Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 1
Distortion of carbon isotope excursion in bulk soil organic 
matter during the Paleocene-Eocene thermal maximum
Allison A. Baczynski1,†, Francesca A. McInerney1,2, Scott L. Wing3, Mary J. Kraus4, Paul E. Morse5,6, 
Jonathan I. Bloch5, Angela H. Chung7, and Katherine H. Freeman7
1Department of Earth and Planetary Sciences, Northwestern University, Evanston, Illinois 60208, USA
2Department of Earth Sciences and Sprigg Geobiology Centre, University of Adelaide, Adelaide, SA 5005, Australia
3Department of Paleobiology, Smithsonian Institution, NHB121, P.O. Box 37012, Washington, DC 20013-7012, USA
4Department of Geological Sciences, University of Colorado–Boulder, Boulder, Colorado 80309, USA
5Florida Museum of Natural History, University of Florida, Gainesville, Florida 32611, USA
6Department of Anthropology, University of Florida, Gainesville, Florida 32611, USA
7Department of Geosciences, Pennsylvania State University, Deike Building, University Park, Pennsylvania 16802, USA
ABSTRACT
The Paleocene-Eocene thermal maximum 
was a period of abrupt, transient global 
warming, fueled by a large release of 13C-
depleted carbon and marked globally by a 
negative carbon isotope excursion. While the 
carbon isotope excursion is often identified in 
the carbon isotope ratios of bulk soil organic 
matter (δ13Corg), these records can be biased 
by factors associated with production, degra-
dation, and sources of sedimentary carbon 
input. To better understand these factors, 
we compared δ13Corg values from Paleocene- 
Eocene thermal maximum rocks in the south-
eastern Bighorn Basin, Wyoming, with those 
derived from leaf wax n-alkanes (δ13Cn-alk). 
While both δ13Cn-alk and δ13Corg records indi-
cate an abrupt, negative shift in δ13C  values, 
the carbon isotope excursions observed in 
bulk organic matter are smaller in magni-
tude and shorter in duration than those in 
n-alkanes. To explore these discrepancies, 
we modeled predicted total plant tissue car-
bon isotope (δ13CTT) curves from the δ13Cn-alk 
rec ord using enrichment factors determined 
in modern C3 plants. Measured δ13Corg val-
ues are enriched in 13C relative to predicted 
δ13CTT, with greater enrichment during the 
Paleocene-Eocene thermal maximum than 
before or after. The greater 13C enrichment 
could reflect increased degradation of autoch-
thonous organic  matter, increased input of 
alloch thonous fossil carbon enriched in 13C, 
or both. By comparing samples from organic-
rich and organic-poor depositional environ-
ments, we infer that microbial degra dation 
rates doubled during the Paleocene-Eocene 
thermal maximum, and we calculate that 
fossil carbon input increased ~28%–63%. 
This approach to untangling the controls on 
the isotopic composition of bulk soil carbon 
is an important development that will inform 
not only future studies of global carbon cycle 
dynamics during the Paleocene-Eocene ther-
mal maximum hyperthermal event, but also 
any study that seeks to correlate or estimate 
duration and magnitude of past events using 
soil organic carbon.
INTRODUCTION
The Paleocene-Eocene thermal maximum was 
an episode of abrupt (≤20 k.y. onset), transient 
(~200 k.y. duration), and significant (5–8 °C) 
global warming ca. 56 Ma that was associ-
ated with the input of thousands of gigatons of 
13C-depleted carbon into the exogenic (ocean, 
atmosphere, biosphere) system (Kennett and 
Stott, 1991; McInerney and Wing, 2011; Sluijs 
et al., 2007; Zachos et al., 2001, 2003). This per-
turbation to the global carbon cycle is recorded 
as a prominent negative carbon isotope excur-
sion in terrestrial and marine carbon archives, 
including organic carbon, carbonate, plant and 
algal lipids, and mammalian tooth enamel (for 
review, see McInerney and Wing, 2011). The 
carbon isotope excursion has been characterized 
as containing three globally recognized phases: 
the initiation or onset, an alternate semistable 
state or body, and the recovery (Bowen et al., 
2006). The exact magnitude and shape of the 
carbon isotope excursion (or excursions) can 
vary significantly in different archive materi-
als and reflect local factors such as depositional 
environment, sedimentation rate, and diagen-
esis (Bowen and Zachos, 2010;  McInerney and 
Wing, 2011). Because each carbon archive mate-
rial is affected by a different suite of conditions 
and preservational biases, understanding the 
diversity of factors that influence the expression 
of the carbon isotope excursion is essential for 
reconstructions of the Paleocene-Eocene thermal 
maximum carbon cycle perturbation (Sluijs and 
Dickens, 2012). The Paleocene-Eocene thermal 
maximum represents the best-known geologic 
analogue to anthropogenic climate change, and 
refining our estimates of the associated atmo-
spheric d13C excursion is critical.
Terrestrial records of the carbon isotope 
excursion are sometimes more complete than 
marine records because they commonly have 
higher rates of sedimentation and are not 
affected by carbonate dissolution (Diefendorf 
et al., 2010; McInerney and Wing, 2011). The 
magnitude of the carbon isotope excursion, 
however, varies widely among different terres-
trial carbon sources. The average carbon isotope 
excursions recorded by plant lipids (–5.1‰), soil 
carbonate (–5.5‰), and tooth enamel (–4.8‰) 
are significantly larger than those derived from 
bulk soil organic matter (–3.5‰; see McIner-
ney and Wing, 2011). Moreover, both the shape 
and magnitude of the carbon isotope excursion 
preserved in bulk soil organic matter can vary 
significantly, even across a geographically lim-
ited field area. For example, bulk soil organic 
matter carbon isotope ratios (d13Corg) from six 
stratigraphic sections across 16 km in the south-
eastern Bighorn Basin record differently shaped 
carbon isotope excursions and underrepresent 
the thickness of the Paleocene-Eocene ther-
mal maximum by 30%–80% (Baczynski et al., 
2013). The isotopic composition of bulk soil 
organic matter is influenced by variable sources 
of input and differential degradation of organic 
GSA Bulletin; Month/Month 2016; v. 128; no. X/X; p. 1–15; doi: 10.1130/B31389.1; 8 figures; 2 tables; Data Repository item 2016097.; 
published online XX Month 2016.
†Present address: Department of Geosciences, Penn-
sylvania State University, Deike Building, University 
Park, Pennsylvania 16802, USA; aab27@ psu .edu.
For permission to copy, contact editing@geosociety.org 
© 2016 Geological Society of America
Baczynski et al.
2 Geological Society of America Bulletin, v. 1XX, no. XX/XX
components, whereas plant biomarker sources 
are simpler and tend to have fewer effects on 
their isotopic composition. High-molecular-
weight n-alkanes with odd-over-even-carbon 
predominance (C25 to C33) are diagnostic bio-
markers for vascular plants in sedimentary rocks 
and are commonly used as paleoclimate proxies 
because of their long residence time and resis-
tance to isotopic exchange (Pedentchouk et al., 
2006; Schimmelmann et al., 2006; Sessions 
et al., 2004).
Here we compare a high-resolution n-alkane 
(C25 to C33) carbon isotope (d13Cn-alk) record with 
d13Corg records from Paleocene-Eocene sediments 
in the southeastern Bighorn Basin, Wyoming. 
We measured d13Corg values in two different 
rock types in the same field area: carbonaceous 
shales and mudstone paleosols (Baczynski 
et al., 2013). We used the compound-specific 
d13Cn-alk record to test the hypothesis that the 
combined influence of organic matter degrada-
tion and allochthonous carbon input distorts the 
pattern of stratigraphic change in d13Corg values 
such that it underestimates the magnitude of 
the carbon isotope excursion and the thickness 
of the carbon isotope excursion body. Specifi-
cally, we used n-alkane d13C values to estimate 
the carbon isotope values of the initial total plant 
tissue (d13CTT). We then constrained the relative 
contributions of soil organic matter degradation 
and allochthonous carbon input required to rec-
oncile the measured d13Corg and estimated d13CTT 
values. The extensive exposures of Paleocene-
Eocene strata in the Bighorn Basin and consid-
erable number of paleontological, sedimento-
logical, and geochemical data sets from decades 
of research provide an excellent setting and 
comprehensive framework in which to explore 
Paleocene-Eocene thermal maximum carbon 
cycling and the terrestrial expression of the car-
bon isotope excursion. Understanding the envi-
ronmental, biogeochemical, and sedimentologi-
cal effects on floodplain d13Corg values is crucial 
for any study seeking to correlate stratigraphic 
sections or estimate duration and magnitude of 
major climatic events.
MATERIALS AND METHODS
Study Area and Sampling
Upper Paleocene and Lower Eocene strata 
in the study area were deposited by fluvial sys-
tems near the southeastern margin of the Big-
horn Basin, Wyoming, which formed during the 
Laramide orogeny (Bown, 1980). These strata, 
comprising the upper Fort Union Formation and 
lower Willwood Formation, are abundant and 
well exposed throughout the Bighorn Basin. 
Hundreds of paleosol samples were collected 
from Paleocene-Eocene thermal maximum 
sections in the southeastern Bighorn Basin, 
Wyoming (Fig. 1). This study focuses on the 
Highway (HW) 16 section south of U.S. Route 
16 because it is the best-studied and strati-
graphically thickest paleosol d13Corg section in 
the southeastern Bighorn Basin. However, the 
same patterns were documented in three other 
trenched sections in the field area (CAB 10, Big 
Red Spit, and North Butte).
Organic-rich rock samples, hereafter referred 
to as carbonaceous shale (or carb shale) sam-
ples, were collected from isolated lenticular 
channel fills where plant macrofossils were 
found. The plant-bearing beds are lens-shaped 
bodies representing infilling of small, aban-
doned channels (5–20 m width; type I units of 
Wing, 1980). The infills are fine grained and 
laminated and preserve relatively complete 
leaves, indicating low depositional energy. 
This makes it likely that the n-alkanes came 
from nearby paleovegetation. Surface mate-
rial was removed from the plant-bearing beds 
to expose consolidated, less-weathered rock. 
Thirty-six rock samples for n-alkane and bulk 
soil organic matter d13C analysis were collected 
from 28 unique stratigraphic levels, spanning 
roughly 100 m of Paleocene–Eocene section 
in the southeastern Bighorn Basin. From their 
position within local stratigraphic sections, 
n-alkane sites from throughout the field area 
have been projected onto a single interpolated 
composite curve that uses meter levels from the 
HW 16 paleosol section (Fig. 2A; Table 1).
Sample Preparation and Analysis
A portion of each sample collected for 
n-alkane d13C analysis was prepared for bulk 
organic matter d13C analysis as described fully 
in Baczynski et al. (2013). Briefly, carbon iso-
tope ratios and weight percent total organic car-
bon (TOC) were measured in duplicate using 
a Costech ECS 4010 combustion elemental 
analyzer coupled to a Thermo Delta V Plus iso-
tope ratio mass spectrometer (IRMS). Carbon 
isotope values are reported in delta notation 
normalized to the international Vienna Peedee 
belemnite (VPDB) standard using acetanilide 1 
(Indiana University) and International Atomic 
Energy Agency (IAEA) reference standards 600 
and CH-3. Isotope ratios and standard devia-
tions of replicate sample analyses are reported 
in Table 1. Replicate measurements of standards 
indicate a measurement precision of 0.1‰, and 
the mean standard deviation on replicate sample 
analyses is 0.1‰.
Lipids were extracted from 40 to 200 g of 
powdered rock using a Microwave Accelerated 
Reaction System (MARS Xpress, CEM Corpo-
ration, Matthews, North Carolina) with 30 mL 
(9:1, v/v) DCM:MeOH (dichloromethane: 
methanol) at 100 °C for 15 min. The total lipid 
extract (TLE) was filtered to remove sediment 
and concentrated under a stream of nitrogen gas 
using a TurboVap LV evaporator (Caliper Life 
Sciences, Walther, Massachusetts). TLEs were 
separated into nonpolar and polar fractions by 
column chromatography using 1 g activated 
silica gel (70–230 mesh) in an ashed Pasteur 
pipet. The nonpolar constituent was eluted with 
4 mL of hexanes and the polar constituent with 
4 mL of DCM:MeOH (1:1, v/v). The n-alkanes 
were identified in the nonpolar fraction using a 
ThermoFisher Trace Gas Chromatograph Ultra 
fitted with a flame ionization detector (FID) 
and coupled to a quadrupole mass spectrometer 
(ThermoFisher DSQII). Average chain length 
(ACL) and carbon preference index (CPI) were 
calculated using FID peak areas for C25 to C37 
n-alkanes (Table 1; Table DR11). Compound-
specific stable carbon isotope ratios were mea-
sured using a gas chromatograph coupled to an 
IRMS interfaced with a gas chromatography/
combustion system (see Supplemental Mate-
rial for detailed methods [footnote 1]). Isotopic 
abundances were determined relative to a refer-
ence gas calibrated with Mix A (n-C16 to n-C30; 
Arndt Schimmelmann, Indiana University). 
Carbon isotope values of samples were normal-
ized to the VPDB scale using the Uncertainty 
Calculator (Polissar and D’Andrea, 2014) and 
are reported in standard delta notation (Table 1); 
standard errors of the mean were calculated 
using the Uncertainty Calculator (Polissar and 
D’Andrea, 2014) and range from 0.1‰ to 0.2‰ 
for individual samples (Table 1).
Model Description and Calculations
Initial total plant tissue carbon isotope values, 
d13CTT, were modeled using the C27, C29, and C31 
n-alkane carbon isotope records by applying a 
lipid-specific enrichment factor (e):
 
ε =
δ13CTT +1000
δ13Cn-alk +1000
−1×103
≈ δ13CTT − δ13Cn-alk




that was calculated from n-alkane and total plant 
tissue carbon isotope values reported in the lit-
erature for modern C3 plants (Fig. 2B; Table 2; 
Bi et al., 2005; Chikaraishi and Naraoka, 2003; 
Collister et al., 1994; Diefendorf et al., 2011). 
1GSA Data Repository item 2016097, detailed 
methods and supplementary figures, is available at 
http:// www .geosociety .org /pubs /ft2016 .htm or by re-
quest to editing@ geosociety .org.
Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 3
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Baczynski et al.
4 Geological Society of America Bulletin, v. 1XX, no. XX/XX
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Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 5
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 1
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e
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4
–
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–
29
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31
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B3
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4-
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§
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3;
 2
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e
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7
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D 
C
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 1
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 1
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 1
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 0
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1
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UH
S0
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W
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4
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dy
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**
CA
B1
-0
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06
§
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28
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S0
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0.
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10
03
#
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10
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6.
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33
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10
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#
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B 
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; 2
8.
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36
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D;
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5
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36
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6
29
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6
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08
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 1
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3
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 1
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10
07
#
H
W
 1
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4.
1
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44
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7
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8
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08
11
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 1
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2
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3.
4
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06
#
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8
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7
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28
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1
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03
06
§
CA
B 
10
; 5
4
Po
st
64
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8
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09
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B;
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6
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st
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B6
-0
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§
CA
B 
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; 7
2.
7
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st
81
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0.
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31
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29
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4.
2
*
δ1
3 C
 v
al
ue
s 
on
 th
e 
Vi
e
n
n
a
 P
ee
de
e 
be
le
m
ni
te
 (V
PD
B)
 sc
ale
 (‰
), c
alc
ula
ted
 us
ing
 th
e U
nc
ert
ain
ty 
Ca
lcu
lat
or 
sp
rea
ds
he
et 
(P
oli
ss
ar 
an
d 
D
’A
nd
re
a,
 2
01
4),
 ex
ce
pt 
for
 th
os
e i
n i
ta
lic
s 
(se
e f
oo
tno
tes
 § 
an
d #
).
† U
nc
er
ta
in
ty
 o
f δ
13
C 
va
lu
es
 re
po
rte
d 
as
 s
ta
nd
ar
d 
er
ro
r o
f t
he
 m
ea
n 
(S
EM
), c
alc
ula
ted
 us
ing
 th
e U
nc
ert
ain
ty 
Ca
lcu
lat
or 
sp
rea
ds
he
et 
(P
oli
ss
ar 
an
d D
’An
dr
ea
, 2
01
4),
 ex
ce
pt 
for
 th
os
e i
n i
ta
lic
s 
(se
e f
oo
tno
tes
 § 
an
d
#).
 
If 
sa
m
pl
e 
wa
s 
an
al
yz
ed
 o
n 
m
ul
tip
le
 in
st
ru
m
en
ts
, l
ar
ge
st
 S
EM
 is
 re
po
rte
d.
§ δ
13
C 
va
lu
es
 a
nd
 u
nc
er
ta
in
ty
 (r
ep
ort
ed
 as
 st
an
da
rd 
de
via
tio
n) 
fro
m 
Sm
ith
 et
a
l. 
(20
07
).
# δ
13
C 
va
lu
es
 c
or
re
ct
ed
 u
sin
g 
in
te
rn
al
 s
ta
nd
ar
ds
 a
nd
 u
nc
er
ta
in
ty
 re
po
rte
d 
as
 s
ta
nd
ar
d 
de
via
tio
n 
(se
e s
up
ple
me
nta
ry 
ma
ter
ial
 [te
xt 
foo
tn
ot
e 
1])
.
*
*
Av
e
ra
ge
 c
ha
in
 le
ng
th
 (A
CL
) c
alc
ula
ted
 fo
r C
25
 
to
 C
33
,
 
fro
m
 S
m
ith
 e
ta
l. 
(20
07
).
††
CP
I—
ca
rb
on
 p
re
fe
re
nc
e 
in
de
x.
Baczynski et al.
6 Geological Society of America Bulletin, v. 1XX, no. XX/XX
Average enrichment factors were calculated 
using only angiosperm data because conifers 
were not abundant at most pre- and post–
Paleocene-Eocene thermal maximum sites and 
are entirely absent from Paleocene-Eocene 
thermal maximum floras. Modern species of 
Metasequoia and Glyptostrobus, the two most 
abundant gymnosperms, produce small quan-
tities of n-alkanes (Diefendorf et al., 2015). 
Therefore, it is likely that angiosperm n-alkanes 
predominate in our samples (Bush and McIner-
ney, 2013; Diefendorf et al., 2011). Moreover, 
unlike some conifers, the average enrichment 
factors for Metasequoia and Glyptostrobus 
are similar to angiosperm enrichment factors 
(Diefen dorf et al., 2015).
The d13C values were binned into five phases: 
pre–Paleocene-Eocene thermal maximum, car-
bon isotope excursion onset, carbon isotope 
excursion body, recovery, and post–Paleocene-
Eocene thermal maximum in order to com-
pare the paleosol bulk soil organic matter and 
predicted d13CTT curves, which have different 
sample resolution (Table 1; Table DR2 [see 
footnote 1]). Phases were identified based on 
floral and faunal biostratigraphy and com-
pound-specific stable carbon isotope records 
from the southeastern Bighorn Basin (Fig. 3; 
see Baczynski et al., 2013).
To investigate the possibility that soil organic 
matter degradation resulted in 13C-enriched 
paleosol and carbonaceous shale d13Corg values, 
we used the method of Wynn et al. (2005) to 
model the kinetic fractionation of carbon iso-
topes in an open system of continuously humi-
fying soil. In this model, carbon isotope ratios 
of organic carbon are explained by Rayleigh 
distillation, where the change in the ratio of 
two isotopes, R, is a function of the fraction 
of soil organic carbon remaining, FSOC, and the 
kinetic fractionation factor during microbial 
decomposition, a:
 Rf = RiFSOC
α−1
. (1)
Equation 1 can be expressed in terms of the 
carbon isotope ratios of the initial (d13Ci) and 
final (d13Cf) organic matter by dividing each 
side by RVPDB and using d13C notation expressed 
in per mil (‰):
 δ13Cf = δ13Ci /1000+1( )× FSOCα−1( ) ×1–1 000( ) . (2)
Soil carbon degradation curves for the pre/
post–Paleocene-Eocene thermal maximum and 
Paleocene-Eocene thermal maximum inter-
vals were modeled using the average predicted 
d13CTT values (see previous) as initial carbon 
isotope values. The shape of the logarithmic 
soil organic matter degradation curve varies 
with fractionation factor, a. A fractionation fac-
tor of 0.998 is typical of microbial respiration 
( Balesdent and Mariotti, 1996; Wynn, 2007; 
Wynn et al., 2005) and was used here, although 
a range of a values was considered. We com-
pared our data to the modeled degradation 
curves, assuming that weight percent TOC is a 
rough proxy for FSOC (Wing et al., 2005).
To estimate the change in soil organic mat-
ter degradation during the Paleocene-Eocene 
thermal maximum, the fraction of soil organic 
carbon remaining was calculated for the car-
bonaceous shales and paleosols:
 
FSOC =
δ13Cf / 1000+1
δ13Ci / 1000+1
1
α−1





 
. (3)
FSOC values were calculated using the average 
predicted d13CTT values for d13Ci values (see pre-
vious), measured d13Corg values for d13Cf values, 
and a = 0.998.
A simple binary mixing equation was used to 
model the effect of allochthonous carbon inputs 
on d13Corg values:
 
Xalloch =
δ13Corg − δ13CTT
δ13Calloch − δ13CTT
,  (4)
TABLE 2. MEASURED n-ALKANE δ13C RATIOS, PREDICTED δ13CTT VALUES AND MEAN AND STANDARD 
DEVIATION OF THE THREE PREDICTED δ13CTT VALUES FOR EACH SAMPLING LEVEL
Sample name
Interpolated
composite curve
meter level (m)
Avg
δ13C27 (‰)
Predicted
δ13CTT (‰)*
Avg
δ13C29 (‰)
Predicted
δ13CTT (‰)*
Avg
δ13C31 (‰)
Predicted
δ13CTT (‰)*
Predicted δ13CTT (‰)†
Avg St. dev
PP0904 –9.8 –30.0 –26.8 –30.9 –26.4 –31.9 –27.1 –26.8 0.4
CAB3-04-06 –0.5 –29.5 –26.3 –30.9 –26.4 –31.6 –26.8 –26.5 0.3
FAS0802 2.4 –30.3 –27.1 –31.0 –26.5 –31.8 –27.0 –26.8 0.3
CAB3-04-07 5.7 –30.3 –27.1 –31.1 –26.6 –31.4 –26.6 –26.8 0.3
SCD C 8.2 –30.6 –27.4 –30.7 –26.2 –31.4 –26.6 –26.7 0.6
FAS0801/FAS0803/SW0801/
SW0803/SW0904
12.4 –30.3 –27.1 –30.9 –26.4 –31.4 –26.6 –26.7 0.3
PS 0901 20.1 –29.8 –26.6 –29.9 –25.4 –30.3 –25.5 –25.8 0.7
NUHS0819-14 21.4 –33.2 –30.0 –34.4 –29.9 –36.1 –31.3 –30.4 0.8
CAB7-04-02 26.6 –34.0 –30.8 –35.5 –31.0 –36.5 –31.7 –31.2 0.5
CAB7-04-03 27.5 –33.0 –29.8 –34.1 –29.6 –35.8 –31.0 –30.1 0.8
CAB1-04-06 28.5 –34.8 –31.6 –35.3 –30.8 –36.3 –31.5 –31.3 0.4
NUHS0817-2 29.9 –32.3 –29.1 –33.7 –29.2 –29.2 0.1
SW1003 33.8 –36.0 –32.8 –36.4 –31.9 –37.2 –32.4 –32.3 0.4
SW1009 36.1 –34.4 –31.1 –34.4 –30.0 –35.6 –30.8 –30.6 0.6
SW1010 36.8 –33.5 –30.3 –34.4 –29.9 –33.8 –29.0 –29.7 0.6
SW0805/SW0906/SW1007 44.1 –34.3 –31.1 –35.3 –30.8 –36.4 –31.6 –31.2 0.4
PP0811 50.7 –33.8 –30.6 –35.0 –30.5 –35.7 –30.9 –30.7 0.2
SW0802/SW1006 51.1 –33.4 –30.2 –34.2 –29.7 –35.0 –30.2 –30.0 0.3
CAB6-04-01.1 52.5 –33.0 –29.8 –33.2 –28.7 –34.5 –29.7 –29.4 0.6
SW0813 53.7 –34.1 –30.9 –34.2 –29.7 –34.0 –29.2 –29.9 0.9
SW0907 55.5 –30.7 –27.5 –31.4 –26.9 –32.2 –27.5 –27.3 0.3
SW0809 55.7 –30.2 –27.0 –31.2 –26.7 –31.9 –27.1 –26.9 0.2
Wing 0902A 56.0 –30.3 –27.1 –30.9 –26.4 –31.7 –26.9 –26.8 0.3
SW1001/SW1011 57.5 –30.8 –27.6 –31.3 –26.8 –31.1 –26.3 –26.9 0.6
SW0817 60.0 –32.5 –29.2 –33.1 –28.6 –32.5 –27.7 –28.5 0.8
SW0306 64.0 –30.0 –26.8 –30.6 –26.1 –32.0 –27.2 –26.7 0.6
SW0905 73.0 –29.8 –26.6 –30.1 –25.6 –31.3 –26.5 –26.2 0.6
CAB6-04-04 81.0 –29.9 –26.7 –30.8 –26.3 –31.8 –27.0 –26.7 0.4
*δ13CTT value was calculated from the δ13Cn-alk value by applying a lipid-specific enrichment factor (ε) based on modern C3 plants (see Methods). The isotopic 
fractionation factors between total leaf tissue (TT) and n-alkane used in these calculations are: ε = 3.3 for n-C27, ε = 4.6 for n-C29, and ε = 4.9 for n-C31
†Average δ13CTT value (n-C27, n-C29, and n-C31).
Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 7
where Xalloch is the proportion of allochthonous 
carbon in a sample, d13Corg is the measured 
bulk organic carbon d13C value, d13CTT is the 
predicted total plant tissue carbon d13C value 
adjusted for soil organic matter degradation, 
and d13Calloch is the estimated allochthonous bulk 
organic carbon d13C value.
Fossil Shark Teeth
Fossil shark teeth were collected from 
Paleocene-Eocene thermal maximum strata 
throughout the field area by prospecting on 
the surface of outcrop or on anthills and using 
screenwashing techniques at localities that had 
demonstrated vertebrate microfossil preserva-
tion. Sediment was collected in bags, soaked in 
buckets to dissolve clays, and poured through 
fine mesh screens. Sorting of the resulting con-
centrate revealed vertebrate fossils, including 
mammals, squamates, and small shark teeth. 
The position and height above the global ellip-
soid of each fossil locality were recorded to sub-
meter precision using a differentially corrected 
global positioning system. We then projected 
the relative vertical position of fossil localities 
into the nearest measured section by calculating 
the difference between locality elevation and 
the elevation of laterally persistent marker beds 
(Baczynski et al., 2013).
RESULTS
Leaf Wax n-alkanes
The nonpolar lipid fractions contained pre-
dominantly high-molecular-weight n-alkanes 
that displayed a strong odd-over-even-carbon 
predominance characteristic of vascular plants 
(Bray and Evans, 1961; Bush and McInerney, 
2013). CPI values range from 2.2 to 5.2 (Table 1; 
Table DR1 [see footnote 1]). The n-alkanes typi-
cally range from n-C23 to n-C33 in most samples, 
and the most abundant n-alkane chain lengths 
are n-C27, n-C29, and n-C31. Average chain length 
ranges from 27.7 to 31.2 (Fig. 3D; Table 1; Table 
DR1 [see footnote 1]). ACL increased ~1.1 units 
during the Paleocene-Eocene thermal maxi-
mum, as previously observed in the Bighorn 
Basin (Smith et al., 2007) and elsewhere (Hand-
ley et al., 2008; Schouten et al., 2007). The car-
bon isotope values of individual n-alkanes vary 
between –28.8‰ and –37.9‰ (Fig. 2A; Table 1). 
Within a particular sample, the carbon isotope 
values tend to become more negative as chain 
length increases, a trend that has also been 
observed in leaf waxes of modern plants (Bi 
et al., 2005; Chikaraishi and Naraoka, 2003; 
Collister et al., 1994; Diefendorf et al., 2011). 
The n-alkane carbon isotope curve shows an 
A B C
δ13CE (‰, VPDB)δ
13Cn-C29 (‰, VPDB)δ
13Corg (‰, VPDB)
90
–10
0
10
20
30
40
50
60
70
80
–12–16 –14
–20
meter
level
Fo
rt 
U
ni
on
 F
m
W
ill
w
oo
d 
Fm
S
E
 B
ig
ho
rn
 B
as
in
 C
om
po
si
te
 S
ec
tio
n 
C
f-3
 b
io
zo
ne
W
a-
0 
bi
oz
on
e
W
a-
1 
bi
oz
on
e
δ13Ccarb (‰, VPDB)
–29 –28–38 –36 –34 –32 –30–23–27–28 –24–25–26 –18 –12 –6–10 –8–16 –14
Copecion
Ectocion
3227 28 29 30 31
D
ACL
Figure 3. Negative carbon isotope excursion (CIE) recorded by different carbon isotope records: (A) bulk soil organic matter (black) and 
pedogenic carbonate (gray; Baczynski et al., 2013), (B) n-C29 alkane (this study), (C) tooth enamel δ13C values (δ13CE; Secord et al., 2012), 
and (D) average chain length (ACL) from the southeastern Bighorn Basin, Wyoming. Gray shading denotes Paleocene-Eocene thermal 
maximum body, which is supported by biostratigraphy (left). VPDB—Vienna Peedee belemnite.
Baczynski et al.
8 Geological Society of America Bulletin, v. 1XX, no. XX/XX
interval of stable d13Cn-alk values in the upper-
most Fort Union Formation, an abrupt negative 
shift in d13Cn-alk values at the base of the Will-
wood Formation, an extended carbon isotope 
excursion body (~30 m), a fairly abrupt recov-
ery (~5 m) to less negative d13Cn-alk values, and 
an interval of stable d13Cn-alk values at or near 
pre-excursion values (Fig. 2A). The magnitude 
of the negative carbon isotopic shift between 
12.4 and 26.6 m is –3.9‰ for n-C25, –3.8‰ for 
n-C27, –4.6‰ for n-C29, and –5.1‰ for n-C31.
Carbonaceous Shales: δ13Corg Ratios 
and Weight Percent TOC
Carbonaceous shale d13Corg ratios range from 
–24.2‰ to –31.4‰ (Table 1; Fig. 2C). Although 
more variable and 13C-enriched relative to the 
d13Cn-alk and predicted d13CTT records, the car-
bonaceous shale d13Corg record typically reflects 
the overall shape of the n-alkane and predicted 
total plant tissue isotope curves (Fig. 4A). Car-
bonaceous shale weight percent TOC varies 
from 0.1% to 2.6% (Fig. 4C; Table 1).
Soil Organic Matter Degradation
To explore the influence of degradation on 
soil organic matter 13C enrichment, we com-
pared the relationship between 1/TOC and 
d13Corg values of the paleosol and carbonaceous 
shale samples to the modeled degradation 
curves (Fig. 5). The modeled degradation curves 
for the pre/post–Paleocene-Eocene thermal 
maximum (blue) and Paleocene-Eocene thermal 
maximum (red) use the average predicted d13CTT 
values for d13Ci, –26.64‰ pre/post–Paleocene-
Eocene thermal maximum and –30.62‰ Paleo-
cene-Eocene thermal maximum, and a = 0.998. 
The logarithmic soil organic matter degradation 
curves fit the higher TOC (lower 1/TOC) car-
bona ceous shale data well (Fig. 5). However, the 
lower TOC (higher 1/TOC) paleosol data do not 
fit well. As TOC decreases (1/TOC increases), 
the paleosol d13C values all converge near –25‰ 
(Fig. 5).
To estimate the change in soil organic mat-
ter degradation during the Paleocene-Eocene 
thermal maximum, the fraction of soil organic 
carbon remaining was calculated for both paleo-
sol and carbonaceous shale samples using Equa-
tion 3. The average FSOC for carbonaceous shales 
is 0.76 ± 0.36 (n = 15; 1s; Table DR2 [see foot-
note 1]) before and after the Paleocene-Eocene 
thermal maximum and 0.38 ± 0.32 (n = 13; 1s) 
during the Paleocene-Eocene thermal maximum. 
The average pre–Paleocene-Eocene thermal 
maximum paleosol FSOC is 0.49 ± 0.11 (n = 53; 
1s), and the average Fsoc value estimated for the 
paleosols during the Paleocene-Eocene  thermal 
 paleosol 
 carb shale
A B C
 predicted δ13CTT 
 carb shale δ13Corg  
 predicted δ13CTT 
 paleosol δ13Corg  
90
–20
–10
0
10
20
30
40
50
60
70
80
–23–33 –31 –29 –27 –25 –23–33 –31 –29 –27 –25 30 0.5 1 1.5 2 2.5
Weight % TOCδ13C (‰, VPDB)δ13C (‰, VPDB)
S
E
 B
ig
ho
rn
 B
as
in
 C
om
po
si
te
 S
ec
tio
n 
(m
)
Figure 4. Measured and predicted δ13C values and weight percent total organic carbon (TOC). (A) Comparison of the measured carbona-
ceous shale δ13Corg (light blue) and predicted δ13CTT (gray) records. Dark-gray shading highlights mismatch between records. (B) Compari-
son of the measured paleosol δ13Corg (red) and predicted δ13CTT (gray) records. Dark-gray shading highlights mismatch between records. 
(C) Weight percent TOC for carbonaceous shale (solid black circles) and paleosol (hollow black circles). Light-gray rectangle indicates 
Paleocene-Eocene thermal maximum interval. VPDB—Vienna Peedee belemnite.
Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 9
maximum is 0.14 ± 0.07 (n = 56; 1s; Table DR2 
[see footnote 1]). Note, however, that this esti-
mate assigns all 13C-enrichment in paleosols to 
organic matter degradation, without consider-
ation of changes in allochthonous inputs.
Allochthonous Carbon Input
Autochthonous–allochthonous mixing lines 
were calculated assuming a constant and small 
input of allochthonous carbon to the total soil 
carbon (0.025 TOC units; Eq. 4; Fig. 5; Table 
DR3 [see footnote 1]). The autochthonous car-
bon component is d13CTT, the predicted total 
plant tissue carbon d13C value, and d13Calloch is 
the estimated allochthonous bulk organic car-
bon d13C value. We used the mean d13Corg value 
(–24.3‰) of 115 Mesozoic samples collected 
from 10 different formations (see Baczyn-
ski et al., 2013) for d13Calloch. The addition of 
allochthonous Mesozoic carbon to Paleogene 
soil carbon should cause d13Corg values to devi-
ate from the modeled logarithmic relationship. 
Moreover, a given flux of allochthonous organic 
carbon would have a larger impact on d13Corg 
values at lower TOC (higher 1/TOC) because 
it would comprise a greater proportion of total 
soil carbon. Supporting evidence for the input of 
allochthonous Mesozoic material into the Paleo-
gene sediments in this study area includes fos-
sil carbon in palynological preparations (Wing 
et al., 2005), Cretaceous-aged zircons, shark 
teeth, and marine dinoflagellates (Baczynski 
et al., 2013).
The proportion of allochthonous carbon was 
calculated using Equation 4, after accounting 
for soil organic matter degradation (see discus-
sion). Approximately 28%–63% of the total 
carbon would need to be allochthonous to pro-
duce the observed dampening and truncation of 
the paleosol d13Corg carbon isotope excursion at 
HW 16 (Fig. 6D). The allochthonous carbon 
estimate depends on the value used for d13CTT 
(Eq. 4). Figure DR1 (see footnote 1) shows how 
the value of d13CTT calculated from the differ-
ent n-alkane chain lengths affects estimates of 
Xalloch. With the exception of a higher degree of 
variability before the Paleocene-Eocene ther-
mal maximum, estimates of Xalloch are quite 
similar throughout most of the record. The 
variability seen before the Paleocene-Eocene 
thermal maxi mum suggests that using the aver-
age, rather than a single chain length, to predict 
d13CTT values is the more conservative approach.
Distribution of Shark Teeth
Shark teeth have been recovered from strata 
corresponding to the body of the Paleocene-
Eocene thermal maximum, and occasionally 
from overlying Willwood strata, throughout the 
southeastern Bighorn Basin field area. The pres-
ence of marine shark teeth in fluvial deposits 
among well-established freshwater and terres-
trial faunal assemblages constitutes strong evi-
dence for the input of reworked fossil material 
into Paleogene sediments (Eaton et al., 1989). 
These fossiliferous deposits typically occur in 
unconsolidated, poorly organized sandstones or 
mudstones composed of sandy clay. The teeth 
and their roots are frequently abraded and nearly 
always missing serrations and cusplets, consis-
tent with high-energy water transport. Most 
specimens are marine lamniform sharks such as 
Eostriatolamia and Scapanorhynchus (Fig. 7). 
Several specimens were identified as members 
of the family Squalidae. The roots of these teeth 
lack the infundibulum diagnostic of the genus 
Squalus or the typical shape of the genus Cen-
trophoroides and are tentatively assigned here 
to Centrosqualus (Fig. 7E). Given that  Squalus 
is the only genus among the Squalidae to cross 
the Cretaceous-Tertiary boundary and that 
marine sharks are a common component of Cre-
taceous assemblages throughout the Bighorn 
Basin (Case, 1987; Eaton et al., 1989; Oreska 
et al., 2013), a Cretaceous sedimentary origin is 
inferred for these specimens.
0 5 10 15 20 25 30 35
–32
–30
–28
–26
–24
–22
–20
–18
13
C o
rg
 (‰
, V
PD
B
)
1/TOC
M
E
S
13C
pre-PETM PETM post-PETM
carb shale
paleosol
pre-PETM degradation curve
PETM degradation curve
allochthonous
allocht
honou
s
mixing
mixing
5 10 15 20 25 30 351
1/Fraction Soil Organic Carbon Remaining (FSOC)
0 5 10 15 20 25 30 35
0
2
4
1/TOC
TO
C
Figure 5. Relationship between measured δ13Corg values and 1/TOC (lower x-axis) for car-
bonaceous shales (filled data points) and paleosols (hollow data points). Pre–Paleocene-
Eocene thermal maximum (PETM) samples are shown in blue circles, Paleocene-Eocene 
thermal maximum are shown in red circles, and recovery/post–Paleocene-Eocene ther-
mal maximum samples are shown in green triangles. The solid logarithmic curves display 
Rayleigh fractionation models for 13C enrichment during soil organic matter decomposi-
tion plotted versus 1/FSOC (upper x-axis). Models were calculated using average predicted 
 total plant tissue (δ13CTT) from pre/post–Paleocene-Eocene thermal maximum (–26.64‰) or 
Paleo cene-Eocene thermal maximum (–30.62‰) for initial δ13C values, with a fractionation 
factor (α) typical of microbial respiration (= 0.998). The dashed lines represent mixing of 
a constant amount of allochthonous carbon (mean δ13Calloch = –24.3‰) with autochthonous 
Paleogene carbon estimated using mean modeled pre/post–Paleocene-Eocene thermal maxi-
mum δ13CTT (–26.64‰) and Paleocene-Eocene thermal maximum δ13CTT (–30.62‰). Rela-
tionship between TOC and 1/TOC is shown below for reference. Range of nearby Mesozoic 
outcrop δ13Corg values is shown to right (shaded gray bar). VPDB—Vienna Peedee belemnite.
Baczynski et al.
10 Geological Society of America Bulletin, v. 1XX, no. XX/XX
DISCUSSION
Hypotheses for Soil Organic Matter 
13C Enrichment
The carbon isotope excursions recorded in 
bulk soil organic matter are smaller in magni-
tude and shorter in duration than those derived 
from n-alkanes from the same field area (Figs. 
2A and 2D). Organic matter in soils can be 
enriched in 13C relative to plant-derived organic 
matter as a result of degradation during pedo-
genic processes and/or the addition of trans-
ported 13C-enriched allochthonous carbon. Both 
of these processes could explain the truncation 
of the d13Corg carbon isotope excursion, the 
reduced overall magnitude of the carbon isotope 
excursion in d13Corg records relative to coeval 
plant lipids, soil carbonate, and tooth enamel 
records, and the high-frequency variability 
(1‰–2‰) between adjacent d13Corg samples 
within local stratigraphic sections (see Baczyn-
ski et al., 2013).
In soils, plant-derived organic matter is 
altered by heterotrophic decomposers, abi-
otic oxidation, and polymerization reactions 
(Balesdent and Mariotti, 1996). Each of these 
degradation pathways results in the release of 
carbon dioxide (under aerobic conditions) or 
methane (under anaerobic conditions), which 
are 13C-depleted with respect to the substrate 
and microbial biomass (Blair et al., 1985; Wynn, 
2007). Studies of depth profiles in modern soils 
reveal a strong logarithmic relationship between 
the fraction of soil organic carbon remaining and 
d13Corg ratios (Wynn, 2007; Wynn et al., 2005). 
Bulk soil organic matter d13C values increase as 
FSOC decreases, resulting from the cumulative 
effects of kinetic fractionation during microbial 
decomposition (Wynn, 2007).
An inverse relationship between d13Corg value 
and FSOC could also result from the mixing of 
isotopically distinct allochthonous and autoch-
thonous organic carbon. Studies have shown 
that the mixing of carbon from multiple sources 
(plant-derived organic matter and ancient 
rock carbon) is ubiquitous in contemporary 
fluvial systems (Blair et al., 2010; Brackley 
et al., 2010; Clark et al., 2013) and would be 
expected in ancient floodplain environments. 
If 13C-enriched allochthonous organic carbon 
were mixed with 13C-depleted autochthonous 
organic carbon during the Paleocene-Eocene 
thermal maximum, then d13Corg ratios would be 
13C-enriched relative to the predicted d13CTT val-
ues. Moreover, the same flux of allochthonous 
organic carbon would increase d13Corg values 
more at low TOC than at high TOC, producing 
an inverse linear relationship between TOC and 
d13Corg ratio.
pr
ed
ic
te
d 
δ1
3 C
TT
+d
eg
 
  p
re
di
ct
ed
 δ
13
C
TT
  
pr
ed
ic
te
d 
δ1
3 C
TT
+d
eg
 
pa
le
os
ol
 δ
13
C
or
g 
pr
ed
ic
te
d 
δ1
3 C
TT
+d
eg
  
ca
rb
 s
ha
le
 δ
13
C
or
g 
 
90 –2
0
–1
001020304050607080
SE Bighorn Basin Composite Section (m)
3
0
0.
5
1
1.
5
2
2.
5
20
0
5
10
15
δ1
3 C
 (‰
, V
P
D
B
)
δ1
3 C
 (‰
, V
P
D
B
)
δ1
3 C
 (‰
, V
P
D
B
)
P
er
ce
nt
 a
llo
ch
th
on
ou
s 
   
   
   
   
  c
ar
bo
n
S
ilt
/C
la
y 
R
at
io
Lo
ca
lit
ie
s 
w
/s
ha
rk
 te
et
h
(s
ha
de
d)
 o
ut
 o
f t
ot
al
 n
um
be
r
of
 fo
ss
il 
lo
ca
lit
ie
s
A
B
C
D
F
E
10
0
–3
3
–3
1
–2
9
–2
7
–2
5
–3
3
–3
1
–2
9
–2
7
–2
5
–3
3
–3
1
–2
9
–2
7
–2
5
–1
00
–5
0
0
50
Fi
gu
re
 6
. E
st
im
at
e o
f t
he
 re
la
tiv
e c
on
tr
ib
ut
io
ns
 o
f s
oi
l o
rg
an
ic
 m
at
te
r d
eg
ra
da
tio
n 
an
d 
al
lo
ch
th
on
ou
s c
ar
bo
n 
to
 δ
13
C
o
rg
 
13
C
 en
ri
ch
m
en
t. 
(A
) T
he
 a
ve
ra
ge
 p
re
di
ct
ed
 δ
13
C
T
T 
cu
rv
e 
(so
lid
 gr
ay
 sq
ua
re
s 
a
n
d 
lin
e) 
an
d t
he
 av
er
ag
e p
re
di
ct
ed
 δ
13
C
T
T 
re
co
rd
 m
od
ifi
ed
 fo
r 
de
gr
ad
at
io
n 
(δ
13
C
T
T+
de
g,
 h
ol
lo
w
 s
qu
ar
es
 a
n
d 
da
sh
ed
 g
ra
y 
lin
e),
 w
ith
 α
 
=
 
0.
99
8,
 p
re
/p
os
t–
Pa
le
oc
en
e-
Eo
ce
ne
 th
er
m
al
 m
ax
im
um
 F
SO
C
 
=
 0
.4
9,
 a
nd
 P
al
eo
ce
ne
-E
oc
en
e t
he
rm
al
 m
ax
im
um
 F
SO
C
 
=
 0
.2
5.
 (B
) C
om
pa
ris
on
 of
 th
e c
ar
b s
ha
le 
δ1
3 C
o
rg
 
re
co
rd
 (li
gh
t b
lue
) a
nd
 pr
e-
di
ct
ed
 δ
13
C
T
T+
de
g 
re
co
rd
 (d
as
he
d g
ra
y).
 D
ar
k-
gr
ay
 sh
ad
ing
 hi
gh
lig
ht
s m
ism
atc
h b
etw
ee
n r
ec
o
rd
s. 
(C
) C
om
pa
ris
on
 of
 th
e p
ale
os
ol 
δ1
3 C
o
rg
 
re
co
rd
 (r
ed
) a
nd
 pr
ed
ic
te
d 
δ1
3 C
T
T+
de
g 
re
co
rd
 (d
as
he
d g
ra
y).
 D
ar
k-
gr
ay
 sh
ad
ing
 hi
gh
lig
ht
s m
ism
atc
h b
etw
ee
n r
ec
o
rd
s. 
(D
) T
he
 p
ro
po
rt
io
n 
of
 a
llo
ch
th
on
ou
s c
ar
bo
n 
ne
ed
ed
 to
 ex
pl
ai
n 
th
e r
em
a
in
in
g 
13
C
 en
ri
ch
m
en
t 
in
 p
al
eo
so
ls 
a
fte
r 
a
cc
o
u
n
tin
g 
fo
r 
de
gr
ad
at
io
n.
 (E
) P
al
eo
so
l s
ilt
-to
-c
la
y 
ra
tio
 fo
r 
sit
e 
H
W
 
16
. (
F)
 N
um
be
r 
o
f v
er
te
br
at
e 
fo
ss
il 
lo
ca
lit
ie
s c
o
n
ta
in
in
g 
sh
ar
k 
te
et
h 
(sh
ad
ed
) o
u
t o
f 
to
ta
l n
um
be
r o
f f
os
sil
 lo
ca
lit
ie
s n
ea
r H
W
 
16
. L
ig
ht
-g
ra
y 
re
ct
an
gl
e 
in
di
ca
te
s P
al
eo
ce
ne
-E
oc
en
e 
th
er
m
al
 m
ax
im
um
 in
te
rv
al
. V
PD
B—
Vi
en
na
 P
ee
de
e 
be
le
m
ni
te
.
Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 11
The paleosol samples tend to have lower 
TOC than the carbonaceous shales (Fig. 4C) 
and would be expected to be more strongly 
influenced by both degradation and alloch-
thonous carbon input. As expected, paleosol 
d13Corg ratios are generally more 13C-enriched 
relative to predicted d13CTT values than the 
carbonaceous shale d13Corg values and become 
even more 13C-enriched in the upper body of 
the Paleocene-Eocene thermal maximum (~40–
53 m; Fig. 4B). Both paleosols and carbona-
ceous shales show greater 13C enrichment dur-
ing the Paleocene-Eocene thermal maximum 
than before and after. Examining the paleosol 
and carbonaceous shale samples separately 
allowed us to evaluate the relative influence 
of degradation and addition of allochthonous 
organic carbon on the carbon isotopic composi-
tion of bulk soil organic matter.
Estimating the Relative Contributions 
of Degradation and Allochthonous 
Carbon to 13C Enrichment
Most of the paleosol samples plot between 
the degradation curves and the allochthonous 
carbon mixing lines, suggesting that 13C enrich-
ment in the paleosols may have resulted from a 
combination of soil organic matter degradation 
and addition of 13C-enriched allochthonous car-
bon (Fig. 5). In contrast, the carbonaceous shale 
data plot along the logarithmic degradation 
curve (Fig. 5) and better match the overall shape 
of the n-alkane and predicted total plant tissue 
isotope curves (Fig. 2), suggesting that they 
contain predominantly autochthonous carbon 
with a negligible contribution of allochthonous 
carbon. This is consistent with the carbonaceous 
shales being collected from isolated lenticular 
channel fills where plant macrofossils and leaf 
wax n-alkanes were found. If we assume that 
there was no significant allochthonous carbon 
input to carbonaceous shales, we can use them 
to estimate the increase in soil organic matter 
degradation during the Paleocene-Eocene ther-
mal maximum, which can then be used to con-
strain extent of organic matter degradation in 
the paleosol samples. Constraining the extent of 
organic matter degradation then allows an esti-
mation of the increase in allochthonous carbon 
required to produce the additional 13C enrich-
ment observed in the paleosols.
Our calculations indicate that, on average, 
FSOC decreased by half during the Paleocene-
Eocene thermal maximum in the organic-
rich carbonaceous shales, suggesting that soil 
organic matter degradation doubled during the 
Paleocene-Eocene thermal maximum (Fig. 8A; 
Table DR2 [see footnote 1]). If we assume the 
same was true for paleosols and use the aver-
age pre–Paleocene-Eocene thermal maximum 
paleosol FSOC value, 0.49 (SD = 0.10, n = 52; 
Table DR2 [see footnote 1]), we predict a Paleo-
cene-Eocene thermal maximum FSOC value 
of 0.25 for the organic-poor paleosol samples 
(Fig. 8B).
To generate a new predicted isotope curve 
that accounts for soil organic matter degradation 
(d13CTT+deg), we used Equation 2 with d13CTT for 
d13Ci and the pre/post–Paleocene-Eocene ther-
mal maximum average FSOC (0.49) or the Paleo-
cene-Eocene thermal maximum estimated FSOC 
(0.25; Fig. 6A). This new predicted d13CTT+deg 
curve allowed us to estimate the 13C enrich-
ment that cannot be explained by soil organic 
matter degradation. We attributed the remain-
ing discrepancy (Figs. 6B–6C) to allochthonous 
carbon input and estimated the proportion of 
allochthonous carbon required to generate the 
measured paleosol d13Corg values using simple 
binary mixing between Paleogene and Meso-
zoic d13Corg end members. We calculated that 
~28%–63% of the total carbon would need to be 
allochthonous to produce the observed dampen-
ing and truncation of the paleosol d13Corg carbon 
isotope excursion at HW 16 (Fig. 6D; see also 
Figs. DR2–DR4 for other sites [see footnote 1]). 
This calculation represents a maximum estimate 
of allochthonous carbon input because it is pos-
sible that the paleosol organic matter experi-
A
B
C D
E
Figure 7. Representative shark specimens recovered from early Eocene, Paleocene-Eocene 
thermal maximum mammal-bearing strata in the southeastern Bighorn Basin, Wyoming. 
Scale bar = 1 cm. (A) UF 322398, Eostriatolamia sp., in mesial, lingual, labial, and distal views; 
(B) UF 322400, Eostriatolamia sp., in mesial, lingual, labial, and distal views; (C) UF 249590, 
Scapanorhynchus lewisii, in lingual and labial views; (D) UF 322399, posterolateral lower 
tooth, Scapanorhynchus sp., in mesial, lingual, labial, and distal views; and (E) UF 322397, 
Centrosqualus sp., in mesial, lingual, and labial views.
Baczynski et al.
12 Geological Society of America Bulletin, v. 1XX, no. XX/XX
enced greater 13C enrichment due to degrada-
tion on better-drained floodplains than in the 
backswamps and pond fills where carbonaceous 
shale was deposited.
The modeled increase in allochthonous 
carbon (of 28%–63%) during the Paleocene-
Eocene thermal maximum body at ~40 m (Fig. 
6D) coincides precisely with significant avul-
sion deposits and rapid sediment accumula-
tion, a relative increase in the abundance of 
small shark teeth reworked from local marine 
Cretaceous deposits (Fig. 6F), and higher silt-
to-clay ratios (Fig. 6E). Moreover, the doubling 
of the rate of degradation during the Paleocene-
Eocene thermal maximum is a reasonable esti-
mate supported by the Arrhenius equation and 
empirical data, which demonstrate that chemi-
cal and biological reaction rates tend to double 
for every 10 °C increase in temperature (Q10 = 
2; Bowen, 2013; IPCC et al., 2007). In addition 
to temperature, soil respiration rates have also 
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.90 1.0
–32
–28
–24
–20
–16
Fraction Soil Organic Carbon Remaining (FSOC)
13
C f
 (‰
, V
PD
B
)
Carbonaceous Shales
Avg FSOC
PETM 13CTT
non-PETM 13CTT
PETM =0.998
non-PETM =0.998
=0.999
=0.997
=0.997
=0.999
PETM FSOC = 50% of pre/post-PETM FSOC
A
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.90 1.0
–32
–28
–24
–20
–16
Fraction Soil Organic Carbon Remaining (FSOC)
13
C f
 (‰
, V
PD
B
)
Paleosols
non-PETM 13CTT
non-PETM =0.998
PETM =0.998
Avg FSOC
PETM 13CTT
Predicted FSOC
=0.999
=0.997
=0.997
=0.999
If PETM FSOC =
50% of pre/post-PETM FSOC
B
Figure 8. Logarithmic curves 
model the change in δ13C values 
as soil organic matter decom-
poses: FSOC decreases from 1, 
and the final δ13C ratio (δ13Cf) 
increases from the initial δ13C 
ratio of the total plant tissue 
(δ13CTT, Eq. 2). Non–Paleocene-
Eocene thermal maximum 
 decay curves are shown in blue 
(δ13CTT = –26.64‰), and Paleo-
cene-Eocene thermal maximum 
decay curves are shown in red 
(δ13CTT = –30.62‰). The shape 
of the logarithmic regression 
curve depends on the fraction-
ation factor, α. The solid curves 
represent α = 0.998, a value 
typical of microbial respira-
tion. The dashed curves depict 
a range of fractionation factors 
from 0.999 to 0.997. (A) Model 
for carbonaceous shales where 
“×” marks the estimated aver-
age pre/post–Paleocene- Eocene 
thermal maximum FSOC (0.76) 
and Paleocene-Eocene ther-
mal maximum FSOC (0.38), 
based on measured δ13Cf val-
ues. (B) Model for paleosols 
where “×” marks the average 
pre–Paleocene-Eocene thermal 
maximum FSOC value (0.49) 
that corresponds to measured 
δ13Cf values, and “○” marks 
the predicted Paleocene- Eocene 
thermal maximum FSOC value 
(0.25) and predicted δ13Cf value, 
if Paleocene-Eocene thermal 
maximum FSOC = 50% of pre/
post–Paleocene-Eocene thermal 
maximum FSOC.
Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 13
been shown to correlate with water availabil-
ity (IPCC et al., 2007). A 5–8 °C temperature 
increase and extreme seasonal precipitation dur-
ing the Paleocene-Eocene thermal maximum 
could have led to a doubling of soil respiration 
rate and the greater loss of soil carbon. It is also 
possible that the reduction in organic matter 
during the Paleocene-Eocene thermal maximum 
resulted from a decrease in productivity rather 
than an increase in degradation (or a combi-
nation of both). However, while a decrease in 
productivity would reduce the total amount of 
autochthonous organic carbon, it would not lead 
to 13C-enriched d13Corg values during the Paleo-
cene-Eocene thermal maximum. Therefore, an 
increase in degradation is the hypothesis pre-
sented here.
δ13Corg Variability in the Bighorn Basin
Paleocurrent studies of Willwood strata sug-
gest that the southern Bighorn Mountains were 
the most likely sediment source for Paleocene 
and Eocene deposits in the southern Bighorn 
Basin (Neasham and Vondra, 1972; Newbury, 
2011; Seeland, 1998). Locally carbonaceous, 
easily eroded, 13C-enriched Cretaceous marine 
shales comprise a major portion of Meso-
zoic strata in the southeastern Bighorn Basin. 
Allochthonous carbon sourced from these 
Cretaceous units could explain both the car-
bon isotope record and presence of Cretaceous 
zircons, marine dinoflagellates, and marine 
shark teeth. Shark teeth have been found in 
conjunction with Wa-0 fauna at all four Paleo-
cene-Eocene thermal maximum sections in the 
southeastern Bighorn Basin (Fig. 6F; see also 
Fig. DR5 [footnote 1]). At each of the Paleo-
cene-Eocene thermal maximum sections, shark 
teeth first appear less than 2 m above where bulk 
organic matter d13C values prematurely return 
to higher d13Corg values and are found through 
the remainder of the Paleocene-Eocene ther-
mal maximum body. Unlike most d13C records 
in which post-recovery d13Corg values remain 
slightly more 13C-depleted than background 
values, post-recovery d13C values at HW 16 
are more 13C-enriched than background values. 
The presence of shark teeth above the Paleo-
cene-Eocene thermal maximum interval at HW 
16 suggests that the d13Corg record was influ-
enced by reworked Cretaceous material even 
after the event. Indeed, petrographic analysis 
of different density organic matter fractions 
showed that the proportion of recycled organic 
carbon increased just prior to the carbon iso-
tope excursion and remained elevated through 
the remainder of the carbon isotope excursion 
body and into the recovery phase at HW 16 
(Bataille et al., 2013).
The bulk soil organic matter d13C records 
and the distribution of shark teeth suggest 
that climate change may have increased the 
amount of allochthonous carbon reworked dur-
ing the Paleocene-Eocene thermal maximum. 
Proxy records from the southeastern Bighorn 
Basin suggest that variability in the intensity 
of rainfall events may have increased dur-
ing the Paleocene-Eocene thermal maximum. 
Evidence for more episodic precipitation 
includes the greater number of crevasse splays 
and sandstone bodies (Kraus et al., 2013) and 
high silt-to-clay ratios (Fig. 6E) in the upper 
Paleocene-Eocene thermal maximum body. 
Changes in fluvial deposition in the Piceance 
Creek basin of western Colorado also suggest 
that flooding and increased sediment flux were 
common during the Paleocene-Eocene thermal 
maximum, perhaps indicating intense episodic 
rainfall events (Foreman et al., 2012). Bowen 
and Bowen (2008) proposed that a decreased 
meridional temperature gradient combined 
with an intensified monsoon may have directed 
seasonal precipitation toward the Bighorn 
Basin. An increase in extreme precipitation 
events could have led to deeper channeling and 
greater erosion/reworking of fossil carbon, par-
ticularly in the upper Paleocene-Eocene ther-
mal maximum body.
Unlike in the southeastern Bighorn Basin, 
d13Corg records from further northwest at Sand 
Creek Divide and Polecat Bench seem to more 
reliably record the full stratigraphic extent of 
the Paleocene-Eocene thermal maximum car-
bon isotope excursion (Magioncalda et al., 
2004; Rose et al., 2012). Spatial variability in 
sedimentation rate and allochthonous carbon 
source may explain the differences in d13Corg 
records across the Bighorn Basin. The southern 
Bighorn Mountains are composed of a crystal-
line core flanked by Paleozoic and Mesozoic 
strata (Love and Christiansen, 1985). Erosion of 
locally carbonaceous 13C-enriched Cretaceous 
marine shales, which comprise a significant 
portion of the strata flanking the basin, coupled 
with exceptionally low sedimentation rates in 
the southern Bighorn Basin could have led to 
an increase in both autochthonous organic car-
bon degradation and input of weathered alloch-
thonous material. These processes would both 
contribute to the dampening and truncation of 
the Paleocene-Eocene thermal maximum d13Corg 
records in the southeastern Bighorn Basin. In 
contrast, the northern Bighorn Basin sediments 
most likely derive from the Beartooth Moun-
tains, which are composed of harder and less-
organic-rich Paleozoic and crystalline rocks. 
Such source rocks would produce a lower flux 
of allochthonous organic carbon to sections like 
the one at Polecat Bench.
CONCLUSIONS
Comparison of bulk soil organic matter and 
n-alkane d13C ratios from Paleocene-Eocene 
thermal maximum sections in the southeastern 
Bighorn Basin, Wyoming, show that d13Corg 
records can significantly underestimate both 
the magnitude of the carbon isotope excursion 
and the thickness of the Paleocene-Eocene ther-
mal maximum body. A predicted total plant 
tissue carbon isotope record was modeled by 
applying lipid-specific enrichment factors to 
the n-alkane curves and then compared to the 
d13Corg records. Both the carbonaceous shale and 
paleosol samples were enriched in 13C relative to 
the predicted curve, but the paleosols were more 
enriched and failed to record the full extent of 
the excursion.
The discrepancy between the measured and 
predicted d13C ratios could result from 13C 
enrichment due to soil organic matter degra-
dation and/or the erosion of older rocks and 
redeposition of 13C-enriched allochthonous car-
bon on Paleocene-Eocene thermal maximum 
floodplains. The carbonaceous shale data were 
used to constrain the minimum amount of soil 
organic matter degradation and suggest that 
degradation rates doubled during the Paleocene-
Eocene thermal maximum. Paleosol 13C enrich-
ment due to soil organic matter degradation 
was estimated by assuming that degra dation 
rates doubled during the Paleocene-Eocene 
thermal maximum interval, and the remain-
ing paleosol 13C enrichment was attributed to 
allochthonous carbon input. A maximum of 
28%–63% of the soil organic carbon would 
need to be allochthonous in order to produce 
the observed 13C enrichment. The estimate of 
allochthonous carbon input increases abruptly 
in the middle of the Paleocene-Eocene thermal 
maximum, coinci dent with the first appearance 
of transported shark teeth in fossil assemblages. 
Organic matter degradation and allochthonous 
carbon input both lead to 13C enrichment of 
soil organic matter and can explain the reduced 
overall magnitude of the carbon isotope excur-
sion in d13Corg records relative to other carbon 
archives, the high-frequency 1‰–2‰ varia-
tions in d13Corg values within local stratigraphic 
sections, and the truncation of the carbon iso-
tope excursion body in bulk organic matter in 
the southeastern Bighorn Basin, Wyoming, 
paleosols. The significant influence of degrada-
tion and eroded fossil carbon on these organic 
matter d13C values demonstrates the importance 
of local environmental, biogeochemical, and 
sedimentological factors in floodplain d13Corg 
values and underscores the caution required 
when using these records to address global car-
bon cycle dynamics.
Baczynski et al.
14 Geological Society of America Bulletin, v. 1XX, no. XX/XX
ACKNOWLEDGMENTS
We thank Doug Boyer for his tremendous effort 
tracing beds and compiling geographic information 
system data throughout the field area over the past 
decade that have contributed to the stratigraphy that 
ties these data together. We are grateful to D. Walizer, 
A. Henderson, and H. Graham for instrument assis-
tance at Pennsylvania State University; J. Ehleringer, 
B. Tipple, M. Berke, and B. Hambach for instrument 
time and assistance in the SIRFER laboratory, Uni-
versity of Utah; R. Leder and A. Müller for their as-
sistance identifying shark specimens; and the many 
students and volunteers who helped with field work 
and laboratory work. Many thanks go to N. Blair and 
B. Sageman for comments and suggestions on early 
drafts of this manuscript. We also thank three anony-
mous reviewers and the editors for their constructive 
comments that greatly improved this manuscript. Ver-
tebrate fossils were collected under Bureau of Land 
Management permits to Bloch (PA04-WY-113, PA10-
WY-185). Funding was provided by National Sci-
ence Foundation awards EAR-0720268 (McInerney), 
EAR-0958717 (McInerney), EAR-0717892 (Wing), 
EAR-0718740 (Kraus), EAR-0640076 (Bloch, Krig-
baum, Secord), EAR-0719941 (Bloch), Initiative for 
Sustainability and Energy at Northwestern (McIner-
ney), and Australian Research Council FT110100793 
(McInerney) and DP13014314 (McInerney).
REFERENCES CITED
Baczynski, A.A., McInerney, F.A., Wing, S.L., Kraus, M.J., 
Bloch, J.I., Boyer, D.M., Secord, R., Morse, P.E., and 
Fricke, H.C., 2013, Chemostratigraphic implications 
of spatial variation in the Paleocene-Eocene thermal 
maximum carbon isotope excursion, SE Bighorn Ba-
sin, Wyoming: Geochemistry Geophysics Geosystems, 
v. 14, p. 4133–4152, doi: 10 .1002 /ggge .20265 .
Balesdent, J., and Mariotti, A., 1996, Measurement of soil 
organic matter turnover using 13C natural abundance, in 
Boutton, T.W., and Yamasaki, S., eds., Mass Spectrom-
etry of Soils: New York, Marcel-Dekker Inc., p. 83–111.
Bataille, C.P., Mastalerz, M., Tipple, B.J., and Bowen, G.J., 
2013, Influence of provenance and preservation on the 
carbon isotope variations of dispersed organic matter 
in ancient floodplain sediments: Geochemistry Geo-
physics Geosystems, v. 14, p. 4874–4891, doi: 10 .1002 
/ggge .20294 .
Bi, X.H., Sheng, G.Y., Liu, X.H., Li, C., and Fu, J.M., 
2005, Molecular and carbon and hydrogen isotopic 
composition of n-alkanes in plant leaf waxes: Organic 
Geochemistry, v. 36, p. 1405–1417, doi: 10 .1016 /j 
.orggeochem .2005 .06 .001 .
Blair, N.E., Leu, A., Muñoz, E., Olsen, J., Kwong, E., and 
Des Marais, D., 1985, Carbon isotopic fractionation in 
heterotrophic microbial metabolism: Applied and En-
vironmental Microbiology, v. 50, p. 996–1001.
Blair, N.E., Leithold, E.L., Brackley, H., Trustrum, N., Page, 
M., and Childress, L., 2010, Terrestrial sources and ex-
port of particulate organic carbon in the Waipaoa sedi-
mentary system: Problems, progress and processes: 
Marine Geology, v. 270, p. 108–118, doi: 10 .1016 /j 
.margeo .2009 .10 .016 .
Bowen, G.J., 2013, Up in smoke: A role for organic car-
bon feedbacks in Paleogene hyperthermals: Global 
and Planetary Change, v. 109, p. 18–29, doi: 10 .1016 /j 
.gloplacha .2013 .07 .001 .
Bowen, G.J., and Bowen, B.B., 2008, Mechanisms of PETM 
global change constrained by a new record from cen-
tral: Utah Geology, v. 36, p. 379–382, doi: 10 .1130 
/G24597A .1 .
Bowen, G.J., and Zachos, J.C., 2010, Rapid carbon seques-
tration at the termination of the Palaeocene-Eocene 
thermal maximum: Nature Geoscience, v. 3, p. 866–
869, doi: 10 .1038 /ngeo1014 .
Bowen, G.J., Bralower, T.J., Delaney, M.L., Dickens, 
G.R., Kelly, D.C., Koch, P.L., Kump, L.R., Meng, J., 
Sloan, L.C., Thomas, E., Wing, S.L., and Zachos, J.C., 
2006, Eocene hyperthermal event offers insight into 
greenhouse warming: Eos, Transactions, American 
Geophysical Union, v. 87, p. 165–169, doi: 10 .1029 
/2006EO170002 .
Bown, T.M., 1980, Summary of latest Cretaceous and Ceno-
zoic sedimentary, tectonic, and erosional events, Big-
horn Basin, Wyoming, in Gingerich, P.D., ed., Early 
Cenozoic Paleontology and Stratigraphy of the Big-
horn Basin, Wyoming: Ann Arbor, Michigan, Univer-
sity of Michigan, p. 25–32.
Brackley, H.L., Blair, N.E., Trustrum, N.A., Carter, L., 
Leithold, E.L., Canuel, E.A., Johnston, J.H., and Tate, 
K.R., 2010, Dispersal and transformation of organic 
carbon across an episodic, high sediment discharge 
continental margin, Waipaoa sedimentary system, New 
Zealand: Marine Geology, v. 270, p. 202–212, doi: 10 
.1016 /j .margeo .2009 .11 .001 .
Bray, E.E., and Evans, E.D., 1961, Distribution of normal-
paraffins as a clue to recognition of source beds: Geo-
chimica et Cosmochimica Acta, v. 22, p. 2–15, doi: 10 
.1016 /0016 -7037 (61)90069 -2 .
Bush, R.T., and McInerney, F.A., 2013, Leaf wax n-alkane 
distributions in and across modern plants: Implications 
for paleoecology and chemotaxonomy: Geochimica et 
Cosmochimica Acta, v. 117, p. 161–179, doi: 10 .1016 /j 
.gca .2013 .04 .016 .
Case, G.R., 1987, A new selachian fauna from the late 
Campanian of Wyoming (Teapot Sandstone Member, 
Mesaverde Formation, Big Horn Basin): Palaeonto-
graphica Abteilung A, v. 197, p. 1–37.
Chikaraishi, Y., and Naraoka, H., 2003, Compound-specific 
dD-d13C analyses of n-alkanes extracted from terres-
trial and aquatic plants: Phytochemistry, v. 63, p. 361–
371, doi: 10 .1016 /S0031 -9422 (02)00749 -5 .
Clark, K.E., Hilton, R.G., West, A.J., Malhi, Y., Gröcke, 
D.R., Bryant, C.L., Ascough, P.L., Robles Caceres, 
A., and New, M., 2013, New views on “old” carbon in 
the Amazon River: Insight from the source of organic 
carbon eroded from the Peruvian Andes: Geochemistry 
Geophysics Geosystems, v. 14, p. 1644–1659, doi: 10 
.1002 /ggge .20122 .
Collister, J.W., Rieley, G., Stern, B., Eglinton, G., and Fry, 
B., 1994, Compound-specific delta-C-13 analyses of 
leaf lipids from plants with differing carbon dioxide 
metabolisms: Organic Geochemistry, v. 21, p. 619–
627, doi: 10 .1016 /0146 -6380 (94)90008 -6 .
Diefendorf, A.F., Freeman, K.H., Wing, S.L., and Graham, 
H.V., 2011, Production of n-alkyl lipids in living plants 
and implications for the geologic past: Geochimica et 
Cosmochimica Acta, v. 75, p. 7472–7485, doi: 10 .1016 
/j .gca .2011 .09 .028 .
Diefendorf, A.F., Leslie, A.B., and Wing, S.L., 2015, Leaf 
wax composition and carbon isotopes vary among 
major conifer groups: Geochimica et Cosmochimica 
Acta, v. 170, p. 145–156, doi: 10 .1016 /j .gca .2015 .08 
.018 .
Diefendorf, A.F., Mueller, K.E., Wing, S.L., Koch, P.L., and 
Freeman, K.H., 2010, Global patterns in leaf C-13 dis-
crimination and implications for studies of past and 
future climate: Proceedings of the National Academy 
of Sciences of the United States of America, v. 107, 
p. 5738–5743, doi: 10 .1073 /pnas .0910513107 .
Eaton, J.G., Kirkland, J.I., and Doi, K., 1989, Evidence of 
reworked Cretaceous fossils and their bearing on the 
existence of Tertiary dinosaurs: Palaios, v. 4, p. 281–
286, doi: 10 .2307 /3514776 .
Finn, T.M., Kirschbaum, M.A., Roberts, S.B., Condon, S.M., 
Roberts, L.N.R., and Johnson, R.C., 2010, Cretaceous–
Tertiary Composite Total Petroleum System (503402), 
Bighorn Basin, Wyoming and Montana: U.S. Geologi-
cal Survey Digital Data Series DDS-69-V, 157 p.
Foreman, B.Z., Heller, P.L., and Clementz, M.T., 2012, Flu-
vial response to abrupt global warming at the Palaeo-
cene/Eocene boundary: Nature, v. 491, p. 92–95, doi: 
10 .1038 /nature11513 .
Handley, L., Pearson, P.N., McMillan, I.K., and Pancost, 
R.D., 2008, Large terrestrial and marine carbon and hy-
drogen isotope excursions in a new Paleocene/Eocene 
boundary section from Tanzania: Earth and Planetary 
Science Letters, v. 275, p. 17–25, doi: 10 .1016 /j .epsl 
.2008 .07 .030 .
IPCC, 2007: Climate Change 2007: The Physical Science 
Basis, in Solomon, S., D. Qin, M. Manning, Z. Chen, 
M. Marquis, K.B. Averyt, M. Tignor, and H.L. Miller, 
eds., Contribution of Working Group I to the Fourth 
Assessment Report of the Intergovernmental Panel on 
Climate Change: Cambridge, UK and New York, USA, 
Cambridge University Press, 996 p
Kennett, J.P., and Stott, L.D., 1991, Abrupt deep-sea warm-
ing, palaeoceanographic changes and benthic extinc-
tions at the end of the Palaeocene: Nature, v. 353, 
p. 225–229, doi: 10 .1038 /353225a0 .
Kraus, M.J., McInerney, F.A., Wing, S.L., Secord, R., Bac-
zynski, A.A., and Bloch, J.I., 2013, Paleohydrologic 
response to continental warming during the Paleocene-
Eocene thermal maximum, Bighorn Basin, Wyoming: 
Palaeogeography, Palaeoclimatology, Palaeoecology, 
v. 370, p. 196–208, doi: 10 .1016 /j .palaeo .2012 .12 .008 .
Love, J.D., and Christiansen, A.C., 1985, Geologic Map 
of Wyoming: Denver, U.S. Geological Survey, scale 
1:500,000.
Magioncalda, R., Dupuis, C., Smith, T., Steurbaut, E., and 
Gingerich, P.D., 2004, Paleocene-Eocene carbon iso-
tope excursion in organic carbon and pedogenic car-
bonate: Direct comparison in a continental stratigraphic 
section: Geology, v. 32, p. 553–556, doi: 10 .1130 
/G20476 .1 .
McInerney, F.A., and Wing, S.L., 2011, The Paleocene-
Eocene thermal maximum: A perturbation of carbon 
cycle, climate, and biosphere with implications for 
the future: Annual Review of Earth and Planetary Sci-
ences, v. 39, p. 489–516, doi: 10 .1146 /annurev -earth 
-040610 -133431 .
Neasham, J.W., and Vondra, C.F., 1972, Stratigraphy and 
petrology of Lower Eocene Willwood Formation, Big-
horn Basin, Wyoming: Geological Society of America 
Bulletin, v. 83, p. 2167–2180, doi: 10 .1130 /0016 -7606 
(1972)83 [2167: SAPOTL]2 .0 .CO;2 .
Newbury, S.S., 2011, The Sedimentary Response to Cli-
mate Change during the Paleocene-Eocene Thermal 
Maximum, Southeastern Bighorn Basin, Wyoming, 
USA [M.S. thesis]: Boulder, Colorado, University of 
Colorado, 91 p.
Oreska, M.P.J., Carrano, M.T., and Dzikiewicz, K.M., 2013, 
Vertebrate paleontology of the Cloverly Formation 
(Lower Cretaceous): I. Faunal composition, biogeo-
graphic relationships, and sampling: Journal of Ver-
tebrate Paleontology, v. 33, p. 264–292, doi: 10 .1080 
/02724634 .2012 .717567 .
Pedentchouk, N., Freeman, K.H., and Harris, N.B., 2006, 
Different response of delta D values of n-alkanes, iso-
prenoids, and kerogen during thermal maturation: Geo-
chimica et Cosmochimica Acta, v. 70, p. 2063–2072, 
doi: 10 .1016 /j .gca .2006 .01 .013 .
Polissar, P.J., and D’Andrea, W.J., 2014, Uncertainty in 
paleohydrologic reconstructions from molecular dD 
values: Geochimica et Cosmochimica Acta, v. 129, 
p. 146–156, doi: 10 .1016 /j .gca .2013 .12 .021 .
Rose, K.D., Chew, A.E., Dunn, R.H., Kraus, M.J., Fricke, 
H.C., and Zack, S.P., 2012, Earliest Eocene mamma-
lian fauna from the Paleocene-Eocene thermal maxi-
mum at Sand Creek Divide, southern Bighorn Basin, 
Wyoming: University of Michigan Papers on Paleon-
tology, v. 36, p. 1–122.
Schimmelmann, A., Sessions, A.L., and Mastalerz, M., 
2006, Hydrogen isotopic (D/H) composition of organic 
matter during diagenesis and thermal maturation: An-
nual Review of Earth and Planetary Sciences, v. 34, 
p. 501–533, doi: 10 .1146 /annurev .earth .34 .031405 
.125011 .
Schouten, S., Woltering, M., Rijpstra, W.I.C., Sluijs, A., 
Brinkhuis, H., and Damste, J.S.S., 2007, The Paleo-
cene-Eocene carbon isotope excursion in higher plant 
organic matter: Differential fractionation of angio-
sperms and conifers in the Arctic: Earth and Planetary 
Science Letters, v. 258, p. 581–592, doi: 10 .1016 /j .epsl 
.2007 .04 .024 .
Secord, R., Bloch, J.I., Chester, S.G.B., Boyer, D.M., Wood, 
A.R., Wing, S.L., Kraus, M.J., McInerney, F.A., and 
Krigbaum, J., 2012, Evolution of the earliest horses 
driven by climate change in the Paleocene-Eocene 
thermal maximum: Science, v. 335, p. 959–962, doi: 10 
.1126 /science .1213859 .
Distortion of carbon isotope excursion in bulk soil organic matter during the Paleocene-Eocene thermal maximum
 Geological Society of America Bulletin, v. 1XX, no. XX/XX 15
Seeland, D., 1998, Late Cretaceous, Paleocene, and early 
Eocene paleogeography of the Bighorn Basin and 
northwestern Wyoming, in Keefer, W.R., and Goolsby, 
J.E., eds., Cretaceous and Lower Tertiary Rocks of the 
Bighorn Basin, Wyoming and Montana, 49th Annual 
Field Conference Guidebook: Casper, Wyoming, Wyo-
ming Geological Association, p. 137–165.
Sessions, A.L., Sylva, S.P., Summons, R.E., and Hayes, 
J.M., 2004, Isotopic exchange of carbon-bound hydro-
gen over geologic timescales: Geochimica et Cosmo-
chimica Acta, v. 68, p. 1545–1559, doi: 10 .1016 /j .gca 
.2003 .06 .004 .
Sluijs, A., and Dickens, G.R., 2012, Assessing offsets be-
tween the delta C-13 of sedimentary components and 
the global exogenic carbon pool across early Paleogene 
carbon cycle perturbations: Global Biogeochemical 
Cycles, v. 26, p. 1–14.
Sluijs, A., Bowen, G.J., Brinkhuis, H., Lourens, L.J., and 
Thomas, E., 2007, The Palaeocene-Eocene thermal 
maximum super greenhouse: Biotic and geochemi-
cal signatures, age models and mechanisms of global 
change, in Williams, M., Haywood, A.M., Gregory, 
F.J., and Schmidt, D.N., eds., Deep-Time Perspec-
tives on Climate Change: Marrying the Signal from 
Computer Models and Biological Proxies: The Micro-
palaeontological Society Special Publication: The 
Geological [London] Society, p. 323–349.
Smith, F.A., Wing, S.L., and Freeman, K.H., 2007, Mag-
nitude of the carbon isotope excursion at the Paleo-
cene-Eocene thermal maximum: The role of plant 
community change: Earth and Planetary Science Let-
ters, v. 262, p. 50–65, doi: 10 .1016 /j .epsl .2007 .07 .021 .
Wing, S.L., 1980, Fossil floras and plant-bearing beds of 
the central Bighorn Basin, in Gingerich, P.D., ed., 
Early Cenozoic Paleontology and Stratigraphy of the 
Bighorn Basin, Wyoming: Ann Arbor, University of 
Michigan, Papers on Paleontology, v. 24, p. 119–125.
Wing, S.L., Harrington, G.J., Smith, F.A., Bloch, J.I., Boyer, 
D.M., and Freeman, K.H., 2005, Transient floral 
change and rapid global warming at the Paleocene-
Eocene boundary: Science, v. 310, p. 993–996, doi: 10 
.1126 /science .1116913 .
Wynn, J.G., 2007, Carbon isotope fractionation during de-
composition of organic matter in soils and paleosols: 
Implications for paleoecological interpretations of 
paleo sols: Palaeogeography, Palaeoclimatology, Palaeo-
ecology, v. 251, p. 437–448, doi: 10 .1016 /j .palaeo .2007 
.04 .009 .
Wynn, J.G., Bird, M.I., and Wong, V.N.L., 2005, Rayleigh 
distillation and the depth profile of C-13/C-12 ratios 
of soil organic carbon from soils of disparate texture 
in Iron Range National Park, Far North Queensland, 
Australia: Geochimica et Cosmochimica Acta, v. 69, 
p. 1961–1973, doi: 10 .1016 /j .gca .2004 .09 .003 .
Zachos, J.C., Pagani, M., Sloan, L., Thomas, E., and Billups, 
K., 2001, Trends, rhythms, and aberrations in global 
climate 65 Ma to present: Science, v. 292, p. 686–693, 
doi: 10 .1126 /science .1059412 .
Zachos, J.C., Wara, M.W., Bohaty, S., Delaney, M.L., 
 Petrizzo, M.R., Brill, A., Bralower, T.J., and Premoli-
Silva, I., 2003, A transient rise in tropical sea surface 
temperature during the Paleocene-Eocene thermal 
maximum: Science, v. 302, p. 1551–1554, doi: 10 .1126 
/science .1090110 .
Science editor: chriStian Koeberl 
aSSociate editor: GreG druSchel
ManuScript received 23 July 2015 
reviSed ManuScript received 21 January 2016 
ManuScript accepted 14 March 2016
Printed in the USA