0.39
0.37
0.37
0.38
0.38
0.15
0.19
0.19
0.19
0.20
3.87
3.84
3.58
3.71
4.35
6.80
6.21
6.52
6.66
5.28
100.00
99.99
100.00
100.00
99.99
). Closest matches are with the tephra layers at Willow
In review)
0.05
0.06
0.06
0.07
0.07
0.53
0.62
0.57
0.62
0.56
0.06
o.to
0.07
0.12
0.10
2.25
2.52
2.85
2.69
2.89
6.37
6.22
6.20
5.96
5.66
99.90
100.00
100.00
99.99
100.01
1.0000
0.9669
0.9621
0.9604
1.0000
0.9618
0.9583
0.9542
0.9541
0.9523
0.9470
0.9455
0.9454
1.0000
0.9832
0.9600
1.0000
0.9639
0.9602
0.9585
0.9309
1.0000
0.9459
0.9374
0.9295
0.9279
66 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 3.?Continued.
Output
no. Field sample no. SiO2 A12O3 Fe2O3 MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer in alluvial deposits at Mud Spring Canyon (1). Closest matches are with the tuffs near Mount Jackson
(2, 5; 2.9 ? 0.16 Ma, S. Weiss, in litt. to M.C.R., 1993), an overlying tephra layer (3), and the tuff (surface) in Clay-
ton Valley, Nevada (7).
EL-35B-MS 76.90 12.93 0.58 0.04 0.04 0.48 0.06 3.34 5.61 99.98
MJ-1A 76.95 13.05 0.56 0.04 0.07 0.50 0.0S 3.26 5.49 100.00
EL-36-MS 77.31 12.71 0.59 0.05 0.06 0.48 0.07 3.53 5.20 100.00
MJ-1B 76.90 13.00 0.59 0.04 0.08 0.49 0.07 3.62 5.20 99.99
D77-27B 77.05 13.00 0.58 0.04 0.04 0.53 0.09 3.45 5.22 100.00
Tephra layer in alluvial deposits at Mud Spring Canyon (1). Closest matches are with the
(2; see reference above), an underlying tephra layer (3), in southeastern Mono Basin (6);
Valley, Nevada (4), a similar tuff at Willow Wash (5; 11.3 Ma; Reheis and Sawyer, 1997),
at Willow Wash (10), -2.57 Ma (Sarna-Wojcicki et al., In review).
tuff near Mount Jackson
with the tuff in Clayton
and with the tephra layer
Tephra layer at Clayton Valley (1). Closest matches are with the ash flow at Mount Adams, Washington (2, 3; ~510
Ka, W. Hildreth, unpub. data), with the tuff above Huckleberry Ridge ash bed at Confidence Hills, Death Valley,
California (4, 5; -2.04 Ma, Sarna-Wojcicki et al., In review), and with the tuff in Burmeister core, Great Salt Lake,
Utah (6, 10; -3 Ma., Williams, 1993).
1
2
4
5
6
7
8
EL-42-CV #1 (major)
SSP-2
C-83-6A
ASW-613 84-27
ASW-61286-14
BUR-915
FLV-WW-52
72.20
72.25
72.50
72.98
73.51
72.02
72.48
5.65
5.68
5.39
5.40
5.53
5.66
6.10
3.22
3.30
3.36
3.17
3.24
3.63
3.19
100.01
99.99
100.01
100.01
100.00
100.01
100.00
1.0000
0.9912
0.9788
0.9765
0.9758
0.9743
0.9697
Tephra layer at Clayton Valley (1). Closest matches are with the La Salida tuff (2, 4, 5, 6), the tephra layer from the
Burmeister core in the Great Salt Lake, Utah (7; 3.06 Ma; Williams, 1993), and the tephra layer in Willow Wash,
California (8).
14.67 2.67 0.27 0.13 0.89 0.31
14.56 2.66 0.24 0.08 0.90 0.32
14.47 2.73 0.27 0.11 0.89 0.29
14.17 2.66 0.26 0.12 0.92 0.33
13.59 2.60 0.27 0.08 0.89 0.29
14.46 2.67 026 0.11 0.87 0.33
13.92 2.70 0,27 0.10 0.92 0,32
Tephra layer at Clayton Valley (1). Closest matches are with the Spearhead Member of Stonewall Flat tuff (2;
7.61 ? 0.03 Ma; and 3; 4.02 ? 0.08 Ma) as well as with similar tuffs in southern Fish Lake Valley (4, 5; Reheis and
Sawwyer, 1997), but Fe and K are slightly different. Also similar to the tephra in Willow Wash (6; <2.81, >2.57
Ma).
0.01 0.06 0.31 0.14
0.01 0.07 0.31 0,12
0.00 0.07 0.32 0.11
0.00 0.08 0.31 0.09
0.00 0.07 0.31 0.10
0.02 0.05 0.35 0.15
Tephra layer at The Gap (1); underlies EL-3-FL. Closest matches are with the tuff of Taylor Canyon, 2.22-2.14 Ma
(Sarna-Wojcicki et al., In review), from Emigrant Pass area, Nevada (3, 8), from Waucoba lake beds (11), and from
Cowan Pumice Mine, California (9, 13).
1
2
3
4
5
6
EL-43-CV
61990-1
B85-195-A
FLV-50-CM
FLV-5-CC
FLV-31-WW
76.61
76.42
76.07
76.86
76.56
77.47
12.17
12.05
12.35
11.75
11.89
12.02
1.97
1.85
1.86
1.87
1.91
2.01
1
3
8
9
11
13
EL-44-FL
FLV-183-EP
FLV-180-EP
TTC-20
W2A
TTC-21
1.0000
0.9784
0.9723
0.9678
0.9662
1
2
3
4
5
6
10
EL-36-MS
MJ-1B
EL-53-ML
D77-27B
FL-KA-3
EL-35B-MS
FLV-35-WW
77.31
76.90
77.27
77.05
77.54
76.90
77.58
12.71
13.00
12.69
13.00
12.27
12.93
12.69
0.59
0.59
0.57
0.58
0.62
0.58
0.60
0.05
0.04
0.04
0.04
0.07
0.04
0.06
0.06
0.08
0.06
0.04
0.04
0,04
0.07
0.48
0.49
0.45
0.53
0.47
0.48
0.49
0.07
0.07
0.08
0.09
0.09
0.06
0.10
3.53
3.62
3.63
3.45
3.47
3.34
3.76
5.20
5.20
5.22
5.22
5.43
5.61
4.65
100.00
99.99
100.01
100.00
100.00
99.98
100.00
1.0000
0.9879
0.9784
0.9728
0.9723
0.9723
0.9652
1
2
3
4
5
6
EL-42-CV #2 (high K)
MA-360 (A)
MA-360 (B)
CH88008
TL88001
BUR-915 (high Fe)
71.48
71.49
71.62
70.92
70.94
73.05
14.53
14.44
14.48
15.64
15.62
14.13
2.73
2.72
2.86
2.71
2.67
2.75
0,26
0.19
0.20
0.2S
0.27
0.25
0.09
0.07
0.08
0.09
0.09
0J0
0.92
0.87
0.91
0.96
0.96
0.86
0.36
0.33
032
0.29
0.30
0.32
5.06
5.12
4.80
5.06
5.01
4.52
4.57
4.77
4.73
4.06
4.15
4.01
100.00
100.00
100.00
100.01
100.01
99.99
1.0000
0.9803
0.9755
0.9601
0.9595
0.9415
4.24
4.24
4.24
3.95
3.91
3.42
4.49
4.94
4.98
5.10
5.25
4.51
100.00
100.01
100.00
100.01
100.00
100.00
1.0000
0.9726
0.9655
0.9539
0.9539
0.9408
76.63
77.20
77.14
77.70
77.35
77.68
13.04
12.70
12.89
12.58
12.48
12.57
0.58
0.56
0.59
0.58
0.60
0.60
0.03
0.05
0.04
0.05
0.03
0.05
0.12
O.U
0.11
0.06
0.08
0.07
0.32
0.31
0.32
0.31
0.34
0.32
0.07
0.07
0.10
0.05
0.08
0.06
3.43
3.23
2.99
3.38
3.30
3.33
5.77
5.78
5.82
5.29
5.74
5.32
99.99
100.01
100.00
100.00
100.00
100.00
1.0000
0.9735
0.9713
0.9703
0.9687
0.9683
NUMBER 33 67
TABLE 3.?Continued.
Output
no. Field sample no. SiO, Al9 Fe-, MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer at Redlich Summit (1). Closest matches are the Wono ash bed of Davis (1978) (27.2 Ka) (Benson et al.,
1997), from a core off the Oregon coast (6), and from Nixon, Nevada (9, 12, 17, 18).
1
6
9
12
17
18
EL-50-RM
F8-90-G21 (37-39 cm)
LD-30
LD-30
NIXON-1 (minor)
NIXON-2 (major)
74.03
73.98
74.08
74.10
74.60
74.54
13.93
14.26
14.03
14.03
13.77
13.81
2.05
2.03
2.11
2.11
2.06
2.10
0.32
0.27
0.30
0.30
0.31
0.30
0.05
0.05
0.06
0,06
0.05
0.06
1.43
1.41
1.38
1.38
1.33
1.34
0.28
0.31
0.32
0.32
0.32
0.30
4.30
4.53
4.51
4.53
4.46
4.46
3.61
3.17
3.21
3.17
3.11
3.10
100.00
100.01
100.00
100.00
100.01
100.01
1.0000
0.9633
0.9619
0.9593
0.9553
0.9534
Tephra layer in southeastern Mono Basin (1). Closest matches are with the tephra layers in Willow Wash section (3,
11; 5.9-5.4 Ma; and 5; <2.57, >2.22 Ma) and various tephra layers (7, 8, 12, 14) of indeterminate age, probably late
Miocene and early Pliocene age.
0.58 0,04 0.07 0.43 0.09
0.56 0.04 0.09 0.42 0.09
0.55 0.06 0.08 0.41 0.06
0.57 0.04 0.08 0.56 0.07
0.58 0.05 0.08 0.39 0.08
0.53 0.06 0.08 0.39 0.09
0.54 0.O5 0.07 0.43 0.08
0.58 0.05 0.06 0.40 0.11
Tephra layer in southeastern Mono Basin (1). Closest matches are with a tephra layer in northeastern Fish Lake Val-
ley (2; <2.57, >2.22 Ma by correlation), a tephra layer at Mud Spring Canyon (3), the tuff near Mount Jackson (5),
and the tephra layer in Yellowjacket Canyon, California (7; <2.57, >2.22 Ma, Sarna-Wojcicki et al., In review).
1
3
5
7
8
11
12
14
EL-53-ML (high K)
FLV-138-WW
FLV-29-WW
86LP39
OT-2
FLV-28-WW
D3SW-1
BE-62
76.86
76.77
76.85
76.57
78.22
77.46
77.93
77.36
12.60
12.96
12.65
12.90
12.23
12.55
11.76
12.59
2.47
2.65
2.20
2.47
2.33
2.17
2.80
1.81
6.85
6.42
7.13
6.75
6.05
6.67
6.33
7.04
99.99
100.00
99.99
100.01
100.01
100.00
99.99
100.00
1.0000
0.9638
0.9582
0.9515
0.9478
0.9436
0.9428
0.9381
1
2
3
5
7
EL-53-ML#l
FLV-63-CS
EL-36-MS
MJ-1B
YJC-4-87
77.27
76.87
77.31
76.90
77.17
12.69
12.94
12.71
13.00
12.92
0.57
0.58
0.59
0.59
0.56
0.04
0.04
0.05
0,04
0.04
0.06
0.09
0.06
0.08
0.10
0.45
0.43
0.48
0.49
0.42
0,08
0.06
0.07
0,07
0.07
3.63
3.77
3.53
3.62
3.45
5.22
5.23
5.20
5.20
5.28
100.01
100.01
100.00
99.99
100.01
1.0000
0.9791
0.9784
0.9749
0.9726
Tephra layer in southeastern Mono Basin (1). Closest matches are with the adjacent tephra layer (2; 2.3 Ma, R.
Fleck, in litt. to A.S.-W., 1997), with the tephra layer at Teels Lake, Nevada (5), and with the tephra layer south of
Rhodes Salt Marsh (7).
1 EL-54B-ML 77.53 12.64 0.53 0.03 0.07 0.41 0.07 3.82 4.91 100.01 1.0000
2 EL-54A-ML 77.38 12.76 0.55 0.04 0.07 0.40 0.08 3.83 4.91 100.02 0.9875
5 TLN-3 77.20 12.87 0.54 0.04 0.09 0.43 0.09 4.03 4.70 99.99 0.9697
7 EL-2-RM 77.54 12.72 0.55 0.03 0.06 0.44 0,06 3.94 4.65 99.99 0.9676
Tephra layer in southeastern Mono Basin (1), dated 2.3 Ma (R. Fleck, in litt. to A.S.-W., 1997). Closest matches are
with the adjacent tephra layer (2) and with the tephra layers south of Rhodes Salt Marsh (6) and at Teels Lake (9).
1 EL-54A-ML 77.38 12.76 0.55 0.04 0,07 0.40 0.08 3.83 4.91 100.02 1.0000
2 EL-54B-ML 77.53 12.64 0.53 0.03 0.07 0.41 0.07 3.82 4.91 100.01 0.9875
6 EL-2-RM 77.54 12.72 0.55 0.03 0.06 0.44 0.06 3.94 4.65 99.99 0.9705
9 TLN-3 77.20 12.87 0.54 0.04 0.09 0.43 0,09 4.03 4.70 99.99 0.9681
Tephra layer in alluvial deposits at McGee Wash (1). Closest matches are with the ash flows in the Silver Peak
Range (5, 8, 20) in the McAfee Creek area (tuff of Cottonwood Springs, 6.1 Ma; R. Fleck, in litt. to A.S.-W., 1997)
and in the Volcanic Hills (22), and with a tephra layer in the Cave Springs Wash area (14; 1.2-0.8 Ma, upper tuffs of
Glass Mountain), all in the Fish Lake Valley area of California and Nevada.
1
5
8
14
20
22
1
4
6
10
16
EL-57-WD
FLV-SP-1
FLV-SP-2
FLV-55-CS
FLV-SP-3
FLV-VH-5B
76.73
76.87
76.79
77.07
77.18
77.29
12.97
12.99
13.07
12.89
12.76
12.69
0.69
0.71
0.71
0.73
0.73
0.70
0.05
0.04
0.05
0.03
0.07
0.05
0.07
0.08
0.07
0.07
0.07
0.07
0.43
0.41
0.44
0.41
0.44
0.43
0.09
0-10
0.10
0.07
0.09
0.10
3.65
3.67
3.73
3.66
3.71
3.85
5.33
5.13
5.02
5.07
4.95
4.82
Tephra layer at McGee Wash (1). Closest matches are with the younger tephra layers of Glass Mountain (4, 6;
0.8 Ma) and similar tephra layers in Fish Lake Valley, Nevada (16), and with the Bishop ash bed (10; 0.76 Ma)
EL-58-WD(l)
BT-14
BT-IC
63CJ-26 (2)
FLV-162-CS
77.65
77.41
77.84
77.93
77.36
12.76
12.77
12.73
12.45
12.89
0.78
0.77
0.77
0.72
0.77
0.04
0.04
0.03
0.03
0.03
0.05
0.04
0.05
0.05
0.07
0.45
0.45
0.42
0.45
0.46
0.07
0.07
0.06
0.06
0.07
3.52
3.53
3.52
3.53
3.40
4.68
4.92
4.57
4.77
4.95
100.01
100.00
99.98
100.00
100.00
100.00
1.2-
100.00
100.00
99.99
99.99
100.00
1.0000
0.9798
0.9768
0.9728
0.9688
0.9682
1.0000
0.9886
0.9820
0.9789
0.9772
68 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 3.?Continued.
Output
no. Field sample no. SiO2 A12O3 Fe2O3 MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer at McGee Wash (1). Closest matches are with the younger tephra layers of Glass Mountain (3; 1.2-0.8
Ma) and with the Bishop ash bed (4, 6, 7, 11, 12; 0.76 Ma).
1
3
4
6
7
11
12
EL-58-WD (2)
BT-2
FLV-69-PA
FLV-4-WP
FLV-S2 (also FLV-148A-CS)
FRIANT 5A
BT-8 (2) (bulk)
77.99
77.64
78.04
77.93
77.81
77.73
77.75
12.31
12.67
12.31
12.49
12.55
12.41
12.48
0.74
0.74
0.74
0.74
0.72
0.75
0.73
0.03
0.03
0.02
0.05
0.03
0,05
0.03
0.03
0.03
0.06
0.04
0.03
0.04
0.0S
0.44
0.43
0.44
0.44
0.43
0.46
0.44
0.07
0.07
0.05
0.06
0.07
0.09
0.05
3.65
3.66
3.77
3.73
3.66
3.66
3.82
4.73
4.73
4.55
4.53
4.70
4.81
4.65
99.99
100.00
99.98
100.01
100.00
100.00
100.00
1.0000
0.9903
0.9882
0.9868
0.9866
0.9854
0.9847
Tephra layer at Drumm Summit in Fairview Valley (1). Closest matches are with the upper Miocene to lower
Pliocene tephra layers (2, 4, 6) and with the Lava Creek B ash bed, except for a high-iron mode found in the Bur-
meister core (5; Williams, 1993; correlated age, 0.65 Ma). Note: 7-element comparison, includes Ti.
1
2
4
5
6
EL-59-FV (2)
VW-CIMA (2)
EL-59-FV
BUR-281.9 (high Fe)
VW-CIMA-1
76.52
76.59
76.47
76.59
76.79
11.99
11.91
12.25
12.10
11.87
1.88
1.88
1.92
1.98
1.84
0.04
0.0(5
0.04
0.02
0.06
0.03
0.04
0.02
0.05
0.04
0.58
0.54
0.58
0.59
0.55
0.17
0.17
0.19
0.16
0.19
2.73
2.70
2.63
2.79
2.69
6.06
6.12
5.91
5.72
5.98
100.00
100.01
100.01
100.00
100.01
1.0000
0.9861
0.9701
0.9694
0.9686
Tephra layer at Drumm Summit, Fairview Valley (1). Replicate of EL-59-FV, above; same comments apply. Note:
7-element comparison, includes Ti.
1 EL-59-FV 76.47
2 VW-CIMA-1 76.79
5 EL-59-FV (2) 76.52
6 EL-66-FV 76.28
7 083089-A 76.95
8 08309-A (2) 77.05
Tephra layers in Truckee River canyon (1,2). Samples 1 and 2 are indistinguishable at analytical precision levels.
Similar to Ishi tuff (4, 10; ~2.5 Ma) and to the tephra layer at 1-80 at 320 mile marker (14; J.D. Davis, unpub. data);
age unknown. Note: 8-element comparison, excludes Mn.
12.25
11.87
11.99
12.22
12.15
12.20
.92
1.84
.88
.86
.84
.84
0.04
0.06
0.04
0.04
0.05
0.05
0.02
0.04
0.03
0.02
0.03
0.03
0.58
0.55
0.58
0.60
0.59
0.62
0.19
0.19
0.17
0.18
0.21
0.19
2.63
2.69
2.73
2.94
2.60
2.46
5.91
5.98
6.06
5.87
5.58
5.55
100.01
100.01
100.00
100.01
100.00
99.99
1.0000
0.9768
0.9701
0.9665
0.9664
0.9652
1
2
4
10
14
EL-60-TC
EL-61-TC
RS-2 (3)
RS-2 (2)
GS-85
77.35
77.34
77.97
77.51
77.68
13.29
13.34
12.48
12.82
12.63
0.83
0.81
0.85
0.81
0.83
0.15
0.14
0.14
0.16
0.15
0.08
0.07
0.05
0.02
0.04
0.69
0.69
0.78
0.74
0.69
0.13
0.11
0.13
0.15
0.16
2.47
2.49
2.81
3.39
3.71
5.02
5.01
4.78
4.40
4.11
100.01
100.00
99.99
100.00
100.00
1.0000
0.9677
0.9446
0.9100
0.9054
Tephra layer at Drumm Summit (1). Matches samples 2, 4, and 5 (this study); generally similar to upper Miocene
and lower Pliocene tephra from Snake River Plain sources, Idaho (8, 9), and with the tephra layer from gravels in
Las Vegas Wash in pre-Chemehuevi Formation (14). Note: 7-element comparison, includes Ti.
1
2
4
5
8
9
14
EL-66-FV
EL-59-FV (2)
EL-59-FV
EL-24-M
VW-CIMA-1
2-8-16J
083089-A (2)
76.28
76.52
76.47
75.58
76.79
76.32
77.05
12.22
11.99
12.25
12.56
11.87
12.28
12.20
.86
.88
.92
.77
.84
.82
.84
0.04
0.04
0.04
0.06
0.06
0.06
0.05
0.02
0.03
0.02
0.03
0.04
0.03
0.03
0.60
0.58
0.58
0.59
0.55
0.62
0.62
0.18
0.17
0.19
0.17
0.19
0.21
0.19
2.94
2.73
2.63
3.17
2.69
2.71
2.46
5.87
6.06
5.91
6.09
5.98
5.94
5.55
100.01
100.00
100.01
100.02
100.01
99.99
99.99
1.0000
0.9680
0.9665
0.9621
0.9592
0.9583
0.9536
Tephra layer at Lone Mountain, Kobeh Valley (1). Matches with the Rye Patch Dam ash bed at Tulelake, California
(3), with other samples of this ash bed at Kobeh Valley (4, 10), and with samples at the type locality of this ash bed
at Rye Patch Dam, Nevada (12, 14). Similar to Desert Springs tuff (ash-flow) in the Bend, Oregon, area (15), which
is probably a near-source site of the Rye Patch Dam ash bed. Note: 6-element comparison, excludes alkalies and
Mn.
0.37 0.06
0.37 0.08
0.36 0.07
0.36 0.09
0.34 0.06
0.39 0.09
0.39 0.09
1
3
4
10
12
14
15
EL-69-KV
TULELAKE-321
EL-70B-KV
AA-04-JOD
JOD-18-7-78B(l)
JOD-18-7-78A (1,2)
DST-3
71.76
70.85
71.81
71.37
71.25
71.34
71.31
14.86
14.80
14.79
14.34
14.82
14.37
14.58
3.32
3.35
3.33
3.35
3.22
3.37
3.19
1.41
1.46
1.39
1.44
1.39
1.45
1.48
0.43
0.43
0.44
0.42
0.43
0.44
0.44
4,35
5.16
4.34
5.01
5.00
5.08
4.40
3.44
3.51
3.46
3.60
3.49
3.47
4.12
100.00
100.01
99.99
99.98
100.00
100.00
100.00
1.0000
0.9900
0.9879
0.9799
0.9775
0.9741
0.9691
NUMBER 33 69
TABLE 3.?Continued.
Output
no. Field sample no. SiO, Al,02^3 Fe9 MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer at Lone Mountain, Kobeh Valley (1); low-Fe fraction. Similar to the tephra layers at Pringle Falls, Ore-
gon, about -180-220 Ka (Herrero-Berver et al., 1994) (2, 3, 6, 7, 8), to the Summer Lake bed NN, Oregon (4)
(Davis, 1985), and to the Columbia Canal ash-flow tuff near Bend, Oregon (10). Note: 8-element comparison, ex-
cludes Mn.
1
2
3
4
6
7
8
10
EL-72-KV#l (lowFe)
PF-88-0 (mafic)
PF-88-0 (silicic-ct)
DR-70
PF-88-P
PF-88-Q (silicic)
PF-88-0 (silicic-b)
BPT-CC
69.86
69.21
68.64
69.70
69.95
68.38
69.36
69.34
15.46
15.37
15.68
15.07
15.17
15.51
15.27
15.31
3.68
3.81
3.85
3.79
3.87
3.85
3.91
3.70
0.79
0.73
0.80
0.67
0.69
0.79
0.74
0.72
0.10
0.12
0.12
0.09
0.11
0.12
0.14
0.13
2.14
2.14
2.27
2.16
2.12
2.23
2.11
2.12
0.70
0.69
0.71
0.68
0.61
0.73
0.65
0.67
4.71
5.40
5.56
5.12
5.00
5.94
5.44
5.70
2.55
2.53
2.38
2.71
2.47
2.37
2.38
2.30
99.99
100.00
100.01
99.99
99.99
99.92
100.00
99.99
1.0000
0.9656
0.9526
0.9518
0.9471
0.9465
0.9465
0.9456
Tephra layer at Lone Mountain, Kobeh Valley (1); high-Fe fraction. Similar to the tephra layer from Honey Lake,
California, stratigraphically close above Huckleberry Ridge ash bed at Honey Lake, California (2); other similar
samples are the upper Miocene tephra from Balcolm Canyon, southern California (3), and the upper Pleistocene
tephra layer from the Tulelake area, California (4). Note: 8-element comparison, excludes Mn.
1 EL-72-KV #2 (high Fe) 61.09 16.78 6.84 2.28 0.14 5.04 1.28 5.02
2 CDGD-1 60.70 16.14 7.71 2.27 0.15 4.92 1.36 5.08
3 MOD-13 61.37 16.03 6.81 2.47 0.15 5.29 1.21 4.85
4 AS W-61085-14 (mafic) 63.69 14.75 7.13 1.84 0.10 4.86 1.26 4.87
Tephra layer in roadcut section, Lone Tree Mine (1). Most similar to the younger Glass Mountain tephra layers, in-
cluding Bailey ash bed, Ventura, California (4), and tephra layer beneath Bishop ash bed, northern Fish Lake Valley,
Nevada (5). Also similar to the Bishop ash bed at Friant, California (6-8).
1.53
1.68
1.82
1.50
100.00
100.01
100.00
100.00
1.0000
0.9568
0.9468
0.9380
1
4
5
6
7
8
EL-76A-LT
PICO-5 (2)
FLV-160-CS
FRIANT 4A
FRIANT-1
FRIANT-2
77.18
77.24
77.40
77.23
77.54
77.53
12.97
12.73
12.82
12.90
12.73
12.86
0.69
0.71
0.73
0.71
0.73
0.68
0,03
0.03
0.01
0.03
0.03
0.03
0.05
0.03
0.06
0,03
0.03
0.03
0.49
0.49
0.50
0.51
0.48
0.47
0.07
0.06
0.07
0.05
0.06
0.06
3.59
3.56
3.50
3.70
3.56
3.40
4.92
5.15
4.91
4.83
4.83
4.94
99.99
100.00
100.00
99.99
99.99
100.00
1.0000
0.9833
0.9806
0.9798
0.9792
0.9791
Tephra layer in roadcut section, Lone Tree Mine (1). Most similar to the younger Glass Mountain tephra layer at
Lake Tecopa, California (3), and to the Bishop ash bed at the southern end of volcanic tableland north of Bishop,
California (11).
1 EL-78-LT 77.29 12.90 0.71 0.03 0.05 0.48 0.05 3.82 4.66 99.99
3 TECO-26 77.61 12.73 0.72 0.03 0,03 0.48 0.06 3.70 4.63 99.99
11 BT-8(2) 77.48 12.68 0.71 0.03 0.02 0.45 0,06 3.85 4.72 100.00
Tephra layer at Lone Mountain, Kobeh Valley (1). Most similar to the Huckleberry Ridge ash bed near Ventura, Cal-
ifornia (2, 4), and in Meade County, Kansas (5, 6); somewhat less similar to the Lava Creek B ash bed from Butte
Valley, California (7), Lake Tecopa, California (8), and Onion Creek, Utah (9). Note: 7-element comparison, in-
cludes Ti.
0.02 0.04 0.60
0.03 0.03 0.60
0.Q2 0.05 0.60
0.02 0.03 0.58
0.02 0.04 0.58
0.03 0.03 0.55
0.02 0,05 0.53
0.02 0.03 0.54
Tephra layer on southwest side of Lone Mountain, Kobeh Valley (1). Closest matches are with the tephra layers of
Cowan Pumice Mine, Blind Springs Valley, California (4, 8), from northeastern Fish Lake Valley, Nevada (3, 6), and
from stratigraphically beneath the Huckleberry Ridge ash bed at Ventura, California (7).
1
2
4
5
6
7
8
9
EL-80-KV
PICO 158
ASW-3-23-75-1
68W98
68W98
#4078
RM880904
67W104
76.73
76.71
76.62
76.65
76.99
76.64
76.63
77.00
12.32
12.25
12.13
12.53
12.30
12.22
12.19
12.23
1.67
1.71
1.83
1.69
1.66
1.65
1.65
1.51
1.0000
0.9885
0.9829
0.10
0.10
0.10
0.10
0.10
0.10
0.10
0.10
3.54
3.62
3.63
3.32
3.22
3.51
3.55
3.58
4.98
4.95
5.02
5.08
5.09
5.27
5.28
4.99
100.00
100.00
100.00
100.00
100.00
100.00
100.00
100.00
1.0000
0.9918
0.9804
0.9793
0.9777
0.9760
0.9714
0.9686
1
3
4
6
7
8
EL-81-KV
FLV-13-CS
TTC-19 (2)
FLV-151-CS
P-169Q
TTC-9A
77.63
77.48
77.54
77.34
77.59
77.51
12.59
12.82
12.57
12.70
12.55
12.83
0.60
0.60
0.59
0.60
0.58
0.59
0.04
0.04
0.05
0.04
0.04
0.05
0.08
0.08
0.07
0.09
0.10
0.10
0.33
0.33
0.33
0.34
0.33
0.34
o.oa0.04
0.06
0.05
0,06
0.07
3.95
3.85
3.91
4.09
3.87
3.87
4.69
4.75
4.89
4.74
4.87
4.65
99.99
99.99
100.01
99.99
99.99
100.01
1.0000
0.9904
0.9883
0.9856
0.9843
0.9841
70 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 3.?Continued.
Output
no.
1
7
9
11
15
Field sample no. SiO, A12O3 Fe2O3 MgO MnO CaO TiO2 Na2O K2O Total (R)
Similarity
coefficient
Tephra layer in core at Lone Tree Mine (1). Closest matches are with the Lava Creek B ash bed near Oreana, Nevada
(2; J.O. Davis, Desert Research Institute, Nevada), the Huckleberry Ridge ash bed in Meade County, Kansas (3, 5,
7; Izett and Wilcox, 1983), and with the Lava Creek B ash bed, Lake Tecopa, California (4), and Ventura, California
(8).
0.02 0.03 0.56
0.03 0.03 0.56
0.02 0.04 0.58
0.03 0.04 0.56
0.02 0.03 0.58
0.02 0.05 0.58
0.02 0,00 0.56
Tephra layer near U.S. Highway 50 in Kobeh Valley, Nevada (1). Closest matches are with the Rye Patch Dam ash
bed at Rye Patch Dam, Nevada (4, 7), at a nearby site in Kobeh Valley (15), and in a core from Tulelake, California
(5); and with Desert Springs tuff, Bend, Oregon (13). Note: 8-element comparison, excludes Mn.
1
2
3
4
5
7
8
EL-92-LT
LM-24
68W98
T89-5/6/9 #1
68W98
68W98
PICO-133C
76.89
76.70
76.99
76.88
76.65
76.67
75.71
12.31
12.62
12.30
12.16
12.53
12.28
12.37
1.66
1.67
1.66
1.67
1.69
1.71
1.59
0.12
0.08
0.10
0.13
0.10
0.12
0.11
3.30
3.30
3.22
3.49
3.32
3.35
3.40
5.12
5.01
5.09
5.05
5.08
5.22
5.25
100.01
100.00
100.00
100.01
100.00
100.00
99.01
1.0000
0.9909
0.9889
0.9856
0.9855
0.9828
0.9806
1
4
5
7
13
15
AA-04 JOD
LD-61
TULELAKE-321
JOD-18-7-78A (1,2)
DST-1
EL-69-KV
71.37
71.07
70.85
71.34
72.69
71.76
14.34
14.70
14.80
14.37
13.39
14.86
3.35
3.32
3.35
3.37
3.21
3.32
0.36
0.37
0.37
0.39
0.38
0.37
0.09
0.09
0.08
0.09
0.07
0.06
1.44
1.45
.46
1.45
1.47
1.41
0.42
0.45
0.43
0.44
0.43
0.43
5.01
4.99
5.16
5.08
4.75
4.35
3.60
3.56
3.51
3.47
3.61
3.44
99.98
100.00
100.01
100.00
100.00
100.00
1.0000
0.9808
0.9804
0.9765
0.9654
0.9629
1
3
7
12
16
D77-2D
63CJ-26(1)
YJC-1-87
PICO-141
BT-8 (2)
77.39
77.37
77.28
77.29
77.48
3.72
3.60
3.74
3.68
3.85
4.89
4.99
4.75
4.91
4.72
99.99
100.00
100.01
99.99
100.00
1.0000
0.9908
0.9872
0.9839
0.9836
Tephra layer at depth of ~175 m in a core from Clayton Valley, Nevada (1). Best matches are with the reworked
Bishop ash at Friant, California (3), the airfall Bishop ash bed at Yellowjacket Canyon near Benton, California (7),
the Bishop ash bed in the Ventura basin, California (12), and the airfall Bishop ash bed at Insulating Aggregates
Quarry north of Bishop, California (16).
12.66 0.73 0.03 0.05 0.45 0.07
12.69 0.73 0.04 0.05 0.45 0.08
12.91 0.73 0.04 0.05 0.44 0.07
12.84 0.73 0.03 0.02 0.42 0.07
12.68 0.71 0.03 0,02 0.45 0.06
Tephra layer in large pit 9.5 m below surface at Columbus Marsh, Nevada (1). Best matches are with the airfall
Bishop ash bed near Westgard Pass, California (7), the Glass Mountain D tephra at Chalk Bluffs north of Bishop,
California, -0.9 Ma (9), the airfall Bishop ash bed at Friant, California (11), and the airfall Bishop ash bed at Insu-
lating Aggregates Quarry north of Bishop, California (15).
29068501 78.06 12.27 0.76 0.04 0,04 0.44 0.07
FLV-4-WP 77.93 12.49 0.74 0.05 0.04 0.44 0.06
BT-2 77.64 12.67 0.74 0.03 0.03 0.43 0.07
FRIANT 5A 77.73 12.41 0.75 0.05 0.04 0.46 0.09
BT-8 (2) bulk 77.75 12.48 0.73 0.03 0.05 0.44 0.05
3.64
3.73
3.66
3.66
3.82
4.67
4.53
4.73
4.81
4.65
99.99
100.01
100.00
100.00
100.00
1.0000
0.9834
0.9826
0.9822
0.9814
about 1330 m; two higher, older barriers, units Q16T and Q15T,
representing different lake cycles that reached altitudes be-
tween about 1330-1350 m and 1350-1370 m, respectively;
and at least one old bar, unit Q14T, that reached a minimum alti-
tude of about 1402 m.
An excellent small outcrop about 4 km north of Thome Bar
(Figures 2, 3) exposes four sequences of lacustrine gravel and
sand that overlie bedrock (Figures 3, 4B), all lying well above
the Sehoo shoreline. The four sequences are separated by pa-
leosols and alluvium. The second youngest lacustrine unit here,
Q13T, contains a tephra (EL-21-WL, Table 2) that is chemically
correlative with the Bishop-Glass Mountain family of tephra
erupted from Long Valley, California. Paleomagnetic measure-
ments on the silty sand layers containing the tephra are not de-
finitive in their polarity; thus, the tephra may be either the
Bishop ash bed dated from 0.76 Ma (normal polarity) or one of
the tuffs of Glass Mountain dated from 1.2 to 0.8 Ma (reversed
polarity; Saraa-Wojcicki et al., 1991). Small fragments offish
bones were found in the ashy layers (G.R. Smith, University of
Michigan, in litt. to M.C.R., 1998). The youngest unit at this
site, Q14T, consists of about 38 m of tufa-cemented, well-bed-
ded pebble and cobble shore gravel that is commonly steeply
bedded (backset beds) and that rises to an altitude of about
1393 m. Beach pebbles continue upward as lag on bedrock to
an altitude of about 1398 m. Because these outcrops and the
outcrops of the oldest morphologic unit at the Thorne Bar site
are similar in cementation and preservation and because they
rise to about the same altitude, we infer that they are equivalent.
The Bishop ash bed, identified by chemical analyses con-
ducted by Davis (1978) and by paleomagnetic data from the
NUMBER 33 71
vehicles
FIGURE 4.?Photographs of Thome Bar area, A, View south toward Thorne Bar (Figure 3); bedrock of the Gillis
Range is on left. Small white dots are vehicles parked on pre-late Pleistocene berm at about 1353 m, 20 m above the
Sehoo shoreline. B, View of outcrop 4 km north of Thorne Bar (Figure 3); people are for scale. East-dipping gravels
in upper part are tufa-cemented backset beds of unit Q14T. Line of coarse boulders at top of trench are shoreface
clasts at base of unit Q12T. Trench exposes paleosols, fan deposits, and lake unit Ql 1T just above basalt outcrop.
present study (data not shown), is present in beach gravel be-
neath Sehoo-aged deposits about 4 km south of Schurz, Ne-
vada (Figure 2). This outcrop indicates that a lake of uncertain
size existed in the Walker Lake subbasin at about 0.76 Ma.
We tentatively assign the tephra-bearing lake deposits of unit
Q13T north of Thorne Bar to this period, but we recognize that
they may be as old as 1 My. Four of the six paleosols formed
in units Q11T and Q12T and the intervening fan deposits (Fig-
ures 3, 4B) have moderately developed to well-developed
argillic horizons that are 20-90 cm thick. This degree of de-
velopment indicates lengthy (>50 thousand years (Ky)) peri-
ods of nondeposition (Reheis et al., 1995) and thus a rela-
tively long stratigraphic record prior to deposition of unit
Q13T.
72 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
The stratigraphy in the Thorne Bar area is shown schemati-
cally in Figure 5. The oldest and highest shoreline in this area,
as revealed by the outcrop north of Thorne Bar, is underlain by
deposits of four different lakes: two lakes (units Q11T and Q12T)
predated deposition of the tephra layer; one lake (QBT) existed
during deposition of the tephra layer; and one lake (Q14T) was
younger than the tephra layer. Units Q11T, Q12T, and Q13T
reached minimum altitudes of about 1355 m, whereas Q14T rose
to a minimum altitude of about 1400 m. In addition, barriers
with preserved morphology at the Thorne Bar site indicate that
three younger lakes (units Q15j, Q16j, and Q17j) successively
reached altitudes of about 1365 m, 1350 m, and 1330 m (the
Sehoo highstand).
20-Mile Beach: At least two localities preserve beach grav-
els along the Wassuk Range front (Figure 2). These localities
are at much higher altitudes than most other sites in the Walker
Lake subbasin. The first site is north of 20-Mile Beach and has
two levels of tufa-cemented beach gravel. The lower level is a
moderately preserved wave-cut bench overlain by 1-2 m of
tufa-cemented beach rock; it extends up to an altitude of about
1400 m (Figure 6A). The upper level is an eroded remnant of
beach rock with foreset beds overlying sheared granitic rocks;
the highest outcrop of beach rock is at 1452 m. The second site
is about 3 km to the south of the first. It has four exposed beds
of beach gravel interbedded with fan gravel above the Sehoo
shoreline; the highest bed caps a ridge at 1475 m (Figure 6B).
The super-elevated beach gravels at these two sites are inter-
preted to reflect absolute uplift of the footwall block of the Was-
suk fault. Demsey (1987) showed that this active normal fault
had about 7 m of displacement during the Holocene, of which
about 20%-25% was the result of uplift of the Wassuk Range.
McGee Wash: The study site at McGee Wash (Figure 2) ex-
tends from the Walker River just below Weber Reservoir east-
ward to U.S. Highway 95 (Figure 7). Exposures indicate that
sediments there are faulted and tilted (Morrison and Davis,
1984); only the larger faults are shown in Figure 7. Dips are
generally greatest on the eastern side near the highway and de-
crease toward the west. Younger units are progressively less de-
22
O 1
s Q-
Lone Tree Hill Thorne Bar area McGee Wash area Campbell Valley Sunshine Amphitheater
QI3L
(beach gravel
as high as -1415 m)
(inset beach gravel
as high as -1350 m)
"\_/-\_/-\_^
QI5T
(inset beach gravel
as high as -1365 m)
~\-/
QI4T
(beach gravel
as high as -1402 m)
(inset beach gravel
as high as -1365 m)
(beach gravel
as high as -1402 m)
Bishop
ash
N?XXXXXBishop QI3T
ash?
"5T R XXXXX
5 R Glass Mtn. Q|-|o U. --?-
 L
o R
ash
Huckleberry ?_.
Ridgeash(?) u"
Tg
XXXXX QTI
XXXXX Tt
QI3C
(inset beach gravel
as high as -1356 m,
possibly to 1378 m)
'\_/-\_/-\_^'
QI2C
(beach gravel
as high as -1378 m,
probably to 1390 m)
R
R XXXXX
R Glass Mtn. QI1C
ash
QI4S
(inset beach gravel
as high as -1378 m)
QI3S
(beach gravel
as high as -1414 m)
(beach gravel
QI2S crops out
to-1396 m)
cm
QTI
R (<2.6 Ma)
Lake-unit
names
Younger lake
deposits of Thorne
Older lake
deposits of Thorne
Lake deposits
of McGee
Lake deposits
of Lone Tree
Lake deposits
of Campbell
Lake deposits
of Sunshine
1111
/\x\
XXX
NorR
EXPLANATION
Buried soil
Unconformity
Tephra layer and
correlation
Paleomagnetic
polarity
Correlation
between units
FIGURE 5.?Schematic strati graphic correlations in the Walker Lake subbasin. Units and symbols are the same as in Fig-
ures 3, 6, 8, and 11. Deposits of Sehoo-aged lake not shown (Q17T in Thorne Bar area, grouped with fan deposits as unit
Qlfy in other site maps). Wavy lines represent unconformities; attached vertical lines are paleosols. N and R indicate
normal and reversed polarities, respectively, of sampled horizons, not shown to scale.
NUMBER 33 73
"Jnr
B
beach ^ ?/-- ?
beach
fa\J_.
??^C? "????"??
Tan besfch
Sehoo berm
FIGURE 6.?Photographs of very high beach gravels on the footwall block of the Wassuk fault (see Figure 2 and
Table 1). A, View south from gravel- and tufa-covered bench at 1399 m to sloping, eroded remnants of tufa-
cemented beach rock on skyline as high as 1452 m. Bush on right is about 1 m high. B, View west of beach gravel
interbedded with fan gravel above Sehoo shoreline at about 1330 m. Beach gravels crop out as high as 1475 m at
top of hill.
formed. The overall structure appears to be that of a block or
horst tilted progressively westward with time. This block was
periodically partly or entirely covered by lakes that prograded
eastward and by alluvial and eolian deposits that prograded
westward. The stratigraphic units at this site are shown sche-
matically in Figure 5.
The oldest units are exposed near the highway and consist of
a conformable sequence (Figures 7, 8) of alluvial deposits (unit
Tt) overlain by about 50 m of well-bedded lacustrine mudstone,
siltstone, and sandstone capped by beach gravel (unit QT1),
which in turn is overlain by a very thick (unmeasured, but more
than 50 m) sequence of alluvium and eolian deposits (unit QTt)
containing numerous carbonate-enriched paleosols. All three of
these units contain tephra layers, but only one, EL-19-WD in
alluvial unit Tt, has been identified; it is tentatively correlated
with an upper Miocene to lower Pliocene tephra erupted from
74
118?51r 118?50'
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
118?49' 118?48'
0
0 1 MILE
CONTOUR INTERVAL: 40 FEET
EXPLANATION
Qes Eolian sand (Holocene)
Qfy
Qlfy
Qce
Qfo
QI3?
Fan deposits (Holocene and late Pleistocene)
Lacustrine and fan deposits, undifferen-
tiated (Holocene and late Pleistocene)
Colluvial and eolian deposits
(Holocene and late Pleistocene)
Deposits of early late(?) Pleistocene lake
Fan deposits (middle Pleistocene)
Deposits of middle Pleistocene lake;
pattern shows beach gravel
Colluvial deposits (middle Pleistocene)
Deposits of middle? and early Pleistocene lakes;
pattern shows possibly correlative beach gravel
Qlo
QTt
EQTE
Tvs,"
58-WD
Lacustrine deposits, undifferentiated
(middle and early Pleistocene)
Alluvial, eolian, and colluvial deposits
(early Pleistocene and Pliocene)
Deposits of early Pleistocene? or Pliocene lake
Alluvial deposits (Pliocene and Miocene)
Volcanic and sedimentary rocks
(Pliocene and Miocene)
Fault; bar and ball on downthrown side; dashed
where inferred, dotted where concealed
Shoreline scarp
Tephra locality and sample number of sites
not within measured sections (Fig. 8)
Line of measured section (see Fig. 8)
Strike and dip of beds
FIGURE 7.?Geologic map of area around McGee Wash (see Figure 2, Table 1), east of Weber Reservoir.
NUMBER 33 75
0_X
LU
CD
O
CO
T3
2
CD T3
"5 WCD to
?e-o
CD C
?P co
COCO
T3i
j2 2k ?
CD
? %>
CO CO
CD "O
o V-
CD
sli
15115o o ?
82s Cn
~ -ao CD
O "DQ--D
- CDCD E
r^ CO -p
E
S
0 0 C
) O 0o o c
i 0 On n ( >?oOc
^
3 ?
Q
CO
co -g
is
CO Q_
9- co
.CD CO
LU
_0)
Q.
E
CO
CO
o
"CD ^"
O) CO1"
? CO
CO
Q_
CD
CDo
inI
COCD
E
CO
Q
CD
1?s
toCO
LU
C
o"T3
CD
C/D
Green, smectite-
rich, well bedded ,5 -5 -i
211818
?1
11
"D^
II
a?i
U
?r- c
| |
ij (S M
3 " a
.2 o cO .5
E 5 2
76
the Snake River Plain (Table 3). In the south-central part of the
map area (Figure 7), erosion between the time of deposition of
alluvial units Tt and QTt has removed lacustrine unit QT1. Unit
QTt apparently extends westward to underlie younger lacus-
trine units around McGee Wash. Three tephra layers crop out in
unit QTt in McGee Wash. Sample EL-57-WD may be of
Pliocene age (Table 3); two paleomagnetic samples (P3, Figure
8) indicate that EL-5-WD has normal polarity. Sample EL-6-
WD has not been correlated with tephra of known age.
Four pre-Sehoo lacustrine units separated by unconformities
overlie unit QTt in McGee Wash. The oldest, Q11M, is a thin
(1-2 m) bed of pebble gravel overlain by silty sand (Figure 8);
this unit is preserved in only a few small outcrops. Unit Q12M is
locally much thicker and crops out around the head of McGee
Wash (Figure 7) to a maximum altitude of about 1390 m (4560
ft). This unit consists of a basal bed of sand and pebble gravel
that grades up into well-bedded silt and sand containing at least
one tephra layer. Two samples (EL-20-WD, EL-58-WD; Tables
2, 3) have been identified as Bishop- or Glass Mountain-like
tephra layers. Paleomagnetic data show a sequence of stable re-
versed magnetization in sediments below the tephra bed, stable
normal magnetization in silt overlying the tephra bed, and sta-
ble reversed magnetization in overlying mudstone (Figures 8,
9). A couple of normal-magnetization samples near the tephra
bed lie in between samples of reversed magnetization. We have
not assessed magnetic mineralogy to determine whether these
normal directions represent some type of remagnetization dur-
ing a later normal-polarity interval. Nevertheless, the pattern of
reversed-normal-reversed magnetizations and ash chemistry in-
dicate that the ash is a Glass Mountain tephra that erupted at
about 1 Ma, just before the Jaramillo Normal Subchron (Sarna-
Wojcicki et al., 1991). If a beach-pebble facies of unit Q12M had
been deposited, it has mostly been removed by erosion, but
remnants of beach gravel that are probably older than that of
unit Q13M are locally preserved capping units QT1 and QTt near
the highway at an altitude of about 1430 m (4700 ft; Figure 7).
Unit Q13M unconformably overlies silt and sand of unit Q12M
around the head of McGee Wash, and it consists mainly of
loose sand and gravel (Figures 7, 8). A buried soil and collu-
vium is locally preserved at the unconformity. Unit Q13M is
poorly exposed, but in aerial photographs it has arcuate mor-
phology resembling that of a shoreline berm. The berm rises to
an altitude of 1402 m (4600 ft), the same altitude as the highest
beach gravel at Thorne Bar (Figure 3).
The youngest of the four pre-Sehoo lacustrine sequences,
unit Q14M, is inset into all older units (Figure 7). It consists
mostly of pebbly sand; in McGee Wash the pebbly sand over-
lies a gray clay layer that Morrison and Davis (1984) assigned
to the Eetza Alloformation. The isolated remnants of this unit
lie at altitudes of 1350-1365 m and may represent faulted de-
posits of one lake cycle or deposits of more than one lake cycle.
Sunshine Amphitheater: Sunshine Amphitheater is a very
large exposure of lacustrine deposits of five different ages
(Figures 2, 10; stratigraphic summary in Figure 5). It is re-
markable for the abundance of vertebrate fossils and the quality
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
A N (UP)
W (Hor)
S (down)
B N (Up)
0.1 A/mW(Hor)-}??i? E
q80
P45
S (down)
N(Up)
5mT
NRM
W (Hor)
20 NRM
H I hE
0.001 A/m
P42
S (down)
FIGURE 9.?Orthogonal demagnetization diagrams (method of Zijderveld,
1967) shown in stratigraphic order for specimens from unit Q12M, section 8b,
McGee Wash. The closed symbols represent projections of the magnetic vector
onto the vertical plane, whereas the open symbols represent projections of the
magnetic vector onto the horizontal plane. Remanent magnetizations are given
in amperes/meter (A/m). Demagnetization steps are in milliTesla (mT). A,
Reversed magnetization in sample P49; B, normal magnetization in sample
P45; c, reversed magnetization in sample P42, from 26 cm above tephra bed
EL-20-WD.
NUMBER 33 77
119"01'W 119?00'30" 119?00' 118?59'30"
Qaf
Qlfy
Qfu
EXPLANATION
Alluvium of the Walker River and fan
deposits, undifferentiated (Holocene)
Lacustrine and fan deposits, undifferentiated
(Holocene and late Pleistocene)
Fan deposits, undiff. (Holocene and Pleistocene)
|o ooj Beach gravel of late middle(?)
l? o s i Pleistocene lake
Qfo Old fan deposits (middle Pleistocene)
'Qfo+/
?rrv
1 Contour interval
1.0 miles 40 feet
Section 14
Qf3
Unit QI3S
12 m omitted
Section 15
^ P34-R
Pale marke
/// ^ P33-R?
Unit QI2S
/P19-R /
Common silt beds
5 m omitted ,
Fossil bones
P17-R?
/Pale; marker silt bed
Pale; locally
interbedded with sand
and pebbly sand
-P16-N
Abundant bones
~-3 Papery; fossil bones
Intermixed old fan deposits and beach gravel
of early middle(?) Pleistocene lake
Deposits of two early Pleistocene lakes sepa-
rated by unconformity (dotted line); pattern
shows transgressive basaltic beach gravel
Alluvial eolian, and colluvial deposits
(early Pleistocene)
Deposits of Pliocene or early Pleistocene lake
Volcanic and sedimentary rocks, mostly
basalt (Pliocene)
Late Pleistocene shoreline
Upper limit of beach gravel
Fault; bar and ball on downthrown side
Line of measured section
Strike and dip of beds
Fossil sample locality
P11-N
Abundant fossil
bones lagged
on surface
Unconformity;
slight dip change
P10-Rcomp.
UnitQI1s
Green
-P13-R
? Section
steps north
Basaltic gravel,
Section 12
Top of bench at -4520 feet
break in slope
Unit QI4S
P31-R?
Unit QTI
e? P30-R
i? P29-R
E? P8-?
Pale, massive
=? P7-R
W
B
? 0
Tan grading up to
brown paper shale
1?P6-R
Brown paper shale
1? P5-R
EXPLANATION
Wi Mudstone
~3 Interbedded mudstone
=3 and sandstone
=3 Siltstone and sandy
^ siltstone
Sandstone and silty sandstone,
locally conglomerate beds
Sandy pebble conglomerate,
well rounded clasts
(nearshore deposits)
Sandstone and conglomerate,
poorly to moderately sorted
Fanglomerate, poorly
sorted and bedded
Steep dips (foreset beds)
Buried soil
Unconformity
P9-R Paleomagnetic sample
and polarity (N, normal;
R, reversed)
Unit QTt
FIGURE 10.?Geologic map of Sunshine Amphitheater (Figure 2, Table 1) and selected measured sections. The
portion of section 14 above the closely spaced contours (top of unit Q12S) was not measured and is schematic.
Question mark (?) in place of N or R indicates polarity uncertain.
78 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
of their preservation; thus far, identified remains include those
of giant sloth, bird, horse, camel, mammoth, and fish. The
abundance of fossil remains at Sunshine Amphitheater may be
the result of its location adjacent to the Walker River near the
point where the river bends in a U-turn (Figure 2). Floods
coming down the north-flowing reach of the river would have
been forced to slow and drop part of their loads; the natural
deposition site would have been on the outside of the bend.
Small faults are exposed, but the units do not appear to be
greatly offset, in contrast to the McGee Wash and Campbell
Valley sites.
The oldest non-bedrock unit, QT1, consists mainly of brown,
finely laminated mudstone capped by pebbly sandstone; the
mudstone has reversed magnetization (Figures 10, 11D). It is
overlain by unit QTt, which has alluvial and eolian deposits
containing carbonate-enriched paleosols. Unit QTt is conform-
ably overlain by well-bedded pebbly sandstone, siltstone, and
mudstone of unit Qlls, which also has reversed magnetization
(Figure lie). A distinctive horse bone (Equus excelsus(!)
Leidy, 1858) from unit QUS appears to constrain the age to less
than 1.4 Ma. The jaw of a Lahontan cutthroat trout (Oncorhyn-
chus clarki (Richardson, 1836)) and a vertebra of a sucker
N(Up)
0.1 A/m--
W (Hor)
P34
S (down)
5 mT
NRM
B (Up)
W (Hor)
0.1 A/m--
S (down)
N (Up) so
W (Hor)
0.001 A/m 20
P10
0.02 A/m -
5 mT w (Hor)
NRM
S (down) S (down)
FIGURE 11.?Orthogonal demagnetization diagrams (method of Zijderveld, 1967) shown in stratigraphic order
for specimens from Sunshine Amphitheater sections. Symbols are the same as in Figure 9. A, Reversed magneti-
zation in sample P34 in unit Q12S, section 15; B, normal magnetization in sample PI 1, lower part of unit Q12S,
section 15; c, reversed magnetization component in sample P10 in unit Qlls, section 15 is indicated by the isola-
tion of southward, negative-inclination (up) vector after the removal of a northward, positive-inclination (down)
vector following demagnetization at and above 40 mT; D, reversed magnetization in sample P30 in unit QT1, sec-
tion 12.
NUMBER 33 79
{Catostomidae sp.; G.R. Smith, University of Michigan, in litt.
to M.C.R., 1996) also were collected from this unit. This age
range is consistent with the assignment of the reversely mag-
netized sediments to the Matuyama Reversed Chron (2.6-0.8
Ma; magnetic-polarity ages of Cande and Kent, 1995). An un-
conformity, locally marked by thin colluvium and a paleosol,
separates units Qlls and Q12S (Figures 10, 12A). Many of the
fossils are associated with this stratigraphic interval, including
mammoth bones (Mammuthus sp., also less than 1.4 Ma) and
the astragalus of a sloth identified as post-Blancan in age (<2.6
Ma; J. Honey, University of Colorado, pers. comm. to M.C.R.,
1995). Unit Q12S extends to an altitude of 1396 m and consists
of poorly indurated siltstone and sandstone that coarsen up-
ward into beach gravel. These deposits also contain fossilized
horse and bird bones. Paleomagnetic measurements indicate
that this unit has reversed magnetization except for an interval
near the base, which has normal magnetization (Figures 10,
11A,B). These data and the presence of a horse bone beneath
the unconformity indicate that the normal polarity interval
probably correlates with the Jaramillo Normal Subchron from
about 1 Ma. Thus, the three oldest lacustrine units at Sunshine
Amphitheater lie within the Matuyama Reversed Chron; unit
Qlls was deposited between 1.4 and 1.0 Ma, and unit Q12S was
deposited beginning at about 1 Ma (Figure 5).
Unit Q12S at the top of the bluff (section 14, Figures 10, 12A)
bears a well-developed argillic paleosol (Figure 12B) overlain
by fan gravel that forms a sloping bench. Exploratory soil pits
on this bench indicate that the fan gravel is capped by a paleo-
sol and is buried in turn by beach sand and gravel of unit Q13S,
another paleosol, and a younger fan gravel; unit Q13S rises to an
altitude of about 1414 m. Unit Q14S consists of sandy beach
gravel inset into all of the older units; it rises to an altitude of
about 1378 m.
Campbell Valley and Vicinity: The interpretation of out-
crops in Campbell Valley (Figures 2, 13) is complicated by a
network of northwest-striking strike-slip faults of the Walker
Lane. Displacement along these faults has juxtaposed lake sed-
iments of different ages, making it difficult to correlate depos-
its of beach gravel that mark shorelines with the finer-grained
sediment that contains age-related information. Three lacus-
trine units are locally exposed (Figures 13, 14). The oldest,
Qllc, consists mainly of mudstone that grades up into beach
sand and gravel; four samples from the mudstone have re-
versed magnetization. Deposits in another outcrop, herein as-
signed to the same unit as the mudstone but possibly belonging
to a younger lake cycle, consist of shoreface gravel that rapidly
grades laterally into fine-grained beds lapping onto subjacent
basalt. A tephra layer, sample EL-15-CA, drapes this facies
change from shoreface into deep-water deposits (Table 2; Fig-
ure 14A). The tephra, a sample from which has normal magne-
tization, is correlated with the Bishop?Glass Mountain family
of tephra (Table 3). The reversed magnetization of sediment
just above and below the tephra layer, however, indicates that
the tephra cannot be Bishop ash on the basis of enveloping re-
versed magnetizations; thus, the tephra is probably a Glass
Mountain tephra from about 1 Ma. Unit Qllc is unconformably
overlain by beach gravel of unit Q12C; a paleosol and colluvium
are locally preserved at the contact. Unit Q12C rises to at least
1378 m, and beach pebbles that probably represent this unit are
present as lag on basalt up to an altitude of 1390 m. The peb-
bles are found up to an abrupt change in slope that resembles a
preserved shoreline angle cut on a small island, or stack (Figure
14B). Beach gravel of unit Q13C is apparently inset into the
older units and forms a berm at an altitude of 1356 m.
Another site in Campbell Valley, between Sunshine Amphi-
theater and Weber Reservoir on the east side of the Walker
River (Figure 2), records shorelines cut into two basalt hills.
Subdued breaks in slope occur on the west sides of these hills,
with the most prominent slope break occurring at an altitude of
1365 m. From a distance these slope breaks resemble flow con-
tacts, but they are actually degraded shoreline angles. Rounded
beach pebbles locally lie below the most prominent slope
break, and they are identical to pebbles that abundantly mark
the Sehoo shoreline-cut on the same hills at 1330 m. Additional
small remnants of slope breaks may also record shoreline an-
gles at altitudes of 1353 m and 1396 m. A high shoreline is fur-
ther supported by an outcrop on the east side of one of the hills.
There, rounded clasts of rhyolite and basalt overlie a rhyolite
flow weathering into angular clasts; the rounded clasts, inter-
preted as beach gravel, crop out as high as 1390 m.
EVIDENCE FROM THE NORTHEASTERN LAHONTAN BASIN
(HUMBOLDT RIVER).?During the Sehoo highstand, the north-
ern subbasins of Lake Lahontan were filled to an altitude of
-1330 m. According to Davis (1982), Lake Lahontan could
have been connected to the Walker Lake subbasin only by the
narrow Adrian Valley (Figures 1, 2). King (1993), however,
thought that Adrian Valley did not exist during the early-mid-
dle(?) Pleistocene. If his theory is correct, high lakes in the
Walker Lake subbasin would have been isolated from Lake
Lahontan during this time. Yet in the foregoing discussion, we
have documented a lake in the Walker Lake subbasin that rose
to an altitude of about 1400 m. This lake could have spilled
over a pair of sills at 1370 m into the northern subbasins of
Lake Lahontan through Rawhide Flats (Figure 2), if the sills
leading into Rawhide Flats were not much higher then than
they are now. If such a connection existed, evidence of high
shorelines north of the Desert Mountains would be required to
establish that the very high shorelines found in the Walker
Lake subbasin were coextensive with Lake Lahontan as a
whole. No such evidence had been previously reported. Rob-
erts (1965), however, suggested that the region around Lone
Tree Hill contained evidence of a lake, but because the sites he
described were above the Sehoo highstand, he inferred that this
lake was separate from Lake Lahontan and must have been im-
pounded by a dam in the Humboldt River canyon east of
Winnemucca. Most of Roberts' sites, except that at Treaty Hill,
consist of fluvial terrace deposits of the Humboldt River.
80 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Figure Section
12B 14 Top of section 15
FIGURE 12.?Photographs of central part of Sunshine Amphitheater. A, View north showing area of sections 14
and 15 (see Figure 10); fence posts show scale. B, Paleosols and fan deposits separating units Q12S and Q13S at
the top of section 14, exposed at top of bluff in A. Trowel indicates top of unit Q12S.
NUMBER 33 81
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82 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
B
FIGURE 14.?Photographs of Campbell Valley area; person is for scale, A, View west of outcrop of tephra EL-15-
CA, showing drape of tephra from shoreface position atop basalt into deeper-water sediments. See Figure 13 for
location. B, View north of shoreline angle cut on outcrop marking shoreline (-1390 m) reached by lake during
deposition of unit Q12C. Hill is just behind highest point visible in A.
Work began in 1997 at a new site on the Humboldt River be-
tween Winnemucca and Battle Mountain (Figure 1), within the
area studied by Roberts (1965). This site has both spectacular
exposures and geomorphic evidence of pre-late Pleistocene
lakes. The outcrops are within large open pits of the Lone Tree
gold mine, currently operated by Newmont Gold Corporation;
as a result of mining operations, the exposures are excellent
but change constantly. Our work has focused on a less fre-
quently remodeled roadcut within the mine on the south flank
of Lone Tree Hill (Table 1) and on samples from a core drilled
southeast of the hill (Figure 15). The data from these sites are
supported by observations and tephra samples from other,
now-destroyed pit walls.
Premining topographic maps show a prominent berm as high
as 1410 m extending south from Lone Tree Hill. The roadcut
sequence (Figures 15, 16) is within this berm. Well-rounded
cobbles are lagged on slopes above the berm as high as about
1415 m. Outcrops of lacustrine deposits form the upper 5-50 m
of the walls of all of the open pits south and west of the hill,
providing a three-dimensional view of kilometers of outcrop.
The lacustrine deposits and older units are cut by three normal
faults that bound the hill on the east, south, and west; thus,
Lone Tree Hill is a horst. From outcrops and core data, the base
of the lake deposits appears to be offset at least 40 m by the
eastern fault, 8 m by the southern fault, and 50 m by the west-
ern fault. The amount of absolute uplift of the horst is un-
known, but it could be considerable in such an active tectonic
setting. A younger, lower berm at about 1370 m in altitude was
present on the northwest side of the hill, but the berm is now
obscured by construction.
Outcrops in Lone Tree Mine and core samples are tentatively
interpreted to represent four lake units separated by unconfor-
mities and paleosols. In the roadcut exposure, unit Q11L is re-
versely magnetized where the sediments are fine-grained; these
sediments interfinger with coarser sand and gravel with steep
foreset beds (Figures 15, 16) containing a tephra layer (sample
NUMBER 33 83
Roadcut sequence, northeast mine area
(highest elevation 1410 m)
QI3L
.o.o. o o
Mn-cemented beds
-Diatomite;S11-LT
. EL78-LT; Bishop (?)
ash, 0.76 Ma
QI2L
S13-LT (projected; approx.
stratigraphic location)
S9-LT
EL76-LT; Glass
Mtn. ash (?), ~1 Ma
Core GT-2
(collar altitude 1366 m) EXPLANATION
Beach gravel?
Greenish-white;
floating pebbles
^ GT2-28.2-LT
<? GT2-54.5-LT
Poor recovery;
mixed with
overlying
mudstone
QI1L
5 ?
Thin, pebbly
transgressive-
sand
- 0
White;
diatomaceous
P89--N?
S8-LT
S7-LT; P88--?
Dropstones
S6-LT
Oolitic marl
'??? P86--?
P84--R cornp.
QTI
QTf
?XX-X-X-
Little recovery
Poor recovery
Greenish-white;
spring deposits?
^ EL94-LT-80.6
i? GT2-89.8-LT
Floating pebbles;
dropstones?
Beach gravel
o o co o
o o co o
Slickensided clay
F?GT2-100.8-LT
??GT2-105-LT
, GT2-109-LT
Mudcracks
5?GT2-114.8-LT
EL92-LT-117.5;
5 Huckleberry Ridge
ash bed?, 2.1 Ma
Floating pebbles;
dropstones?
'?GT2-125.5-LT
Lost core
Mudstone
Interbedded mudstone
and sandstone
Siltstone and sandy
siltstone
Sandstone and silty
sandstone, locally
conglomerate beds
Sandy pebble
conglomerate, well
rounded clasts
(nearshore deposits)
Sandy cobble
conglomerate, well
rounded clasts
(beach deposits)
Sandstone and con-
glomerate, poorly to
moderately sorted
Fanglomerate, poorly
sorted and bedded
Steep dips
(foreset beds)
Calcareous cement,
locally tufa; in fine-
grained beds, indicates
marly strata
Buried soil
x x x | Tephra layer and
EL76-LT samP|e numDer
P86--R
S11-LT,
GT2-28.2-LT
Paleomagnetic sample
and polarity
Ostracode sample;
number after GT-2
is core depth in feet
-GT2-131.5-LT
^-GT2-142.5-LT
Greenish-white;
mudcracks
? GT2-151.7-LT
Total depth 152 m
-10
-20
-30
?c5
FIGURE 15.?Measured section and core description of lacustrine deposits in Lone Tree mine on the flanks of
Lone Tree Hill (see Figure 1, Table 1 for location). Question mark (?) in place of N or R indicates polarity uncer-
tain. Correlation lines between measured section and core are only tentative. Note scale difference between road-
cut sequence and core GT-2.
84 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
/?>?
unconformity
FIGURE 16.?Photographs of central portion of roadcut sequence in Lone Tree Mine. A, East-dipping foreset
gravel beds and diatomite in unit Q12L overlie a horizontal unconformity and paleosol that formed on west-dip-
ping foreset beds and tephra EL-76A-LT (Table 3) in unit Q11L. B, Close-up of fossiliferous sand layer above
tephra (pencil for scale); sand contains abundant fish bones {Catostomus sp.; G.R. Smith, University of Michi-
gan, in litt. to M.C.R., 1998) and ostracode shells.
EL-76A-LT, Tables 2, 3) with Bishop-Glass Mountain-type
chemistry. A sandy layer above the tephra layer (sample S9-LT,
Figures 15, 16B) contains abundant ostracodes and small bones
of Catostomus sp. (sucker; G.R. Smith, University of Michi-
gan, in litt. to M.C.R., 1998). Unit Q12L in the roadcut consists
entirely of sandy gravel, which is commonly steeply bedded,
and aim thick diatomite bed (Figure 16A); the unit overlies an
unconformity and a paleosol. In two places, gravel clasts of
unit Q12L are embedded in a matrix of tephra (sample EL-78-
LT), which also has Bishop-Glass Mountain-type chemistry.
These characteristics indicate that units Q11L and Q12L are from
about 1 Ma and 0.76 Ma, respectively. It is possible that the
NUMBER 33 85
lower tephra could be the Bishop ash because the sediment en-
closing the tephra was too coarse-grained for paleomagnetic
determination, but this is unlikely because this coarse sediment
interfingers laterally with fine sediment that is reversely mag-
netized, as noted above. Also, the presence of a paleosol sepa-
rating the units indicates that the upper tephra was probably not
reworked from the lower tephra. The youngest and highest lake
unit in the roadcut is undated.
Correlation of the roadcut exposures to the core samples is
problematic because of differences in grain size, fault displace-
ment between the core and the roadcut, and alteration in the
core sediments that prevented analysis of most of the tephra
samples. However, the core should contain a more complete re-
cord of lacustrine deposition because of its location on a sub-
siding fault-block. A tephra layer (sample EL-92-LT) from the
lowest lake unit in the core contains shards that are chemically
similar to the Lava Creek B ash bed and to the Huckleberry
Ridge ash bed, both of which erupted from the Yellowstone
area (Izett, 1981) 0.66 Ma and 2.1 Ma, respectively (Pluhar et
al., 1992; Sarna-Wojcicki et al., In review). Ostracodes that be-
came extinct by about 0.6 Ma (Elkocythereis bramletti (Dickin-
son and Swain, 1967); R. Forester, USGS, in litt. to M.C.R.,
1998) are present throughout the core, indicating the older age
is likely correct. The older lake unit (from 2.1 Ma) in the core
is not present in the roadcut (Figure 15).
Environmental conditions interpreted from ostracodes
throughout most of the core and from unit Q11L in the roadcut
indicate this area had shallow to moderate water levels, rela-
tively saline alkaline water, and cool-season runoff as well as
warm summers (R. Forester, USGS, in litt. to M.C.R., 1998).
An unusual suite of ostracodes in sample S-13-LT from unit
Q12L (Figure 15) includes several highly ornate, endemic spe-
cies that are thought to have evolved in a long-lived, stable lake
(discussed in Forester, 1991) and a species that can live in a rel-
atively deep-water lake (Candona aff. C. caudata Kaufmann,
1900). We speculate that the endemic species, which clearly
did not evolve locally because they are not present in any other
samples, could have arrived during a sudden influx of water
from a previously undrained lake basin upstream. The Can-
dona shells suggest deeper water prevailed during the time of
deposition of unit Q12L than during that of Q11L. A few samples
from the core and the roadcut exposures contained ostracodes
and diatoms (J. Bradbury, USGS, in litt. to M.C.R., 1998) that
live in springs. This is consistent with sediments in the core
and outcrops that are interpreted as spring deposits and with the
presence of modern springs at Lone Tree Hill.
Other localities near Lone Tree Hill provide additional evi-
dence of pre-late Pleistocene highstands. On the north side of
the Humboldt River, a saddle between the two summits of
Treaty Hill (Figure 1, Table 1) has the form of a beach berm
and consists of very well-rounded pebble to cobble gravel of
diverse lithologies. The saddle is at an altitude of 1402 m; peb-
ble lag extends up the south hillslope to as high as 1414 m.
This gravel is unlikely to be a terrace deposit of the Humboldt
River because the shape and orientation of the berm seem to
preclude deposition by the river, and because the river does not
flow in an antecedent course across a rising block at this loca-
tion. However, some tectonic uplift of Treaty Hill, like that in-
ferred at Lone Tree Hill, is certainly possible. A lower ero-
sional bench on the east side of Treaty Hill also may represent a
shoreline at about 1360 m, as noted by Roberts (1965). West of
Lone Tree Hill, at Iron Point on the Humboldt River, scattered,
very well-rounded pebbles lie on gently sloping basalt surfaces
at altitudes from 1350 m to 1370 m (Table 1).
Northeast of Treaty Hill is a broad, very flat surface that
slopes gently westward for several kilometers from an altitude
of 1362 m to 1344 m (Table 1) and is underlain by well-
rounded gravel and sand. This surface is interrupted by a sub-
dued break in slope at 1356 m. A drainage ditch exposes a
well-developed argillic soil that formed on these gravel depos-
its and that is buried beneath nearly 2 m of loess. Although ex-
posures are few and poor, the evidence suggests that this sur-
face and the underlying deposits may represent pre-late
Pleistocene deltaic deposits of the Humboldt River graded to a
lake level that reached as high as about 1356 m.
EVIDENCE FROM OTHER SUBBASINS OF LAKE LAHONTAN.?
Soda Springs Valley (Figures 1, 2) extends southeast of Walker
Lake. Modern elevations of the valley floor are below 1400 m;
small basins in the valley are as low as 1330-1352 m, and sills
are between 1377 m and 1396 m. It is likely that valley-floor
altitudes are higher at present than in the past because of allu-
vial-fan sedimentation. These sediments accumulate because
rates of vertical displacement along the active strike-slip faults
in this part of the Walker Lane belt are low (Yount et al., 1993).
Thus, the 1400 m high lake-level recorded at several sites in
the Walker Lake subbasin indicates that the pre-late Pleistocene
lake should have flooded Soda Springs Valley. Lacustrine de-
posits of Miocene to late Pliocene age are preserved along the
range fronts on both sides of Soda Springs Valley and south of
Rhodes Salt Marsh (discussed below); however, younger lacus-
trine deposits have not been found.
Several sites along the Truckee and Carson Rivers in can-
yons below Reno and Carson City, respectively (Figure 1, Ta-
ble 1), consist mainly of rounded, exotic clasts overlying bed-
rock. The clasts may represent times when rivers were graded
to lake levels far above the Sehoo shoreline, or they may repre-
sent terraces stranded by uplift of the mountain blocks through
which the rivers flow. Arcuate geomorphic features with
rounded cobbles occur in the Truckee River canyon at an alti-
tude of about 1400 m, and one roadcut exposes sediment that
may be lacustrine or deltaic at about 1355 m. At one locality on
the Carson River, the rounded clasts are found on slopes pro-
tected from river erosion and above a bedrock sill where the
Sehoo-aged lake spilled northward out of the canyon. The
clasts at this site are more likely to reflect high lake-stands
rather than fluvial gravels.
Good evidence of high, old shorelines exists at three sites in
the Smoke Creek Desert (Figure 1, Table 1). These old shore-
86 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
lines were first observed by Mifflin while conducting field
work during the 1970s. The southern site is near Burro Moun-
tain in the Smoke Creek Desert and lies 1-2 km west of an ac-
tive Quaternary fault that bounds the western edge of the basin
(Weick, 1990). Two subtle benches and risers are underlain by
basalt and rhyolite flows (Figure 17A). The upper bench termi-
nates upward at a slope break at an altitude of about 1400 m,
and well-rounded small pebbles, many flattened and discoidal
in shape, are scattered over the bench at altitudes between 1395
and 1400 m. The lower bench, further east and closer to the
fault, is at about 1375 m and is blanketed with abundant well-
rounded pebbles and a few fragments of lacustrine tufa (Figure
17B). A few kilometers to the north, another site also has abun-
dant well-rounded pebbles of similar shape at about 1390 m.
Alluvial interbeds that could have weathered to provide
rounded gravel were not observed within the volcanic flows at
these sites. It is possible that the well-rounded pebbles at these
sites could represent deposition by an ancient stream prior to
uplift along the Bonham Ranch fault zone (Weick, 1990); three
other nearby sites that were field-checked by M.C.R. were so
interpreted on the basis of their lithologic composition and
broad range of clast sizes, which were similar to those of the
upper bench
K^.v^EJ-.-^
FIGURE 17.?Beach-gravel sites in the Smoke Creek Desert (see Figure 1, Table 1). A, East flank of Burro Moun-
tain. View is to the west from the lower shoreline bench at 1375 m to the upper bench at 1400 m. B Typical scat-
ter of beach pebbles at Burro Mountain site; remnants of lacustrine tufa also present on lower bench. Lens cap is
for scale. C, View northwest to Freds Field site. Subtle benches are cut on basalt flows above the Sehoo shoreline
on photo right. D, Well-rounded and polished basalt clasts make up the majority of block streams on the shoreline
benches of the Freds Field site. Rock hammer is for scale.
NUMBER 33 87
local modern streams, namely, Smoke Creek and Buffalo
Creek.
The eastern site in the Smoke Creek Desert (Figure 1; Table
1, Freds Field) consists of two benches and risers underlain by
basalt flows (Figure 17c), forming subdued topography above
a steep cliff on which the Sehoo shorelines are cut. The
benches are laterally continuous and are covered with basalt
clasts closely fitted together in pavements interspersed with
rougher surfaces. The risers are also covered with basalt clasts
that form matrix-poor stone stripes. More than 50 percent of
the clasts on the benches and risers are polished and rounded
(Figure 1 7D) rather than angular or subangular, as is typical of
clasts in colluvial block fields. This polishing and rounding re-
quires vigorous transport in water. The upper bench at this site
rises gradually from an altitude of about 1370 m to a subtle
slope break at about 1400 m; the lower bench rises from a flat
surface at 1360 m to a distinct slope break at about 1365 m.
The Freds Field site appears to be unaffected by faulting, and it
is located well away from possible locations of ancestral
streams that might have carried rounded gravel. The Sehoo
shoreline at Freds Field bears clasts of more varied lithologies
than do the upper benches. This difference possibly results
from the topographic setting; east-flowing longshore currents
in the inferred higher lakes might have been channeled along
the cliff bearing the Sehoo shoreline but below the higher
benches (Figure 17c).
Summary: Very High Shorelines of Lake Lahontan
Several outcrops in the Walker Lake subbasin, one locality in
the northeastern Lahontan Basin, and two localities in the
Smoke Creek Desert subbasin document highstands of a lake
during the early and middle Pleistocene that rose far above the
Sehoo and Eetza(?) shorelines. A conservative correlation
among all the sites in the Walker Lake subbasin yields a total of
six or seven lacustrine units whose shorelines exceeded that of
the Sehoo lake. Three of the units date from about 1 Ma to less
than 2.6 Ma; one tentatively contains the Bishop ash bed; and
three are probably younger than the Bishop ash bed, but they
definitely date from less than 1 Ma. At Lone Tree Hill, the old-
est lake unit is from about 2 Ma, two middle units are probably
from about 1 Ma and 0.76 Ma, and the highest unit was depos-
ited less than 0.76 Ma. Other lake units may be present in the
core. In the area north of Lone Tree Hill, erosional features and
possible deltaic deposits are interpreted to indicate an even
younger and lower highstand.
Based upon these correlations, we assign informal names to
the deposits that comprise a lake unit?that is, the fine- and
coarse-grained lacustrine sediments deposited during the rise
and fall of lakes during a pluvial climatic phase. Naming the
lake units permits easy revision should a new pluvial phase be
identified, and it allows changes in the correlation of lake units
at a site if new age information is obtained. This is particularly
likely where correlations are made chiefly on the basis of alti-
tude. Variations in altitudes of such old landforms and deposits,
especially as they represent minimum altitudes of lacustrine
highstands, are expectable in the tectonically dynamic land-
scape of the Basin and Range. Without independent dating of
beach gravels at different sites, we cannot be confident that
they represent the same highstand.
Stratigraphic relations among four sites (Figure 5) permit a
conservative correlation of the two youngest pre-Sehoo units;
these units always crop out as beach gravel inset into all older
units. The maximum altitude of the youngest pre-Sehoo unit
(Q16T) at Thorne Bar is about 1350 m; it is characterized by a
relatively well-preserved barrier and is henceforth referred to
as "younger lake deposits of Thorne." The delta plain between
1360 m and 1345 m in altitude, north of the Humboldt River
near Treaty Hill, is provisionally correlated with this lake unit.
Maximum altitudes of the next older lake units range from
1356 m to 1378 m at the three tectonically deformed sites
(units Q14M, Q13C, and Q14S). Unit Q15T at the Thorne Bar site
has not been deformed, and its maximum altitude is about 1370
m, which corresponds with the shoreline angle cut on basalt
hills north of Weber Reservoir (Figure 2). We infer that these
units represent a highstand at about 1370 m and are herein
termed "older lake deposits of Thorne." The lower of the two
shoreline benches at two sites in the Smoke Creek Desert sub-
basin is correlated with these deposits.
The beach-gravel units overlying older units containing
tephra at the sections north of Thorne Bar and at McGee Wash,
Campbell Valley, and Lone Tree Hill are also those that rise to
the highest altitudes at each site: about 1400 m at Thorne Bar
and McGee Wash (units Q14T and Q13M); 1390 m at Campbell
Valley (unit Q12C); and 1414 m at Lone Tree Hill. These data
suggest, but do not prove, that these units represent the same
highstand at about 1400 m altitude. We also tentatively assign
unit Q13S (altitude 1414 m) at Sunshine Amphitheater to this
period, as well as the higher of the two shoreline benches at
sites in the Smoke Creek Desert subbasin. This unit is named
"lake deposits of McGee" for the arcuate berm composed of
beach deposits at that locality.
The tephra layer within unit Q13T north of Thorne Bar is ten-
tatively correlated with the Bishop ash bed. The correlation is
based upon the chemical composition (Table 3) and strati-
graphic position of the tephra layer (Figure 3). The tephra-ma-
trix bed in unit Q12L at Lone Tree Hill (Figures 15, 16) has the
chemical composition of the Bishop ash and overlies unit Q11L
containing probable Glass Mountain tephra from 1 Ma. This
stratigraphic relation strengthens the interpretations that (1)
the tephra layers near Thorne Bar and Schurz (discussed
above) are the Bishop ash, and (2) this unit represents a sepa-
rate lake cycle, herein termed "lake deposits of Lone Tree"
(Figure 5).
Unit Qllc at Campbell Valley, unit Q12M at McGee Wash,
and unit Q11L at Lone Tree Hill contain a Glass Mountain
tephra that erupted at about the time of the Jaramillo Subchron,
which occurred about 1 Ma. That would correlate this unit with
SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Morrison's (1991) younger pre-Eetza unit west of Weber Res-
ervoir (Figure 2). Unit Q12S at Sunshine Amphitheater also was
deposited during and after the Jaramillo Subchron. We name
this unit "lake deposits of Campbell" for the tephra in lacus-
trine sediments at this locality.
Paleomagnetic data from Sunshine Amphitheater show that
the two oldest lacustrine units there lie within the Matuyama
Reversed Chron. These units may be correlative to older units
at Thorne Bar, McGee Wash, and Lone Tree Hill, and to Morri-
son's (1991) older pre-Eetza unit west of Weber Reservoir.
Unit Qlls is from between 1 and 1.4 Ma, based on paleomag-
netic data and fossil ages, and is herein named "lake deposits of
Sunshine." Unit QT1 is from more than 1.4 Ma but less than 2.6
Ma (timescale of Cande and Kent, 1995). The lowest lake unit
in the core from Lone Tree Mine (Figure 15) may be equivalent
to unit QT1 in the Walker Lake subbasin. Because unit QT1
seems to represent different environmental conditions than the
younger lakes (paper shale at Sunshine Amphitheater, gypsifer-
ous mudstone at McGee Wash) and may include deposits of
different ages, we do not assign it a pluvial-lake name.
The combined shoreline sites can be used to infer a partial
history of connections among the Lahontan subbasins (Figure
1). Shoreline evidence from Freds Field in the Smoke Creek
subbasin and from Thorne Bar in the Walker Lake subbasin, as
well as deltaic deposits of the Humboldt River northeast of
Treaty Hill, indicate the presence of highstands at altitudes of
about 1400 m, 1370 m, and 1350 m (Figure 18), implying that
the northern and southern subbasins have periodically merged
into a single lake since the early-middle Pleistocene. In the
Lone Tree area, faulted deposits of the highest shoreline, as-
signed to the lake deposits of McGee, are as high as 1414 m in
altitude on the footwall block. The lake deposits of Lone Tree
and Campbell represent minimum highstands at about 1355 m
in the Walker Lake subbasin and at 1405 m and 1390 m, re-
spectively, on the footwall block of Lone Tree Hill. Total dis-
placement in the last one million years is at least 40 m on two
of the faults at Lone Tree Hill, and the orientation of the faults
suggests that uplift of Lone Tree Hill could account for a sig-
nificant portion of the displacement. Thus, the discordance in
altitudes still permits the possibility of a contiguous Lake
Lahontan from the middle Humboldt River to Walker Lake as
early as 1 Ma.
The presence of bones of the Lahontan cutthroat trout (R.R.
Miller, University of Michigan, in litt. to M.C.R., 1996) in unit
Qlls at Sunshine Amphitheater (Figure 10) requires a drainage
connection (but not necessarily a continuous lake) between the
Walker Lake subbasin and the northern subbasins at some time
prior to 1 Ma. A continuous lake at 0.76 Ma is supported by the
appearance of Candona caudata, which can live in colder and
thus deeper water than the ostracodes that were preserved in
older deposits at Lone Tree Mine (R. Forester, USGS, in litt. to
M.C.R., 1998). Possibly these older deposits (>0.76 Ma) reflect
damming of the Humboldt River above its canyon east of Win-
nemucca, as suggested by Roberts (1965). These data do not
refute King's (1993) interpretation that the Walker River
flowed into an isolated Walker Lake subbasin in the early
Pleistocene, but they may push back the timing of integration
to before 1 Ma.
The general correspondence of shoreline altitudes of the lake
deposits of McGee and the older and younger lake deposits of
Thorne in the Walker Lake subbasin and in the northern subba-
sins argues for little or no regional northward tilting during the
middle Pleistocene, although a few meters of tilt during the late
Pleistocene (Mifflin, 1984) are possible. The shorelines, how-
ever, have been significantly affected by local faulting, tilting,
and isostatic rebound. For example, the proposed correlations
among Lahontan sites (Figure 5) imply that the -1400 m
shoreline altitude recorded near Thorne Bar has been raised
about 14 m at Sunshine Amphitheater, Lone Tree Hill, and
Treaty Hill and lowered about 10 m in Campbell Valley; the
sense of these displacements corresponds with those of
mapped faults at the four localities (Reheis and Morrison,
1997). Very high beach gravels are preserved at two sites on
the west side of Walker Lake (Table 1, Figures 2, 6) on the
footwall block of the Wassuk normal fault. If the highest rem-
nant of these gravels corresponds to the 1400 m shoreline, it
records 50-75 m of absolute uplift of the Wassuk Range since
the early-middle Pleistocene.
OTHER BASINS IN WESTERN AND CENTRAL NEVADA
Previous workers inferred the existence of ancient drainage
connections or contiguous lakes extending southeastward from
the Walker Lake subbasin to Fish Lake Valley (Figure 1). Such
"early pluvial" connections were first suggested by Hubbs and
Miller (1948) on the basis of remnant populations of Lahontan
fish, including Siphateles bicolor obesus (Girard, 1856) (tui
chub) in Fish Lake Valley and S. b. obesus and Rhinichthys os-
culus (Girard, 1856) (speckled dace) in the basin of Lake Toiy-
abe in northern Big Smoky Valley. The lowest outlet to Lake
Toiyabe at present is south toward Lake Tonopah through an al-
luvium-filled valley about 60 m above the late Pleistocene
shoreline of Lake Toiyabe. Mifflin and Wheat (1979) specu-
lated that Rhodes and Columbus Salt Marshes and Clayton Val-
ley once acted as overflow basins for drainage spilling south
from the Lahontan Basin because they contain anomalous con-
centrations of various salts relative to their drainage basins.
Pre-Pleistocene Lakes
LAKES SOUTH OF WALKER LAKE SUBBASIN.?We infer that
the 1400 m level of Lake Lahontan extended 60 km southeast
from Walker Lake through Soda Springs Valley to the south
end of Rhodes Salt Marsh (Figure 1). No definite deposits of
middle Pleistocene age have been identified in these valleys.
Older lacustrine deposits crop out along the range fronts both
east and west of Mina, Nevada, and at one site these deposits
contain a probable late Miocene tephra layer (EL-24-M, Table
NUMBER 33 89
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90 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
3) and late Tertiary diatoms (J. Bradbury, in litt. to M.C.R.,
1996). Younger deposits, probably dated at about 2.5-2.2 Ma
based upon tephra correlations (layers EL-2-RM and EL-26-
RM, Table 3) and diatoms, crop out along the highway south of
Rhodes Salt Marsh up to an altitude of about 1423 m (Figures
1, 19A). These deformed, latest Pliocene lake deposits may ei-
ther lap onto or underlie old fan gravel that blankets the hills to
the south (Figures 1 9B, 20) that, in turn, lie above a thick se-
quence of lacustrine sediment beneath Redlich Summit, the
present-day divide between Rhodes and Columbus Salt
Marshes (Reheis and Morrison, 1997). Lacustrine deposits in-
vestigated by J. Liddicoat (Columbia University) at the south
end of Teels Marsh, west of Rhodes Salt Marsh, contain a
tephra layer identical to tephra EL-2-RM in the roadcut at the
south end of Rhodes Salt Marsh (Table 3).
Lacustrine deposits occupy a wide area north of Redlich
Summit (Figures 19B, 20). They had been previously mapped
(Stewart et al., 1994) but had not been studied in detail. Recon-
naissance mapping and several measured sections for this study
indicate the presence of two or more lacustrine units separated
by unconformities or periods of desiccation. Outcrops north-
east of the highway expose an upward sequence of deltaic
sandstone, diatomaceous siltstone, green massive mudstone,
and nearshore sandstone with oscillatory ripple marks. Chemi-
cal analyses of many tephra layers indicate correlations with
tephra ranging in age from late Miocene to late Pliocene (e.g.,
EL-29-RM, EL-34-RM, and EL-50-RM, Table 2). Tephra lay-
ers near the top of one section where lacustrine deposits are in-
terbedded with fan deposits exhibit inverse grading of pumice
clasts, reflecting a beaching line (shoreline; Sarna-Wojcicki
FIGURE 19.?Photographs of lacustrine deposits south of Rhodes Salt Marsh (see Figure 1, Table 1). A, Roadcut
exposes two sets of late Pliocene lacustrine sediments containing tephra layers (Tables 2, 3) separated by fluvial
or deltaic deposits. B, View northeast near Redlich Summit of late Miocene to late Pliocene lacustrine deposits
filling divide between Rhodes and Columbus Salt Marshes (power poles included for scale).
NUMBER 33 91
118? orw 118? OO1 117?
Section 5 Section 7 Section 6
15 m omitted
15 m omitted
Unit Qfo
Reddish; cross-
bedded, some
debris flows
Unit Tim
Common thin,
40 m omitted altered tephra
layers
Unit Tim
Dip change;
? minimum thick-
nesses below /
EL32-RM/
Green; massive,
saline^'many small
folds' and faults
Reddish; cross-
bedded, some
debris flows
EL31-RM
Gray; thin-bedded,
common oscillatory
ripple marks
Unit TIs
Coarsens upward
Probably same as
tephra EL47-RM
Green
Probably same as
tephra EL46-RM
Green
EL51-RM
Reddish;
mudcracks
Green; ripple
marks
EL33-RM
EL50-RM
Unit Tim
EL49-RM
EL48-RM
Green, Unit Tld
diatomaceous
***.'EL45-RM
EXPLANATION
p~-^| Mudstone
Igidgij Interbedded mudstone
N=^=H and sandstone
Siltstone and sandy
siltstone
Sandstone and silty
sandstone, locally
conglomerate beds
Sandstone and con-
glomerate, poorly to
moderately sorted
Fanglomerate, poorly
sorted and bedded
| xxx| Tephra layer and
EL31-RM sample number
&
? 5
Qfy
38?14'N
Qfm
Tu
Ta
TIs
38? 13'
EXPLANATION
Fan deposits (Holocene)
Fan deposits (late and
middle Pleistocene)
Fan deposits (middle and
early Pleistocene)
Tertiary sediments, undiffer-
entiated (late Pliocene)
Tertiary alluvial deposits
(late Pliocene)
Tertiary lacustrine deposits (late
Miocene? and Pliocene).
Locally divided into:
Nearshore sandstone and
conglomerate
Diatomaceous sandstone
and siltstone
Mudstone
Deltaic sandstone and
conglomerate
Stratigraphic relation of
surface and underlying unit
Tephra locality
Line of measured section
Fault; bar and ball on
downthrown side
Strike and dip of beds
1 2 KMJ
0 1 MILE
CONTOUR INTERVAL: 80 FEET
FIGURE 20.?Preliminary geologic map and mea-
sured sections of area northwest of Redlich Summit
between Rhodes and Columbus Salt Marshes (see
Figure 1, Table 1).
92 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
and Davis, 1991; Reheis et al., 1993). These deposits indicate
that a lake intermittently occupied the present divide between
Rhodes and Columbus Salt Marshes up until about 2 Ma.
Fish Lake Valley is located south of Columbus Salt Marsh
and presently is part of its drainage area (Figure 1). Pluvial
Lake Rennie existed in Fish Lake Valley during at least four in-
tervals from the late Pliocene to the middle Pleistocene (Reheis
et al., 1993). Lacustrine deposits in the valley include mud-
stone, siltstone, sandstone, and conglomerate. Numerous
tephra layers (e.g., EL-3-FL, EL-4-FL, and EL-44-FL, Table 3)
within the lacustrine deposits document lake episodes from
about 3.4 to 2.8 Ma and from about 2.2 to 2.0 Ma. Deposits of
similar age do not crop out in Columbus Salt Marsh.
Davis (1981) and Kunasz (1970) used data obtained during
field work and data from well records from a lithium-extraction
operation to show that Clayton Valley (east of Fish Lake Val-
ley, Figure 1) was periodically filled with fresh to brackish
lakes. Exposures on the east side of the valley indicate at least
two such intervals occurred (Davis, 1981). A recent analysis of
tephra layers (samples EL-42-CV and EL-43-CV, Table 3)
from the lower interval suggests the interval occurred between
3 and 2 Ma. Thus, the record of Pliocene lakes in Clayton Val-
ley is in accord with those in Rhodes Salt Marsh and Fish Lake
Valley (Figure 18).
The parallel histories of Pliocene lakes south of Rhodes Salt
Marsh and in Clayton Valley and Fish Lake Valley, combined
with the lack of sills or the presence of very low sills between
Columbus Salt Marsh and Fish Lake, Clayton, and Big Smoky
Valleys (Figure 18), permits the interpretation that the lakes in
these basins were interconnected during the Pliocene. In addi-
tion, we infer that a Miocene to Pliocene lake basin that at least
included Rhodes and Columbus Salt Marshes was disrupted by
faulting and (or) compression along the northeast-striking
Huntoon Valley fault system (Stewart, 1988; dePolo et al.,
1993) that lifted the area of Redlich Summit, which is presently
about 35 m above the highest known shoreline altitude in Co-
lumbus Salt Marsh (discussed below) and about 95 m above
the 1400 m high shoreline of Lake Lahontan.
Early to Middle Pleistocene Lakes
PLUVIAL LAKE COLUMBUS-RENNIE.?Reheis et al. (1993)
showed that a large lake, informally named Lake Rennie, peri-
odically existed in Fish Lake Valley during the Pliocene (dis-
cussed above) to middle Pleistocene. They noted that thick,
shallow-water deposits of Bishop ash indicate the existence of
a large lake at about 760 Ka with a highstand of about 1460 m
in elevation; a lake of uncertain size probably also existed at
about 1 Ma. These outcrops of lacustrine sediment are sup-
ported by data recorded in drilling logs for privately owned wa-
ter wells. The logs, copies of which are stored at Carson City,
Nevada, by the Water Resources Division of the U.S. Geologi-
cal Survey, record the occurrence of numerous thick beds of
clay and sandy clay that are inferred to be deep-water lacus-
trine deposits. Reheis et al. (1993) speculated that Lake Rennie
and Lake Columbus at times were contiguous through The
Gap, the present-day drainage connection between Fish Lake
Valley and Columbus Salt Marsh.
New discoveries in Columbus Salt Marsh (Figure 1) indicate
the presence of at least four lacustrine units that are higher and
older than the late Pleistocene shoreline (altitude 1400 m; Mif-
flin and Wheat, 1979; Reheis et al., 1993) of Lake Columbus.
No Quaternary faults have been mapped in this subbasin
(Dohrenwend, 1982a; Stewart et al., 1994). Outcrops in a small
hill on the south side of the subbasin expose three units com-
posed of beach gravel, lacustrine sand, and silt; each unit is
separated by paleosols that formed in alluvial or eolian deposits
(Figures 21, 22). Paleomagnetic measurements yielded mainly
indeterminate polarity as a result of diagenetic and pedogenic
alteration, but one sample from unit Q12CM has reversed mag-
netization. The uppermost unit in this section, Q13CM, consists
of tufa-cemented gravelly foreset beds that underlie a promi-
nent berm, which protrudes from the south side of the hill at an
altitude of about 1435 m (Table 1). This altitude and the well-
preserved morphology are similar to those of beach berms ini-
tially identified by J.O. Davis (Desert Research Institute, in litt.
to M.C.R., 1989) on the southeast side of the subbasin at
Coaldale Junction (Reheis and Morrison, 1997). Very well-
rounded gravel clasts (in part reworked from a nearby Tertiary
conglomerate) and tufa are present as lag on the slope above
the berm to the top of the hill at an altitude of 1452 m. The
clasts and tufa suggest a yet higher lake stand that could be
equivalent to unit Q13CM or to one of the buried lacustrine units.
A similar gravel-covered bench on the west side of Columbus
Salt Marsh extends to an altitude of 1457 m. These altitudes are
very similar to the shoreline altitude of pluvial Lake Rennie
(1460 m), and we infer that they record the highest shoreline of
the combined lakes at about 760 Ka. Evidence of the youngest
pre-late Pleistocene lacustrine unit here has been destroyed by
gravel operations, but in 1991, excavations exposed gravelly
foreset beds in a berm at an altitude of about 1417 m. This alti-
tude corresponds to those of small berms on the north side of
the subbasin (Figure 1, Table 1); surface soils on these berms
are moderately developed with argillic horizons that are not
present on the late Pleistocene deposits of Lake Columbus.
A deep pit on the north side of the subbasin exposes lacus-
trine sand and marl containing probable Bishop ash (sample
29068501, Tables 2, 3; collected by J. Whitney and P. Glancey,
USGS) that are stratigraphically beneath late Pleistocene beach
deposits of Lake Columbus (Figure 21A). The tephra layer is
impure, and no unconformities or buried soils lie between it
and the overlying late Pleistocene deposits. Thus, the tephra
may have been reworked into much younger deposits and may
not truly represent the age of the older lake unit.
In summary, deposits and berms at several locations around
Columbus Salt Marsh indicate there had been at least four
lake cycles that were older and higher than the late Pleis-
tocene shoreline: the two oldest lakes rose in altitude to at
least 1420 m, and one of these two rose to 1457 m; an inter-
mediate-aged lake rose to about 1435 m; and a younger lake
NUMBER 33 93
LakePletelocene
^JE^ deposits
e
FIGURE 21.?Pre-late Pleistocene lacustrine features in Columbus Salt Marsh (see Figure 1). A, View of west
side of deep pit on north side of valley; small dog and 1 m pick below tephra are for scale. Late Pleistocene lake
gravel at top of pit overlies a sequence of debris-flow deposits that, in turn, overlie middle Pleistocene lake
gravel containing probable Bishop ash (Tables 2, 3, no. 29068501) at bottom of pit. B, View north from borrow
pit on southwest side of valley (see Figure 1) to shorelines (arrows) cut on basalt outcrops.
rose to 1417 m. One of the two oldest lakes probably occurred
about 760 Ka.
PLUVIAL LAKE CLAYTON.?The floor of Clayton Valley has
been cored numerous times by the USGS (Pantea et al, 1981)
and by mining companies (reported in Kunasz, 1970; Davis,
1981). No exposures of lacustrine deposits of Pleistocene age
have been found. Most of the wells bottomed in the so-called
"ash aquifer," a thick (as much as 10 m), coarse-grained tephra
layer that lies at depths between about 100 and 300 m below the
playa surface. This tephra layer contains dissolved lithium salts
in a large brine pool, and it is these salts that are exploited by
the mining companies. Sediments in the cores consist of alter-
nating beds of sand, mud, gypsum, and halite, and include sev-
eral coarse tephra layers and glass-rich intervals. The mud is in-
terpreted to reflect deeper, freshwater conditions in the lake,
whereas the gypsum and halite beds are interpreted to reflect
periods of brackish water and desiccation (Davis, 1981). In one
core, Kunasz (1970) interpreted the absence of gyspum and ha-
lite below about 70 m to indicate freshwater conditions. In this
core, this depth approximately corresponds to the interval in-
cluding and below the ash aquifer. A sample of the ash aquifer
obtained by Davis (1981) from a depth of 175 m in another core
hole is tentatively identified as the Bishop ash (sample D77-2D,
Table 3). In summary, sediment beneath Clayton Valley records
a long history of intermittent lacustrine deposition, including
one relatively freshwater lake at about 760 Ka.
PLUVIAL LAKE RUSSELL (MONO BASIN).?The Mono Basin
is a structural depression that probably began forming about
4-3 Ma (Gilbert et al., 1968). Freshwater lacustrine sediments
interbedded with volcanic flows deposited during this time
94 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Columbus Salt Marsh
Section 3
EXPLANATION
Unit
Q'3CM
Colluvium
Top of bench at 1435
m; beach pebbles
and tufa to 1452 m
Moderate soil;
gypsum in matrix
P39-?
P38-?
P37--R
Strong argillic
soil
P35-?
P36--?
? 0
P37--R
Siltstone and sandy
siltstone
Sandstone and silty
sandstone, locally
conglomerate beds
Sandy pebble
conglomerate, well
rounded clasts
(nearshore deposits)
Sandy cobble
conglomerate, well
rounded clasts
(beach deposits)
Sandstone and con-
glomerate, poorly to
moderately sorted
Steep primary dips
(foreset beds)
Calcareous cement,
locally tufa
Buried soil
Unconformity
Paleomagnetic
sample and polarity
(?, polarity
uncertain)
FIGURE 22.?Measured section above borrow pit on southwest side of Colum-
bus Salt Marsh (see Figure 1, Table 1), above late Pleistocene shoreline.
crop out on the east side of the basin. These deposits are com-
posed of diatomaceous sand and silt that contain abundant fos-
sils, including molluscs and fish. Lake Russell has a well-
known late Pleistocene history, as established by studies per-
formed by Russell (1889), Putnam (1949), and Lajoie (1968).
The highstand shoreline during the late Wisconsin was at an al-
titude of 2155 m; however, the last overflow of Lake Russell
southward into Adobe and Owens Valleys (Figure 1) occurred
at a highstand of about 2188 m. This lake cycle is thought to be
equivalent to the Tahoe glaciation. Reheis et al. (2002) dis-
cussed the shoreline altitudes and changing drainage direction
of Lake Russell since the late Pliocene.
Miller and Smith (1981) and Hubbs and Miller (1948) used
fossils from the Mono Basin and modern fish populations in
other areas (no fish presently live in Mono Lake) to show that
fish had migrated southward into the Mono Basin from the
Lahontan Basin, probably during the Pliocene. Fish later mi-
grated southward from the Mono Basin into Owens Valley and
Death Valley, probably using the overflow route through
Adobe Valley (Figure 1). Russell (1889) hinted, and Hubbs and
Miller (1948) proposed, that the earlier connection between the
Lahontan and Mono Basins had been via the East Walker River
and Aurora Valley over a sill at the northeast corner of Lake
Russell that is presently at an altitude of 2237 m.
Recent discoveries around the north and east sides of the
Mono Basin (Figure 1, Table 1) include (1) lacustrine deposits
near the former southern outlet of Lake Russell, representing at
least one lake cycle that greatly exceeded the 2188 m high
shoreline, (2) beach gravels on the north side at altitudes as
high as 2268 m at one site and 2252 m at another site, and (3)
an abandoned valley, now partially filled by a basalt flow, that
lies north of the putative northern sill.
The lacustrine deposits near the southern outlet lie at an alti-
tude of about 2205-2220 m and include two tephra layers
within a coarsening-upward sequence of silt, sand, and beach
gravel. Chemical analysis indicated that these tephra (samples
EL-53-ML and EL-54-ML) are latest Pliocene in age (2.5-2.0
Ma; Table 3). A lake that at one time stood much higher than
these deposits, however, is documented by beautifully rounded
beach pebbles, large chunks of tufa-cemented beach rock, and
snail coquina that are found on slopes as high as 2243 m in alti-
tude directly above the tephra-bearing deposits. Several north-
east-striking faults displace these beach-gravel remnants; lo-
cally the maximum altitude of the beach pebbles is 2277 m.
Beach pebbles within a similar range of altitudes in areas with
little evidence of Quaternary faulting (Dohrenwend, 1982b) on
the north side of the Mono Basin (Figure 1, Table 1) corrobo-
rate the prior occurrence of very high levels of Lake Russell. In
addition, the similarity of the beach pebbles and their outcrop
appearance (including a preserved berm at one locality) to
those of the last two highstands indicates that these deposits are
Pleistocene in age. The landscape must have been very differ-
ent when Lake Russell reached these altitudes because there is
no present-day topographic barrier high enough to enclose a
lake for several kilometers east and west of the outcrops near
the southern outlet. Gilbert et al. (1968) reported numerous,
closely spaced faults that have cut basalt flows from 3 Ma, and
examination (this study) of aerial photographs has revealed ev-
idence of recent faulting activity. It is likely that faulting acted
to lower the terrain in this area, which eventually permitted
Lake Russell to drain southward.
These old lacustrine deposits around the Mono Basin repre-
sent a lake that was high enough to have overflowed northward
across the sill in the northeastern corner at its present altitude of
2237 m. It is difficult if not impossible to reconstruct prefault-
ing altitudes in this tectonically active part of the basin (Dohren-
wend, 1982b); however, the andesite-cored sill locally bears
patches of subrounded pebble to cobble gravel, suggesting a
former drainage way. A sample of this andesite (EL-KA-2) was
dated by conventional K-Ar techniques at 1.44 ? 0.06 Ma (E.
McKee, USGS, in litt. to M.C.R., 1995) (Reheis et al., 2002).
North of the sill, the drainage is northward into the west-
draining Mud Spring Canyon, which is a tributary to the East
Walker River (Figure 1). Mud Spring Canyon appears to be an
abandoned, underfit drainage; it carries water only intermit-
tently, but its floor is as much as 300 m wide and is incised as
much as 50 m below old alluvial surfaces. The alluvial depos-
its exposed in Mud Spring Canyon contain several thick, coarse
tephra beds. Chemical analyses of glass in these beds (samples
EL-35A-MS, EL-35B-MS, and EL-36-MS) indicate ages of
2.9-2.5 My (Tables 2, 3). The upper part of the canyon, near
NUMBER 33 95
the intersection of the northward drainage from the andesite
sill, has been filled twice by basaltic eruptions and subse-
quently reincised by discharge from ancient Lake Russell dur-
ing the early Pleistocene (Reheis et al., 2002).
The high beach gravels, drainage relations, and ages of lake
sediments and volcanic flows on the east side of Mono Basin
support the following scenario: an ancestral Lake Russell
reached higher altitudes than its last overflow level (2188 m)
during the late Pliocene and early Pleistocene. At least once
during this time, the lake reached an altitude between 2237 m
(the elevation of the sill) and 2277 m (the highest preserved
beach gravel) and overflowed to the north and thence west
down a precursor of Mud Spring Canyon into the East Walker
River. Fish fossils (Miller and Smith, 1981) imply that a con-
nection to the Lahontan Basin existed during the late Pliocene.
The age of the andesite underlying the sill suggests that north-
ward overflow also occurred after about 1.5 Ma. Thus, the
Mono Basin periodically may have been a tributary to the
Walker Lake basin in the late Pliocene and early to middle(?)
Pleistocene (Reheis et al., 2002).
PLUVIAL LAKE JONATHAN (KOBEH VALLEY).?The large
modern drainage basin of Lake Diamond includes several val-
leys to the southwest (Figure 1). When Lake Diamond fills and
overflows, this drainage basin becomes a tributary to the Hum-
boldt River and thus to Lake Lahontan (Tackman, 1993). Most
of the drainage basin feeds first into Kobeh Valley, which lies
just upstream of Diamond Valley and is separated from it by a
very narrow, steep bedrock canyon called Devils Gate (Figure
23A). Lake sediments in Kobeh Valley (Figures 23B, 24) repre-
sent an internally drained lake that breached its sill at Devils
Gate and became a tributary to Lake Diamond sometime after
670 Ka (Davis, 1987). We propose to name this former lake
"Lake Jonathan" for Jonathan O. Davis, who discovered it two
years before his untimely death. Lacustrine deposits in Kobeh
Valley are extensively exposed in the eastern part of the valley
near Lone Mountain, which formed an island in Lake Jonathan
(Figures 1, 23B). Reconnaissance field work combined with
data from Davis (1987) indicates a long and complex lacustrine
history, which is worthy of a far more detailed investigation.
Two measured sections expose deposits of at least three lake
units. In section 16, normally magnetized sediments of two of
these lake units are separated by an unconformity and collu-
vium (Figures 24, 25). The upper lake unit contains the Rye
Patch Dam ash bed (dated at 670 Ka) (AA-04, Tables 2, 3),
which crops out in several other localities near Lone Mountain
(samples EL-69-KV, shown in Figure 23c, and EL-70-KV, not
in tables). Ostracodes in the ash bed are indicative of alkaline,
moderately deep to deep lakes (R. Forester, USGS, in litt. to
M.C.R., 1998). This lake unit was the last deposited before the
sill was breached, and it is represented by several relatively
well-preserved, carbonate-cemented berms (e.g., Figure 23B)
in eastern Kobeh Valley. The lower lake unit contains an al-
tered tephra layer that we believe is the same as that identified
as a Bishop-like ash by Davis (1987); correlation to the Bishop
ash is supported by the paleomagnetic data.
In section 18, three lake units are separated by unconformi-
ties. The two oldest units contain ostracodes that are indicative
of alkaline, shallow to moderately deep lakes (R. Forester,
USGS, in litt. to M.R.C., 1998). Of these two, the oldest unit
is reversely magnetized, whereas the next younger, or middle,
unit contains reversely magnetized samples near the base and
normally magnetized samples above (Figure 25), suggesting
that the unit spans the Matuyama-Brunhes boundary. Sample
EL-81-KV taken from a tephra layer in the middle unit, how-
ever, is best correlated, on the basis of glass chemistry (Table
3), with one of the several tuffs of Taylor Canyon layers that
erupted from Long Valley caldera 2.2-2.1 Ma (Sarna-
Wojcicki et al., In review). This age assignment is supported
by deposits that are unconformably beneath the youngest lake
unit (dated at 0.67 Ma; see section 16, Figure 24) on the east
side of Lone Mountain. This lake unit contains a tephra (sam-
ple EL-80-KV, Tables 2, 3) that correlates with the Huckle-
berry Ridge ash bed (dated at 2.1 Ma). If this correlation is ac-
cepted, then (1) the ~14 m of normal-magnetization sediment
in the middle lake unit may have been deposited during the
Reunion Normal Subchrons within the Matuyama Chron; (2)
the lower lake unit in section 16 also may represent this time
period, so the correlation to the Bishop ash bed is incorrect;
and (3) the unconformity between these deposits and the
youngest lake unit represents about 1.5 million years (My)
rather than 100 Ky.
A third outcrop north of Lone Mountain (Figure 1, Table 1)
exposes three deposits of beach gravel separated by unconfor-
mities and paleosols; the uppermost gravel belongs to the lake
unit dated at 670 Ka. The development of paleosols on the
lower two beach gravels is moderate (weak argillic horizons
and stage II?III carbonate), and together they probably can ac-
count for no more than 200 Ky of subaerial exposure. These re-
lationships suggest that these two lake units are no older than
early Pleistocene, and that two or more lake units may be miss-
ing from section 18.
OTHER BASINS IN CENTRAL NEVADA.?Two independent
basins in east-central Nevada held Lake Hubbs and Lake New-
ark during the latest Pleistocene at shoreline altitudes of about
1902 and 1848 m, respectively (Figure 1, Table 1). Mifflin and
Wheat (1979) observed higher shorelines at the south ends of
both lakes and attributed them to regional northward tilting;
however, berms at similar altitudes on the north and east sides
of both basins were located during field reconnaissance for this
study in 1997, so northward tilting can be ruled out. Three lev-
els of beach gravels that are increasingly indistinct, hence
older, with altitude are preserved above the late Pleistocene
shorelines of both lakes (Table 1). The highest levels are repre-
sented only by rounded beach gravel lagged on bedrock slopes.
Relative preservation of berm morphology and properties of
relict soils (Figure 26) indicate significant age differences
among the lowest three shorelines (including that of the late
Pleistocene) with preserved berm crests on the southeast side
of Lake Newark. Soil properties, including solum thickness
and pedogenic CaCO3 content, suggest that the penultimate
96 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Devils Gate
Lone Mountain Location of
photo C
FIGURE 23.?Photographs of lacustrine depos-
its of Lake Jonathan in Kobeh Valley, A, View
southeast from top of bench in B on Lone
Mountain showing Devils Gate, the bedrock
channel that formed when Lake Davis over-
topped the divide between it and Lake Dia-
mond. B, View northwest showing apron of
pale-colored lake sediment capped by beach
gravel as high as 1920 m flanking the "island"
of Lone Mountain. C, Close-up of these sedi-
ments showing thick tephra layer with soft-sed-
iment deformation (shovel is about 40 cm
long); this layer (EL-69-KV) is identical in
appearance to one in another outcrop identified
as 670 Ka Rye Patch Dam ash (Tables 2, 3, no.
AA-04 JOD) by Davis (1987).
shoreline probably occurred at least 100-50 Ka and the next
older likely occurred at least 200 Ka by comparison to soils of
known age in the Basin and Range area (e.g., Reheis et al.,
1995). The soil chronosequence that formed on shorelines of
different ages in the Newark Basin has been studied in detail by
Redwine (In prep.).
Several levels of well-preserved, nested shoreline features in
Diamond Valley indicate highstands of pluvial Lake Diamond
NUMBER 33 97
Section 16
(U.S. Hwy 50)
Section 18
(SW Lone Mtn.
EXPLANATION
=:^=*? P77--N
AA04-JOD;
Rye Patch/dam
ash bed/~670 ka
-N, S5-KV
-KV
P65B--N?
P63A,B~N, S3-KV
P60A.B--N, S2-KV
P61A--N
(Bishop ash?
Davis, 1987)
S1-KV
P79--N
P78--N, S10-KV
EL-81-KV;tuffof
Taylor Canyon,
ca. 2.2 Ma
P76--R, S9-KV
P75--R
P74--R, S8-KV
Yellowish-gray; shallow lacustrine
EL72-KV
P73--R
P72--R
S7-KV; P-71-?
Yellowish-gray; cut-and-fill, fluvial
S6-KV
Mudstone
Interbedded mudstone
and sandstone
Siltstone and sandy
siltstone
Sandstone and silty
sandstone, locally
conglomerate beds
Sandy pebble
conglomerate, well
rounded clasts
(nearshore deposits)
Sandstone and con-
glomerate, poorly to
moderately sorted
-?- _,_ I Calcareous cement,
^1 locally tufa
Buried soil and
unconformity
x x x IEL81-KV Tephra layer and
' sample number
-5
-L 0
FIGURE 24.?Measured sections of lacustrine deposits of Lake Jonathan in Kobeh Valley (see Figure 1, Table 1).
Symbols are the same as in Figure 9. Question mark (?) after N indicates polarity uncertain.
as much as 25 m above its late Pleistocene shoreline (Figure 1;
Tackman, 1993). Relict soils indicate that the highest of these
is likely several hundred thousand years old (Mifflin and
Wheat, 1979); however, the decrease with time in the levels of
pluvial Lake Diamond is likely the result of progressive ero-
sion of the sill between Diamond Valley and Huntington Creek,
a tributary to the Humboldt River.
Pine Valley, northwest of Diamond Valley, was an internally
drained, fault-controlled basin with a long record of lacustrine
sedimentation (Smith and Ketner, 1976; Gordon, 1990). Dur-
ing much of its history, probably spanning the late Pliocene and
early Pleistocene, the lake was alternately alkaline and rela-
tively fresh. Sometime after the deposition of the Lava Creek B
ash bed (660 Ka), which lies near the top of the lake beds, the
northern sill was breached and the valley began draining into
the Humboldt River via Pine Valley Creek. Gordon (1990) in-
ferred that Lake Pine was never a deep freshwater lake. His
opinion was based upon the alkaline character of the sediments
and the reported absence of beach gravels. In field reconnais-
sance for the present paper, however, an approximately 10 m
thick bed of well-rounded beach gravel, locally steeply bedded,
was found at the top of the lacustrine sequence in the northeast-
ern corner of the valley. We suggest that Pine Valley, like
Kobeh Valley, was breached and integrated with the Humboldt
River drainage when an exceptionally large lake spilled over a
sill composed of easily eroded Tertiary tuff and sedimentary
rocks at, or shortly after, 660 Ka.
Discussion
AGES OF MIDDLE PLEISTOCENE SHORELINES
IN THE WESTERN GREAT BASIN
The ages of deposits that mark the very high shorelines through-
out the Lahontan Basin and in the basins of Lakes Columbus,
Hubbs, and Newark can only be estimated because the units every-
where consist of gravel and sand that have not been directly dated.
98 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
C N (UP)
W (Hor)
W (Hor)
S (down)
B N (Up)
S (down) NRM
5mT
W (Hor)
NRM
W (Hor)
0.002 A/m - -
0.01 A/m
P76
S (down) S (down)
NRM
FIGURE 25.?A-D, Orthogonal demagnetization diagrams (method of Zijderveld, 1967) shown in stratigraphic
order for specimens from section 18 in Kobeh Valley. A, reversed magnetization in sample P73 from the oldest
exposed lake unit; B, reversed magnetization in sample P76 from the lower part of the middle lake unit; C, normal
magnetization in sample P77 from the middle lake unit; D, normal magnetization in sample P78 from the upper
part of the middle lake unit. This sample is from just above tephra EL-81-KV, correlated with -2.2 Ma tephra of
the tuff of Taylor Canyon (Table 3).
The lake deposits of McGee overlie the lake deposits of
Campbell deposited about 1 Ma at three sites in the Walker
Lake subbasin, and they overlie the lake deposits of Lone Tree
deposited about 0.76 Ma (Figure 5). At Sunshine Amphithe-
ater, two layers of fan gravel and three argillic paleosols inter-
vene between the lake deposits of Campbell and the highest
beach gravel, an interval likely representing at least 200 Ky.
These relationships and the very high lake level represented by
the lake deposits of McGee suggest a correlation to marine ox-
ygen-isotope stage (OIS) 16, one of the largest deviations to-
ward glacial (and hence pluvial) conditions (Figure 18). We in-
terpret the unit to be equivalent to the Rye Patch Alloformation
NUMBER 33 99
FIGURE 26.?Photographs of soils formed on beach gravels on south-
east side of Lake Newark (see Figure 1, Table 1). A, Late Pleistocene
berm at 1845 m (trowel is 28 cm long). B, Berm at 1853 m (divisions
on tape = 10 cm intervals). C, Highest berm at 1858 m. Soils are sig-
nificantly more developed with age: late Pleistocene soil contains lit-
tle pedogenic carbonate; soil formed on next higher berm has stage
III carbonate horizon below 55 cm; soil formed on highest berm has
strongly cemented Bkm horizon below 60 cm.
100 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
deposited about 0.7-0.6 Ma (Reheis and Morrison, 1997)
along the Humboldt River. At this site, sublacustrine deltaic de-
posits containing the Rye Patch Dam (-0.67 Ma; A. Sarna-
Wojcicki, in litt. to M.C.R., 1997) and Lava Creek B (0.66 Ma;
Izett et al., 1992) ash beds crop out at an altitude of about 1285
m (Morrison, 1991). If the lake deposits of McGee (at an alti-
tude of 1400 m) correlate with the Rye Patch Alloformation,
the deltaic deposits at Rye Patch Dam must represent low-level
phases of the lake cycle. The correlation to OIS 16 is bolstered
by lake sediments in Kobeh and Pine Valleys (Figure 1). The
youngest lake sediments and highest beach gravels of Lake
Jonathan in Kobeh Valley are dated at 0.67 Ma by the Rye
Patch Dam ash (Figures 23c, 24; Davis, 1987). The youngest
lake sediments and highest beach gravels in Pine Valley are
dated at 0.66 Ma by the Lava Creek ash bed (Smith and Ketner,
1976; Gordon, 1980; and this study).
We infer, based upon the rare occurrence and poor preserva-
tion of lag beach gravel on bedrock, that the highest shoreline
deposits of Lakes Columbus, Hubbs, and Newark also repre-
sent OIS 16 (Figure 18). From its degree of soil development,
the highest shoreline of Lake Diamond (Mifflin and Wheat,
1979) also may have formed at this time, possibly as a result of
overflow from Lake Jonathan (Tackman, 1993).
Beach deposits of the -1370 m and -1350 m high shorelines
of Lake Lahontan (the upper and lower lake deposits of Thome,
respectively; Figure 5) and of the two intermediate shorelines
of Lakes Columbus, Hubbs, and Newark, locally retain the
forms of berms and wave-cut benches. In the Walker Lake sub-
basin, these beach deposits are inset into older lake units. The
relationships of these beach deposits and their morphologic
preservation cause us to conclude that the upper and lower lake
deposits of Thorne represent units that are younger than those
correlated to the Rye Patch Alloformation but are as old or
older than the last highstand represented by the upper part of
the Eetza Alloformation, which is thought to be equivalent to
OIS 6 (Morrison, 1991). Thus, these two sets of shorelines
probably represent pluvial episodes corresponding to OIS 14 to
perhaps as young as OIS 6 (-530-130 Ka). The upper lake de-
posits of Thorne and the higher of the two intermediate shore-
lines of the other lakes may be equivalent to OIS 12 (-430 Ka),
a relatively large glacial episode in the marine record (Figure
18). This age assignment is consistent with the morphology of
relict soils formed on these shorelines at the Thorne Bar site
(Adams and Wesnousky, 1999) and at the south end of Lake
Newark (Figure 26c).
The -1350 m high shoreline is only represented in the
Lahontan Basin by the lower lake deposits of Thorne (Figures
5, 18) at Thorne Bar and possibly along the Humboldt River,
but other lake basins contain numerous remnants of a lower in-
termediate shoreline 7?17 m above the late Pleistocene high-
stand (Figure 1, Table 1). Properties of relict soils at Thorne
Bar (Adams and Wesnousky, 1999) and in the basins of Lakes
Columbus and Newark (Figure 26B) indicate ages from 100 to
50 Ka or earlier. Thus, these shorelines may represent pluvial
lakes associated with OIS 6 (an intense glacial episode at about
150-130 Ka) or possibly OIS 8 (a less intense glacial episode at
about 250 Ka), but they are not likely to be as old as OIS 10
(-340 Ka; Figure 18). If the lower lake deposits of Thorne are
as young as OIS 6, it seems that more remnants of this shore-
line should have been found in the Lahontan Basin. Assign-
ment of the 1350 m high shoreline of Lake Lahontan to OIS 6
also conflicts with Morrison's (1964, 1991) evidence that the
highest Eetza shoreline is at about the same altitude as the 1330
m Sehoo shoreline; however, soil data of Adams and Wes-
nousky (1999) suggests that Lake Lahontan shorelines around
the 1330 m level may all be Sehoo in age. Recently published
uranium-series ages of tufa from the Pyramid Lake area of
Lake Lahontan show that the lake achieved at least an inter-
mediate level (-1207 m in altitude) almost continuously from
about 400 to 170 Ka (Szabo and Bush, 1996) and may have in-
termittently stood much higher, but the timing and altitude of
the highstands are unknown.
CAUSES OF VERY HIGH SHORELINES
DURING THE EARLY AND MIDDLE PLEISTOCENE
The records of lakes in Nevada described herein stand in
contrast to records presently available for lakes in the south-
western Great Basin and for Lake Bonneville in the eastern
Great Basin; however, the records in the different regions are
based upon different types of evidence. For example, interpre-
tations of the records of Lakes Owens, Searles, and Bonneville
depend largely upon stratigraphic evidence from cores and to a
lesser extent from outcrops, whereas those of the present study
depend upon shoreline deposits. The records of Lakes Owens
and Searles show that these were fresh, overflowing lakes dur-
ing periods that correspond with OIS 16, 12, 10, and 6 (Jannik
et al., 1991; Smith et al., 1997). Flow-through lakes, however,
do not leave adequately detailed evidence of changes in lake
size because lake size is controlled by sill height. Oviatt et al.
(1999) reinterpreted the middle Pleistocene record of the Bonne-
ville Basin (east of the present study) using data from cores and
concluded that four freshwater lakes were present during the
Brunhes Chron. The oldest of these lakes just preceded the
Lava Creek ash (660 Ka), and the younger three were inferred
to have occurred during OIS 12, 6, and 2 (Oviatt et al., 1999).
In contrast to Pleistocene lakes in northern Nevada, the largest
of the Bonneville lakes formed during OIS 2. The large size of
this late Pleistocene lake is not directly related to climate; it
was probably caused by an increase in water volume from the
capture of the Bear River drainage (Bouchard et al., 1998).
The descending bathtub-ring character of several Nevada
shoreline sequences of this study suggests that pluvial periods
became increasingly drier from the early-middle to the late
Pleistocene. Climate, tectonics, drainage changes, and changes
in moisture sources have probably interacted to produce the ob-
served changes in pluvial lake levels in the western Great Ba-
sin, as discussed below, but their relative importance is difficult
to determine and probably has changed with time.
NUMBER 33 101
Changes in the position of the jet stream induced by chang-
ing configurations of the Laurentide ice sheet could affect both
precipitation and temperature in the western United States.
Global circulation models show that the ice sheet induces split-
ting of the jet stream, and the movement of the two branches
and their associated storm tracks appears to account for pluvial-
lake fluctuations in the Basin and Range during the late Pleis-
tocene (e.g., Thompson et al., 1993). Previous configurations
(especially altitudes) of pre-late Pleistocene ice sheets are too
poorly known to estimate their effects, and they may have dif-
fered little from that of the late Wisconsin (except in the area
west of the Mississippi River; summarized in Hallberg, 1986).
Geologists have long believed that uplift of the Sierra Ne-
vada since the late Pliocene, estimated to be about 1 km since 3
Ma and 300 m since 0.8 Ma (Huber, 1981; Winograd et al.,
1985), has produced an increasing rain shadow to the east
through time. In contrast, recent studies have suggested that the
range was elevated before the middle Miocene and has since
been decreasing in average elevation (Small and Anderson,
1995; Weraicke et al., 1996). Nevertheless, several lines of evi-
dence indicate progressive warming and drying of at least the
southern Great Basin during the last two million years. These
include the depletion in deuterium content and the increase in
818O content of ground water (Winograd et al., 1985, 1997), the
decline in the water table (Winograd and Szabo, 1988), and the
decrease in pluvial-lake size (this study). In addition, Whitney
and Harrington (1993) used relict colluvial deposits in southern
Nevada to infer that winter temperatures in the southern Great
Basin were at least 1-3? C colder in the early and middle Pleis-
tocene than in the late Pleistocene. The apparent discrepancy
between the two views of Sierran uplift may be resolved by
Small and Anderson's (1995) model. Their model suggests that
geomorphic, rather than tectonic, uplift of the Sierra Nevada
permitted 30-200 m of uplift per one million years during the
Quaternary and that the uplift was in the maximum (as opposed
to the average) altitude of slowly eroding parts of the Sierran
crest. In addition, there is strong evidence that the Transverse
and Cascade Ranges, to the south and north of the Sierra Ne-
vada, respectively, also have increased in altitude during this
period (Priest, 1990; Hildreth and Lanphere, 1994; Page et al.,
1998; Spotila et al., 1998). Nevertheless, relative uplift of the
western Cordillera does not account for the extremely high lake
levels that we correlate with OIS 16, but it may help to explain
the progressive decrease in lake levels since then.
Regional-scale lithospheric subsidence may have decreased
the average altitudes in the northern Great Basin following up-
lift associated with emplacement of the plume head of the Yel-
lowstone hot spot centered in northern Nevada (Parsons et al.,
1994). This possible effect seems contradicted by recent inter-
pretations of fossil flora that suggest elevations in Nevada were
much higher at about 15 Ma but had decreased to about the
same as the present by about 12 Ma (Wolfe et al., 1997).
More moisture may have crossed the Sierra Nevada in the
early and early-middle Pleistocene because of the presence of
Lake Clyde, a tectonically dammed lake that filled the Great
Valley of California (Sarna-Wojcicki, 1995). Pacific air masses
that gave up moisture while passing over the Coast Ranges
could have been recharged by water that evaporated from this
large lake (30,000-50,000 km2). The added moisture could
have increased the snow pack on the Sierra Nevada and on
ranges in the Great Basin farther east. Lake Clyde overtopped
its sill, rapidly incised its outlet, and drained shortly after 660
Ka (Sarna-Wojcicki, 1995). This moisture effect helps explain
the high lake levels of OIS 16, but it does not account for the
intermediate shoreline altitudes of pluvial Lakes Lahontan, Co-
lumbus, Newark, Hubbs, and possibly Gale that formed be-
tween OIS 16 and 2.
Drainage changes have altered the size of the Lahontan
drainage basin through time. For example, Lake Russell prob-
ably discharged into the Walker River in the late Pliocene and
early Pleistocene (Figure 1; Reheis and Morrison, 1997; Reheis
et al., 2002). The addition of water from Lake Russell could
help to account for high lake levels during the early Pleistocene
in the Walker Lake subbasin, if it were not integrated with the
Lahontan Basin to the north during part of this period.
Other lake basins that are now tributaries to Lake Lahontan
were previously isolated. Smith and Ketner (1976) showed that
Pine Valley (PV, Figure 1), presently a tributary to the Hum-
boldt River, was internally drained until sometime after about
660 Ka (following deposition of the Lava Creek B ash bed).
Davis (1987) suggested that the lake in Kobeh Valley (Lake
Jonathan) was internally drained until sometime after about
670 Ka (following deposition of the Rye Patch Dam ash bed).
Recent field work (this study) confirms this proposal and sug-
gests that Lakes Jonathan and Diamond (Figure 1) became inte-
grated with each other and with the Humboldt River when Lake
Jonathan spilled over its eastern sill shortly after 670 Ka. The
sudden appearance of endemic ostracodes (sample S13-LT,
Figure 15) in lake deposits of Lone Tree at Lone Tree Mine
could represent a rapid influx of water from a long-lived, stable
lake not previously part of the drainage basin (R. Forester,
USGS, in litt. to M.C.R., 1998). The newly discovered highest
shorelines of Lakes Hubbs and Newark (Table 1) are very close
to modern sill levels between these two basins and between the
Newark and Lahontan Basins (Figure 18); thus, it is possible
that these lakes were temporarily tributaries to Lake Lahontan
during this old (OIS 16?) lake cycle. Beach gravel at elevations
3-7 m higher than the late Pleistocene shoreline of Lake Gale,
and at the same altitude as its present northern sill (Figure 1,
Table 1), indicate that this lake once discharged northward into
the Lake Franklin basin, as previously suggested by Hubbs et
al. (1974). In addition, other pluvial lakes in east-central Ne-
vada that were isolated during the late Pleistocene, such as
Lakes Gilbert, Franklin, and Clover (Figure 1; Mifflin and
Wheat, 1979), could have potentially discharged into the
Lahontan Basin if their water levels were about 10?40 m above
their late Pleistocene levels (estimated from modern topogra-
phy and not accounting for changes caused by faulting and sed-
imentation). These three basins as well as the Lake Gale basin
presently contain relict fish populations that indicate former
102 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
drainage connections with the Lahontan Basin via the Hum-
boldt River, as discussed below (Hubbs and Miller, 1948).
Gravel pits north of the northeastern corner of Lake Clover ex-
pose thick deposits of probable beach gravel that indicate the
lake level was as much as 32 m above the late Pleistocene high-
stand (Table 1). In addition, gravel pits on the west side of
Wells, Nevada, expose a 5-10 m thick sequence of deltaic
gravel and sand beds, including steeply dipping foreset beds.
These beds lie directly north of the lowest sill around the Lake
Clover basin and are interpreted to indicate overflow of the
Lake Gale-Lake Franklin-Lake Clover chain of lakes into a
small basin around Wells that is now the headwaters of the
Humboldt River (Reheis, 1999). Other overflows and basin
connections are also possible (Figure 1).
IMPLICATIONS FOR MIGRATION OF AQUATIC SPECIES
IN THE WESTERN GREAT BASIN
The evidence presented in this paper for connections be-
tween presently disjunct lake basins permits geological and
chronological constraints to be placed on the Pliocene and Qua-
ternary dispersion of fish and molluscs in the western Great Ba-
sin. The pioneering work of Hubbs and Miller (1948) and sub-
sequent refinements by Hubbs et al. (1974), Miller and Smith
(1981), and Smith et al. (2002) established that many of these
isolated basins contain (or once contained) species of relict fish
that are related to each other, to species in the modern Lahontan
Basin, and to species in the Columbia River basin to the north
and the Owens Lake-Death Valley area to the south. Similar
studies of springsnails (Hershler and Sada, 2002) and leeches
(Hovingh, 1997) have suggested similar connections, although
many of the migrations are probably much older than the pe-
riod of time discussed in this paper. Research conducted in the
last 20 years on material from fossil localities has shown that
many modern species of fish and springsnails are very similar
to ancestral species that were widely distributed in western
North America during the Pliocene. As a result, speciation pre-
viously attributed to rapid evolution since about 12 Ka (Hubbs
and Miller, 1948) is now more commonly attributed to inherit-
ance from widespread Tertiary populations followed by extinc-
tions and intermittent population-mixing caused by wet-dry
climatic cycles and tectonic disruptions during the Pleistocene
(e.g., Miller and Smith, 1981; Taylor and Smith, 1981; Taylor,
1985; Minckley et al., 1986; Hershler and Sada, 2002; Smith et
al., 2002).
This paper documents large, previously unknown lakes in the
western Great Basin that were intermittently connected during
the early and middle Pleistocene. This long lacustrine history
may allow the observed divergence between some species to be
averaged over half a million years rather than over several
thousand years, thereby greatly reducing the apparent discrep-
ancy in evolutionary rates. At the least, this new information
provides alternative hypotheses about rates of evolution from
independent evidence, as urged by Smith et al. (2002). In some
cases, these older Pleistocene lakes may require revision of pre-
vious conclusions that were based upon incorrect age assign-
ments. For example, Taylor and Smith (1981) reported on fossil
snails and fish from an outcrop in the Mopung Hills in the
Lahontan Basin. They noted that the fish were probably ances-
tral to Lahontan forms but were unlike known Pliocene fossils
from the Honey Lake area; therefore, they postulated that a bar-
rier had existed between the two subbasins. These lacustrine
deposits in the Mopung Hills had been assigned a Pliocene age
by Willden and Speed (1974); lacking chronologic dating,
these authors assumed incorrectly that indurated lacustrine de-
posits older than the late Pleistocene Lake Lahontan sediments
must be Pliocene in age! Similar deposits previously mapped as
Pliocene at McGee Wash and Sunshine Amphitheater (Greene
et al., 1991) are now known to represent early to middle Pleis-
tocene deposits of Lake Lahontan.
PLIOCENE AND EARLY PLEISTOCENE CONNECTIONS.?Late
Miocene to latest Pliocene lacustrine deposits that fill the
present divide between Rhodes and Columbus Salt Marshes
(Figures 1, 20) reflect the presence of a former waterway. This
connection could have allowed fish to migrate from the Lahon-
tan Basin into Fish Lake Valley, where a relict population of a
Lahontan tui chub, Siphateles bicolor obesus, survived until
historic settlement of the valley (Hubbs and Miller, 1948).
Thus, this fish could have migrated into Fish Lake Valley no
later than about 2 Ma, after which movement along the
Huntoon Valley fault system disrupted the ancient lake basin.
Siphateles b. obesus and Rhinichthys osculus lariversi Lugaski,
1972 (a speckled dace related to the Lahontan dace; Smith et
al., 2002), are found in the basin of Lake Toiyabe, in northern
Big Smoky Valley, and may have arrived by the same route. Al-
ternatively, Hubbs and Miller (1948) postulated that these fish
could have migrated southward into Big Smoky Valley from
Grass Valley (pluvial Lake Gilbert, Figure 1) by diversion of
tributary streams in the divide area between these two valleys.
The modern distribution of a springsnail (Pyrgulopsis gibba
Hershler, 1995) in the Lahontan Basin, Grass Valley, and Big
Smoky Valley supports the latter route (Hershler and Sada,
2002).
The combination of high remnants of lacustrine deposits con-
taining Lake Lahontan-related fish fossils (Miller and Smith,
1981), drainage relationships, and ages of sediments and flows
on the east side of the Mono Basin support the interpretation
that the basin was a tributary northward to the Walker Lake ba-
sin (Figure 1) in the late Pliocene and early to middle(?) Pleis-
tocene (Reheis et al., 2002). Sometime after 1.3 Ma (the age of
basalt filling a channel that carried lake discharge northward),
this outlet was abandoned and highstands of Lake Russell be-
gan to overflow to the south, permitting fish to move into the
Owens River drainage. The creation of the new outlet probably
was caused by intense faulting in the area of the southern out-
let, as evidenced by mapping studies conducted during this
study. The relatively youthful appearance of many of the fault
scarps suggests that this faulting may have followed the erup-
tion of the Bishop ash (760 Ka) (Sarna-Wojcicki et al., 1991)
from Long Valley caldera (Bailey et al., 1976); several studies
NUMBER 33 103
have concluded that the rates of vertical slip along fault zones
in the area increased markedly after this eruption (dePolo,
1989; Gillespie, 1991; Reheis and Sawyer, 1997). A deep core
taken on the floor of Owens (dry) Lake indicates that Siphate-
les bicolor, Prosopium sp. (whitefish), and Oncorhynchus cf.
clarki (trout) were present in Owens Lake by 730 Ka (Firby et
al., 1997); however, fossils identified as Chasmistes (lake
sucker), similar to C. batrachops Cope, 1883, from localities in
Oregon and the Snake River Plain, were reported in sediments
of early Pleistocene age near China Lake, south of Owens Lake
(Miller and Smith, 1981). These fossils suggest that some fish
were spread widely along the east slope of the Sierra Nevada
by the late Pliocene (Reheis et al., 2002). Three species of
springsnails in Owens Valley are most closely related to a
Lahontan species (Hershler and Sada, 2002); they may have ar-
rived by migration via the Mono Basin, as discussed above.
Another species {Pyrgulopsis wongi Hershler, 1995) lives in
several isolated basins along the eastern side of the Sierra Ne-
vada south of the Carson River (Figure 1), suggesting the oc-
currence of a widespread pre-Quaternary population that was
subsequently isolated by tectonic disruptions.
MIDDLE PLEISTOCENE CONNECTIONS.?The implications of
Lake Lahontan having a shoreline at an altitude of 1400 m,
herein estimated to have occurred about 650 Ka (within OIS
16), are staggering (Figure 1). For example, such a lake would
have (1) inundated Granite Springs and Kumiva Valleys, (2)
backed up the Truckee and Carson Rivers to submerge all or
part of present-day Reno and Carson City, Nevada, (3) backed
up the Humboldt River at least to present-day Battle Mountain,
and (4) extended 60 km southeast from Walker Lake to Rhodes
Salt Marsh. If allowances are made for local tectonics and for
sedimentation in the last half million years, the lake also would
have flooded into Dixie and Fairview Valleys and would have
inundated many other flat-lying valleys whose floors are now
only 5-30 m above the 1400 m level.
Such projections allow specific hypotheses to be proposed
and tested. For example, Hubbs and Miller (1948) postulated
an "early pluvial" connection between Lake Dixie and Lake
Lahontan (Figure 1) on the basis of a subspecies of S. b. obesus
that inhabits a present-day spring in Pleasant Valley, between
Dixie Valley and Grass Valley (an arm of Lake Lahontan). The
modern distribution of the springsnail P. gibba also supports
this connection (Hershler and Sada, 2002). As shown in Figure
1, the 1400 m level of Lake Lahontan probably would have
flooded Dixie, Fairview, and Pleasant Valleys (the present low-
est sill height is only 8 m above this lake level; Reheis and
Morrison, 1997), providing easy access for migrating fish. An
enigmatic outcrop in eastern Fairview Valley provides support-
ing evidence for such a connection. At this site, possible beach
gravel extends up to an altitude of 1384 m; this gravel rests un-
conformably on fine-grained lacustrine deposits containing
tephra of probable early Pliocene age (samples EL-59-FV and
EL-66-FV, Tables 2, 3). Further searches should be made in
Dixie and Fairview Valleys for better sites to confirm their pos-
tulated connection to Lake Lahontan.
It is also likely that Lake Lahontan, while at its 1400 m high-
stand about 650 Ka, flooded Granite Springs and Kumiva Val-
leys. These valleys are located within a large "island" in the
center of the Lahontan Basin, and each held a small, isolated
pluvial lake during the Sehoo highstand of Lake Lahontan (Fig-
ure 1; Mifflin and Wheat, 1979). No search for deposits of the
1400 m highstand has been done in this internally drained area;
however, pre-late Pleistocene inundation is supported by the
existence of leeches (Erpobdella sp.) in present-day springs
within the "island" that are identical to species in springs of Ea-
gle Lake, Madeline Plains, and Winnemucca Lake (P. Hovingh,
Utah Museum of Natural History, in litt. to M.C.R., 1997)
within the Lahontan Basin. The springsnail Pyrgulopsis gibba
also lives in both the Lahontan and Granite Springs Basins
(Hershler and Sada, 2002).
Several basins in central Nevada show evidence of much
higher lake levels in the past. Results of the present study com-
bined with previous research show that two of these basins,
Kobeh Valley (Davis, 1987; Tackman, 1993) and Pine Valley
(Smith and Ketner, 1976), were internally drained basins con-
taining pluvial lakes until at least about 670 and 660 Ka, re-
spectively, and that they subsequently integrated with the
Lahontan Basin. It is likely that shortly after 660 Ka, very high
stands of Lake Jonathan (Kobeh Valley), and perhaps of Lake
Pine (Pine Valley), triggered overflows that caused the integra-
tion of these drainage basins with the Lahontan Basin. If so,
fish presently living in these valleys, including populations of
S. b. obesus and R. osculus (Hubbs et al., 1974), may not have
arrived until after 660 Ka.
Hubbs et al. (1974) proposed that Lakes Hubbs and Newark
may have once been connected to the Humboldt River (Figure
1) by overflow from Lake Hubbs to Lake Newark through Beck
Pass and thence from Lake Newark northward to Huntington
Creek. Their best evidence for overflow of Lake Newark came
from the presence of two subspecies of S. bicolor in modern
springs of Newark Valley; no fish were found in Long Valley
(Lake Hubbs). They observed what they considered to be wave-
cut terraces nearly up to the height of the sill (1891 m in alti-
tude) at the north end of Newark Valley. Although Mifflin and
Wheat (1979) reported no shorelines this high, field work for
the present study found definite beach gravel as high as
1887-1890 m in altitude at the north end of Newark Valley,
only 1-4 m below the modern divide (Table 1); there is also a
large drainage channel leading north into Huntington Creek that
begins just below the summit. Beach gravel is preserved at least
as high as about 1927 m in Long Valley, which is 25 m above
the late Pleistocene shoreline of Lake Hubbs, and possibly as
high as 1940 m, which is about 6 m below the modern sill at
Beck Pass. The highest shorelines of Lake Hubbs and Lake
Newark may have formed at the same time as those of the 1400
m highstand of Lake Lahontan (Figure 1) and of Lake Jonathan
and Lake Pine, i.e., at about 650 Ka (Figure 18). The highest
shoreline of Lake Newark is so close to the present northern di-
vide that it is very likely that the lake overflowed northward
into Huntington Creek at this time, providing a migration route
104 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
for the fish presently in Newark Valley. We speculate that Lake
Hubbs may have overflowed into Lake Newark at the same
time. These hypotheses can be tested by more detailed field
work and mapping.
We further speculate that other isolated fish-bearing basins in
central and eastern Nevada could have been populated by fish
and springsnails during pre-late Pleistocene overflows (Figure
1). These basins held Lakes Gilbert (Grass Valley: R. osculus,
P. gibba), Gale (Butte Valley: Relictus solitarius Hubbs and
Miller, 1972), Franklin (Ruby Valley: R. solitarius, Pyrgulopsis
leporina Hershler, 1998), and Clover {R. osculus and S. bi-
color) (Hubbs et al., 1974; Hershler and Sada, 2002; Smith et
al., 2002). We acknowledge, however, that some of the fish
species?especially R. solitarius?are so different from their
nearest relatives in the Lahontan Basin that they could have re-
sulted from divergence from fish that were widely distributed
across Nevada in the Pliocene, when conditions were much
wetter and drainage basins were configured differently (For-
ester, 1991). The modern divide between Grass Valley and
Crescent Valley (part of the Lahontan drainage basin) is 36 m
above the late Pleistocene shoreline of Lake Gilbert (Figure 1;
Mifflin and Wheat, 1979). As noted by Hubbs et al. (1974), the
altitude of this divide has probably been affected by faulting.
Lakes Gale, Franklin, and Clover form a chain of lakes with
very low divides that are less than 15m above their late Pleis-
tocene shorelines (Figure 1). Both Hubbs et al. (1974) and Mif-
flin and Wheat (1979) noted the presence of a channel leading
north from the present divide between Butte and Ruby Valleys;
reconnaissance for the present study found beach gravels at al-
titudes at and below the altitude of the lip of this channel in
northern Butte Valley (Figure 1, Table 1). These authors ex-
plained that overflow from Lake Gale to Lake Franklin in pre-
late Pleistocene time enabled the same fish to inhabit these two
lakes. The divide between Ruby and Clover Valleys is even
lower, and it seems very likely that Lake Franklin also over-
flowed into Lake Clover, although Hubbs et al. (1974) found no
evidence of a connecting channel and thought that such an
overflow had not occurred because of the disparity in fish spe-
cies between Ruby and Clover Valleys. Nevertheless, R. soli-
tarius, which apparently evolved from a Lake Lahontan-type
dace (Hubbs et al., 1974), had to have reached Ruby and Butte
Valleys by some means, and the obvious route was by discharge
from Lake Franklin to Lake Clover and thence northward into
the Humboldt River. Recent discovery of probable beach gravel
that lies 32 m above the late Pleistocene shoreline of Lake Clo-
ver as well as deltaic deposits at Wells, just north of the present
lowest sill of Lake Clover (Table 1), suggests that a discharge
had occurred. Such a connection is also supported by the occur-
rence of the springsnail P. leporina in both Ruby Valley and the
upper Humboldt River drainage (Hershler and Sada, 2002).
Summary
The large study area and the amount of detail presented in
this paper is dauntingly complex. In the following summary of
results, we present a speculative but coherent framework of hy-
potheses that provides context for the results and a springboard
for future integrative studies to test the hypotheses. We encour-
age other researchers to pursue these tests.
During the Pliocene to middle Pleistocene, pluvial lakes in
the western Great Basin repeatedly rose to levels much higher
than those of the well-documented later Pleistocene. Early mi-
grations of aquatic species were facilitated by connections
among Pliocene lakes, which were subsequently disrupted by
tectonic movements. The geologic evidence for these Pliocene
connections supports recent work summarized by Smith et al.
(2002) and by Hershler and Sada (2002), indicating that many
of the modern distributions of aquatic species result from pre-
Quaternary dispersions.
The parallel histories of Pliocene lakes in Rhodes Salt Marsh,
Clayton Valley, and Fish Lake Valley, combined with the lack
of sills or very low sills between Columbus Salt Marsh and
Clayton, Fish Lake, and Big Smoky Valleys, support the inter-
pretation that these lakes were interconnected during the
Pliocene. In addition, we infer that a Miocene to Pliocene lake
basin that at least included Rhodes and Columbus Salt Marshes
was disrupted by faulting and (or) compression. The timing of
these connections indicates that fish and possibly springsnails
must have migrated into Fish Lake Valley from the north before
about 2 Ma. The populations of fish and springsnails in north-
ern Big Smoky Valley may have arrived by the same route or
across a higher divide from Grass Valley (Lake Gilbert) to the
north (Hubbs and Miller, 1948), or they may be relict from
older, more widespread ancestral populations (e.g., Taylor,
1985;Minckleyetal., 1986).
The combination of high remnants of lacustrine deposits con-
taining Lake Lahontan-related fish fossils (Miller and Smith,
1981), drainage relationships, and ages of sediments and flows
on the east side of the Mono Basin indicate that the Mono Basin
periodically was a tributary to the Walker River in the late
Pliocene and early Pleistocene, permitting Lake Lahontan fish
to enter the Mono Basin. Some of these fish then migrated
southward into the Owens River drainage when a new, lower
outlet opened sometime after 1.3 Ma and possibly as recently as
after 760 Ka. Some fish, however, may have already reached
the area south of Owens Lake before the early Pleistocene
(Miller and Smith, 1981).
Several outcrops in the Lahontan Basin document highstands
of a lake that rose far above the Sehoo and Eetza(?) shorelines
during the early and middle Pleistocene. A conservative corre-
lation among all the known sites yields a total of seven lacus-
trine units whose shorelines exceeded that of the Sehoo lake;
three of the units postdate the lake unit containing a tephra pro-
visionally correlated to the Bishop ash bed dated from 0.76 Ma,
and three predate the tephra-containing unit but are younger
than 2.6 Ma. The Walker Lake subbasin became integrated with
the northern Lahontan subbasins probably by 1 Ma (lake depos-
its of Campbell) but possibly as recently as about 650 Ka (lake
deposits of McGee). The youngest three units represent middle
Pleistocene shorelines stepped progressively downward with
time. The lake deposits of McGee, representing the highest
Lake Lahontan shoreline preserved at an altitude of 1400 m, is
correlated with the Rye Patch Dam Alloformation from 670 to
630? Ka (Morrison, 1991) and with marine OIS 16. The
younger two shorelines, at altitudes of about 1370 and 1350 m,
are tentatively correlated with OIS 12 (-430 Ka) and OIS 10
(-340 Ka) or younger glacial stages, respectively.
NUMBER 33 105
Similar records of undated shorelines descending from an
early-middle Pleistocene highstand are preserved in the basins
of Lakes Columbus-Rennie, Hubbs, and Newark. Very high
stands also were achieved shortly after about 660 Ka by Lake
Jonathan and Lake Pine. The geographic distribution of these
lakes indicate that the causes of the very high stands were not
limited in their effects to western Nevada; how far east the
higher lake levels of middle Pleistocene age can be found is un-
known, but they have not been reported in the Bonneville Basin
of Utah (Oviatt et al., 1999). Complex interactions among cli-
mate, tectonics, drainage changes, and moisture sources are re-
sponsible for the observed differences in pluvial lake levels in
the western Great Basin. These interactions could be tested by
global circulation models.
At about 650 Ka, pluvial lakes in Nevada reached such high
levels that numerous interbasin thresholds were likely over-
topped. In the cases of Lake Jonathan and Lake Pine, these
overflows caused the basins to become permanently or periodi-
cally part of the Lahontan drainage basin. In other cases, such
as Granite Springs and Dixie Valleys, Lakes Hubbs and New-
ark, and perhaps simultaneously Lakes Gilbert, Gale, Franklin,
and Clover, the overflows were temporary. These connections
provided potential migration routes for relict populations of
Lake Lahontan-type fishes and springsnails that presently re-
side in internally drained basins, especially those species and
subspecies that apparently have diverged little from their mod-
ern relatives in the Lahontan Basin. This scenario presents a
startling vision of a western Great Basin in which climatic and
tectonic effects conspired to integrate presently isolated basins
into one enormous drainage system during at least one brief pe-
riod during the Quaternary. Much work remains to be done to
test this hypothetical vision.
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Pleistocene Lakes and Paleoclimates (0 to 200 Ka)
in Death Valley, California
Tim K. Lowenstein
ABSTRACT
The existence of an ancient lake in Death Valley, California,
named Lake Manly by geologists in the early part of the twentieth
century, was suggested on the basis of shoreline features (gravel
deposits, erosional terraces, and carbonate tufas) and basin-center
muds and evaporites. Radiocarbon dating of black lacustrine muds
and uranium-series dates on shoreline tufas later established that
permanent lakes occurred in Death Valley during at least two dif-
ferent periods: marine oxygen isotope stage 2 (26 thousand to
about 10 thousand years ago (Ka)) and marine oxygen isotope
stage 5e/6 (>200 to -120 Ka).
The sources of inflow water to Death Valley during wet pluvial
periods were the Amargosa River, the Owens River system, and
the Mojave River. The timing of spillover from the Owens River
system via Panamint Basin into Death Valley is still uncertain;
inflow from the Mojave River is interpreted to have occurred dur-
ing the last 18 Ky, but more precise timing and the significance of
the Mojave River inflow is unknown.
A borehole core (DV93-1) taken from Badwater Basin, Death
Valley, in 1993, is composed of interbedded salts and muds. Core
DV93-1 contains a 186 m long, 200 thousand years (Ky) record of
closed-basin paleoenvironments and paleoclimates, which is inter-
preted on the basis of sedimentology, ostracode assemblages,
homogenization temperatures of fluid inclusions in halite. The
lacustrine sediments in core DV93-1 correlate with dated shoreline
tufa. The 200 Ky paleoclimate record is dominated by two cycles
of dry and/or warm conditions followed by wet and cold condi-
tions that occurred on a 100 Ky time scale. Of significance are the
major lacustrine phases of 35?10 Ka and 186-120 Ka (marine
oxygen isotope stages 2 and 5e/6, respectively), which represent
dramatically colder and wetter conditions than those of modern
Death Valley. Of the two major lake periods, the penultimate gla-
cial lakes were deeper and lasted far longer than those of the last
glacial.
Tim K. Lowenstein, Department of Geological Sciences, State Univer-
sity of New York at Binghamton, Binghamton, New York 13902, USA;
e-mail: lowenst@binghamton. edu.
Reviewers: D.E. Anderson, Center for Environmental Sciences and
Education, Northern Arizona University, Flagstaff, Arizona 86011,
USA; David B. Madsen, Environmental Sciences, Utah Geological
Survey, Salt Lake City, Utah 84114, USA.
Previous Work on the Ancient Lakes of Death Valley
The first published evidence of an ancient lake in Death Val-
ley, California, is from Noble (1926), who briefly described
strandlines along a basalt hill in southern Death Valley (Shore-
line Butte) and in gravel deposits at Mormon Point (Figure 1).
These former shorelines were, according to Noble (1926), 122
to 153 m, respectively, above the low point of Death Valley.
The ancient lake in Death Valley was named Lake Manly by
Means (1932) for William Manly, who traveled through Death
Valley in 1849?and survived. Blackwelder (1933) described
further evidence of former Lake Manly, including the mud de-
posits underlying the modern surface salt crust in the center of
Death Valley, and at least 12 erosional terraces on Shoreline
Butte that were up to 95 m above sea level and contained
rounded gravels and carbonate coatings. Less pronounced ter-
races from several locations on the southwest and eastern mar-
gins of Death Valley (Park Village and Mormon Point, respec-
tively) also were described by Blackwelder (1933), as well as
two localities where small tufas were interpreted as possible
shoreline deposits of a former lake.
Blackwelder (1933, 1954) speculated that Lake Manly ex-
isted during a glacial period. The poor preservation of shoreline
features compared with the well-documented features from
Lake Bonneville and Mono Lake led him to conclude that Lake
Manly existed during the Tahoe glaciation (early Wisconsin,
now dated at ~74 to -49 thousand years ago (Ka)) and that a
shallower lake also may have existed during the late Wisconsin
Tioga glacial (dated at -30 to -12 Ka (Phillips et al., 1990;
Phillips et al., 1996)). Blackwelder (1933) suggested that Lake
Manly was a relatively short-lived lake because of the relative
paucity of tufa and the poor development of erosional shoreline
features in Death Valley.
Hunt and Mabey (1966) followed previous workers when
they concluded that Lake Manly had been synchronous with
Wisconsin (Tahoe and Tioga) glaciation in the Sierra Nevada,
and they added several more observations to the growing list
of shoreline features. The lacustrine shoreline and near-shore
109
110 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 1.?Map of Death Valley, California, and location of
core DV93-1 (indicated by the white star).
35?37'
117?30' W
Saline pan f.?"??;:'-/--N Mud flat Elevation <305 m, predominantlyalluvial fan and delta sediments
features described by Hunt and Mabey (1966) included shin-
gled gravels along the north and east sides of Death Valley that
were up to 116m above sea level; a large gravel bar composed
of well-sorted and crossbedded gravels, near Beatty Junction at
46 m above sea level; terraces near Park Village at elevations
between sea level and 46 m above sea level; calcite-cemented
gravels along the Black Mountains between Badwater and
Mormon Point, most commonly between sea level and 61 m
above sea level; and numerous horizontal shoreline features of
Shoreline Butte.
Hooke (1972) found important new information on the an-
cient lakes in Death Valley from borehole cores taken in the
center of Badwater Basin and from shoreline carbonate tufas
plastered on the Black Mountains. Hooke (1972) was the first
to publish quantitative information on the timing of the ancient
lakes in Death Valley, which he derived from radiocarbon dat-
ing of organic carbon from the Badwater Basin cores. The
Badwater Basin cores contained brown mud, 3-15 m thick,
overlying 1-2 m of olive green mud, which graded into black
mud below. The black muds were interpreted as lake deposits
that were never subaerially exposed (Hooke, 1972). Radiocar-
bon dates from two samples near the top of the black mud in-
terval were 11,900 yr and 12,980 yr. A date of 21,500 yr was
obtained from lower in the black mud interval. Hooke (1972),
therefore, proposed that a perennial lake existed in Death Val-
ley from about 26 to 10.5 Ka. Radiocarbon dates from the ba-
sin-center lacustrine muds led Hooke (1972) to interpret the
shorelines 90 m above sea level as latest Wisconsin in age,
11-10 Ka, and he named the youngest shoreline features the
Blackwelder stand of Lake Manly. Hooke (1972) also studied
carbonate tufas from various locations in the Black Mountains.
He tentatively inferred the presence of three generations of tu-
fas on the basis of their coloration and weathering characteris-
tics: latest Wisconsin (Blackwelder stand), early Wisconsin,
and pre-Wisconsin.
Later work on dating the Black Mountain tufas by Hooke
and Lively (1979) and by Hooke and Dorn (1992) has led to
new interpretations of the Blackwelder stand. Hooke and
Lively (1979) dated 34 tufas by uranium-series (U-series)
methods. Although there were problems with the possible loss
NUMBER 33 111
of uranium from samples, and also possible 230Th contamina-
tion, a majority of the dates fell between 225 and 135 Ka,
which led Hooke and Lively (1979) and Hooke and Dorn
(1992) to conclude that the Blackwelder stand was much older
than previously thought and likely coincided with Illinoian gla-
ciations and marine oxygen isotope stage 6. In addition, two tu-
fas from the high shorelines of neighboring Panamint Valley
(Figure 2) were dated by Hooke and Lively (1979). They, too,
were apparently Illinoian in age, 138?7 Ka and 137?4 Ka, al-
though loss of uranium and 230Th contamination may have af-
fected the samples. Finally, Dora (1989) and Dora et al. (1989)
performed cation-ratio dating of desert varnish and concluded
that the high Blackwelder stand receded 130-120 Ka.
Dora et al. (1990) used radiocarbon dating of rock varnish to
estimate the beginning of the recession of the latest Pleistocene
lake in Death Valley. They obtained dates of 12,970 yr from
Mormon Point (3 m above sea level) and 12,630 yr from Park
Service Housing, Park Village (5 m below sea level) (Dorn et
al., 1990). These dates were consistent with the radiocarbon
dates obtained by Hooke (1972) for the existence of a perennial
lake in Death Valley between 26 and 10 Ka.
Most recently, Meek (1997) carefully measured the shoreline
elevations at Shoreline Butte and Mormon Point. The promi-
nent 94.7 m bench at Shoreline Butte was interpreted as an Illi-
noian-age (marine oxygen isotope stage 6) shoreline, marking a
lake 1606 km2 in area and having a volume of 176 km3. Meek
(1997) interpreted the shorelines around 5 m above sea level as
late Wisconsin in age (marine oxygen isotope stage 2). The
highest shoreline in Death Valley, 135.4 m above sea level at
Shoreline Butte, was interpreted by Meek (1997) as being older
than marine oxygen isotope stage 6. Meek (1997) discussed the
difficulties involved in correlating shorelines from different lo-
calities because the timing and magnitude of faulting are un-
known. For example, the highest shoreline features at Mormon
Point (104-113 m) and the uppermost shoreline at Shoreline
Butte (135.4 m) may represent different lake phases or they
may represent the same lake phase with faulting of approxi-
mately 26 m between locations (Meek, 1997).
D.E. Anderson and Wells (1997) and Anderson (1998) inves-
tigated late Wisconsin-age lake sediments from 10 shallow
cores taken in southern Death Valley. On the basis of sedimen-
tary textures and accelerator mass spectrometry (AMS) radio-
carbon ages, Anderson (1998) identified three late Pleistocene
highstands at -26, -18, and -12 Ka, which are in complete
agreement with the work of Hooke (1972).
Lake Manly was fed by inflow from the Amargosa River.
This river currently drains areas to the east of Death Valley in
southwestern Nevada and is a source of inflow water into the
southern end of Death Valley. In the Pliocene and Pleistocene,
Lake Tecopa, with an estimated area of 235 km2, may have
been the terminal lake of the Amargosa River drainage (Figure
2) (Morrison, 1991). Correlation of tephra layers indicates the
breaching of Lake Tecopa may have occurred between 170 and
150 Ka (Morrison, 1991). After that time, the Amargosa River
flowed directly into Death Valley.
Additional inflow sources that may have supplied water to
Death Valley during wet pluvial periods were (1) the Owens
River system to the west of Death Valley, which receives water
from the Sierra Nevada, and (2) the Mojave River from the
south, which receives the drainage from the San Bernardino
Mountains (Figure 2) (see for example, Blackwelder, 1933,
1954; Smith, 1979, 1984; Smith and Pratt, 1957; Jannik et al.,
1991; Brown and Rosen, 1995). The Owens River system de-
rived waters from the Sierra Nevada and eventually reached
Death Valley via a system of lakes and spillways, namely
Owens Lake, China Lake, Searles Lake, Panamint Lake (Lake
Gale), and finally Death Valley (Lake Manly) via Wingate Pass
in southwestern Death Valley (Gale, 1914; Blackwelder, 1954;
Smith and Pratt, 1957; Smith, 1979, 1984; Jannik et al., 1991).
The inferred connection between Panamint Lake and Lake
Manly at Wingate Pass is controversial because evidence of
erosion by spillover flow from Panamint Lake into Death Val-
ley is equivocal. Hooke (1972) interpreted four hills composed
of gravels as delta deposits, implying they were formed by wa-
ter spillover from Panamint Lake. Tufa deposits, along with
erosional benches and shallow-water gravel deposits, have
been described by R.S.U. Smith (1978) from various locations
in Panamint Valley between elevations of 502 and 627 m above
sea level. Some of these shoreline features are at or above the
elevation of Wingate Pass (605 m), which is strong evidence
for spillover into Death Valley, ignoring the effects of faulting.
Fitzpatrick and Bischoff (1993) obtained U-series dates on
samples from the high shorelines of Panamint Lake but argued
that the dates may be unreliable because of post-depositional
mobilization of uranium. They concluded that the shoreline
features are between 300 and 30 Ka in age but probably do not
represent late Pleistocene marine oxygen isotope stage 2 (Fitz-
patrick and Bischoff, 1993). Their isochron plots suggest ages
between 95 and 55 Ka; therefore, although water clearly en-
tered Death Valley via spillover from Panamint Lake, the tim-
ing of that spillover is still uncertain.
Inflow of water from the Mojave River into the south end of
Death Valley was discussed by Brown and Rosen (1995), who
concluded that Lake Manly may have received inflow from the
Mojave drainage area only in the latest Pleistocene. This con-
clusion was based on the age of Pleistocene Lake Mojave (now
Silver Lake playa near Baker, California) to the south of Death
Valley (younger than 22 Ka) and on the supposition that for
water to overflow into Death Valley from the Mohave River
drainage area, Lake Mojave itself must have been at or above a
spillover level (Figure 2). This interpretation is supported by
work on the Manix Basin by Meek (1989; pers. comm., 1996).
From the study of the sediments at Afton Canyon, near Lake
Manix, Meek (1989; pers. comm., 1996) concluded that Lake
Manix was the terminal lake in the Mojave River system for
most of the Pleistocene until Afton Canyon was eroded some-
time after 18 Ka. At that point, waters from the Mojave River
112 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Present-day river
Kilometers
119? 118? 117? 116? 115?
FIGURE 2.?Map of pluvial lakes, modern lakes and playas,
Pleistocene drainage patterns, and modern drainages of the
Owens River, Amargosa River, and Mojave River systems.
Modified from Blackwelder (1954), Smith (1984), Jannik et al.
(1991), and Brown and Rosen (1995).
reached the Silver Lake and Soda Lake playas. During excep-
tionally high levels of Silver Lake (Lake Mojave), water may
have overflowed into the southern end of Death Valley. There
are no features reported from Death Valley indicating inflow
of large amounts of water at approximately 18 Ka, but this pe-
riod does coincide with a wet period in Death Valley when the
basin contained a perennial lake (see "Sedimentology," be-
low).
Additional information on the timing of inflow waters from
the Mojave River into the southern end of Death Valley comes
from studies of Lake Dumont conducted by K.C. Anderson and
Wells (1996, 1997). The Dumont Basin (10 km2) is a small
playa located between Silver Lake playa and the southern arm
of the Amargosa River (Figure 2). K.C. Anderson and Wells
(1996, 1997) found lacustrine sediments in two 13 m long
cores taken in Dumont Basin. They identified Lake Dumont I
(25.3-30 Ka) and Lake Dumont II (18-19.8 Ka) on the basis of
radiocarbon dates obtained from ostracode-bearing lacustrine
muds. K.C. Anderson and Wells (1996) concluded that Lake
Dumont II broke through its sill sometime after 18 Ka, which
established a direct connection between the Mojave River, the
Amargosa River, and Death Valley. That date coincides with
the timing of the erosion of Afton Canyon upstream in the Mo-
jave River system (Meek, 1989; pers. comm., 1996). It is not
known whether Lake Dumont I overflowed into Death Valley
between 30 and 25 Ka.
Analysis and Interpretation of Data
from Core DV93-1 from Death Valley
A 6.4 cm diameter core, DV93-1, was drilled to a depth of
185.5 m through Badwater Basin, Death Valley, during April
and May of 1993 (Figure 1). The drill site was near the center
of the saline pan, 3.2 km northwest of Badwater Springs. The
core is composed of interbedded salt and mud. Virtually all salt
layers were completely recovered, but retrieval of the mud in-
tervals was about 60 percent successful. The core records the
200 Ky record of closed-basin paleoenvironments and paleocli-
mates in Death Valley, as interpreted on the basis of sedimen-
tology, ostracode species, fluid-inclusion homogenization tem-
peratures in halite, saline mineralogy, and correlation with
shoreline tufa.
NUMBER 33 113
U-series
dating
(ka) Environment
Salt pan and
saline mud flat
Perennial
salinelake
Salt pan
with
ephemeral
lakes
Salt pan
with
ephemeral
lakes
Shallow
saline lake
97.1 ?6.8 ar|d salt pan
U-series
dating
(ka) Environment
Primary halite
Cavity-filling halite
Core not recovered
119.6 ?2.6
Mud flat
Perennial
salinelake
150
160
170
180
Perennial
lake
Perennial
146 + 6 saline _
lake
Perennial
lake
.3 ?8.5 Perennial
saline i
lake
Salt pan
and
shallow
saline lake
Mud flad
TpT Vertical dissolution pipes in halite
Mud
Mud with mudcracks
FIGURE 3.?Stratigraphic column, Core DV93-1, with uranium-series ages, sediment types, and interpreted dep-
ositional environments.
DATING CORE DV93-1.?The core was dated using the
23OTh/234U method, which determines the time uranium was in-
corporated into a sample (Luo and Ku, 1991; Ku et al, 1998).
The total dissolution isochron technique was used to account
for initial detrital 230Th in the samples (Luo and Ku, 1991).
Samples from 12 stratigraphic intervals were dated (Figure 3).
The samples contained preserved primary halite as well as in-
corporated clay, carbonate (predominantly calcite), and sulfate
minerals (mainly gypsum or glauberite, CaSO4-Na2SO4). Au-
thigenic 23OTh/234U and 234U/238U ratios used in calculating the
age of a halite sample were determined from linear plots (isoch-
rons) of 23OTh/232Th versus 234U/232Th and of 234U/232Th versus
114 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
238U/232Th of a series of subsamples taken from each of the 12
stratigraphic intervals. These activity ratios were counted using
alpha particle spectrometry (Ku et al., 1998). The given age er-
rors are one standard deviation, which were derived from
counting statistics and the fitting of isochrons. The age errors
are less than 11 percent, except for the uppermost sample from
7.7 m, where the large error results from the relatively large
amount of detrital 230Th initially in this relatively young sam-
ple. The reliability of the dating technique can be seen from the
general increase in age down the core. Two samples from
around 60 m (Figure 3) with very close ages, but with very dif-
ferent amounts of detrital material, and hence different 230Th/
232Th ratios, supports the validity of the isochron approach. All
dates discussed that were not obtained directly from the U-se-
ries isochron method were calculated by interpolation, assum-
ing constant accumulation rates for each sediment interval.
Sedimentation rates averaged over the entire 186 m of core
DV93-1 are lm/lKy (thousand years). Salts, predominantly
halite, were deposited rapidly, 1.7-3.8 m/lKy, whereas muds
accumulated more slowly, at rates of 0.4-1.0 m/1 Ky.
Carbonate shoreline tufas from the east side of Death Valley,
between Badwater Springs and Mormon Point were carefully
chosen to avoid samples with impurities, weathering rinds, ex-
tensive cements, or recrystallization textures (see "Shoreline
Carbonate Tufas," below). Samples were cleaned and dis-
solved, and 238U, 234U, 232Th, and 230Th were counted by alpha
particle spectrometry (Ku et al., 1998). The ages of tufas, along
with one standard deviation errors derived from counting sta-
tistics, are shown in Figure 4c. For most samples, the high ac-
tivity ratios of 23OTh/232Th, greater than 15, suggest that the
ages of samples should be close to the time of tufa formation
(Kuetal., 1998).
SEDIMENTOLOGY.?Paleoenvironments present in Death
Valley over the past 200 Ky were interpreted by comparing
sedimentary structures and petrographic features described
from modern closed-basin sediments with those observed in
core DV93-1 (Figures 3, 4A) (Hardie et al., 1978; Lowenstein
and Hardie, 1985; Smoot and Lowenstein, 1991; Roberts et al.,
1994; Li et al., 1996). The cored sediments contain a record of
closed-basin paleoenvironments ranging from hot, dry mud-
flats, similar to the modern floor of Death Valley, to temperate
climate, deep perennial lakes.
Disrupted brown muds overlain by a 0.25 m thick surface ha-
lite crust (0-7.7 m) indicate Death Valley has been a mudflat
and saline pan during the last 10 Ky (Figures 3, 4A). Further
details are lacking because most sedimentary structures from
this interval were destroyed during coring.
A perennial saline lake existed in Death Valley for an -25 Ky
period, from 10 to 35 Ka (core depths of 7.7-18 m) (Li et al.,
1996). Interbedded mud and halite cycles accumulated during
the early perennial lake stage, followed by interlayered the-
nardite (Na2SO4) and mud. A cap of chevron halite formed
during the latest lake stage. These deposits record a generally
wet climate, with fluctuating lake levels and salinities. The lack
of synsedimentary dissolution features in evaporite layers indi-
cates that accumulated salts were permanently protected from
dilute flood waters by the saline lake, which suggests that lakes
existed more or less continually, without drying out, for 25 Ky.
The mud layers are structureless, contain rare ostracodes, and
are black in color but change to gray green after exposure to
air. Halite textures include cumulate settle-out layers and bot-
tom-grown crusts, which are interlayered with black muds (Li
et al., 1996). Thenardite layers from 10 to 13 m are apparently
diagenetic in origin but may have first formed as mirabilite
(Na2SO4-10H2O), as observed in the winter in Great Salt Lake
and Owens Lake (Spencer et al., 1984; Spencer et al., 1985a;
Spencer et al., 1985b; Smith et al., 1987). The age of the the-
nardite is ~22 to ~16 Ka in core DV93-1. This age coincides
with the last glacial maximum and suggests that Na2SO4 crys-
tallized as mirabilite during winter cold water conditions, prob-
ably at temperatures below -10? C (Smith et al., 1987). Halite
from the uppermost saline lake interval (7.7-10 m) contains
chevrons, raft textures, and pisoids. These are interpreted to
have accumulated in a shallow, perennial lake that ultimately
dried out about 10 Ka (Li et al., 1996).
Saline-pan evaporites present from 18 to 60 m indicate that
Death Valley was commonly desiccated between ~35 to -60
Ka, but it did receive enough inflow water to maintain shallow
saline lakes at times and to supply the solutes required to accu-
mulate salts rapidly, at rates of 1.7 m/1 Ky (Figures 3, 4A) (Li
et al., 1996). Saline-pan halites consist of poorly bedded halite,
with little preservation of primary depositional textures. A dis-
tinctive feature of halite layers is a vertical pipework fabric ce-
mented with halite, which signifies multiple episodes of sub-
aerial exposure, dissolution of surficial halite by dilute
floodwaters, and later cementation of cavities by shallow
groundwater brines. Five thin halite intervals between 43 and
52 m (Figure 4A) (~50-54 Ka), which carry primary features
such as chevrons and rafts, indicate that shallow saline lakes
existed in Death Valley during this time.
Clayey silts, some massive and some with mudcracks and
sand-patch textures, which are diagnostic of deposition on dry
mudflats, are the dominant sediment type from 60.5 to 109 m
(Figures 3, 4A) (60-120 Ka) (Smoot and Lowenstein, 1991;
Roberts et al., 1994; Smoot and Castens-Seidell, 1994; Li et al.,
1996). Sand patches are wind-blown or water-deposited sedi-
ments, which filled depressions in efflorescent crusts on the
mudflat surface (Smoot and Castens-Seidell, 1994). Later
flooding dissolved the efflorescent salt crusts, which caused
collapse of the overlying sediments and produced internal frac-
tures, faulting, and rotation of sand patches. Saline minerals
consist of crystals of halite, glauberite (CaSO4Na2SO4), and
gypsum (CaSO4-2H2O) that formed during early diagenetic
displacive growth from groundwater brines. Interbedded rem-
nants of saline-pan halite and a total lack of lacustrine fossils,
such as diatoms and ostracodes, reinforce the interpretation that
these muds were from mudflats, not perennial-lake deposits.
One halite interval, interpreted to have been deposited in shal-
NUMBER 33 115
AGE (ka)
0 20 40 50 60 80 100 110 120 130 150 170 190
TTiiir 1 I I
DEPTH(m) 1 n i n r
CO
0 10 20 40 60 80 100 120 j,
Deep Lake
PerennialSaline Lake
Saline PanMudflat
120 140 0 180 (A)
"g L. Sappaensis
o L. Staplini
2 L. Ceriotuberosa
to C. Caudata
I I II Illl?II II INI II 1 I III I
I I I III Illl II Mill H-H h-
-H 1 H-H 1
(B):
-166 ?6-
High Shoreline
Tufa Ages
(ka)
^i^178?15 1^,7* 1?
162.5 ?2.5* . 185?15 i-216*15-
15
Gypsum
Glauberite
Thenardite
50
co 40
10
iiiiiiiiiiiiini mi i iiiiiiiiiiiiiiiniii
in iiiiiiini niiii iiiniiiniiiiiiniiii
(E)-
(F)
DEPTH (m)
,9 MCO (D
' O
20 40 60
0 20 40 50
80 100. 120_ 140 160^ 180 (Q)
60 80 100 110 120 130 150 170
-7?
190
AGE (ka)
FIGURE 4.?Paleoclimate record, core DV93-1, during the last 200 Ky. A, Lake levels and paleoenvironments,
based on relative ranking, from driest to wettest: mudflat, saline pan, perennial saline lake, and deep lake. The
deepest lakes are interpreted from the occurrence of ostracode species Candona caudata and Limnocythere ceri-
otuberosa (R.M. Forester, pers. comm., 1995). B, Ostracodes. C, Uranium-series ages of high-shoreline (about 90
m above sea level) carbonate tufas, Black Mountains, Death Valley. D, Maximum homogenization temperatures
of fluid inclusions in halite (Ty^x )? E, Sulfate minerals, determined by x-ray diffraction analysis of bulk sam-
ples. F, Weight percent calcite, determined by x-ray diffraction analysis of the >1 \\rs\ size insoluble fraction of
the core. G, Uranium-series ages, interpolated ages, and marine oxygen isotope stages.
low saline lake and saline-pan settings, occurs from 78 to 87 m
(Figure 3). Uranium-series dates from the top and bottom of
the interval are 98.6 Ka and 97.1 Ka, respectively, which sug-
gests rapid evaporite accumulation.
Perennial-lake sediments were deposited during a 66 Ky pe-
riod, between 186 and 120 Ka (Figures 3, 4A). Dark lacustrine
muds (core depth of 127-161 m) were overlain by saline-lake
halites (core depth of 109-127 m), which were interlayered
with muds (Roberts et al., 1994, 1997). The perennial-lake
muds were massive to vaguely layered and dark gray when first
exposed to air, but they changed color to gray green over time.
Halite layers at 152 m (166 Ka) and at 137.5 to 139.5 m (146
Ka) indicate high salinities and probable shallower lake waters
occurred during these intervals. Perennial-lake halites are gen-
erally clear and lack dense, fluid-inclusion banding; they have
sorted cumulate textures produced by crystallization at the air-
brine interface as rafts, skeletal crystals, and cubes (Roberts et
al., 1994; Roberts and Spencer, 1995). It is likely that lake lev-
els and water salinities commonly fluctuated over the entire 66
Ky lacustrine phase (Roberts et al., 1997). The upper 18 m of
the perennial-lake succession (109-127 m) consist of interlay-
ered subaqueous halite (settle-out cumulates or vertical crystal
116 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
fabrics formed by slow crystallization at the brine bottom) and
mud that record two major shallowing events between 120 and
128 Ka, and final desiccation at 120 Ka.
The bottom 25 m of the core, 186 to 192 Ka (core depth of
161-186 m) is predominantly layered halite and mud that
formed in saline pan and shallow saline lake environments. The
halite layers contain primary fluid-inclusion banded-chevron
textures and cumulate settle-out layers that crystallized at the
bottom and top of a shallow saline lake (Roberts and Spencer,
1995). Vertical pipework fabrics, cemented with halite, indicate
exposure of salt layers to dilute floodwaters above the water ta-
ble, which produced the diagnostic vertical pipes (Roberts et
al., 1994). The lowest one meter of the core is composed of
mudcracked silty mud and is interpreted as a mudflat deposit.
OSTRACODES.?Ostracodes from core DV93-1 provide in-
formation on the lower limits of water salinities of the two
long-lived perennial lakes in Death Valley during the last gla-
cial, 35-10 Ka, and the penultimate glacial, 186-120 Ka (ma-
rine oxygen isotope stages 2 and 5e/6, respectively) (Figure
4B). Perennial-lake muds contain ostracode species Candona
caudata Kaufman, 1900, Limnocythere ceriotuberosa De-
lorme, 1967, L. sappaensis Staplin, 1963, and L. staplini
Gutentag and Benson, 1962, which, although not abundant, in-
dicate that the salinity of the Death Valley lakes in which the
ostracodes lived was typically below 10,000 parts per million
and at times was less than 3000 parts per million for the salinity
sensitive species C. caudata (Forester, 1983, 1987). Such
lower-salinity lake phases are inferred from ostracode occur-
rences at 10, 14, and 18 m (approximately 16, 25, and 34 Ka,
respectively) for the last glacial lake and 113-116, 125-126,
138-143, 157, and 162 m (approximately 121-123, 127-129,
146-154, 179, and 186 Ka, respectively) for the long-lived
penultimate glacial lake. Unfortunately, information on the last
glacial lake is incomplete because muds from depths of 9-14 m
(~ 12-25 Ka) were partially lost during coring.
FLUID-INCLUSION PALEOTEMPERATURES.?The paleotem-
perature component of the Death Valley climate record was ob-
tained from homogenization temperatures of fluid inclusions in
subaqueously formed halites from core DV93-1, which record
brine temperatures during salt precipitation (Goldstein and
Reynolds, 1994; Roberts and Spencer, 1995; Roberts et al.,
1997; Lowenstein et al., 1998) (Figure 4D). Shallow saline lake
temperatures, in turn, correlate closely with air temperatures in
modern settings (Carpelan, 1958; Hammer, 1986; Roberts and
Spencer, 1995; Roberts et al., 1997; Lowenstein et al., 1998).
Crystals of halite with textures diagnostic of crystallization in a
lacustrine environment (cumulates, chevrons, bottom-grown
crusts) and with initially one-phase liquid inclusions were
chilled in a freezer at temperatures of-5? C to -20? C to nucle-
ate vapor bubbles without freezing the inclusions (Roberts and
Spencer, 1995). These crystals, now with abundant two-phase
liquid-vapor inclusions, were slowly heated to the temperature
at which the vapor phase disappeared and the fluid inclusions
"homogenized" back to one liquid phase. The temperatures of
homogenization record the original brine temperatures at
which the crystal grew in the Death Valley saline lakes. Maxi-
mum homogenization temperatures of fluid inclusions in halite
(ThMAx) were used as a record of maximum brine temperature
and maximum air temperature during halite crystallization.
This method was used by Roberts and Spencer (1995) and Lo-
wenstein et al. (1998) for the precipitation of halite from Death
Valley in 1993. Halites in 66 stratigraphic intervals of core
DV93-1, from depths of 0 to 183.5 m (0-192 Ka) have fluid-in-
clusion homogenization temperatures commonly below the
modern rhMAX of 34? C (maximum brine and air temperature
during halite precipitation, late April and early May, 1993)
(Roberts et al., 1997; Lowenstein et al., 1998). Lacustrine ha-
lites deposited during the last glacial period, 35-10 Ka, have
low fluid-inclusion maximum homogenization temperatures
(ThMAX = 19-30? C), which suggests brine temperatures were
~4?-15? C below modern late April and May values. Ephem-
eral saline lake halites precipitated from 60 to 35 Ka have
rhMAX between 23? C and 28? C, which are 6?-ll? C below
the modern T^MAX values. A relatively low temperature inter-
val, 10? C below the modern, existed ~50 to -47 Ka. Limited
data from fluid inclusions in two halite intervals, -92 and -80
Ka in age, suggest generally cool conditions, 6?-7? C below
modern late April to May temperatures. Relatively higher
rhMAX values, to 34? C and 35? C in halite dated as -100 Ka
and to 32? C in halite dated as 120 Ka, may record climate re-
gimes more similar to the modern. These temperature, how-
ever, are still below modern mid-summer temperatures in
Death Valley (average air temperatures in Death Valley are
39? C in July and 37? C in August; average maximum air tem-
peratures are 46? C in July and 45? C in August).
Generally colder conditions are recorded in most fluid inclu-
sions in halite for the 186-120 Ka perennial-lake sequence,
where rhMAX values range from 25? C to 32? C, but only 3 of
22 stratigraphic intervals have T^MAX greater than 30? C (Rob-
erts et al., 1997). A combination of homogenization tempera-
tures of fluid inclusions in halite and petrographic evidence of
possible pseudomorphs of hydrohalite (NaCl-2H2O) suggest
that temperatures in the perennial lakes may have dropped be-
low 0? C at times and that temperatures during this period
probably averaged 10?-15? C below the modern (Roberts et
al., 1997). In the bottom 20 m of core DV93-1 (164-184 m,
186-192 Ka), homogenization temperatures are relatively high.
Of 19 stratigraphic intervals analyzed, 14 have ThMAX values
greater than or equal to 30? C, which is similar to those ob-
tained from halite precipitated in modern Death Valley in late
April and early May.
SALINE MINERALOGY.?Chemically precipitated minerals (sul-
fates and carbonates) in the Death Valley core provide informa-
tion on Pleistocene water inflow sources and climate (Figure
4E,F) (Li et al., 1997). Abundant glauberite (CaSO4-Na2SO4)
and gypsum (CaSO4-2H2O) and relatively small amounts of
calcite (13%?35% of the >1 |J,m size insoluble fraction) are as-
sociated with dry periods (mudflat deposits of 120-60 Ka and
NUMBER 33 117
10-0 Ka and saline-pan sediments of 192-186 Ka). In contrast,
scarce CaSO4-bearing minerals but relatively abundant calcite
(20%-50% of the >1 |j,m size insoluble fraction) are associated
with halite and mud layers from wetter periods (186-120 Ka,
60-10 Ka). Different mixing ratios of inflow waters between
wet and dry periods are interpreted to be responsible for the re-
lationship between mineral assemblages and climate.
Modern Death Valley, which is representative of ancient dry
periods, is characterized by Na-Cl-SO4 brines that are pro-
duced by mixing two basic types of inflow waters (Li et al.,
1997): (1) Na-HCO3-rich and Na-Cl-SO4-HCO3-rich meteoric
waters from the Amargosa River, from springs, and from
groundwaters from northern and central Death Valley; and (2)
Na-Ca-Cl-rich springs and groundwaters from southern Death
Valley, possibly related to volcanism, hydrothermal activity,
and a 15 km deep magma body (de Voogd et al., 1986; Serpa et
al., 1988; Li et al., 1997). During dry periods, relatively abun-
dant Ca-rich spring inflow removes HCO3 as calcite during the
early stages of brine evolution (Jones, 1966; Hardie et al.,
1978; Li et al., 1997). Further evaporative concentration pro-
duces gypsum and glauberite from the remaining Ca. During
wetter periods, increased discharge of meteoric HCC^-rich
Amargosa River water and basin-margin spring waters re-
moves most Ca from the brine via precipitation of calcite, and,
with low calcium concentrations, CaSO4-minerals are not
abundantly formed during further evaporative concentration.
Such Ca-poor, Na-Cl-SO4-rich brines precipitate thenardite
(Na2SO4) during later stages of brine evolution.
SHORELINE CARBONATE TUFAS.?Shoreline carbonate tu-
fas provide information on the depths of Pleistocene lakes in
Death Valley (Figure 4c). Shoreline tufas, most prominent ap-
proximately 90 m above sea level, form horizontal terraces en-
crusting bedrock and gravel on the eastern side of Death Valley
(Black Mountains, between Badwater Springs and Mormon
Point) (see "Previous Work on the Ancient Lakes in Death Val-
ley"; see also Hooke, 1972). Tufas were collected from loca-
tions previously described by Hooke (1972), including, from
north to south, Badwater, Goblet, Westilt, and Mormon Point
(see Figure 1 for locations) (Ku et al., 1998). Relatively fresh
hand-samples of tufas are porous and fine-grained with mas-
sive or radiating fabrics. In thin section, tufas commonly form
crusts around tubular molds, which are up to several millime-
ters in length and range in diameter from tens of microns to
about 1 mm. The tubular molds are surrounded by coatings
composed of dark micrite, and then calcite microspar, which
may have a radial fabric. These textures probably represent car-
bonate precipitation around plant material, such as green algae
and moss. A second petrographic texture exhibited by tufas
consists of clusters of dark micrite, tens of microns in diameter,
organized into millimeter-scale "bundles of grapes" similar in
appearance to the "arborescent shrubs" of Demicco and Hardie
(1994). Yet a third tufa fabric consists of bladed calcite crys-
tals, up to 1 mm long, grown from thin carbonate crusts. The
calcite crystals have upward-radiating growth fabrics, with the
number of calcite crystals decreasing away from the original
growth substrate. This fabric is interpreted to have formed by
direct chemical precipitation of calcite. Elemental mapping of
samples for Mg and Sr using an electron microprobe indicates
that chemical zonation of these elements coincides with tex-
tural and mineralogical (calcite and minor aragonite) zones in
the tufas, which confirms their unrecrystallized primary origin.
The above petrographic and geochemical studies indicate that
the tufas selected for dating have not recrystallized nor have
they undergone major later cementation, both of which could
complicate interpretation of age dates (Benson et al., 1990; Ku
etal., 1998).
Eleven samples of tufa from the 90 m shoreline were dated
by U-series methods. Of the 11 samples, 7 are 185-150 Ka,
which closely agrees with the results (200-120 Ka) obtained by
Hooke and Lively (1979) and Hooke and Dorn (1992) on sam-
ples from the same locations and the same general elevations.
The dates also correlate with perennial lake sediments ob-
served in core DV93-1 from 186 to 128 Ka; thus they serve as
an independent check confirming the reliability of the dating
methods. These tufas indicate lake depths of up to 335 m, if
faulting is ignored and if 130-160 m of basinal sediments are
added to reach equivalent-age lacustrine sediments in core
DV93-1. Four high-shoreline tufas, from 73 to 90 m above sea
level, dated between 194 and 216 Ka, are older than any sedi-
ments from core DV93-1. These oldest tufas suggest the exist-
ence of yet a third perennial lake stage in Death Valley at ap-
proximately 200 Ka, which is consistent with earlier age-dating
(Hooke and Lively, 1979; Hooke and Dorn, 1992).
Summary and Conclusions
The 200 Ky paleoclimate record of Death Valley from core
DV93-1 is dominated by two dry and/or warm and wet/cold cy-
cles that occurred on a 100 Ky time scale (Figure 4A,D). These
cycles began with mudflat deposits (192 Ka to bottom of core,
and 120?60 Ka). Wetter conditions, colder climates, or both,
produced greater effective moisture, with saline-pan and
ephemeral-saline-lake sediments (192-186 Ka, 60-35 Ka)
overlying the mudflat sediments. Eventually, enough water was
supplied to Death Valley to sustain perennial lakes, fluctuating
from fresh to saline (186-120 Ka, 35-10 Ka, respectively). The
perennial-lake phase ended, the lakes desiccated (120 and 10
Ka), and dry conditions led to the accumulation of mudflat de-
posits, completing the cycle. Of particular importance are the
major lacustrine phases from 35 to 10 Ka and from 186 to 120
Ka (marine oxygen isotope stages 2 and 5e/6, respectively),
which represent markedly colder and wetter conditions than
those in modern Death Valley. Of the two perennial-lake peri-
ods, the penultimate glacial lake was deeper and lasted far
longer than the last glacial lake.
In summary, the Death Valley climate record is divided into
six periods (Figure 4). (1) From 192 to 186 Ka, the climate
was dry, dominated by saline pans and shallow ephemeral
118 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
lakes. Inflow and evaporation were sufficient to allow the
rapid accumulation of 25 m of salt in 6 Ky. ThMAX values are
30? C or greater for 14 of 19 stratigraphic intervals, indicating
that temperatures were similar to the modern late April to early
May ThMAX values. (2) The climate from 186 to 120 Ka was
relatively cool to cold, with abundant inflow to continually
maintain a series of fluctuating perennial lakes, at times
greater than 300 m deep and having a salinity less than 3000
parts per million. Temperatures during this period probably av-
eraged 10?-15? C below the modern. If the effects of faulting
are ignored, the ages of high-shoreline tufa indicate that Death
Valley at that time commonly had lakes hundreds of meters
deep. During the latest phase of this 66 Ky lacustrine period,
from 128 to 120 Ka, reduced inflow or more saline inflow pro-
duced hypersaline conditions, which ultimately led to basin
desiccation at 120 Ka. Slightly higher fluid-inclusion tempera-
tures (ThMAX = 32? C), comparable to the modern late April
to early May ThMAX? mark the terminal lake stage at 120 Ka.
(3) From 120 to 60 Ka, a dry period dominated by mudflat en-
vironments prevailed. Limited data from fluid inclusions in
two halite intervals suggest generally cool conditions, 6?-7? C
below modern late April to May temperatures, ~92 and ?80
Ka; however, ThMAX from 10 halite intervals around 100 Ka
are close to the modern late April to early May values, which
suggests a warm climate but still cooler than the modern at 100
Ka. (4) From 60 to 35 Ka, the climate was cold, with moderate
aridity; saline pans dominated, but there was enough inflow to
accumulate thick salts relatively rapidly and to maintain shal-
low saline lakes approximately 54-50 Ka. Homogenization
temperatures (ThMAX between 23? C and 28? C) are 6?-ll? C
below the modern late April to May values. A relatively low
temperature interval, 10? C below the modern, existed -50 to
-47 Ka. (5) From 35 to 10 Ka, the climate was cold and wet;
plentiful inflow fed a perennial lake that was commonly saline
but at times became up to 90 m deep (Dorn et al., 1990) with
salinities less than 3000 parts per million. Overall, for this 25
Ky period, fluid inclusion ThMAX data suggest brine tempera-
tures were 4?-15? C below modern late April to May values.
During the latest lake phase, from 15 to 11 Ka, ThMAX values
of halites were only 4? C below modern values, whereas
ThMAX values decreased to 7? C below the modern late April to
May temperatures in halites from ~11 to ~10 Ka, signifying
cooler temperatures during this period. (6) From 10 to 0 Ka,
dry arid conditions featured mudflat and/or saline-pan environ-
ments, minor inflow of meteoric waters, and a water table nor-
mally below the surface.
Acknowledgments
The study of core DV93-1 was a collaborative project funded
by the National Science Foundation Global Change Research
Program, grant EAR-9218717. Jianren Li and Chris Brown
(State University of New York at Binghamton) and Sheila Rob-
erts and Wenbo Yang (University of Calgary, under the direc-
tion of Ronald Spencer) studied the core as part of their gradu-
ate research. Teh-Lung Ku and Shangde Luo (University of
Southern California) dated core DV93-1 and shoreline tufas.
Several members of the United States Geological Survey were
involved in various phases of the project: Art Clark, drilling;
Rick Forester, analysis of ostracodes and modern waters, and
interpretation of the paleoclimate record; Platt Bradbury, analy-
sis of diatoms; and Zell Peterman, strontium isotope analyses
of sulfates and carbonates. To all these colleagues, many
thanks!
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Bonneville Basin Lacustrine History:
The Contributions of G.K. Gilbert and Ernst Antevs
Charles G. Oviatt
ABSTRACT
The careful observations and eloquent writings of Grove Karl
Gilbert in the late 1800s laid the framework for most subsequent
studies of lacustrine history in the Bonneville Basin. By the time
of the 1942 symposium on the glacial and postglacial history of
the Great Basin, his work had been largely unchallenged by the
scientific community. Ernst Antevs, who participated in that sym-
posium, reinterpreted some of Gilbert's observations and contrib-
uted both a global perspective and a hypothesis involving
atmospheric circulation to studies of Lake Bonneville and other
Great Basin Quaternary lakes. Fifty years after the 1942 sympo-
sium, the contributions of Gilbert and Antevs are still at the core of
state-of-the-art research on Great Basin Quaternary lakes.
Introduction
Systematic studies of ancient lakes in the Bonneville Basin
began with the work of Grove Karl Gilbert, who published his
monograph on Lake Bonneville in 1890. By the time of the
1942 symposium on the glacial and postglacial history of the
Great Basin, Gilbert's reconstruction of the lake's history had
been reinterpreted by only a few people (Sack, 1989). Ernst
Antevs, who contributed to the 1942 symposium, was one of
those few. In subsequent decades, many new chronologies and
hypotheses about the lake have been proposed. This paper is a
brief review of the knowledge of the Lake Bonneville basin as
of 1942, with emphasis on the contributions of G.K. Gilbert and
Ernst Antevs, and a critique of those contributions from the per-
spective of the late 1990s.
ACKNOWLEDGMENTS.?Ellis Yochelson and an anonymous
reviewer provided useful suggestions for improving the manu-
script. I am grateful to all my friends, colleagues, professors,
Charles G. Oviatt, Department of Geology, Thompson Hall, Kansas
State University, Manhattan, Kansas 66506-3201, USA.
Reviewers: Ellis Yochelson, Department of Paleobiology, National
Museum of Natural History, Smithsonian Institution, Washington,
D.C. 20560-0137, USA; and one anonymous reviewer.
and students, from whom I have learned about Lake Bonneville
and the Great Basin. I dedicate this paper, however, to the one
person who has taught me more than all the others and who
continues to inspire me with his intelligence, curiosity, knowl-
edge, and good humor?Don Currey.
G.K. Gilbert
Lakes began forming in the Bonneville Basin during the
Miocene with the initiation of tectonic extension in the region
(Miller, 1997). Gilbert (1890) noted a few exposures of Tertiary
lake deposits in the Bonneville Basin, but very little was known
about them at the time. Based on the meager evidence he had
available, Gilbert made an insightful comment concerning the
Tertiary history of the basin. He thought it was "possible that
the Bonneville Basin continuously held a lake which, as the
land rose and fell unequally, was expanded and contracted, now
in one direction, now in another" (Gilbert, 1890:101). Later
studies have confirmed this general statement, but a detailed
history of Tertiary (primarily Neogene) lakes is just beginning
to emerge (Miller, 1997; Thompson et al., 1997; Davis, 2002).
Gilbert's great contributions were the careful study of sedi-
mentary deposits and shorelines of Lake Bonneville and the
proposal of hypotheses to explain its history. Gilbert (1890) re-
ported geomorphic and stratigraphic evidence that suggested
Lake Bonneville had risen twice to high levels in the late Pleis-
tocene after a long period of low lakes and alluvial-fan deposi-
tion on piedmont slopes. His conclusions were based on his in-
terpretation of sediments deposited on the bottom of Lake
Bonneville (the Bonneville beds). Two of his study sites were
along the Old River Bed, and one was near the town of Leam-
ington, in west-central Utah (Figure 1). Gilbert defined the
stratigraphic units at these sites by measuring the sediment lay-
ers, which he then traced or correlated from outcrop to outcrop
throughout the basin. Gilbert argued that the shorelines of the
ancient lake and the stratigraphy of the offshore deposits to-
gether tell the story of rising and falling lake levels.
OLD RIVER BED.?The Old River Bed is an abandoned river
valley cut into the Bonneville beds. Gilbert demonstrated that it
121
122 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 1.?Map of Lake Bonneville at its highest stage, the Bonneville shore-
line. SLC = Salt Lake City. Arrows point to the areas at Old River Bed and
Leamington illustrated in Figures 2 and 4, respectively. A heavy line marks the
Old River Bed strait, which is a narrow constriction between Keg Mountain
(KM) to the west and the Simpson Mountains (SM) to the east.
was formed during the regressive (falling) phase of the lake by
a river that flowed from the Sevier Basin in the south to the
Great Salt Lake basin in the north (Figures 1, 2). The exposures
of the Bonneville beds and the shorelines along the Old River
Bed are some of the finest in the basin, and Gilbert took advan-
tage of this by studying them for a total of about two weeks
during several different trips in 1879 and 1880 (Hunt, 1982;
Oviatt and McCoy, 1988).
Gilbert measured and described stratigraphic sections of the
Bonneville beds at two sites at the Old River Bed. The Lower
FIGURE 2.?Map of the Old River Bed in the vicinity of Gilbert's Lower and
Upper River Bed sections.
River Bed section was located approximately where the Over-
land Stage, or Pony Express, Road crossed the Old River Bed
(Figure 2). At this section, Gilbert defined two stratigraphic
units, the Yellow Clay and the White Marl. Because both of these
units are fine grained, he interpreted them as deep-water depos-
its. The Yellow Clay was the thicker of the two and occurred be-
low the White Marl, so he concluded that it was deposited during
a relatively long period of deep-lake conditions, called the Yel-
low Clay Epoch, prior to the White Marl Epoch (Figure 3).
The White Marl was deposited during the most recent deep-
water period and included the time of formation of the promi-
nent Bonneville, Provo, and Stansbury shorelines. A thin sand
unit containing gastropod shells separates the Yellow Clay from
the White Marl. Gilbert interpreted the base of this sand as
marking an unconformity, or period of time when the lake was
below the level of the section and was not depositing sediments
there. In his interpretation, the sand and gastropods were depos-
ited as the lake rose during the White Marl Epoch. Gilbert re-
ferred to the Lower River Bed section as a typical section of the
Bonneville beds, which therefore recorded the major events in
the history of the lake (Figure 3).
Gilbert's Upper River Bed section displayed the Yellow Clay,
the White Marl, and several other stratigraphic units. The Yel-
low Clay and the White Marl are overlain at this section by an
NUMBER 33 123
FEET
1000
850
400
330
YELLOW CLAY\ ALLUVIUM
YELLOW CLAY
YELLOW CLAY
WHITE CLAY ALLUVIUM
EROSION AND
FIRST GRAVEL
UNCONFORM.
WHITE MARL & UROSION AND '
LOWER SAND 1 SECOND GRAVEL/ UPPER SAND
THIRD GRAVEL,
EROSION OF OLD RIVER BED,
AND LATER ALLUVIUM
WHITE MARL EROSION & ALLUVIUM
A B C
BONNEVILLE SHORE
LINE
LEAMINGTON
UPPER RIVER BED
LOWER RIVER BED
WATER OF GREAT
SALTLAKE
FIGURE 3.?Modification of "Diagram of Lake Oscillations inferred from Deposits and Erosions" from Gilbert
(1890, fig. 30). The horizontal axis can be read as time, with a variable scale, or as vertical distance (also with
variable scale) in each of the three stratigraphic sections (Lower River Bed, Upper River Bed, and Leamington).
The White Clay at Leamington is equivalent to the White Marl. Letters A, B, and C represent the three lake high-
stands interpreted by Gilbert (A = Yellow Clay Epoch; B = White Marl Epoch; C = minor regressive-phase fluc-
tuation).
additional sand unit, termed the Upper Sand, which Gilbert con-
sidered as representing a short-lived, minor rise in lake level
(Figure 3).
Of primary importance in Gilbert's interpretation of the Up-
per River Bed section was a gravel unit, which he referred to as
the First Gravel. The First Gravel is more than 15 m thick along
the east side of the exposures at the Upper River Bed section
(Figure 4), but it thins abruptly to the west to less than 1 m, thus
leading Gilbert to refer to the gravel as a "wedge" in his field
notes (Hunt, 1982:128).
Gilbert called the First Gravel a "bar" in his field notes and
described large quantities of tufa cementing the gravel (Hunt,
1982:128), thus indicating that he regarded the gravel as having
been deposited in a lake environment. In United States Geo-
graphical Survey (USGS) Monograph 1, Gilbert (1890:195)
wrote that "the form of the [gravel mass] described, the associ-
ated tufa, and the preponderance of boulders of local derivation,
indicate shore action, but it is possible that an interlacustrine
[that is, between the Yellow Clay and White Marl Epochs] river
was the agent of transportation." References in other places in
his monograph to thick alluvial deposits at this horizon suggest
that by the time Monograph 1 was published Gilbert had de-
cided that the First Gravel at the Upper River Bed section was
alluvial rather that lacustrine gravel, despite the field evidence
he had observed that the gravel had a lacustrine origin.
Regardless of whether the First Gravel was interpreted as al-
luvial or lacustrine, it indicated to Gilbert that the lake fell to a
low level relative to the deep-lake intervals of the Yellow Clay
and White Marl Epochs (Figure 3). The tops of both River Bed
sections represented the final regression of the lake and local
deposition of alluvium.
LEAMINGTON.?The section at Leamington, Utah, is approx-
imately 140 m higher in altitude than the Old River Bed sec-
tions, and it is close to the Bonneville shoreline (Figure 1). At
Leamington, Gilbert noted stratigraphic units that he correlated
with the Yellow Clay and White Marl. They were thicker and
sandier than at the Old River Bed. Because Leamington is lo-
cated at the mouth of the Sevier River where it enters the
Sevier Desert, the river delivered large volumes of fine-grained
sediment (silt, sand) to the lake margin in this area (Figure 5).
Gilbert interpreted gravel between the Yellow Clay and the
White Marl (e.g., White Clay in Figure 3) as alluvium, and he
correlated this gravel with the unconformity at the Lower River
Bed section and with the First Gravel at the Upper River Bed
section. Leamington is too high in altitude to record events af-
ter the lake fell from the Bonneville shoreline (Figure 3).
GILBERT'S CONCLUSIONS.?By combininsg the stratigraphic
evidence from the Old River Bed and Leamington sections with
information from the shorelines and other geologic evidence in
the basin, Gilbert deduced a lake history illustrated in Figure 6.
124 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 4.?Aerial photograph of Gilbert's Upper River Bed section and surrounding area. A = Yellow Clay and
White Marl exposed in the section described by Gilbert (1890); B = lacustrine gravel spit composed of Gilbert's
First Gravel (note the eastward-recurved distal end of the spit; longshore current direction was from the north);
C = eolian sand dunes composed of Gilbert's Upper Sand.
FIGURE 5.?Gilbert's (1890) fig. 28, about which he stated, "Section showing succession of Lacustrine and Allu-
vial Deposits at Leamington, Utah. 1. Paleozoic sandstone. 2. The Yellow Clay (Lower Bonneville). 3. Wedge of
alluvial gravel. 4. The White Marl [White Clay of Figure 3] (Upper Bonneville). 5. Recent alluvial gravel. 6.
Bonneville shore notch, with recent talus."
NUMBER 33 125
FIGURE 6.?Modified from Gilbert's (1890) fig. 34, about which he stated, "The upper and lower horizontal lines
represent the horizons of the Bonneville shore and the surface of Great Salt Lake. Horizontal distances represent
time, counted from left to right. The curve represents the height of the oscillating water surface, and the shaded
area indicates ignorance." Gilbert's curve shows the Yellow Clay Epoch (YC), the White Marl Epoch (WM), the
Provo shoreline (PS), and the regressive-phase fluctuation he deduced from the Upper Sand at the Upper River
Bed section (US).
The Yellow Clay Epoch occurred after a long period of shallow
lakes and alluvial-fan deposition in the basin. By Gilbert's esti-
mate, the Yellow Clay lake did not overflow, and it persisted
much longer than the White Marl lake. Gilbert's interpretation
of the First Gravel at the Upper River Bed section and the
gravel at Leamington as alluvium led him to conclude there had
been a profound lake-level lowering between the Yellow Clay
and the White Marl Epochs (Figure 6). This idea strongly influ-
enced Antevs and most later workers.
According to Gilbert, the White Marl Epoch included the rise
to overflow at Red Rock Pass, Idaho (Figure 1), and the forma-
tion of the Bonneville shoreline. Catastrophic failure of the al-
luvial-fan dam at Red Rock Pass resulted in the Bonneville
Flood and the subsequent formation of the Provo shoreline. A
minor rise in lake level during the regression from the Provo
shoreline at about the altitude of the Upper River Bed section
deposited the Upper Sand; the lake subsequently dropped to
levels near that of the modern Great Salt Lake. Gilbert appar-
ently thought the Stansbury shoreline formed during the regres-
sive phase of Lake Bonneville, that is, later than the formation
of the Provo shoreline, although he did not explicitly state this
in his monograph. This reconstructed history was the starting
point for the work of Ernst Antevs.
Ernst Antevs
During the first half of the twentieth century, Ernst Antevs
published a number of papers concerning Lake Bonneville and
Great Basin Quaternary history. He also contributed to the 1942
symposium. Antevs was convinced that lake fluctuations in the
Bonneville Basin were correlated with growth and decay of
Northern Hemisphere ice sheets and mountain glaciers. Despite
the lack of direct evidence for the ages of either the glacial or
lacustrine deposits, he presented careful arguments for ways in
which they simultaneously varied in size during the late Pleis-
tocene (Figure 7). Evidence from the mouths of Little Cotton-
wood and Bells Canyons in the Wasatch Mountains near Salt
Lake City indicated that Lake Bonneville reached its highest
level, the Bonneville shoreline, shortly after the glaciers in
those canyons reached their maximum down-valley extent (An-
tevs, 1948), so that correlation with mountain glaciers in the Si-
erra Nevada and ice sheets in Canada was likely.
Antevs spent less time in the field examining the stratigraphic
sections at the Old River Bed and Leamington than did Gilbert,
but he reinterpreted Gilbert's descriptions at the three primary
sections, focusing especially on those at the Old River Bed. An-
tevs thought the Yellow Clay Epoch was correlated with an old
period of glaciation in the Rocky Mountains and Sierra Nevada
(Bull Lake and Tahoe glaciations, respectively), and he did not
include it in his discussion of Lake Bonneville history in his
symposium paper (Figure 7) (Antevs, 1948). His interpretation
of the depositional environment of the Yellow Clay as having
been a deep lake was essentially the same as Gilbert's.
Western Ice
(Cordilleran,
Keewatin)
Eastern Ice
(Labrador,
Patricia) o ?
Q. CD
FIGURE 7.?Fluctuations of Lake Bonneville (L.B.), Lake Lahontan (L.L.) in
western Nevada, and Northern Hemisphere ice sheets (modified from Antevs,
1948:171). D.L. = Dendritic Lake; L.P. = Lake Provo; L.S. = Lake Stansbury;
T.L. = Thinolite Lake. Letters B and C represent lake highstands and correla-
tions with Gilbert's interpretations of Lake Bonneville history (see Figure 3).
126 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Antevs regarded the regressive-phase lake rise, which Gilbert
had deduced from the Upper Sand at the Upper River Bed sec-
tion, as a major event in the history of the lake. He suggested
that during this event the lake reoccupied the Provo shoreline
and overflowed for a second time at Red Rock Pass, Idaho. In
addition, Antevs thought the Stansbury shoreline formed during
the interval between the two Provo occupations. He then corre-
lated fluctuations of Lake Bonneville with fluctuations of Lake
Lahontan in western Nevada, and with the Cordilleran and Lau-
rentide ice sheets (Figure 7).
Antevs (1948) argued that during the late Pleistocene the
large North American ice sheets, particularly the Cordilleran
ice sheet, attained a size large enough to deflect the polar front
far to the south of its mean position in historic time. As a conse-
quence, frequent storms carried moisture from the Pacific
Ocean across the Sierra Nevada to the Great Basin to create the
large pluvial lakes. Thus, the filling and emptying of the lake
basins were correlated with the expansion and contraction of
the Northern Hemisphere ice sheets. He also thought that the
astronomical hypothesis of climate change formulated by Milu-
tin Milankovitch was probably valid (Antevs, 1938). Added to
this was his fundamental belief that major climatic events had
to be global in extent (e.g., Antevs, 1955). Accordingly, he was
eager to find correlative climatic events across North America
and in northern Europe.
Late Twentieth Century Views
Machette and Scott (1988) and Sack (1989) have reviewed
how interpretations of Lake Bonneville have changed during the
second half of the twentieth century. In this paper, I focus only
on current interpretations of the work of Gilbert and Antevs.
With a few possible exceptions, all the lake sediments ob-
served by Gilbert, including the Yellow Clay and the White
Marl, are now recognized to have been deposited during a sin-
gle major lake cycle between about 30 thousand years (Ky) be-
fore present (B.P.) and 10 Ky B.p. (Scott et al., 1983; Oviatt,
1987; Oviatt et al., 1994). This is the Bonneville lake cycle, or
Lake Bonneville. Some Pleistocene lake cycles older than 30
Ky B.P. have been given separate names (e.g., Cutler Dam, Lit-
tle Valley, Pokes Point; Figure 8).
Massive shoreline deposits, erosional shorelines, deltas, and
offshore sediments are prominent in the Bonneville Basin.
Most of these were almost certainly produced during the
Bonneville lake cycle. In a refinement of Gilbert's initial work,
prominent shorelines named and mapped by him have been
dated and linked to the offshore deposits, and an accurate radio-
carbon chronology has been established that can now be used in
many ways as a predictive model (Figure 9) (Oviatt et al., 1992;
Oviatt, 1997). The major shorelines have been mapped with
precision (Currey, 1982), and Currey (1990) has shown how the
extensive knowledge about the Bonneville Basin serves as a
prime example of the landforms and processes typical of semi-
desert basins.
In addition to the accumulation of abundant numerical-age
data and numerous other advances in understanding the Bonne-
ville lake cycle, Gilbert's sections at the Old River Bed and
Leamington have been reinterpreted. The Yellow Clay is now
recognized to be a shallow lacustrine deposit; it consists essen-
tially of fine-grained deltaic sediment deposited early in the rise
of Lake Bonneville (Oviatt, 1987). During deposition of the
Yellow Clay the Sevier River flowed across the flat floor of the
Sevier Desert, through the Old River Bed strait, and transported
sand, silt, and clay to the margin of the rising lake in the Great
Salt Lake basin ("Main Body" in Figure 1). In light of this in-
terpretation, Gilbert's dilemma with explaining why the Yellow
Clay is so much thicker and contains so much more silt and
sand than the White Marl (Gilbert, 1890:200-209), is re-
solved?the Yellow Clay and the White Marl were deposited in
different depositional environments and should not be expected
to look similar. Although deposits similar to the Yellow Clay
are present at many localities in the basin, they represent dis-
connected shallow-water depositional environments near the
mouths of large rivers rather than a single, widespread, deep-
water lake deposit. In contrast, the White Marl is the basin-
wide stratigraphic unit that records the deep-water, or offshore,
history of Lake Bonneville (Oviatt, 1997).
60
to
8 408
ro 0
A
CO
? 8
as
A
100 200 300 400
age (ka)
500 600 700
FIGURE 8.?Schematic diagram of middle and late Pleistocene lake cycles in the Bonneville Basin. Modified
from Machette and Scott (1988, fig. 3) and Oviatt and Miller (1997, fig. 2).
NUMBER 33
1600
.? 1500
cg
CO
?. 1400
CDTJ
0
# 1300
CO
127
1200
lake level
o carbonate
? wood
? organics
A basaltic ash
Stansbury
Pilot Valley
U2U3 ri Bonnevillec
Gunnison
Gilbert
35 30 25 20 15
age (ka)
10
FIGURE 9.?Chronology of Lake Bonneville fluctuations, modi-
fied from Oviatt et al. (1992), Oviatt (1997), and Oviatt and
Miller (1997). Elevations have been adjusted for isostatic
rebound. Ul, U2, and U3 are transgressive-phase fluctuations
(Oviatt, 1997). Radiocarbon ages were obtained for dissemi-
nated organic carbon, wood or charcoal, carbonate (shell or
tufa), and basaltic ashes.
Figure 4 is an aerial photograph of the area around Gilbert's
Upper River Bed section. It shows the prominent recurved
lacustrine spit composed of Gilbert's First Gravel, which he had
decided might be alluvium. The spit was deposited by long-
shore currents in Lake Bonneville after deposition of the Yel-
low Clay as the lake began to rise, following formation of the
Stansbury shoreline (Figure 9); it does not represent river depo-
sition during a lowstand between two high lake stands, as Gil-
bert suggested in his monograph. At the Lower River Bed sec-
tion the unconformity between the Yellow Clay and the White
Marl, which Gilbert and Antevs interpreted as representing a
long period of erosion and alluvial deposition, represents a
short period (possibly less than 500 years) when the lake level
fell about 45 m and the Stansbury shoreline formed (Oviatt,
1987; Oviatt et al., 1990). The gravel wedge between the Yel-
low Clay and the White Marl at Leamington, which Gilbert in-
terpreted as alluvium, is now interpreted as lacustrine gravel
deposited during a fluctuation in lake level during the trans-
gressive phase of the lake (U2 in Figure 9; Oviatt et al., 1994).
Figure 4 also shows well-developed eolian sand dunes com-
posed of Gilbert's Upper Sand. Although Gilbert recognized
that the surface of this sand "may have gained or lost [thick-
ness] by the action of the wind, which throws its surface into
waves, and has caused it to bury at the north the exposure of the
lower formations" (Gilbert, 1890:196), he thought that the
lower part of the sand was lacustrine. Because no features of
lacustrine origin have been found in the sand, his interpretation
of a minor lake rise, and Antevs' suggestion of a second trans-
gression to the Provo shoreline, are unlikely.
Despite these changes in interpretation of Gilbert's strati-
graphic sections, which are the results of building on his work
and the work of others and the availability of superior technol-
ogy for field and laboratory investigations, Gilbert laid the
groundwork for all subsequent research on Lake Bonneville.
Most of his observations from his horseback trips into the
desert were highly accurate, even if some of his interpretations
have been replaced by alternative hypotheses. It is clear that he
thoroughly understood the lake beyond the context of his own
century, and many of his conclusions have not been improved
upon even after 100 years of subsequent research. One of Gil-
bert's many long-lasting contributions to science was to dem-
onstrate that paleoclimatic information can be obtained from
studies of ancient lake deposits.
Antevs also made his mark on Great Basin paleoclimate re-
search. Modern research has demonstrated empirically the gen-
eral validity of Antevs' conviction that the major climate
changes that resulted in large pluvial lakes were global in
scope. Although the details of his correlations with Lake
Lahontan and with glacial events have not survived scientific
tests, the urge to find global connections persists today. With
refinements in glacial chronologies and modeling, including the
realization that the Laurentide ice sheet was even more influen-
tial than he thought, Antevs' general hypothesis that ice sheets
diverted storm tracks into the Great Basin to feed the large plu-
vial lakes is well supported by computer simulations and field
data (COHMAP members, 1988; Kutzbach et al., 1993;
Thompson et al., 1993).
Future Work
Since the days of Gilbert and Antevs, scientific advance-
ments have resulted in major changes in interpretations, but the
solid foundation laid by Gilbert and the refinements made by
Antevs half a century later are still at the heart of our knowl-
edge of Lake Bonneville and late Pleistocene climate change in
the Great Basin (e.g., Oviatt, 1997; Benson, 1997). Despite
many advances in understanding the history of Lake Bonneville
and Great Salt Lake, a number of important questions remain
only partially answered. These include questions about the tim-
ing and duration of regression from the Provo shoreline and the
128 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
age of the Provo shoreline itself; the timing and character of the
initial rise of Lake Bonneville between about 30 and 20 Ky
B.P.; the chronology of fluctuations of Great Salt Lake during
the last 10 Ky; the precise relationships between Lake Bonne-
ville and Great Salt Lake and climate, including the relative im-
portance of temperature and precipitation or other variables in
determining lake size; and the details of pre-Bonneville lake
cycles and their connections to global climate. These questions
and others are being pursued in the Bonneville Basin, and,
thanks in part to the powerful early contributions of Gilbert and
Antevs, reliable answers are likely to be forthcoming in the
near future.
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1938. Climatic Variations during the Last Glaciation in North America.
Bulletin of the American Meteorological Society, 19:12-16.
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Utah, 38(20): 167-191.
1955. Geologic-Climatic Dating in the West. American Antiquity, 20:317-
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Benson, L.V.
1997. Keynote Address: Global Change Modulation of Great Basin
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1993. Climatic Changes in the Western United States since 18,000 yr B.P.
In H.E. Wright, Jr., J.E. Kutzbach, T. Webb III, W.F. Ruddiman, F.A.
Street-Perrott, and P.J. Bartlein, editors, Global Climates since the
Last Glacial Maximum, pages 18-153. Minneapolis: University of
Minnesota Press.
Fluvial Linkages in Lake Bonneville Subbasin Integration
Dorothy Sack
"... and so lakes developed in nearly all the basins. Of these lakes the great majority over-
flowed into adjacent basins ..."
E.Blackwelder (1948:12)
ABSTRACT
Several late Pleistocene lakes in the Great Basin, including Lake
Bonneville, consisted of component subbasins that were united
into a single large lake only when water levels were high enough
to breach the subbasins' thresholds. During the last deep-lake
cycle, but before and after their period of integration with the
major water body, isolated closed-basin lakes existed in many of
those subbasins. This study focuses on two of Lake Bonneville's
seven major subbasins, Puddle Valley and Tule Valley, because
they display distinctive geomorphic and sedimentologic evidence
of the subbasin integration process. In both cases, Lake Bonneville
burst through material at the dividing threshold, and water poured
into the adjoining subbasin as a fluvial current of high discharge.
Where the jet spread out onto the subbasin's piedmont, it deposited
a gravel expansion bar over which gravel was transported in
large-scale subaqueous bedforms. Fluvial braid bars in their initial
stage of development provide a good analogue for the general
form of the inflow bars and help identify the fluvial processes of
inflow-bar formation. Details of the geomorphology and sedimen-
tology of the inflow bars reflect the site-specific conditions under
which they formed. The ability to recognize inflow bars in the geo-
morphic and stratigraphic record will help researchers identify
subbasin integration events and will, therefore, contribute to
improved accuracy in reconstructing the geomorphic and paleocli-
matic history of other paleolake basins.
Introduction
For more than a century, scientists have periodically pub-
lished comprehensive inventories of late Pleistocene lakes in
western North America (e.g., Russell, 1885, 1896; Meinzer,
Dorothy Sack, Department of Geography, 122 Clippinger Labs, Ohio
University, Athens, Ohio 45701, USA.
Reviewers: B.G. Bills, National Aeronautics and Space Administra-
tion, Goddard Space Flight Center, Greenbelt, Maryland 20771, USA;
R.M. McCoy, University of Utah, Salt Lake City, Utah 84112, USA.
1922; Hubbs and Miller, 1948; Snyder et al, 1964; Mifflin and
Wheat, 1979; Williams and Bedinger, 1984). Although these in-
ventories have been compiled for various specific physio-
graphic or political boundaries, each included all or a large part
of the Great Basin. Information provided about Great Basin
paleolakes in that literature ranged from their approximate loca-
tions and relative extent, as delineated on small-scale maps; to
measurements of various lake attributes, such as depth and area;
to observations regarding possible overflow episodes for the
predominantly closed-basin lakes. Not surprisingly, those con-
tributions emphasized the maximum extent of the paleolakes
and any associated overflow that made the lakes tributary to
lower closed-basin lakes, or to the sea, via long, connecting
stream courses. The emphasis on maximum lake size, however,
obscured the fact that, like most large modern lakes (Sly, 1978),
several late Pleistocene lakes in the Great Basin were assem-
blages of subbasins united into a single lake only during periods
of relatively high water levels (Table 1). Many of the lacustrine
TABLE 1.?Examples of Great Basin paleolakes that consisted of component
subbasins.
Paleolake
Lake Bonneville
Lake Lahontan
Lake Chewaucan
Fort Rock Lake
Lake Alkali
Lake Warner
Lake Malheur
Lake Alvord
Lake Surprise
Lake Searles
Lake Panamint
Little Mojave Lake
Lake Mojave
Location
UT, NV, ID
NV
OR
OR
OR
OR
OR
OR
CA,NV
CA
CA
CA
CA
Open-basin
interval
Yes
No
No
?
No
?
No
?
No
No*
No*
No
No
Reference
Gilbert, 1890
Russell, 1885
Waring, 1908; Allison, 1945
Waring, 1908; Allison, 1979
Waring, 1908
Waring, 1908, 1909; Weide, 1974
Russell, 1903; Waring, 1909
Waring, 1909
Russell, 1884
Gale, 1915
Smith, 1976
Thompson, 1929
Thompson, 1929
"Overflowed to another closed-basin lake by means of a long fluvial connec-
tion.
129
130 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
subbasins that constituted arms or embayments of a major lake
at its higher levels contained independent closed-basin lakes
before and after their interval of integration.
In order to reconstruct accurate paleohydrologic and paleo-
climatic models for such a region of closed-basin lakes, it is
important to distinguish those water-level fluctuations con-
trolled primarily by geomorphic factors from those induced
more directly by climatic variables. A shoreline that is well
represented in the geomorphic or stratigraphic record, for ex-
ample, is not necessarily the result of static climatic condi-
tions; nor does a change in lake level always indicate a climati-
cally induced shift in effective moisture (Sack, 1994c). In
addition to climatic factors that affect lake hydrology, geomor-
phic factors influencing lake hydrography are also reflected in
the lake's shoreline record. Variations in lake size that create
multiple shorelines may be the result of changes in drainage
basin area caused by such factors as tectonism, volcanism,
mass movement, stream diversion, or overflow of upstream
lake basins. If the geomorphically induced water-level change
is rapid, a vertical zone mostly devoid of shoreline evidence,
i.e., a null zone, may result (Sack, 1990, 1995). It is also possi-
ble that a well-developed shoreline in a closed-basin lake
marks the elevation of an intrabasin threshold rather than an in-
terval of climatic equilibrium (Eardley et al., 1957; Benson,
1978; Currey and Oviatt, 1985; Benson and Paillet, 1989). For
example, when, under conditions of increasing effective mois-
ture, the water level in one subbasin reaches the elevation of an
intra-basin sill, it may remain essentially at that elevation until
overflow fills the adjoining closed basin to the threshold level
and a single integrated closed-basin lake is formed. The coa-
lescing of subbasin lakes into a single water body by succes-
sive overflows from one subbasin to another can create in the
confluent lake basin a complex set of shoreline evidence and
can result in the amalgamated lake basin having not just a sin-
gle lake-level chronology but, rather, a complicated assem-
blage of subbasin chronologies (Davis, 1982). Properly allo-
cating the contributions to paleolake-level changes influenced
primarily by geomorphic factors versus those responding di-
rectly to climatic variables leads to a more accurate reconstruc-
tion of a paleolake basin's climatic and geomorphic history.
At its maximum extent, Lake Bonneville was the largest of
the standing water bodies that accumulated in the Great Basin
during the late Pleistocene (Figure 1). The Bonneville lacustral
cycle lasted from about 30 thousand years ago (Ka) to 12 Ka
(Figure 2) (Oviatt et al., 1992), but considerable shoreline and
sediment evidence of it remains visible in the landscape (Sack,
1992). The white marl, for example, is a distinctive Lake
Bonneville marker bed that was deposited widely around the
basin during the deeper phases of Lake Bonneville spanning an
interval of at least 8 thousand years (Ky) (Gilbert, 1890; Oviatt,
1987). The most prominent of the relict Lake Bonneville shore-
lines are the Stansbury, Bonneville, and Provo. The Stansbury
shoreline was formed between about 22 and 20 Ka, during the
transgressive phase of Lake Bonneville (Currey et al., 1983;
-45?N
-40 >^
-35? /
? Great Basin
? Lake Bonneville
D Great Salt Lake
0 300 km
120?
i
M
115? 110? W
I l
FIGURE 1.?Map showing Lake Bonneville at its greatest extent relative to the
Great Basin.
1550
1450
1350
1250
PS
- BS
BF
SS
-
^??-r
Provo shoreline
Bonneville shoreline
Bonneville flood
Stansbury shoreline ,
SS /
J
B5
/
RF
PS
K
\
30 25 20
Age (Ka)
15 10
FIGURE 2.?Generalized hydrograph of Lake Bonneville (after Burr and Cur-
rey, 1988; Oviatt etal., 1992).
Oviatt et al., 1990). After Stansbury shoreline time, the lake
continued rising until it reached the lowest point on its drainage
divide, perhaps as late as 15 Ka (Oviatt et al., 1992). With its
upper limit constrained by this external threshold, the lake con-
structed the Bonneville shoreline (Gilbert, 1890). Bonneville
shoreline time ended about 14.5 Ka when failure of material at
the threshold triggered the Bonneville Flood (Gilbert, 1890;
Malde, 1968; Oviatt et al., 1992; O'Connor, 1993). The lake re-
NUMBER 33 131
stabilized at the level of a bedrock sill about 110 m below the
Bonneville shoreline, and the conspicuous Provo shoreline was
created during a period of a few hundred years while the lake
was controlled by this lower external threshold (Gilbert, 1890;
Burr and Currey, 1988; Oviatt et al., 1992). Closed-basin con-
ditions finally resumed when Lake Bonneville regressed from
the Provo shoreline because of climatic reasons, with the lake
cycle ending roughly 2 Ky later (Oviatt et al., 1992).
The Bonneville lake basin consists of seven major closed
subbasins each of which was integrated into the megalake for a
different period of time (Figure 3; Table 2). Lake Bonneville
originated in the lowest subbasin, that of Great Salt Lake. Rush
Valley was part of the Great Salt Lake basin at the start of the
Bonneville lacustral cycle but was isolated from it after Bonne-
ville shoreline time because of the barricading effect of the
Stockton Bar complex, which Lake Bonneville had built across
their connecting strait (Gilbert, 1890). Integration of the Great
TABLE 2.?The major Lake Bonneville subbasins and selected characteristics.
Subbasin
Great Salt Lake
Great Salt Lake Desert
Puddle Valley
Sevier Desert
Tule Valley
Cedar Valley
Rush Valley
Approximate
modern threshold
elevation (m)
Lowest basin
1285
1362
1399
1446
1507
1585
Mode of integration
Origination basin
Spillover from Lake Bonneville
Spillover from Lake Bonneville
Overflow to Lake Bonneville
Spillover from Lake Bonneville
Overflow to Lake Bonneville (?)
Origination basin, isolated later
Salt Lake Desert, Puddle Valley, and Tule Valley occurred
when transgressing Lake Bonneville spilled over successively
higher thresholds into those adjacent subbasins (Currey, 1980;
Currey and Oviatt, 1985; Sack, 1990). Although Currey and
Oviatt (1985) maintained that Cedar Valley was also integrated
by spillover from Lake Bonneville, overflow may actually have
113?W 112?W
-42?N Idaho
to Utah
Puddle
Valley Great Salt
\p Lake _Wyqming
Salt Lake City
Great Salt
Lake Desert
?: /
?-Rush ? / .?
Valley / ',
CedarValley
Sevier
Desert
Maximum Extent of
Lake Bonneville
60 km
40 mi
Tule
Valley
38?N-
FlGURE 3.?Map showing the seven major subbasins of Lake Bonneville.
132 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
been from the high Cedar Valley lake basin, which has a rela-
tively small capacity, into Lake Bonneville. On the other hand,
it seems clear that the Sevier Desert basin initially drained
northward over the lowest point on its divide into Lake Bonne-
ville by means of a river channel tens of kilometers long (Cur-
rey et al., 1983; Oviatt and McCoy, 1988). Lake Bonneville
eventually transgressed well above that threshold, converting
the connecting river bed into a strait and the Sevier Desert trib-
utary basin into its Sevier Desert arm.
Regardless of the direction of overflow, spillover from one
subbasin to an adjacent one is essentially a fluvial process.
Even if both basins initially contain independent closed lakes,
differences in such variables as elevation, relief, hypsometry,
precipitation, surface runoff, and ground-water budgets will
cause the water bodies to rise at different rates (Benson, 1978;
Benson and Mifflin, 1986). The first one to overtop their con-
necting sill will discharge a current of water flowing down-
slope by the force of gravity toward the base level of the re-
ceiving basin. Depending upon specific geomorphic and
hydrologic circumstances, this fluvial current may occupy a
long or short channel, flow on a continuous or intermittent ba-
sis, and have a total existence of brief or extensive duration.
In both the Puddle Valley and Tule Valley subbasins of Lake
Bonneville, a well-exposed example of a distinctive deposi-
tional landform created during integrating inflow is found adja-
cent to the subbasin threshold (Currey, 1980; Currey and
Oviatt, 1985; Sack, 1990). Combined with the results of valley-
wide investigations, a comprehensive analysis of the geomor-
phology, sedimentology, and morphostratigraphy of these fea-
tures, referred to in this paper as inflow bars, provides insights
into the nature of the subbasin integration process and the tim-
ing of Lake Bonneville's integration of the two valleys. Results
of this investigation, moreover, are useful beyond the Bonne-
ville Basin. Recognizing similar evidence in the geomorphic
and stratigraphic record of other Quaternary and perhaps even
pre-Quaternary lake basins will help researchers to identify
subbasin integration events and, thereby, to reconstruct more
accurately the geomorphology, chronology, and paleoclimate
of other paleolake basins.
Previous Work
The fact that climate is not the only variable affecting the
size of closed-basin lakes has been recognized in several previ-
ous paleolake studies. More than a century ago Gilbert
(1890:262) emphasized the significance of "topographic con-
figuration" as well as "meteorologic conditions" with regard to
changing lake levels. Stream diversion is probably the most
frequently mentioned nonclimatic influence on closed-basin
lake levels (e.g., Gilbert, 1890; Bright, 1963; King, 1978;
Davis, 1982; Benson and Paillet, 1989). In their report on late
Pleistocene lakes in Nevada, however, Mifflin and Wheat
(1979) stressed the importance of recognizing episodes of ex-
ternal drainage in a lake's history because pluvial hydrologic
indexes calculated for shorelines controlled by external thresh-
olds are only minimum estimates, which limits their usefulness
in making interbasin hydrologic comparisons. Smith and
Street-Perrott (1983:192) noted that the size, shape, and hydro-
logic regime of a lake basin, and therefore the position of its
water plane, may change as the result of various nonclimatic
processes, including tectonism, volcanism, landsliding, sedi-
ment infilling, stream capture, and the cutting of an external
outlet. Such alterations can make interpreting the paleoclimatic
signal from closed-basin lakes especially challenging
(Street-Perrott and Harrison, 1985).
The Great Basin paleolake literature contains some discus-
sion of the nature of constituent subbasins and intrabasin sills
and their role in lake history. The second largest late Pleis-
tocene lake in the Great Basin, Lake Lahontan, has thus far
been studied more intensively from the subbasin perspective
than any other Great Basin paleolake, probably because even
when portrayed at its largest extent it maintains the strong ap-
pearance of discrete water bodies separated by narrow straits.
The importance of the subbasin nature of Lake Lahontan was
first identified by Russell (1885), who recognized that various
component basins would have been united or isolated at differ-
ent lake levels. In another early work, Jones (1925:37) ob-
served that the Lake Lahontan intrabasin sills consist of allu-
vium, and he examined precipitation differences among the
"subsidiary basins." Later, those subbasin lakes that would ex-
pand mainly by direct inflow from perennial tributary streams
were distinguished from those without perennial streams,
which would rise instead by spillover from adjoining subbasins
(Benson, 1978). More recently, Benson and Paillet (1989) used
seven major Lake Lahontan subbasins to demonstrate the influ-
ence of topography on lake level. They concluded that total
lake-surface area is a better indicator of climate change than
lake level or volume. Since the late 1970s, subbasin events and
sill elevations have figured prominently in efforts to recon-
struct the chronology of Lake Lahontan (Benson, 1978, 1991;
King, 1978; Davis, 1982, 1983; Thompson et al., 1986; Benson
and Thompson, 1987; Benson et al., 1992). It was in this period
that Davis (1982:54) summarized so well the importance of
component basins to reconstructing the chronology of a paleo-
lake: "there is no 'Lahontan chronology'; there is a complex
family of chronologies, which are fewer when the water is
high, and more numerous when it is low."
Gilbert (1890) briefly mentioned constituent subbasins and
intrabasin sills in his extensive monograph on Lake Bonneville.
Gilbert (1890:20) observed that, because of coalescence of ex-
panding lakes, the number of independent endoreic drainage
basins in the entire Great Basin decreased during the late Pleis-
tocene lake epoch, with Lake Bonneville becoming the largest
of the "confluent lakes." Desiccation at the end of the lake cy-
cle left the Lake Bonneville basin redivided into numerous, in-
dependent closed basins (Gilbert, 1890). According to Gilbert
(1890:186), the strongly developed geomorphic characteristics
of one of the lake's most conspicuous closed-basin-phase
NUMBER 33 133
shorelines, the Stansbury, suggest that it was controlled "by an
outflow or its equivalent" rather than by climatic equilibrium.
He considered the possibility that sill-controlled inflow to a
component subbasin "might have served as a regulator, causing
the oscillating lake to linger at a particular level each time it
rose" (Gilbert, 1890:186-187).
In a 1957 paper, Eardley et al. (pages 1155-1164) presented
some hypotheses regarding Lake Bonneville subbasins and
shorelines. Those authors emphasized the notion that the ex-
pansion of Lake Bonneville into a subbasin increased the lake's
surface area, resulting in greater evaporation from the lake. If
precipitation surplus over evaporation was modest during the
lake's transgressive phase, which they believed was generally
the case, expansion in surface area by subbasin integration
would create a temporary balance between precipitation and
evaporation, and a water-level stillstand would result. The
longest stillstand, and hence the best shoreline evidence, would
be associated with the largest and most arid receiving basins.
During the regressive phase, with precipitation slightly less
than evaporation, abandonment of a subbasin would cause a
sudden decrease in evaporation from the lake surface. The re-
sulting temporary balance between precipitation and evapora-
tion would again cause a stillstand in lake level. Eardley et al.
(1957) used sill and shoreline elevation data to reject Gilbert's
(1890) suggestion that the Stansbury shoreline of Lake Bonne-
ville was controlled by an intrabasin sill. On the other hand,
they hypothesized that the stillstand that created the Gilbert
shoreline of Great Salt Lake was caused by Great Salt Lake
spilling over its threshold into the Great Salt Lake Desert.
More recent publications that mention Lake Bonneville sub-
basins and interior sills identify the landforms created by in-
flow. In a study of the coastal geomorphology near Great Salt
Lake, Currey (1980) suggested that during the regressive phase
of the last deep-lake cycle inflow from the desiccating Lake
Bonneville briefly maintained a lake in the Puddle Valley sub-
basin. According to Currey (1980:76), a low-level shoreline
marks the small Lake Puddle, and a "lobate, steep-fronted
delta" indicates the point of high-velocity inflow. A subsequent
analysis of shoreline evidence, however, describes rapid filling
of Puddle Valley by Lake Bonneville inflow during the trans-
gressive phase of the lake cycle and concludes that the
low-level Puddle Valley shoreline was formed by an indepen-
dent Lake Puddle that existed during the regressive phase of
the lake cycle after the subbasin lost its surface connection with
Lake Bonneville (Sack, 1995). Shoreline evidence supporting
the existence of an isolated Lake Puddle during the transgres-
sive phase of the last lake cycle, before Lake Bonneville inte-
gration, has not been reported.
In their outline of the fluctuations of Lake Bonneville, Cur-
rey and Oviatt (1985) included probable transgressive-phase
stillstands caused by the lake spilling into the Cedar Valley and
Tule Valley subbasins. Those authors attributed a scoured zone
close to the Cedar Valley threshold to high-velocity inflow, but
they were unable to identify a shoreline in Lake Bonneville that
resulted from that inflow event. Currey and Oviatt (1985:11)
found inflow into Tule Valley to be "clearly marked by a
bed-load delta," and they tentatively assigned a particular, un-
named Lake Bonneville shoreline to the associated stillstand.
Sack (1990, 1994a, 1994b) later suggested that the Tule Valley
inflow "delta" may be a compound feature, and she attributed
the 29 m high Tule Valley null zone, a basinwide sector of the
piedmont that lacks substantial shoreline evidence, to rapid in-
flow from Lake Bonneville. She also reported on the existence
during the last lake cycle of an isolated Lake Tule both before
and after the subbasin's interval of integration with Lake
Bonneville.
Data and Interpretation
PUDDLE VALLEY
GENERAL SETTING.?Puddle Valley is located about 100 km
west of Salt Lake City, Utah (Figure 3). It is bordered by the
Lakeside Mountains to the east, Cedar Mountains to the south,
Grassy Mountains to the west, and by smaller ridges and knolls
to the north (Figure 4). The 400 km2 closed drainage basin is
entirely surrounded by (inset into) the Bonneville Basin. Spe-
cifically, it adjoins the Great Salt Lake subbasin, which lies to
threshold
and
inflow area
subbasin
divide
FIGURE 4.?Regional map of Puddle Valley.
134 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
the north, east, and southeast, and the Great Salt Lake Desert
subbasin, which lies to the west and southwest. Elevations in
the subbasin range from 1316 m on the valley floor to 2019 m
at the highest peak in the Lakeside Mountains. There is no nat-
urally occurring perennial or intermittent surface water in Pud-
dle Valley today.
The lowest point on the Puddle Valley drainage divide, i.e.,
its threshold, lies in an unnamed pass in the northeastern part of
the valley (Figure 4). A paved highway to Lakeside, Utah, ex-
tends through the pass. The pass has an approximate elevation
of 1362 m, which is below the modern local elevations of the
Bonneville (1625 m), Provo (1501 m), and Stansbury (1380 m)
shorelines and above the post-Bonneville Lake Puddle level
(1338 m) (Currey, 1982; Sack, 1995). On the Puddle Valley
(south) side of the threshold, the pass area consists of a north-
south-trending, trough-shaped valley, which is about 1.2 km
long and 0.5 km wide (Figure 5). The sides of the trough con-
ca\vy ^nave\>een Teworkeo. into alluvial fans; eolian, fluvial, and
FIGURE 5.?Aerial photograph of the Puddle Valley threshold area, including
the unnamed pass (P), Lake Bonneville shoreline deposits (S), the trough-like
depression (T), and the inflow bar (IB).
alluvial-fan sediments dominate the floor of the trough. The
floor of the confined trough slopes about 0.7? toward Puddle
Valley, but the modern topography is locally irregular primarily
because of eolian action.
INTEGRATION INTERVAL.?The approximate duration of
Puddle Valley's integration into the Bonneville lacustrine sys-
tem can be estimated by plotting the rebound-corrected eleva-
tion of the Puddle Valley threshold on a detailed hydrograph of
Lake Bonneville. The following formula is used to convert
modern Bonneville Basin elevations to their values prior to hy-
dro-isostatic rebound (Currey and Oviatt, 1985):
Za = Zr-{Zr~l200\zb-\552),(Zfc-1200) (Equation 1)
where Za is the adjusted rebound-free (estimated original) ele-
vation, Zr is a given rebounded (present) elevation, Zb is the
local modern elevation of the Bonneville shoreline, 1552 is
1200 is an empirical constant (all elevations are in meters).
For the Puddle Valley threshold, 1362 m and 1625 m are used
m Equation 1 for Zr and Zb, respectively. Plotting the result-
ing Za (1334 m) on the Lake Bonneville hydrograph of Oviatt
et al. (1992) yielded an estimate of 25.8 to 12.2 Ka for the in-
terval of Puddle Valley integration (Figure 6). The latter value
is a minimum-limiting age estimate because some hydro-iso-
static rebound had already occurred by this stage of Lake
Bonneville (Currey and Burr, 1988); the elevation of the pass
at the time of re-isolation would have been somewhat higher
than 1334 m.
GEOMORPHOLOGY AND SEDIMENTOLOGY OF THE INFLOW
BAR.?The depositional landform that marks the inflow of
Lake Bonneville into Puddle Valley (Currey, 1980) lies imme-
diately south of the Puddle Valley threshold area, just beyond
the southern end of the trough-shaped valley (Figures 4, 5, 7).
Field work and the interpretation of aerial photographs reveal
that the inflow landform is a large-scale, tongue-shaped bar on
1550
1450
??g 1350
UJ
PS Provo shoreline
BS Bonneville shoreline
BF Bonneville flood
SS Stansbury shoreline
30 10
FIGURE 6.-
Bonneville.
-Estimated interval of Puddle Valley's integration with Lake
NUMBER 33 135
which are found linguiod, flow-transverse bedforms, i.e.,
three-dimensional subaqueous dunes (Figure 8) (Ashley,
1990). Both the bar and the subaqueous dunes are composed of
gravel. The feature ranges in elevation from 1341 m at the dis-
tal (southern) base of the bar to 1357 m at the crest of the high-
est bedform. The bar, which is about 1.5 km long and up to 1.2
km wide, consists of a stoss-side pan bordered by a continuous
lateral and lee-side rim. Along the lee side the rim has a maxi-
mum lakeward slope of 15?. The bar complex is mantled with
Lake Bonneville marl, clastic lacustrine fine-grained sedi-
ments, and postlacustrine eolian sandy silt. The maximum
measured thickness of this post-inflow sequence is 1.9 m. Over
the ridges of the linguoid bedforms, where the post-inflow se-
quence is thinner, the surface is gravelly. Postlacustrine eolian
fines thicken dramatically toward the northeastern portion of
the feature; yellowish gray eolian sandy silt found along the
northern end of the left lateral rim is at least 1.0 m thick.
A measured cross section (Figure 9) shows that the three-
dimensional dunes are moderately asymmetrical and com-
A-A' location of cross section
I I bedforms
swales
bedform crest
0 200 400 600 m
FIGURE 7.?Map-view diagram of the Puddle Valley inflow bar, showing line
of measured section, A-A'.
FIGURE 8.?Large-scale subaqueous gravel dunes overlain by light-colored
lacustrine and eolian fines exposed in a road cut along the measured section in
Puddle Valley.
pound, with medium-scale dunes superposed on large-scale
dunes, which, in turn, are superposed on very large dunes (Ash-
ley, 1990). Measured lengths (L) range from 6.1 to 143.3 m,
heights (H) from 0.2 to 3.2 m, and the vertical-form index {LI
H) from 17.8 to 62.2. The equation that expresses the height
versus length relationship of the Puddle Valley subaqueous
dunes,
H = 0.0445 L0.90 (Equation 2)
B1353 m
n
1351 m-
1349 m-
[ | Post-inflow fines
[ ? :'."l Inflow gravel B'
V.E.=10x
20 40 60 80 100 120 m
FIGURE 9.?Measured cross section through gravel dunes on the Puddle Valley
inflow bar. Location of measured section is shown on Figure 7. (V.E. = vertical
exaggeration.)
136 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 3.?Metric height (H) vs
dunes.
Equation
H = 0.074L0-77
H = 0.0052L1 5
H = 0.0677L0 898
H = 0.0445L090
H = 0.0409L113
. length (L) relationships for subaqueous
Source
Allen (1968)
Baker(1973)
Ashley (1990)
Puddle Valley (this paper)
Tule Valley (this paper)
is similar to those derived from "giant current ripples" by Allen
(1968) and Baker (1973:50) and is close to the equation deter-
mined from almost 1500 bedforms, reported in Ashley (1990)
(Table 3).
The subaqueous dunes on the Puddle Valley inflow bar con-
sist of cross beds that, at the measured section, dip 9? to 24? to
the east-southeast. Orientation data collected from the 10 larg-
est clasts at 14 different sediment-sample sites along A-A' also
show a current direction to the east-southeast (Figure 10). This
direction of dip for beds and clasts reflects the position of the
line of section to the east of the main axis of flow (Figure 7).
East-southeast is perpendicular to the crestal trend of the sub-
aqueous dunes along A-A'.
Sampled clasts were predominantly rounded or subrounded,
and the 10 largest had an average longest axis (A-axis) of 19
cm. The overall 10 largest clasts observed anywhere along the
length of the cross section had a mean A-axis of 35 cm. In ad-
dition to the local Paleozoic limestones, dolomites, cherts, and
sandstones, gravel-size clasts of tufa-cemented conglomeratic
beachrock, tufa, and laminated marl were noted in the sample
pits and elsewhere along the cross section. A gravel-size rip-up
clast of marl that was collected from one of the sediment-sam-
ple pits contained an ostracode assemblage that appears to be
of pre-Bonneville age (R.M. Forester, pers. coram., 1992).
FIGURE 10.?Representative sediment-sample pit dug in the Puddle Valley
gravel dunes. Note size of clasts and dip of beds. Shovel handle is approxi-
mately 0.5 m long.
INFLOW DISCHARGE.?In order to reconstruct the paleodis-
charge during inflow from Lake Bonneville to Puddle Valley
several variables have to be determined. Depth of the inflow
current can be estimated from the following regression equa-
tion derived for particle diameters between 5 and 330 cm
(Knox, 1987):
d = 0.0001DL215 ?'57 (Equation 3)
where d is the competent flow depth (in meters), D is the inter-
mediate axis (B-axis) of the coarsest particles moved (in mm),
and S is the dimensionless energy slope. For the Puddle Valley
data, D = 133.0 mm, which is the mean of the B-axes of the 10
largest clasts sampled from the A-A' cross section. S, the hy-
draulic slope, is estimated with 0.0000747, the value derived
from field measurement of the topographic slope just upflow
from the gravel dunes. When substituted into Equation 3 these
values yield a flow depth, d, of 8.36 m. From flow depth and
from the representative particle diameter, a roughness value, n,
is calculated with Limerinos' (1970) equation,
0.113d1/6 (Equation 4)
with D now expressed in meters. ?or relatively wide and shal-
low channels it is generally considered acceptable to substitute
flow depth for hydraulic radius in the Manning equation (Le-
opold etal., 1964),
,2/3cl/2
a S (Equation 5)
where v is flow velocity. Using the depth and slope values from
above and using n = 0.034, as calculated from Equation 4, ve-
locity of flow during dune formation would have been 1.05 m/s.
Paleodischarge, Q, is calculated as
Q = vwd, (Equation 6)
where w is the channel width. Setting approximate channel
width to 1200 m, which is the width of the bar at its proximal
end, and using the values of d and v calculated above, dis-
charge becomes 10,500 m3/s (=370,000 ft3/s). If the maximum
height of the subaqueous dunes, 3.2 m, is substituted for depth
in Equations 4 through 6, a discharge of 2000 m3/s (~73,000
ft3/s) results. This latter value is considered a minimum esti-
mate of discharge during inflow.
TIMING AND NATURE OF THE INFLOW EVENT.?The white
marl is a prominent Lake Bonneville marker bed that was de-
posited during the deeper levels of Lake Bonneville (Gilbert,
1890; Oviatt, 1987). Its location on top of the inflow bar indi-
cates the bar was not created during the regressive phase of
Lake Bonneville, as was previously implied by Currey (1980).
There is no evidence of subaerial alteration of the dunes, and
the contact between the inflow gravel and marl appears con-
formable (Figure 9); therefore, it is concluded that the inflow
bar was created when the transgressing Lake Bonneville spilled
NUMBER 33 137
into Puddle Valley about 25.8 Ka, rather than during a previous
lake cycle.
The dimensions of the bedforms, the size of constituent
clasts, and the presence of rip-up fragments of marl, tufa, and
beachrock, together with the discharge estimate, indicate that
spillover was a powerful event. It seems likely that most of the
material transported in the inflowing current was eroded from
the threshold area. From the rounded to subrounded shape of
the gravels and the present lack of bedrock outcrop at the
threshold it is inferred that just prior to inflow the pass con-
sisted mostly of unconsolidated or semiconsolidated deposits
rather than coherent bedrock. This unconsolidated material
may have had alluvial and colluvial as well as lacustrine com-
ponents.
No geomorphic or stratigraphic evidence has yet been found
to indicate the presence of an isolated water body in Puddle
Valley just before inflow from Lake Bonneville. If there was
one it must have been quite small, occupying only the lowest
part of the subbasin. In the absence of pre-integration Lake
Puddle shorelines near the inflow landform, there is no reason
to conclude that the landform was created by anything other
than the strictly fluvial processes of the inflow current per se,
shortly before being submerged under the rising Puddle Valley
arm of Lake Bonneville. If inflow discharge was indeed within
the 2000 to 10,500 m3/s estimated range, the volume of Puddle
Valley below the sill (=4 km3) would have filled in roughly
four to 23 days.
TULE VALLEY
GENERAL SETTING.?Tule Valley lies between the House
and Confusion Ranges in the southwestern part of the Bonne-
ville Basin, approximately 210 km southwest of Salt Lake City
(Figures 3, 11). Like Puddle Valley, it is currently an indepen-
dent desert basin of interior drainage with only ephemeral
streamflow. Several springs, however, are found on the Tule
Valley floor and along its piedmont junction. Tule Valley has
greater relief (1340-2950 m) and a larger area (1425 km2) than
Puddle Valley. Climatic conditions are essentially the same on
the two subbasin floors, but because of elevation differences,
somewhat cooler and wetter conditions exist in the ranges adja-
cent to Tule Valley compared with those bordering Puddle Val-
ley (Stevens and Brough, 1987; Sack, 1990).
The lowest point on the drainage divide between Tule Valley
and the main body of Lake Bonneville is located at Sand Pass
in northeastern Tule Valley (Figures 11, 12). Sand Pass has an
elevation of 1446 m, which is approximately 21m below the
local elevation of the Provo shoreline. A channel-like depres-
sion slopes about 1.2? from the threshold to the Tule Valley in-
flow bar. The Sand Pass threshold area consists of a compli-
cated assemblage of primarily coastal sediments deposited near
the elevation of the Provo shoreline. In addition to Provo
shoreline deposits, white marl, coastal gravel from the trans-
gression of Lake Bonneville, and indurated remnants of a par-
\ 113.5 W/ \\ p. I )
Great Salt YU \ V
Lake Desert JV I )
VVNVK
Snake | /
Valley I Tule
) Valley /
( Confusion J /
/ Range / JP
' inflow
subbasin
divide
0 5 mi
0 10 km
( ^\^
Sand
Pass
/
\ Juab Co.
S MillardCo.1 \
House ^-v
Range ^i
1
/^ Sevier
V. Desert
\ 1 39?N-
\
FIGURE 11.?Regional map of Tule Valley.
tially eroded sand and gravel coastal landform, interpreted as
an oxygen isotope stage (OIS) 6 barrier complex, are found in
the threshold area. Remnants of the barrier complex consist of
in-place as well as large, partially collapsed blocks of ce-
mented, cross-bedded sandstone and conglomerate beachrock
exposed on both the east and west sides of the channel-like de-
pression. The central portion of the barrier, representing most
of its original length, has been completely eroded away. Lacus-
trine tufa deposited after the barrier was breached fills tension
cracks between semicollapsed blocks of the cross-bedded sand-
stone and conglomerate (Figure 13). Interpretation of the bar-
rier complex as an OIS 6 feature is supported by amino-acid
analyses (conducted by the Amino Acid Geochronology Labo-
ratory (AGL), University of Massachusetts) of gastropod shells
(AGL2083, AGL2084, AGL2085) collected from indurated
coastal deposits inferred to be topsets of the barrier. In addition,
along the west side of the channel the beachrock lies beneath
138 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Tufa
FIGURE 12.?Aerial photograph of the Sand Pass area showing Sand Pass (SP),
the channel-like depression extending from the pass (C), remnants of cement-
ed, cross-bedded sandstone and conglomerate beachrock interpreted as
pre-Bonneville coastal deposits (B), and the Tule Valley inflow bar (IB).
Cross-bedded sandstone and conglomerate beachrock also crop out along the
west side of the channel-like depression.
Fkil'Kr 13.?Lacustrine tufa deposited in a tension crack separating a col-
lapsed block of the cross-bedded conglomeratic beachrock from an intact por-
tion of the beachrock.
26795, 13C adjusted; 17,300 ? 320 yr B.P., unadjusted), deter-
mined on gastropod shells collected from the top of a Lake
Bonneville gravel spit located 45 km south-southwest of Sand
Pass, shows that Tule Valley was integrated with Lake Bonne-
ville at least by 17.7 Ka (Sack, 1990). The 13.5 Ka estimated
age of reisolation of Tule Valley from Lake Bonneville is sup-
ported by a radiocarbon age of 13,430 ? 130 yr B.P.
(Beta-47754, 13C adjusted; 13,020 ? 130 yr B.P., unadjusted)
determined from gastropod shells found in fine-grained calcar-
eous sediments overlying the Tule Valley inflow bar. The ap-
proximately 10 cm thick fine-grained unit containing those
gastropod shells and the overlying 60 cm thick gravel are inter-
preted as sediments of independent Lake Tule deposited after
Tule Valley's reisolation from Lake Bonneville.
GEOMORPHOLOGY AND SEDIMENTOLOGY OF THE INFLOW
BAR.?Like the Puddle Valley example, the inflow landform in
Tule Valley consists of a large-scale linguoid bar that displays
large transverse bedforms, a stoss-side pan, and a lateral and
lee-side rim (Figures 12, 15). The Tule Valley inflow bar is ap-
the transgressive Lake Bonneville sequence of coastal gravel
and white marl.
INTEGRATION INTERVAL.?As with Puddle Valley, the inter-
val of Tule Valley's integration with Lake Bonneville can be
estimated by plotting the rebound-corrected elevation of the
threshold on the Lake Bonneville hydrograph. Using the mod-
ern elevation of Sand Pass, 1446 m, as an approximation of the
pre-integration threshold elevation (Zr) and 1583 m for the lo-
cal elevation of the Bonneville shoreline (Zb), Equation 1
yields a rebound-corrected elevation (Za) of 1426 m. Accord-
ing to the hydrograph of Oviatt et al. (1992), the Bonneville
Basin landscape at 1426 m was submerged by Lake Bonneville
from about 19.4 to 13.5 Ka (Figure 14). These estimates are
consistent with stratigraphic observations and numeric ages ob-
tained in association with Tule Valley lake sediments. A radio-
carbon age of 17,710 ? 320 yr before present (B.P.) (Beta-
1550
1450
?| 1350
LU
1250
PS Provo shoreline
BS Bonneville shoreline
BF Bonneville flood
SS Stansbury shoreline
BS
30
BF
PS
Tule Valley
integration
interval
25 20
Age (Ka)
10
FIGURE 14.?Estimated interval of Tule Valley's integration with Lake Bonne-
ville (a.s.l. = above sea level).
NUMBER 33 139
proximately 2 km long and 1.5 km wide, with elevations rang-
ing from 1391 m at the foot of the distal edge to at least 1426 m
on the right lateral rim. The lakeward slope off the lee-side rim
is about 4.5?. In contrast to the arcuate distal end of the Puddle
Valley feature (Figure 7), the Tule Valley inflow bar displays in
planimetric view a much straighter lee-side rim. The Tule Val-
ley bar is overlain by Bonneville white marl, by coastal sedi-
ments of the post-Bonneville independent Lake Tule (Sack,
1990), and in some places by significant accumulations of post-
lacustrine subaerial deposits. Compared with the Puddle Valley
bar, morphologic details of the Tule Valley bar are more ob-
scured by postlacustrine fluvial, eolian, and mass-wasting pro-
cesses. Nevertheless, the gravel dunes in Tule Valley also ap-
pear to be three-dimensional.
A cross section was measured along the right lateral margin
of the bar where a gravel pit south of Sand Pass Road and a
stream cut north of the road provided good exposures (Figures
15, 16). The representative portion of the measured cross sec-
tion presented in Figure 16 shows subaqueous dunes overlain
by marl and capped by coastal deposits of regressive Lake
Tule. The lower portion of the marl sequence is Lake Bonne-
ville white marl, but the upper part of the unit is interpreted as
Bonneville marl reworked by Lake Tule after Tule Valley's re-
isolation from Lake Bonneville.
Dunes from the entire measured section are large (Ashley,
1990), ranging in length (L) from 15.9 to 61.6 m, in height (H)
from 0.9 to 4.8 m, and in vertical-form index (L/H) from 10.0
to 30.2. The relationship between length and height is de-
scribed by the equation
A-C location of cross section
? bedform & rim crests
[-.. j eolian deposits
I | pan
400 m
2000 ft
H = 0.0409L1.13 (Equation 7)
FIGURE 15.?Map-view diagram of the Tule Valley inflow bar, including the
location of the measured cross section (A-C).
Although the exponent in this equation is somewhat larger than
that reported by Ashley (1990) for a very large data set, it is
smaller than the exponent that Baker (1973:50) found for "giant
current ripples" created by the Lake Missoula Flood (Table 3).
The subaqueous gravel dunes studied in detail along the right
margin of the inflow bar (Figure 15) consist of cross beds that
dip 12? to 31? to the southwest. An exposure at the southeast
corner of the bar, however, exhibits gravel bedform cross beds
that dip 15? to 20? to the southeast.
The 10 largest clasts were investigated from eight sediment-
sample sites along the entire right-lateral exposure. Measured
clasts had an average A-axis of 18 cm and showed dips to the
1415 ?
c 1410H
"CD
LLJ
1405-
B
Regressive Lake Tule deposits
White and reworked marl
Inflow gravel V.E. = 7x
50
I
100
I
150 200
Distance (m)
FIGURE 16.?Cross-sectional diagram of large-scale subaqueous dunes exposed along the southwest edge of the
Tule Valley inflow bar. See Figure 15 for location of B-C. (V.E. = vertical exaggeration.)
140 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
southwest or west. Most of the sampled gravels were sub-
rounded or subangular. Numerous boulders were noted in dune
gravel along the distal edge of the bar (Figure 17). A-axis mea-
surements of the 10 largest clasts observed along the entire
length of the lee-side rim ranged from 100 to 200 cm, with a
mean length of 120 cm. Dune and bar gravels consist primarily
of limestones, dolomites, and quartzites derived from the adja-
cent mountain ranges, but clasts of conglomerate and sandstone
beachrock identical to the indurated remnants of the breached
barrier found adjacent to Sand Pass are common. The 10 larg-
est clasts of sandstone and conglomerate beachrock found in
the inflow bar, up to 2.5 km downflow from the breached bar-
rier, had an average A-axis of 100 cm.
INFLOW DISCHARGE.?The large subaqueous dunes and
large size of constituent gravel reveal that inflow from Lake
Bonneville was a high-velocity event. Paleodischarge can be
reconstructed using the same technique as was employed with
the Puddle Valley data (Equations 3-6). With a representative
clast diameter, D, of 115 mm, which is the average B-axis of
the 10 largest sampled clasts from the Tule Valley gravel dune
sediment pits, and an approximate slope near the pan of
FIGURE 17.?Boulder-size clast of conglomerate beachrock like that found in
outcrop near Sand Pass (Figure 12) but located at the distal edge of the Tule
Valley inflow bar. Lens cap has a 50 mm diameter.
0.0036, depth is estimated to be 0.77 m. These values yield a
roughness coefficient of 0.038 and a velocity of 1.33 m/s. Us-
ing the 1450 m maximum width of the bar for channel width
results in an estimated discharge of 1500 m3/s (* 52,000 ft3/s).
The depth of 0.77 m calculated from Knox's (1987) empirical
equation, however, is much smaller than the maximum bed-
form height, 4.8 m. Substituting 4.8 m for flow depth yields a
velocity estimate of 5.17 m/s and a discharge estimate of
36,000 m3/s (?1,271,000 ft3/s).
TIMING AND NATURE OF THE INFLOW EVENT.?The con-
formable sequence of white marl overlying preserved subaque-
ous dunes verifies that the Tule Valley inflow bar is of trans-
gressive Lake Bonneville age and not from a previous lake
cycle. According to the above estimates of the Tule Valley in-
terval of integration into Lake Bonneville, the inflow bar was
deposited before 17.7 Ka and probably about 19.4 Ka.
The common occurrence within the subaqueous dunes of
conglomerate and sandstone beachrock clasts derived from the
breached barrier located near the divide helps support the no-
tion that the barrier was created during a pre-Bonneville lake
cycle and eroded during the spillover of Lake Bonneville into
Tule Valley. Apparently, the relict pre-Bonneville barrier that
crossed the Sand Pass threshold area failed suddenly either
when the hydraulic load from Lake Bonneville exceeded the
barrier's strength, thereby initiating inflow, or shortly after the
lake began to spill over the barrier into Tule Valley.
Tule Valley definitely contained an independent, transgress-
ing lake, called Lake Tule, when Lake Bonneville spilled over
Sand Pass (Sack, 1990). Shorelines created by Lake Tule are
well exposed around the valley. Their pre-integration relative
age is determined from the conformable relationship between
the shoreline gravel and the Bonneville white marl. The highest
Lake Tule shoreline lies at an approximate elevation of 1416
m. Above that level, shoreline evidence is lacking in a vertical
null zone stretching to about 1446 m, the elevation of the Sand
Pass threshold. A continuous series of Lake Bonneville shore-
lines extends from about 1446 m to the Bonneville shoreline at
approximately 1583 m. When Lake Bonneville stood at the
brink of spilling into Tule Valley, it occupied the highest eleva-
tion it had yet achieved in the lake cycle. At that same time,
Lake Tule was probably also at its highest level; therefore, the
1416 m Lake Tule shoreline should mark the position of the in-
dependent water body when Lake Bonneville began flowing
over Sand Pass (Sack, 1990). Exposures of this highest Lake
Tule shoreline are found adjacent to the lee-side rim on both
the east and west sides of the Tule Valley inflow bar. The distal
end of the Tule Valley inflow bar was apparently deposited at
the margin of the water body in Tule Valley. At the estimated
inflow discharge of 1500 to 36,000 m3/s, the water level in
Tule Valley would have risen the 30 m from the Tule Valley
shoreline to Sand Pass (<22.5 km3) in roughly seven to 174
days.
NUMBER 33 141
Discussion
The linguoid bars created by inflow of Lake Bonneville into
Puddle and Tule Valleys appear very similar to the bars that
Rundle (1985a, 1985b) observed in the first stage of fluvial
braid-bar formation (Figure 18). According to Rundle, that ini-
tial tongue-shaped braided stream deposit, referred to in this
paper as the "Rundle tongue," is an ephemeral, constructional
form deposited downstream from a comparatively deep, con-
fined channel reach. It is created under conditions of radiating
flow where water discharging from the confinement mixes
with calmer water beyond the orifice. Turbulent mixing along
the contact between the stream and quiet water "saps the mo-
mentum of the stream and reduces its velocity progressively
toward the core" (Rundle, 1985a:9-10). Flow is retarded along
the edges of the spreading jet relative to its axis, causing the
flow to turn and spread away from the axis. As flow velocity
decreases, the load is deposited downstream from the orifice
of the confined channel. The resulting deposit, a kind of ex-
pansion bar, is linguoid, consists of an upstream pan and a lat-
eral and downstream rim, and displays a steep lee-side slope.
In addition, Rundle (1985b) observed that some bars have
multiple concentric rims, formed by successive sheets of
gravel being transported over the pre-existing pan surface. The
complete Rundle tongue is a transient feature in braided
streams because it is soon dissected by chutes at falling- and
low-water stages. This dissection of the tongue is, according to
Rundle, the mechanism of braiding. Remnants of the Rundle
tongue preserved between the chutes constitute longitudinal
braid bars.
In many respects the Puddle Valley and Tule Valley inflow
bars are analogous to the initial stage of the braid bar. Like the
Rundle tongue, the inflow bars extend downcurrent from the
orifice of a confined channel, which is the sloping, trough-like
depression leading from each of the thresholds. The Rundle
x
X
11
i
\\
\<
^?-?-?
rim
I\
rim
pan
^**^ orifice
__
orifice
X'
1
X1
FIGURE 18.?Plan and profile views of the ideal tongue-shaped bar that forms
where water flowing as a fast jet leaves an orifice and spreads out into a region
of quieter water (after Rundle, 1985b). The Puddle Valley and Tule Valley
inflow bars closely resemble this model.
tongue and the inflow bars consist of a pan and a well-devel-
oped rim. The subaqueous dunes of the inflow bars are similar
to Rundle's (1985b) multiple concentric rims, and both show
that bedload was being transported across the top of the lin-
guoid bars. A major chute, like those that dissect the initial
braided-stream tongue structure, has breached the Puddle Val-
ley inflow bar on its western side (Figure 7).
Although the Lake Bonneville inflow bars and the Rundle
tongue are similar, they do have some differences. The pan on
the Puddle Valley example appears smaller than that on the
ideal Rundle tongue because of the presence of numerous sub-
aqueous dunes on the bar (Figures 7, 18). Unlike the Rundle
tongue, which forms where a jet spreads out into quieter water,
Lake Bonneville's inflow bars are deposits of the rapid and ex-
panding inflow current alone. Deposition of the Lake Bonne-
ville inflow bars occurred where the inflow current spread out
on the piedmont and was not caused by the presence of lake
water in the receiving basins. Deposition of the Tule Valley in-
flow bar, in fact, ended bluntly where the inflowing current
struck the rising Tule Valley water body. The water level in
Tule Valley rose so rapidly that inflow current expansion
ceased virtually at the moment of contact.
The presence of a large receiving lake in Tule Valley and the
absence of one in Puddle Valley is regarded as responsible for
the difference in planimetric shape of the distal end of the bars
(Figures 7, 15). The flattened distal end of the Tule Valley bar
reflects a sudden decrease in velocity that occurred where the
inflowing jet hit Lake Tule. In Puddle Valley there is no evi-
dence to indicate that the inflow current interacted with a
pre-existing water body. The Puddle Valley inflow bar was de-
posited by the jet of water spreading out onto the piedmont of
an initially dry subbasin, or, perhaps more likely, onto the
slope well above a very small Puddle Valley lake or marsh. By
not meeting rising lake waters until after the bar was com-
pletely formed, the Puddle Valley current expanded symmetri-
cally outward from the axis of flow, giving the resulting de-
posit a more arcuate shape at its distal end. As a result,
compared with the Tule Valley example, the Puddle Valley bar
is regarded as having the more ideal inflow-bar shape.
In contrast to the Rundle tongue of braided fluvial environ-
ments, which is dissected after its formation by falling water
levels (Rundle, 1985a, 1985b), both Lake Bonneville inflow
bars are mostly preserved because their formation was fol-
lowed by rising instead of falling water levels. As the expand-
ing water bodies in the subbasins rose rapidly to equilibrate
with the main body of Lake Bonneville, no major reworking of
the features occurred; however, for 13.6 and 5.9 Ky after their
respective formation, the Puddle Valley and Tule Valley inflow
bars existed at the bottom of a water column that was as much
as 275 m and 183 m deep, respectively. During the sublacus-
trine period the inflow bars were blanketed with pelagic white
marl.
142 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Conclusions
The Rundle tongue reported from a braided stream system
(Rundle, 1985a, 1985b) provides a good analogue for the gen-
eral form of the studied Lake Bonneville inflow bars; however,
details of the geomorphology and sedimentology of the Lake
Bonneville inflow bars reflect the site-specific physical condi-
tions under which they formed. The large size of transported
clasts, the presence of three-dimensional subaqueous dunes, and
the paleohydraulic reconstructions support the notion that in
both studied subbasins Lake Bonneville spillover was a high-
discharge event, probably resulting from an outburst through
material at the threshold. Field evidence in Tule Valley indicates
that Lake Bonneville inflow burst through a relict OIS 6 barrier
complex at Sand Pass. Because such relict "divide barriers" (in-
cluding tombolos) created during the Bonneville lake cycle are
common in the landscape today in passes throughout the lake
basin, and considering the nature of the dune gravel, it is proba-
ble that before integration the unnamed pass separating Puddle
Valley from the main body of Lake Bonneville was also crossed
by some kind of divide barrier. Both of the studied inflow bars
date from the transgressive phase of Lake Bonneville and are
relatively well preserved because the subbasin water level rose
rapidly over them on the transgression and fell relatively
quickly past them during the regressive phase of the lake cycle.
Their original surface form was subdued under a mantle of
lacustrine fine-grained sediments, regressive coastal deposits,
and later subaerial sediments.
The geomorphic and hydrologic circumstances within Puddle
Valley and Tule Valley at the time of Lake Bonneville spillover
were such that the integrating fluvial current was a very
short-lived but continuous flow, although some initial minor
spillover may have occurred. Water flowed through a fairly
confined region adjacent to the threshold as a fast jet. Very
coarse sediment provided by threshold failure and spillway ero-
sion was transported by that current. Where the jet emerged
from the orifice of the confined zone, it spread out. Flow diver-
gence just outside of the orifice formed the pan. Beyond the
pan, bedload was transported in the form of large-scale sub-
aqueous dunes. The spreading current caused a comparative de-
crease in flow velocity especially near the margins of the flow,
resulting in deposition of a fluvial bar. The tongue shape of the
bar reflects this velocity decrease along the edges of the spread-
ing jet relative to its axis, which effected a rotational motion in
the flow (Bates, 1953).
Considering the fluvial processes by which they formed, the
inflow landforms in Puddle and Tule Valleys are most appropri-
ately referred to as bars. The ideal, linguoid bar form was more
effectively developed in Puddle Valley than in Tule Valley be-
cause in the latter subbasin expansion of the fluvial current
ceased abruptly where it struck the Tule Valley lake waters.
Rather than being responsible for formation of the inflow land-
form, interaction of the inflow current with the Tule Valley lake
actually terminated the process of bar formation. For this rea-
son, and because no evidence has been found to indicate the ex-
istence of a transgressive-phase Lake Puddle anywhere near the
Puddle Valley inflow landform at the time of its deposition, the
term delta, although originally applied to these landforms (Cur-
rey, 1980; Currey and Oviatt, 1985; Sack, 1990), is no longer
considered the most appropriate term.
Studying Lake Bonneville as a complex system of compo-
nent subbasins focuses attention on aspects of its geomorphol-
ogy, sedimentology, and chronology that have been generally
overlooked by studying the lake as one simple, coherent unit.
This subsystem approach may be, and to some extent already
has been (e.g., Benson, 1978), profitably applied to other large
paleolake basins, which typically consist of several subbasins.
Because lakes and lacustrine deposits are sensitive indicators of
climate change, paleolake investigations can make important
contributions to reconstructing paleoclimates. The accurate
modelling of paleoclimates, moreover, is essential for the re-
finement of general climate models. Lake-level changes, how-
ever, are influenced by tectonism, drainage-basin changes, and
subbasin integration as well as by climate change. Because
lacustrine inflow bars are evidence of subbasin integration, rec-
ognizing them will help paleoenvironmental researchers to ac-
curately reconstruct the hydrographic history of a paleolake,
thereby contributing to an improved understanding of the true
paleoclimatic signal.
Acknowledgments
Acknowledgment is made to the Donors of The Petroleum
Research Fund, administered by the American Chemical Soci-
ety, for the partial support of this research. The research that led
to this paper was also partly funded by the National Geographic
Society and the Graduate School of the University of Wiscon-
sin-Madison. I am grateful to D.R. Currey, G.C. Nadon, B.G.
Bills, R.M. McCoy, and C.G. Oviatt for suggestions, helpful
discussions, and comments on earlier drafts of this manuscript. I
appreciate the many hours of field assistance provided by T.
Burr, S. Kaplan, J. Killian, G. Nadon, A. Reid, and T. Sabin.
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Environment and Paleolimnology of Owens Lake,
California: A Record of Climate and Hydrology
for the Last 50,000 Years
J. Platt Bradbury andR.M. Forester
ABSTRACT
Diatoms and ostracodes from lake-sediment cores beneath
Owens Lake playa in southeastern California document an almost
continuous paleolimnological record of limnological, hydrologi-
cal, and climatological changes for the last 50 thousand years
(Ky), based upon an extrapolated radiocarbon chronology. Close-
interval sampling enabled chronological resolution at a century-
scale or less. Modern ecologic studies of diatoms and ostracodes,
as well as climatic and hydrologic processes affecting the Owens
Lake basin, were used to track environmental changes. The middle
Wisconsin to Holocene record indicates that climates wetter than
those of modern times existed 80 percent of the time, but that
changes between wet and dry climates occurred rapidly?within
300 years to a century or less in some instances. The ostracode
record indicates that polar air masses permanently resided at the
latitude of Owens Lake during parts of the full-glacial period,
whereas the diatom record and heavy-mineral analyses imply that
extensive alpine glaciers (on the scale of the full-glacial period)
did not exist in the Owens Lake drainage before 22 thousand years
ago (Ka) and that the effects of glaciation on sedimentation in this
area ended by 13 Ka.
During oxygen-isotope stage 3 (-50-25 Ka) at Owens Lake,
summer-season freshwater planktonic diatoms correlate with
geochemical evidence related to increased winter precipitation in
marine cores at similar latitudes on the California margin. These
data document southward shifts and increased strength of the
Aleutian low-pressure system (Aleutian Low) in the northeast
Pacific Ocean and its role in determining the hydrologic balance of
J. Platt Bradbury, 5784 Horseradish Gulch, Golden, Colorado 80403,
USA. R.M. Forester, United States Geological Survey, P.O. Box 25046,
Federal Center, Denver, Colorado 80225, USA.
Reviewers: Larry V. Benson, United States Geological Survey, P.O.
Box 25046, Federal Center, Denver, Colorado 80225, USA; S. Fritz,
Lehigh University, Bethlehem, Pennsylvania 18015; Robert Her shier,
Division of Mollusks, National Museum of Natural History, Smithson-
ian Institution, Washington, D.C. 20560-0118, USA; G.I. Smith,
United States Geological Survey, MenloPark, California 94025, USA.
western North America as well as its effect on the California Cur-
rent. Such correlations indicate that the Owens Lake paleolimno-
logical record responded to the major dynamics of climate
variation in the northern hemisphere.
Introduction
Thick, fossiliferous lacustrine deposits underlying Owens
Lake (Smith and Pratt, 1957) in southeastern California have
been cored and analyzed to extract millennium-scale pale-
olimnologic and paleoclimatic information for the Owens
Lake basin and surrounding region spanning the last 800 thou-
sand years (Ky) (Bradbury, 1997a; Smith and Bischoff, 1997).
In this paper, we interpret climate-controlled paleolimnologi-
cal changes, documented in diatom and ostracode records at
relatively high chronological resolution, between the middle
Wisconsin (-50 thousand years ago (Ka)) and the late Ho-
locene (0.27 Ka). Sample resolution averages about 250 years
for diatoms and about 75 years for ostracodes. In addition, we
synthesize ecological, hydrological, limnological, and clima-
tological data for the lake and its drainage to provide a back-
ground for interpreting the fossil record. We compare the
Owens Lake paleolimnological history to local packrat mid-
den records of past vegetation and to the glacial records in the
Sierra Nevada. Our data provide an independent chronology
for the timing and magnitude of Sierra Nevada glaciation that
augments conclusions of Bischoff et al. (1997b). Finally, we
evaluate the middle Wisconsin paleolimnology of Owens
Lake in terms of paleoceanographic and paleoclimatic changes
registered in marine cores off the California coast spanning
the same time interval as at Owens Lake. This detailed, multi-
proxy record of past climate and hydrology from Owens Lake
provides important insights to the timing, magnitude, and
character of expected future climate changes at Yucca Moun-
tain, a proposed nuclear waste repository site 140 km east of
Owens Lake.
145
146 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Modern Owens Lake Basin
SETTING
Owens Lake (playa floor elevation, 1080 m; area, 182 km2)
occupies a graben immediately east of the highest part of the
Sierra Nevada (Mt. Whitney, 4418 m), near the intersection of
coordinates 36.5?N and 118.0?W (Figure 1). Drainage from the
eastern flank of the Sierra Nevada contributes the major vol-
ume of water to the Owens River. Although streams from the
White and Inyo Mountains east of Owens Lake also recharge
groundwater to the basin, the streams rarely reach the Owens
River (Hollett et al., 1991). Today, most of the water in the
Owens River is diverted to Los Angeles, and the lake basin is
almost dry.
The Owens Lake playa is surrounded by desert saltbush
(A triplex) and shadscale (Sarcobatus) plant communities. Salt
grass (Distichlis) grows on alkaline soils near the playa,
whereas grasses (Poaceae), sedges (Scirpus), and reeds (Phrag-
mites), along with riparian communities of cottonwoods (Popu-
lus) and willows (Salix), characterize the valley floor vegeta-
tion close to the Owens River (Hollermann, 1973). Shrub
communities of creosote bush (Larrea) grow south of the playa
lake, whereas sagebrush (Artemisia) and blackbrush (Coleog-
yne) dominate upstream and to the north of the basin. Scattered
stands of Utah juniper (Juniperus osteosperma) and single-leaf
pinyon (Pinus monophylla), as well as sagebrush, occupy sur-
rounding mountain slopes as low as 1600 m elevation (Litwin
et al., 1997). Along the Owens Valley, these conifers can form
open pygmy conifer forests above 2100 m elevation (Koehler
and Anderson, 1994). Montane and subalpine forests of pine
(Pinus) and fir (Abies) occupy elevations between 2500 m and
the tree line at 3600 m (Hollermann, 1973).
38>o
37C
Owens
Lake
36U -
present playas
and lakes
Pleistocene lakes
? Pleistocene drainage
packrat middens
coresOL-90/1+2
& OL 84/B
marine records
119
FIGURE 1 ?Index map showing Owens Lake and Owens River drainage system in southeastern California and
locations of marine cores along the California margin. B = Bishop, C = Crowley Reservoir, HC - Hot Creek, I -
Independence, LP = Lone Pine, LV = Long Valley, T = Tinemaha.
NUMBER 33 147
CLIMATE
Owens Lake has a wet-winter climate regime. Moisture is
delivered to the west coast of California by the southward mi-
gration of the Polar Front (and the associated Aleutian Low) as
Arctic high-pressure systems expand and coalesce in the win-
ter. As the Subtropical (Hawaiian) High weakens and moves
southwestward, westerly air flow drives storms toward the Si-
erra Nevada where heavy snowfall accumulates; however, be-
cause Owens Lake lies to the east of the Sierra Nevada, a pro-
nounced rain shadow affects the basin. Mean annual pre-
cipitation is only about 14 cm; about 80 percent of which falls
in the winter (Smith and Bischoff, 1997). Annual temperatures
average 15.6? C. Lake evaporation rates are high (1.5 m/yr),
and arid conditions prevail, especially during the summer when
temperatures are high (Figure 2). West of Owens Lake, up to
the crest of the Sierra Nevada, precipitation, mostly as snow,
accumulates rapidly to local annual amounts greater than 150
cm. Today, annual precipitation increases at a rate of about 23
cm per 1000 m elevation in this area (Smith and Bischoff,
1997).
Eighty percent of the variation in California precipitation re-
lates to the direction and speed of the Polar Front jet stream
over the California coast (Bell and Basist, 1994). Interannual
variation of precipitation near Owens Lake occurs as irregular,
circa-decadal cycles (Figure 3). These cycles may reflect un-
stable air-sea interactions in the North Pacific that ultimately
govern the speed and direction of the jet stream (Latif and Bar-
nett, 1994). Heavy precipitation years may or may not co-
incide with the El Nino-Southern Oscillation pattern. Substan-
tial summer precipitation at Owens Lake occurs mostly when
tropical cyclones bring moisture from the Gulf of California
max. temperature
min. temperature
IV V VI VII VIII IX X XI XII
MONTH
precipitation
FIGURE 2.?Thirty-year (1961-1990), monthly normals (mean values) of pre-
cipitation, and maximum and minimum temperatures at Independence, Califor-
nia. Data are from the National Weather Service as contained on the World
Weather Disk, 1994 edition, version 1.4.
and the adjacent Pacific Ocean. Such storms often characterize
those years when the previous winter precipitation in the area
was low.
HYDROLOGY
SURFACE WATER.?In today's climate, precipitation at
Owens Lake is limited and has minor relevance to hydrology
1980
1970
1960
1950
1940
1930
1920
1910
1900
1890 i " " i " " i " " i
10 20 30 40
precipitation (cm)
50
FIGURE 3.?Variation of annual precipitation at Independence, California
(water years 1899-1980). A water year begins on 1 October and ends on 30
September of the following year. Silhouette with horizontal lines represents
summer (May to October) precipitation. Data are from the State of California,
Department of Water Resources.
148 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
and limnology. Nevertheless, wet winters in the Owens Valley
are positively correlated with heavy snowfall in the Sierra Ne-
vada (Figure 4), and snowmelt runoff dominates flow in the
Owens River. Consequently, there is a positive relationship be-
tween winter precipitation and upper Owens River discharge
40
if.
o
8c
o
pend
?ac
oop
CDCM
O
=
?
X
n (cmo
as
+->'Q.
'oa>
35
30
25
20
15
10
':
;
: ?
: V
:
Ml?r
i
?-1 'l-
? _i|l i
?
i _
1 ?
?
?
?
31-1969
0 10 20 30 40 50 60 70 80 90
precipitation (cm) in snow
maximum yearly value (1926-80)
Rock Creek, Bishop, Calif.
FIGURE 4.?Relationship between annual maximum water content of snow
averaged for three Rock Creek snow course stations (elevations of 3049 m,
2759 m, and 2652 m) 30 km WNW of Bishop, California (37?27.0'N,
118?44.5'W), and winter precipitation [December (XII) to March (III)] at Inde-
pendence, California (elevation 1195 m). Snow data are from the State of Cali-
fornia Department of Water Resources, California Cooperative Snow Surveys.
Precipitation data sets span 1926-1980. The 1969 data point identifies data for
that year.
(Figure 5). Before diversion of water from the Owens River to
Los Angeles, the level of Owens Lake was related to Sierra Ne-
vada winter precipitation and the resulting discharge of snow-
melt down the Owens River. For example, the abundant local
precipitation and snowfall during the winter of 1968-1969
(Figures 3-5) so greatly exceeded the capacity of the Los An-
geles Aqueduct that the excess runoff was delivered to Owens
Lake, raising the level of the almost dry playa to a depth of 2.4
m (Friedman et al., 1976). Two years later, Owens Lake had al-
most desiccated again, and dissolved salts reprecipitated over
the playa floor.
The unusually abundant precipitation that occurred during
the winter of 1968-1969 provided a perspective for evaluating
the capability of surface water to fill and maintain Owens Lake
by indicating how sensitive this basin is to increases in effec-
tive moisture. Runoff from snowmelt and precipitation at the
town of Lone Pine (19 km upstream from Owens Lake)
reached 26.8 cm (more than twice the normal) that season and
added 390 hm3 to Owens Lake (Friedman et al., 1976). At that
time, the water diverted to Los Angeles equaled 350 hm3; if
this water had not been diverted, it would have provided a po-
tential total discharge to Owens Lake of 740 hm3 (Smith and
Street-Perrot, 1983). This potential total discharge was 87 per-
cent of that required to maintain a full and overflowing Owens
Lake at present temperatures and evaporation rates. Smith and
Street-Perrot (1983) also calculated that the annual discharge
of Owens River would have to be 210 percent of historic values
(410 hm3) to fill the basin to overflowing. The relative dis-
charge of the Owens River and its tributaries during water year
1969 (October 1968 through September 1969) was 180 percent
of the historic discharge. After the wet winter (January and
February) and spring of 1969, evaporative loss exceeded sur-
face-water input, and Owens Lake desiccated by 1971. Never-
reO
8c
0)?o
c2.
Q)75
C
,/yrE
co
cL
'o2
Q.
50
40
30
? 20
10
- ?
???
?
? ?
? ?
?
m ,
?
?
??
100
-1969
150 200 250 300 350 400 450
Upper Owens River discharge (hm /yr)
FIGURE 5.?Relationship between annual precipitation at Independence, California, and full natural flow dis-
charge on the upper Owens River (Long Valley to Tinemaha) (see Figure 1). Data span 1962-1992. Discharge
data are from the State of California Department of Water Resources, Division of Flood Management. The 1969
data point identifies all data for that year.
NUMBER 33 149
theless, the magnitude of the anomalous precipitation during
the winter of 1968-1969 was almost the amount needed to fill
the lake, especially if mean annual temperatures had been sev-
eral degrees lower (Smith and Street-Perrot, 1983, tbl. 10.2).
The major dissolved ions in high-elevation water from the
Sierra Nevada and White Mountains along the Owens River
drainage are Ca and HCO3 (Table 1). The ratio of total alkalin-
ity (mostly HCO3, in meq/L) to Ca is usually <2 (see Table 1).
Sodium (Na), Mg, and Cl are present in low concentrations
compared with the other major ions, except in White Mountain
drainages that have somewhat higher Mg concentrations de-
rived from dolomite dissolution. Concentration of total dis-
solved solids (TDS) in high-elevation water is low, mostly less
than 200 mg/L. As high-elevation surface water flows down
the Owens Valley, it becomes enriched in Na and total alkalin-
ity (HCO3) that the river gains from groundwater input, which
is modified by interaction with volcanic rock. Locally, water
from highly alkaline, thermal springs (Table 1) enters tributar-
ies to the Owens River, making water in the river more alka-
line.
GROUNDWATER.?The input of groundwater via springs and
seeps to the immediate Owens Lake groundwater basin (39 km
in length) was estimated to be 28.6 hm3 (Lopes, 1987). Hollett
et al. (1991) estimated spring and seep discharge along the
Owens Valley aquifer north of Owens Lake (108 km in length)
to be 40.7 hm3. If these figures are approximately correct, they
indicate that proportionally more spring discharge occurs in the
immediate Owens Lake groundwater basin than along the
Owens Valley. The combined discharge, about 69 hm3, is about.
20 percent of the average natural discharge of the Owens River
(370 hm3) at Lone Pine (Gale, 1914). Accordingly, groundwa-
ter input could have an important effect in determining the vol-
ume of Owens Lake were it not for withdrawal of water to the
Los Angeles Aqueduct.
Springwater and groundwater chemistries along the Owens
Valley have two distinct categories: low-salinity water domi-
nated by Ca and HCO3 and alkalinity-enriched water of modest
to high salinity (Hollett et al., 1991). Groundwater from rela-
tively shallow, permeable, hydrogeologic units that is chiefly
recharged through alluvial fans by Sierra Nevada runoff char-
acterizes the low-salinity water. This type of groundwater is
typical of the west side of Owens Lake and of the aquifer asso-
ciated with the Owens River delta at Owens Lake. Deep and
presumably old groundwater associated with active volcanic
processes or with ancient deposits from Owens Lake is saline
and high in Na, HCO3, and Cl. This alkalinity-enriched
groundwater characterizes the east side of Owens Lake where
it rises along permeable fracture zones and mixes to variable
degrees with fresher, shallow groundwater sources associated
with alluvial fans (Font, 1995).
Large, active springs, such as Casa Diablo Hot Springs and
others along Hot Creek, discharge into the Owens River system
in the Long Valley caldera area (Figure 1) and can supply large
amounts of dissolved solids to surface water that ultimately
flows to Owens Lake (Smith, 1976a; Menking et al., 1997).
Crowley Reservoir, on the Owens River, receives a mixture of
alkaline-enriched springwater and relatively low-salinity, cal-
cium bicarbonate water from high-elevation Sierra Nevada
sources. We estimate that about 10 percent of Crowley Reser-
voir water input comes from springs and seeps in the Long Val-
ley caldera region. This estimate was based upon the percent of
Cl in the Hot Creek thermal springs (Smith, 1976b) and in the
high-elevation Sierra Nevada water sources (Table 1).
The input of alkaline-enriched groundwater via springs and
seeps is important in the hydrochemistry of Owens Lake, espe-
cially during times of low surface-water discharge. Low-eleva-
tion water derived from spring discharge is dominated by Na,
HCO3, and Cl. Sulfate is variable, and Ca and Mg are at low
concentrations. The alkalinity to Ca ratio (alk/Ca) is large,
generally ?2, and the TDS are high, often greater than 500
mg/L (Table 1). During evaporation, the low-elevation, alka-
line-enriched springwater and seep water becomes increasingly
more concentrated with solutions of Na and HCO3. This hydro-
chemistry characterized Owens Lake in the late nineteenth cen-
tury (Gale, 1914), because surface-water input was low and al-
kaline-enriched springwater input was proportionally high
during arid climate conditions. Owens Lake hydrochemistry is
predisposed toward Na, HCO3, and Cl enrichment from the
chemical character of its groundwater input, and dissolved sol-
ids in the lake water can undergo considerable concentration
before saline minerals begin to precipitate and change the basic
hydrochemistry. In such a system, Ca entering the lake via the
Owens River would soon precipitate as CaCO3.
LIMNOLOGY
The earliest limnological observations concerning Owens
Lake are from the 1861 Nevada-California border survey
(Woodward, 1961). The cited description (on 27 March 1861)
indicated an approximately oval or pear-shaped lake about 26
km long and 15 km wide, with an extensive, shallow littoral
area. A white alkali efflorescence bordered the lake margin;
this was replaced farther upslope by a belt of salt grass (Dis-
tichlisl) and rushes (probably Scirpus). High winds made the
water turbulent and turbid. In the middle to late nineteenth cen-
tury, Owens Lake was a closed basin about 15m deep and alka-
line. Between 1872 and 1905, the lake received only about 25
percent of its natural input because of water divergence for irri-
gation in the Owens Valley (Gale, 1914). Nevertheless, the in-
flow was sufficient to maintain a perennial, if saline, lake about
6 m deep. The construction of the Los Angeles Aqueduct in
1913 removed the principal source of water to the lake, and by
1923 it desiccated (Smith, 1976b; Tyler et al., 1997). If filled to
its spill point today (1145 m), Owens Lake would be about 65
m deep with an area of 622 km2 and a volume of 22,100 hm3
(BischoffetaL, 1997a).
In its low-stage, Holocene phase, Owens Lake would have
had a seasonal (early summer) input of freshwater that would
150 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
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have soon mixed throughout the water column because of the
lake's large fetch and comparatively shallow depth. As a conse-
quence of snowmelt and runoff, levels in Owens Lake would
have fluctuated seasonally, as has been documented at Mono
Lake (Mason, 1967). The presence of abundant oolites in the
predesiccation sediments of Owens Lake implies that condi-
tions of calcite supersaturation and shallow depths sufficient
for wind-generated currents to agitate sediment grains pre-
vailed at that time (Smith, 1997).
The salinity of historic Owens Lake varied between 60 and
>200 g/L TDS as a result of short-term climate variation and
increasing use of water for irrigation. The lake water was a so-
dium bicarbonate plus carbonate brine that supported abundant
brine flies (Ephydra), brine shrimp (Artemia), and assorted
microorganisms (infusoria) (Lee, 1906). Fishes, such as the
Owens sucker (Catostomus fumeiventris Miller), now living in
the Owens River might have survived in the lake during peri-
ods of freshwater inflow and stratified lake conditions until
high salinities eliminated the populations, as documented by
Smith etal. (1987).
The lake could have stratified thermally, at least for short pe-
riods of calm, warm weather (Friedman et al., 1976), but at its
shallow-depth stages, the water column usually would have
mixed completely, as reported in 1861. Most of the time, there-
fore, bottom-water temperature would have been close to the
mean monthly air temperature (Smith et al., 1987). Possibly,
short-lived periods of chemical stratification also existed dur-
ing the runoff season when less dense, freshwater could float
on top of dense, saline water. Owens Lake was almost certainly
a warm, polymictic system?a lake mixing throughout the year
without ice cover (Lewis, 1983). Pyramid and Walker Lakes in
Nevada, 250 and 400 km north of Owens Lake, respectively, do
not freeze today, although they are fresher and deeper than the
historic Owens Lake. At high stages (e.g., 30+ m in depth),
Owens Lake probably would have resembled the warm mo-
nomictic Pyramid and Walker Lakes, which circulate in the
winter and thermally stratify in the summer.
DIATOM ECOLOGY
Diatoms certainly lived in Owens Lake in the late nineteenth
and early twentieth centuries, even when lake salinity was al-
most 100 g/L, but apparently no information exists from those
times. The modern diatom flora of Mono Lake may be repre-
sentative of the diatom communities of the historic Owens
Lake, at least in part. Historically, Mono and Owens Lakes had
similar chemistries and high alk/Ca ratios (Lee, 1906). The
present Mono Lake salinity of about 100 g/L (Table 1) is of the
same order of magnitude as the salinity for historic Owens
Lake before the diversion of water from the Owens River, al-
though considerable variation in salinity characterized both
lakes in the early twentieth century (Lee, 1906; Gale, 1914). In
littoral samples between 0.15 and 10 m depths, Kociolek and
Herbst (1992) listed Anomoeoneis sphaerophora (Ehrenberg)
152 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Pfitzer, Navicula crucialis (O. Miiller) Frenguelli, Nitzschia
frustulum (Kutzing) Grunow, Nitzschia latens Hustedt, and a
new species, Nitzschia monoensis Kociolek and Herbst, as the
numerically important diatoms in Mono Lake. The same, or
closely related, species occur in the Holocene sediments of
Owens Lake.
Diatoms in the Owens Lake sediments may be grouped into
taxa having ecological associations with freshwater (<1 g/L
TDS) and saline water (> 1 g/L TDS). Many species found in
the cores overlap in their salinity tolerances, so these associa-
tions need to be used carefully. For example, although most of
the taxa considered to live in freshwater have optima well be-
low 1 g/L TDS, they may inhabit waters two to three times
more saline. Furthermore, saline taxa, which often tolerate (or
even prosper in) a broad range of salinities above 5 g/L TDS,
may be present in waters with less than 1 g/L TDS.
Diatoms occupy different physical habitats in the photic zone
of lakes: planktonic diatoms remain in open water and are sus-
pended by turbulence; benthic semiattached diatoms are loosely
attached to substrates, such as submerged vascular plants;
benthic motile diatoms are capable of moving over and through
the mud in illuminated areas of the lake bottom; and benthic ad-
nate diatoms are firmly attached by stalks or pads of mucilage
to substrates in the shallowest parts of lakes and can withstand
strong currents. In modern lakes, these habitats are more or less
distributed along a water depth gradient (deeper to shallower)
controlled by the photic zone and the effect of currents.
As with salinity, diatoms overlap in habitat, especially sea-
sonally. Semiattached species may become planktonic when
there is sufficient turbulence, as do some benthic motile species
of Campylodiscus, Nitzschia, and Surirella. Planktonic species
often settle to the lake bottom until favorable conditions return
to support them in the open-water photic zone. These criteria
and habitat options characterize three principal ecological
groups that have more or less coherent ecological distributions
(Figures 6-8).
1. Saline diatoms (Figure 6). Benthic species: Amphora
coffaeiformis (Agardh) Kiitzing, Anomoeoneis sphaerophora,
Campylodiscus clypeus Ehrenberg, Navicula crucialis, Navic-
ula pygmaea Kiitzing, Nitzschia frustulum, Nitzschia monoen-
sis, Nitzschia pus ilia Grunow, Rhopalodia acuminata Kram-
mer, Surirella hoefleri Hustedt, Surirella ovalis Brebisson,
Surirella striatula Turpin. Planktonic species: Chaetoceros
muelleri Lemmermann, Cyclostephanos sp., Cyclotella me-
neghiniana Kiitzing, Cyclotella quillensis Bailey.
2. Freshwater benthic diatoms (Figure 7). Attached spe-
cies: Achnanthes lanceolata (Brebisson) Grunow, Amphora
ovalis (Kutzing) Kutzing, Amphoraperpusilla Grunow, Cocco-
neis placentula Ehrenberg, Cymbella muelleri Hustedt, Epi-
themia adnata (Kutzing) Brebisson, Epithemia turgida (Ehren-
berg) Kiitzing, Fragilaria capucina Desmazieres,
Rhoicosphenia curvata (Kutzing) Grunow, Rhopalodia gibba
(Ehrenberg) O. Muller. Semiattached species: Fragilaria
vaucheriae (Kutzing) Peterson, Pseudostaumsira brevistriata
(Grunow) Williams & Round, Staurosira construens (Ehren-
berg) Williams & Round, Staurosirella pinnata (Ehrenberg)
Williams & Round, Synedra rumpens Kutzing, Synedra ulna
(Nitzsch) Ehrenberg.
3. Freshwater planktonic diatoms (Figure 8). All species of
Aulacoseira and Stephanodiscus; Asterionella formosa Has-
sall, Cyclotella bodanica Grunow, Cyclotella ocellata Pantoc-
sek, some Cyclostephanos spp., Fragilaria crotonensis Kitton,
Tabellaria flocculosa (Roth) Kutzing.
Because of the great variety and abundance of freshwater
planktonic diatoms in the Owens Lake cores, the ecological dy-
namics and limnoclimatic relationships of these species are im-
portant in the interpretations of past environments in Owens
Lake.
Stephanodiscus species generally characterize phytoplankton
blooms in the spring or fall (Kilham et al., 1986; Bradbury,
1988). In the Owens Lake drainage, Stephanodiscus oregoni-
cus (Ehrenberg) Hakansson and the small Stephanodiscus spe-
cies (S. minutulus (Kutzing) Round and S. parvus Stoermer and
Hakansson) are indicators of major hydrologic input in the
spring, probably associated with melting of the Sierra Nevada
snowpack and discharge into the Owens River. In March 1996,
S. oregonicus was numerically dominant within the phy-
toplankton communities in Klamath Lake, Oregon (Figure 8).
Stephanodiscus niagarae Ehrenberg is probably a fall species
because it was numerically dominant in the September 1994
bloom at Crowley Reservoir, upstream from Owens Lake
(J.P.B., unpublished data).
Except for Stephanodiscus niagarae, the Stephanodiscus
species common in the Owens Lake cores cannot be perfectly
related to existing taxa in the Western United States. In general,
the fossil Owens Lake species are more robust and more
heavily silicified than modern relatives, perhaps as a result of
distinctive nutrient conditions that existed in ancient Owens
Lake (Theriot and Stoermer, 1981). Descriptions of some Qua-
ternary fossil Stephanodiscus taxa, however, fit the morpholog-
ical characteristics of the fossil Owens Lake diatoms fairly
well, although much variation exists in the assemblages. For
example, the fossil taxa Stephanodiscus subtransylvanicus
Gasse and S. asteroides Gasse (Gasse, 1980) are morphologi-
cally close to S. oregonicus (and to each other), but they tend to
be larger, heavier, and more coarsely structured than modern
representatives of S. oregonicus in the western United States.
Because S. asteroides compares favorably with the Owens
Lake fossil Stephanodiscus taxa, this epithet is used for this
species complex in the core studies. Because of its general sim-
ilarities to S. oregonicus, S. asteroides was assumed to occupy
a similar ecological niche as S. oregonicus, i.e., blooming in
the spring shortly after ice breakup when circulation and river
inflow provide abundant nutrients.
Asterionella formosa blooms in low-conductivity lakes when
phosphorus to silicon ratios are low, usually in the early sum-
mer or early fall (Kilham and Kilham, 1978). Its presence in
Owens Lake would indicate conditions of comparatively dilute
freshwater and seasonal nutrient supplies governed by phy-
toplankton succession and lake stratification. Asterionella for-
NUMBER 33 153
Site
TDS (mg/L)
,55- *&f y </ y / / /? /? ^ ^
BENTHIC DIATOMS 96'442 10'650 9'154 7'985 5'110 1>118 632 247 133 128
Nitzschia monoensis
Navicula crucial is
Nitzschia pusilla
Nitzschia frustulum
Amphora coffaeiformis
Anomoeoneis sphaerophora
Surirella spp. aff. S. ovalis
Surirella striatula
Campylodiscus clypeus
Rhopalodia acuminata
PLANKTONIC DIATOMS
Chaetoceros muelleri
Cyclotella quillensis
Cyclostephanos sp.
Cyclotella meneghiniana
FIGURE 6.?Estimated abundance of saline-habitat diatoms from modern aquatic environments in the western
Great Basin. Sites along the abscissa are arranged in order of decreasing salinity (see Table 1).
mosa was common in phytoplankton collections taken in June
1996 from Sierra Nevada lakes and reservoirs above an eleva-
tion of 2000 m (J.P.B., unpublished data).
Little is known about the ecology and nutrient requirements
of the Cyclostephanos species, although some prosper in highly
eutrophic and often turbid freshwater systems (Hakansson and
Kling, 1990). The species found in the sediments of Owens
Lake presently lives in Utah Lake (500-1300 mg/L TDS) dur-
ing the winter (under ice) and early spring (Squires and Rush-
forth, 1986; they initially identified this diatom as Stephanodis-
cus dubius). In the Owens Lake system, this taxon also
probably represented winter or very early spring diatom pro-
duction. A similar diatom lives in hypereutrophic Upper Cow
Lake in Malheur County, Oregon, and in Diaz Lake, about 4 km
northwest of the Owens Lake playa (J.P.B., unpublished data).
Aulacoseira granulata (Ehrenberg) Simonsen and A. am-
bigua (Grunow) Simonsen bloom when wind-driven mixing of
the water and a deepening of the thermocline supply nutrients,
generally phosphorus and silica, from the hypolimnion and
when turbulence maintains the diatoms in the photic zone.
154 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Site
TDS
(mg/L)
<<r
10,650 5,110 632 247 223 133 128 122 104 41 38 32 22
Rhoicosphenia curvata
Amphora ovalis
Cocconeis placentula
Amphora perpusilla
Fragilaria capucina*
Epithemia adnata
Epithemia turgida
Pseudostaurosira
brevistriata*
Achnanthes lanceolata
Rhopalodia gibba
Staurosira construens*
Cymbella muelleri
Staurosirella pinnata*
Fragilaria vaucheriae*
Synedra ulna*
Synedra spp. aff.
S. rumpens*
scale
FIGURE 7.?Estimated abundance of benthic diatoms from modern freshwater aquatic environments in the west-
ern Great Basin. Sites along the abscissa are arranged in order of decreasing salinity (see Table 1). Taxa with
asterisks are semiattached species.
These diatoms are indicators of warm, open-water (or at least
high light) conditions and abundant nutrients. Aulacoseira
granulata often lives in shallow lakes where nutrients are
readily available during windy periods (Brugam, 1983). In
Owens Lake, these species represented favorable limnological
environments in the summer, perhaps because summer precipi-
tation maintained the lake against excessive salinities from
evaporation or, more likely, because substantial discharge from
the Owens River persisted through the summer and into the
fall. The modern limnology of Crowley Reservoir and Lower
Twin Lake may serve as reasonable analogs for Aulacoseira-
rich intervals of the Owens Lake core. These lakes are 1000 m
or more above Owens Lake and are in moister, cooler climatic
zones that have abundant hydrologic input from the Owens
NUMBER 33 155
Site
elevation
(m)
Stephanodiscus
oregonicus
Aulacoseira granulata
Stephanodiscus
niagarae
Stephanodiscus
pan/us
Stephanodiscus
minutulus
Aulacoseira ambigua
Fragilaria crotonensis
1,262 1,518 2,067 2,158 2,162 2,174 2,317 2,323 2,323 2,378 2,731
scale
Asterionella formosa
Tabellaria flocculosa
Cyclotella bodanica
Aulacoseira subarctica
Cyclotella ocellata
FIGURE 8.?Estimated abundance of planktonic freshwater diatoms from modern aquatic environments in the
western Great Basin. Sites along the abscissa are arranged in order of increasing elevation (see Table 1).
River. Both lakes have low summer turbidity that promotes
high-light-dependent diatoms.
Freshwater diatoms disperse to and proliferate in new habi-
tats rapidly and easily. In addition to water transport, diatoms
may be carried by birds and wind because many species have
resistant resting spores or accumulate oils to retard desiccation.
The ability of diatom communities to rapidly respond to habitat
changes makes them excellent paleolimnologic indicators.
OSTRACODE ECOLOGY
Our interpretation of the fossil ostracode data is based on the
chemical preferences and biogeographic distribution of these
species, which we derived from mostly unpublished United
States Geological Survey (USGS) ostracode habitat and chemi-
cal data sets. The data were collected by USGS staff, Donald F.
Palmer and Alison J. Smith (Kent State University), and by B.
Brandon Curry (Illinois State Geological Survey) from springs,
wetlands, lakes, and a few streams, totaling about 800 localities
throughout the United States. Additional information on ostra-
code distribution and ecology came from more than 3500 Ca-
nadian sites (Delorme, 1989).
Candona caudata Kaufmann, Cytherissa lacustris Sars, Lim-
nocythere ceriotuberosa Delorme, and L. sappaensis Staplin,
the most common ostracodes in the Owens Lake cores, now
live in lakes that have varied chemical, thermal, and hydrocli-
156 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
matic characteristics (Forester, 1986, 1987, 1991; Delorme,
1989; Forester et al., 1994; Smith and Forester, 1994; R.M.F.,
unpublished data). Candona caudata also commonly inhabits
streams, whereas L. sappaensis lives in some spring pools and
wetlands. Limnocythere ceriotuberosa and L. sappaensis com-
monly live in slightly to very saline water, respectively, that has
high alk/Ca ratios, whereas Cytherissa lacustris and Candona
caudata more typically live in freshwater habitats (Forester,
1983, 1986; Bradbury et al., 1989; Forester et al., 1994).
In the Owens Lake drainage, Candona caudata lives in the
Owens River and also in a number of lakes at higher elevation
(>2000 m) in the Sierra Nevada. Limnocythere sappaensis
lives in saline springs around Owens Lake, such as Dirty Socks
Spring and the Keeler springs on the eastern margin of Owens
Lake, and in wetlands and shallow lakes having high alk/Ca ra-
tios (>7) along the Owens River (e.g., Big Alkali Lake and Lit-
tle Lake). Limnocythere ceriotuberosa lives in Diaz Lake,
which varies seasonally in water volume and amount of TDS
(as a result of controlled input from the Los Angeles Aqueduct)
and has a moderately elevated alk/Ca ratio of just more than 3.
This species has not been collected elsewhere in the Owens
River drainage.
Cytherissa lacustris does not live in the Owens River drain-
age nor in any lake south of Canada that was sampled during
this study, except for rare populations in the Great Lakes and in
Yellowstone Lake. In Canada and Alaska, C. lacustris lives in
dilute, oligotrophic, cold-water lakes of the boreal forest. De-
lorme (1989) reported the upper TDS tolerance of C. lacustris
at 300 mg/L. The level of TDS alone, however, does not ac-
count for the distribution of C. lacustris. This species also re-
quires lakes that have interseasonal chemical stability and low,
or at least predictable, thermal variation.
The biogeographic distribution, temperature, and chemical
preferences of these four ostracode species provide a means of
understanding the probable climate and limnological history of
Owens Lake. The presence of Cytherissa lacustris in core sedi-
ments implies there once was an almost continuous residence
of polar air in the Owens Lake region. Today, this air mass re-
mains in the boreal regions of Canada and southern Alaska.
The presence of polar air in the Owens Lake drainage during
past climates reflects the continental glaciation that existed in
northern North America, which maintained polar air south of
the ice-sheet margin. The subsequent disappearance of C.
lacustris from the fossil record implies the retreat of polar air
northward and the start of deglaciation. From a limnological
standpoint, the presence of C. lacustris implies, by analogy,
that a dimictic lake with a long ice-cover season existed in the
Owens Lake basin.
Candona caudata may be present with Cytherissa lacustris,
but Candona caudata has a wider distribution in lakes and
streams well south of Canada and Alaska. The southernmost
modern record of Candona caudata is in southern Nevada, and
its upper TDS limit is about 3 g/L (Bradbury et al., 1989). At
Owens Lake, the presence in the fossil record of Candona cau-
data without Cytherissa lacustris implies the generally fresh-
water conditions and temperature regimes that are characteris-
tic of the Great Basin today also occurred in the past. To
maintain freshwater habitats in the shallow lake basin, flow in
the Owens River needed to have been sufficient on a seasonal
basis to offset potentially high valley-bottom evaporation rates;
therefore, the presence of Candona caudata in the fossil record
indicates a positive hydrological balance that, at least, partly
reflects greater precipitation at that time.
Limnocythere ceriotuberosa usually lives in lakes that have
strong seasonality in volume and associated chemistry. It has
been collected alive in lakes that have a TDS level as low as
400 mg/L and as high as 12 g/L. Lakes such as Pyramid and
Walker Lakes, which seasonally freshen and increase in size
from surface flow but then shrink and become more saline once
surface flow decreases, are common settings for this ostracode.
Often these limnological changes in such lakes represent a sea-
sonal riverine hydrological shift from surface flow derived
from snowmelt to base flow derived from shallow aquifers.
Generally, such lakes and rivers are in semiarid, temperate set-
tings. Increased abundance of L. ceriotuberosa in Owens Lake
deposits suggests surface-water discharge into the Owens
River and then to Owens Lake varied seasonally. Limnocythere
ceriotuberosa often appeared in the Owens Lake fossil record
during transitions between glacial and interglacial climates.
Limnocythere sappaensis commonly lives in wetlands and
lakes that are saline or that maintain a saline, alkaline-enriched
solute composition during the year (Forester, 1983, 1986). The
species lives in wetlands and lakes that have salinities as low as
650 mg/L and as high as 110 g/L. The solute and salinity char-
acteristics of its habitat indicate that L. sappaensis would have
lived in Owens Lake during arid interglacial periods when al-
kaline base flow predominated in the Owens River, which re-
sulted in a generally low-volume lake with alkaline, saline hy-
drochemistry. Present-day Owens River water, if combined
with alkaline spring discharge and allowed to enter Owens
Lake, could support L. sappaensis.
Like diatoms, ostracodes disperse rapidly and easily, and
they proliferate quickly. Because of their larger size, most adult
and juvenile ostracode transport is by birds. Ostracode eggs,
some of which are sticky and require desiccation to hatch, may
be transported by fish, birds, insects, and other terrestrial ani-
mals, or perhaps by wind. Ostracode eggs apparently survive
transport through the digestive tracts of most transport animals.
Fossil Record from Owens Lake Basin
METHODS
The paleolimnological history of Owens Lake is based upon
three continuous sediment cores. Core OL-84-B was raised in
1984 from near the center of the Owens Lake playa (Figure 1)
(36?23.91'N, 117?56.142'W) using a modified Livingstone
piston corer (Lund et al., 1996; Benson et al., 1997). It bot-
tomed at a depth of 11.6 m. Cores OL-90-1 and OL-90-2 were
NUMBER 33 157
obtained about 3 km southwest of core OL-84-B using a truck-
mounted, hydraulic, split-spoon sampler (Benson et al, 1996).
Cores OL-90-1 and OL-90-2 bottomed at depths of 32.8 and
28.2 m, respectively. For diatom and ostracode analyses, a 34
m composite section (OL-90-1+2) containing the upper 28.16
m of core OL-90-2 and the basal 5.91 m of core OL-90-1 was
constructed from a magnetic susceptibility correlation at 3 cm
intervals (Benson et al., 1996).
One-half cubic centimeter splits of magnetic susceptibility
samples from core OL-84-B and from the upper 16 m of core
OL-90-2 provided material for diatom studies. The remaining
samples from cores OL-90-2 and OL-90-1 were excised with a
spatula and placed in labeled, air-tight plastic bags. Approxi-
mately 0.5 cm3 subsamples were processed for diatoms by us-
ing concentrated HNO3 in a water bath at 100? C. After the res-
idue was washed free of acid and dissolved solids, it was
settled onto coverslips and mounted in naphrax (Bradbury,
1997a). Diatom data were presented as relative abundance. In
samples that had diatom valve counts less than 100, an artifi-
cial count of 100 was used to calculate percentages, thereby un-
derweighting relative diatom abundances in samples of low di-
atom concentration. Silt-sized heavy-mineral grains of
magnetite, zircon, and monazite were tallied from the diatom
microslides on the basis of color, transparency, and crystal
form. Heavy-mineral grain concentration is presented as the
number of grains per millimeter of microscope transect at
xlOOO magnification. Heavy-mineral grain identifications were
verified by petrography.
Estimated abundance of diatoms in modern samples are de-
noted as follows: dominant = more than 10 specimens of a
taxon in every field of view examined; abundant = taxon oc-
curred in every field of view; common = taxon present in every
fourth field of view; few = taxon present in one of 10 fields of
view; rare = taxon present in one of 20 fields of view; and
absent = not seen.
Sections of about 3 cm thickness were cut from core OL-84-
B between 1.63 m and 9.11 m to obtain ostracode samples.
Sections of about 2 cm thickness were cut from cores OL-90-
1+2. The sediment was weighed, disaggregated, and washed
through sieves to concentrate the ostracodes, which were re-
ported as number of adult valves per gram of sediment.
CHRONOLOGY
Polynomial-age models were based upon 40 radiocarbon
dates from the total organic fraction using accelerator mass
spectroscopy (AMS) (Benson et al., 1996; Benson et al., 1997).
These provided the chronologies for cores OL-84-B and OL-
90-1+2. The radiocarbon dates indicate hiatuses or periods of
very slow deposition at 2.25 m and 9.13 m in core OL-84-B. A
hiatus at 9.13 m in core OL-84-B correlates with one at 6.0 m
in core OL-90-2. In both cores, this hiatus separates lithostrati-
graphic units C and D of Lund et al. (1996) and represents
about 2000 years of missing time (and sediment?) between
about 15.5 and 13.5 14C Ka in core OL-90-2 (Benson et al.,
1996). In core OL-84-B, the C-D transition occurs 30 cm
above a radiocarbon date of 15.02 ?0.1 Ka, indicating that, in
this core, the hiatus may span only from 14.7 to ~13.5 Ka. Un-
calibrated radiocarbon dates used in this paper are reported as
thousands of years ago (Ka). According to these chronologies,
samples from cores OL-90-1+2 date from 52.5 to 12.7 Ka.
Samples from core OL-84-B date between 17.3 and 0.27 Ka.
CORES OL-90-1+2
DIATOMS.?Freshwater planktonic diatoms numerically
dominate most of the OL-90-1+2 record. Saline diatoms only
appear near the base of the core at an extrapolated age of about
52.5 Ka (Figure 9). The saline diatom assemblage is an indica-
tor of environmental conditions, including low lake levels, im-
portant hydrologic contributions from alkaline, mildly saline
springs along the Owens River, and seasonal but reduced input
of surface water from the Owens River and tributaries marginal
to Owens Lake.
Cy do Stephanos spp., Stephanodiscus asteroides, S. niaga-
rae, and Aulacoseira ambigua rapidly replaced the saline dia-
tom assemblage by 52.0 Ka and continued to characterize
Owens Lake until about 41 Ka when Stephanodiscus sp. cf. S.
carconensis Grunow became the predominant diatom in the
lake. Benthic freshwater diatoms (Figure 10) also formed an
important part of the assemblage. Pseudostaurosira, Stauro-
sira, and Staurosirella species, hereafter referred to as
"Fragilaria" species because they once resided in that genus,
fluctuated but increased in abundance from 52 Ka until 41 Ka.
This trend indicates a progressively greater representation of
shallow-water habitats through time and may imply either a
gradual shallowing of the lake caused by sedimentation or the
establishment of marsh and shallow, freshwater habitats around
Owens Lake and upstream along the Owens River. Other
benthic diatom ecological groups had less regular fluctuations.
These fluctuations in benthic freshwater diatoms, as well as the
fluctuations of the planktonic freshwater diatoms, partly reflect
percentage control when one or another species, usually a
planktonic diatom, became especially abundant, causing the
other taxa to decrease to proportionally lower percentages.
Nevertheless, high percentages of both groups document lim-
nological changes, which were possibly climate mediated, at
comparatively short intervals. Aulacoseira spp. and Stephano-
discus niagarae assemblages alternated with minor, but ap-
proximately regular, increases in S. asteroides (Figure 9), per-
haps indicating alternations of greater and lesser Sierra Nevada
precipitation, respectively. In particular, the five regular pulses
of S. asteroides between 50 and 43 Ka (Figure 9) indicate a cy-
clicity of 1.4 Ky, a millennial-scale rhythm (within the limits of
chronological control) that has also been observed in marine
cores that reflect discharge of cold, ice-bearing water into the
North Atlantic Ocean (Bond et al., 1997).
Aulacoseira ambigua, A. granulata, and Stephanodiscus nia-
garae bloom today in the summer or early fall in intermediate-
158 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
IftilfiitiitLdtlhWlin^^
3 2
?2 a
-g a
o
o
rj-
O
100
o
oCD
ent
2CD
Q.
iato
?u
nico
c
"5.
a
;ntriio
is, major
diatorr
_c"I
?S3
oK
55-s:
d
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wens Lai
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96B
NUMBER 33 159
55
50 0 10
i
30
FIGURE 10.?Percentage of minor planktonic and benthic freshwater diatoms versus age in cores OL-90-1 and
OL-90-2 from Owens Lake. Planktonic araphid spp. represent Fragilaria crotonensis, F. vaucheriae, F.
capucina, Synedra acus, and S. rumpens. "FragilaricT spp. include members of Staurosira, Pseudostaurosira,
and Staurosirella. Attached diatoms mostly represent Cymbella, Epithemia, Cocconeis, Achnanthes, Gompho-
nema, and Rhoicosphenia. Motile diatoms mostly represent Navicula and Nitzschia.
elevation lakes (1000 to 1300 m above Owens Lake) in the Si-
erra Nevada (Figure 8). These diatoms require freshwater con-
ditions (generally < 1000 mg TDS/L) during the summer; thus,
their presence implies the existence of temperate to warm-tem-
perate environments and water bodies that have sufficient hy-
drologic input to offset evaporative losses. Consequently, these
analogs indicate periods of temperate, but moist, climate at
Owens Lake between 52 and 24 Ka (Figure 9.). Although the
episodes characterized by Aulacoseira spp. do not form a par-
ticularly coherent signal, their waxing and waning indicates
multi-millennial scales of climate variability that also has been
documented in the North Atlantic Ocean (Heinrich, 1988).
After 40 Ka, Stephanodiscus sp. cf. S. carconensis Grunow,
hereafter referred to as S. "carconensis," dominates the diatom
record of Owens Lake. Its coherent stratigraphy defies an ob-
vious interpretation; however, the persistence of associated
minor diatom taxa before, throughout, and after the S. "car-
conensis" zone indicates there was no major change in lake
ecology. As with other large species of Stephanodiscus, the
morphology of S. "carconensis" indicates a planktonic habitat,
160 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
but it provides no evidence about seasonal dynamics. This
taxon is now apparently extinct; rare records of S. carconensis
Grunow from modern diatom collections probably represent
contamination or misidentification (Bradbury, 1991).
Other paleolimnological and paleoecological data provide
insights into the limnology of Owens Lake when S. "carconen-
sis" predominated (43-25 Ka). High organic-carbon values
(2%-4%) in this interval (Benson et al., 1996) as well as the
great abundance of diatoms in the sediment (Bradbury, 1997a)
indicate lake productivity had been moderate to high and may
have resulted in periodic bottom-water anoxia. The absence of
ostracodes in samples from the S. "carconensis'" zone also
could reflect seasonal anoxia in Owens Lake at that time. Lo-
wenstein et al. (1997) reported halite crystallization air temper-
atures at 34-25 Ka (obtained from core samples from Death
Valley) that were 8?-15? C below modern air temperatures in
Death Valley. These lower crystallization air temperatures
likely reflect the cool temperatures during springs and sum-
mers when halite generally crystallizes from this system. The
modest heavy-mineral content between 43 and 27 Ka in cores
OL-90-1+2 (Figure 11) indicates decreased alpine glacial ac-
tivity compared with later in the record. Consequently, the vol-
ume of Owens River discharge may have been low during this
period. Pollen evidence (Woolfenden, 1996) also indicates
cool, but relatively dry, summer climates prevailed, whereas
the presence of Yucca brevifolia Engelman (Joshua tree) in as-
sociation with Juniperus osteosperma (Torrey) Little in the Al-
abama Hills packrat midden at 31 and 26 Ka (Koehler and
Anderson, 1994) indicates warmer winters than today, at least
at those times. Together, these data indicate that from 43 to 25
55
OL-90/1+2
magnetite
OL-84/B
monazite + zircon magnetite monazite + zircon
5 0
grains / mm traverse
12 0
grains / mm traverse
FIGURE 11.?Concentration of heavy-mineral grains versus age in cores OL-90-1+2 and OL-84-B.
NUMBER 33 161
Ka Owens Lake probably stratified during the summer after
productive fall and spring seasons when lake circulation would
have provided abundant nutrients to support large diatom
blooms. Perhaps S. "carconensis" predominated during the fall
circulation phase in monomictic Owens Lake just as the simi-
lar-sized S. niagarae does in Crowley Reservoir today.
By 24 Ka, Stephanodiscus asteroides and CycloStephanos sp.
returned to predominance, whereas S. "carconensis" decreased
to low numbers and Aulacoseira species disappeared (Figure
9). Melosira variansl Agardh became common throughout the
remainder of the record (Figure 10). The peak in abundance of
Stephanodiscus asteroides and the presence of large numbers
of heavy-mineral grains in the sediments at about 20 Ka mark
the existence of full-glacial conditions (Figure 11). At that
time, Owens Lake probably had sufficient winter ice cover to
behave as a dimictic lake, with the water circulating after ice
breakup and again before freezing, as do many temperate lakes
today. The diatom assemblages in Owens Lake during the full-
glacial period are typical of species that have late winter (dur-
ing and after ice breakup) and spring production. Presumably,
glacier meltwater entering the basin during the summer and fall
rendered the lake too turbid to support diatom photosynthesis
during those seasons (Bradbury et al., 1994). The diatom iden-
tified as Melosira sp. cf. M. varians may represent a thick-
walled form, perhaps a resting spore, of extant M. varians. This
form apparently has not been found in modern aquatic systems,
although a similar diatom has been found in late Pleistocene
deposits at Kenai, Alaska (McLaughlin and Stone, 1986, pi. 7:
fig. 114). Melosira sp. cf. M. varians probably occupied the
same habitat as M. varians, which today lives in shallow, lit-
toral regions of lakes and low-gradient reaches of streams and
rivers. The variant's presence in the Owens Lake stratigraphy
during the full- and late-glacial periods probably relates to in-
creased input from shallow-water riverine and lake-margin wa-
ter sources.
Late-glacial limnologic changes at Owens Lake (18-15.5
Ka) were marked by an increase in Asterionella formosa. This
increase in abundance implies a favorable open-water environ-
ment (less turbid) existed during the summer and early fall, ac-
cording to the distribution of this species in Sierra Nevada
lakes today. For example, A. formosa numerically dominated
plankton net hauls in Mary, Lundy, and Grant Lakes during
June 1996, whereas it was numerically subdominant (38%) in
the September 1994 collection in Upper Twin Lake (J.P.B., un-
published data).
A hiatus interrupts the OL-90-1+2 sediment record at 15.5
Ka. Sedimentation resumed shortly before 13 Ka with abun-
dant Cyclotella ocellata and Stephanodiscus niagarae (Figures
9, 10), indicating a correlation to the post-hiatus interval of
core OL-84-B.
OSTRACODES.?The OL-90-1+2 ostracode record begins at
about 52.5 Ka with large numbers of Limnocythere ceriotube-
rosa and!, sappaensis (Figure 12). These species are indica-
tors of alkaline, saline, shallow water, and their presence sug-
gests that the climate of Owens Lake and its immediate
drainage was arid to semiarid at that time. Under such climates,
strong, basin-floor evaporation rates, and decreased snowpack
in the Sierra Nevada, the Owens River would have been unable
to maintain a deep and overflowing lake system. Like today, al-
kaline-rich springs along the Owens River and around the lake
would have contributed substantially to the hydrochemistry of
the lacustrine environment. Closed-basin conditions at this
time also were documented by geochemical data in another
Owens Lake core (Bischoff et al., 1997a).
Limnocythere bradburyi Forester, a warm-climate, hydro-
chemical and ecological equivalent of L. ceriotuberosa, be-
came common during one short interval after the predominance
of L. sappaensis. Its abundance documents a brief period (~300
yrs) when the lake became deeper as a result of increased dis-
charge from the Owens River. Because this ostracode is cold
sensitive (Forester, 1985; Smith and Forester, 1994), it indi-
cates that increased discharge began when climates were still
warm (winter temperatures above 0? C).
At 52 Ka, Limnocythere ceriotuberosa and Candona caudata
characterize the ostracode record. Their abundance documents
an abrupt increase in Owens River discharge into the basin
(Figure 12). As lake levels rose and salinity decreased, the vari-
able lacustrine environments preferred by L. ceriotuberosa dis-
appeared, and C. caudata and small numbers of Limnocythere
sp. aff. L. itasca Cole remained. The change in abundance from
L. ceriotuberosa to C. caudata identifies a change from highly
seasonal discharge to a more continual discharge. The more
consistent discharge probably, in part, reflected low summer air
temperatures, which caused slower snow melt and decreased
evaporative loss, and resulted in a spilling lake that had a stable
volume. Candona caudata and Limnocythere sp. aff. L. itasca
make pulse-like reappearances in the record for the next 10 Ky,
approximately matching the minor fluctuations of Stephanodis-
cus asteroides described above. Although not perfectly cyclic,
the average recurrence times for these pulses is 1.4 Ky, match-
ing the millennial-scale climate oscillations recorded in high-
resolution marine cores from the North Atlantic Ocean (Bond
et al., 1997). In this case, the pulses of C. caudata and, to a
lesser extent, of Limnocythere sp. aff. L. itasca apparently mark
brief episodes of cooler temperatures and increased river input
to Owens Lake.
Between 40 and 24 Ka, the Owens Lake record lacks ostra-
codes (Figure 12), perhaps because unfavorable (anoxic) lake-
bottom conditions precluded ostracodes or dissolved their cal-
careous carapaces in the sediments. At 24 Ka, however, Can-
dona caudata reappears in the core with large numbers of
Cytherissa lacustris. This assemblage persisted until about 21
Ka and reappeared briefly in small numbers at about 19 Ka.
These ostracodes, particularly C. lacustris, indicate the limno-
climatic conditions were very cold but reasonably stable. Ac-
cording to the modern distribution of C. lacustris, its presence
in Owens Lake during the full-glacial period implies that polar
air masses were present year-round at the latitude of Owens
162 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
10-,
15
20
25
30
35
40-
45-
50-
55
0 100 200 0 60 0 100 0 50 0 20 0 5
valves / gram
FIGURE 12.?Concentration of ostracode valves versus age in cores OL-90-1+2 from Owens Lake.
Lake. The residence of such cold air masses would have main-
tained seasonally stable climatic and limnologic environments.
CORE OL-84-B
DIATOMS.?Freshwater planktonic diatoms are predominant
in core OL-84-B from the base of the core (extrapolated age of
17.3 Ka) until the break in sedimentation at 14.7 Ka (Figures
13?15). Stephanodiscus asteroides, Astehonella formosa, and
Cyclostephanos sp. characterize this assemblage (Figure 13)
and indicate that fresh, open-water limnological conditions pre-
vailed in the spring (S. asteroides), summer (A. formosa), and
winter to early spring (Cyclostephanos). The large Stephano-
discus species, S. "carconensis," also is present in this late-gla-
cial assemblage, but it may have been reworked from sedi-
ments surrounding Owens Lake that predated the full-glacial
period (Figure 9).
Freshwater benthic diatoms also are conspicuous in the late-
glacial part of core OL-84-B and imply there was substantial
water input from littoral, marginal areas of the ancient lake
(Figure 13). Collectively, benthic diatoms make up 46 percent
of the late-glacial assemblage. The relationship between lake
depth and benthic diatom percentages among lakes in the east-
ern Sierra Nevada has considerable scatter (Figure 16) because
of the variable efficiency of the lakes' currents to redeposit
benthic diatoms. Nevertheless, shallow lakes generally contain
greater numbers of benthic diatoms because illuminated, shal-
low-water habitats characterize these systems. The high per-
centages of benthic diatoms in the post-glacial period of core
OL-84-B imply lake depths generally were shallower than
30 m; these freshwater benthic diatoms also indicate that
Owens Lake must have spilled at this time. The shallow lake
depths probably resulted from high, full-glacial sedimentation
NUMBER 33 163
-(
20 0 20 0 10 0 10 0
percent
20
FIGURE 13.?Percentages of diatom ecologic groups and selected benthic diatoms versus age in core OL-84-B.
See Figure 10 for explanation of silhouette headers.
rates, which decreased the basin's water capacity after the
Tioga glaciation. An erosional hiatus between 14.7 and 13.5
Ka, as reflected in the sediments from cores from this part of
the Owens Lake basin, may have resulted from increased resus-
pension of sediments from the bottom of the shallow lake. Ben-
son et al. (1996) suggested lake desiccation and subsequent de-
flation as the cause of this erosional hiatus. Although diatom
stratigraphy indicates modest lake depths, it does not document
the increased salinity that would have preceded desiccation;
however, deflation may have removed all traces of that epi-
sode. Alternatively, a flood of fresh, sediment-laden water
could have scoured previously deposited sediments from the
Owens Lake basin (J.P. Smoot, pers. comm. to J.P.B., 1997). A
sediment unit called the Parting Mud in Searles Lake, which is
down-river from Owens Lake, documents the continuous depo-
sition of dolomite, aragonite, and calcite laminae between 17.1
and 10.5 Ka (Smith, 1979; Benson et al., 1990), which would
have required at least some overflow from Owens Lake. The
origin and importance of the hiatus needs further study. The
widespread hiatus in Owens Lake records may reflect a combi-
nation of morphometric, climatic, hydrologic, and possibly tec-
tonic events.
After the erosional hiatus (13.4 Ka), the Owens Lake basin
received silty sediment containing the same freshwater plank-
tonic diatom species present before the hiatus. Oscillating peaks
of Stephanodiscus asteroides characterize this assemblage.
Stephanodiscus niagarae initially became abundant, and Cyclo-
tella ocellata increased shortly thereafter (Figure 14). These di-
atoms currently bloom in the summer or early fall in Sierra Ne-
vada lakes (Figure 8) and indicate that Owens Lake was less
164 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
FIGURE 14.?Percentages of selected planktonic freshwater diatoms versus age in core OL-84-B. See Figure 10
for explanation of silhouette headers.
turbid and received more continuous Owens River discharge
during the summer and fall than it did before 13.4 Ka. Melting
alpine glaciers during the summer may have contributed sub-
stantially to the Owens River discharge. This lacustrine phase
lasted until shortly after 11.0 Ka and may document the final
disappearance of large alpine glaciers in the Sierra Nevada.
Two benthic freshwater diatoms, Nitzschia radicula Hustedt
and Cymbella formosa Hustedt, characterize the sediments ini-
tially deposited in Owens Lake after 13.4 Ka (Figure 13). Cym-
bella formosa is morphologically and ecologically similar to
Cymbella muelleri, which is widely distributed in shallow-wa-
ter habitats of lakes and in streams east of the Sierra Nevada
(Figure 7). This species usually lives firmly attached to the
submerged stems of aquatic vegetation or to other suitable sub-
strates (wood, stones) in mucous mats or coatings as an adapta-
tion to resist dislodgment by currents (Peterson, 1996). Nitz-
schia radicula lives among submerged plant stems on the
bottom of shallow, macrophyte-rich ponds, marshes, and small
lakes (Cumming et al., 1995).
The presence of these benthic diatom taxa in core OL-84-B
after 13.4 Ka implies there was substantial hydrologic input to
Owens Lake from the Owens River and marginal habitats (pos-
sibly marsh environments at the head of Owens Lake or up-
stream in the Owens Valley), because these diatom species
would not have lived in large numbers in the deep, open-water
habitats of the lake itself. Perhaps these redeposited riverine
and marsh diatoms recorded the diminishing effects of a flood
that eroded the bottom of Owens Lake and initiated the hiatus.
After 12.5 Ka, planktonic freshwater diatom assemblages
predominated until 10.9 Ka when saline, shallow-water diatom
NUMBER 33 165
30 0
percent
FIGURE 15.?Percentages of selected saline diatoms versus age in core OL-84-B. Saline benthic species include
Amphora coffaeiformis, Anomoeoneis sphaerophora, Navicula crucialis, N. halophila (Grunow) Cleve, Nitz-
schia fmstulum, Rhopalodia gibberula (Ehrenberg) O. Miiller, R. acuminata, and Surirella hoefleri.
to:
Ar\ '
IK '
20 -i
0-
I 1 ?
?
?
? ^
?
?
???
0 10 20 30 40 50 60 70 80 90 100
benthic diatoms in surface sediment (%)
FIGURE 16.?Percentage of benthic diatoms in surface sediment versus water
depth in selected Sierra Nevada freshwater lakes.
assemblages suddenly appeared (Figure 15). Campylodiscus
clypeus, a benthic species of alkaline and moderately saline
habitats, predominated; the assemblage also included small
percentages of benthic Surirella ovalis and S. striatula and
planktonic Cyclotella meneghiniana and C. quillensis. Collec-
tively, this assemblage indicates lake salinities were 3-10 g/L
and alk/Ca ratios were >4. This short-term (500 year) saline
event ended by 10.4 Ka when freshwater planktonic diatoms
returned to predominance for about 800 years (Figures 14, 15).
This freshwater phase in Owens Lake may correlate with sedi-
mentological evidence for an early Holocene lake at Silver
Lake playa, about 200 km southeast of Owens Lake, that ex-
isted just before 9.3 Ka (Enzel et al., 1989).
After 9.6 Ka, the record of saline diatoms in Owens Lake be-
comes more continuous (Figure 15). A short interval (9.4-8.5
Ka) of mixed freshwater and saline diatom assemblages, includ-
ing large numbers of Chaetoceros sp., documents a transition to
166 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
higher lake salinity and indicates arid climate. Today, Chaetoc-
eros is an open-water planktonic diatom that is common in Pyr-
amid and Walker Lakes (Figure 6). Its wide salinity tolerance
(Fritz et al., 1993) may indicate that Owens Lake remained
fairly deep during this short interval. The freshwater planktonic
diatom {Stephano discus spp.) component of this assemblage
(Figure 14) could have prospered in the lake seasonally when
spring runoff sufficiently diluted salinities or during short peri-
ods when the lake spilled and became fresh.
After 8.5 Ka, benthic saline species, such as Campylodiscus
clypeus, persisted in the extensive shallow-water habitats cre-
ated by low lake levels. By 8 Ka, saline Nitzschia species (Nitz-
schia latens, N. pusilla, N. monoensis), which presently inhabit
very saline, shallow-water habitats in Mono Lake and similar
environments (Kociolek and Herbst, 1992; Blinn, 1993), be-
came prominent in response to a substantial decrease in Owens
River flow, a proportional increase in alkaline spring discharge,
and high evaporation rates. Owens Lake remained saline and
shallow until about 6.5 Ka when, according to the radiocarbon
chronology, the lake desiccated and clastic sedimentation
ended (Benson et al., 1997). Sedimentation of oolitic sand re-
sumed after 4.3 Ka, and freshwater and saline diatoms are
sparsely represented in core OL-84-B by fragments of corroded
valves (Figures 13-15). The presence of freshwater planktonic
and benthic diatoms in samples from after about 8 Ka may
have resulted from the reworking of diatomaceous highstand
deposits left behind as Owens Lake shrank. These reworked di-
atoms do, however, indicate episodic fluxes of water into the
basin that may reflect flooding during very wet years (Smith et
al., 1987). Saline benthic diatoms, including the Nitzschia spe-
cies of very saline habitats, returned between 4 and 3 Ka.
In-situ diatom productivity certainly occurred throughout the
late Holocene in Owens Lake, but, generally, high lake alkalin-
ities and salinities combined with episodic desiccation likely
destroyed sedimented diatoms by dissolution and breakage.
OSTRACODES.?Except for isolated, presumably reworked
valves of Cytherissa lacustris from highstand, full-glacial de-
posits (Figure 17), core OL-84-B has no substantial late-glacial
(15?12.5 Ka) ostracode record. Decreased inflow from the
Owens River and a climate regime of increased summer insola-
tion would have promoted lake stratification, and seasonal an-
oxia may have eliminated ostracode habitats in the central re-
gions of the Owens Lake basin. The presence of ostracodes,
however, would have indicated that shallow and saline lake
conditions existed between 15.5 and 13.5 Ka, but those depos-
its could have been removed by erosion.
At 11.5 Ka, Limnocythere ceriotuberosa appeared (Figure
17); its presence indicates this was a time of variable lacustrine
conditions associated with seasonal filling and freshening of
Owens Lake that was followed by lake evaporation, concentra-
tion, and shrinkage. Larger pulses of L. ceriotuberosa, accom-
panied by Candona caudata, occurred from 11.0 to 10.5 Ka
and from 9.9 to 9.5 Ka. These ostracode pulses approximately
coincided with the shallow-water, moderately saline lake envi-
ronments evidenced by the diatoms, particularly "Fragilaria,"
Campylodiscus clypeus, Surirella spp., and Cyclotella spp.
(Figures 13-15). The maximum salinity tolerance of Candona
caudata (3 g/L) and the low (or absent) representation of Lim-
nocythere sappaensis also indicate Owens Lake had only mod-
est salinity at those times, although the diatoms and L. ceriotu-
brosa can tolerate considerably higher salinities. These pulses
in the fossil record probably record episodes when the Owens
River flooded the basin to produce short-lived, reasonably
fresh, shallow lake systems. Near the mouth of the Owens
River, at an elevation of 1097 m, coarse deltaic sands that have
a freshwater molluscan fauna, including the river clam An-
odonta californiensis Lea (dated at 11.4 Ka), also indicate riv-
erine environments occurred at the north end of Owens Lake
(Koehler, 1995).
The large number of Limnocythere sappaensis by 9.4 Ka in-
dicates increased salinity and a high alk/Ca ratio. The lake may
have become shallower, although this ostracode can live in
deep lakes as well. The absence or low abundance of ostra-
codes between 9.4 and 8.8 Ka in the section of core having
abundant Chaetoceros may indicate a deeper lake was present
at that time, perhaps with seasonally anoxic bottom water.
Progressively larger pulses of Limnocythere ceriotuberosa
characterized the millennium between 8.8 and 7.8 Ka (Figure
17), documenting a return to lacustrine environments that had
variable depths and salinities. Large quantities of Heterocypris
appeared near the end of this episode. Today, this species re-
sides in springs and seep-fed marsh environments. Its presence
between 8.8 and 7.8 Ka implies that Owens Lake had shrunk
considerably and was surrounded, at least near the mouth of the
Owens River, by marsh communities. The marshes probably
were dominated by Scirpus, which presently grows abundantly
along the Owens River, within the valley, and along the mar-
gins of playa springs (Hollermann, 1973).
Limnocythere sappaensis dominated the ostracode assem-
blage from 7.7 to 6.5 Ka. Occasional Cytherissa lacustris in
this interval resulted from the reworking of highstand lake sed-
iments deposited during the last glacial maximum (Carter,
1997). During this time, saline, alkaline springwater issuing
along the Owens River and around Owens Lake probably had a
major effect on the hydrochemistry of the lake, although sur-
face water from Sierra Nevada snowmelt continued seasonally
to enter the basin. Nitzschia pusilla and N. monoensis (Figure
15), which tolerate highly saline environments, such as at
Mono Lake today, coexisted with L. sappaensis.
Limnocythere sappaensis remained prominent after the hia-
tus in sedimentation (6.1-4.3 Ka) (Figure 17), as did benthic
saline diatoms and Nitzschia spp. (Figure 15). These saline lake
assemblages persisted until about 3 Ka, when diatoms occurred
only as rare, reworked fragments and ostracodes disappeared
from the Owens Lake record; however, in core OL-92-3, oolitic
sediment, radiocarbon dated between 5.3 and 5.0 Ka (Bischoff
et al., 1997c), had abundant L. sappaensis, which indicates
NUMBER 33 167
Y*?
hiatus
50 0 10 0 1 0200 0 10 0
valves / gram
FIGURE 17.?Concentration of ostracode valves versus age in core OL-84-B.
300
some sedimentation occurred in Owens Lake during the middle
Holocene drawdown, at least at the site of core OL-92-3.
Discussion
The diatom and ostracode records from the middle Wiscon-
sin to the late Holocene at Owens Lake provide the most con-
tinuous paleontological evidence for changing limnological
and climatic environments in the southern Great Basin. Be-
cause this important record relates to the glacial and vegeta-
tional history in this region, it can provide a chronological
framework for local glaciation and for isolated vegetation as-
semblages found in packrat middens. Conversely, the nonla-
custrine evidence of climate change from packrat middens and
glacial records can refine climate interpretations from lacus-
trine microfossils.
RECORD OF GLACIATION.?Owens Lake cores OL-90-1+2
and OL-84-B record the timing and intensity of Sierra Nevada
alpine glaciation in the Owens River basin. The substantial
quantity of heavy-mineral fragments (magnetite and zircon +
monazite) in diatom slides (Figure 11) indicates that glacial
erosion of the Sierra Nevada granitic plutons occurred. These
minerals were incorporated into the base load of the Owens
River. The middle Wisconsin period (~54 to -25 Ka), or oxy-
gen isotope stage (OIS) 3, for the most part, had small, isolated
heavy-mineral peaks that may represent minor high-elevation
glaciation at 52-51, 44, 40, 35, and 26 Ka. The absence of the
boreal ostracode Cytherissa lacustris during these periods,
however, indicates that severe glacial climates did not affect
Owens Lake. In fact, the presence of the warm-season diatoms
Aulacoseira spp. during most of these periods seems inconsis-
tent with extensive glaciation. These minor peaks in heavy-
168 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
mineral concentration may simply have recorded years or de-
cades of exceptional runoff down the Sierra Nevada drainages.
Nevertheless, geochemical evidence of glacial-rock flour from
Owens Lake core OL-92-1 (Bischoff et al., 1997b), taken in the
same area of the basin, approximately correlates with most of
the heavy-mineral peaks shown in Figure 11 and supports the
interpretation that these peaks represent minor Sierra Nevada
glaciations.
The heavy-mineral peaks (Figure 11) approximately corre-
late with the magnetic-susceptibility and organic-carbon evi-
dence of glaciation recorded by Benson et al. (1996). The mag-
netic-susceptibility and organic-carbon records, however,
contained more peaks, which were interpreted by Benson et al.
(1996) as glacial advances, than are present in the heavy-min-
eral record. Uncertain equivalencies between radiocarbon and
calendar years make correlations to 36Cl-dated moraines (Phil-
lips et al., 1996) problematic; however, the potential glacial ad-
vance at about 51 Ka, which perhaps correlates to the younger
Tahoe moraines, also may correlate with the double peak in
heavy minerals between 52 and 51 Ka (Figure 11). The Tioga 1
stage moraines, which were dated to 32-30 calendar Ka (Phil-
lips et al., 1996), may correspond to heavy-mineral peaks at
27-25 Ka in core OL-90-1+2, assuming an approximate 4 Ky
correction to equate the radiocarbon and calendar chronologies.
The episodes defined by high concentrations of heavy minerals
generally correspond to perennial lake stages in Searles Lake.
The mineralogical evidence of a 5? C cooling in Searles Lake
at 26-25 Ka (Smith, 1979:89) might corroborate the occur-
rence of a glacial period at that time.
According to the heavy-mineral record from cores OL-90-
1+2 (Figure 11), glaciation during OIS 2 (the last major glacial
period) began by 23 Ka. The heavy minerals continued to in-
crease, but with sharp fluctuations, until about 18 Ka when
their values decreased, again with strong fluctuations between
greater and lesser quantities. In contrast to the Tioga 1 mo-
raines, the Tioga 2 and 3 (glacial maximum) moraine dates
(Phillips et al., 1996) correspond closely to the heavy-mineral
record of glaciation. The large and persistent peak in magnetite
grains centered on 22 Ka probably corresponds to the extensive
Tioga 2 moraines identified and dated by Phillips et al. (1996).
The several sharp fluctuations in heavy-mineral concentrations
afterward may provide much more detail than the Tioga 3 mo-
raine sequences, but transportation of heavy minerals may have
occurred during episodes of glacial advance and retreat. The
sharp fluctuations in heavy-mineral concentrations correlative
to the Tioga 3 glacial deposits generally resemble and partly
correlate with the sharp fluctuations in the magnetic-suscepti-
bility record reported by Benson et al. (1996). The Tioga 4 gla-
cial advance at about 15 Ka does not stand out as an especially
distinctive peak of large concentrations of heavy minerals in
cores OL-90-1+2, although such a peak exists at this time in
core OL-84-B (Figure 11).
CLIMATE AND LIMNOLOGICAL EVIDENCE FROM PACKRAT
MIDDENS.?Radiocarbon-dated packrat middens near Owens
Lake document paleoclimate records from fossil vegetation as-
semblages. The closest midden to Owens Lake is from the
northeast side of the lake (Figure 1), about 10 m above the
highstand elevation (1145 m). Juniperus osteosperma domi-
nates the full-glacial levels of this midden (23.0-14.5 Ka), but
J. scopulorum (Rocky Mountain juniper) and Pinus mono-
phylla Torrey and Fremont form large components of the as-
semblage at midden levels corresponding to 17 Ka (Koehler
and Anderson, 1994).
Juniperus scopulorum Sargent now lives 2100-2900 m
above Owens Lake (Koehler and Anderson, 1994), but it lived
at least 900 m lower during the full-glacial period. The pres-
ence oiP. monophylla and J. osteosperma at lower elevations
indicates that 17 Ka precipitation was greater by a factor of two
to three and temperature was about 7? C lower than today, ac-
cording to modern lapse rates (Smith and Bischoff, 1997). To-
day in the Great Basin, pinyon-juniper woodland vegetation
does not characterize regions where mean January tempera-
tures fall below -1? C (Thompson, 1990). Modern mean Janu-
ary temperature at Independence, California, 100 m above
Owens Lake, is about 5? C (Figure 2), so a 7? C decrease would
be somewhat below the tolerance of pinyon-juniper woodland.
A full and overflowing Lake Owens, however, could have
served as a local thermal buffer against very low (and high)
temperatures and thereby may have created suitable conditions
for the establishment and growth of P. monophylla and J.
osteosperma during the full-glacial period.
Packrat midden sites in the Alabama Hills, at elevations of
1264-1460 m and about 17 km northwest of Owens Lake (Fig-
ure 1), extend the fossil vegetation record of the Owens Valley
to 31.4 Ka (Koehler and Anderson, 1995). Full-glacial middens
in the Alabama Hills do not contain P. monophylla, perhaps be-
cause of the elevation of this locality. Nevertheless, the pres-
ence of Yucca brevifolia (Joshua tree) and abundant J. osteo-
sperma at 31.4, 25.7, and 21.1-17.8 Ka indicates winter
temperatures generally were above freezing. This association
now grows on the east side of Walker Pass, about 90 km south-
west of Owens Lake, at an elevation of about 1500 m. As sug-
gested above, perhaps the large body of water in the Owens
Lake basin during this episode helped mitigate severe winter
temperatures in the Alabama Hills. The cold temperature epi-
sode from 24 to 21 Ka, implied by the presence of the boreal
ostracode Cytherissa lacustris (Figure 12), fell between epi-
sodes of mild winter temperatures, as evidenced by Y. brevifo-
lia in the Alabama Hills packrat middens.
The principal climate change documented by macrofossils
from packrat middens in the Alabama Hills occurred at 13.3 Ka
when most elements of the modern Sarcobatus scrub plant
community appeared (Koehler and Anderson, 1995). Juniperus
osteosperma, however, persisted in the midden sequence, indi-
cating cooler and wetter conditions, compared with modern en-
vironments, prevailed until 9.5 Ka. Diatoms and ostracodes in
the late-glacial and early Holocene Owens Lake indicate that
freshwater and presumed overflowing lake conditions existed
NUMBER 33 169
until 11 Ka. Afterward, lake levels and salinity fluctuated, re-
flecting variable, but decreasing, input from the Owens River.
The first major appearance of Limnocythere sappaensis, an in-
dicator of shallow and saline aquatic habitats, coincided with
the disappearance of J. osteosperma and with the complete
dominance of desert shrubs and herbs in the middens in Ala-
bama Hills at 9.5 Ka.
MIDDLE WISCONSIN LINKAGES TO MARINE RECORDS ON
THE CALIFORNIA MARGIN.?Isotopic, magnetic, and geochem-
ical records from Owens Lake have been tentatively correlated
with climate changes recorded in North Atlantic Ocean marine
sediments (Benson et al., 1996) and Greenland ice cores (Ben-
son et al., 1997). These correlations could imply that hemi-
spheric-scale climate changes affected far-separated regions,
but the interactive climatic and oceanographic mechanisms that
might link changes in Owens Lake to changes in the North At-
lantic Ocean and in Greenland were unspecified.
Gardner et al. (1988) proposed a marine-terrestrial paleocli-
mate linkage for the past 20 Ky between the lacustrine record
at Clear Lake in California (latitude 39?N) and an adjacent ma-
rine record (core V1-80-P3) on the continental slope off north-
ern California (latitude 38.43?N) (Figure 1). Benthic coastal
and reworked freshwater diatoms in the marine core from the
full-glacial period document lower sea levels and greatly de-
creased deep-water upwelling than today. These marine condi-
tions correlate with cooler climates in coastal California where
Pinus was the dominate vegetation. An episodic, fluctuating
transition to marine planktonic diatoms indicates that up-
welling off the California coast characterized the establishment
of Holocene conditions. On land, warming temperatures fa-
vored abundant growth of Quercus (oak) and the local presence
of Sequoia (redwood). Coastal fogs produced by upwelling
may have been responsible for providing suitable conditions
for Sequoia.
Upwelling of nutrient-rich marine water on the California
margin relates to the strength and persistence of the Hawaiian or
Subtropical High, whereas increased winter and spring precipi-
tation that causes runoff to California coastal regions reflects
the predominance of the Aleutian Low. These two pressure sys-
tems affect present-day terrestrial climates and ocean dynamics,
although their effects vary from north to south (Gardner et al.,
1997). Their important role in glacial, Holocene, and modern
marine and terrestrial climate records in western North America
implies that similar variations in the strength of the Aleutian
Low and the Hawaiian High may also account for earlier pale-
oceanographical and paleolimnological changes.
The middle Wisconsin, or OIS 3, is represented by the ma-
rine core ODP-893 (latitude 34.29?N) from the Santa Barbara
Basin (Kennett and Venz, 1995) and by core F2-92-P34 (lati-
tude 35.03?N) from the California margin off Point Conception
(Gardner et al., 1997) (Figure 1). Both cores have sufficient
chronological control for comparison to the Owens Lake pale-
olimnological record between 55 and 20 Ka. Perhaps the most
distinctive paleolimnological evidence for climate change dur-
ing this interval in the Owens Lake core is the distribution of
the summer-season planktonic diatoms Aulacoseira granulata
and A. ambigua. The most logical explanation for the presence
of these species relates to past episodes of exceptional snowfall
in the Sierra Nevada. During these episodes, high-volume run-
off down the Owens River would have persisted into late sum-
mer or fall; this would have provided suitable freshwater lim-
netic conditions for Aulacoseira blooms. Intervals in which
these diatoms were uncommon or absent represent, by this rea-
soning, episodes in which precipitation was less but still suffi-
cient to maintain freshwater in Owens Lake, at least during
spring and early summer.
Using extrapolations from uncorrected radiocarbon chronol-
ogies for marine core F2-92-P34 (Gardner et al., 1997) and for
Owens Lake, a correlation between small fluctuations in the
percent of aluminum present and the abundance of summer
Aulacoseira species is evident (Figure 18). Increased concen-
trations of aluminum and other geochemical evidence of detri-
tal input within the coastal marine cores indicate greater runoff
from the Coast Ranges and the Sierra Nevada (Dean et al.,
1997), and the same process may explain the approximately co-
eval proliferation of Aulacoseira in Owens Lake.
Gardner et al. (1997) suggested that increased fluxes of detri-
tal material to the coastal sediments during the full-glacial pe-
riod were caused by the same moisture source that supplied the
glaciers in the Sierra Nevada and the lakes in the Great Basin.
An intensification and southerly displacement of the Aleutian
Low and the storms associated with the Polar Front jet stream,
probably forced by a southward expansion of polar air masses,
are the most probable reasons for this moisture source. The
same climatic mechanisms can be used to explain similar, but
less intense, variations during OIS 3.
In core F2-92-P34, intervals that are not characterized by in-
creased percentages of aluminum instead contain evidence of
increased anoxia, such as high values of molybdenum and or-
ganic matter (Dean et al., 1997). In other presumably correla-
tive marine cores, anoxic conditions excluded burrowing ma-
rine organisms, and zones of laminated sediments resulted.
Increased productivity (upwelling) of the California Current
apparently provided sufficient marine productivity and organic
material that, once deposited, promoted the formation and
maintenance of anoxic environments.
The history of upwelling intensity of the California Current
also can be deduced by the fluctuating presence of the plank-
tonic foraminifera Globigerina bulloides D'Orbigny in a core
from ODP-893 in the Santa Barbara Basin (Kennett and Venz,
1995). According to the oxygen-isotope chronology developed
from this core, fluctuations of G. bulloides (Figure 18) were
antithetical to the aforementioned evidence of increased pre-
cipitation {Aulacoseira and Al), suggesting that the same cli-
matic conditions that increased upwelling also decreased win-
ter precipitation along the California coast. In this case, the
strength and persistence of the Hawaiian High was the likely
candidate because clockwise winds generated around this high-
170 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
20-
Aulacoseira (summer) % Globigerina bulloides
20
55 55
Point Conception
F2-92-P34
Owens Lake
OL-90/1+2
0 50
Santa Barbara Basin
ODP #893
FIGURE 18.?Chronological comparison of middle Wisconsin episodes of abundant Aulacoseira species from
Owens Lake cores OL-90-1+2, percentage fluctuations of Al from core F2-92-P34 on the California margin west
of Point Conception, and percentage of Globigerina bulloides from ODP site #893 in the Santa Barbara basin
(Figure 1). Aluminum data are from W.E. Dean (USGS). Foraminifera data are from Kennett and Venz (1995).
pressure center would have driven the California Current to the
south, which is what happens today. Because this atmospheric
circulation pattern would have blocked moisture from the trop-
ical Pacific Ocean, it also would have reduced the annual pre-
cipitation reaching the Sierra Nevada and, consequently,
Owens Lake.
Admittedly, correlations of the paleoceanographic and lacus-
trine evidence for past climate change in cores from coastal
California and from Owens Lake are not perfect, and they rely
on chronologies of uncertain accuracy. Nevertheless, this inter-
pretation is based upon operationally consistent climatic, ter-
restrial, and oceanographic systems and processes that logi-
cally force large changes between glacial and interglacial
environments.
Conclusions
High-resolution diatom and ostracode records from cores
from Owens Lake have revealed a sensitive and comparatively
simple record of Sierra Nevada precipitation for the last 55 Ky.
At one extreme, high precipitation, combined with cold tem-
peratures associated with the southward expansion of polar air
masses, produced glaciation in the Sierra Nevada (Hostetler
and Clark, 1997). The Owens Lake basin filled with detrital
sediment containing the remains of freshwater diatoms and bo-
real ostracodes. At the other extreme, scant winter precipitation
failed to offset high evaporation rates, and the Owens Lake ba-
sin desiccated to playa or near playa conditions. Intermediate
limnological and climatic conditions existed at other times.
These conditions included episodes of saline lakes of shallow
to moderate depths and regimes of greater than modern precip-
itation under temperatures warm enough to preclude extensive
glaciation. During the last 50 Ky, the predominant paleocli-
matic scenario at Owens Lake included greater precipitation
than at present and probably cooler than modern temperatures.
During this time, overflow conditions existed for about 40 Ky
(50-10 Ka), and only during the last 10 Ky of the Holocene did
Owens Lake remain below its spill level for long periods of
NUMBER 33 171
time. The paleolimnological record of the last 50 Ky is consis-
tent with the longer record of Owens Lake reported by Brad-
bury (1997a, 1997b), which documents overflow conditions for
about 80 percent of the last 800 Ky.
The high-resolution study of limnological and climatic vari-
ations at Owens Lake also indicates rapid changes between
positive and negative hydrologic balance. For example, the
change from closed- to open-basin conditions about 52 Ka oc-
curred in less than 300 years. Changes in diatom and ostracode
assemblages document the step-like progression of this shift.
Holocene changes from fresh to saline diatom floras about 10
Ka indicate that a shift in the hydrologic balance occurred
within a century or less. Assuming a reasonable chronological
accuracy, this rate of change in salinity was faster than what
Bischoff et al. (1997a, fig. 5) predicted for simple evaporative
concentration of Owens River water during current climates;
thus, saline springwater discharge entering Owens Lake must
have had an important hydrochemical effect during arid cli-
mate periods. The apparent correlation of the Owens Lake re-
cord with the paleoceanographic records indicates that climatic
conditions at Owens Lake were not just local, but were part of
a global system that involved major reorganizations of atmo-
spheric and oceanic circulation. The Owens Lake fossil record
supports the contention that climate change can be rapid and
that future climate in the Owens Lake area could be wetter
rather than drier.
Acknowledgments
S.P. Lund (University of Southern California) is particularly
thanked for providing access to the cores used in this study.
L.V. Benson (United States Geological Survey (USGS)) pro-
vided much of the core chronology, and he and S.P. Lund of-
fered important insights into the history of limnologic and cli-
matic changes in the Owens Lake basin. R.S. Thompson, J.L.
Bischoff, W.E. Dean, G.I. Smith, J.P. Smoot, R.L. Reynolds
(all USGS), and Vera Markgraf (University of Colorado) pro-
vided helpful discussion of points raised in this paper. M.G.
Paquette and K.L. Conrad (both USGS) prepared the samples
for study, and Claire Carter (USGS) identified and counted the
ostracodes. This study is part of the USGS Yucca Mountain Fu-
ture Climates Project and the USGS Climate History Program.
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Biogeography and Timing of Evolutionary Events
among Great Basin Fishes
G.R. Smith, T.E. Bowling, K. W. Gobalet, T. Lugaski,
D.K. Shiozawa, andR.P. Evans
ABSTRACT
Hubbs and Miller's Great Basin works are model studies of the
evolution of isolated endemic aquatic organisms and of hydro-
graphic history. They investigated spatial and temporal barriers
that permitted evolutionary divergence of fishes. They also used
fish evidence to identify the locations and possible timings of past
aquatic connections among basins. Hubbs and Miller focused on
morphological differences among fish populations in related
basins and hypothesized that their evolutionary divergence
occurred over only hundreds to thousands of years. Their conclu-
sions depended upon the assumption that divergence began after
the most recent time of homogeneous ancestral populations, which
was assumed to be the time of closure of the most recent hydro-
graphic connection. This approach and its key assumptions are
standard in studies of evolutionary rates, but they require reexami-
nation.
Our data suggest that rates of morphological change in oligo-
genic and ecophenotypic traits of Great Basin fishes may be rapid
in the short term (thousands of years), but Miocene to Pleistocene
fossils of western North American fishes indicate slow changes
over millions of years. Tests of hypotheses about evolutionary
rates require independent estimates of both time and degree of
divergence. Our evolutionary-rate estimates are based upon
regression of percent sequence divergence (of mitochondrial
DNA) between pairs of lineages against millions of years of diver-
G.R. Smith, Museum of Zoology, University of Michigan, Ann Arbor,
Michigan 48109, USA. T.E. Dowling, Department of Zoology, Arizona
State University, Tempe, Arizona 85287, USA. K. W. Gobalet, Depart-
ment of Biology, California State University, Bakersfield, California
93311, USA. T. Lugaski, deceased, formerly of the Mackay School of
Mines, University of Nevada, Reno, Nevada 89557, USA. D.K.
Shiozawa, Department of Zoology, Brigham Young University, Provo,
Utah 84602, USA. R.P. Evans, Department of Zoology, Brigham Young
University, Provo, Utah 84602, USA.
Reviewers: Robert Hershler, Division of Mollusks, National Museum
of Natural History, Smithsonian Institution, Washington, D.C., 20560-
0118, USA; Edward O. Wiley III, Department of Ecology and Evolu-
tionary Biology, Division of Biological Sciences, University of Kan-
sas, Lawrence, Kansas 66045, USA; and one anonymous reviewer.
gence of sister clades, estimated from dates of first appearances in
the fossil record. The results indicate rates of molecular evolution
in the range of approximately one-half percent (salmonids) to one
percent (cyprinids and cyprinodonts) sequence divergence per one
million years. The ages of some forms thought to represent ten
thousand years of postpluvial evolution are more likely to be hun-
dreds of thousands of years old, according to molecular/fossil esti-
mates. Estimated divergence times of fishes, based upon DNA,
provide unique data on the nature and timing of barriers and
aquatic connections among basins.
Great Basin fish faunas exhibit high endemism, but they have
lower diversity than predicted by pluvial cycles of isolation and by
rapid evolution in allopatry. We hypothesize that uplift causes high
stream gradients and, therefore, decreased local habitat diversity
because of rapid stream runoff. Small and isolated populations on
the basin floors are prone to allopatric differentiation, but they also
are prone to high extinction rates; therefore, isolation and stabiliz-
ing selection in desert basins rarely produced new species that sur-
vived extinction events.
Dedication
This paper is dedicated to Carl L. Hubbs, Robert R. Miller,
and Dwight W. Taylor, whose studies of aquatic organisms in
western North American desert basins set the twentieth century
standard for aquatic biogeography and conservation.
Introduction to Great Basin Ichthyology
Carl L. Hubbs began his studies of fishes in the Great Basin
in 1915 while on an expedition to the Bonneville Basin as a
field assistant to J.O. Snyder (Miller and Shor, 1997). Prior to
the Bonneville trip, Snyder had surveyed fishes in the Lahontan
Basin and in the Oregon Lakes, Owens River, and Mohave
River basins (Snyder, 1908, 1912, 1914, 1917a, 1917b, 1918,
1919, 1924). Hubbs continued to conduct exhaustive explora-
tions of the Great Basin in 1934, 1938, and 1942; with his wife,
Laura, and their children, Earl, Frances, and Clark, he collected
and identified more than 118,000 specimens. Hubbs was joined
by his student, Robert Miller, in 1938, and between them they
175
176 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
produced about three dozen papers on Great Basin fishes.
Miller's efforts resulted in numerous papers on fishes of the
Death Valley system (e.g., 1946a, 1946b, 1950, 1972) and on
broader issues relating to evolution and conservation of Great
Basin fishes (1959, 1961a, 1961b, 1965, 1981). Miller contin-
ued conducting this field work with his wife, Frances Hubbs
Miller, and their five children, and he extended the project to
include a survey of the fishes of Mexico.
The landmark synthesis of the systematics and biogeography
of fishes of the Great Basin was published by Carl Hubbs and
Robert Miller in 1948(b), which was expanded from a sympo-
sium paper read in 1942. The monograph described the geo-
graphical limits of more than 100 hydrographic units, the eleva-
tions of ancient lake levels, and the biogeographic relationships
of fishes in each basin. The authors' goal was to show how
convincingly physiographic deductions are supported by
knowledge of the distribution of fishes in the Great Basin.
The Hubbs-Miller team and Elliot Blackwelder (1933) both
used fish distributions to infer past aquatic connections, which
were based upon the principles of Jordan (1905, chapters 16,
17). The interaction of time and barriers framed the questions
in the Hubbs-Miller inventory of Great Basin fish evolution.
Their descriptions and classification of those fishes were in-
tended to answer questions about hydrographic connections
among basins, barriers between basins, and rapidity of evolu-
tion. Although they chose not to speculate about evolutionary
mechanisms (1908b: 119), they offered frequent hypotheses
about amounts and rates of evolution. They later expanded
upon their ideas about fish speciation (Hubbs, 1940, 1941a,
1941b; Miller, 1946a, 1950).
Hubbs and Miller's most exciting idea was declared in the
introduction to their 1948 monograph (Hubbs and Miller,
1948b:24):
Some of the isolation and speciation that is dated from Pleistocene time may be
much more recent. Now disconnected waters within many of the basins of
western North America may have been connected and their fish fauna made
uniform not many centuries ago. From a consideration of the evidences on cli-
matic changes ... it seems likely that such connections may have been estab-
lished several times during the last three millennia. We believe, however, that
most of the local differentiation and endemism among the fishes of the isolated
desert waters dates from the end of the Glacial period. Hutchinson (1939) con-
cluded that the similar differentiation of now isolated fish populations in Kash-
mir and Indian Tibet dates from the period of maximum glaciation. ...
In correlation with Post-pluvial isolation, a limited degree of speciation has oc-
curred here and there, but data are too meager to warrant any attempt to use the
rate of species evolution in dating the Pluvial waters. ... A considerable body
of observations and mathematical data indicates that small, closely inbreeding
populations may speciate at a much more rapid rate than was previously
thought possible (Wright, 1940a, 1940b, 1942; Huxley, 1943).
The authors treated their geological and ichthyological data
in two ways: they used geological data to infer rates offish dif-
ferentiation, and they used degree of morphological divergence
to infer the times of geological barriers. The usual assumption
was that fish populations were uniform at the end of the most
recent connection between aquatic habitats. This led to the de-
duction that rapid speciation occurred in the postpluvial inter-
val of isolation, about 11,000 years.
Carl Hubbs, Robert Miller, and Laura C. Hubbs (1974) ex-
panded upon their previous work and the work of others in
their subsequent monograph on hydrographic history and relict
fishes of the North-Central Great Basin. Laura Hubbs mea-
sured hundreds of fish specimens from different basins and cal-
culated the mean and standard deviation for the populations in
each basin, providing replicated comparisons of times and
amounts of divergence. Hubbs and Miller (1948a, 1972) de-
scribed many additional localized cyprinid species and subspe-
cies, which provided further evidence for postpluvial diver-
gence in geographic isolation (see Appendix).
Hubbs and Miller sought to answer a question that still moti-
vates biogeographers (e.g., Ricklefs and Schluter, 1993); that
is, What are the appropriate scales of the spatial and temporal
interactions that bring about species production? Hubbs and
Miller (1948b) emphasized differences among and within ba-
sins, namely, different taxonomic levels, numbers of endemic
forms, and local diversity in relation to regional diversity as in-
dicators of evolutionary history: "The Lahontan fauna ... com-
prises, we estimate, about 30 forms, as compared with perhaps
28 in the Bonneville and about 24 in the Death Valley system.
The relatively high number of total forms is correlated with the
isolation of local populations, for the Lahontan system has be-
come greatly disrupted since Pluvial times" (Hubbs and Miller,
1948b:32).
Hubbs and Miller (1948b:41) also recognized an important
role for hybridization in evolution: "Since the waters have
dwindled, and the lake and creek types have been forced into
cohabitation, the two have hybridized very extensively. We
have referred to this specific breakdown in connection with our
analysis of a similar case of hybridization in the Mohave
Desert (Hubbs and Miller, 1943)." The effects of hybridization
in reducing diversity and erasing evidence of evolutionary his-
tory are peculiarly prominent among fishes in the western
United States (Smith, 1992a; Dowling and Secor, 1997).
Our goal is to continue the comparison of fish and hydro-
graphic information so as to answer questions about timing and
origins of fish diversity. Evidence presented herein for timing
of fish divergence comes from the fossil record (G.R. Smith,
1981a, 1981b; Smith et al., 1982) and from molecular studies
(Dowling et al., 2002; Shiozawa and Evans, unpublished). In
this paper, we summarize gene sequence divergence data in the
context of the fossil record of Great Basin fishes. We seek to
shed light on the rates of evolutionary changes in Great Basin
drainages in the spirit of the other Great Basin naturalists of 50
years ago (Hall, 1946:59-67; Durrant, 1952:490-518; LaRiv-
ers, 1962; and the vast body of literature published by Hubbs
and Miller).
Our understanding of freshwater fish distributions benefits
from the fact that freshwater fish are more geographically con-
strained than most other organisms (Cope, 1868; Jordan, 1905).
The principles thought to influence distribution and evolution
NUMBER 33 177
of fishes have remained rather constant during the last 130
years, but estimates of timing of geological and evolutionary
events vary by five orders of magnitude (compare Cope, 1868,
with Hubbs et al., 1974). The Great Basin offers a unique op-
portunity to study the time scales of evolutionary processes be-
cause of its numerous, isolated basins that have replicated fish
habitats and histories. In addition, well-documented evidence
exists for environmental gradients, pluvial cycles, and ages of
sediments that preserve fossil ancestors of Holocene fishes
(e.g., Morrison, 1965, 1991; Benson and Thompson, 1987;
Benson et al., 1990; Oviatt et al., 1992; Reheis and Morrison,
1997; Smith and Bischoff, 1997). These factors provide the ba-
sis for Cope's (1883) and Hubbs and Miller's (1948b) assump-
tion that fish evidence can contribute much to the study of hy-
drographic history. Our attempts to fulfill this promise have
been guided further by Dwight Taylor's methods and his prin-
ciple (1960) that hydrologic, topographic, and climatic features
change more rapidly than do lineages of organisms.
The Late Cenozoic fossil record in the Great Basin reflects
the climatic and faunal revolution between the Paleogene and
the Neogene. For example, the temperate Miocene to Holocene
fish faunas of the Great Basin (as now known) share only one
family (Catostomidae) with the subtropical faunas of the west-
ern North American Eocene in Wyoming, Utah, and Colorado
(Grande, 1984). Transitional Great Basin fish records are
Cope's (1872) Amyzon mentalis and Trichophanes hians from
the Osino, Nevada, Oligocene coal beds and Early Tertiary fos-
sils similar to Diplomystus in Elko County, Nevada. Our dis-
cussion herein, however, will be limited to the late Cenozoic
record. Publications on fossil Empetrichthys, fundulids, and
cyprinodontids by Miller (1945b) and Uyeno and Miller (1962)
included reviews of many pre-Pleistocene records of modern
lineages in the southwestern Great Basin. LaRivers (1962)
compiled the first account of endemic and introduced fishes in
Nevada, including what was known at that time about ecology
and fossil fishes.
Newberry (1871, and references therein) was the first to
study Great Basin fishes from a geological context, but he was
soon followed by Jordan (1878a, 1878b, 1891), Cope (1883),
Jordan and Evermann (1896), and Tanner (1936). J.O. Snyder
and Hubbs and Miller, in works cited above, were the next to
dominate the field. Uyeno and Miller (1962), Miller (1981),
M.L. Smith (1981), and G.R. Smith (1966, 1978, 1981a,
1981b) discussed fossils and described patterns of differentia-
tion, species density, and zoogeography of western North
American fishes among basins. Minckley et al. (1986) dis-
cussed this history in detail, then compiled current geological
literature on the Cenozoic tectonics of western North America
and hypothesized broad aspects of zoogeographic history of
western fishes. They looked for congruence between area cla-
dograms and fish cladograms then available and concluded that
the modern fauna occupied western tectonic units in the Oli-
gocene before they were separated from Atlantic Ocean tribu-
taries by the Miocene cycle of tectonic uplift. Minckley and
Deacon (1991) described critical conservation issues and ex-
tinction threats to the modern fauna. Behnke (1992) provided
original analysis and summarized the available information on
recent trouts. Sigler and Sigler (1987, 1996) summarized the
modern fauna, including introduced fishes.
The fishes discussed in the present paper inhabit or inhabited
the hydrographic Great Basin, as defined, for example, by
Hubbs and Miller (1948b) as being the region of Holocene in-
ternal drainage bounded by barriers shared with the present
drainages of the Colorado, Sacramento, Klamath, and Colum-
bia Rivers (Figure 1). The Great Basin is not static; rather, it is
a mosaic whose boundaries and external connections have
shifted with tectonic, volcanic, and climatic changes across
time scales ranging from 10 thousand years (Ky) to 10 million
years (My) during the last 17 My (Minckley et al., 1986; and
citations below). Many hydrographic connections have been
hypothesized. For example, in the middle Miocene, there were
large lakes in northern Nevada and Utah that drained through
the Snake River Plain southwest across the Sierra Nevada to
the Pacific coast drainage (Stokes, 1979). Later, the Lahontan
Basin (which still drained to the Pacific Ocean) was connected
to the Snake River Plain. At different times, the White River
(Hubbs and Miller, 1948b), Las Vegas Creek (Haynes, 1967),
Death Valley-Owens River systems (Miller, 1946a, 1948), and
perhaps even the Bonneville Basin (Ives, 1948) and upper
Snake River were variously connected to the Bouse Embay-
ment of the lower Colorado River. As the Sierra Nevada was
uplifted, it is possible that a north-to-south drainage system in-
cluded basins in Oregon and possibly Washington (Smith et al.,
2000). There is good evidence that 2.5 million years ago (Ma)
the Snake River's course changed with its capture by the Co-
lumbia River through Hell's Canyon (Wheeler and Cook,
1954). At undetermined times in the late Cenozoic, there were
connections from the Mono Lake basin to the Lahontan Basin,
from Owens River west to the Pacific coast drainage, and later,
south to Mexico. In addition, Death Valley was connected to
the lower Colorado River as well as to the Basin and Range
Province in Mexico (Minckley et al., 1986; Hale, 1985). Fif-
teen to 20 thousand years ago (Ka) the basins of internal drain-
age were probably restricted mainly to northern and eastern
Nevada and sometimes to the Death Valley system. At about
that time, the Bonneville Basin and the basins of Catlow and
Fort Rock Valleys and of Harney and Malheur Lakes (hereaf-
ter, Harney-Malheur) were sometimes tributaries to the Snake
River (Taylor and Bright, 1987; Bisson and Bond, 1971;
Minckley et al., 1986), whereas the White River was occasion-
ally connected to the lower Colorado River (Taylor, 1983). The
sizes and connections of basins fluctuated with climatic mois-
ture through the late Pliocene (Smith et al., 1982) and the Pleis-
tocene (Morrison, 1991; DiGuiseppi and Bartley, 1991).
The hypotheses presented in this paper do not assume the
above hydrographic outline, but they suggest that connections
and disconnections, such as outlined above, were more frequent
than commonly assumed. Our approach is to use fish fossils,
178 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
37'
121* 119' 117* 115
? Miocene
A. Pliocene
# Pleistocene
1 ' i
-35"
FIGURE 1.?Fossil fish localities in the Great Basin and adjacent drainages. Ages of fossil fish are indicated by
symbols defined on the figure. Fossils are listed in the Appendix, where they are indicated by a cross. Heavy
lines demarcate major drainage divides.
phylogenetic analysis, biogeography, and genetic distances as
primary evidence for the times and locations of hydrographic
change.
Material and Methods
Our hypothesized drainage connections are based upon dis-
tribution ranges of fish clades, species, or genetically similar
populations, especially species and subspecies that span drain-
age divides that are no longer passable. We assume that ancient
fishes dispersed only through their habitats, because this
method is supported by testable generalizations, in contrast to
ad hoc hypotheses for dispersal offish eggs, such as on birds'
feet (e.g., Brown and Rosen, 1995; Spencer and Patchett,
1997). Likely geological mechanisms for transfer include lake
spillover (Hubbs and Miller, 1948b), river capture, erosional
beheading of streams, diversion of channels by erosion, and di-
version of drainage by uplift or volcanic eruptions (Bishop,
1995). The degree of isolation of basins constrains dispersal,
which leads to predictions that are testable by geographic, geo-
NUMBER 33 179
logic, genetic, and cladistic analyses. The large number of adja-
cent basins with similar habitats permits analysis of replicated
histories of similar local faunas in addition to isolated case his-
tories. It is important that we attempt to use fish data that are
independent of geological assumptions, in the spirit of Hubbs
and Miller (1948b) and Taylor (1960), so as to contribute non-
circular conclusions useful to geology.
Records of Holocene fishes from the Great Basin were com-
piled primarily from the records of Hubbs (and his students and
successors), which have been deposited in the Division of
Fishes, Museum of Zoology, University of Michigan, and from
the literature. Late Cenozoic fossil fish data (Figure 1; Table 1)
were compiled from collections in the Museum of Paleontol-
ogy and the Museum of Zoology, University of Michigan; the
W.M. Keck Museum, Mackay School of Mines, University of
Nevada, Reno; the Museum of Vertebrate Paleontology, Uni-
versity of California at Berkeley; the Department of Biology,
California State University at Bakersfield; and the paleontolog-
ical literature. (Lists of comparative material are available from
authors G.R.S. and K.W.G.) Records of new distributions are
noted in the species accounts in the Appendix. These accounts
include information on specimen characteristics, collecting lo-
calities, and names of the collectors when the specimens were
not deposited at the above-listed institutions (Appendix; Table
5). Ancestral-descendent pairs are recognized by shared apo-
morphies that are not shared with the next closest sister group.
The probability of discovering ancestral-descendant pairs is
thought to be low in marine and terrestrial systems (Foote,
1996), but in the Great Basin it is greatly enhanced by short
geological time frame, highly constrained geography, the small
number of taxa, and the salient osteological characteristics
among fish taxa.
Cladistic methods used to estimate branching sequences for
cyprinid fishes are from Dowling et al. (2002), Simons and
Mayden (1997, 1998), and Cavender and Coburn (1992); for
Catostomidae, Smith (1992b); for Prosopium, Smith and Todd
(1993); for trout, Shiozawa and Evans (unpublished); and for
Cyprinodontidae, Echelle and Dowling (1992). Molecular
methods follow those used by Echelle and Dowling (1992),
Shiozawa and Evans (1995), and Dowling and Naylor (1997).
Cladograms were calculated from mitochondrial DNA
(mtDNA) sequence data, using the parsimony algorithm "Phy-
logenetic Analysis Using Parsimony" (PAUP*; Swofford,
1990), version 4.0 d64 by D. Swofford. Informative characters
among the species of Cyprinidae are the shared base pairs of
the mitochondrial cytochrome b gene, relative to outgroups.
North American and Eurasian cyprinids were used as out-
groups for the analysis of western North American Cyprinidae.
Informative cladistic characters among Oncorhynchus clarki
and among Cyprinodon are shared restriction-site differences
in mtDNA, relative to outgroups. Oncorhynchus mykiss (rain-
bow trout) is the outgroup for the cladistic analysis of subspe-
cies of Oncorhynchus clarki (cutthroat trout). Cyprinodon mac-
ularius and several species of southwestern Cyprinodon from
outside the Great Basin served as outgroups for the cladistic
analysis of Cyprinodon species of the Death Valley system.
DATA USED TO ESTIMATE THE RATE OF EVOLUTIONARY
CHANGE PER UNIT TIME.?The relative homogeneity of molec-
ular distances (Tables 2-4) among species with a common an-
cestor suggested a source of independent evidence to break the
logical circularity inherent in using geological events to date
evolution and in using evolutionary differentiation to date geo-
logical events. Previous workers carefully contrasted geologi-
cal and ichthyological data, but it has been difficult to find in-
dependent data to constrain the time of geographic isolation.
The traditional method (Hutchinson, 1939) defines a relevant
interval of evolutionary time as beginning with the last possible
period of hydrographic connection. This method assumes that
one panmictic and homogeneous form at the time of connec-
tion evolved into two isolated and differentiated forms in the
present. For example, it is assumed that a uniform Pliocene
population before the Panamanian land bridge evolved into Pa-
cific Ocean and Atlantic Ocean populations isolated by the
Panamanian isthmus during the last 3.5 My (Bermingham et
al., 1997). This method is limited to investigation of only one
possible time for the origin of divergence, with bias toward the
most recent.
It has been deemed parsimonious to choose the time of most
recent connection as the beginning of the divergence because
to choose an earlier time seemed to add an unnecessary ad hoc
hypothesis, i.e., more time than is minimally necessary. Dis-
covery of the amount of time required for the observed diver-
gence is, however, the original question, and to assume the
shortest of several possible times does not test the hypothesis.
It is important to test the alternative possibility, namely, that
multiple populations were present prior to the time of last con-
nection and that the forms observed in the Holocene have been
accumulating genetic differences for a time period longer than
since the last connection. This alternative is important for stud-
ies such as the Panamanian separation of Atlantic and Pacific
marine faunas, as well as for freshwater fishes in the Great Ba-
sin. Divergence times of clades must ultimately be based upon
ages determined from the stratigraphic context of fossils?usu-
ally some combination of mammalian biostratigraphy (Wood-
burne, 1987), potassium-argon, 40argon-39argon, or other ra-
diogenic dates (e.g., Baksi et al.,1992), magnetostratigraphy
(Negrini et al., 1987), fission-track ages (Kimmel, 1982), and
field studies of geomorphology and geochemistry (e.g., Alli-
son, 1940, 1982; Hubbs and Miller, 1948b; Bright, 1966, 1967;
Benson and Thompson, 1987; Benson et al., 1990; Curry,
1990; Morrison, 1991; Reheis and Morrison, 1997).
We compared data on divergence of mtDNA molecules
among fishes (Kocher and Carleton, 1997) with dates of rele-
vant fossils to calculate rates of molecular evolution (Collins et
al., 1996). Molecular differences among homologous mito-
chondrial genes in these taxa are expressed as pairwise percent
sequence divergence (Gillespie, 1991). The difference values
are proportions of the specified part of the genome that differ in
180 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 1.?Late Cenozoic fossil fish localities in the Great Basin. Numbers refer to indexed taxa mentioned in the
Appendix; a question mark (?) = uncertain identification or age.
Locality, by state
Nevada
Bullfrog Mine near Beatty, Nye County, 23 Ma?
Esmeralda, Esmeralda County
Siebert Tuff, Nye County, 17 Ma
Buffalo Canyon, Churchill County, 15 Ma
Middlegate Formation, Churchill County
Stewart Valley, Mineral County, 12 Ma
Virgin Valley, Humboldt County
Truckee Formation, Churchill County
Red Rock locality, Churchill County
Rabbit Hole, Pershing County
Sahwave Mountains, Pershing County
Humboldt Formation, Elko County, 9.5 Ma
Coal Valley, Mineral County
Alum Mine, Nye County
Jersey Valley, Pershing County
Brady Pocket, Churchill County, Hemphillian?
Carson Valley, Douglas County
Oregon
Trout Creek, Harney County
Utah
Cache Valley Formation, Salt Lake group
California
Secret Valley, Lassen County
Madeline Plains, Lassen County
White Hills, Airport Lake, Inyo County
Mono Basin
Honey Lake
Idaho
Salt Lake group at Georgetown
Nevada
Carson Valley, Douglas County, early Blancan
Carson Valley, Douglas County, late Blancan
Pine Nut Ridge, Douglas County
Duck Valley, Washoe County
Carson Sink, Churchill County
Mopung Hills, Churchill County
Fish Cave, Churchill County
Oregon
Fort Rock Lake, Lake County, Blancan
Utah
Junction Hills, Salt Lake group, Hemphillian
Park Valley, Box Elder County, Hemphillian?
California
Manix Lake, San Bernardino County
Searles Lake, San Bernardino County
China Lake, San Bernardino County
Lake Mohave or Soda Lake, San Bernardino County
Owens Lake Core, Inyo County
Nevada
Humboldt River Canyon, Pershing County
near Fallon, Churchill County
Carson Valley, Irvingtonian
Sunshine Amphitheater, northwestern Walker Lake,
pre-Lahontan
Lahontan Cave deposits
Smith Creek Cave
Indexed taxa
MIOCENE
117
5,27,28,58,88, 120
122
6,88
88
27,88
6,58, 124
8,29,80,89, 128
87
88
116, 128
123,130
6,88
88
8
56
15,30,88
129
3
PLIOCENE
1,30,69,91, 124
1,25,267,30,91
12,62
1,13, 14,30,63,74
9,25,60,65,68,70,91
3,59
2, 15,30,64,75
2, 15,64,92
6
10,62
92
9,72,93, 124
124
7, 59, 62, 90
3,23,26,30,57,61,69,81
125
PLEISTOCENE
12,129
12
12,62
12
13,73,82,92
10,92
60,72
15,30,64
72,92
124
23,99
Source1
Lugaski, UNRM
LaRivers, 1962, UNRM
Lugaski, 1978, UNRM
T. Lugaski, UNRM
D.I. Axelrod, UCD
Lugaski, UNRM
Bell, UMMP
J.R. Alcorn, UMMP
LaRivers, UNRM
Lugaski, UNRM
T. Kelly, MVP
LaRivers, UNRM
Lugaski, UNRM
Charles Dailey, MVP, UMMP
C. Arnold, UMMP
J.S. Williams, USU
UMMZ, Wagner et al., 1997, MVP
Wagner etal., 1993, MVP
Gobalet; UMMP
Miller and Smith, 1981; Gobalet; UMMP
Taylor and Smith, 1981, UMMP
S.A. Bilby, ISU
T. Kelly, Gobalet, MVP
T. Kelly, Gobalet, MVP
T.Kelly, 1994, MVP
Taylor and Smith, 1981, UMMP
R. Miller, UMMP
Allison and Bond, 1983, OSU
P. McClellan, MVP
T. Cavender, UMMP
Buwalda, 1914; Reynolds, 1991, 1994
UMMP
Leatham, UMMP
G.I. Smith, UMMP
Gobalet, MVP
J.R. Alcorn, UMMP
T. Kelly, Gobalet, MVP
M. Reheis, USGS
Jordan, 1924, MVP
Mead etal., 1982
NUMBER 33 181
TABLE 1.?Continued.
Locality, by state
Oregon
Chewaucan, Lake County
Utah
Hot Springs, Salt Lake County
Black Rock, Salt Lake County
Black Rock, Tooele County
Homestead Cave
Thatcher Basin
Old River Bed
Nevada
Nevada AEC Test Site, Nye County, Miocene?
California
Indexed taxa
11
23, 66, 84, 85, 125, 126
23,84,85,98, 125, 126
23
23, 66, 84, 85, 87?, 98, 126
61
62, 82, 98
STRATIGRAPHY UNCERTAIN
119
Titus Canyon, Pliocene? (stratigraphy uncertain) 111, 117
Titus Canyon Formation, Oligocene?
Furnace Creek, Death Valley, Pliocene?
117
118
Source'
Gobalet and Negrini, 1992
Smith etal., 1968, UMMP
Smith etal., 1968, UMMP
J.H. Madsen, Jr., UMMP
Broughton, UU
Bright, UMMP
Oviatt, UU
Richard Wyman, UMMP
Miller, 1945b
Lugaski, UNRM
Miller, 1945b
'Published source and/or collector and location of specimens. Abbreviations: ISU = Idaho State University
Museum, Pocatello; MVP = Museum of Vertebrate Paleontology, University of California, Berkeley; OSU =
Oregon State University, Corvallis; UCD = University of California, Davis; UMMP = University of Michigan
Museum of Paleontology, Ann Arbor; UMMZ = University of Michigan Museum of Zoology; UNRM =
Mackay School of Mines, University of Nevada, Reno; USGS = United States Geological Survey; USU =
Department of Geology, Utah State University; UU = University of Utah.
accumulated mutations, corrected for back mutation (Tamura
and Nei, 1993). We were encouraged to pursue this method by
the homogeneity of values of pairwise comparisons of sets of
lineages, each derived from a common ancestor (Tables 2-4).
This homogeneity suggested that the rate of molecular evolu-
tion was roughly constant for a given gene system in a given
clade. The relative-rate test of Sarich and Wilson (1967) was
used to evaluate the precision of the time estimate. It is based
upon the assumption that all taxa derived from one node
(branch point) on a cladogram have evolved while accumulat-
ing mostly neutral mutations for the same amount of time (e.g.,
Figure 2; Table 2), regardless of the number of intervening
branch points; therefore, all the distances (in a homologous
molecule) from the sister clade to all the members of the clade
in question should be identical. This degree of similarity is a
measure of the constancy of the rate of molecular change and,
therefore, an estimate of the precision of a molecular estima-
tion of geological time. In other words, the homogeneity of the
measured distances between sister branches within a clade is an
indication of how constant the rate of molecular change has
been. It is important to stress that these distance values were
tested in conjunction with ages of fossil fish that we interpret as
ancestors; distance values were not converted to rates of mo-
lecular change based upon molecular clock assumptions from
mammals or other organisms.
If we accept the accumulated differences (genetic distance)
as numerators in a rate estimate for these fishes, we then need
an independent denominator measured in millions of years to
yield the amount of molecular change per one million years
(Collins et al., 1996). The denominator, evolutionary time, is
estimated herein as millions of years since the time from first
appearance of a fish lineage to the present. Relative times of
lineage origins, recognized by possession (in fossils) of evolu-
tionarily derived states that diagnose appropriate groups
(Lundberg, 1993), are underestimated because occurrences of
fossils are usually minimum estimates of the age of their lin-
eage (Minckley et al., 1986). This problem exists because the
first members of a lineage are not only rare, they are also not
easily recognized (assuming that apomorphies accumulate
gradually at many points along branches between nodes).
Where fossils are limited to infrequently preserved deposi-
tional environments, the sample sizes are not adequate to eval-
uate statistical error around the estimated age of the first occur-
rence of a lineage. The denominator error from underestimated
times of first appearance results in overestimated rates of
change (rate = percent divergence/time). Errors in the numera-
tors, e.g., from back mutations, yield underestimated rates, un-
less genetic distances are adequately corrected (Tamura and
Nei, 1993). The potential consequences of these errors are
readily apparent from inspection of Figure 14, in which numer-
ator error is corrected, but denominator error is not; therefore,
our rates of molecular change (0.5%-1%/My) are likely to be
overestimates, approximating upper bounds. If the oldest of
these lineages are as old as the Oligocene, as suggested by
Minckley et al. (1986), correct rate estimates could be 0.3%-
0.7%/1 My.
The pairwise difference values for lineages in a clade, such
as the Cyprinodontidae, Salmonidae, or Cyprinidae, are re-
gressed against geological divergence-time estimates (Figure
14). The regression coefficient then gives the percent sequence
182 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
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Ctenopharyngodon idella
U
NUMBER 33 183
i i i i i i r
35 30 25 20 15 10 5
% sequence divergence
Phoxinus neogaeus, Eastern N.A.
Mylopharodon, Sacramento Dr.
Lavinia, Sacramento Dr.
Hesperoleucus, Sacramento Dr.
Siphateles obesus, Lahontan Dr.
Siphateles o. pectinifer, Pyramid L.
Gila (Moapa) coriacea, White Dr.
Gila robusta, Colorado Dr.
Gila elegans, Colorado Dr.
Gila seminuda, Virgin Dr.
Gila purpurea, RioYaqui
Gila atraria, Bonneville-Snake Dr.
Gila cypha, Colorado Dr.
Gilajordani, White Dr.
Gila orcutti, Coastal S California
Gila nigrescens, Guzman Basin
Gila pandora, Rio Grande
Acrocheilus, Columbia Dr. -
Klamathella coerulea, Klamath Dr.
Ptychocheilus lucius, Colorado Dr.
Ptychocheilus grandis, Sacramento Dr.
Eremichthys, Lahontan Dr.
Relictus, L Franklin-L Waring, NV
Ptych. oregonensis, Columbia Dr. -
Orthodon, Sacramento Dr.
Snyderichth. S Bonneville Dr.
Snyderichth. Snake-Bear R.
Lepidomeda molispinus, Virgin Dr.
Lepidomeda vittatus, Little Colorado Dr.
Rhinichthys osculus, western U.S.
Mylocheilus, Columbia Dr. -
Pogonichthys, Sacramento Dr.
lotichthys, Bonneville
Richardsonius balteatus, Bonneville-
Columbia Dr.
FIGURE 2.?Mitochondrial DNA phylogeny of western Cyp-
rinidae, based upon parsimony cladistic analysis of the cyto-
chrome b gene. This tree is extracted from a larger analysis of
88 cyprinid taxa (Dowling, unpublished). The relative rates test
(see "Methods") is based on the expectation that all of the sister
species to any outgroup should have the same percent sequence
divergence. The variation in those distances (Tables 2-4) is a
measure of the variation in relative rates of change and the lack
of precision in any inference. The depth to each node is scaled
to age in millions of years, assuming that 1 My corresponds to
1% average distance between clades (Table 2). Rank order of
cladistic nodes and distances is preserved within 1%, except
that the Relictus-Eremichthys distance is 15%.
divergence per one million years for that gene and that clade
(Figure 14), with an indication of the reliability of that estimate
(Table 2). Once we have estimated the rate of change for a gene
in a taxon, we can return to the cladogram and can estimate the
amount of absolute time elapsed between sister clades in differ-
ent basins, as estimated by genetic differences. Cladograms
(without time) can then be recast as phylogenies (with estimates
of time), as in Figures 2, 10, and 11. The estimated rate depends
upon the correction factor for back mutations at all base pair
sites. For example, the Tamura-Nei correction, used herein,
gives 50 percent higher sequence-divergence values and older
branch times (for the older cyprinids) than p distances or the
Kimura two-parameter correction (Tamura and Nei, 1993).
In the analyses below, both morphological and molecular
cladistic estimates of rates of change are presented for the Cyp-
rinidae (Figures 2, 3), but the molecular cladistic estimate is
used to frame the relative-rate test. Correlations between the
node sequences of the molecular trees and percent sequence di-
vergence are positive and significant; correlations between
node sequences of the morphological trees and percent se-
quence divergence are not significant.
Different kinds of hydrographic connections between basins
affect molecular-rate estimates differently. For example, lacus-
trine spillovers of lakes into rivers that transferred low-eleva-
tion suckers, minnows, pupfishes, lacustrine trouts, and white-
fishes occurred in sometimes datable pluvial times during
glacial cycles (e.g., Lake Bonneville's spillover, McCoy,
1987), but headwater stream captures that transferred cold-wa-
ter, high-gradient mountain suckers, salmonids, and sculpins
occurred more frequently (e.g., between the Bear, Snake, and
Green Rivers). Therefore, large, lowland fishes will generally
contribute to older time estimates, whereas headwater fishes
contribute to younger estimates. This is independent of the
question of whether headwater fishes evolve more rapidly
(Smith, 1992b).
There are two alternative general hypotheses for times of di-
vergence of Great Basin fishes. The first alternative is that a
population was devoid of genetic structure prior to the most re-
cent connection. Alternatively, differentiation may date not
from the most recent connection and disconnection, but from a
previous period of high water in the Pleistocene or Pliocene
(see discussions of Whitehorse and Bonneville cutthroat trout
and Death Valley pupfishes, below). Consideration of a broader
range of possible isolation phenomena, along with dated fossils
and molecular data, enable us to test alternative hypotheses
about timing of divergence.
Several additional problems may detract from the generality
of these methods. In particular, fish fossils are limited, and
some possess few apomorphic characters. More complete mo-
lecular cladograms for more taxa are also needed. In addition,
mtDNA data may have suffered introgressive loss of evolution-
ary information in basins that frequently changed connections
184 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Richardsonius
Phoxinus, etc.
Hesperoleucus
Lavinia
Orthodon
Mylocheilus
Acrocheilus
Pogonichthys
Mylopharodon
Ptychocheilus
Gila atraria
lotichthys
Siphateles
Relictus
Eremichthys
Gila
Plagopterins
Gila (Moapa)
Rhinichthys
FIGURE 3.?Morphological phylogeny of western Cyprinidae after Coburn
and Cavender (1992). This tree was extracted from a consensus tree, which was
based upon 100 equally parsimonious trees representing a larger sample of
cyprinid fishes.
among allopatric populations (Smith, 1992a; Echelle and
Dowling, 1992; Dowling and Secor, 1997).
Cladistic patterns and estimated amounts of molecular diver-
gence are presented in the "Results." Distinct population units
are listed and indexed in the Appendix, with their diagnostic
characters, location, fossil ages, and brief history. Statements
about the kinds of characters and the qualitative degree of dif-
ferentiation seen between sister units are included to enable
comparisons with our molecular- and fossil-based hypotheses
about rates of evolution.
Authors of recent species and subspecies names are given in
the Appendix, following Robins et al. (1991). Citations and au-
thors of species names of fossils are given in the Appendix. In-
stitutional abbreviatons are defined in Table 1.
Acknowledgments
Robert Miller and his family and students collected thou-
sands of fossil fish bones. Dwight W. Taylor was important in
field and theoretical development of this project, but our in-
debtedness to him does not imply his agreement with all of our
methods or conclusions. The late Carl L. Hubbs and Ira LaRiv-
ers collected fossils and compiled information. Ray Alcorn,
Clyde Barbour, Robert Behnke, Sue Ann Bilby, Beverly Black,
Anthony Boudreau, Bob Bright, Jack Broughton, Lee Carlson,
David Catania, Charles Dailey, Jonathan Davis, R.L. Elder,
John Fortuna, Kallie Hardin, Geoff Hoetker, Pat Holroyd, Tho-
mas J. Kelly, Britt Leatham, Susan Lindstrom, David Madsen,
James Madsen, Rose Malamma, Carl Mann, Todd Martin,
Kelly McGuire, Carrie Muerer, Robert Negrini, Doug Nelson,
Kevin Padian, Phil Pister, Marith Reheis, Bill Robin, Mark
Roeder, Gary Scoppettone, Keith Smith, Ray Smith, Ralph
Stearley, Scott Stine, Steve Stone, Joanna Strain and family,
Dwight Taylor, Paul Wagner, David Whistler, David White,
Russ Wickwire, Richard N. Williams, Todd Yeoman, and many
other family members, students, and friends collected speci-
mens or provided comparative material and information. Shane
Webb provided goodeid mtDNA data, examined fundulid fos-
sils for goodeid characters, and helped with computations.
Donald Buth provided fish DNA for molecular studies. We are
especially grateful to Robert Miller for granting us permission
to mention fossils that he collected and for directing us to use-
ful literature. The California Academy of Science and the Mu-
seum of Paleontology, University of California at Berkeley,
loaned fossils for study. Research was supported by grants
from the National Science Foundation (NSF) and the U.S. Bu-
reau of Reclamation to Dowling. The U.S. Fish and Wildlife
Service, Utah Division of Wildlife Resources, Nevada Depart-
ment of Wildlife, U.S. Forest Service, and the Wyoming De-
partment of Game and Fish supported research by Shiozawa.
California State University, Bakersfield University Research
Council, and the Far Western Anthropological Research Group,
Inc., Davis, California, supported research by Gobalet. Bonnie
Miljour created the figures. Robert Miller, Catherine Badgley,
Robert Hershler, Ed Wiley, and anonymous reviewers read and
improved the manuscript.
Results
Fifty-two recent species?19 Cyprinidae, 12 Catostomidae,
eight Salmonidae, five Cyprinodontidae, four Goodeidae, and
four Cottidae?comprise the Holocene Great Basin fish fauna.
Twenty-seven of these species are endemic. Four species
among the cyprinids, goodeids, and cottids have been driven to
extinction, and two species or subspecies of salmonids have
been extirpated from the Great Basin as a result of habitat de-
struction caused by human activities (Minckley and Deacon,
1991). Among the 52 species (especially cyprinids, salmonids,
and cyprinodonts), more than 50 additional isolated popula-
tions have been recognized as separate subspecies. Sixty-four
fossil forms, from the late Miocene to the present, represent
many recent lineages and several clades that are now locally
extinct (for example, Gasterosteidae and Centrarchidae). Fossil
NUMBER 33 185
fishes are now known from 55 Great Basin localities (Figure 1;
Table 1); several hundred additional fossil localities occur at
the northern edge of the Great Basin, on the Snake River Plain.
Extensive analyses of mtDNA divergence have been carried
out on recent representatives of these fossil forms: Dowling et
al. (2002) on Cyprinidae (minnows), Echelle and Dowling
(1992) on Cyprinodontidae (pupfishes), and Shiozawa and
Evans (1995) on Salmonidae (trouts).
CYPRINIDAE.?Minnows, family Cyprinidae, are the most
diverse and widespread fishes in the Great Basin; the 19 Ho-
locene lineages include 10 endemic species and more than 20
endemic subspecies. Three species of minnows, Gila atraria,
Siphateles bicolor, and Rhinichthys osculus, have been espe-
cially persistent and have differentiated into many distinct sub-
species and local populations. Unfortunately, small minnows
are under-represented in the Great Basin fossil record, so they
contribute little to our estimates of rates of molecular evolu-
tion. At least five lineages, Ptychocheilus, Mylocheilus in-
flexus, M. robustus, Mylopharodon, and Lavinia (exclusive of
Hesperoleucus), lived in the Great Basin in the Pliocene and
Miocene but are absent there today. Siphateles is the best repre-
sented genus in the fossil record, occurring in the Miocene to
Pleistocene of the Lahontan Basin and Mohave River drain-
ages.
Fossil and living Cyprinidae in the Great Basin are herein as-
signed to three ancient clades, which are based on our molecu-
lar analyses (Figure 2; see also Simons and Mayden, 1997,
1998). These three clades are older than the Great Basin, and
molecular evidence indicates that they probably date back to
the Oligocene (Table 2; Appendix). The first of the three clades
(Figure 2) in the Great Basin (including fossils) is the large
western North American chub group that includes Mylopharo-
don, Lavinia, Hesperoleucus, Siphateles, Gila, Acrocheilus,
Klamathella, Ptychocheilus, Eremichthys, and Relictus (Ap-
pendix, numbers 1-32). These are related to Orthodon in the
Sacramento drainage system and to Phoxinus in eastern North
America. This "Western Chub clade" has four subgroups in the
Great Basin: (1) the Mylopharodon-Lavinia-Hesperoleucus
and Siphateles clade (Appendix, numbers 1-22); (2) the
Gila-Acrocheilus-Klamathella clade, with Ptychocheilus lu-
cius as its sister group (Appendix, numbers 23-29); (3) the
Relictus-Eremichthys clade, including Ptychocheilus grandis
(Appendix, numbers 31-32); and (4) the Ptychocheilus oregon-
ensis clade (Figure 2; Appendix, number 30). Whether or not
Ptychocheilus is actually paraphyletic, as indicated by Figure
2, is problematical. Ptychocheilus cf. oregonensis is one of the
oldest cyprinids in the region (Smith et al., 2000).
The second major cyprinid group in the Great Basin is the
Snyderichthys-Lepidomeda (spinedace) clade (Appendix,
numbers 33-35). It is related to the other spinedaces, which in-
habit tributaries of the Bouse Embayment (see below). Sister
groups of this clade include Margariscus and Couesius of
northern and eastern North America.
The third major minnow clade includes western American
outliers of the huge eastern North American shiner clade. West-
ern forms include Iotichthys (Appendix, number 36), Richard-
sonius (Appendix, numbers 37-39), Rhinichthys (Appendix,
numbers 40-55), and Mylocheilus (Figure 2; Appendix, num-
bers 56-58). This group also includes Pogonichthys of the Sac-
ramento drainage system. An alternate cladistic hypothesis
(Figure 3), based upon morphological analysis, shows different
interpretations of relationship (Coburn and Cavender, 1992).
Great Basin cyprinids are listed (Tables 1, 2), numbered, and
reviewed in the Appendix in approximately the sequence given
in Figure 2. Typical distribution patterns of minnows (illus-
trated by Gila and Siphateles) are shown in Figure 4.
Fossil minnows in the Great Basin date back to the middle
Miocene or earlier (Cavender, 1986; Reynolds, 1994). Many
cyprinid genera are early members of the Great Basin fauna
(Appendix, numbers 5, 14, 15, 26-28, 30, 56-58). The Gila
group is widespread and especially diverse in all of the Great
Basin as well as in the Colorado River drainage, the Rio
Grande, and northern Mexico. Siphateles is centered in the
Lahontan Basin and Sacramento River drainages, and Gila is in
the drainages surrounding the Lahontan Basin. The Great Basin
part (White River) of the lower Colorado River (Bouse Embay-
ment) drainage is inhabited by Lepidomeda, Moapa, and Gila,
as well as by Catostomus clarki and the empetrichthyine good-
eids. Snyderichthys copei of the Bonneville Basin and upper
Snake River drainages is cladistically part of the Plagopterin
group and provides evidence of Pliocene and Pleistocene drain-
age connections (Dowling et al., 2002). The southern Califor-
nia coastal form, Gila orcutti, is the sister group to Gila atraria
of the Great Basin and to G. jordani and G. cypha of the Colo-
rado River drainage.
Electrophoretic variation in Gila atraria of the Bonneville
Basin was studied by Rosenfeld (1991), who concluded that ac-
cumulated differences were influenced more by natural selec-
tion in local springs, lakes, and streams than by vicariance and
time. He estimated that 14,500 years was too short an interval
for the accrual of qualitative differences between hierarchical
groups of vicariant populations. Populations of Gila atraria
from northern and southern Bonneville basins and from the
Bear River have diverged extensively in their mtDNA; they
have not remained uniform during the last 2 My (sequence
divergence = 0.009-0.024) despite recent and probably fre-
quent sympatry. The northern Bonneville form of G. atraria
shares an unexpectedly high proportion of its mtDNA with G.
cypha and G. jordani of the Colorado River (percent sequence
divergence = 1.1% and 0.7%, respectively; Table 2), support-
ing the hypothesis that a connection existed from Lake Bonne-
ville through Meadow Valley Wash, Nevada, to the Colorado
River drainage (Ives, 1948).
A separate biogeographic pattern encompasses the northern
Great Basin and the Columbia River and Snake River drain-
ages. This pattern is marked by the distributions of Kla-
mathella, Siphateles, Acrocheilus, Lavinia, Mylopharodon, and
186 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
121? ?119- 117* 115
Miocene
Pliocene
Pleistocene
-35s
FIGURE 4.?Distribution of Holocene and fossil Siphateles, Gila, and certain relatives in the Great Basin. Map
shading indicates different species or subspecies ranges. Ages of fossil fish are indicated by symbols defined on
the figure. Numbers refer to species or subspecies listed in the Appendix.
Mylocheilus in a Columbia-northern Great Basin-Pit River
track. At a lower cladistic hierarchical level, this biogeographic
pattern broadens to include the Sacramento River drainage,
with Ptychocheilus grandis, Hesperoleucus, and Orthodon
(Figures 2, 3), and the Colorado River drainage, with Pty-
chocheilus lucius. The several Great Basin records of My-
locheilus and Mylopharodon are evidence of a late Miocene
connection between west-central and northwestern Nevada and
the Snake River Plain. Mylocheilus caurinus (Table 2, number
58) is the closest living relative to Mylocheilus robustus of the
Mio-Pliocene Snake River Plain (Appendix, number 58) and to
Mylocheilus heterodon from the Ringold Formation of Wash-
ington (Smith et al., 2000).
Pairwise distances among clades that are based upon a parsi-
mony cladistic analysis enable evaluation of relative rates of
molecular change (Table 2). The homogeneity of these rates
among minnows can be evaluated by comparing the standard
deviation (s.d.) around the mean distances. The 26 species esti-
mated to share a common ancestor with Phoxinus have a mean
distance between each of them and Phoxinus of 0.22 (s.d. =
0.022). The 66 combinations of sister species to Mylopharo-
don, Lavinia, and Hesperoleucus have a mean distance of 0.15
NUMBER 33 187
(s.d. = 0.02); the 38 combinations of Siphateles sister species
have a mean distance of 0.12 (s.d. = 0.02); the 30 species of the
Gila orcutti-G. atraria group have a mean distance of 0.08
(s.d. = 0.02); the 26 sister species to Acrocheilus and Kla-
mathella have a mean distance of 0.08 (s.d. = 0.02); the 216
combinations of the spinedaces have a mean distance of 0.34
(s.d. = 0.03); and the 210 combinations of the western repre-
sentatives of the shiner clade have a mean distance of 0.35
(s.d. = 0.05). With the rate of divergence estimated at about 1%
per one million years, the standard deviations of group or bar-
rier ages is usually on the order of 10%, but rarely more than
25%, of the estimated age (Table 2). The distances and inferred
timing of cyprinid patterns span the time from Miocene to the
last postpluvial period.
CATOSTOMIDAE.?Catostomids are the most ancient of the
fish groups present in the Great Basin (Cope, 1872; Smith,
1966, 1992b; Minckley et al., 1986). They are also among the
largest and hardiest aquatic survivors of the Great Basin's his-
torically variable climate and ecology. Catostomidae are the
second most diverse and abundant family in the Great Basin,
with 12 species, but the maximum number of sympatric forms
is three or four. Great Basin suckers are part of a primarily
western American monophyletic group (Figure 5), the Catosto-
mini, whose sister group (and outgroup in the analysis) is the
78 Catostomus clarki, Virgin-White
77 C. discobolus, Colorado-Snake-Bear
C. santaanae, Los Angeles
79 C. platyrhynchus, Great Basin-Columbia
C. plebeius, Rio Grande
C. leopoldi, Yaqui
C. wigginsi, Sonora
C. microps, Pit
C. rimiculus, Rogue, Klamath
76 C. columbianus, Columbia
C. catostomus, Missouri-Columbia
72 C. tahoensis, Lahontan
73 C. fumeiventris, Owens
C. snyderi, Klamath
66 C. ardens, Bonneville-U Snake
67 C. macrocheilus, Columbia-L Snake
C. latipinnis, Colorado
C. occidentalis, Great Valley
C. bernardini, Yaqui
C. insignis, Lower Colorado
C. commersoni, Mississippi
Xyrauchen texanus, Colorado
59 Deltistes luxatus, Klamath
Chasmistes brevirostris, Klamath
60 Chasmistes cujus, Lahontan
61 Chasmistes liorus, Bonneville
FIGURE 5.?Estimate of the phylogenetic tree for Chasmistes, Deltistes, and Catostomus in the Great Basin and
their relatives in the Catostomini (from Smith, 1992b). Not all species of Chasmistes were included in the cladis-
tic analysis. Numbers refer to the species accounts in the Appendix.
188 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Moxostomatini of the eastern United States and Mexico (Smith,
1992b).
The Catostomini range from medium-sized, benthic, head-
water species in the subgenus Pantosteus (genus Catostomus)
(Figure 7), to large-sized benthic predators in low-gradient
streams (other Catostomus, Figure 6), to large-sized plankti-
vores in large lakes (Chasmistes, Figure 7). The fossil record of
catostomids is diverse; it includes plesiomorphic forms as well
as forms that were morphologically more specialized than
present-day species. Fossil Catostomini are known from the
Miocene (10 Ma) to the Holocene, with Great Basin diversity
reaching a peak of four species in the northern Bonneville Ba-
sin (and upper Snake River and upper Colorado River basins)
in the late Pleistocene (see Appendix, numbers 59-79). Ca-
tostomids are tetraploid fishes that show abundant hybridiza-
tion and extreme variation, but no indication of rapid speciation
(Smith, 1992b).
There are two clades of Catostomini in the Great Basin
(Smith, 1992b). The first is the genus Chasmistes, the lake suck-
ers. Four Holocene species of Chasmistes are localized in Utah
^37'
miles
113'
? Miocene
A Pliocene
# Pleistocene
' ' iin*
-35'
FIGURE 6.?Native distribution of Catostomus ardens (Utah, 66), C. macrocheilus (Idaho, 67), C. fumeiventris
(southern California, 73) and C. tahoensis (Nevada, 72) in the Great Basin and some of their relatives outside the
Great Basin. Shading indicates drainage basins occupied by species (as numbered in the Appendix). Fossil
records and Holocene occurrences are indexed to the numbered species accounts in the Appendix. Ages of fossil
fish are indicated by symbols defined on the figure. These are low-gradient fishes that required a major river or
lake connection to cross a drainage divide.
NUMBER 33 189
76 columbianus
77 discobolus
78 clarki gg?|
79 platyrhyncus | |
-351
121"
FIGURE 7.?Distribution of the mountain suckers: Catostomus platyrhynchus (Utah, Idaho, Wyoming, Nevada,
79), C. columbianus (Idaho, Oregon, 76), C. clarki (southern Nevada, 78), and C. discobolus (Utah, Wyoming,
Idaho, Colorado, 77) in the Great Basin and of their relatives (C. rimiculus, Oregon; C. santaanae, California)
outside the Great Basin. Ages of fossil Chasmistes and Deltistes in the Great Basin are indicated by symbols
defined on the figure. Numbers refer to the species accounts in the Appendix.
and Pyramid Lakes, in Klamath River (Figure 7), and in the
headwaters of the Snake River. There are also numerous Mio-
cene to Holocene fossil Chasmistes specimens in the Great Ba-
sin (Miller and Smith, 1981; Figure 7). Silicified gill-raker
plates with fimbriate distal ends have been found in the Mono
Lake basin. Fimbriate distal ends are an apomorphy for Chas-
mistes. Their presence in Mono Lake deposits indicates this ad-
aptation to planktivory occurred 4-3 Ma (Gobalet, unpublished
data). The sister genera to Chasmistes are Xyrauchen and Del-
tistes (Appendix, number 59) of the Colorado River and the Kla-
math system, respectively. The modern lake sucker in Utah
Lake, Chasmistes liorus mictus Miller and Smith (1981), shows
intense introgression from Catostomus ardens. The introgres-
sive event is remarkable because Chasmistes and Catostomus
diverged 10-4 Ma (both genera are morphologically distinct in
Pliocene sediments of the Snake River Plain; Smith et al., 1982).
The second clade is the genus Catostomus, which includes
two ecological groups that were formerly classified as subgenera
190 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
(Smith, 1966; Smith and Koehn, 1971). As now understood,
the subgenus Pantosteus (mountain suckers) (Appendix, num-
bers 76-79) is monophyletic, but recognition of Pantosteus
makes the subgenus Catostomus paraphyletic (Smith, 1992b).
The more plesiomorphic species of catostomids are usually re-
stricted to lowlands in basins (Figure 6), where they inhabit
low-gradient rivers or lakes in a vicariant pattern. These spe-
cies are C. ardens in the Bonneville Basin, C. macrocheilus in
the Harney-Malheur Basin (and Columbia River basin), C.
warnerensis in the Warner Lakes basin, C. tahoensis in the
Lahontan Basin, C. fumeiventris in the Owens River basin,
and C. shoshonensis in Wall Canyon of Surprise Valley (Ap-
pendix, numbers 66-75; Figures 5, 6). Fossil Catostomus (10
Ma to Holocene) are known from the north edge of the Great
Basin, which is along the southern margin of the Snake River
Plain.
The mountain suckers (C. columbianus, C. platyrhynchus, C.
discobolus, C. clarki) are morphologically derived (Smith,
1992b). They inhabit high-gradient streams and stream head-
waters, and they have been broadly dispersed by many headwa-
ter transfers (Figure 7; Smith, 1966). Catostomus platyrhyn-
chus has been transferred from the Bonneville Basin to the
Lahontan Basin and to the drainages of the Snake/Columbia,
Colorado, and Missouri Rivers. Catostomus discobolus is na-
tive to the upper Colorado River, including the Grand Canyon,
and has been transferred among the Colorado River and Snake
River drainages to the Bear and Weber Rivers in the northern
Bonneville Basin. Catostomus clarki is found in the drainages
of the Gila, Virgin, and White Rivers, but not the Colorado
River in and above the Grand Canyon, and it helps define the
Pliocene drainage of the Bouse Embayment of the Gulf of Cal-
ifornia during the time when it might have been blocked from
its connection with the upper Colorado River drainage (Taylor,
1983). Catostomus santaanae, in the Los Angeles area (Figure
7), is the sister species to the Colorado River species, C. disco-
bolus and C. clarki. Catostomus santaanae represents the pop-
ulation that lived in the lowest Colorado Paleoriver when it
flowed to its Los Angeles delta, prior to 9 Ma. Catostomus co-
lumbianus inhabits both the Harney-Malheur Basin and the
Columbia River drainage and helps to define the overflow
through the South Fork of the Malheur River to the Snake
River across the Voltage lava flow (Piper et al., 1939), which
occurred as recently as about 9000 years B.P. (Gehr and New-
man, 1978).
ICTALURIDAE.?North American catfishes of the genus
Ameiurus were rare in the Great Basin. Two Miocene forms are
known (Appendix, numbers 80, 81), one from the Truckee For-
mation in Nevada and one from formations in the Salt Lake
group in northern Utah. Their relatives include fossils in the
Pliocene Columbia River, Miocene and Pliocene Snake River,
and Miocene to Holocene rivers in eastern North America
(Lundberg, 1992; Smith et al., 2000). These occurrences pro-
vide additional evidence for former Great Basin connections to
the Snake River Plain and to south-central Washington before
the formation of the modern Columbia drainage. The Miocene
form from northern Utah is most similar to bullheads of the
Snake River Plain, supporting a Hemphillian or older connec-
tion between the Salt Lake group and Lake Idaho (Taylor,
1966; Stokes, 1979). No Ameiurus are known from the south-
western United States or Mexico, despite the diversity of other
genera of catfishes in Mexico (Lundberg, 1992).
SALMONIDAE.?Eight native species of salmonids apparently
inhabited the Great Basin in recent times: rainbow trout (Onco-
rhynchus mykiss), cutthroat trout {Oncorhynchus clarki, with
many subspecies), bull trout (Salvelinus confluentus), four
whitefishes (Prosopium spp.), and possibly the Arctic grayling
(Thymallus arcticus). Additional trouts and char are represented
in the fossil record. Salmonids are cool-water fishes that inhabit
mountain streams and deep lakes. Their distribution patterns in-
dicate that they dispersed among basins via headwater-stream
captures as well as by discharge from pluvial lakes. During in-
terglacial times (such as the present), desiccation and warming
of aquatic habitats restricts these fishes to small populations in
a few cool lakes, rivers, and headwaters. During pluvial stages
of the last 3 My, however, there were numerous, large popula-
tions (see Appendix) in large, deep lakes and large rivers.
Salmonids from the Miocene, Pliocene, and Pleistocene are
known to have occurred in lake basins, such as the Salt Lake
group, Fort Rock Lake basin, Lahontan Basin, Honey Lake ba-
sin, Bonneville Basin, and Owens Basin. Holocene salmonids
are found in drainages west, north, and east of the Great Basin,
and in a few drainages at high elevations in Mexico.
COREGONINAE: There are six recent species of whitefishes
in the genus Prosopium of the subfamily Coregoninae. Four
live in the Great Basin; three of these are endemic to Bear
Lake in Utah and Idaho (Appendix, numbers 82-85). The other
two, P. cylindraceum and P. coulteri, inhabit eastern to north-
western North America and Siberia (Smith and Todd, 1993).
At least one fossil species, P. prolixus Smith (1975), is com-
mon in Pliocene sediments of the Snake River Plain. The mor-
phologically based, cladistic estimate of relationships of the
species in the group is shown on a distribution map (Figure 8).
Unpublished molecular data (Shiozawa) indicates a more com-
plicated pattern than that depicted in Figure 8. Prosopium illus-
trates intralacustrine as well as allopatric speciation in Bear
Lake and in subbasins of the Bonneville-Bear system. As pre-
dicted by the intralacustrine speciation hypothesis, the differ-
entiated characters among the species of Prosopium are mostly
associated with feeding: the shape of the mouth and jaw and
the numbers of teeth, gill-rakers, and pyloric caeca (Smith and
Todd, 1993). The phylogenetic pattern indicates that the most-
derived forms in the genus evolved in the Bonneville Basin
and Bear Lake.
THYMALLINAE: Grayling {Thymallus) have not been re-
garded as part of the Great Basin fauna. There are no fossil or
Holocene specimens documenting their presence. Inclusion in
the present list is based upon the apparently knowledgeable
identifications in the journal of John Townsend (1978), who
recorded collecting grayling in the Bear River in 1834 (Appen-
dix, number 86).
NUMBER 33 191
123? 121*
-43
prolixum, Pliocene, Idaho
coulteri, North, L. Superior
cylindraceum, North
williamsoni, West
spilonotum, Bear Lake
abyssicola, Bear Lake
gemmifer, Bear Lake -35
109? 107'
FIGURE 8.?Native distribution of Prosopium (shown with shading) in the Great Basin and nearby drainages,
with estimate of the phylogenetic tree (from Smith and Todd, 1993). Shading indicates drainage basins occupied
by species (as numbered in the Appendix). Symbols representing fossil ages are defined in Figure 1.
SALMONINAE: Settlers in the nineteenth century found
abundant trout in many northern Great Basin lakes and
streams (map, Figure 9), but these were quickly overexploited
(Behnke, 1992). By the 1930s, introduced hatchery fish began
to cloud the possibilities for interpreting salmonid and hydro-
graphic history (Behnke, 1981, 1992), and unfortunately mu-
seum specimens that predate these introductions are sparse.
Modern protein and mtDNA studies are, however, discrimi-
nating native from introduced stocks and are adding consider-
able historic information (Leary et al., 1984, 1987). Trouts are
adapted to cold lakes and streams and to high gradients.
Stream forms survived arid times at high elevations and gla-
cial times in large lakes. There are indications that, in addition
to trouts (Oncorhynchus), at least two char (Salvelinus) were
native to the northern Great Basin; this evidence is reviewed
in the Appendix (number 87). The fossil record of Salvelinus
is older than the fossil record of Oncorhynchus. Salvelinus
was in the Great Basin (central Nevada) and northern Idaho in
the middle Miocene and in the Snake River Plain in the late
Miocene.
The late Miocene Oncorhynchus of the Truckee Formation
differs in many characters from modern species of Oncorhyn-
chus. It possesses synapomorphies that are shared by O. clarki,
indicating it is more advanced than the basal Mexican golden
192 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
123? 12T 119* 117"
121? ?119* 117* 115
Miocene
Pliocene
Pleistocene
' i
-35'
FIGURE 9.?Distribution of Holocene and fossil species and subspecies of rainbow trout {Oncorhynchus
mykiss, stipple pattern) and cutthroat trout (O. clarki, shaded patterns) in the Great Basin and adjacent drain-
ages. Shading indicates drainage basins occupied by species and subspecies (as numbered in the Appendix).
Squares = Miocene; triangles = Pliocene; bullets = Pleistocene.
trout, O. chrysogaster. The characteristics of the Truckee trout
indicate that it existed before the differentiation of rainbow
trout and cutthroat trout (Stearley and Smith, 1993), while its
habitat was connected to the Pacific Ocean. As early as late
Miocene and throughout the Pliocene, the trout, ^Oncorhyn-
chus ("Rhabdofario") lacustris (Cope), of the lower Snake
River drainage and northwestern Nevada, shows a composite
of O. clarki (cutthroat trout) and O. mykiss (rainbow trout)
characters, e.g., O. clarki dentaries, intermediate maxillae, and
O. mykiss premaxiilae. These characters diagnose either the
ancestor prior to the clarki-mykiss divergence or the ancestral
redband trout {Oncorhynchus mykiss gairdneri). The distinc-
tion cannot be made at this time because intermediates be-
tween O. clarki and O. mykiss still exist, especially in the Co-
lumbia Basin. Characters of these trout subsequently sorted
into O. mykiss and O. clarki elsewhere, but they continue to be
introgressed or unsorted in redband trout throughout their
range in the lower Columbia River, Klamath River, and Sacra-
mento River drainages (Leary et al., 1984, 1987; Berg, 1987;
Allendorf and Leary, 1988), possibly throughout much of the
last 3 My. A sample of hatchery-produced O. mykiss differs
from O. clarki by 3.7% mtDNA (Table 3).
NUMBER 33 193
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Oncorhynchus clarki, the cutthroat trout, has been in the
Great Basin longer than proposed by Behnke (1981, 1992), ac-
cording to the Miocene fossil record. Molecular data indicate
that interbasin differentiation of O. clarki subspecies goes back
nearly 5 My, into the Pliocene (Table 3).
The Great Basin trouts form three clades. The Bonneville
Basin clade is cladistically basal, being the sister group to the
remainder of the species (Figure 10). The Snake River clade is
the next-most basal group, and it includes the Bear River form.
Although the Bear River is now a tributary to the northern
Bonneville Basin, it was part of the Snake River drainage prior
to 20-15 Ka (McCoy, 1987; Curry, 1990; Bouchard et al.,
1998). The Lahontan Basin and southern Oregon subspecies
are a sister group to the westslope trout subspecies, O. c. lewisi.
Coastal cutthroat trouts, O. c. clarki, are an intermediate clade
between the Colorado River plus Snake River groups and the
Westslope plus Oregon-Lahontan groups (Figures 9, 10).
Fort Rock Valley, west of the Harney-Malheur Basin in Lake
County, Oregon, was the site of Fort Rock Lake, also called
Fossil Lake, which was a tributary to the Deschutes and Co-
lumbia Rivers prior to blockage of the outlet by postpluvial
lava flows (Allison, 1940, 1982). Pliocene and Pleistocene sed-
iments in the Fort Rock Valley contain trouts, minnows, and
suckers. A Pliocene salmon from diatomaceous sediments of
the Fort Rock Formation, Oncorhynchus sp. (either nerka, keta,
or gorbuscha), was reported by Cavender and Miller (1972).
The trout specimens in this basin (Oncorhynchus sp.; Appen-
dix, number 90) are osteologically similar to redband trout and
were considered to date from the late Pleistocene by Allison
and Bond (1983). They based their conclusion upon shared
characters of the palatine, maxilla, hyomandibular, and orbito-
sphenoid. More than 50 dentaries from Fossil Lake were exam-
ined for the present study. Forty-seven maxillae have a mixture
of modern O. clarki (from Pyramid Lake before introductions)
features and features of Pliocene O. lacustris (pre-clarki and
pre-mykiss) of the Snake River Plain, whereas seven maxillae
resemble O. m. gairdneri (redband trout). One premaxilla re-
sembles that of O. clarki and seven resemble that of O. mykiss,
but it is not clear whether the differences represent variability
in one species. Palatines from Fossil Lake are deep, like that of
O. lacustris, and the hyomandibular is angled posteroventrally,
also like that of O. lacustris.
The Pliocene fish fauna of Secret Valley, Modoc County,
California (Oncorhynchus (Appendix, number 91), Kla-
mathella, Ptychocheilus, Lavinia, Chasmistes, and possibly Ac-
rocheilus and Cottus), and the Pliocene fish fauna of the Snake
River Plain are very similar; this suggests the two drainages
were connected prior to the Pliocene (Wheeler and Cook, 1954;
Miller and Smith, 1981; Wagner et al., 1997). A similar fauna
from the nearby Alturas Basin, Modoc County, California, con-
tains Klamathella, Ptychocheilus, Chasmistes, and Catostomus
(Wagner et al., 1997). These similarities suggest former hydro-
graphic connections (not necessarily contemporaneous) among
the Pit River, the Snake River, and the Klamath Basin.
194 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
mykiss
c. subsp. (99)
c. subsp. (99)
c. Utah (98)
c. subsp. (100)
c. bouvieri
c. bouvieri
c. bouvieri
c. pleuriticus
c. pleuriticus
c. clarki
c. clarki
c. lewisi
c. lewisi
c. lewisi
c. henshawi (93)
c. subsp. (95)
c. seliniris (94)
c. subsp. (96)
c. subsp. (97)
Trout Cr
Deaf Smith
Sevier Dr.
Bear R
Yellowstone
Snake R
Sedge Cr
Current Cr
Sheep Cr
Mack
Flynn
FishL
Wing
Cougar
Washburn
Humboldt
Silver King
Willow Cr
Whitehorse Cr
w
s
T
04 03 02 01 0
% sequence divergence
FIGURE 10.?Estimate of the phylogenetic tree of Oncorhynchus clarki (cutthroat trout) and its outgroup,
Oncorhynchus mykiss, based on restriction fragment analysis of mtDNA. Branch lengths may be read as esti-
mated time if divided by the calculated rate of 0.5% sequence divergence per 1 My. Abbreviations: BONNEV -
Bonneville Basin, SNAKE = Snake River drainage, COLO = Colorado River drainage, COAST = Coastal
drainages, WEST = west slope drainages, LAH = Lahontan Basin, OR - Oregon Lakes drainages, southeast
Oregon.
CYPRINODONTIDAE.?The Great Basin pupfishes are the
showpiece of Great Basin ichthyology. They were the focus of
R.R. Miller's doctoral research, the results of which were pub-
lished in his 1948 monograph. This monumental systematic
and experimental work illustrated the relationship between
isolation and differentiation in relict waters. Pupfishes of the
Great Basin's Death Valley system (Miller, 1946a) have been
the subject of extensive ecological and physiological studies
(Brown and Feldmeth, 1971; Soltz and Naiman, 1978; Feld-
meth, 1981), behavioral observations (Liu in Turner, 1974),
biochemical comparisons (Turner, 1984; Echelle and Echelle,
1993), and molecular analyses (Echelle and Dowling, 1992;
Parker and Kornfield, 1995). Relationships of these species
are now understood to be more complex than formerly
thought. Cyprinodon radiosus (Appendix, number 101) of the
Owens River is the sister species to C. macularius of the lower
Colorado River, and these two species are the sister group to
the clade that includes C. salinus, C. nevadensis, C. diabolis,
and relatives, in the Salt River, the Amargosa River, Devil's
Hole, and Ash Meadows (Figure 11; Appendix, numbers
102-110). This Death Valley and Ash Meadows clade possibly
includes Cyprinodon fontinalis of northern Chihuahua, Mex-
ico (Figure 11), but this is not consistent with geography or
scale morphology (Miller and Smith, 1986; Minckley et al.,
1986). Hydrographic connections among basins and with
drainages in Mexico are indicated by Cyprinodon mtDNA
analysis (Figure 11; Table 4; Echelle and Dowling, 1992; Ech-
elle and Echelle, 1993) and molluscan data (Hershler and
Pratt, 1990). Study of the data usually regarded as indicative
of the relationship between isolation and divergence reveals
interesting conflicts?molecular, biochemical, and morpho-
logical divergence and reproductive isolation are incongruent
NUMBER 33 195
C. eximius
C. variegatus
C. meeki
Whitefin west
Whitefin east
Palomas
C. m. eremus
C. m. macularius
C. radiosus
C. fontinalis
C. salinus
C. diabolis
C. n. mionectes
C. n. shoshone
C. n. amargosae
C. nevadensis sbsp
Rio Grande
Gulf Coast, Miss.
RMezquital,Mex.
RYaqui,Mex
Guzman Dr. Mex.
Palomas, Chih.,Mex.
Organ Pipe N.M., Ariz.
Colorado R., Calif.
Owens R., Calif.
Chihuahua, Mex.
Salt R., Death V.
Devils Hole, Death V.
Ash Meadows, Death V.
Shoshone Spr., Death V.
Amargosa R., Death V.
Ash Meadows, Death V.
08 07 06 05 04 03 02
% sequence divergence
01
FIGURE 11.?Phylogenetic relationships (strict consensus tree) of certain western North American Cyprinodon,
based upon parsimony analysis of mtDNA restriction site data (Data and trees from Echelle and Dowling, 1992).
The Death Valley species are C. C. nevadensis, C. salinus, and C. diabolis; C. milleri has C. salinus mtDNA.
Cyprinodon radiosus is the Owens River species, and C. macularius is in the lower Colorado River and vicinity.
The other species are primarily from Mexico (Echelle and Dowling, 1992). The depth to each node may be
scaled to age in millions of years, assuming that 1 My corresponds to 1% of the average percent sequence diver-
gence between clades.
with each other and with evidence for timing of divergence
(Turner, 1974; Echelle and Dowling, 1992; Echelle and Ech-
elle, 1993). Cladistically analysed molecular data indicate that
Cyprinodon has been in the Death Valley system for about 3
My. Biogeographic relationships of Cyprinodon (and good-
eids) of the Great Basin are to the south, in Mexico, especially
if C. fontinalis, from Chihuahua, is related to the Death Valley
clade (Figure 11). The diversity of Cyprinodon species within
the Great Basin, taken with its sister species, C. macularius
Baird and Girard, displays a geographic pattern that extends
from Death Valley and the Owens River to across the Gila
River and the lower Colorado River drainages (Minckley et
al, 1986).
Pliocene connections from the lower Colorado River to the
Owens, Mohave, and Amargosa Rivers occurred about 3 Ma,
according to molecular data (Figure 11; Table 4). The connec-
tions possibly involved China Lake, Searles Lake, Panamint
Valley, and Wingate Pass to Lake Manly, now Death Valley,
according to fish and other evidence (Miller, 1946a; Hubbs and
Miller, 1948a; Benson et al., 1990). The Pliocene Owens-Mo-
have drainage connected to the lower Colorado River drainage
(Blackwelder, 1933, 1954; Hale, 1985) near the region of Lud-
low Pass (at 594 m), Bristol Lake, and Cadiz Lake. Brown and
Rosen (1995) found no evidence of a through-flowing connec-
tion across the present topography, and they suggested alterna-
tive dispersal processes, routes, or times; however, rounded
boulders in a large paleochannel, which are evidence of a
through-flowing stream, are present near Ludlow (Hale, 1985;
G.I. Smith, pers. comm. to G.R.S., 1999). Encroachment of the
Gulf of California toward Death Valley (Durham and Allison,
1960) and the presence of foraminifera correlated with the
Bouse Formation in cores taken from the dry lakes of Danby
and Cadiz in southeastern San Bernardino County, California
(P.B. Smith, 1960, 1970), suggest a connection existed to the
estuary of the lower Colorado River. The most compelling fish
evidence for a connection between the Mohave and Colorado
Rivers is the distribution of Cyprinodon macularius of the
lower Colorado drainage and C. radiosus of the Owens River.
Mitochondrial DNA of these two species indicates they had a
common ancestry between 2 and 3 Ma (Table 4).
196 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
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G00DEIDAE.?The poolfishes and springfishes, Empetrich-
thys and Crenichthys, respectively (Appendix, numbers
112-115), are Great Basin endemics in the subfamily Empet-
richthyinae (Webb, 1998) of the family Goodeidae (Parenti,
1981; Grant and Riddle, 1995). Empetrichthys and Crenichthys
are oviparous and were formerly classified as cyprinodontids,
but Parenti's (1981) phylogenetic evidence demonstrated their
relationship to the live-bearing goodeids of the Mexican Pla-
teau. Empetrichthyines resemble fundulids in osteology, num-
bers of fin rays, and tooth form, which distinguish them from
the cyprinodontids, but unlike fundulids, they lack pelvic fins.
Fossil Empetrichthys were discovered from the Pliocene near
Los Angeles, California, and from the Pleistocene in southern
Nevada (see Appendix, number 112). Paleontological studies
by M.L. Smith (in Miller and Smith, 1986) and molecular stud-
ies by Webb (1998) indicate middle to late Miocene goodeid
vicariance between the White River and the lower Colorado
River drainages and the Mexican Plateau, probably involving
tributaries of the Bouse Embayment and the Gulf of California.
The distribution of fossil and Holocene Empetrichthys, Fundu-
lus, and Gasterosteus, along with species of Pyrgu I ops is (hy-
drobiid snails), provide evidence that late Miocene lowland Pa-
cific coast tributaries flowed westward from west-central
Nevada and Death Valley across a lower, proto-Sierra Nevada
(Bell, 1979; Hershler and Pratt, 1990) about 6 Ma, perhaps
through the (reversed) Mohave River valley (Weldon, 1982).
FUNDULIDAE.?Most species of killifish in the family Fun-
dulidae live in eastern North America, but there are representa-
tives in Mexico and California. Several fossil forms were de-
scribed (Miller, 1945b) before it was recognized (Parenti,
1981) that the Goodeidae of the Mexican Plateau, relatives of
Fundulidae, are represented in the Great Basin. The Nevada
fossils (Appendix, numbers 116-123) have some observable
characters that suggest they could be related to the Empetrich-
thyinae or the Goodeidae (Miller, 1948:100), except that none
of the Nevada fossils lack pelvic fins, as do the empetrichthy-
ines.
The ancient diversity of fundulids in the southwestern Great
Basin is remarkable, considering the paucity of other genera
(except Siphateles and Gasterosteus) in the fossil record from
there (M.L. Smith, 1981) and the absence of Fundulus in the
Great Basin today. The presence of fundulids from northeast
(Humboldt Formation) and west-central (Truckee Formation)
Nevada and Death Valley through most of the Miocene and
Pliocene establishes a benchmark of antiquity that is important,
given that the ages of origin of Great Basin fish groups are un-
derestimated (see "Material and Methods" and the Appendix).
COTTIDAE.?The Cottidae, like the Salmonidae, are low-
temperature, usually high-gradient fishes. They occupy moun-
tain streams and cold lakes. Their fossil record in the Great Ba-
sin goes back to the Pliocene and Pleistocene in the Lahontan
Basin (C. beldingi) and to the late Pleistocene in the Bonneville
Basin (C. bairdi, C. extensus). Their phylogenetic relationships
are to the north and west. More species of northwestern
NUMBER 33 197
sculpins occur in the northern Bonneville Basin and Bear River
drainages than are currently acknowledged in the literature, ac-
cording to new discoveries by Shiozawa and others (unpub-
lished). Sculpins are not found with cyprinodontoids in western
North America, unlike their occasional sympatry in the eastern
United States, reflecting the more extreme habitat gradients
(temperature and stream profiles) in the west.
GASTEROSTEIDAE.?The Miocene three-spined stickleback
{Gasterosteus doryssus; Appendix, number 128) occurs in the
Great Basin from the Truckee Formation in west-central Ne-
vada from Hazen (Bell, 1974) to the Sahwave Mountains
(LaRivers, 1964). It represents a fauna that colonized the basin
prior to final elimination of the drainage outlet to the Pacific
coast by the Sierra Nevada uplift (Bell et al., 1985). The Ne-
vada Gasterosteus is the most significant fossil fish lineage in
the Great Basin, if not in Cenozoic ichthyology in general, be-
cause of the abundant, beautifully preserved samples in annual
laminae of diatomite and the detailed studies of these samples
by Bell and his students (Bell and Haglund, 1982; Bell, 1994,
and references therein). In brief, the study of more than 8000
specimens from a 110,000-year time sequence has shown rap-
idly fluctuating but gradual change, with no stasis (Figure 12).
Bell (1994) concluded that much of the change can be attrib-
uted to selection, but some seems uncorrelated with environ-
mental change.
Gasterosteus aculeatus aculeatus (Appendix, number 129),
the modern, anadromous, fully plated, three-spined relative,
was reported from the late Miocene Monterey Formation at
Lompoc, California, by Bell (1977), strengthening the hypothe-
sis that inland freshwater populations of sticklebacks were in-
dependently derived from a widespread anadromous form
(Bell, 1979). Comparison of these sticklebacks demonstrates
that new species can arise repeatedly, as restricted peripatric
and paraphyletic derivative populations, without effects on the
ecology, morphology, or long-term history of the widespread,
ancestral species. Bell's point demonstrates the logical difficul-
ties attending Hennig's rule (1966), which requires changing a
0 10 20 30 40 50 60 70 80 90 100 110
Time ( x 1000 years)
FIGURE 12.?Variation in mean number of dorsal spines in Gasterosteus
dorysus. Variation shows rapid short-term changes and slow long-term trend
(after Bell et al., 1985).
species name with each branching event in a clade's history, if
the parent species did not change while daughter lineages were
produced (Bell, 1979).
CENTRARCHIDAE.?Sunfish occupied lakes and streams in
western North America from Idaho, Washington, and Oregon to
Nevada and northern Utah (Salt Lake group) during the Miocene
and Pliocene. The Nevada specimens suggest that the climate
during the middle to late Miocene was warm and that the basin
floor was at low elevations, but floristic data indicate the floor
was at higher elevations (Wolfe et al., 1997). The combination
suggests local topographic extremes. There was more diversity
during the Miocene than is recognized in the currently defined
genera Archoplites and Plioplarchus. An undescribed form ex-
isted in the Miocene of Oregon (Appendix, number 130). Mod-
ern centrarchids are restricted to Archoplites of the Great Valley,
California, and to a radiation of nine genera and more than 30
species in the Atlantic coastal drainages of North America.
Discussion
FISH BIOGEOGRAPHY AND GEOLOGICAL TIME
RELATIONSHIP BETWEEN PHYSIOGRAPHIC AND BIOLOGICAL
EVIDENCE.?Hubbs and Miller's search for congruence be-
tween physiographic and ichthyological histories is evident
throughout their 1948b monograph. They noted, for example,
the local morphological distinctiveness of Siphateles, Rhinich-
thys, and Cyprinodon as evidence of postpluvial separation of
previously uniform gene pools (Hubbs and Miller, 1948b:66,
83, 91). The seemingly parsimonious assumption that faunas
were uniform at the time of the last habitat connection has
rarely been questioned (Kocher and Stepien, 1997), but even in
1948, it was in conflict with the fossil record of modern fishes,
which, when known, extended back beyond Pleistocene events
to the Pliocene or Miocene (Hubbs and Miller, 1948b:25-27).
Taylor (1960) noted the antiquity of many modern lineages
of mollusks and fishes and suggested that they are much older
than modern physiography, ice ages, and pluvials. The corol-
lary of this suggestion is that the tectonic, volcanic, and glacial
changes in the landscape occur more rapidly than the evolving
changes in lineages on that landscape. Taylor's principle sug-
gests that we cannot assume that the observed amount of spe-
cies and subspecies evolution automatically occurred in the in-
terval since the most recent possible geological changes and
vicariance events. Multiple possible vicariant times should be
among the tested alternatives. For example, the assumption that
cladogenesis of Holocene birds originated in the Pleistocene
was refuted by Klicka and Zink (1997), who demonstrated that
numerous supposedly Pleistocene species date back to the
Pliocene. Our interpretations of fossil fish specimens combined
with DN A analyses and phylogenetic analyses of Holocene
species indicate that the ages of clades and vicariance barriers
have usually been underestimated, thereby challenging the fre-
quent assumption that the most recent possible connection was
the operative one.
198 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Phylogenetic estimates for the Cyprinidae (Figures 2, 3), Ca-
tostomidae (Figure 5), Coregoninae (Figure 8), Oncorhynchus
(Figure 10), and Cyprinodontidae (Figure 11) are presented in
the figures with key words for taxon geography to provide
background for examining the relative timing of historic hy-
drographic connections among basins. Traditional studies
(Hubbs and Miller, 1948b; Smith, 1966, 1978; Hubbs et al.,
1974; Minckley et al., 1986) and the new phylogenies enable
the formulation of hypotheses concerning the relative ages of
barriers.
According to the phylogenies, moderately old and prominent
barriers separated the Great Basin from its neighboring drain-
ages to the east and the Sacramento Basin to the west. Phyloge-
netic indicators of the barrier between the Great Basin and the
Colorado Basin, for example, are the relative antiquity of the
nodes between the Gila atraria group and the G. robusta group
(Figures 2, 3), between Xyrauchen and Chasmistes, between
Catostomus ardens and C. latipinnis, between C. platyrhyn-
chus and C. discobolus (Figure 5), and between Oncorhynchus
clarki Utah and O. c. pleuriticus (Figures 9, 10). The impor-
tance of the Sierra Nevada as a barrier is indicated by the sepa-
ration of Catostomus occidentalis and C. microps from C.
tahoensis (Figure 5), and by Oncorhynchus mykiss from O.
clarki (Figures 9, 10). A strong barrier separated the Bonne-
ville and Lahontan Basins along the original line of separation
in the Great Basin, as illustrated by the relative distance be-
tween nodes, for example, between nodes for Gila and Siphate-
les (Figures 2, 3), C. ardens and C. tahoensis (Figure 5), and O.
c. Utah and O. c. henshawi (Figures 9, 10). The moderately
strong separation of the Lahontan Basin and Colorado Basin
drainages is illustrated by the nodal distances between Gila and
Siphateles (Figures 2, 3), C. tahoensis and C. latipinnis (Figure
5), and O. c. henshawi and O. c. pleuriticus (Figures 9, 10).
The weakest, most recent barriers separated the Lahontan Ba-
sin and Death Valley systems, and the subdrainages within the
Bonneville, Lahontan, White River, and Death Valley systems
(most examples are similar species or subspecies not shown on
the phylogenies), but all of these barriers have been breached
by headwater-stream captures, indicated by some of the exam-
ples below and in Figures 2, 3, 5, and especially Figure 10.
Fish distributions suggest that much of the drainage from the
area of the Great Basin flowed to the Pacific Ocean at times
during the Miocene, Pliocene, and Pleistocene, even though the
Basin and Range topography began forming more than 15 Ma.
The fish data presented herein indicate that these hydrologic
confluences probably had an episodic history that began much
earlier than the most recent known connection.
MIOCENE.?Geological and paleontological studies indicate
that at times during the Miocene, the Death Valley, Lahontan,
and Bonneville systems drained westward to the Pacific Ocean
(Bell, 1974, 1994; Stokes, 1979; Taylor, 1985; Minckley et al.,
1986; G.I. Smith, pers. comm. to G.R.S., 1999). Some Mio-
cene biogeographic patterns, however, were dominated by
north-south systems (Taylor, 1985) that were congruent with
the north-south structures that resulted from convergence of
the west coast of North America with the Pacific Plate (Figure
13). Early hydrographic connections occurred between the
Humboldt lacustrine system and the western Snake River Plain
and southeastern Oregon, according to Stokes (1979) and evi-
dence from the distribution of Centrarchidae. Later in the
Hemphillian, the Salt Lake group, with a relatively rich fish
fauna in the northern Bonneville Basin, was probably con-
nected to the eastern Snake River Plain (McClellan, 1977). The
fish fauna was related to, but specifically different from, that of
the western Snake River Plain (McClellan and G.R. Smith, In
press). The presence of Lavinia in the lower Salt Lake group
indicates there were earlier connections to the western Snake
River Plain and the Pacific coastal drainages. Lavinia is now
restricted to the Sacramento drainage. The mixture of lowland
fish faunas and upland floras in Nevada during the Miocene
(Wolfe et al., 1997) suggests that evolution of regional topo-
graphic relief took place subsequently (Orr, 1982; Pierce and
Morgan, 1992).
LATE MIOCENE TO PLIOCENE.?The southern Lahontan Ba-
sin fossils from the late Miocene and Pliocene are indicative of
low elevations, but the northern Lahontan Basin fossils are in-
dicative of higher elevations. Elevations of the Snake River
Plain were low in the late Miocene, but uplift of the eastern
Snake River Plain at this time occurred in the path of the Yel-
lowstone hot spot (Pierce and Morgan, 1992).
A connection from the Mono Lake area of the Great Basin to
the San Joaquin River ended when basalt dammed the V-
shaped channel near Deadman Pass in the Devil's Postpile
quadrangle (Huber and Rinehart, 1965, 1967), causing cessa-
tion of river flow 3.2 Ma (Dalrymple, 1964; Huber, 1981). The
channel is now at about 8400 ft (2580 m) elevation, having
been raised 3120 ft (950 m) by uplift of the Sierra Nevada
(Dalrymple, 1964; Huber, 1981), so the estimated elevation of
the active channel was 5280 ft (1620 m; G.I. Smith et al., 1983;
G.I. Smith, pers. comm. to G.R.S., 1999). The outlet possibly
continued from the Great Valley across the Diablo Uplift to de-
posit the Monterey Fan (Cole and Armentrout, 1979). Cyprinid
and gasterosteid fish connections require lower elevations dur-
ing the late Miocene, based upon their present elevational re-
strictions.
PLIOCENE.?Our phylogenetic data suggest the occurrence
of some ancient drainage connections that had not been previ-
ously noted (Figure 13). For example, the similarity of mor-
phology and DNA (see below) of Lepidomeda from the Virgin
River drainage and Snyderichthys from the upper Snake River
and the Bear River drainages suggests there were connections
from the upper Snake River to the Virgin River-White River
drainage (Dowling et al., 2002). Keyes (1917), in a discussion
of the origins of Lake Bonneville, proposed that the headwaters
of the Snake River were formerly a tributary to the Virgin
River. Ives (1948) presented evidence that Lake Bonneville
drained from its Escalante arm to Meadow Valley Wash (a trib-
utary to the Virgin River) prior to its drainage through Red
NUMBER 33 199
35
>?? ? ? ? Miocene
'.??' Pliocene
-35'
121* ?119*
? Pleistocene
in*
FIGURE 13.?Paleogeographic map showing major Miocene, Pliocene, and Pleistocene hydrographic con-
nections, based upon our current interpretations of age and location and supplemented by the conclusions of
Cole and Armentrout (1979), Taylor (1985), Taylor and Bright (1987), Ives (1948), and G.I. Smith (pers.
comm. to G.R.S., 1999). Heavy dashed lines and arrows represent captures, overflows, or rivers. See "Discus-
sion" and Appendix for the estimated timing of some of the barriers represented by lightweight lines.
Rock Pass to the upper Snake River. Hubbs and Miller (1948b)
rejected these hypotheses as unsupported by fish evidence.
Possible connections require serious consideration now be-
cause of new cyprinid and molluscan evidence. Miocene and
Pliocene distributions of freshwater snails, namely, Valvata
idahoensis Taylor, Lutralimnea dineana (Taylor), L. gentilis
Taylor, and Tryoniaprotea (Gould), led Taylor (1985:285, 296,
317) to propose late Miocene and possibly Pliocene connec-
tions (not necessarily contemporaneous) from the upper Snake
River, through the Bonneville area, to the lower Colorado
River drainage.
The White River canyon in southern Nevada is surrounded
by possible drainage connections, as indicated generally by
Taylor (1985, figs. 15, 25), leading to the Bouse Embayment of
the Gulf of California in the late Miocene or early Pliocene
(Shaffiqullah et al., 1980; Buising, 1990) and in the Pleis-
tocene (DiGuiseppi and Bartley, 1991). The Colorado Paleo-
river was diverted from its delta in the Los Angeles Basin to
200 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
the Salton trough about 9 Ma (Howard, 1996). The deposi-
tional elevation of the Bouse Formation (near the Arizona-Ne-
vada border) is usually cited as at sea level, constraining the
time of uplift of the Colorado Plateau, but this is currently in
dispute. Bouse Formation sediments and shells show ratios of
strontium isotopes intermediate between values expected for
marine and freshwater, indicating this area had a mixture of
water sources with different strontium isotopes, according to
Spencer and Patchett (1997). These authors' contention that
the formation was deposited in a lake, possibly at an elevation
of 400-700 m, is rejected on the basis of fossil and biogeo-
graphic data; the Bouse Formation contained a fauna of fora-
minifera, barnacles, brackish-water fishes, and marine mol-
lusks (Metzger, 1968; Smith, 1970) that could not have been
supported for 3 My by dispersal on birds' feet, as suggested by
Spencer and Patchett (1997). The hypothesis that the Bouse
Formation was lacustrine (Spencer and Patchett, 1997) has
been tested and, by inference, rejected by the absence of any
such marine faunas in saline lakes today (Kidwell in Spencer
and Patchett, 1997).
Pliocene fluvial connections between the northern pre-
Lahontan Basin and the Alvord Basin of Oregon are indicated
by geological data (Reheis and Morrison, 1997) and fish dis-
tributions (Appendix, numbers 9, 22, 91, 96, 97). DNA diver-
gence data (below; Table 3) suggest that reconnections 0.2
Ma allowed trout populations in these basins to once again
mix.
PLIO-PLEISTOCENE.?Connections among the Mono Lake,
Owens River, Death Valley, Mohave River, and lower Colo-
rado River drainages are indicated by cyprinodontid fish distri-
butions (Miller, 1948) and geological data (Gale, 1915; Morri-
son, 1965; G.I. Smith, 1978; Benson et al., 1990). The
connection between the Mohave River and the Colorado River
basins through Bristol, Cadiz, and Danby Basins lacks support
from geomorphology, according to Brown and Rosen (1995),
but the connection suggested by present Cyprinodon distribu-
tion is supported by sedimentary data (Hale, 1985) and by ma-
rine or saline-lake foraminifera from cores in Danby and Cadiz
dry lakes (Smith, 1960). The foraminifera occur in green clays
that are currently within a few hundred feet above or below sea
level.
PLEISTOCENE.?Pleistocene connections occurred on a land-
scape more like that of the present (Figure 13). Additional evi-
dence for exchange between the Lake Bonneville and Virgin
River drainages, possibly in the Pleistocene, is seen in Catosto-
mus clarki from the headwaters of the Virgin River. These fish
have unique predorsal scale and gill-raker counts that increase
clinally toward their relatives across the divide in the Bonne-
ville Basin (Smith, 1966). Numerous possible stream captures
are associated with the major north-to-south valley west of
Hurricane Fault (Averitt, 1964; Anderson and Mehnert, 1979)
in southern Utah (along Highway 91). The drainage connection
between the Bonneville Basin and the Colorado River drainage
is also suggested by intergrades between Catostomus platy-
rhynchus and C. clarki in Shoal Creek (Smith and Koehn,
1971), which is in the southernmost part of the Bonneville Ba-
sin near the outlet proposed by Ives (1948).
Small headwater-stream captures are required to explain the
broad northern distributions of headwater inhabitants, such as
Catostomus platyrhynchus, C. discobolus (Figure 7), Onco-
rhynchus clarki subspecies (Figure 9), and Coitus spp. (Appen-
dix, numbers 124, 125). Numerous examples of northern con-
nections can be recognized on topographic maps and in the
field. For example, in Wyoming, two such possible connections
may be related to the distribution of Catostomus platyrhynchus,
C. discobolus, and Coitus bairdi in the Green, Bear, and Snake
Rivers. Willow Creek, formerly tributary to Muddy Creek of
the Green River (T.I3 N., R.I 19 W., Ogden sheet, United
States Geological Survey (USGS) 1:250,000 series), was cap-
tured by the headwaters of the Bear River 11 mi (17 km) south-
east of Evanston, Wyoming. Twin Creek, of the Bear River be-
tween Fossil Butte and Kemmerer, Wyoming (T.21 N., R.I 17
W., Ogden sheet), apparently captured the head of Hams Fork
of the Green River. We hypothesize that these captures could
have carried Catostomus discobolus from the Green River
drainage to the Bear River, from whence they could have trans-
ferred to the Snake River at Soda Springs (Bright, 1967) and to
the Weber River in upper Chalk Creek (Hansen, 1969, 1985).
MLTOCHONDRIAL DNA EVIDENCE FOR THE TIMING
OF PHYLOGENETIC AND HYDROGRAPHIC EVENTS
Analysis of DNA divergence data from fishes offers a means
of constraining ages of barriers and uplift in the Great Basin.
Estimated rates of evolution of mtDNA genes in Cyprinidae,
Salmonidae, and Cyprinodontidae (Figure 14) vary from about
0.5% sequence divergence per one million years in salmonids
to about 1 % per one million years in cyprinids and cyprinodon-
tids. These values are similar to those reported for snails of the
marine gastropod genus Nucella (see Collins et al., 1996) and
for goodeid fishes (Webb, 1998) but differ from the estimate of
Martin and Palumbi (1993). The relative-rate tests (e.g., Table
2) show standard deviations of 10%?25% of the mean values of
percent sequence divergence, or roughly 2-5 My for mean ages
of 7-38 My (Table 2) for basal branches in the cyprinid data.
Standard deviations are 15%?30% in the salmonid data, allow-
ing a rough idea of the imprecision of the estimates. These are
overestimates of rates of molecular change because the fossil
ages in the denominators of the rate calculations are underesti-
mates. If the rates are underestimated by 10%?30%, elapsed
times inferred from rates of molecular change are expected to
be underestimates (by as much as 0.5-1.2 My in 4 My for
spinedaces, for example). This means that the connection indi-
cated by the Snake River Snyderichthys relationship to the Vir-
gin River Lepidomeda (percent sequence divergence = 4%)
probably occurred within the range of 4-5 Ma.
In general, divergence times based upon the rates calculated
herein are old, as predicted by Minckley et al. (1986). Great
Basin cyprinids in the sister clade to Phoxinus of eastern
North America differ from Phoxinus by an average of 22%
NUMBER 33 201
.50 -I
8 .40c
CD
.30-
CD
.20-
CD
CD
CO
y = 0.011x + 0.05
R2 = 0.28
? ?
Cyprinidae
10 15
Million years before present
CD y = 0.0047X + 0.007
R2 = 0.65
4 6 8 10
Million years before present
12 14
CDO
c
CD
> .15 i
CD -if
CD
CT
CD
CO
.05-
y = 0.097x + .036
R2 = 0.69
Cyprinodontidae
5 10 15
Million years before present
20
FIGURE 14.?Rates of molecular change in Great Basin and Snake River plain Cyprinidae, Salmonidae, and
Cyprinodontoidei. Estimated ages of fossils are given in the Appendix and Table 5. DNA sequence divergences
are from Tables 2-5; R2 is an estimate of the percent of the variance explained by the regression.
sequence divergence (Table 2), implying that separation oc-
curred 22 Ma, or in the early Miocene. Rhinichthys osculus
from the Virgin River and R. atratulus of eastern North Amer-
ica differ by 22%. Snyderichthys and the Lepidomeda group
differ from Semotilus of eastern North America by 20%.
Closer relatives across the continental divide generally differ
by less. For example, Clinostomus and Richardsonius differ
by 12% sequence divergence, and eastern and northern North
202 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
TABLE 5.?Fossil cyprinid, salmonid, and cyprinodontoid taxa and ages of samples of first appearances used
in the estimations for the rates of molecular change in Figure 14. Fossil taxon numbers correspond to num-
bers in the Appendix; unnumbered taxa are from outside of the Great Basin. Ages in parentheses are addi-
tional fossil records for that taxon. (Fm. = formation.)
Fossil taxon
CYPRINIDAE
1. Lavinia
2. Lavinia
3. Lavinia
5. Mylopharodon
5. Mylopharodon
5. Mylopharodon?
6. Siphateles
6. Siphateles?
7'. Siphateles
8. Siphateles
9. Siphateles
10. Siphateles bicolor obesa
14. Siphateles sp.
15. Siphateles sp.
23. Gila a. atraria
25. Klamathella
26. Acrocheilus
26. Acrocheilus
27. Gila
28. Gila
29. Gila
Gila spp.
30. Ptychocheilus
37. Richardsonius
56. Mylocheilus
57. Mylocheilus
58. Mylocheilus
SALMONINAE
87. Salvelinus
88. Oncorhynchus
89. O. clarki (Truckee Fm.)
90. 0. clarki (Ft. Rock Basin)
91.0. clarki (Honey Lake)
92. O. clarki (Carson Sink)
92. O. clarki (Mopung Hills)
92. O. clarki (Humboldt River Canyon)
97. O. clarki (Whitehorse Basin)
97. O. clarki (Sunshine Amphitheater)
98. O. clarki (Sevier Basin)
98. O. clarki (Trout Creek)
O. nerka1
O. keta1
O. rastrosus^
O. keta[
CYPRINODONTOIDEI
Chapalichthys encaustus2
Ameca splendens2
Alloophorus robustus1
Goodea atripinnis2
Girardinichthys multirad2
Tapatia2
111. Cyprinodon breviradius
113. Empetrichthys erdisi
122. Fundulus lariversi
Sister taxon
Mylopharodon
Mylopharodon
Mylopharodon
Siphateles
Lavinia
Eastern minnows
Gila
Eastern minnows
Gila
Gila
Gila
Mylopharodon
Gila
Gila atraria subsp.
Gila
Klamathella
Klamathella
Siphateles
Siphateles
Siphateles
Ptychocheilus
Gila
Clinostomus
Richardsonius
Richardsonius
Richardsonius
Hucho
O. clarki + O. mykiss
O. mykiss
O. mykiss
O. mykiss
O. clarki
O. clarki
O. clarki
(Snake River)
(Sevier River)
O. gorbuscha
O. gorbuscha
O. nerka
O. nerka
Chapalichthys pardalis
Xiphophorus variatus
Ameca, etc.
Ataeniobus. toweri
G. viviparus
other Girardinichthyini
C. spp.
Crenichthys
Goodeidae
Age (My)
6(3.5,3)
5(3)
6
12
4.5
12
6(6)
15
3
6(6)
3(3)
2 (0.6)
8
5(3,1)
0.01
6
6
7
12
12
12
6
15(6,6,8)
3.5
7
6
12
14
14
6
3.5
3.5
3.5
3
0.7
0.15
1
0.01
0.01
6
6
12
6
0.25
0.25
0.25
0.25
5.0
6.0
5.5
4.0
17
% seq. div.
8.1
8.1
8.1
15.0
8.1
42
11.5
36
11.5
11.5
11.5
3.1
15.0
11.5
0.9
6.2
6.4
6.4
11.5
11.5
11.5
10.3
10.4
12.0
21
21
21
4.8
10
4
4
4
1
1
0.1
0.1
0.1
1.3
1.0
5
3
5
6
0.6
4.7
6.2
9.1
7.1
12.2
2.9
7.0
22
'Data from G.R. Smith (1992a).
2Data from M.L. Smith (1980) and Webb (1998).
NUMBER 33 203
American Couesius differ from Lepidomeda by 16% (Dowl-
ing et al., 2002), implying that separations occurred in the
early to middle Miocene. Great Basin Gila differ from Colo-
rado River and Rio Grande Gila by 4%-9% (except for north-
ern Gila atraria, G. jordani, and G. cypha, which differ by
0.4%?1.1% sequence divergence, implying there have been
relatively recent introgressive transfers of mtDNA; Table 2).
Similarly, Great Basin Gila and their relatives in the Colorado
River, Rio Grande, Rio Yaqui, and in waterways in nearby
valleys in the Basin and Range Province of northern Mexico
differ among themselves by 5%?9%, implying lower eleva-
tions and more connections existed to the south 5-9 Ma (Ta-
ble 2; Dowling, unpublished data). Great Basin Gila differ
from Columbia-Snake Acrocheilus and Klamathella by 6%
sequence divergence, but Great Basin Siphateles differ from
the same taxa by 10%-ll%, implying separation occurred
well back in the Miocene. Differences between conspecific
populations of Gila and Snyderichthys within the Bonneville
Basin range up to 2.4%, indicating subspecific divergence be-
gan in the late Pliocene.
Populations of Oncorhynchus clarki (cutthroat trout, Table 3)
differ from O. mykiss (rainbow trout and redband trout) by
about 4% sequence divergence, implying separation occurred
about 8 Ma (assuming rates of molecular evolution of about
0.5% sequence divergence per 1 My). Populations of Onco-
rhynchus clarki in the Great Basin (except for Deaf Smith
Creek and Bear River) differ from populations of the west
slope (Columbia River) and the Snake-Yellowstone-Bear Riv-
ers by l%-2% sequence divergence, implying separation oc-
curred about 4-2 Ma. Bonneville Basin populations are hetero-
geneous. The Deaf Smith Creek (northern Bonneville)
population differs from all other trouts by about 2% more than
do other Bonneville trout populations (indicating rapid change
occurred in the Deaf Smith Creek population because it is not
cladistically basal, Figure 10). The Bear River (most northern
Bonneville) population differs from the upper Snake River
population by only 0.3%-0.4% sequence divergence, which
implies separation took place about 0.7 Ma, which is roughly
consistent with part of the Bear River chronology of Bouchard
et al. (1998). Populations of the Bonneville and Lahontan Ba-
sins differ from each other by an average of 2% sequence di-
vergence, implying they have been separated for about 4 My.
Outside the Bonneville Basin, variation within basins is usually
less than 0.1%. Lahontan Basin populations differ from those
from the Oregon Lakes by only 0.1% sequence divergence, im-
plying they have been separated for about 200 Ky.
Great Basin pupfish populations are related to Cyprinodon
macularius and its relatives of the Colorado River and Mexico
(Echelle and Dowling, 1992; Echelle and Echelle, 1993; Figure
11, Table 4). The Owens pupfish, C. radiosus, is a member of
the C. macularius clade but not the Death Valley clade; C. ra-
diosus differs from C. macularius by 2.7%-2.8% sequence di-
vergence. Death Valley pup fishes differ from C. macularius
and C. radiosus by about 3% sequence divergence, implying
separation occurred about 3 Ma. Cyprinodon fontinalis of Mex-
ico may be a member of the Death Valley clade. These relation-
ships require multiple connections between the Death Valley
system and drainages in Mexico, through the Colorado River
drainage in the Plio-Pleistocene. Cyprinodon salinus of the
Salt River in Death Valley differs from other Death Valley
pupfishes by 1.3% sequence divergence, implying separation
took place a little more than 1 Ma. Cyprinodon diabolis and
several subspecies of C. nevadensis form a clade whose popu-
lations differ from each other by about 0.2%-0.6% sequence
divergence, implying separations occurred 0.2-0.6 Ma. Miller
(1948:85-86) concluded that the divergences of C. salinus, C.
diabolis, and C. nevadensis date from about the middle Pleis-
tocene, which is consistent with the molecular differences
given above (Table 4). For divergence of some Death Valley
populations to have been as rapid as hypothesized by Hubbs
and Miller (1948b:82)?as short as 10 Ky?would require un-
realistically high rates of sequence divergence.
SPECIES DIVERSITY
Fishes of the Great Basin are not diverse, except in endemic
species and subspecies (Hubbs and Miller, 1948b). Most of
these taxa are explainable as relicts of repeated cycles of ex-
tinction or introgression during dry periods between Pleis-
tocene pluvials. The subspecies are the product of rapid change
in morphology in isolated, small populations in rapidly chang-
ing environments (Hubbs and Miller, 1948b). But low diversity
in a landscape so divided by barriers (permitting divergence
and recolonization following divergence) is not consistent with
most hypotheses about geographic and climatic control of spe-
ciation processes (Hubbs and Miller, 1948b; Mayr, 1963; Vrba,
1993). Great Basin fossil and Holocene fish data indicate, how-
ever, that periods of aridity and insularity, coupled with high
barriers that prevent recolonization after local extinctions, have
eliminated more species than were produced by the numerous
stages of isolation; therefore, the processes that cause species
divergence have not matched extinction at the spatial and tem-
poral scales of isolation that have occurred in the Great Basin.
High extinction rates have been predicted for small, isolated
populations (MacArthur and Wilson, 1967; Jablonski, 1980).
To determine whether speciation has been influenced by these
factors, we considered the processes that have operated in Great
Basin fish populations in relation to isolation by pluvial cycles.
Two unusual kinds of data lend themselves to empirical anal-
yses of these two problems. First, the Basin and Range Physio-
graphic Province is divided into dozens of pseudo-replicated
fish habitats isolated by mountain ranges, making a highly con-
strained physiographic region. This permitted us to trace lin-
eages with rather well-controlled histories. Second, the region
has a more detailed fossil fish record (Figure 1) than is avail-
able for most modern fish faunas, and this record is supported
204 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
by a reasonably well-known stratigraphic context (e.g., Benson
et al., 1990; Morrison, 1991; Oviatt et al., 1992) and by mam-
malian stratigraphy (e.g., Stirton, 1936; Lindsey et al., 1984;
Lundelius et al., 1987; Repenning, 1987; Tedford et al., 1987;
Woodburne, 1987). These circumstances enabled us to begin to
address some of the most difficult problems in evolutionary
biogeography: Over what time scales do speciation processes
occur? How long does it take for differentiation of races, sub-
species, and species to occur in isolation? Are the rates of dif-
ferentiation interval-dependent (Gingerich, 1986), i.e., is the
concept of a general rate inappropriate (Bookstein et al., 1978;
Bookstein, 1986; Bell, 1994)? Consideration of questions
about species and their origins requires a species concept that is
open to testable hypotheses concerning processes of origina-
tion.
SPECIES CONCEPTS AND SPECIATION HYPOTHESES.?Three
species concepts stand out as most relevant to models of fresh-
water fish speciation in the Great Basin: phylogenetic, biologi-
cal, and evolutionary. The phylogenetic species concept (Crac-
raft, 1983) focuses on diagnosable, monophyletic, and
irreducible units that maximize the recognition of biotic diver-
sity. Many more phylogenetic species could be recognized than
are currently acknowledged in the Great Basin. The phyloge-
netic species concept is often rejected because its lower limits
are not well constrained (Mishler and Donoghue, 1982) and be-
cause many species are not monophyletic. Local demes, such
as the subspecies and races of Cyprinodon nevadensis, Siphate-
les bicolor, and Rhinichthys osculus, could be recognized as
species under this model. Moritz et al. (1995) argued for use of
mutually and reciprocally unique differences as species criteria.
This concept requires no information about time, monophyly,
or reproductive isolation, so it is independent of evolution of
reproductive isolation. The addition of reproductive and histor-
ical tests to proposed species enables more complete evaluation
of speciation models.
The biological species concept (Mayr, 1963) recognizes
groups of populations that are reproductively isolated from
other such groups. The focus on reproductive isolation makes
this concept central to most speciation process models. Species
so defined often contain geographically variant populations, so
that in the Great Basin the concept is realistic and objective but
is operational only by inference. This definition is sometimes
applied too restrictively to accommodate some Cyprinidae, Ca-
tostomidae, and Salmonidae, for which there is evidence of
gene flow between historically well-differentiated lineages that
diverged several million years ago (Smith, 1992a). The biolog-
ical species concept is improved for our purposes by integra-
tion with the ecological species concept of Van Valen (1976,
1988), that is, accommodation of limited gene exchange be-
tween long-term, ecologically independent lineages.
The evolutionary species concept (Simpson, 1951; Wiley,
1981; Mayden and Wood, 1995) requires evidence that lin-
eages have separate evolutionary histories, usually diagnosed
in a cladistic context. This concept does not require (or rule
out) evidence for reproductive isolation as evidence for lineage
independence. The evolutionary species concept is applicable
to Great Basin fish lineages to the extent that introgressed but
still-separate lineages are not excluded because of non-mono-
phyly or temporary lapses in independence.
The time required for acquisition of genetically based ecolog-
ical or reproductive isolation as a guarantor of lineage indepen-
dence is a matter of special concern in the study of speciation
in the Great Basin because of the large number of diagnosable,
but possibly ephemeral, allopatric populations. Western North
American fishes show a high degree of allopatry, but they also
provide unusually frequent examples of gene exchange be-
tween species of long-standing independence (Hubbs, 1955).
Great Basin species of minnows, suckers, trouts, pupfishes, and
sculpins are especially consistent with species concepts and
origination models that require reproductive independence, but
not necessarily isolation. Modest introgressive influence does
not contradict this definition, if the morphologically and eco-
logically different lineages are not amalgamated through hy-
brids with high fitness. Reproductive independence and repro-
ductive isolation are often assumed to be synonymous and
mutually required for species recognition, but in many species,
it is worth making a distinction. Evidence for introgression is
usually taken to be evidence against independence (by zoolo-
gists, not botanists; Wagner, 1983; Rieseberg and Wendel,
1993), but if forms are ecologically different enough to be dis-
ruptively selected, introgression does not necessarily reduce
two lineages to one.
For example, the ecologically distinct pairs Catostomus
ardens and Chasmistes liorus of Utah Lake belong to clades
that diverged 3-6 Ma. Introgression occurred in the Provo
River at the time of the drought in the early 1930s (Miller and
Smith, 1981), but morphological stability of C. ardens and eco-
logical persistence of (modified) C. liorus indicate sufficient
reproductive separateness between these forms to warrant spe-
cies rank. In contrast, some pairs of populations of Siphateles
bicolor obesus and S. b. pectinifer in the Lahontan Basin failed
to stay separate through recent times (Hubbs and Miller,
1948b). These different outcomes are ecology dependent (Van
Valen, 1976, 1988) in that the reproductive isolation is main-
tained only by spawning-site selection based upon ecological
characteristics?depth, current, light, temperature, and sub-
strate. Evolution of lineage independence, therefore, is partly a
function of ecological context (Smith, 1992a; Smith et al.,
1995). The introgression hypotheses of Hubbs and Miller
(1948b) and Smith (1992a) suggest a Pleistocene history of al-
ternation between habitat-rich pluvials and habitat-poor, arid
interpluvials, which could have resulted in cyclic adaptation to
special habitats in wet times, followed by selection of morpho-
logically generalized phenotypes during times of reduced habi-
tat diversity. The result is introgressive destruction of some
(but not all) independent lineages.
We therefore define species as lineages whose separate his-
tory can be reasonably inferred from diagnostic morphological,
NUMBER 33 205
ecological, genetic, or behavioral apomorphies. This definition
recognizes that species are individuals (Ghiselin, 1974, 1997),
and it is meant to be consistent with Ghiselin's general defini-
tion (1997:99): "Biological species are populations within
which there is, but between which there is not, sufficient cohe-
sive capacity to preclude indefinite divergence." Our wording
acknowledges extrinsic properties and emphasizes history. It
might also be peculiarly ichthyological?D.S. Jordan (quoted
by Joseph Grinnell, 1924:225) said that "a species is not
merely a group of individuals distinguished from other groups
by definable features;" it is "a kind of animal or plant which
has run the gauntlet of the ages and has persisted." The intrin-
sic bases for maintaining historical lineage separateness may
be ecological (different habitats), reproductive (different times
or places of spawning), or developmental (different genes for
growth), but they are ultimately genetic and controlled by long-
term selection (Fisher and Ford, 1947; Wright, 1948). The
traits that we focus upon herein are associated with reduced hy-
brid fitness because of high mortality or low fertility (Dowling
et al., In prep.). Tests of intrinsic reproductive separation are
usually inferred from (1) the presence or absence of traits
unique to sympatric relatives, (2) apomorphic courtship dis-
plays, (3) assortative mating, (4) failure to produce fertile or vi-
able offspring, or (5) reduced fitness of hybrid products. Tests
of long-term persistence require paleontological or molecular
biogeographic data that provide a time context in which to
evaluate persistence or stability for a period longer than a post-
glacial fragment of a glacial cycle. We suggest that molecular
divergence measures can contribute to this test, although a
sharp, defining demarcation in time or sequence divergence is
not suggested. The criteria presented above provide appropriate
evidence for accepting or rejecting species assignment under
both the evolutionary and the biological species concepts.
These considerations provide a context for analyzing the tem-
poral scale of morphological and genetic differentiation in rela-
tion to speciation.
The above definition and the patterns of Great Basin fishes
through geologic time suggest avoidance of the chronospecies
concept, which is popular in paleontology but incompatible
with cladistic, evolutionary, or biological views of species
(Krishtalka and Stucky, 1985; Ghiselin, 1997). Great Basin
species are interpretable as individuals that may change
through time, sometimes anagenetically, and sometimes
through recombination of genes from related populations. In ei-
ther case, the individuality of ancestral-descendant lineages
through time is paramount (Ghiselin, 1974, 1997), and it seems
unhelpful to the study of speciation to change lineage names to
indicate anagenetic changes. To name chronological species
recognizes only stratigraphic diversity, not species diversity.
Similarly, it seems unhelpful to change the name of a sister
taxon that has given rise to a peripatric daughter taxon without
any change in the morphology of the original population (Bell,
1979).
SPECIES FORMATION IN THE GREAT BASIN.?The Great Ba-
sin fishes are a product of cycles of expanded and contracted
habitats, resulting in cycles of differentiation of local popula-
tions but only modest persistence of lineages through time.
Here the speciation process involves three steps (Smith et al.,
1995): (1) separation of populations by hydrographic or habitat
barriers (or microhabitat difference in lakes); (2) accumulation
of genetic differences in growth, development, physiology,
ecology, or behavior; and (3) accumulation of some minimum
combination of genetically based differences sufficient to cause
hybrids to suffer reduced fitness. Operationally, the products of
these process are recognized when lineage independence seems
reasonably supported by reciprocally nonoverlapping molecu-
lar, morphological, or ecological apomorphies that indicate an
intrinsic basis for lineage separateness. Confidence in and pre-
dictions from these hypotheses are subject to tests by paleonto-
logical or biogeographic circumstances that indicate persis-
tence through more than one environmental cycle.
The Cenozoic tectonic and Pleistocene hydrographic history
of the Great Basin created hundreds of allopatric isolates with
the potential to become species, as demonstrated by Hubbs and
Miller (1948b) and Hubbs et al. (1974). We conclude, however,
that the same conditions (small isolates) and processes (long
isolation and short cycles of habitat fluctuation) that caused the
initial differentiation also caused extinction of the majority of
potential species. Evidence of this extinction is seen in the de-
clining late Cenozoic fish diversity, depauperate local faunas,
and absence of the vast species numbers expected from scores
of so-called "species-pump" cycles.
Examples of surviving vicariant pairs in the Great Basin are
nevertheless notable. They include Lepidomeda albivallis and
L. altivelis in the White River Valley and in Pahranagat Valley,
Nevada; Richardsonius egregius and R. balteatus in the Lahon-
tan and the Bonneville-Columbia Basins; Catostomus tahoen-
sis and C. fumeiventris in the Lahontan and Owens Basins;
Cyprinodon salinus, C. milleri, C. nevadensis, and C. diabolis
in the Death Valley system; Crenichthys baileyi and C. nevadae
in the White River valley and in Railroad Valley; Empetrich-
thys latos and E. merriami in Pahrump Valley and Ash Mead-
ows; and Cottus extensus and C. echinatus in Bear Lake and
Utah Lake. Most of the other species in the Great Basin have
sister taxa outside the basin or have complex distribution pat-
terns requiring dispersal. Examples of stages of differentiation
similar to the above are found among the subspecies and races
of Siphateles bicolor, Rhinichthys osculus, Gila atraria, and
Oncorhynchus clarki. These forms are early and perhaps tem-
porary examples of vicariant speciation, as described above.
Our DNA divergence data and the hypothesis that speciation
usually involves the long-term accumulation of multiple ge-
netic differences suggest that, on average, species-level differ-
entiation (beyond the subspecies and races of Hubbs and
Miller) requires a longer time than the duration of the Pleis-
tocene.
206 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Lacustrine speciation requires special explanation. Examples
include the species of Prosopium in Bear Lake and the Bonne-
ville Basin, and Siphateles bicolor pectinifer in the Lahontan
Basin. The trophic and lacustrine adaptations of these sympat-
ric forms indicate that they might have differentiated in intrala-
custrine micro-allopatry (Smith and Todd, 1984) involving
evolution of different habitats and spawning times. Indepen-
dent evidence suggests that intralacustrine speciation is much
more rapid than speciation in most other circumstances (Mayr,
1984; McCune et al., 1984; Smith et al., 1995), possibly be-
cause it is driven by competitive selection (Rosenzweig, 1978).
Variability also might be enhanced by relaxed selection during
lake transgression and population expansion (McCune, 1990).
There is a widely cited hypothesis that lacustrine speciation
is driven by cyclic allopatry, with low stages that separate allo-
patric populations, which diverge, and high stages, which fos-
ter dispersal and widespread sympatry (e.g., Mayr, 1963; Mc-
Cune et al., 1984). The history of the Great Basin includes
numerous pluvial cycles (dozens, if pluvial cycles tracked gla-
cial cycles) during the last 3 My (Imbrie and Imbrie, 1979;
Benson et al., unpublished data; G.I. Smith et al., unpublished
data). The pattern of species abundance and distribution in
Great Basin fishes does not conform to the lacustrine fluctua-
tion model of lacustrine speciation (Mayr, 1984; Smith and
Todd, 1984) because extinction, not speciation, has dominated
Great Basin history (Smith, 1978), perhaps because the arid
parts of the cycle have been so extreme.
The fossil and molecular evidence for an ancient presence (a
criterion suggested by Ricklefs and Schluter, 1993) of most of
the species in the Great Basin also suggests that extinction has
been more frequent than species production. The abundance of
races and subspecies in the Great Basin suggests that frequent
origin and extinction jointly act on populations with limited
ranges. Species with stable habitats or large geographic ranges,
by contrast, tend to be long-term survivors (Jablonski, 1980).
The pattern of morphological change through time, if viewed
on the million-year time scale, would appear to be consistent
with the hypothesis that most morphological changes in species
occurred at the time of initial divergence (Stanley, 1975; Gould
and Eldredge, 1977). Viewed on the thousand-year time scale,
we see early morphological response to changing environ-
ments, but slow accumulation of characters that may be capa-
ble of sustaining long-term individuality.
TEMPORAL SCALE OF SPECIES DIFFERENTIATION IN THE
GREAT BASIN.?These cases suggest a model in which Great
Basin fish differentiation was distributed broadly across pluvial
cycles of the last 3 My. Through time, some fraction of the new
forms went extinct, and some species and subspecies intro-
gressed into relatives during extinction episodes. Some of these
diverged again during pluvial and early postpluvial episodes.
The net result is abundant geographic variation cycling through
time, as seen in the fossil record, combined with low species
diversity, as seen in Great Basin fishes today (Miller, 1959).
The fossil record indicates that throughout the million-year
time scale, net changes have been slow. Geographic variation
shows that during the thousand-year time scale, some morpho-
logical changes in oligogenic and ecophenotypic traits have
been rapid (Hubbs and Miller, 1948b; Bell, 1994). It is possible
that bony characters seen in the fossil record are multigenic and
evolve more slowly than oligogenic meristic, color, and shape
characters that differentiate subspecies. It is also likely that
body shape, scale, and fin characters fluctuate, perhaps ecophe-
notypically, which is consistent with the random-walk hypothe-
sis (Bookstein et al., 1978). Fluctuating variations observed
during a time scale of hundreds of years (Bell and Haglund,
1982; Bell et al., 1985; Bell, 1994) and decades support this
hypothesis (Figure 12). Small, rapid changes are expected in
short-term studies, in contrast to slow changes expected in
long-term studies (Gingerich, 1986), in part because genetic
variation is consumed by selection (Lewontin, 1974).
These differences in amounts of change as a function of time
are observed in empirical studies in the laboratory as well as in
comparisons of paleontological rates (Gingerich, 1986). Ran-
dom fluctuation was not one of the hypotheses considered by
early students of Great Basin fishes: Hubbs (1940:211) saw
"orderly adjustment under the control of environment" not
"aimless wandering of genes through the organic world." Stud-
ies of morphological fluctuations in the fossil record (Figure
12; Bookstein et al., 1978; Chernoff, 1981; Bookstein, 1986;
Bell, 1994) and complex biogeographic patterns in the Great
Basin suggest a model of rapid adaptive response to changing
environments but minimal long-term trends because of the high
amplitude and periodicity of the fluctuations.
Conclusions
The Great Basin fish fauna is the most depauperate ichthyo-
fauna in North America, despite being situated on a landscape
ideal for allopatric speciation. Comparison of morphological
subspecies differences (Hubbs and Miller, 1948b; Hubbs et al.,
1974) with paleontological studies of populations through time
indicates that abundant, newly differentiated forms usually fail
to become long-term members of Great Basin biodiversity.
This fauna has not significantly diversified despite repeated
Pleistocene fluctuations in lake levels. The absence of the di-
versity predicted by repeated environmental change leads to
two conclusions: (1) extinction of small populations is a more
potent long-term force than speciation in fluctuating environ-
ments, such as occurred in the Pleistocene Great Basin; and (2)
speciation processes are favored in small populations (Wright,
1940a; Mayr, 1963; Gould and Eldredge, 1977), but during a
time scale of thousands of years, most of this diversity be-
comes extinct if individuals in small populations cannot dis-
perse among drainages.
A related aspect of these fishes is that western salmonids, ca-
tostomids, cyprinids, and cyprinodontids are unexpectedly
lacking in intrinsic reproductive isolation among morphologi-
NUMBER 33 207
cally and ecologically differentiated lineages, some of which
have separate evolutionary histories going back several mil-
lions of years (Smith, 1992a). Fluctuations in abundance have
caused frequent hybridization (Hubbs and Miller, 1943; Hubbs,
1955). The frequent exchange of genes through introgression
appears to have maintained some genetic compatibility, even as
it provided variation for adaptations to environmental change.
The extreme isolated allopatry and resulting divergence of
Great Basin fishes permit some calibration of the chronology of
hydrologic history of basins, as sought by Hubbs and Miller.
Molecular divergence estimates, in conjunction with ages of
fossil ancestors, enable the calculation of DNA divergence-rate
estimates in salmonids (0.5% per 1 My) and in cyprinids and
cyprinodontids (1% per 1 My). These results permit percent se-
quence divergence to be used to calculate a rough chronology
of hydrologic barriers between individual basins, groups of ba-
sins, and drainages neighboring the Great Basin. The Rocky
Mountain barrier between Great Basin and eastern North Amer-
ican fishes was crossed at low elevations multiple times be-
tween 22 and 12 Ma, according to cyprinid sequence diver-
gences. Major connections among Great Basin, Colorado River,
Rio Grande, Rio Yaqui, and adjacent drainages in the Basin and
Range Province of northern Mexico date from 9 to 5 Ma, based
upon cyprinid divergences. Great Basin and Columbia-Snake
divergences of cyprinids imply that separations occurred early,
11-6 Ma, but that many reconnections occurred during the in-
terim. Two levels of differences between lowland and upland
taxa of the Lahontan and Bonneville Basins imply that major
confluences existed prior to 12-6 Ma and minor headwater con-
nections occurred 4-0.2 Ma. Estimates based upon trout DNA
sequence divergences imply that the Lahontan Basin headwa-
ters were connected to the Columbia River drainage about 0.2
Ma, but Lahontan Basin populations were connected to each
other more recently than 0.2 Ma. The Bear River populations
were connected to the Snake River populations 0.3 Ma, but the
trouts of the Bonneville and Lahontan Basins were connected
4-2 Ma. Cyprinodontoids of the Death Valley system were con-
nected to the lower Colorado River basin and to Mexico several
times in the last 2-3 My and were among basins in the Death
Valley system during the last 0.1-1.3 My. In summary, paleon-
tological and molecular evidence indicates that most Great Ba-
sin fish lineages are older than previously thought.
Appendix
The fish species and subspecies of the hydrographic Great
Basin, both fossil and Holocene, are listed below in approxi-
mate phylogenetic sequence and are numbered to correspond to
references to them in the text, tables, and figures. Forty-one
forms known only from fossils are indicated by a dagger (|).
An asterisk (*) indicates that the account of the Holocene lin-
eage contains a pertinent fossil record.
CYPRINIDAE
Western Chub Group
Mylopharodon-Lavinia-Hesperoleucus Subgroup
1. ^Lavinia sp. (Secret Valley hitch). This small species is
diagnosed by pharyngeal teeth with corrugated grinding sur-
faces and by dentaries with flared lateral edges (Smith, 1975,
fig. 19B). Idadon of the Snake River Plain is herein interpreted
as a synonym of Lavinia of the Great Valley ("1" in Table 2).
Idadon and Lavinia share the unique features of corrugated
pharyngeal teeth and flared dentaries, although these features
are more extremely developed in Idadon. Flared dentaries col-
lected from elsewhere on the Madeline Plains, northeastern
California, have been identified as Lavinia and Acrocheilus by
Wagner et al. (1997). The Secret Valley dentaries are more
flared than those from the Snake River Plain. Although highly
variable, the Secret Valley dentaries are more similar to Acro-
cheilus than to Lavinia (Wagner et al., 1997), but they lack the
abrupt angles of Acrocheilus. {Acrocheilus teeth are not present
among UMMZ specimens from Secret Valley.) The Secret Val-
ley fossils are assigned an age of ~3.5 Ma; this assignment was
based on their correlation with specimens from the Snake River
Plain.
Lavinia + Hesperoleucus is the sister group to Mylopharodon
in the cladogram (Figure 2). Lavinia and Mylopharodon are
known from apparent first appearances on the Snake River Plain
about 6 Ma and 4 Ma, respectively; these ages correlate with the
sequence divergence value of 8.1% (Tables 2, 5) in Figure 14.
Molecular evidence (Table 2) indicates a sequence diver-
gence of 15% between Siphateles bicolor and Lavinia. (See
number 14, below.)
2. ^Lavinia sp. (Carson Valley hitch). A partial left den-
tary from late Blancan sediments (~3 Ma) and three caudal ver-
tebrae from early Blancan sediments (~5 Ma) were collected in
Carson Valley by T. Kelly. The dentary is 10.5 mm long and
represents a fish about 18 cm in standard length (SL). The
specimen is diagnosed as Lavinia by the flared gnathal ramus,
coronoid process at midpoint of the bone, mesially curved ante-
rior end, and position of eight lateralis pores. The caudal verte-
brae are dorsoventrally compressed, with a thick, single longi-
tudinal ridge on the lateral surface of the centrum.
3. ^Lavinia sp. (Cache Valley hitch). Specimens collected
from Cache Valley near Logan, Utah, by J. Stewart Williams of
Utah State University, appear to be large Lavinia. This conclu-
sion is based upon the slightly corrugated grinding surfaces on
208 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
the pharyngeal teeth, the flared dentaries, and the high fin-ray
counts. The distinct morphology and the absence of Pliocene
indicators with these fossils suggest a pre-Blancan, late Mio-
cene age (~6 Ma). McClellan (1977) reported Idadon (herein
treated as a synonym of Lavinia) from the Junction Hills fauna
of the Salt Lake group, near Fielding, Box Elder County, Utah.
The fossils appear to be Hemphillian in age by correlation with
adjacent faunas. Cyprinids similar to Lavinia or Mylopharo-
don, but not to Gila, have been collected from the Bear River
valley near Georgetown, Idaho, by Sue Ann Bilby and have
been deposited in the collections of the Idaho State Museum of
Natural History. This record supports the age estimate of 6 Ma
assigned to the sequence divergence value of 8.1% in Figure
14.
4. Hesperoleucus symmetricus (Baird and Girard) (Califor-
nia roach). This species lives in the Great Basin only in the
Warner Lakes system, near Adel, Lake County, Oregon. It is
most similar to Hesperoleucus sp. in Goose Lake and the Pit
River drainage (from where it might have been introduced
(Miller, 1959)). Although sometimes placed in Lavinia because
of protein and molecular similarities, the morphological differ-
ences and fossil record suggest that Lavinia and Hesperoleucus
are long-different lineages that share molecular similarity (they
differ by only 1.3% sequence divergence) probably because of
introgression.
5. *Mylopharodon Ayres (western chubs). Chubs with
molariform teeth occurred in the Great Basin and in the Snake
River Plain from the middle Miocene through the Pliocene, and
they remain in the Great Valley of California. Fossils of this ge-
nus are rare in the Great Basin. |Mylopharodon{l) doliolus
Smith and Kimmel was a large chub with widely spaced,
rounded molariform teeth (Smith et al., 1982, fig. 4E). It was
placed in Mylopharodon because of its molariform teeth, elon-
gate posterodorsal process of the pharyngeal arch, and general
shape of the pharyngeal arch. This form is from the Esmeralda
Formation in Silver Peak Quadrangle, Nevada, and may be as
old as 12 + 2 My. It lacks the derived characters of all My-
locheilus and Mylopharodon (of the Snake River Plain, Great
Basin, Sacramento Basin, northeastern California, and south-
eastern Oregon) and, therefore, could represent the stem group
leading to Mylopharodon and Siphateles. Mylopharodon lived
in the Honey Lake basin (collected by Todd Yeoman of Sierra
College) and the Snake River Plain during the Pliocene. ^Mylo-
pharodon hagermanensis Uyeno first appeared at Hagerman,
Idaho, on the Snake River Plain about 4.5 Ma. Mylopharodon
conocephalus (Baird and Girard) of the Great Valley (number 5
in Table 2) and Siphateles (of the Lahontan Basin) differ by
10.2% sequence divergence (Figure 2; Table 2). The Miocene
fossil could also be the earliest western relative of eastern cyp-
rinids, from which Mylopharodon conocephalus differs by
42% sequence divergence (Table 2; Dowling, unpublished
data).
Siphateles Subgroup
6. ^Siphateles sp. (chub). Chubs with a single row of pha-
ryngeal teeth on each arch are widespread in the late Miocene
to Holocene fossil record of Nevada. Whole fossils and numer-
ous separate bones similar to Siphateles were discovered in
Barstovian (middle Miocene, -15 Ma) beds of Buffalo Canyon,
Churchill County, Nevada, by Ron and Zachary Jones. These
are the oldest known cyprinid bones in the Great Basin, so they
provide an estimate (underestimate) of the time of vicariance
with the eastern clade of North American shiners, from which
Siphateles differs by more than 30% sequence divergence (Ta-
ble 2; Dowling, unpublished data). A complete but undescribed
specimen similar to Siphateles is known from the Nevada Test
Site in southern Nye County, Nevada. It is believed to be Mio-
cene in age. Several complete specimens and numerous bones
are known from the late Miocene beds of the Virgin Valley For-
mation, Humboldt County, Nevada. Unidentified cyprinid fos-
sils are known from Pine Nut Ridge (Douglas County, late
Hemphillian to Blancan) and from the Coal Valley Formation.
The late Hemphillian fossils are estimated to be 6 My old and
may indicate the branch point with Gila, which differs from
Siphateles by an average of 11.5% (s.d. = 2.2%) sequence di-
vergence (Figure 14; Table 2).
7. f Siphateles breviarchus (Cope) (Fort Rock chub). This
is the common cyprinid in the Pliocene sediments of Fossil
Lake, Lake County, Oregon. It has robust pharyngeals that are
similar to those of Mylopharodon. It represents a stage of evo-
lution prior to diversification of the Siphateles bicolor com-
plex. The date of its divergence from Mylopharodon is prob-
ably about 10 Ma based on sequence divergence (Figure 14;
Table 2).
8. tSiphateles traini (Lugaski) (chub). This fossil form,
from the late Miocene lacustrine beds in Jersey Valley, Persh-
ing County, Nevada (Lugaski, 1979), is morphologically
somewhat similar to S. bicolor. The zeolite beds in which S.
traini occurs suggest that the fish lived in a shallow, saline, al-
kaline lake of pH 9, much like Pyramid Lake, Washoe
County, Nevada, in which S. b. obesus currently lives
(Lugaski, 1979). Michael Bell and others have collected a
large cyprinid from the late Miocene part of the Truckee For-
mation, near Hazen, Nevada (see number 29, below), with ro-
bust proportions and fin-ray counts that suggest a relationship
to this group. These late Miocene sites are estimated to be 6
My old (Figure 14).
9. ^Siphateles bicolor (Girard) (tui chub). This polytypic
species is widespread in the Lahontan Basin drainage as S. bi-
color obesus, and it also extends into Oregon's lake basins as S.
b. bicolor, and into the Harney Lake and lower Columbia River
basins, where it has been referred to S. b. columbianus. Siphat-
eles bicolor is diagnosed by a single row of pharyngeal teeth.
Cyprinid bones found in indurated sediments of Pyramid and
Winnemucca Lakes and other areas of the Lahontan Basin are
probably Gila bicolor obesus. Fossils related to this species are
NUMBER 33 209
known from the Pliocene of Honey Lake, Lassen County, Cali-
fornia, and Mopung Hills, Churchill County, Nevada, and from
numerous Holocene Indian middens reported in the archaeo-
logical literature. The Blancan fossils occurred before Pleis-
tocene tectonism began at the Honey Lake site, possibly 3 Ma.
Samples from the Lahontan Basin, from the Owens River,
Mono Lake, Mohave River, and Oregon Lakes basins, and
from the Columbia River, analyzed by Dowling, differ from
each other by 3.1% of the cytochrome b sequence (Table 2).
Siphateles bicolor obesus is widespread throughout the
streams and springs in the Lahontan Basin system, where it
typically has 10-19 short gill rakers, a 5-4 pharyngeal tooth
formula (Hubbs et al., 1974), and distinctive scales. Many un-
named, geographically distinct forms, in addition to the named
taxa summarized below, vary in gill raker number or morphol-
ogy, scale counts, size, proportions, nuptial characters, and
color patterns. For example, Sulfur Spring, in Diamond Valley,
Eureka County, Nevada, contains a dwarf form with long gill
rakers, complete lateral line, long pectoral fins, and distinctive
colors; it is included with S. b. obesus because a downstream
population has features that are intermediate between it and the
Humboldt drainage forms (Hubbs et al., 1974).
Hubbs and Miller (1948b) reported distinct subspecies in the
valleys of Lake Dixie and Lake Toiyabe, in Big Smokey Val-
ley, and in Railroad Valley in central Nevada, south of the
Lahontan Basin drainage (Figure 4). Another distinct form
lived on the east side of the valley of pluvial Lake Toiyabe in
Lander and Nye Counties, Nevada, along with Rhinichthys os-
culus lariversi (Lugaski, 1972). (Forms on the west side of the
valley were apparently transplanted into the springs around the
turn of the century; both populations of S. bicolor obesus are
now extinct.)
Just about every isolated population of S. bicolor in the Ore-
gon Lakes has a distinctive combination of characters involv-
ing size, body shape, fin shape, gill raker length and number,
and color (Hubbs and Miller, 1948b; Hubbs et al., 1974).
Siphateles bicolor eurysoma (Williams and Bond, 1981) was
named from the Guano Basin, Nevada. The Catlow Valley
chub in Malheur County and the Guano Valley chub in south-
eastern Lake County, Oregon, are recognized as Siphateles bi-
color eurysoma Williams and Bond (1981) (Hubbs and Miller,
1948b). Siphateles bicolor in the Oregon basins vary from
forms with more than 16 gill rakers (Catlow Valley, Guano
Lake, and Warner Lakes basins) to forms with fewer than 16
gill rakers (basins of the lower Columbia River, of Harney,
Klamath, Fort Rock (Fossil), Alkali, and Summer Lakes, and
of Lakes Abert and Chewaucan) (Minckley et al., 1986).
10. ^Siphateles sp. (chub). Pliocene(?) fossil Siphateles
from Duck Valley, Washoe County, Nevada, is estimated herein
to date from 2 Ma. Specimens like the modern Siphateles of the
Lahontan Basin occur in Humboldt Canyon below the Lava
Creek tephra, which was deposited 665 ? 10 Ka (Izet, 1981;
Izet et al., 1992) and the Brunhes-Matuyama geomagnetic re-
versal, 780 Ka (Baksi et al., 1992). Oregon and Nevada speci-
mens differ by 3.1% sequence divergence (Table 2). G.I. Smith
and colleagues of the United States Geological Survey (USGS)
recovered a specimen from the Searles Lake core (R.R. Miller,
pers. comm. to G.R.S., 1999).
11. jSiphateles bicolor subsp. (chub). Gobalet and Negrini
(1992) reported a distinct but undescribed fossil form of S. bi-
color from the Chewaucan Basin of southern Oregon; the form
was stated to be more than 10 Ky old. Modern specimens of
Siphateles from Oregon and Nevada differ from each other by
3.1% sequence divergence (Table 2). Modern representatives in
the Chewaucan, Alkali, and Fort Rock Basins (Figure 4) are
similar to each other (Hubbs and Miller, 1948b).
12. * Siphateles mohavensis Snyder (Mohave tui
chub). This distinct chub was originally described as a spe-
cies from the Mohave River in the valley of pluvial Lake Mo-
have. It has eight anal-fin rays, uniserial teeth, and 18-29 gill
rakers. Its numerous gill rakers characterize a plesiomorphic
condition retained from its lacustrine ancestor (Hubbs and
Miller, 1948b). Otherwise, it is similar to tui chubs (Miller,
1973). Fossils related to this form were reported from sedi-
ments of China, Manix, and Searles Lakes by Buwalda (1914),
Blackwelder and Ellsworth (1936), and Flint and Gale (1958).
Numerous isolated bones collected by D.F. Hewett and D.W.
Taylor from Lake Manix beds represent this species and have
osteological characters like the Siphateles figured by Miller
(1973). Teeth and vertebrae recovered by B. Leatham from
"Lake Mohave" (also called Soda Lake) north of Baker, Cali-
fornia, dated at 20 Ky B.P., are most likely from this species.
Late Pliocene tui chub bones associated with Catostomus and
Chasmistes from the White Hills near Airport Lake, Inyo
County, California, may be the same as the Soda Lake form.
13. *Siphateles bicolor snyderi (Miller) (Owens tui chub).
This chub inhabits springs and creeks of the Owens River
drainage, Inyo and Mono Counties, California. Small lateral
line scales, distinctive lateral radii on the scales, few dorsal-
and anal-fin rays, a distinctively shaped pharyngeal arch, and
usually 10-14 gill rakers diagnose this form. It is most similar
to Siphateles bicolor obesus in the East Walker River, from
which Miller (1973) presumed it was derived. Fragments as-
sumed to represent the Owens River Siphateles were recovered
throughout the Pleistocene Owens Lake core (Firby et al.,
1997). Holocene Indian middens along the nearby Alabama
Hills in Inyo County, California, contain abundant remains of
this tui chub and Catostomus fumeiventris, but no other fishes.
14. ^Siphateles sp., (Mono Basin chub). Pliocene fossil
bones from Mono Lake basin and Miocene to Pleistocene sedi-
ments of Carson Valley are clearly plesiomorphic Siphateles.
The dentaries from both sites are similar to each other and are
sufficiently distinct from other Siphateles to be regarded as a
single, different species (see number 15, below; to be described
by K. Gobalet). The Mono Basin sample of Siphateles now in-
cludes hundreds of identifiable specimens. They are similar to
Gila crassicauda but are different from other Gila and other
Siphateles in the distinct robustness of the anterior ramus of the
210 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
dentary and in the anterior position of the dorsal mental fora-
men; the anterior edge of the coronoid process is angled more
posteriorly in Gila. The masticatory platform of the basioccipi-
tal is similar to Mylopharodon in the obtuse angle of the poste-
rolateral processes at the origin of the caudal process; this part
of the bone matches Siphateles rather than Mylopharodon. The
pharyngeal arches bear a single row of teeth, which are robust,
somewhat rounded, and reach large size (to 9 mm). These teeth
indicate the fossil species had a larger maximum size than the
modern record of 42 cm in total length for this species (Sigler
and Sigler, 1987). We estimate the age of these fossils to be
about 8 My old (based upon the morphological differences de-
scribed here), which places them closer to the branch point
with Mylopharodon than with Gila. A large, straight dentary
and molariform pharyngeal teeth attributed to Mylopharodon
by Miller and Smith (1981) might be large Siphateles with long
jaws and rounded teeth.
15. ^Siphateles sp. (Carson Valley tui chub). Abundant
Carson Valley materials are known from the early Blancan
Buckeye Creek local fauna, the late Blancan Fish Spring Flat
local fauna, and the early Irvingtonian Topaz Lake local fauna,
Douglas County, Nevada. They are distinctive in their robust
and posteriorly angular pharyngeals, somewhat like the Mono
Basin forms. The dentary shows a more anterior position of the
dorsolateral mental foramen than other forms, except those
from Mono Basin. The masticatory plates of the basioccipitals
are rounded ventroposteriorly, like fossil Siphateles from the
Mono Basin and Humboldt River canyon and from Holocene
forms from some Lahontan Basin localities and the Sacramento
River drainage. The Carson Valley forms are estimated to have
lived 5, 3, and 1 Ma.
16. ^Siphateles bicolor pectinifer (Snyder) (lake min-
now). This lacustrine tui chub was described by Snyder
(1917a) and placed in the distinct genus Leucidius. It can be dis-
tinguished by its long, numerous gill rakers (29-36) and by its
usually 5-5 pharyngeal-teeth pattern. It is sympatric with S.
obesus (Snyder, 1917a; Hubbs et al., 1974:143-150). Extensive
introgression has eliminated recognizable individuals of this
form in Eagle Lake and has reduced its distinctiveness in Lake
Tahoe, Walker Lake, and Pyramid Lake. The presence of hybrid
pharyngeals in ancient Indian middens (Hubbs, 1961; Smith,
1985) is evidence that the parental forms hybridized but main-
tained their differences for thousands of years (Hubbs, 1961).
Remnant populations of S. b. pectinifer in different, uncon-
nected habitats suggests that the ecologically significant differ-
ences are genetically based (Hubbs, 1961; Hubbs et al., 1974).
17. Siphateles bicolor vaccaceps (Bills and Bond) (Cowhead
Lake tui chub). This unique population was discovered by
Hubbs and Miller (1948b) in Cowhead Lake, Modoc County,
California. It is characterized by its small size (less than 12 cm
SL), numerous, short gill rakers (19-25), and small scales. It is
most similar to the unnamed chub in Warner Valley (to which
Cowhead Lake was tributary in pluvial times), but that form
has longer gill rakers (Bills and Bond, 1980).
18. Siphateles bicolor newarkensis (Hubbs and Miller)
(Newark tui chub). This chub is the only native fish in the
springs of the main basin of pluvial Lake Newark, in White
Pine County, Nevada. It is characterized by having few, short,
thick gill rakers, a usually complete supratemporal canal, long
pectoral fins, and small size (less than 10 cm SL) (Hubbs and
Miller, 1972; Hubbs etal., 1974).
19. Siphateles bicolor euchilus (Hubbs and Miller) (big-
mouth tui chub). This is the sister form to S. b. newarkensis,
with which it shares many features. It is found in springs of
Fish Creek (Little Smokey) valley in the southwestern part of
the valley of pluvial Lake Newark (southeast corner of Eureka
County, Nevada). It is characterized by relatively large size
(10-15 cm SL), especially anteriorly; large fins, especially the
pectorals; and large scales (Hubbs and Miller, 1972; Hubbs et
al., 1974).
20. Siphateles bicolor isolatus (Hubbs and Miller) (Clover tui
chub). This represents one of two species found in Indepen-
dence Valley, Elko County, Nevada, in the valley of pluvial
Lake Clover. Siphateles b. isolata is small, with intermediate-
length gill rakers reduced to about 11. It shares characteristics
with S. b. newarkensis and S. obesus of nearby Bishop Creek of
the Lahontan Basin drainage (Hubbs and Miller, 1972; Hubbs
etal., 1974).
21. Siphateles alvordensis (Hubbs and Miller) (Alvord
chub). This species inhabits creeks in the Alvord Basin of
Malheur County, Oregon, and adjacent Humboldt County, Ne-
vada. It is characterized by uniserial teeth, small, embedded
scales with radii in all fields, usually seven dorsal and anal fin
rays, and numerous, short gill rakers (16-22) (Hubbs and
Miller, 1972).
22. Siphateles boraxobius (Williams and Bond) (Borax Lake
chub). This small fish occurs in Borax Lake, Alvord Basin,
Malheur County, Oregon. The diagnosis includes a large head,
large eyes, small caudal peduncle, and a reduced lateral line.
The species shares general characters with S. alvordensis, but it
has fewer gill rakers, fewer lateral line scales, and fewer pecto-
ral-fin rays, like many dwarf populations of Siphateles (Will-
iams and Bond, 1980).
Gila-Acrocheilus-Klamathella Subgroup
23. *Gila atraria (Girard) (Utah chub). This variable,
abundant species is widespread in the Bonneville Basin and up-
per Snake River above Shoshone Falls. It is distinctive in its
2,5-4,2 pharyngeal teeth with grinding surfaces, 9 dorsal-fin
rays, 8 anal-fin rays, and 8-16 gill rakers. It is locally divergent
and polymorphic in its mtDNA. The Bear River population di-
verged from the Bonneville Basin populations more than 2 Ma,
well before the complex Late Quaternary interactions between
the Bear River and Lake Bonneville (Bouchard et al., 1998).
The fossil record of Gila atraria is known back only to the late
Pleistocene Black Rock Canyon, Smith Canyon, and Home-
stead Cave faunas in the Lake Bonneville drainage basin, Utah
NUMBER 33 211
and Nevada (Smith et al., 1968; Mead et al., 1982; Broughton,
2000a, 2000b).
24. Gila jordani Tanner (Pahranagat roundtail chub). This
species lives in Ash Springs, Crystal Lake, and (formerly) Hiko
Springs, all in Pahranagat Valley, pluvial White River system,
Lincoln County, Nevada (Tanner, 1950). It has the unique com-
bination of Gila robusta morphology and G. cypha and G.
atraria mtDNA (Table 2); this unique combination is diagnos-
tic for the species. Gila jordani is sympatric with Lepidomeda
altivelis, Rhinichthys osculus, Catostomus clarki, and Crenich-
thys baileyi grandis (LaRivers, 1962). It has been regarded as a
subspecies of Gila robusta, with which it was presumably con-
nected via the pluvial White River during the last 10 Ky
(Hubbs and Miller, 1948b). The principle of individuality of
distinct forms of hybrid origin, as applied to Gila seminuda
(DeMarais et al., 1992), suggests that although Gila jordani has
been separated for only 35 Ky, it should be recognized as a spe-
cies with a unique history because of its unique combination of
apomorphies.
Gila (or Moapa) coriacea, the Moapa dace, is a localized rel-
ative of Gila robusta (Figure 2) and G. intermedia that inhabits
warm springs in the Moapa River, which is the former lower
course of the pluvial White River channel in northern Clark
County, Nevada. It is not a current member of the Great Basin
fauna, but it is geographically peripheral and is a relative of
Gila jordani. Moapa was named by Hubbs and Miller (1948b)
who considered it to be related to Agosia, Rhinichthys, and Gila.
According to the molecular cladogram (Figure 2), it is a mem-
ber of the Gila clade. A Pliocene or Pleistocene fossil opercle of
the Gila group was collected by G.R. Smith from sediments ex-
posed in a roadcut near the town of Moapa, Nevada.
25. *Klamathella sp. (chub). Klamathella fossils are found
in Pliocene sediments at Honey Lake, Madeline Plains, and Al-
turus, California. The fossils are diagnosed by their Gz7a-like
pharyngeal bones with two rows of pharyngeal teeth. Fossils
from the Salt Lake group (~6 Ma) at Junction Hills, Box Elder
County, Utah, were reported by McClellan (1977); they are
similar to fossil Klamathella from the Snake River Plain. Kla-
mathella was formerly a subgenus of Gila, which was erected
for G. coerulea. We hereby raise it to generic rank on the basis
of molecular evidence that Klamathella and Acrocheilus are
sister groups (Figure 2). Separation of Klamathella from Gila
makes Gila monophyletic.
26. *Acrocheilus alutaceus Agassiz and Pickering
(chiselmouth). This is a Snake River and Columbia River
species that also lives in the Harney-Malheur Basin. The basin
was cut off from the Malheur River, a tributary to the Snake, by
rather young volcanics (less than 1 Ma, Luedke and Smith,
1982). McClellan (1977) found Acrocheilus teeth in sediments
of the Salt Lake group in the Junction Hills, Box Elder County,
Utah. Fossils reported as Acrocheilus from the Madeline Plains
in northeastern California (Wagner et al., 1997) may be highly
variable Lavinia.
In the molecular phylogeny (Figure 2), Acrocheilus is the
sister of Klamathella coerulea, and together they are sister to
the Gila clade. Acrocheilus and Klamathella differ by 6.4% in
their cytochrome b sequence (Table 2); fossils show that the
apomorphic jaw-shape of Acrocheilus was well established on
the Snake River Plain 7 Ma (Smith et al., 1982), indicating di-
vergence from Klamathella (plesiomorphic) occurred prior to
that date.
27. "fGila turneri (Eastman) (chub). This fossil form is
known from the Miocene Esmeralda Formation of south-cen-
tral Nevada. Specimens are also known from slightly older
(-12 Ma) lacustrine beds of Stewart Valley, Mineral County,
Nevada. It is characterized by a slender caudal peduncle and 10
anal-fin rays. Gila turneri was assigned to Richardsonius by
LaRivers (1962), but, based upon the large anal fin, it could be
related to Colorado River Gila, which differ from other Gila by
an average of about 6% sequence divergence.
28. f Gilail) esmeralda LaRivers (chub). This species from
the Miocene (-12 Ma) Esmeralda Formation may be related to
Siphateles. It has 8 anal-fin rays, the dorsal fin in advance of
the pelvics, and a deep caudal peduncle (Lugaski, 1977, 1979;
Smith and Miller, 1985). Gila and Siphateles differ by an aver-
age of about 11% sequence divergence (Table 2).
29. f Gila{l) sp. (chub). A large chub, greater than 40 cm
long, with 8 anal-fin rays, 9 pelvic-fin rays, long dentaries, and
blunt, slightly hooked pharyngeal teeth with no grinding sur-
face was collected by Michael Bell from the Truckee Forma-
tion southeast of Ferndale, Nevada, in the Hemphillian Quarry
(N81) along with numerous sticklebacks (~6 Ma). A similar
fish was collected nearby by J. Platt Bradbury from the Eagle-
Picher Mine in the Clarendonian section (~12 Ma) of the
Truckee Formation (Axelrod, 1956). The Bell specimen has
dentaries like those of Gila or Siphateles, but it has extremely
short articular angulars, like those of Ptychocheilus or Mylo-
pharodon. These forms could be related to Siphateles from
Mono Lake and could be near the stem group for Gila, Pty-
chocheilus, Siphateles, and Mylopharodon, judging from the
shape of the angulars and the dentaries. Siphateles, Mylopharo-
don, Ptychocheilus, and most Gila differ from each other by an
average of about 10%?12% sequence divergence.
Ptychocheilus?Relictus?Eremichthys Subgroup
This group is not monophyletic in the analysis from which
Figure 2 was drawn, because Ptychocheilus is polyphyletic in
that tree. The possibility that Ptychocheilus is paraphyletic is
supported by the existence of fossils from 15 Ma bearing Pty-
chocheilus apomorphies (Smith et al., 2000), but the issue is in
need of further study.
30. * Ptychocheilus oregonensis (Richardson) (northern pike-
minnow). This Snake River-Columbia River species also
lives in the Harney-Malheur Basin. Ptychocheilus occurred in
the Pliocene Mono Basin, Alturus Basin, Madeline Plains, and
Secret Valley, California (Wagner et al., 1997). It is known from
Miocene (15 to ~6 Ma) and Pliocene (5.5 to 3 Ma) deposits in
212 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
the Columbia River and Snake River basins (Smith et al., 1982;
Smith et al., 2000). Numerous Miocene to Pliocene sites in Car-
son Valley, Douglas County, Nevada (Kelly, 1994), have
yielded Ptychocheilus dentaries, quadrates, vertebrae, a tooth,
and a neurocranium. During this study a single vertebra was
identified from Miocene deposits in Brady Park, Churchill
County, Nevada. McClellan (1977) identified long, slender, ca-
niniform pharyngeal teeth from the Miocene Salt Lake group in
the Junction Hills, Box Elder County, Utah, as Ptychocheilus.
Living Ptychocheilus oregonensis differ from their cladistic rel-
atives by an average of 11.2% (s.d. = 2.1%) sequence diver-
gence (Figure 14; Table 2). The oldest known fossil Pty-
chocheilus bears a well-constrained date of 15 Ma from the
Columbia Basin (Smith et al., 2000).
31. Relictus solitarius Hubbs and Miller (relict dace). This
small fish is endemic to Ruby and Butte Valleys (pluvial lakes
Franklin and Gale) and to Goshute and Steptoe Valleys (pluvial
lakes Waring and Steptoe); each of the four valleys compose
parts of both Elko and White Pine Counties, Nevada (Hubbs
and Miller, 1972). Relictus displays the western spring mor-
photype: small size (usually less than 10 cm SL), small
rounded fins, no barbels, reduced lateral line, and scales not in
orderly rows. It also has distinctive, cone-like nuptial tubercles
on the head and pectoral fins (Hubbs and Miller, 1972). Ac-
cording to mtDNA evidence, it may be cladistically, but dis-
tantly related to Eremichthys (sequence divergence = ~15%),
but its sequence is phenetically more similar to Gila robusta
(sequence divergence = -11%) (Table 2). The data indicate that
it is an ancient relict (Figure 2), as concluded by Hubbs and
Miller (1972).
32. Eremichthys acros Hubbs and Miller (desert dace). This
is a highly localized endemic of Soldier Meadows, in the north-
western part of the Lahontan Basin. It is diagnosed by its small
size (usually less than 6 cm SL), laterally broad pharyngeal
arches, uniserial teeth, and specialized sheaths on the jaws
(Hubbs and Miller, 1948b). According to parsimony cladistic
analysis of mtDNA data, its closest relatives are Relictus (se-
quence divergence = 15%) and Ptychocheilus (sequence
divergence = 14%), but it also shares 13% similarity with sev-
eral species of Gila. These data indicate that it is an ancient
relict, as suggested by Hubbs and Miller (1948b).
Snyderichthys-Lepidomeda Group
33. Lepidomeda albivallis Miller and Hubbs (White River
spinedace). This species was found in springs and their efflu-
ents in the upper White River valley of White Pine and Nye
Counties, Nevada (Hubbs and Miller, 1960), but it is now ex-
tinct. It differs from its sister species, Lepidomeda altivelis and
L. mollispinis, from elsewhere in the Virgin-White drainage,
by 0.5 % sequence divergence (Table 2). It lived with Rhinich-
thys osculus, Catostomus clarki, and Crenichthys baileyi.
34. Lepidomeda altivelis Miller and Hubbs (Pahranagat
spinedace). This localized species, now extinct, lived in the
Ash Spring and Pahranagat Lake area of the White River val-
ley, in Lincoln County, Nevada (Hubbs and Miller, 1960). It
differed from its sister species, Lepidomeda albivallis, by hav-
ing (on average) a more oblique mouth, a sharp snout, less dark
pigment, and a larger and more pointed dorsal fin with a stron-
ger spine.
35. Snyderichthys copei (Jordan and Gilbert) (leatherside
chub). This apparently paraphyletic species is endemic to the
Bonneville Basin, upper Snake River, and Little Wood River
drainages in Utah, Wyoming, and Idaho. Its generic affinities
have been ambiguous (Miller, 1945a), but Hubbs and Miller
(1960) suggested a relationship to Lepidomeda. Snyderichthys
shares mtDNA similarity at the level of 5% sequence diver-
gence with the Virgin River Lepidomeda, which provides an
important clue to its relationships and to western biogeography.
Snyderichthys and the Virgin River and White River Lepi-
domeda spp. differ from Lepidomeda vittata of the Little Colo-
rado River and from more plesiomorphic members of the pla-
gopterin clade (Simons and Mayden, 1997; Dowling et al.,
2002) by 8%-9% sequence divergence (Figure 2; Table 2).
This pattern of relative similarity in mtDNA indicates the rela-
tionships among the northern and southern members of this
species are complex (Dowling et al., 2002). The Bear River and
Snake River populations are more similar to Virgin River
drainage Lepidomeda (5% sequence divergence) than to Sevier
River and other Bonneville Basin Snyderichthys (11%-12% se-
quence divergence; Table 2), suggesting there has been inde-
pendent loss of spines or introgressive transfer of mtDNA in
Snyderichthys. This pattern of mtDNA similarity supports
other evidence for a former drainage connection between the
upper Snake River and the Virgin-White drainage of the
Mio-Pliocene Bouse Embayment (see "Discussion"). Addi-
tional evidence of this connection is the presence of the Muddy
Creek Formation from the Lake Mead area to the southwestern
Escalante Desert of the Bonneville Basin, prior to 4 Ma (Spen-
cer Reber, pers. comm. to G.R.S., 1999).
lotichthys-Richardsonius?Rhinichthys?Mylocheilus Group
36. Iotichthys phlegethontis (Cope) (least chub). This small
species is endemic to the Bonneville Basin. It was formerly
known from eastern tributaries to the Great Salt Lake south to
Snake Valley and Juab County, Utah. It shares a unique mor-
phological character, namely retarded closure of the posterior
cranial fontanel by the parietal, with Gila atraria (Coburn and
Cavender, 1992:346), but it shares its mtDNA haplotype with
Richardsonius (sequence divergence 9% with R. balteatus;
18% with G. atraria) (Figure 2; Table 2). Incongruent morpho-
logical and molecular characters suggest there has been intro-
gressive transfer of either morphology or mtDNA from either
Gila atraria or Richardsonius (respectively) to Iotichthys (Fig-
ure 2 versus Figure 3).
37. *Richardsonius balteatus balteatus (Richardson) (Co-
lumbia redside shiner). This subspecies is characterized by its
long anal fin. It lives in the Columbia River drainage and in the
Great Basin. In the latter, it inhabits the Silvies River of the
NUMBER 33 213
Harney-Malheur Basin (Bisson and Bond, 1971). We infer that
it dispersed north from the Columbia Basin postglacially be-
cause it is not in northern refugia. Although Richardsonius is a
solely western North American genus, it is closest to eastern
North American Clinostomus (sequence divergence = 12%;
Dowling, unpublished data). There is a fossil record of the R.
balteatus lineage from the Pliocene (~3.5 Ma) of the Snake
River Plain (Smith et al., 1982).
38. * Richardsonius balteatus hydrophlox (Cope) (Bonneville
redside shiner). This subspecies differs from R. b. balteatus
in having fewer anal-fin rays. It is found in the Bonneville Ba-
sin, in isolated creeks of the Harney-Malheur Basin, and above
barrier falls in the Snake River drainage and in the South Fork
of the John Day River in Oregon (Bisson and Bond, 1971). The
presence of relicts of this subspecies above various barrier falls
in the Columbia River drainage suggests that the plesiomorphic
morphology of this subspecies represents an earlier form than
its sister subspecies, R. b. balteatus. It presumably dispersed
after connection of the Snake River to the Columbia River
about 2 Ma, but before formation of the falls and other causes
of isolation in the late Pleistocene. A fossil right pharyngeal
arch of R. balteatus, presumably R. b. hydrophlox, was recov-
ered by G.R.S. from Wisconsinan deposits in Homestead Cave,
northern Bonneville Basin (Broughton, 2000a).
39. Richardsonius egregius (Girard) (Lahontan red-
side). This small shiner and its cohabitant in the Lahontan
Basin, Catostomus tahoensis, are the only widespread endemic
species of the Great Basin. Their absence from surrounding ba-
sins is (weak) evidence against low-elevation drainage connec-
tions between the Lahontan Basin and other basins in the last
several million years. Richardsonius egregius is abundant in di-
verse habitats, but it has not differentiated as much as Siphate-
les or Rhinichthys. It differs from its sister species, R. baltea-
tus, primarily in its smaller body size and smaller fins.
40. Rhinichthys cataractae (Valenciennes) (longnose
dace). This one of the most widely distributed fishes in North
America, especially in the eastern and northern parts of the
continent. Within the Great Basin, it is found in the northeast-
ern Bonneville Basin, specifically the Provo River, and north-
ward into Utah, Idaho, Wyoming, and the Harney-Malheur Ba-
sin of Oregon.
Rhinichthys {Apocope) osculus (Girard) (speckled
dace). This is a ubiquitous, polytypic species in the Great Ba-
sin and surrounding waters; its subspecies are summarized in
the entries to follow. Geographic and ecophenotypic variation
is extreme in Great Basin representatives of this species
(Hubbs et al., 1974). The general pattern of variation usually
trends toward one of two ecological morphotypes. The lotic
morphotype has a well-developed frenum and barbels, 2,4-4,2
pharyngeal teeth, small scales, medium-sized fins, and a fusi-
form body shape that is adapted to rheophilic life. The lentic
morphotype is a reductively evolved form; it has a reduced fre-
num, barbels, and lateral line; 1,4-4,1 teeth; small, rounded
fins; small body size; and inhabits quiet, isolated waters
(Hubbs et al., 1974; Miller, 1984:15). The following subspecies
represent only part of the variation in the group.
41. Rhinichthys osculus carringtoni (Cope) (Bonneville
speckled dace). This subspecies name applies not only to
Bonneville Basin, upper Snake River, and most Harney-Mal-
heur Basin populations, but also to the lower Snake River and
some Columbia Basin populations. This morphotype is prob-
ably an early form of the genus. The features of the lotic mor-
photype, listed above, characterize R. o. carringtoni; however,
the Harney-Malheur Basin fish are smaller and have reduced
barbels, especially in Warm Springs (Bisson and Bond, 1971).
Hubbs and Miller (1948b) suggested that the Warner Basin
contains three subspecies ofR. osculus.
42. Rhinichthys osculus adobe (Jordan and Evermann)
(Sevier speckled dace). This dace of the southern Bonneville
Basin is distinctive in its large head, large, horizontal mouth,
flat snout, and many pigment differences (Hubbs et al., 1974).
43. Rhinichthys osculus velifer Gilbert (Pahranagat terete
dace). This dace inhabits the swift water outflow of Ash
Springs, in Pahranagat Valley, Lincoln County, Nevada. The
fish is slender, with barbels, a frenum, and large, falcate fins
(Miller, 1984:15,17-18).
44. Rhinichthys osculus robustus Rutter (western speckled
dace). This name was applied to the form of Rhinichthys os-
culus of the Sacramento River, Lahontan Basin, and some asso-
ciated drainages by Snyder (1917a) and Hubbs et al. (1974). It
is characterized by possession of barbels and usually 1,4-4,1
pharyngeal teeth. Prior to 3.2 Ma, the Sacramento and Lahon-
tan drainages were connected by a pass at the head of San
Joaquin Canyon in Mono and Madera Counties, California
(G.I. Smith, pers. comm. to G.R.S.). The form in Soldier
Meadows, Humboldt County, Nevada, shows possible evi-
dence of introgression?it usually has the normal tooth formula
of 1,4-4,1, but a unique specimen has 0,5-4,0, like sympatric
Eremichthys acros (Hubbs and Miller, 1948a: 17).
45. Rhinichthys osculus nevadensis (Gilbert) (Amargosa
speckled dace). This fish lives in Ash Meadows, Nye County,
Nevada. It is small, with barbels, large scales, an incomplete
lateral line, 1,4-4,1 teeth, and no frenum. Its closest relative
was Rhinichthys osculus deaconi (Las Vegas dace), which was
endemic to Las Vegas Creek and Las Vegas Springs, which
were tributary to the Colorado River (Hubbs and Miller, 1948a;
Miller, 1984). The Amargosa and Las Vegas drainages are con-
nected by the Indian Springs Basin, a large section of which
was captured by, and is now a tributary to, the Amargosa River
(Gilluly, 1929:682; Miller, 1946a). The similarity of the sub-
species R. o. deaconi and R. o. nevadensis indicates a conflu-
ence of fish habitat and gene flow occurred between the Las
Vegas and Indian Springs Valleys, probably in the pass occu-
pied by the old railroad bed and Highway 95 east of Indian
Springs.
46. Rhinichthys osculus deaconi Miller (Las Vegas speckled
dace). This subspecies (now extinct) was endemic to Las Ve-
gas Creek and Las Vegas Springs, which were tributary to the
214 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
Colorado River in pluvial times. It was characterized by a
uniquely rounded anal fin, large scales, small pectoral fins,
1,4-4,1 teeth, no barbel or frenum, and an incomplete lateral
line (Miller, 1984).
47. Rhinichthys osculus reliquus Hubbs and Miller (Grass
Valley speckled dace). This form lived in the valley of pluvial
Lake Gilbert, Lander County, Nevada, and is probably extinct.
It lacked barbels and had a much-reduced lateral line (Miller,
1984).
48. Rhinichthys osculus oligoporus Hubbs and Miller (Clover
Valley speckled dace). Clover Valley is in the valley of plu-
vial Lake Clover, Elko County, Nevada. Rhinichthys osculus
oligoporus lacks barbels and has a reduced lateral line (Miller,
1984).
49. Rhinichthys osculus lethoporus Hubbs and Miller (Inde-
pendence Valley speckled dace). This subspecies lives in In-
dependence Valley, a tributary to the valley of pluvial Lake
Clover, Elko County, Nevada. It was apparently derived from
R. o. robustus of the Lahontan Basin drainage (Hubbs et al.,
1974). It is small and compressed, lacks a frenum and barbels,
and has a reduced lateral line.
50. Rhinichthys osculus lariversi Lugaski (Big Smokey
speckled dace). Described as a new species by Lugaski
(1972), this form was regarded as a subspecies by Hubbs et al.
(1974) and Miller (1984). It lives in the valley of pluvial Lake
Toiyabe in Lander and Nye Counties, Nevada. This fish has
62-70 lateral scale rows, 1,4-4,1 pharyngeal teeth, 8 dorsal-fin
rays, 7 anal-fin rays, no frenum, and variable development of
barbels and lateral line. The Lahontan form ofR. o. robustus is
its closest relative; Hubbs and Miller (1948b) hypothesized that
R. o. lariversi gained access to the valley of Lake Toiyabe
through a stream connection with the valley of Lake Gilbert, on
its northern border. The taxon is unusual in that it coinhabits
the springs on the east side of the valley with a local form of
Siphateles bicolor obesus. In most areas of the western Great
Basin only a single form of Rhinichthys or Siphateles exists in
a given spring, or Siphateles lives in the spring and Rhinichthys
lives below the spring outflow.
51. Rhinichthys osculus moapae (Williams) (Moapa
dace). This subspecies was described from the Moapa River,
Clark Covmty, Nevada. It is a rheophile aivd has barbels, a fre-
num, falcate fins, and silver coloration (Williams, 1978).
52. Rhinichthys osculus subsp. (Meadow Valley Wash
dace). This inhabitant of springs in Lincoln County, Nevada,
is a large dace with small scales, a complete lateral line, a fre-
num, barbels, 1,4-4,1 pharyngeal teeth, and 37-39 vertebrae
(Miller, 1984).
53. Rhinichthys osculus subsp. (Pahranagat speckled
dace). The subspecies occurs in Lincoln and Clark Counties,
Nevada. It has rather large scales, small fins, an incomplete lat-
eral line, and reduced number and size of barbels, but it has
2,4-4,2 pharyngeal teeth (Miller, 1984).
54. Rhinichthys osculus subsp. (White River shortfin
dace). The dace that lives in quiet waters of the White River
valley in White Pine County, Nevada, is an undescribed sub-
species with robust body, small rounded fins, no barbels, no
frenum, and 2,4-4,2 pharyngeal teeth (Miller, 1984).
55. Rhinichthys sp. (Deep Creek dace). This endemic spe-
cies has been collected in the western Bonneville Basin, from
Fish Springs to Snake Valley. It differs from other robust Rhin-
ichthys osculus in color, body shape, meristics, and 2,4-4,2
pharyngeal teeth (R.R. Miller and G.R. Smith, unpublished
data).
56. ^Mylocheilus inflexus (Cope) (Chalk Hills chub). This
fossil form of Mylocheilus from Idaho and Nevada is similar to
Mylopharodon (Appendix, number 5) in the possession of mo-
lariform teeth. The two lineages may have hybridized in the
past, according to molecular and fossil evidence (Dowling and
Smith, unpublished data). Mitochondrial DNA data indicate
that Mylocheilus caurinus (number 58 in Table 2) belongs in
the shiner clade with Richardsonius (sequence
divergence = 33%, Table 2). Miocene(?) chub pharyngeals
from Brady Pocket (Churchill County, Nevada, collected by
Charles Dailey) are similar to specimens of M. inflexus from
the late Miocene (~7 Ma) Snake River Plain (Smith et al.,
1982, fig. 4A). This suggests a Miocene connection existed be-
tween central Nevada and the Snake River Plain.
57. jMylocheilus sp. (chub). A pharyngeal arch and teeth
of a Mylocheilus, intermediate between M. inflexus and M. ro-
bustus, were collected from the Salt Lake group at Junction
Hills, Box Elder County, Utah (McClellan, 1977). Correlation
with other Mylocheilus suggests that the Junction Hills fauna is
Hemphillian in age. The presence of this form and fossil
Ameiurus in the fauna agrees with molluscan evidence for a
Miocene connection and subsequent isolation between waters
in northern Utah and the western Snake River Plain (Taylor,
1966).
58. ^Mylocheilus robustus (Cope) (baby-jaw chub). This
species was described from late Miocene and Pliocene pharyn-
geal arches from the Snake River Plain; the arches are large
and have molar teeth. Similar arches with molariform teeth are
in beds dated from the late Clarendonian to Hemphillian in the
Virgin Valley Formation, Nevada. A single tooth fragment
from the Barstovian to Hemphillian age in Esmeralda Forma-
tion equivalent beds of Stewart Valley, Nevada, may represent
this species (but see Mylopharodon from the Esmeralda Forma-
tion). Mylocheilus robustus is replaced in the modern Snake
River and Columbia River drainages by Mylocheilus caurinus.
CATOSTOMIDAE
59. fDeltistes sp. (Fort Rock Lake sucker). Rare maxillae
and dentaries from Pliocene localities at Fossil Lake, Fort Rock
Basin, Oregon, have characters that separate these fish from the
abundant Chasmistes in the basin. The maxilla has a posteriorly
directed premaxillary process and a general shape similar to
Deltistes luxatus of Klamath Lakes basin. The ventral keel of
the maxilla is prominent anteriorly; the dorsal keel is weak,
NUMBER 33 215
with a long convex slope anteriorly. The dentaries are shorter
and more robust than those of the Chasmistes clade, with the
gnathal ramus deflected mesially and lacking a ventral concav-
ity. An almost complete Weberian apparatus, similar to that of
the Deltistes clade of the Snake River Plain, was collected by
Sue Ann Bilby from sediments of the Salt Lake group, near
Georgetown, Idaho.
60. * Chasmistes cujus Cope (cui-ui). This is a large plank-
tivore in Pyramid Lake that formerly occupied Lake Lahontan
and its postpluvial remnants. It has the most specialized mor-
phology among the Holocene species in the genus. Fossil den-
taries, vomer, pterotic, and vertebrae from sediments of the
Humboldt River canyon confirm the presence of Chasmistes in
the pre-Lahontan system in the middle Pleistocene. Pleistocene
fossils were collected from near Fallon, Nevada, by J.R. Al-
corn. A similar form is found in Pliocene sediments at Secret
Valley and Honey Lake, California (see below; Taylor and
Smith, 1981). Chasmistes cujus is known from widespread
Lahontan Basin Indian middens, such as those in Humboldt
Cave, Lovelock Cave, Fishbone Cave, and Hidden Cave, Ne-
vada, and in Karlo, California (Hubbs and Miller, 1948b:41;
Follett, 1963, 1967, 1980; Miller and Smith, 1981; Smith,
1985). These specimens were probably transported by Ameri-
can Indians from Pyramid and Walker Lakes.
61. ^Chasmistes liorus Jordan (June sucker). This species
is now endemic to Utah Lake, but it was probably a Pleistocene
inhabitant of Lake Bonneville. It is apparently not closely re-
lated to Chasmistes muriei, the form that survived in the upper
Snake River (Miller and Smith, 1981). Chasmistes that were
not diagnostic to species were collected from the Thatcher Ba-
sin near Preston in southeastern Idaho by R.C. Bright (1967).
Chasmistes fossils were collected by McClellan (1977) from
sediments of the Salt Lake group at Junction Hills, Box Elder
County, Utah.
Chasmistes muriei of Jackson Hole, Wyoming, was native to
the upper Snake River drainage, where one specimen was col-
lected in the 1930s, but it is now apparently extinct. The mor-
phology of C. muriei is similar to that of C. batrachops (ex-
tinct, Fossil Lake) or C. brevirostris (Klamath drainage).
Pliocene fossils from the Teewinot Formation, upper Snake
River drainage in Teton County, Wyoming, are similar to C.
muriei in the Snake River and to C. brevirostris in the Klamath
Basin (Miller and Smith, 1981).
62. ^Chasmistes batrachops Cope (Fossil Lake suck-
er). This sucker was described from Pliocene or Pleistocene
sediments of Fossil Lake, Fort Rock Basin, Oregon. Its dentary
and maxillary osteology is generalized. The name may be ap-
plicable to similarly generalized bones from the Pliocene
Glenns Ferry Formation of Owyhee County, Idaho (not Chas-
mistes spatulifer); Mono Basin, Mono County, California;
China Lake and White Hills, Inyo County, California; and
Duck Valley, Washoe County, Nevada (Miller and Smith,
1981). Chasmistes bones are variable in Fossil Lake, perhaps
because Pliocene and Pleistocene remains might be mixed as
lag clasts on deflation surfaces in the basin.
63. ^Chasmistes sp. indeterminant (Mono Basin lake
sucker). This lake sucker is found in Pliocene sediments of
the Mono Basin, California, with Ptychocheilus, Siphateles,
Catostomus, and Pliocene mollusks (Taylor, 1966). Skulls from
Mono Basin resemble skull bones of C. batrachops from Fossil
Lake, Oregon. Dentaries and maxillae from Mono Basin are
similar to those from Secret Valley, California. Dentaries are
similar to those of C. batrachops from Fossil Lake (UMMP
74488) in the medially flattened ridge, in the depth of the
groove between the gnathal ridges, and in the anterior recess
caudal to the symphysis and gnathal ridges. The position of the
median mental foramen and the end of the Meckelian groove
are similar to those of C batrachops from the Snake River
Plain (UMMP 72458). Maxillae are similar to those from Fos-
sil Lake (UMMP 774490) in the narrow, pointed dorsal keel.
64. ^Chasmistes sp. (Carson Valley lake sucker). An early
Blancan form from Carson Valley, Douglas County, Nevada, is
known from dentaries and a maxilla; late Blancan fossils in-
clude an opercle, dentaries, and a vomer; an Irvingtonian den-
tary extends the record into the Pleistocene. The axial ridge of
the vomer is broad and flat, as in Chasmistes and Xyrauchen,
but lacks the postrostral constriction of Catostomus. The den-
taries are more slender than those of Catostomus but are similar
to those of Chasmistes. The medial ridge posterior to the sym-
physis resembles that of C. batrachops, but it is less pro-
nounced than those of the Mono and Lahontan Basins and
those of C. spatulifer of the Snake River Plain.
65. ^Chasmistes sp. (Honey Lake sucker). This species was
collected from Pliocene sediments of "The Island" of Honey
Lake by Taylor and Smith (1981) and Charles Dailey. The
Honey Lake sucker is intermediate between C. cujus of the
Lahontan Basin system and C. spatulifer of the Snake River
Plain (Miller and Smith, 1981). It is found with Oncorhynchus
cf. "flacustris, Gila, Siphateles, and Catostomus. Similar bones
are known from Duck Valley, Washoe County, Nevada, and
from the Alturus Basin, the Madeline Plains, and Secret Valley
of northeastern California.
66. * Catostomus ardens Jordan and Gilbert (Utah suck-
er). This species is native to the Bonneville Basin drainage in
Utah and to the upper Snake River drainage in Utah, Idaho, and
Wyoming. It is an old lineage that is related to Catostomus
macrocheilus of the lower Columbia drainage and to Catosto-
mus latipinnis of the Colorado River drainage (Smith, 1992b;
Figure 5). During the drought in the early 1930s, Catostomus
ardens hybridized with Chasmistes liorus on spawning grounds
in the Provo River, Utah, resulting in introgression that is still
seen in Chasmistes liorus specimens (Miller and Smith, 1981).
Late Pleistocene fossils were reported from the Hot Springs lo-
cality, Salt Lake County, Utah (Smith et al., 1968); from
Homestead Cave, Tooele County, Utah (Broughton, 2000a),
and from the Old River Bed between northern Bonneville and
Sevier Basins, Millard County, Utah (Oviatt, 1984). A small,
216 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
nearly complete Weberian apparatus, similar to that of Catosto-
mus ardens, was collected from sediments of the Salt Lake
group near Georgetown, Idaho, by Sue Ann Bilby and has been
deposited in the Idaho State Museum of Natural History.
67. *Catostomus macrocheilus Girard (largescale suck-
er). This Holocene species was a lower Columbia River (Rin-
gold Formation) sucker in the Pliocene; fossil evidence indi-
cates that it spread to the Snake River basin in the late Pliocene
(~2.8 Ma), after capture of the Snake River at Hell's Canyon
(Smith et al., 2000). It spread north of the Columbia as the gla-
ciers receded. It persists in the Harney-Malheur Basin as a
relict of the former Snake River outlet through Crane Gap or
Malheur Gap. Fossil ages indicate that the colonization of the
Harney-Malheur Basin could have occurred after the species
immigrated south through Hell's Canyon and prior to (or even
after, until about 9 Ka) occlusion of the outlet to the South Fork
of the Malheur River by the Voltage lava flow (Piper et al.,
1939), 0.04-1 Ma (Gehr and Newman, 1978; Luedke and
Smith, 1982).
68. ^Catostomus cristatus Cope (sucker). This fossil spe-
cies is best known from the Pliocene of the Snake River Plain.
A complete maxilla was collected by Charles Dailey in de-
formed Pliocene sediments on "The Island" of Honey Lake.
The species is diagnosed by the slender anterior neck of the
maxilla in the vicinity of the ventral maxillary muscle, anterior
to the dorsal and ventral crests (Taylor and Smith, 1981, pi. 19:
fig. 3) and the long, gently curved gnathal ramus of the dentary
(Taylor and Smith, 1981, pi. 19: figs. 5, 6).
69. ^Catostomus sp. (Secret Valley sucker). This form dis-
plays characters of C. cristatus (Pliocene) and C. macrocheilus
(Holocene) of the Snake and Columbia Basins, respectively. As
in C. macrocheilus, the dentaries are short, relatively straight,
and wide anteriorly, with strong ridges for the labial cartilage
(Smith, 1975, fig. 121). The maxillae have medium-high dorsal
and ventral ridges and a general shape like those of the Wall
Canyon Catostomus; however, the ridge for insertion of the pa-
latine ligament (the anterior edge of the bone), although mostly
vertical, has a small, horizontal, J-shaped curve that resembles
the condition in C. shoshonensis (Smith, 1975, fig. 12A and
points xpy on fig. 12G). McClellan (1977) collected similar
sucker bones from the Salt Lake group at the Junction Hills,
Box Elder County, Utah. Suckers from the Alturus Formation
may be correlated with the Secret Valley form (Wagner et al.,
1997). Age of the Alturus beds is correlated with the Pliocene
Glenns Ferry Formation by a tephra at Crowder Flat Road near
Alturus, California, which bears the chemical signature of a
Pliocene ash of the Glenns Ferry Formation (Swirydczuk et al.,
1982).
70. f*Catostomus shoshonensis Cope (Shoshone suck-
er). This species (including its synonym, C. ellipticus) is
from the Miocene and Pliocene Snake River Plain (Smith et al.,
1982) and from Pliocene sediments at Honey Lake, California
(Taylor and Smith, 1981). Catostomus shoshonensis is diag-
nosed by an elliptical maxilla with the attachment of the pa-
latine ligament high on the neck of the anterodorsal process
(Smith, 1975, fig. 12A). This form is represented by the relict
sucker of Wall Canyon, Surprise Valley, Nevada; this conclu-
sion is based upon the shape of the maxilla.
71. Catostomus warnerensis Snyder (Warner sucker). This
species lives in the Warner Valley in eastern Lake County, Ore-
gon, and is similar to Catostomus shoshonensis of the Snake
River Plain. It is also osteologically similar to the form in Wall
Canyon, Surprise Valley, Nevada.
72. *Catostomus tahoensis Gill and Jordan (Tahoe suck-
er). This Holocene sucker lives only in the Lahontan Basin.
Fossils have been found in Pliocene sediments at Mopung
Hills, Churchill County, Nevada (Taylor and Smith, 1981), in
middle Pleistocene, pre-Lahontan sediments of the Humboldt
River canyon, and in Holocene Indian midden material in the
western part of the Lahontan Basin (Thomas, 1985:172, and
references therein). The maxilla has a thick anterior segment in
the vicinity of the ventral maxillary muscle attachment, with
low dorsal and ventral keels that are similar to maxillae oiPan-
tosteus (mountain suckers). The dentary has a mesially in-
flected gnathal ramus suggestive of the plesiomorphic moun-
tain suckers. Catostomus tahoensis was a basal member of the
Pantosteus clade in the cladogram of Smith and Koehn (1971).
This is supported by the maxillary shape of Holocene and fossil
maxillae, which are Pantosteus-like (Taylor and Smith, 1981,
pi. 19: 1, 2). Additional material representing C. tahoensis was
collected by Marith Reheis (USGS) from sediments deposited
~1 Ma at Sunshine Amphitheater (northwest of Walker Lake)
and Lone Tree Mine, Nevada, and by J.R. Alcorn from near
Fallon, Nevada.
73. * Catostomus fumeiventris Miller (Owens sucker). This
species is osteologically similar to C. tahoensis, except in de-
tails described by Miller (1973). It is less derived than the
Mopung Hills form of C. tahoensis. Fragments of Catostomus
fumeiventris{l) were recovered from the Pleistocene Owens
Lake core (Firby et al., 1997). Holocene Indian midden materi-
als from along the Owens River near the Alabama Hills in Inyo
County, California, are rich in remains of the Owens sucker
and Siphateles bicolor snyderi.
74. ^Catostomus sp. (Mono Basin sucker). This Pliocene
fish lived with Chasmistes sp., Ptychocheilus sp., and Siphate-
les sp. Its dentary has an abrupt anterior edge of the coronoid
process, which is similar to that of Catostomus fumeiventris of
Owens Valley (Miller, 1973) and to C. tahoensis (see Reheis et
al., 2002). The maxilla is similar to that of Catostomus crista-
tus of the Miocene and Pliocene Snake River Plain, Idaho
(Smith, 1975, fig. 12G), and Secret Valley, California.
75. ^Catostomus sp. (Carson Valley sucker). Pliocene sedi-
ments in Carson Valley contain remains of a Catostomus asso-
ciated with Siphateles, Ptychocheilus, Lavinia, Chasmistes,
and Oncorhynchus. The dentary of this form is more similar to
C. occidentalis than to C. tahoensis or C. fumeiventris in the
weak peripheral labial ridge and the modest scalloping of the
lateral surface of the precoronoid ramus. These characters sug-
NUMBER 33 217
gest a pre-Lahontan connection existed that allowed popula-
tions of C. occidentalis of the Sacramento River drainage to
mix with the Carson Valley sucker of the pre-Lahontan drain-
age. According to G.I. Smith (pers. comm. to G.R.S., 1999),
such a connection existed at the head of the San Joaquin Can-
yon prior to being blocked by volcanics 3.2 Ma.
76. *Catostomus columbianus (Eigenmann and Eigenmann)
(bridgelip sucker). This species of mountain sucker (subge-
nus Pantosteus) is native to the Columbia Basin and further
north, but within the Great Basin it is restricted to the Har-
ney-Malheur Basin. It normally has small scales, but in the
Harney-Malheur Basin it has large scales and similarities to
isolated headwater populations found in the Columbia-Snake
Basin (Smith, 1966; Bisson and Bond, 1971). Catostomus co-
lumbianus is represented in the Pliocene of the Snake River
Plain. Miller and Smith (1981) described the Pliocene speci-
mens as a new species, C. arenatus, but, because no evidence
of lineage splitting has been found, there is no need for the fos-
sil name.
77. Catostomus discobolus Cope (bluehead sucker). This
species of mountain sucker (subgenus Pantosteus) arose in the
upper Colorado River drainage, but it is also native to the We-
ber River, Bear River, and upper Snake River drainages (Figure
7; Smith, 1966). It is the sister species to C. clarki of the lower
Colorado River drainage; the two species are the most morpho-
logically apomorphic in the genus. Catostomus discobolus is
the most widespread mountain sucker in the upper Colorado
River drainage, indicating this was the area of its origination
and subsequent dispersion. Populations spread from the head-
waters of the Green River to the headwaters of the Bear River
via stream capture at the east end of the Uinta Mountains, per-
haps before the upper Bear River was diverted from the Snake
River to Lake Bonneville (Bright, 1967; Bouchard et al, 1998).
It is reported from Pleistocene deposits in Homestead Cave
near Great Salt Lake by Broughton (2000a). Stream capture in
the headwaters of Chalk Creek, Summit County, Utah (Hansen,
1969) could have transferred C. discobolus from the Bear River
to the Weber River.
78. Catostomus clarki (Baird and Girard) (desert suck-
er). This sister species to C. discobolus occupies headwaters
of the lower Colorado River drainage below the Grand Canyon.
It inhabits Great Basin streams that were formerly connected to
the Colorado River in the Pahranagat Valley-White River
drainages in eastern Nevada and co-occurs with species of
clades Lepidomeda, Rhinichthys, and Crenichthys. It is clinally
variable in the Virgin River drainage. Northern populations
possess high scale counts similar to those of C. platyrhynchus
in the adjacent southern Bonneville Basin (see "Discussion").
79. * Catostomus platyrhynchus (Cope) (mountain suck-
er). This species of the subgenus Pantosteus lives in the
Bonneville and Lahontan drainages, but probably early in the
Pleistocene it spread north to the Snake and Columbia drain-
ages, then east to the Colorado and Missouri drainages (Smith,
1966). Pleistocene fossils are unknown from the Great Basin,
but specimens have been collected by J.D. Stewart from the
Pleistocene of northwestern Nebraska. It is the sister to the
clade of three species, C. santaanae (of the Los Angeles Ba-
sin), C. clarki (lower Colorado River drainage), and C. disco-
bolus (upper Colorado River drainage). Differences between
the Bonneville Basin and Lahontan Basin forms are slight. The
population of C. platyrhynchus in Shoal Creek, Washington
County, Utah, shows introgressed characters from C. clarki of
the Virgin River drainage, possibly as a result of genetic con-
tact at the time of the outlet of Lake Bonneville to Meadow
Valley Wash in the middle Pleistocene (Ives, 1948).
ICTALURIDAE
80. ^Ameiurus hazenensis Baumgartner (Truckee bull-
head). This is one of only two Late Cenozoic catfishes of the
Great Basin. It is known only from late Miocene sediments of
the Truckee Formation, near Hazen, Nevada (Baumgartner,
1982). Ameiurus hazenensis is related to A. peregrinus of the
Chalk Hills Formation, Oregon, to A. vespertinus of the Glenns
Ferry Formation, Idaho (Lundberg, 1992), to Ameiurus sp.
from the Junction Hills, Box Elder County, Utah (see number
81, below), and to Ameiurus reticulatus from the Ringold For-
mation of Washington (Smith et al., 2000).
81. ^Ameiurus sp. (Junction Hills bullhead). McClellan
(1977) collected spines of a small bullhead, possibly of Hemp-
hillian age, from the Junction Hills, Box Elder County, Utah.
SALMON ID AE
COREGONINAE
82. *Prosopium williamsoni (Girard) (mountain white-
fish). This widespread western North American mountain
species lives in the Bonneville, Harney, Lahontan, and sur-
rounding basins (Figure 8). Fossil scales of Prosopium (not
necessarily williamsoni) were taken from the Old River Bed
units of Lake Bonneville (Oviatt, 1984) and the Pleistocene
part of the Owens Lake core (Firby et al., 1997). The earliest
known form of Prosopium, i.e., P. prolixus, occupied Lake
Idaho on the Snake River Plain throughout much of the
Pliocene (Smith, 1975). Cold adaptations, cladistic relation-
ships, and paleogeography indicate that the genus Prosopium is
northern in origin.
83. Prosopium abyssicola (Snyder) (Bear Lake white-
fish). This whitefish is now endemic to Bear Lake, a part of
the Bear River drainage, which is a tributary to Lake Bonne-
ville. It is not represented among the abundant Lake Bonneville
fish bones examined by Smith et al. (1968) or Broughton
(2000a, 2000b).
84. *Prosopium spilonotus (Snyder) (Bonneville white-
fish). This species is now a Bear Lake endemic. Prosopium
spilonotus existed in late Lake Bonneville (Smith et al., 1968;
Broughton, 2000a, 2000b).
218 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
85. *Prosopium gemmifer (Snyder) (Bonneville cis-
co). Prosopium gemmifer is a specialized planktivore that is
now endemic to Bear Lake, but a similar form also occurred in
Lake Bonneville (Smith et al., 1968; Broughton, 2000a,
2000b). The two forms of P. gemmifer differ significantly in
bones that reflect the shape of the head and jaws. The Bear
River was a tributary to the Snake River prior to the late Pleis-
tocene diversion of the Bear River into the Bonneville Basin at
Soda Springs (Bright, 1967; Bouchard et al., 1998).
THYMALLINAE
86. Thymallus arcticus (Pallas) (Arctic grayling). Grayling
have not been previously acknowledged as native to the Great
Basin fauna, but they were recognized among fishes collected
in 1834 from the Bear River near Georgetown, Idaho, by John
Kirk Townsend (1978:89). Townsend was a physician and zo-
ologist working with Thomas Nuttall, a British botanist from
Harvard University who was en route to the Columbia River.
Nuttall was familiar with Great Lakes fauna, which includes
grayling and char. Townsend was the first trained zoologist to
cross the American continent, collecting specimens and record-
ing observations of natural history. Townsend recorded gray-
ling with trout and whitefish in the headwaters of the Green
River (1978:84 (June 30)) and contrasted them in a way that
implies that he and Nuttall knew the difference between gray-
ling and whitefish: "The river, here, contains a great number of
large trout, some grayling, and a small, narrow-mouthed white-
fish, resembling a herring. They are all frequently taken with
the hook." On the Bear River, about 12 miles south (upstream)
of Soda Springs, Idaho, on 6 July 1834, Townsend (1978:89)
wrote: "Trout, grayling, and a kind of char are very abundant
here?the first very large." On the previous day, he recorded:
"This is a fine stream of about 150 feet in width, with a move-
able sandy bottom." (This was prior to the overgrazing, ero-
sion, and destruction of these aquatic habitats that commenced
three to four decades later.) Grayling are now native, nearby, to
the upper Missouri River in Montana, in streams that were trib-
utaries to the Hudson Bay during preglacial and early glacial
times (Menely et al., 1957; Howard, 1960; Lemke et al.,
1965:15). Grayling had access from the upper Missouri River
to the North Platte River, the upper Green River basin (Sears,
1924; Bradley, 1936), the upper Snake River, the Bear River,
and the Bonneville Basin through headwater-stream captures
northeast of the Uinta Mountains (Hansen, 1969; see "Discus-
sion"). There are no Holocene or fossil specimens from the
Great Basin in the collections we examined.
SALMONINAE
87. *Salvelinus confluentus (Suckley) (bull trout). Town-
send (1978), in the account of his (1834) westward journeys,
reported "charr" from the Bear River, as pointed out by Ros-
tlund (1951). That the upper Bear River was sufficiently cold
and stable in 1834 to support grayling, char, and large trout
(Townsend, 1978:88; see Thymallus arcticus, above) has rarely
been considered. In 1934, Steven D. Durrant, of the University
of Utah, discovered a population of S. confluentus in the Jar-
bidge drainage, a tributary to the Snake River drainage, north
of the edge of the Great Basin (Miller and Morton, 1952; Cav-
ender, 1978). Hubbs and Miller (1948b) captured specimens in
the Lost River system on the eastern Snake River Plain, an area
that shares most of the Bonneville Basin fauna. Bull trout could
have colonized the Bonneville Basin from the Snake River
while Lake Bonneville was a tributary to the Snake River
(Bright, 1967). Salvelinus confluentus might have gained ac-
cess to the Snake River via headwater transfer from the Salmon
River (Behnke in Cavender, 1978) and then to the Bear River
when it was a tributary to the Snake River. Char also could
have occupied the Great Basin throughout the late Cenozoic.
One premaxilla with diagnostic characteristics of S. confluen-
tus was identified by Jack Broughton from abundant fossils in
Homestead Cave (Broughton, 2000a, 2000b). We (GRS and
KWG) have examined the maxilla and concur in the identifica-
tion; we are searching for additional bones to confirm the re-
cord. In the middle Miocene, about 14 Ma, Salvelinus was in
central Nevada (Churchill County, collected by J.R. Alcorn and
Ted Cavender; UMMP57107) and in the Columbia River
drainage of Idaho (Smiley and Rember, 1985) as Hucho sp.
(Smith and Miller, 1985). It lived in the Snake River Plain in
the late Miocene (referred to as "\Hucho larsoni) (Kimmel,
1975; Smith et al., 1982). Native Salvelinus are now extinct in
the Great Basin.
88. ^Oncorhynchus cynoclope (LaRivers) (Rabbit Hole
trout). This trout was named for Rabbit Hole, Pershing
County, Nevada, by LaRivers in 1964. "Salmo esmeralda" was
named by LaRivers in 1966; the description was based upon
fragmentary material from the late Miocene Alum Mine, Nye
County, Nevada, in the upper portion of the Esmeralda Forma-
tion. It is not clear how these trouts are related to each other or
to a trout from the Miocene Truckee Formation of western Ne-
vada. The oldest name is used until more material enables more
detailed comparisons. Salmonid bones from Stewart Valley,
Mineral County, Nevada, are slightly older than the specimens
from the Esmeralda Formation (15-12 Ma, radiometric date);
they were found with plants, insects, mammals, and other
fishes. A late Hemphillian salmonid dentary was collected by
Tom Kelley with mammals from the Coal Valley Formation,
Mineral County, Nevada. A trout from Carson Valley, Nevada,
is known from vertebrae of individuals about 45 and 54 cm in
standard length. The perforated pattern on the centra of the ver-
tebrae and the smooth connections of the arches to the centra
indicate the vertebrae are different from recent salmon and
trout. A dentary of a similar salmonid was collected by Lily
Lugaski with cyprinid bones and fossil plants from the Buffalo
Canyon Formation (-15 Ma), Churchill County, Nevada.
89. ^Oncorhynchus sp. (Truckee trout). The Truckee For-
mation produced a small, nearly complete trout collected by
NUMBER 33 219
Michael Bell from a diatomite mine near Hazen, Nevada. The
late Miocene diatomite deposit also produced Fundulus, Gila,
Ameiurus (Baumgartner, 1982), and a rich record of Gasteros-
teus evolution described by Bell (1994) (see number 128, be-
low).
90. *Oncorhynchus mykiss gairdneri Richardson (redband
trout of the inland Columbia River and Oregon Lakes
basins). The redband trout of the Columbia Basin between
the Cascade Range and upstream barrier falls are closely re-
lated to the redband trout of the Oregon Lakes division of the
Great Basin and those of Goose Lake in Oregon and California
(Figure 9; Behnke, 1992, map). Unusual variation exists within
and among the many populations of the redband subspecies O.
m. gairdneri. The fish typically have 130-170 large scales in
the lateral line (mean = 135-160), 63-66 vertebrae, zero to
many basibranchial teeth, 18-24 (mean = 20) gill rakers, 30-50
pyloric caeca, 58 chromosomes, intermediate-sized spots and
parr marks, variable cutthroat marks, and light-colored tips on
the dorsal, anal, and pelvic fins (Behnke, 1992). Some of this
variation (especially in basibranchial teeth and cutthroat
marks) is shared with cutthroat trout, O. clarki. Many popula-
tions of redbands were originally regarded as O. clarki (Cope,
1889; Snyder, 1908; Hubbs and Miller, 1948b). In particular,
the trout of the Fort Rock Basin frequently have a high basi-
branchial tooth count (Behnke, 1992) and are more like O.
clarki than O. mykiss, as were their Pliocene and Pleistocene
ancestors in the basin (see below). The general trend of red-
band characters is, however, less like O. clarki and more like
O. mykiss. Cope (1879) and Behnke (1992) identified redband
trout populations as O. mykiss.
All authors have agreed that the populations in question
show composites of cutthroat troat and rainbow trout charac-
ters, but each has offered different ideas about whether the
mixture was plesiomorphic or the product of ancient or post-in-
troduction introgression. Gilbert and Evermann (1894) and
Needham and Behnke (1962) regarded redband trout to be the
product of introgression. Redband trout have been reported
(Behnke, 1992) from the McCloud drainage (northern Califor-
nia home of hatchery rainbows), from the Goose Lake drainage
of the Sacramento system, and from five desert basins of the
"Oregon Lakes system" (Snyder, 1908; Allison, 1940; Hubbs
and Miller, 1948:67), namely, the Harney-Malheur, Fort Rock
Valley, Catlow Valley, Lake Chewaucan, and Warner Lakes.
These are west and north of the basins containing O. clarki,
namely, Alvord and Whitehorse Basins (Legendre et al., 1972;
Legendre, 1976; Smith, 1978; Behnke, 1992). Hubbs and
Miller (1948b) described the close relationships of trouts, min-
nows, and suckers to the pluvial lakes in these basins.
The Harney-Malheur and Catlow Valley basins (central Har-
ney County, Oregon) were occupied by Malheur and Catlow
Lakes, respectively, in the late Pleistocene and possibly many
times earlier. The Harney-Malheur Basin fishes are derived
from two different colonizations, as documented by Bisson and
Bond (1971): (1) the John Day River connection with the Co-
lumbia River to the north; and (2) the Malheur River connec-
tion east to the Snake River, dated to have occurred less than 1
Ma (Luedke and Smith, 1982). The Harney-Malheur Basin
possibly received drainage from Catlow Valley in pluvial
times, but this neighbor to the south is now isolated (Allison,
1982) and nearly fishless. Catlow Valley lost the heads of its
eastern tributaries to capture by Donner and Blitzen Creeks of
the Malheur Lake basin on the west side of Steens Mountain in
southeast Harney County. The trouts in these two drainage ba-
sins are similar, but those from Catlow Valley have larger
scales, as expected for the more southern fish in a cline.
Behnke (1992) regarded the native inhabitants of these basins
to be redband trouts. This conclusion is supported by testimony
from older inhabitants of Steens Mountain, indicating that na-
tive redband trout were in Home, Three-Mile, and Skull Creeks
in the Catlow Basin (west of Steens Mountain) as early as
1917. By contrast, these residents noted, cutthroat trout were
native on the east side of Steens Mountain (letter from Stanley
Jewett to Carl Hubbs, 13 Jan 1937). The Catlow Valley red-
band trout in the UMMZ (collected by S. Jewett in 1936-1937)
have 20-22 gill rakers, large, posteriorly concentrated spots,
and no basibranchial teeth.
Holocene redband trout examined by Behnke from the Fort
Rock drainage were variable, with an intermediate frequency
of basibranchial teeth that is expected in products of introgres-
sion or plesiomorphic character retention. As a consequence of
this mixture of characters, we conclude that the Fort Rock trout
(ancient and modern) are not cladistically diagnosable as O.
clarki or O. mykiss. They possess a mixture of the characters of
O. mykiss and O. clarki from the various Oregon lakes. Appar-
ently, as O. clarki and O. mykiss were diverging elsewhere in
the Pliocene and Pleistocene, a mixture of ancestral characters
were only partially sorted out in the Oregon Lakes system, pos-
sibly under the influence of gene flow from the north and
south. The population remained mixed in the Fort Rock Basin.
A relative of f O. lacustris also apparently inhabited the Fort
Rock Basin in the Pliocene. The reports of Pacific salmon in
the Fort Rock Basin (Hubbs and Miller, 1948b) are based upon
bones of redband trout (Allison and Bond, 1983).
The redband trouts of the Warner Lakes (pluvial Lake Cole-
man; Allison, 1982) and nearby Chewaucan Basin in southern
Lake County, Oregon, tend to have slightly fewer vertebrae and
more gill rakers, indicating closer relationships to the Goose
Lake (Sacramento drainage) redband trout (Behnke, 1992).
91. Oncorhynchus mykiss aquilarum Snyder (Eagle Lake
trout). These fish were recognized by Snyder (1917a) and
Behnke (1972, 1992) as a population of rainbow trout, Onco-
rhynchus mykiss, that possibly transferred from the Pit River
drainage to Eagle Lake by a stream capture in Lasson County,
California. Hubbs and Miller (1948b) identified O. m. aqui-
larum as a hybrid between native O. clarki and introduced O.
mykiss gairdneri, redband trout.
Fossils resembling Pliocene Oncorhynchus lacustris of the
Snake River Plain were collected by Taylor and Smith (1981)
220 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
and by Charles Dailey of Sierra College in deformed Pliocene
sediments at "The Island" of Honey Lake, south of Eagle Lake
in the northwestern part of what is now the Lahontan drainage.
Although the Honey Lake fossils are not necessarily directly
related to the Eagle Lake trout, both O. m. aquilarum and O.
lacustris have mixtures of O. clarki and O. mykiss characters.
The morphological and geographical circumstances are impor-
tant because the diagnostic fossil maxillae (round in cross sec-
tion) indicate the Pliocene presence of O. lacustris or O.
mykiss characters in the area. The Honey Lake sediments also
contain two species of Catostomus and one species each of
Chasmistes, Mylopharodon, Klamathella (with two rows of
pharyngeal teeth), and Siphateles (with a single row of teeth).
The exact hydrographic relationship of the highest stages of
Eagle Lake to the Lahontan Basin is hidden in several possible
shorelines above and below the rim of Eagle Lake basin,
which have been summarized by Hubbs and Miller
(1948b:37-38). Observations on recent shorelines and connec-
tions are not necessarily relevant to the history of these fishes,
however. It is possible that the rainbow trouts called O. m.
aquilarum are relicts of the ancient redband trout of the Snake
River-Sacramento River drainage of the early Pliocene and
are part of the series of redband trouts of the Oregon Lakes
System, described above. Two Oncorhynchus maxillae from
Secret Valley (Madeline Plains, Modoc County, California)
sediments (Wagner et al., 1997) establish the presence of a
trout similar to the O. clarki-O. mykiss series north of Eagle
and Honey Lakes in the Pliocene. A left dentary from Savage
Canyon in the Mineral Mountains, Mineral County, Nevada
(MVP 58685), most closely resembles O. lacustris.
92. *Oncorhynchus clarki (Richardson) (cutthroat
trout). These are inland trout of the Rocky Mountains (Fig-
ure 9). They are the only native trout of the Bonneville and
Lahontan Basins, as well as of the Alvord and Whitehorse Ba-
sins located east of Steens Mountain, Oregon (Figure 9). But
as noted above, redband populations west of Steens Mountain
and in parts of the lower Columbia River basin possess O.
clarki characters in varying amounts. Cutthroat trouts possess
more primitive character states than other Oncorhynchus, and
their morphological characters are predominant in the older
(ancestral) fossil record. This assessment was based upon a
cladistic analysis (Stearley and Smith, 1993), which included a
trout specimen collected by Michael Bell from the late Mio-
cene-Pliocene Truckee Formation (see number 89, above) and
other trouts from this portion of western Nevada (LaRivers,
1964, 1966). Oncorhynchus clarki fossils occur in Pliocene
strata in the Carson Sink, Churchill County, Nevada, portion of
the Lake Lahontan area. Fossil bones of O. clarki were found
in Pliocene sediments in the Mopung Hills, in the central part
of the Lahontan Basin drainage (associated with Pliocene mol-
lusks, Taylor and Smith, 1981); in Savage Canyon, Cedar
Mountains, Mineral County, Nevada (UCMP 58685); and in
pre-Lahontan sediments in Sunshine Amphitheater northwest
of Walker Lake, Nevada, by Marith Reheis (USGS). A large
sample of cutthroat trout bones from the Humboldt River can-
yon, in sediments below the Lava Creek tephra (665 Ka), is
herein diagnosed as O. clarki by the following characters: the
anterodorsal strut of the posttemporal is tapered, not spatulate;
the median shelf of the anterior portion of the dentary ap-
proaches horizontal; and the ascending wing of the premaxilla
is high anteriorly. Trout scales were recovered from the Owens
Lake core at levels dated at 730 and 695 Ka (Firby et al.,
1997). Molecular relationships of O. clarki are discussed be-
low (numbers 93-99) and are shown in Figures 10 and 14; in-
terpopulation genetic distances are shown in Tables 2 and 3.
93. ^Oncorhynchus clarki henshawi Gill and Jordan (Lahon-
tan cutthroat trout). The Lahontan cutthroat trout inhabits
Pyramid Lake and its tributaries, except the Humboldt River
(inhabited by an unnamed subspecies) and Silver King Creek
of the East Fork of the Carson River (O. c. seliniris, Paiute
trout) (Behnke, 1992:111). These several populations form a
group of three subspecies that are related to the (unnamed) Al-
vord Basin and Whitehorse Basin subspecies (see numbers 96,
97, below), according to Behnke (1992) and our molecular
data (Figure 10; Table 3).
Oncorhynchus clarki henshawi differs from its closest rela-
tives in the Lahontan Basin in having large spots over most of
the body compared with the Paiute trout, O. c. seleniris, which
lacks spots, and it has more gill rakers and smaller scales than
the Humboldt River subspecies. The other characters of this
trio are similar to those of the Whitehorse Basin and Alvord
Basin trouts (Figure 10; Behnke, 1992). The five forms differ
among themselves by only 0.1% sequence divergence, accord-
ing to our analysis of their mtDNA restriction sites (Table 3).
Originally, different spawning runs separated stocks in Pyra-
mid Lake (Snyder, 1917a); an additional name, O. c. tahoensis
(Jordan and Evermann (1896)), was applied to the "great
trout" or "silver trout" from deep waters of Pyramid Lake, but
Behnke (1981, 1992) found no morphological differences.
Different size and color, if genetic and if characteristic of dif-
ferent spawning runs, would be evidence of significant (prob-
ably intralacustrine) differentiation.
Oncorhynchus clarki henshawi is among Holocene fossils
from shore-face sediments around Pyramid and Winnemucca
Lakes and from various Indian midden sites reported in the ar-
chaeological literature. These records provide documentary
evidence and osteological material of the Holocene form for
study; otherwise, museums lack adequate specimens from the
period before the nineteenth century introductions, hybridiza-
tion, and destruction.
94. Oncorhynchus clarki seleniris (Snyder) (Paiute cutthroat
trout). This trout is limited to a few headwaters in Silver
King Creek of the East Carson drainage in Alpine County, Cal-
ifornia. It differs from O. c. henshawi only in the absence of
spots. There is no evidence pertaining to the time of its isola-
tion, except for the 0.1% sequence divergence (Table 3) from
the other Lahontan Basin trouts. We assume that cutthroat
trout inhabited the stable environments of the East Carson
NUMBER 33 221
drainage and other mountain streams in the Lahontan Basin
throughout the Miocene to Pleistocene because of the fossil re-
cord of 0. clarki and its ancestors in the basin, extending back
to the Miocene.
95. Oncorhynchus clarki subsp. (Humboldt cutthroat
trout). Behnke (1992) pointed out the distinctiveness of cut-
throat trout of the Humboldt River, relative to those of the
Walker Lake, Truckee River, and Carson Valley drainages.
The Humboldt cutthroat trout tends to have larger scales and
fewer gill rakers, as expected for fish inhabiting warm, fluvial
waters. It also has distinctive mtDNA (Williams, 1990, cited in
Behnke, 1992), except in the Quinn River, where trout with
Humboldt morphology have mtDNA characteristic of trout of
the Lahontan Basin drainage. Behnke (1992:119) analyzed
possible isolation dates in the context of the main dates of
drainage separations, 13-10 Ka (Benson and Thompson,
1987), and he concluded that the divergence of the Humboldt
form from O. c. henshawi dates from an earlier time, which is
consistent with the 0.1% sequence divergence. The type speci-
men of the Pliocene O. cynoclope (LaRivers) is from the range
of this subspecies, indicating an ancient local ancestry for the
clade, but it lacks sufficient characters to clarify the ancestry
of this subspecies.
96. Oncorhynchus clarki subsp. (Alvord cutthroat
trout). These trout were collected by Hubbs from Virgin
Creek in 1933 and Trout Creek in 1934, in southeastern Har-
ney County, Oregon. Hubbs' field notes indicate that hatchery-
raised rainbow trout were introduced into Virgin Creek as
early as fall, 1933. Behnke (1992) indicated that Alvord cut-
throat trout were similar to Lahontan cutthroat trout (0. c. hen-
shawi) but had fewer spots, rosier sides and opercles, larger
scales, fewer pyloric caecae, and fewer basibranchial teeth
(Behnke, 1981, 1992). The rosy color is reminiscent of red-
band trout (0. m. gairdneri) (Tol and French, 1988). Mito-
chondrial DNA restriction sites indicate that the Alvord trout
is the sister subspecies to the Whitehorse trout (0. clarki
subsp.) (Figure 10). Protein allele frequencies of Alvord trout
(Tol and French, 1988) did not differ from frequencies of 0. c.
henshawi at five diagnostic loci, but all of the specimens
showed introgressive influence from Oncorhynchus mykiss.
Tol and French (1988) suggested that Alvord trout were de-
rived from the Lahontan (Quinn) drainage to the south, via
Summit Lake, or were introduced. Behnke (1992:123) con-
cluded that the Alvord trout was derived from Lahontan cut-
throat trout in the late Pleistocene, assuming that a Wiscon-
sinan connection was the first to provide access. This is
consistent with mtDNA divergence of 0.1% (Table 3). The
other fish lineage in the Alvord Basin, Siphateles alvordensis,
may have been derived from the Lahontan Siphateles in the
Pliocene. It is probable that both species would have used an
existing connection; if one colonized earlier, it would probably
have been the trout. Older Alvord "Trout Creek" sediments
(now elevated) contain Miocene sunfish (Smith and Miller,
1985), but these were deposited in an earlier basin and are un-
related to the trout problem.
97. Oncorhynchus clarki subsp. (Whitehorse cutthroat
trout). This form lives in the Coyote Lake drainage in south-
ern Harney County, Oregon. It bears polymorphic mtDNA
haplotypes, namely the Lahontan and Quinn (not Humboldt)
haplotype and a unique haplotype (Williams, 1990, in Behnke,
1992). Its mtDNA is similar to that of Alvord trout. White-
horse trout share most morphological characters (except fewer
pyloric caeca) with Humboldt cutthroat trout and (except for
fewer gill rakers) Lahontan trout (Behnke, 1992). Behnke
(1992) contrasted the Whitehorse trout with the morphologi-
cally dissimilar Alvord trout and determined that the White-
horse trout has fewer gill rakers, smaller scales, more pyloric
caeca, more basibranchial teeth, and more spots than the Al-
vord trout. The two are similar in their mtDNA, however. The
distribution of morphological characters, as well as the iso-
lated position of the Whitehorse Basin, suggests that the
Whitehorse trout could represent the most plesiomorphic form
in the Lahontan?Humboldt?Alvord series.
The Whitehorse trout and Alvord trout characters require a
connection to 0. clarki of the Humboldt River drainage. Ac-
cording to sedimentary data (see Minckley et al., 1986:562;
Reheis and Morrison, 1997), the lake beds of Alvord and
Whitehorse Lakes (1231 m; 4040 ft) would have been inun-
dated by one of two middle Pleistocene Eetza highstands of
Lake Lahontan (with shoreline altitudes more than 1400 m
(4600 ft)). The mtDNA molecules of the five subspecies of 0.
clarki from the Lahontan Basin and Whitehorse/Willow Creek
drainages are similar to each other at the 0.1% level. The esti-
mated rate of divergence (0.5% per 1 My; Figure 14) suggests
that these populations were connected 200 Ka. (The most re-
cent Eetza highstand is dated as just prior to the 150 Ka Wads-
worth tephra (Reheis and Morrison, 1997).) Subsequent char-
acter sorting is responsible for the similarities and differences
between Alvord trout and Whitehorse trout. Headwater capture
has been important to the distribution of this form (Behnke,
1979). Lower Whitehorse Creek and its tributary, Willow
Creek, flow north out of the Trout Creek Mountains to the dry
lake floor 12 miles east of the Alvord desert. The drainage pat-
tern (USGS 1:250,000 series, Adel Sheet, 1955) indicates two
captures: (1) the headwaters of Whitehorse Creek were cap-
tured by McDermitt Creek (T. 40 S., R. 39. E.), a tributary to
the Quinn River (Humboldt River drainage); and (2) the head-
waters of Trout Creek (T. 39 E., R. 37 W.) were apparently
headwaters of Willow Creek (T. 39 S., R. 37 E.) prior to their
capture at Flagstaff Butte by Trout Creek, a tributary to Alvord
Lake.
The Lahontan-Alvord-Whitehorse clade (Figures 9, 10) is
the sister group to the Westslope cutthroat trout, Salmo clarki
lewisi, from the Columbia Basin (e.g., Fish Lake, Wing
Creek, and Cougar Creek, Idaho), from which they differ by
0.7%-0.9% sequence divergence (implying they were hydro-
graphically connected 1.6 Ma). Oncorhynchus c. lewisi is
222 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
represented by relict populations throughout the Columbia
River drainage, including as far south as the Boise River
drainage near Atlanta, Idaho (UMMZ 234336 with cutthroat
trout marks and O. clarki osteology). The distribution and
mtDNA of Lahontan Basin and Columbia Basin trouts sup-
port an early Pleistocene connection between the Lahontan
Basin and Columbia River drainages, as hypothesized by Re-
heis and Morrison (1997), but the other species in the two
faunas are too different to support a direct spillover from
Lake Lahontan to the Snake River drainage. Together, O. c.
lewisi and O. c. henshawi and their relatives are the sister
group to Oncorhynchus clarki clarki, from which they differ
by 1.3%-1.8% sequence divergence, implying separation oc-
curred about 3 Ma (Figure 10; Table 3).
98. * Oncorhynchus clarki Utah Suckley (Bonneville cut-
throat trout). The Bonneville Basin contains 41 populations
of cutthroat trout (Duff, 1988) differentiated into three kinds:
Bonneville, Snake Valley, and Bear River (Loudenslager and
Gall, 1980; Martin et al., 1985; Behnke, 1992). They were
characterized by Behnke (1992) as subdued in coloration,
133-183 (160) lateral scale rows, 33^6 (38) scales above the
lateral line, 16-24 (19) gill rakers, 25-54 (35) pyloric caeca,
and 1-50 basibranchial teeth in at least half the population
(Hickman, 1978; Duff, 1988). The counts are basically those
of the Yellowstone cutthroat trout, O. c. bouvieri. Mitochon-
drial DNA restriction-site analysis indicates that the Bonne-
ville trouts, exclusive of the Bear River forms, are the plesio-
morphic sister group to all other O. clarki (Figure 10),
differing from them by 1 %-4% sequence divergence (Table
3). The Deaf Smith Creek (Salt Lake County, Utah) population
differs from other Bonneville cutthroat trouts in its mtDNA
(sequence divergence >2% from other Bonneville cutthroat
trouts and 3%-4% from all other cutthroat trouts; Table 3).
The Bonneville cutthroat form shares a protein allele with the
Colorado River subspecies (Martin et al., 1985; Shiozawa and
Evans, unpublished data) and has representatives in Santa
Clara River headwaters in Pine Valley in southwestern Utah
(see "Discussion"). Fossil cutthroat trout are known from the
late Pleistocene Black Rock Canyon locality (shoreline dated
at 11,000-13,000 years B.P.) in northeastern Bonneville Basin
(Smith et al., 1968); from Homestead Cave, 10,160-11,170
years B.P. (Broughton, 2000a, 2000b); and from the late Pleis-
tocene Old River Bed between the northern Bonneville and
Sevier Basins (Oviatt, 1984).
99. * Oncorhynchus clarki subsp. (Snake Valley cutthroat
trout). The Snake Valley cutthroat trout forms tend to have
more spots than other Bonneville cutthroat trouts. Behnke
(1992) suggested that trouts were possibly absent from the
Bonneville Basin until the late Pleistocene, and Duff (1988)
postulated they had an insufficient time of isolation to become
a distinct species. We interpret the cladistic primitiveness of
the diversity in the Bonneville Basin and the level of diver-
gence of the Bonneville cutthroat trout forms to indicate an an-
cient, possibly Miocene, presence of cutthroat trout in what
became the Bonneville Basin. Fossils from 12 Ka or older
have been reported from Smith Creek Canyon in Snake Valley
(Meadetal., 1982).
100. Oncorhynchus clarki subsp. (Bear River cutthroat
trout). The Bear River form shares proteins and mtDNA
with the Yellowstone cutthroat trout, which is consistent with
its former connections to the Snake River, between 140 Ka and
80 Ka (Bouchard et al., 1998). The Bear River and Thatcher
Basin were most recently diverted to the Bonneville Basin sys-
tem about 20 Ka (Bright, 1966, 1967; Bouchard et al., 1998).
The Bear Lake relative of the Bear River form has more py-
loric caecae, an azure blue color that is dense on the snout but
diminishes posteriorly, and yellowish fins (Nielson and
Lentsch, 1988). The Bear Lake form is distinct in its life his-
tory, reproductive separation, and coloration (Nielson and
Lentsch, 1988; Behnke, 1992). The Bear River form is the sis-
ter group to the Yellowstone River and Snake River popula-
tions (Figure 10), differing from them at the 0.3%-0.4% level
in its mtDNA restriction sites (Table 3). These populations dif-
fer from O. clarki Utah of the main Bonneville Basin system at
the 1.4%?3.5% level, indicating their separate evolution
through most of the Pleistocene.
CYPRINODONTIDAE
101. Cyprinodon radiosus Miller (Owens pupfish). This
species inhabits the Owens River valley, which is a tributary to
Owens Lake, Mono County, California. It is the sister species
to C. macularius of the lower Colorado River drainage (Ech-
elle and Dowling, 1992; Echelle and Echelle, 1993; Figure
11); it differs from that species by 2.8% sequence divergence
of mtDNA (Table 4). Cyprinodon macularius and C. radiosus
are related to various species of pupfishes from the Death Val-
ley system (Miller, 1946a, 1948) and northern Mexico (Figure
11). Cyprinodon radiosus is similar to C. macularius in having
seven pelvic-fin rays, dorsal origin at the midbody, and dis-
rupted vertical bars in females, but it shares an apomorphic
scale structure with the Death Valley species (Miller, 1943b,
1948). It is additionally characterized by having large scales
and a large, deep body, like C. nevadensis; its long dorsal and
anal fins and long prehumeral length resemble those of C. di-
abolis. Male coloration is unique, with amber-colored margins
on the dorsal and anal fins and a pale or amber-colored (not
black) terminal band on the caudal fin.
Cyprinodon nevadensis Eigenmann and Eigenmann
(Amargosa pupfish). This is a geographically variable and
widespread pupfish in the Death Valley system. It is found in
numerous springs from Ash Meadows in Nye County, Nevada,
to Shoshone, Tecopa, and Saratoga Springs and other localities
along the Amargosa River near the boundary of Inyo and San
Bernardino Counties, California (Miller, 1948). The species is
large and deep, with large scales, preorbital scales; broad,
rounded, incised outer tooth-cusps; six pelvic-fin rays; and the
dorsal-fin origin usually behind the midbody. Males are deep
NUMBER 33 223
blue with a dark terminal border on the pale caudal fin; fe-
males have continuous vertical bars. Protein and mtDNA evi-
dence shows the C. nevadensis group (Figure 11) to be related
to C. salinus, C. diabolis, and C. milleri, with evidence of pos-
sible introgressive homoplasy (Echelle and Dowling, 1992;
Echelle and Echelle, 1993).
102. Cyprinodon nevadensis nevadensis Eigenmann and Eig-
enmann (pupfish). This subspecies, from Saratoga Springs in
southern Death Valley, California, is characterized by its ro-
bust body shape; high, narrow scales; numerous radii; and
dense reticulations (Miller, 1948).
103. Cyprinodon nevadensis amargosae Miller (pup-
fish). This subspecies is from the Amargosa River in San
Bernardino County, California, including Death Valley. Small
scales with few radii characterize it. Miller (1948) recognized
several isolated populations as different races.
104. Cyprinodon nevadensis callidae Miller (pupfish). This
subspecies, from Tecopa Hot Springs along the Amargosa
River in Inyo County, California, is diagnosed by large scales,
narrow interorbital distance, and a short caudal peduncle
(Miller, 1948).
105. Cyprinodon nevadensis shoshone Miller
(pupfish). This subspecies is from Shoshone Springs, Inyo
County, California. It is characterized by a narrow, slender
body and relatively straight ventral profile (Miller, 1948).
106. Cyprinodon nevadensis mionectes Miller (pup-
fish). This subspecies is from Big Spring, Ash Meadows,
Nye County, Nevada. Miller (1948) diagnosed it by its small
size, few scales and fin rays, narrow body, posterior dorsal fin,
long head, and arched predorsal profile. There are several dis-
tinct races.
107. Cyprinodon nevadensis pectoralis Miller (pup-
fish). This subspecies is from near Lovell's Spring, Ash
Meadows, Nye County, Nevada. Small scales, large number of
pectoral-fin rays, and robust body shape characterize it (Miller,
1948).
108. Cyprinodon salinus Miller (Salt Creek pupfish). This
isolated species inhabits Salt Creek canyon, northern Death
Valley, Inyo County, California (Miller, 1943a). A median
ridge on the teeth, slender body, small scales, and scaleless
preorbital region characterize it. It is similar to C. nevadensis
and C. milleri, with which it shares small fins, especially the
pelvics, posterior position of the dorsal fin, small head and
eyes, and dark blue color. Miller (1946a) outlined geological
evidence that the divergence of C. salinus and C. nevadensis
occurred in the interval since isolation of the Salt River by the
desiccation of Lake Manly, which was estimated at that time to
be about 20 Ky (but see C. milleri, below). It is related to C.
nevadensis and is the sister group to C. milleri, according to
protein evidence (Echelle and Echelle, 1993). Molecular data
indicate that C. salinus is the sister group to the other Death
Valley and Ash Meadows species (Echelle and Dowling, 1992;
Figure 11). Cyprinodon salinus does not differ from C. milleri
in its mtDNA, but these two differ from diabolis + nevadensis
by 0.13% sequence divergence, indicating more than 10 Ky of
divergence.
109. Cyprinodon milleri LaBounty and Deacon (Cottonball
Marsh pupfish). This distinctive species was discovered liv-
ing under most severe conditions in a marsh near Salt Creek,
several miles southeast of the type locality of C. salinus in
Death Valley, Inyo County, California (LaBounty and Deacon,
1972). The population lives below sea level in waters that ex-
perience a temperature range of 0?-40? C. The fish can sur-
vive a salinity range of 0-78 g/L. Large teeth, reduced or ab-
sent pelvic fins, reduced lacrimal pores, short, slender caudal
peduncle, small fins, and small scales diagnose the species.
The large, tricuspid teeth have a broad, spatulate central cusp
with no ridge on the outer face and a slim central shaft, like
that of C. macularius Baird and Girard but with deeper inci-
sions separating the outer cusps. It shares the absence of preor-
bital scales with C. salinus and C. diabolis; it shares small
scales and a slender body with C. salinus. Cyprinodon milleri
and C. salinus differ at 4 of 31 protein loci (Rogers genetic dis-
tance <0.06) according to Echelle and Echelle (1993), but the
mtDNA of C. salinus and C. milleri do not differ. The docu-
mentation of morphological differences by LaBounty and
Deacon (1972) showed C. milleri to be the most phylogeneti-
cally derived species in Death Valley. LaBounty and Deacon
(1972) and Minckley et al. (1986) hypothesized an isolation
interval of only a few hundred years, assuming no heterogene-
ity at the time of the most recent hydrological connections
documented by the geological studies of Hunt and Mabey
(1966) and Hunt et al. (1966). This is consistent with the 0%
sequence divergence and with geography, but it is in conflict
with protein and morphological differences. Introgressive ho-
moplasy of mtDNA is possible, in light of this inconsistency.
The time of divergence is therefore ambiguous, but this is a
possible example of rapid evolution.
110. Cyprinodon diabolis Wales (Devils Hole pup-
fish). Devils Hole is a small cave fault that opens at an
elevation of 2488 ft, in Ash Meadows, Nye County, Ne-
vada. Its pool is the sole range of the 200-400 individuals
of Cyprinodon diabolis. This small species has no pelvic
fins, long dorsal and anal fins, a posterior dorsal fin, a
convex caudal fin, and no preorbital scales. Mature males
have a black terminal band on the caudal fin; mature fe-
males lack vertical cross bars and show reversal of some
sexually dimorphic characters (Miller, 1948). The popula-
tion has lost individual territorial behavior (Liu in Turner,
1974). Mitochondrial DNA data (Figure 11) indicate that
Cyprinodon diabolis is related to the C. nevadensis clade,
and it shares most of its characters (e.g., large scales) with
C. nevadensis; however, it shares two strikingly nonreduc-
tive characters, increased dorsal- and anal-fin rays, with
C. radiosus. Turner (1974) electrophoretically analyzed
31 presumptive loci and found C. diabolis to be similar to
C. nevadensis and C. radiosus at the 0.968 and 0.936 lev-
els of similarity, respectively. Morphological and mtDNA
224 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
homoplasies in combination suggest reticulate evolution
in the history of C. diabolis (Echelle and Dowling, 1992).
111. ^Cyprinodon breviradius Miller (pupfish). This fossil
form was found with Fundulus curryi in Titus Canyon on the
east side of Death Valley, Inyo County, California (Miller,
1945b). The age was estimated by McAllister to be Miocene
(Miller, 1981), but it could be Pliocene (see number 117, be-
low). It is similar to C. macularius, but it has a slender body,
like C. salinus, and 20 caudal vertebrae, which is several more
than normal for the Death Valley group. On the basis of these
apomorphies, the fossil is assumed to be ancestral to the ances-
tor of the C. fontinalis?salinus?diabolis?nevadensis clade in
Figure 11. This clade differs from its sister species by an aver-
age of 2.9% sequence divergence.
GOODEIDAE
112. *Empetrichthys merriami Gilbert (Ash Meadows pool-
fish). This species, now extinct, was taken in five separated
springs in Ash Meadows, Nevada. It was a small, robust fish
with a deep body, large head, narrow mouth, and molariform
pharyngeal teeth (Miller, 1948; Uyeno and Miller, 1962). Ap-
parently always rare, it was possibly eliminated by scientific
collecting (Miller, 1948:101).
113. *Empetrichthys latos Miller (Pahrump poolfish). This
species lived in three springs in Pahrump Valley, Nye County,
Nevada. Each spring had a distinct subspecies. Miller (1948)
and Uyeno and Miller (1962) distinguished it from C. merri-
ami by its slender body, broad and weak mouth, pointed teeth,
and smaller scales. Springs that were inhabited by the three
subspecies are geographically close to each other, and the taxa
are morphologically similar. They differ from each other in
body and fin shapes. A fossil relative, ^Empetrichthys erdisi
(Jordan), from Pliocene sediments in the Ridge Basin, Los An-
geles County, California, had conical teeth, smaller scales, and
more vertebrae (Uyeno and Miller, 1962). Empetrichthys pha-
ryngeals were identified by W.L. Minckley from the Glendale
Formation (locality in Van Devender and Tessman, 1975) in
southern Nevada (Minckley et al., 1986:582).
114. Crenichthys baileyi (Gilbert) (White River spring-
fish). This species inhabits springs in the Pahranagat Valley
and tributaries to the White River, in Clark, White Pine, Nye,
and Lincoln Counties, Nevada. Geographic variation was rec-
ognized, and four subspecies were named by Williams and
Wilde (1981).
115. Crenichthys nevadae Hubbs (Railroad Valley spring-
fish). This species differs from C. baileyi primarily in color-
ation. It lives in Hot Springs in Railroad Valley and tributary
Duckwater Valley, Nye County, Nevada. This drainage was
probably connected to the Pahranagat Valley and the White
River system through the Penoyer and Desert Valleys (Hubbs
and Miller, 1948b).
FUNDULIDAE
116. f Fundulus nevadensis (Eastman) (killifish). This is a
late Miocene species (from at least 6 Ma) from the Truckee
Formation in the Lahontan Basin near Hazen, Nevada. It is a
long, slender fish with 11 or 12 dorsal-fin rays, the dorsal fin
well in advance of the anal fin, 10-13 anal-fin rays, and small,
circular scales in more than 65 lateral rows (Eastman, 1917;
Miller, 1945b). LaRivers (1964) reported it in correlated beds
in the Sahwave Mountains, Pershing County, Nevada.
117. ^Fundulus curryi Miller (killifish). This species, from
the east side of Death Valley, was similar in some ways to Em-
petrichthys (Miller, 1948:100). It is a short (44-66 mm SL),
deep fish with 14 or 15 dorsal-fin rays, the dorsal fin much
closer to the anal fin than to the pelvic fins, 15 or 16 anal-fin
rays, a broadly rounded anal fin, and scales like those of F. eu-
lepis (Miller, 1945b). Numerous specimens of Fundulus have
been collected from Titus Canyon, Death Valley, California.
Their ages are uncertain. The lower part of the Titus Canyon
Formation has been dated as early Oligocene; the date was
based upon the presence of Mesohippus, titanotheres, hydrac-
odont rhinoceros, artiodactyls, and rodents, as well as on ra-
diogenic dates (Saylor, 1994). But if F. curryi was collected
near Cyprinodon breviradius Miller, it could be Miocene or
Pliocene (Miller, 1945b). An old specimen of Fundulus was
recently collected at the Bullfrog Mine, southeast of Beatty,
Nevada. It was found in a dark gray to black shale beneath a
late Oligocene to early Miocene volcanic unit dated at 23 Ma.
118. ^Fundulus eulepis Miller (killifish). This species is-
from beds of Pliocene age in Death Valley near Furnace Creek.
The specimens are relatively long (~9 cm SL), with 13 or 14
dorsal- and anal-fin rays, the dorsal fin in advance of the anal
fin by about one-third of the distance from the anal fin to the
pelvic-fin origins, and about 45-50 lateral scale rows (Miller,
1945b).
119. ^Fundulus spp. (killifishes). Undescribed samples of
Fundulus have been collected from three localities on the Ne-
vada Test Site by Richard V. Wyman (of the University of Ne-
vada, Las Vegas) and others. Specimens studied so far are
large (up to 9 cm SL), with conical teeth, 13 rays in the dorsal
and anal fins, the dorsal fin slightly in advance of the anal fin,
6 pelvic-fin rays, moderately small scales on the body, scaled
cheeks, and 20 caudal vertebrae. The caudal peduncle is mod-
erately long, like that of F. eulepis. They lack evidence for
modification of the anal fin, as seen in Goodeidae. Cyprinodon
and Fundulus in the Miocene and Pliocene of the Great Basin
are similar in morphology and diversity to modern cyprinod-
ontoids (Empetrichthyines, Cyprinodon, Fundulus) in the
southwestern United States. The fossil Fundulus add to the ev-
idence of widespread, warm, low-gradient habitats during the
Miocene.
120. ^Fundulus spp. (killifishes). Fundulus have been col-
lected from two localities in the Esmeralda Formation, or beds
reported to be Esmeralda Formation northwest of Death Val^
NUMBER 33 225
ley. These locations are in different and younger rocks than the
Siebert tuff in which Fundulus lariversi was found.
121. ^Fundulus davidae Miller (killifish). A single specimen
(Cal. Tech. 10276) of this species was obtained through Lore
David from an oil driller. The specimen is from the Mohave
Desert in northwestern San Bernardino County, California, but
the formation of origin was not known. Miller (1945b) judged
that the fossil was of Pliocene or early Pleistocene age. The
specimen is short and moderately deep, with 11 or 12 dorsal-
and anal-fin rays, 6 pelvic-fin rays, a large head, 40-45 lateral
scale rows, and scale features like those of F. eulepis. Some of
the characters are similar to those of Empetrichthys erdisi
(Miller, 1945a). Specimens figured by Pierce (1959) from the
Miocene Barstow Formation near Yermo in the Mohave
Desert, might be Fundulus (Uyeno and Miller, 1962); the sup-
posed fossil fish eggs reported from this locality were shown to
be ostracods (Hubbs and Miller, 1962).
122. fFundulus lariversi Lugaski (killifish). This species
was described from the Siebert tuff (which formed 17 Ma),
from Tonopah, Nye County, Nevada. It is characterized by a
slender body, conical teeth, dorsal fin over anal fin and half
way between snout and caudal tip, 11 dorsal-fin rays, 13 anal-
fin rays, 12 pectoral-fin rays, pelvic fin reduced to seven rays,
17 caudal-fin rays, and 31 vertebrae. This species is similar to
Empetrichthys (Lugaski, 1978), but it possesses a pelvic fin.
123. ^Fundulus sp. (killifish). Fundulus was collected from
the middle Miocene Humboldt Formation north of Wells, Ne-
vada. The specimen is deposited at the Mackay School of
Mines, Unversity of Nevada, Reno.
COTTIDAE
124. *Cottus beldingi Eigenmann and Eigenmann (Paiute
sculpin). This sculpin occupies the Lahontan, northern
Bonneville, and Columbia-Snake drainages. Jordan (1924) re-
ported fossils from Pleistocene cave deposits in the Lahontan
Basin. Taylor and Smith (1981) reported the species from
Pliocene rocks at Mopung Hills. A basioccipital diagnosed only
as Cottus has been identified (by K.W.G.) from Pliocene sedi-
ments in Secret Valley, California, and a cottid parasphenoid
was collected by Julian Humphries from Virgin Valley, Nevada.
An angular-articular from middle Pleistocene sediments of the
Humboldt River gorge unquestionably represents Cottus beld-
ingi. R.R. Miller collected specimens from Pliocene diatomites
at Fish Cave, Churchill County, Nevada. This has apparently
been a Lahontan Basin cool-water species since a time when
the Snake River Plain had a strikingly different sculpin fauna
(Smith, 1987) and the central Columbia Basin had a warm,
lowland fauna lacking sculpins (G.R. Smith et al., 2000).
125. *Cottus bairdi Girard (mottled sculpin). This is the
abundant, widespread sculpin of the Snake-Columbia and
Bonneville drainages (as well as of eastern North America).
Two different forms are found in the Harney-Malheur drainage
(Bisson and Bond, 1971). The two morphs possibly colonized
from the Columbia River drainage (by stream capture from the
John Day River) and from the upper Snake River drainage
(prior to separation of the Harney-Malheur drainage from the
Snake River drainage). Pleistocene fossils were reported from
the Black Rock and Hot Springs localities of Lake Bonneville
(Smith et al., 1968). A Pliocene or Pleistocene Cottus was col-
lected from Park Valley, Box Elder County, Utah (Felix, 1956),
by Ted Cavender.
126. * Cottus extensus Bailey and Bond (Bear Lake scul-
pin). The Bear Lake sculpin is now endemic to that lake,
which is on the Utah-Idaho border, but it was common in Lake
Bonneville in the late Pleistocene (Smith et al., 1968; Brough-
ton, 2000a, 2000b).
127. Cottus echinatus Bailey and Bond (Utah Lake scul-
pin). This extinct species is known only from Utah Lake,
Utah. It has not been found among Lake Bonneville fossils, but
it probably shared a common ancestor with Cottus extensus
(Bailey and Bond, 1963).
GASTEROSTEIDAE
128. ^Gasterosteus doryssus (Jordan) (Truckee stickle-
back). The fossil three-spined stickleback (Gasterosteus do-
ryssus) occurs in the Great Basin in the Miocene Truckee For-
mation in west-central Nevada from Hazen (Bell, 1974) to the
Sahwave Mountains (LaRivers, 1964). It represents a lowland
fauna from the late Miocene (Bell et al., 1985) of Nevada when
the drainage was west to the Pacific Ocean. Thousands of spec-
imens have been collected and analyzed as a premier example
of rates of evolution; this data has been summarized by Bell
(1994, and references therein; Figure 12).
129. * Gasterosteus aculeatus Linnaeus (threespine stickle-
back). Threespine sticklebacks are recorded from late Pleis-
tocene sediments in the Mohave Basin in association with an-
cient Lakes Manix and Thomson and Mohave River sediments
(Roeder, 1985; Jefferson, 1991; Reynolds and Reynolds, 1991;
Bell, 1994). Recently collected, living specimens of Gasteros-
teus aculeatus williamsoni from high elevation in the Mohave
Basin have some males with a unique melanistic nuptial pat-
tern, suggesting that they are native (Bell, 1982). But trout
(with stray sticklebacks) were transplanted by fisheries biolo-
gists in that region in the 1930s or earlier, raising the possibility
that the modern population is introduced (Leo Shapovilov, Cal-
ifornia Department of Fish and Game, in litt. to R.R. Miller,
University of Michigan, Museum of Zoology; Miller and
Hubbs, 1969).
CENTRARCHIDAE
130. ifArchoplites sp. (perch). Late Miocene sunfish occur in
or near the Great Basin at several localities. These include Park
Valley, Box Elder County, Utah (Felix, 1956); the Salt Lake
group, (MVP 117073, collected by Patrick McClellan); Wind-
ermere Hills near Elko, Nevada; the Snake River Plain in Idaho
226 SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES
and Oregon; and Trout Creek, Harney County, Oregon (Smith
and Miller, 1985). Trout Creek specimens have been referred to
the genus Plioplarchus Cope. Two partial dentaries of large
sunfish were collected by J.A. Shotwell from the Juntura For-
mation, outside the edge of the Great Basin in Malheur County,
Oregon. Other specimens are from Miocene sediments in Kla-
math Lake, Harney and Malheur Counties. These centrarchids
are diverse and under study.
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