SMITHSONIAN CONTRIBUTIONS TO THE EARTH SCIENCES • NUMBER 19

Mineral Sciences Investigations 1974–1975

Brian Mason

EDITOR

ABSTRACT

Mason, Brian, editor. Mineral Sciences Investigations 1974–1975. Smithsonian Contributions to the Earth Sciences, number 19, 125 pages, 48 figures, 37 tables, 1977.—Nine short contributions from the Smithsonian’s Department of Mineral Sciences for 1974 and 1975 are gathered together in this volume. These contributions comprise an account of the composition of garnet xenocrysts from three kimberlite pipes in Arizona and New Mexico; a catalog of major element chemistry of abyssal volcanic glasses, and the application of these data to determine magma compositions; descriptions of the Harleton (Texas), St. Mary’s County (Maryland), and Ras Tanura (Saudi Arabia) chondritic meteorites; a comparative study of eight chondrite meteorites from India and Pakistan; geochemical data on separated components of the Allende carbonaceous chondrite; and a mineralogical and chemical study of silicate inclusions in the El Taco mass of the Campo del Cielo iron meteorite.

OFFICIAL PUBLICATION DATE is handstamped in a limited number of initial copies and is recorded in the Institution’s annual report, Smithsonian Year. SERIES COVER DESIGN: Aerial view of Ulawun Volcano, New Britain.

Library of Congress Cataloging in Publication Data

Main entry under title:

Mineral sciences investigations, 1974–1975.

(Smithsonian contributions to the earth sciences; no. 19)

“Nine short contributions from the Smithsonian’s Department of Mineral Sciences.”

Includes bibliographies.

Supt. of Docs. no.: SI 1.26:19

1. Meteorites—Addresses, essays, lectures. 2. Petrology—Addresses, essays, lectures. 3. Mineralogy—Addresses, essays, lectures. I. Mason, Brian Harold, 1917– II. National Museum of Natural History. Dept. of Mineral Sciences. III. Series: Smithsonian Institution. Smithsonian contributions to the earth sciences; no. 19.

QE1.S227  no. 19  [QE395]  550′.8s  [549′.1]  76–608124

Contents

MINERALOGY

COMPOSITION OF GARNET XENOCRYSTS FROM THREE KIMBERLITE PIPES IN ARIZONA AND NEW MEXICO, by George S. Switzer

PETROLOGY

ABYSSAL BASALTIC GLASSES AS INDICATORS OF MAGMA COMPOSITIONS, by Gary R. Byerly, William G. Melson, Joseph A. Nelen, and Eugene Jarosewich

A CATALOG OF THE MAJOR ELEMENT CHEMISTRY OF ABYSSAL VOLCANIC GLASSES, by William G. Melson, Gary R. Byerly, Joseph A. Nelen, Timothy O’Hearn, Thomas L. Wright, and Tracy Vallier

METEORITES

THE HARLETON, TEXAS, CHONDRITE, by Roy S. Clarke, Jr., Edward P. Henderson, and Brian Mason

EIGHT L-GROUP CHONDRITES: A COMPARATIVE STUDY, by Ananda Dube, Becky J. Fredriksson, Eugene Jarosewich, Joseph A. Nelen, Albert F. Noonan, Jane O’Keefe, and Kurt Fredriksson

THE RAS TANURA, SAUDI ARABIA, CHONDRITE, by Brian Mason

GEOCHEMICAL DIFFERENCES AMONG COMPONENTS OF THE ALLENDE METEORITE, by Brian Mason and Philip M. Martin

THE ST. MARY’S COUNTY, MARYLAND, CHONDRITE, by Albert F. Noonan, Eugene Jarosewich, and Roy S. Clarke, Jr.

MINERALOGICAL AND CHEMICAL COMPOSITIONS OF SILICATE INCLUSIONS IN THE EL TACO, CAMPO DEL CIELO, IRON METEORITE, by Frank Wlotzka and Eugene Jarosewich

Mineral Sciences Investigations
1974–1975

Composition of Garnet Xenocrysts from Three Kimberlite Pipes in Arizona and New Mexico

George S. Switzer

ABSTRACT

Microprobe analyses were made of a large number of garnet xenocrysts from three kimberlite pipes in Arizona and New Mexico in a search for garnets having a composition similar to those found as inclusions in diamond. At none of the localities were any xenocrysts found having compositions lying within the diamond inclusion field in Ca-Cr, Ca-Mg-Fe, and Ca-Mg-Cr plots. From this one could conclude that none of the three kimberlite pipes examined is diamondiferous. If they were prospected on the basis of garnet composition, Garnet Ridge is the one most likely to contain diamond.

George S. Switzer, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560.

Introduction

Garnet xenocrysts in kimberlite may be derived from three types of xenoliths: from peridotites and eclogites of mantle origin, and disaggregated crustal xenoliths. Consequently, garnets from any one kimberlite pipe may vary widely in composition. This is reflected in the wide range of colors frequently seen in garnets from kimberlite. These colors may include red, orange, yellow, violet, purple, pink, brown, with all gradations between.

Of particular interest are the garnets of high pressure type considered to be of mantle origin, and especially those having a composition similar to garnets found as inclusions in diamond. Gurney and Switzer (1973) have shown that there are small compositional differences between garnets found in common garnet peridotite xenoliths in kimberlite, and those found as inclusions in diamond. These differences are that the diamond inclusions have higher magnesium and lower calcium content associated with high chromium. In their study of garnets from the Finsch diamond mine, Gurney and Switzer found some Mg-rich, Cr-rich, Ca-poor garnets that fell within the composition field of garnets previously only reported as inclusions in diamond. These garnets are considered to have a deeper provenance than the Mg-garnets commonly found as xenocrysts in kimberlite and in garnet peridotite xenoliths. Their presence in a kimberlite should be diagnostic of the presence of diamond, although it was not demonstrated whether or not there was any quantitative relationship. This discovery showed that N. V. Sobolev, et al. (1969) were correct when they suggested that kimberlite heavy mineral concentrates had not been adequately studied to reveal their full compositional range.

An extension of the work of Gurney and Switzer has been carried out by Lawless (1974). He found garnets with compositions similar to those found

TABLE 1.—Composition of garnet inclusions in diamond used to delineate the “diamond inclusion field”

as inclusions in diamond in the heavy mineral concentrates of the Bobbejan diamond mine, Kao kimberlite pipe, Koffiefontein diamond mine, Newlands kimberlite pipe, Premier diamond mine, and Sekretariskop kimberlite pipe. He found that diamond-inclusion-type garnets, although rare, are more plentiful than expected. All of the kimberlites studied by Lawless are diamond-bearing, ranging in diamond content from two carats per 100 metric tons to 100 carats per 100 tons. He was unable, however, to draw any conclusions as to whether or not the frequency with which the diamond-inclusion-type garnets occur is related to diamond content of the kimberlite, because to prove this would require the examination of an extremely large number of garnets.

The present study is a further extension of this work to two kimberlite occurrences in northeastern Arizona, Garnet Ridge and Buell Park, and to nearby Green Knobs in New Mexico. These pipes were selected because, although extensively collected at and well known for many years, no diamond finds have ever been reported. They have not been systematically prospected for diamonds, but it seems safe to assume that they are at least very diamond-poor, if not diamond-free. A specific search was made for diamonds at Buell Park by Schmitt, Swann, and Smith (1974:682), who report: “Exhaustive search through hundreds of anthills (much to the consternation of the ants) by the authors and other interested parties has revealed no diamonds.”

The three kimberlite intrusions studied are among eight scattered over a wide area in northeastern Arizona and nearby parts of Utah and New Mexico. These kimberlites have been described by many workers, including Gregory (1916, 1917); Reagan (1927); Balk and Sun (1954); Balk (1954); Malde (1954); Watson (1960, 1967); Malde and Thaden (1963); O’Hara and Mercy (1966); Gavasci and Kerr (1968); McGetchin (1968); Gavasci and Helmstaedt (1969); Watson and Morton (1969); McGetchin and Silver (1970); Stuart-Alexander, Shoemaker, and Moore (1971); and Schmitt, Swann, and Smith (1974).

Procedure

At kimberlite pipes that are being operated as diamond mines, a garnet sample is readily obtained from the heavy mineral concentrate from which the diamond is extracted. To obtain a large sample of garnet from localities such as those studied here would normally be difficult, for the garnet is very

FIGURE 1.—Location of the three kimberlite pipes studied.

FIGURE 2.—Location of the Garnet Ridge kimberlite pipes.

FIGURE 3.—The calcium content of Garnet Ridge garnets plotted against chromium content.
× garnets from Garnet Ridge      Δ garnet inclusions in diamond

sparsely distributed through the weathered kimberlite. Fortunately, the garnet has been substantially concentrated in numerous ant hills that dot the outcrops.

Samples of this natural concentrate were collected at the three localities. The garnet was further concentrated in the laboratory and sorted into color groups. Microprobe analyses were then made of 811 grains, for Ca, Mg, and Cr, using an ARL EMX microprobe. Most of these same grains were later analyzed on a nine channel ARL SEMQ microprobe. Operating conditions were 15KV, 0.15μA. Standards used were: for Si, Al, Mg, Ca-pyrope from Kakanui, New Zealand; for Ti-hornblende from Kakanui, New Zealand; for Cr-chromite from the Stillwater Complex, Montana. The complete analyses were corrected by the method of Bence and Albee.

Diamond Inclusion Field

The eclogitic garnets that have been found as inclusions in diamond are low in chromium content and, therefore, indistinguishable from garnet from

FIGURE 4.—Ca-Mg-Cr diagram.
× garnets from Garnet Ridge      Δ garnet inclusions in diamond

TABLE 2.—Analyses of Garnet Ridge garnets (from ant hills) (digits in parentheses identify grain number)

eclogite xenoliths. Of greater significance are the peridotitic garnets found as inclusions in diamond, characterized by high Mg, low Ca and Al, and high Cr. Analyses of these garnets (Table 1) from Gurney and Switzer (1973, table 4) and Lawless (1974, table 54), have been used to delineate the “diamond inclusion field” in the plots of garnet compositions for the three localities studied.

Garnet Ridge

Garnet Ridge was named for abundant garnet xenocrysts found in four kimberlite pipes that intrude Jurassic age rocks about 35 miles [56 km] northeast of Kayenta, Arizona (Figure 1). Three of the pipes lie in a cluster just north of U.S. Highway 164, about 10 miles [16 km] east of Dinnehotso (Figure 2). A fourth pipe lies at the ridge crest 2½ miles [4 km] west of the other three.

Garnet was collected at all four pipes, but since preliminary work failed to reveal any differences between them, the results were combined.

A total of 361 garnet xenocrysts were analyzed for Ca, Mg, and Cr, of which 250 complete analyses were made. Complete analyses for 31 are given in Table 2 to show the complete compositional range encountered. The highest chromium content (6.5% Cr_2O_3) was found in the dark lilac color group. All of the partial analyses are plotted on the Ca-Cr diagram, and the 250 complete analyses are plotted on the ternary diagrams (Figures 3–5).

GARNET MASSES.—Abundant in all of the Garnet Ridge pipes (but absent at Buell Park and Green Knobs) are pieces of brown, nearly monomineralic, polycrystalline garnet, sometimes of large size. Gregory (1916) reported finding pieces of this material

up to four inches [10 cm] in diameter. The writer found numerous fragments up to 2 cm in size, and in Pipe No. 2 a piece 11 × 7 × 5 cm weighing 515 grams. Texturally it is equigranular massive, with grain size 0.5–1 mm. Analyses of two specimens are given in Table 3. The first (column 1) contains minor rutile. The second (column 2) contains rutile and ilmenite.

XENOLITHS.—Xenoliths of two kinds are very abundant in the Garnet Ridge pipes: (1) breccia blocks of sedimentary rocks, some several hundred feet long, have subsided as much as 1200 feet into the pipes, and (2) pebble to boulder size crystalline rocks from more than 4000 feet below (Malde and Thaden, 1963).

The wide variety of garnetiferous xenoliths found in the Garnet Ridge pipes has been described by Malde and Thaden (1963), O’Hara and Mercy (1966), and Watson and Morton (1969). Of particular interest are the rare eclogite xenoliths. An unusually large example, found by the writer in Pipe No. 2, is a well-rounded nodule with dimensions 16 × 13 × 8 cm and weighing 2.55 kg. Analyses of the garnet from this, and from another smaller eclogite xenolith from the same pipe, are given in Table 3, columns 3 and 4. In columns 5–8 are given analyses of eclogite garnets from Garnet Ridge by O’Hara and Mercy (1966) and Watson and Morton (1969).

Buell Park

The Buell Park kimberlite pipe is about 30 km north of Fort Defiance, Arizona (Figure 6). This pipe, diameter 4.5 km, is apparently the largest in the world.

FIGURE 5.—Ca-Mg-Fe diagram.
× garnets from Garnet Ridge      Δ garnet inclusions in diamond

The Buell Park pipe now forms a basin underlain largely by kimberlite tuff breccia (Schmitt, Swann, and Smith, 1974). A ring-dike of olivine-bearing minette forms a ridge, known as Peridot Ridge, that curves across the southeast quadrant of the basin. It was from ant hills on the slopes of this ridge that the samples were collected for this study. Garnet-bearing xenoliths are rare and no eclogite xenoliths have been found (Schmitt, Swann, and Smith, 1974).

Of a total of 227 xenocrysts analyzed for Ca, Mg, and Cr, complete analyses were made of 179. Complete analyses for 37 of these are given in Table 4 to show the compositional range found in Buell Park garnets. The highest chromium content (5.1% Cr_2O_3) was found in xenocryst having a pink color. Thirty lilac color grains gave chromium contents ranging from 1.4% to 2.6%. All of the partial analyses are plotted on the Ca-Cr diagram, and the complete analyses are plotted on the ternary diagrams (Figures 7–9).

Garnet xenocrysts of varying color are present in kimberlite tuff matrix, and they have the same mixed origin as the detrital grains in the soil.

TABLE 3.—Analyses of garnet from Garnet Ridge xenoliths

Analyses of four such grains picked from a single hand-specimen of dark gray-green, friable kimberlite tuff are given in Table 5. The one with the highest chromium content is dark red in color.

Green Knobs

Green Knobs is a kimberlite pipe lying 6 km northeast of Buell Park in New Mexico, near the northern end of Red Lake (Figure 6). It has a diameter of approximately 0.8 km and contains only kimberlite tuff-breccia and abundant xenoliths (Schmitt, Swann, and Smith, 1974). Gneissic, gabbroic, and ultrabasic xenoliths are abundant, but none collected by the writer contained garnet. O’Hara and Mercy (1966) also report finding no garnetiferous xenoliths, other than one small fragment of clinopyroxene rock containing 2% of pale pink garnet.

A total of 236 garnet grains were analyzed for Ca, Mg, and Cr. Complete analyses were made of 174 of these same xenocrysts, of which 29 analyses are given in Table 6 to show the compositional range found. The highest chromium content, 5.9%

FIGURE 6.—Location of the Buell Park and Green Knobs kimberlite pipes.

FIGURE 7.—The calcium content of Buell Park garnets plotted against chromium.
× garnets from Buell Park       Δ garnet inclusions in diamond

FIGURE 8.—Ca-Mg-Cr diagram.
× garnets from Buell Park      Δ garnet inclusions in diamond

TABLE 4.—Analyses of Buell Park garnets (from ant hills) (digits in parentheses identify grain number)

FIGURE 9.—Ca-Mg-Fe diagram.
× garnets from Buell Park      Δ garnet inclusions in diamond

FIGURE 10.—The calcium content of Green Knobs garnets plotted against chromium content.
× garnets from Green Knobs      Δ garnet inclusions in diamond

TABLE 5.—Analyses of Buell Park garnets in Kimberlite tuff

Constituent

1

2

3

4

SiO_2

38.6

39.5

41.6

41.4

TiO_2

0.12

<0.1

<0.1

<0.1

Al_2O_3

21.5

22.2

23.4

21.5

Cr_2O_3

0.11

<0.1

0.43

4.1

FeO *

25.2

29.0

12.4

9.0

MnO

0.57

0.47

0.28

0.30

MgO

7.5

8.6

17.3

18.5

CaO

6.5

1.8

4.6

5.6

100.1

101.6

100.0

100.5

ATOMIC PROPORTIONS ON THE BASIS OF 12 OXYGENS

Si

2.98

3.01

3.00

2.97

Ti

0.01

0.00

0.00

0.00

Al

1.96

1.99

1.99

1.82

Cr

0.01

0.00

0.03

0.23

Fe

1.63

1.85

0.75

0.54

Mn

0.04

0.03

0.02

0.02

Mg

0.86

0.98

1.86

1.98

Ca

0.54

0.15

0.36

0.43

*Total iron calculated as FeO.

Cr_2O_3, was in a grain of dark red color. All of the partial analyses are plotted on the Ca-Cr diagram, and the complete analyses are plotted on the ternary diagrams (Figures 10–12).

As at Buell Park, garnet xenocrysts of several colors may be seen in the kimberlite tuff. Analyses of six garnet grains picked from a single hand-specimen are given in Table 7. The one with the highest chromium content, 2.5% Cr_2O_3, is red in color.

Discussion

In garnets from garnet periodotites the calcium and chromium contents increase sympathetically. Plots of calcium against chromium for garnet xenocrysts from Garnet Ridge, Buell Park, and Green Knobs (Figures 3, 7, 10) show that garnets from the peridotite suite predominate in these kimberlites. The points having between 4% and 10% calcium but little chromium, and therefore plotting close to the × axis, are eclogite garnets. At Garnet

TABLE 6.—Analyses of Green Knobs garnets (from ant hills)

TABLE 7.—Analyses of Green Knobs garnets in kimberlite tuff

TABLE 8.—Garnet from garnet peridotite xenoliths from South African kimberlite pipes (Gurney and Switzer, 1973, table 1)

FIGURE 11.—Ca-Mg-Cr diagram.
× garnets from Green Knobs       Δ garnet inclusions in diamond

Ridge and Green Knobs a few xenocrysts with between 1% and 2% calcium and very low chromium are high in ferrous iron and are almandines derived from basement igneous or metamorphic rocks.

At Garnet Ridge a few scattered points to the left of the garnet peridotites trend approach in composition the low calcium-high chromium content characteristic of diamond inclusion garnets. None fall well within the diamond inclusion field, however, and it is not possible to say that they have a deeper origin than the common garnet peridotite garnet.

Figure 13 is a Ca-Mg-Fe plot of garnets from three rock groups: (1) eclogite xenoliths from Garnet Ridge, (2) metamorphic rocks from Garnet Ridge, and (3) garnet peridotite xenoliths from South Africa kimberlites. The analyses are given in Tables 3 and 8.

Comparison of Figure 13 with the Ca-Mg-Fe plots of garnet xenocrysts from Garnet Ridge, Buell

FIGURE 12.—Ca-Mg-Fe diagram.
× garnets from Green Knobs       Δ garnet inclusions in diamond

Park, and Green Knobs shows again that most of the garnet in these kimberlites have a composition characteristic of garnets from garnet peridotite and were probably derived from disaggregated garnet peridotite xenoliths.

The provenance of the eclogite xenoliths at Garnet Ridge presents an interesting problem. Since they are found in kimberlite pipes, they would fall in group A of Coleman, et al. (1965). Instead, as pointed out by Watson and Morton (1969), the garnets in Garnet Ridge eclogites have a composition that places them in Coleman’s group C, suggesting that the Garnet Ridge xenoliths came from a glaucophane schist facies in the basement rocks. Other factors, however, discussed in detail by Watson and Morton, lead them to conclude that the Garnet Ridge eclogite has come from the deep crust or upper mantle. O’Hara and Mercy (1966) concluded that the Garnet Ridge eclogite xenoliths are extremely sodic and not comparable to pre-

FIGURE 13.—Ca-Mg-Fe diagram. (The analyses are given in Tables 3 and 8.)
• eclogite xenoliths from Garnet Ridge
○ other garnetiferous rocks from Garnet Ridge
□ garnet peridotite xenoliths from South African kimberlite pipes

viously described material from any source.

The large garnet masses found at Garnet Ridge have a composition similar to garnet from the eclogite xenoliths, indicating that they came from the same source, perhaps from garnet cumulates.

The Ca-Mg-Cr plot of Garnet Ridge xenocrysts (Figure 4) shows three analyses that lie on the border of the diamond inclusion field, but none within it. For Buell Park and Green Knobs there are no analyses in or near the diamond inclusion field.

It is not possible to draw any broad conclusions from these data for two reasons: (1) the number of xenocrysts analyzed from each locality is small, and (2) the true extent of the diamond inclusion field is not yet known because of the limited data available.

From this study one could conclude that none of the three kimberlite pipes examined is diamondiferous. If they were prospected on the basis of garnet composition, Garnet Ridge is the one most likely to contain diamond, while on the basis of size Buell Park is the most likely, since in any cluster of diamondiferous pipes the largest is also the most economic.

Literature Cited

Balk, R.

1954.    Kimberlitic Tuff Plugs in Northeastern Arizona. Transactions of the American Geophysical Union, 35:385.

Balk, R., and M. S. Sun

1954.    Petrographic Descriptions of Igneous Rocks. In J. E. Allen and Robert Balk, editors, Mineral Resources of Fort Defiance and Tohatchi Quadrangles, Arizona and New Mexico. New Mexico Bureau of Mines and Mineral Resources Bulletin, 36:100–118.

Coleman, R. G., D. E. Lee, L. B. Beatty, and W. W. Brannock

1965.    Eclogites and Eclogites: Their Differences and Similarities. Geological Society of America Bulletin, 76:482–508.

Gavasci, A. T., and H. Helmstaedt

1969.    A Pyroxene-rich Garnet Peridotite in an Ultramafic Breccia Dike at Moses Rock, Southeastern Utah. Journal of Geophysical Research, 74:6691–6695.

Gavasci, Anna T., and Paul F. Kerr

1968.    Uranium Emplacement at Garnet Ridge, Arizona. Economic Geology, 63:859–875.

Gregory, H. E.

1916.    Garnet Deposits of the Navajo Reservation, Arizona and Utah. Economic Geology, 11:223–230.

1917.    Geology of the Navajo Country: A Reconnaissance of Parts of Arizona, New Mexico and Utah. United States Geological Survey Professional Paper, 93:93–95, 102, 146–147.

Gurney, J. J., and G. S. Switzer

1973.    The Discovery of Garnets Closely Related to Diamonds in the Finsch Pipe, South Africa. Contributions to Mineralogy and Petrology, 39:103–116.

Lawless, P. J.

1974.    Some Aspects of the Geochemistry of Kimberlite Xenocrysts. Masters Thesis, University of Capetown.

Malde, H. E.

1954.    Serpentine Pipes at Garnet Ridge, Arizona. Science, 119:618.

Malde, H. E., and R. E. Thaden

1963.    Serpentine at Garnet Ridge. In Irving J. Witkind and Robert E. Thaden, editors, Geology and Uranium-Vanadium Deposits of the Monument Valley Area, Apache and Navajo Counties, Arizona. United States Geological Survey Bulletin, 1103:54–61.

McGetchin, T. R.

1968.    The Moses Rock Dike: Geology, Petrology and Mode of Emplacement of a Kimberlite-bearing Breccia Dike, San Juan Co., Utah. Ph.D. dissertation, California Institute of Technology, Pasadena.

McGetchin, T. R., and L. T. Silver

1970.    Compositional Relations from Kimberlite and Related Rocks in the Moses Rock Dike, San Juan County, Utah. American Mineralogist, 55:1738–1771.

Meyer, H. O. A., and F. R. Boyd

1972.    Composition and Origin of Crystalline Inclusions in Natural Diamonds. Geochimica et Cosmochimica Acta, 36:1255–1278.

Meyer, H. O. A., and D. P. Svizero

1975.    Mineral Inclusions in Brazilian Diamonds. Pages 785–795 in Ahrens, Dawson, Duncan, and Erlank, editors, Physics and Chemistry on the Earth. Volume 9. Oxford: Pergaman Press.

O’Hara, M. J., and E. L. P. Mercy

1966.    Eclogite, Peridotite and Pyrope from the Navajo Country, Arizona. American Mineralogist, 51:336–352.

Prinz, M., D. V. Manson, P. R. Hlava, and K. Keil

1975.    Inclusions in Diamonds: Garnet Lherzolite and Eclogite Assemblages. Pages 797–815 in Ahrens, Dawson, Duncan, and Erlank, editors, Physics and Chemistry of the Earth. Volume 9. Oxford: Pergaman Press.

Reagan, A. B.

1927.    Garnets in the Navajo Country. American Mineralogist, 12:414.

Schmitt, H. H., G. A. Swann, and D. Smith

1974.    The Buell Park Kimberlite Pipe, Northeastern Arizona. Pages 672–698 of part 2 in Karlstrom, Swann, and Eastwood, editors, Guide to the Geology of Northern Arizona. Boulder, Colorado: Geological Society of America.

Sobolev, N. V., M. A. Gnevushev, L. N. Mikhaylovskaya, S. I. Futergendler, Ye. I. Shemanina, Yu. G. Lavrent’yev, and L. N. Pospelova

1971.    Composition of Garnet and Pyroxene Inclusions in Ural Diamonds. Doklady Akademi Nauk S.S.S.R., 198.190–193.

Sobolev, N. V., Yu. G. Lavrent’yev, L. N. Pospelova and E. V. Sobolev.

1969.    Chrome Pyrope from Yakutian Diamonds. Doklady Academi Nauk S.S.S.R., 189:133–136.

Stuart-Alexander, D. E., E. M. Shoemaker, and H. J. Moore

1971.    Geologic Map of the Mule Ear Diatreme, San Juan Co., Utah. U.S. Geological Survey Miscellaneous Geologic Investigations Map I–674.

Watson, K. D.

1960.    Eclogite Inclusions in Serpentine Pipes at Garnet Ridge, Northeastern Arizona. Geological Society of America Bulletin, 71:2082–2083.

1967.    Kimberlite Pipes of Northeastern Arizona. Pages 261–269 in P. J. Wyllie, editor, Ultramafic and Related Rocks. New York: John Wiley and Sons, Inc.

Watson, K. D., and D. M. Morton

1969.    Eclogite Inclusions in Kimberlite Pipes at Garnet Ridge, Northeastern Arizona. American Mineralogist, 54:267–285.

Abyssal Basaltic Glasses as Indicators of Magma Compositions

Gary R. Byerly, William G. Melson,
Joseph A. Neleu, and Eugene Jarosewich

ABSTRACT

A review of the literature on alteration of abyssal basalts shows that often portions of the glassy rind are the most representative, or least altered, material present. Separating this glass from the palagonite and other alteration products provides the material closest in chemistry to the original magma chemistry. Use of the electron microprobe allows for analysis of only the freshest portions of the glass. Analyzing glasses by electron microprobe, with glass standards internally mounted in the sample discs, is shown to be a very precise technique. Examples of the homogeneity of dredged and cored abyssal basalts, based on glass analyses by electron microprobe, are presented.

Gary R. Byerly, William G. Melson, Joseph A. Nelen, and Eugene Jarosewich, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560.

Introduction

Pillow basalts are the most common igneous rock recovered from dredges or cores taken from the ocean depths. Chemical analyses of these basalts have played key roles in many petrogenetic and plate tectonic models and yet considerable controversy still exists over the degree to which these analyses reflect primary igneous processes rather than abyssal alteration processes. Many workers have investigated the variation in chemistry within single pillows (Paster, 1968; Corliss, 1970; Hart 1969; Melson, Thomson, and Van Andel, 1968; Muehlenbachs and Clayton, 1972; Matthews, 1971; Moore, 1965; Melson and Thompson, 1973; Shido, Miyashiro, and Ewing, 1974; and Hart, Erlank, and Kable, 1974; Melson, 1973; Thompson, 1973) or in abyssal basalts as a function of age (Hart, 1973; Hekinian, 1971). All workers agree that basalt is rapidly altered in the ocean-bottom environment but there is no general agreement on the mechanisms of alteration, the products formed, or the basaltic material that might be most representative of the primary chemistry. Many of these studies have concluded that alteration varies from a highly altered glassy rim to slight alteration in pillow cores, making the core the most popular candidate for chemical analysis to represent the primary chemistry. It is the theme of this paper that portions of the glassy rim often remain unaltered and when these portions can be analyzed separately they are the material most representative of the primary igneous chemistry. Since most abyssal basalts are sparsely phyric (usually less than 5 percent) the glass will be very close to the chemistry of the total rock. Late stage gravity or flow differentiation of phenocrysts are reflected in the whole rock chemistry and may confuse the primary partial melting or differentiation trends. Glass chemistry appears more homogeneous in single igneous events and apparently is much less effected by late stage igneous differentiation. We present a model for the alteration of pillow basalts and a technique for electron-microprobe analysis of basaltic glasses. Data is presented on the homogeneity of basaltic glass within a single thin section and within single eruptive flows.

Model for Alteration of Abyssal Pillow Basalts

Most basaltic pillows range in size from 10 cm to several meters in diameter and are characterized by a concentric zonal structure consisting of an outer glassy selvage (up to 2 cm in thickness), an intermediate variolitic zone (0.5 to 2 cm in thickness, predominately radiating intergrowths of submicron sized plagioclase and pyroxene), and a microcrystalline core. The degree and type of alteration varies considerably between these three zones.

Three stages of alteration are commonly recognized: a high temperature or deuteric stage beginning on extrusion, an intermediate temperature stage during cooling of the pillow, and the slow, low temperature stage of weathering and hydration. Each stage varies in alteration process, products formed, and zone of material most affected. Furthermore, it is commonly impossible to assign a given product to one of these three alteration stages. Deuteric processes lead to some chloritization of olivine, pyroxene, and interstital glass and probable oxidation of iron oxides in pillow cores. These alterations take place with late stage magmatic volatiles, but probably without addition of seawater (Paster, 1968; Corliss, 1970). Also important during this stage in the development of contraction joints and cracks that later admit seawater into the interior of the pillow.

The second stage of alteration begins as the pillow cools to below approximately 400° C. Sea-water begins circulating through cracks and joints where most of the alteration occurs. This results in formation of serpentine and iron oxides (Paster, 1968), and a variety of clay minerals, though it is difficult to determine the temperature at which these may have formed (Melson and Thompson, 1973).

The third stage of alteration involves the slow weathering process once the pillow has cooled down to normal ocean floor temperatures. Resultant products are hydra ted glass and further development of clay minerals. The thickness of the hydrated rim increases as a function of age of the pillow (Moore, 1966). Hydrated glass, or palagonite, is predominately a cryptocrystalline mixture of clay minerals such as smectite (Melson and Thompson, 1973; Melson, 1973). Several studies have shown that the process of hydration in the glass is a “surface reaction.” No diffusion gradients are apparent from microprobe traverses across glass-palagonite boundaries. Progressive hydration of the fresh glass proceeds into the pillow from the outer surface and along cracks and joints connecting the interior of the pillow allowing chemical exchange with sea water. Because erosion does not continuously expose a fresh surface for hydration this process is impeded by the crust of palagonite formed. The presence of empty vesicles, often lined with primary sulfides, near the palagonite glass interface is evidence of the barrier of the fresh glass to diffusion of water and partial hydration (Melson and Thompson, 1973). Only vesicles connected to the surface by fractures show alteration.

Corliss (1970) summarized a model for the structural and textural controls on alteration. To a very large extent the diffusion of material to and from reaction sites and exchange with sea water controls alteration. The glassy margin of pillow fragments prevents diffusion into the interior of the glassy rind and remains largely unaltered itself because of its lack of grain boundaries. Joints and cracks act as high diffusion conduits to supply sea water to the alteration reactions in the variolitic and crystalline portions of the pillow. Grain boundaries in these zones further aid in the distribution of material feeding alteration reactions and act as nucleation sites for deposition of alteration products.

Chemical variations caused by the above reactions are complex and, in spite of many studies, are not generally agreed upon. Several points, however, seem particularly noteworthy. Early deuteric processes probably result in oxidation of iron and perhaps small losses of iron and magnesium to joints and cracks (Paster, 1968). The intermediate stage of cooling results in an addition of water and potassium to the pillow interiors. Low temperature reactions add water, potassium, and iron, while depleting magnesium and calcium. Oxidation of iron also occurs. Using the above criteria, published analyses of glass-crystalline pairs from pillow fragments show an increase in apparent alteration of the crystalline material relative to glass (Paster, 1968, Corliss, 1970; Melson, 1973; Thompson, 1973; Shido, Miyashiro, and Ewing, 1974; and Hart, Erlank, and Kable, 1974). Muehlenbachs and Clayton (1972) present oxygen isotope data on glass-crystalline pairs which indicate less alteration of the glass.

This review indicates that glass, carefully separated from palagonite and other alteration products, provides the material closest in chemistry to the original liquid chemistry of abyssal pillow basalts.

Electron-Microprobe Analysis

An electron-microprobe technique is a rapid and precise technique capable of analyzing the freshest portions of glassy material that might otherwise be almost entirely altered. Small subsamples of glassy material are crushed to less than 1 mm diameter for microscopic examination. Those chips that appear freshest and display the finest glassy luster are mounted for electron microprobe analysis. The chips are mounted in 5 mm holes drilled around the edge of a 25 mm diameter by 6 mm thick lexan disc. Ten unknown glass chips are located around the edge of the disc with the glass standard in a central hole. Using the technique described by Moreland (1968), the discs are polished down to a 0.5 micron finish. The discs are carbon coated, using a Veeco V300 coater, modified by an attached timer. The process is automatically cut off after 1.25 seconds. The conditions were arrived at experimentally and appear adequate (Table 1).

Early analyses were performed with a three spectrometer ARL-EMX electron microprobe. Currently all glass analyses are made with an ARL-SEMQ electron microprobe. This instrument has six fixed spectrometers for Si, Al, Fe, Mg, Ca, and K, and three scanning spectrometers, which are tuned to Na, Ti, and P for basaltic glasses. Use of nine spectrometers enables rapid analysis of a spot, eliminating changing spectrometer positions during analyses. This feature is also particularly useful for samples that degrade in the beam. Sample points with low or high counts for any of the nine elements can be immediately spotted and a re-analysis made at an adjacent spot. A defocused beam of about 50 micron diameter at 15 kilovolts and 0.30 microamp beam current is used. The beam current is visually monitored at all times, using a digital display, to insure against short period instrument drift. The defocused beam helps decrease volatilization of light elements. Counting time on each spot is 10 seconds. Each analysis consists of an average of five spots.

Since the six fixed spectrometers do not lend themselves to detuning, some other means of measuring background has to be established. In our application a standard with average atomic number relatively close to that of the unknowns is selected for background determinations. Since that background is very low for major and some minor elements the error due to the background variation is insignificant. For very low concentration, however, care must be exercised when this approach is taken. For this study quartz was used for determination of background for Al, Fe, Mg, Ca, Ti, Na, and P, and corundum for Si. In Table 2 backgrounds and counts are compared between scanning and fixed spectrometer for potassium. The potassium content of abyssal basaltic glasses ranges widely and may reach very low concentrations. Results in Table 2 show that the average values on quartz and corundum determined above and below peak using scanning spectrometers agree well with those determined on the analytical x-ray spectral line and also with those determined on the fixed spectrometer. The difference is well within the statistical variation of the background counts when samples with close average atomic numbers are used.

In addition to the background correction the count data are further reduced by applying the

TABLE 1.—Effect of variable carbon coating time on probe results (see text for coating procedure)

TABLE 2.—Background for potassium from scanning and fixed spectrometers (15 kv, 0.30 μ amp beam current, and average of ten 10-second counts)

TABLE 3.—Analyses from poorly coated disc (only background and Bence-Albee corrections have been made)

TABLE 4.—Analyses from Table 3 have been normalized to the “best” chemical value of the internal standard VG–2

TABLE 5.—Analyses of the disc used in Tables 3 and 4 after cleaning and recoating (only background and Bence-Albee corrections have been made)

TABLE 6.—Analyses from Table 5 have been normalized to the “best” chemical value of the internal standard VG–2

Bence-Albee (1968) correction procedures. The empirical correction factors, used in this procedure, work especially well when standards with compositions close to the unknowns are used. VG–2, a deep sea basaltic glass (analyzed by E. Jarosewich), appears well suited as a standard for most ocean ridge basalts and glasses. It is used as the primary standard for Si, Al, Fe, Mg, Ca, Na, and Ti. Adularia and apatite standards are used respectively for K and P. For high-alkali and/or high-silica glasses, a Hawaiian tholeiite standard (VG-A99) or a rhyolite standard (VG–568) are used now as primary standards.

These refined data are then normalized by the following method: The VG–2 chip mounted in the center of each disk, as mentioned above, is analyzed after the ten unknown basaltic glasses. The probe value of this VG–2 chip is then compared to the wet chemical value and factors are established, if necessary, to correct the probe value of each element to the chemical value. The ten unknowns in the disk are then normalized by applying these factors. This procedure corrects for most operational errors such as long-time probe drift (poor current bleed-off, because of imperfect carbon coating or poor contact, etc.). VG–2 has been shown to be very homogeneous and is well suited for this purpose. Examples of the normalization procedure are given in Tables 3–6. In Table 3 a run was made on a disc with a bad carbon coating (porous due to poor surface adhesion). The oxide sums are systematically low. Normalizing the VG–2 analysis of that disc to its wet chemical value produces a set of normalization factors. Table 4 gives the normalized analyses and the factors used. Recoating the sample disc produced the analyses in Tables 5 and 6. If an analysis has been performed well then normalization is no more than 0.01% to 0.02%. Analyses that sum to less then 98.5 are generally deleted from the data base and set aside for reanalysis.

Melson, et al. (1976) compare the glass analyses of this laboratory to analyses of the same glasses at the MIT Laboratory (Frey, Bryan, and Thompson, 1974). The comparisons show systematic differences between the two laboratories. Total iron as FeO is somewhat higher, while Na_2O and MgO are lower than in the analyses reported by Frey, Bryan, and Thompson (1974). The Na_2O value of the standard glass VG–2 used by Frey, Bryan, and Thompson is too low by 0.15 wt. Use of this value in their correction procedures may be the explanation for some of these interlaboratory variations. Further comparisons should help resolve some of the systematic differences.

Homogeneity of Basaltic Glass

Microprobe analyses of abyssal basaltic glasses demonstrate their homogeneity on the thin section and outcrop scales. XRF analyses of similar bulk material shows considerably more inhomogeneity. To test the homogeneity of basaltic glass in a single pillow two thin sections were made of a Juan de Fuca Ridge sample (Melson, 1969) from palagonite and glass and from the variolitic zone. Microprobe analyses were made in eight areas of the two thin sections as shown in Table 7. The three areas within the glassy zone compare very favorably to one another. Comparing these to the glass within 100 microns of the palagonite interface shows only sodium deviating at the 95% confidence level. This value of sodium at the palagonite glass interface is just outside the 95% confidence level and should be studied further. Tentatively, this may be related to an incipient stage in the alteration of glass. Comparison of the variolite to glass chemistry shows a definite loss of iron and magnesium and a gain of sodium in the variolites relative to glass. There is no change in potassium so this is interpreted as a deuteric alteration of the variolitic zone.

TABLE 7.—Compositional variations within glassy and variolitic zones

TABLE 8.—Homogeneity of basaltic glasses (mean and 95% confidence interval (2 sd) on examples from text; numbers in parentheses indicate total number of analyses; each analysis represents five 10-second point counts, except for USNM 111230, where the 91 analyses are for single 10-second point counts)

Constituent

Standard glass, VG–2

(77)

USNM #111230

(91)

Galapagos DeSt. 5

(76)

DSDP Site 238

(21)

SiO_2

50.81 ± 0.70

50.30 ± 1.04

50.13 ± 0.88

49.47 ± 0.73

Al_2O_3

14.06 ± 0.30

15.68 ± 0.51

16.21 ± 0.54

15.14 ± 0.41

FeO

11.84 ± 0.34

10.57 ± 0.23

9.02 ± 0.38

11.15 ± 0.17

MgO

6.71 ± 0.22

7.50 ± 0.51

8.43 ± 0.37

7.58 ± 0.34

CaO

11.12 ± 0.30

11.07 ± 0.27

12.50 ± 0.37

12.61 ± 0.17

Na_2O

2.62 ± 0.08

3.04 ± 0.10

2.20 ± 0.08

2.45 ± 0.07

K_2O

0.19 ± 0.04

0.26 ± 0.05

0.06 ± 0.03

0.07 ± 0.03

TiO_2

1.85 ± 0.12

1.97 ± 0.11

1.03 + 0.12

1.22 ± 0.09

P_2O_5

0.17 ± 0.04

0.26 ± 0.06

0.07 ± 0.02

0.08 ± 0.02

Total

99.37

100.65

99.68

99.77

Analyses of 76 glassy pillow fragments from the Galapagos Ridge demonstrate the homogeneity of glasses within a single dredge haul. Table 8 gives the mean and two standard deviations for the glass analyses. Comparing the standard deviation for this dredge against that of VG–2 we find very close agreement. The increase in 2σ for Al, Mg, and Ca is partially due to the higher contents of these oxides in the dredged glasses. Similarly the lower 2σ for K and P are due to lower concentrations. The extreme homogeneity of the glasses from this dredge haul, which yielded several hundred pounds of glassy pillow basalt, is not unusual in the dredges that have been extensively analyzed by this laboratory (Melson, et al., 1976).

A second example of the homogeneity of glass in pillow basalts is demonstrated by variations in chemistry in a 70-meter interval of core from Leg 24, Site 238, DSDP (Melson, et al., 1976). Glass from each glassy horizon was analyzed and, along with XRF analyses (Dmitriev, 1974) from the same core, plotted in Figure 1 as a function of depth. The glasses are clearly much more homogeneous than the XRF analyses of the crystalline material. Fe, Mg, and, to a lesser extent, Ca are depleted in the crystalline material while K and Na are enriched. These chemical variations in the crystalline material and the concurrent homogeneity of the glasses again document the superiority of glass analyses.

In Table 8 the mean and standard deviations of the Galapagos dredge and DSDP Site 238 are compared to VG–2. The standard deviations are very close to being only a function of analytical error and differences in concentration. This indicates that these analyses reflect single eruptive events and that on a larger scale glass analyses could be very important in chemical stratigraphy. Melson, et al. (1976) present data from this laboratory indicating the extent to which the chemically homogeneous groups from one DSDP site or dredge location can be distinguished and compared to groups within or between locations. This technique seems quite promising because of the precision of the analyses used. A data file of chemically distinct groups of glass analyses is being built, which will be used in studying regional differences in basaltic chemistry.

Preliminary data on subaerial glasses indicate a homogeneity similar to that of abyssal basaltic glasses. The Hawaiian tholeiite glass standard (VG–A99) and the rhyolite glass standard (VG–568) used here have been checked extensively by probe for homogeneity. Two suites of fractionally crystallized Hawaiian tholeiite glasses are currently being studied at this lab. Using a least squares liquid line of descent to compare the glass analyses they

FIGURE 1.—Variation in glass and whole rock chemistry with depth in a 70-meter interval at DSDP Site 238.
• microprobe analyses of glass
XRF analyses of whole rocks

appear very homogeneous within the limits imposed by the precision of the microprobe technique (T. L. Wright of the U.S. Geological Survey, pers. comm.).

In summary, basaltic glasses often contain volumes which remain unaltered after much of the remaining rock has undergone some alteration. Analyzing these separately yields homogeneous results both within single pillow fragments and within large single eruptive events. An electron-microprobe technique has been presented that minimizes many sources of variations by the inclusion of internal standards with all unknowns. Using this rapid and precise technique much more information on the chemical diversity of the midocean ridges is being discovered.

Literature Cited

Bence, A. E., and A. L. Albee

1968.    Empirical Correction Factors for the Electron Microanalyses of Silicates and Oxides. Journal of Geology, 76:382–401.

Corliss, J. B.

1970.    Mid-Ocean Ridge Basalts, I: The Origin of Submarine Hydrothermal Solutions; II: Regional Diversity Along the Mid-Atlantic Ridge. Ph.D. dissertation, University of California, San Diego.

Dmitriev, Leonid V.

1974.    Petrochemical Study of the Basaltic Basement of the Mid-Indian Ridge: Leg 24, Djibouti to Mauritius. Pages 767–779 in L. Musich, editor, Initial Reports of the Deep-Sea Drilling Project, 24. San Diego: Scripps Institution of Oceanography.

Frey, F. A., W. B. Bryan, and G. Thompson

1974.    Atlantic Ocean Floor: Geochemistry and Petrology of Basalts from Legs 2 and 3 of the Deep-Sea Drilling Project. Journal of Geophysical Research, 79:5507–5527.

Hart, R.

1973.    A Model for Exchange in the Basalt-Seawater System of Oceanic Layer II. Canadian Journal of Earth Sciences, 10:799–816.

Hart, S. R.

1969.    K. Rb, Cs Contents and K/Rb, K/Cs, Ratios of Fresh and Altered Submarine Basalts. Earth and Planetary Science Letters, 6:295–303.

Hart, S. R., A. J. Erlank, and E. J. D. Kable

1974.    Sea Floor Basalt Alteration: Some Chemical and Sr Isotopic Effects. Contributions to Mineralogy and Petrology, 44:219–230.

Hekinian, R.

1971.    Chemical and Mineralogical Differences between Abyssal Hill Basalts and Ridge Tholeiites in the Eastern Pacific Ocean. Marine Geology, 11:77–91.

Matthews, D. H.

1971.    Altered Basalts from Swallow Bank, an Abyssal Hill in the NE Atlantic, and from a Nearby Seamount. Philosophical Transactions of the Royal Society of London, A268:551–571.

Melson, W. G.

1969.    Preliminary Results of a Geophysical Study of Portions of the Juan de Fuca Ridge and Blanco Fracture Zone. Coast and Geodetic Survey Technical Memorandum, 6.

1973.    Basaltic Glasses from the Deep Sea Drilling Project, Chemical Characteristics, Compositions of Alteration Products, and Fission Track “Ages.” Transactions of the American Geophysical Union, 54:1011–1014.

Melson, W. G., and G. Thompson

1973.    Glassy Abyssal Basalts, Atlantic Sea Floor near St. Paul’s Rocks: Petrography and Composition of Secondary Clay Minerals. Bulletin of the Geological Society of America, 84:703–716.

Melson, W. G., G. Thompson, and Tj. Van Andel

1968.    Volcanism and Metamorphism in the Mid-Atlantic Ridge, 22°N. Journal of Geophysical Research, 73:5925–5941.

Melson, W. G., T. L. Vallier, T. L. Wright, G. Byerly, and J. Nelen

1976.    Chemical Diversity of Abyssal Volcanic Glass Erupted along Pacific, Atlantic, and Indian Ocean Sea-Floor Spreading Centers. In G. H. Sutton, M. H. Manghnani, and R. Moberly, editors, Geophysics of the Pacific Ocean Basin and Its Margin. Washington, D.C.: American Geophysical Union.

Moore, J. G.

1965.    Petrology of Deep-Sea Basalt near Hawaii. Science, 263:40–52.

1966.    Rate of Palagonitization of Submarine Basalt Adjacent to Hawaii. United States Geological Survey Professional Paper, 550–D:D163–D171.

Moreland, G.

1968.    Preparation of Polished Thin Sections. American Mineralogist, 53:2070–2074.

Muehlenbachs, K., and R. N. Clayton

1972.    Oxygen Isotope Studies of Fresh and Weathered Submarine Basalts. Canadian Journal of Earth Sciences, 9:172–184.

Paster, T. P.

1968.    Petrologic Variations within Submarine Basalt Pillows of the South Pacific-Antarctic Ocean. Ph.D. dissertation, Florida State University, Tallahassee.

Shido, F., A. Miyashiro, and M. Ewing

1974.    Compositional Variation in Pillow Lavas from the Mid-Atlantic Ridge. Marine Geology, 16:177–190.

Thompson, G.

1973.    A Geochemical Study of the Low-Temperature Interaction of Sea Water and Oceanic Igneous Rocks. Transactions of the American Geophysical Union, 54:1015–1018.

A Catalog of the Major Element Chemistry
of Abyssal Volcanic Glasses

William G. Melson, Gary R. Byerly, Joseph A. Nelen,
Timothy O’Hearn, Thomas L, Wright, and Tracy Vallier

ABSTRACT

A data set is presented on the major element chemistry of abyssal volcanic glasses analyzed by electron microprobe in this laboratory. Information on the location, depth, age and tectonic setting is also included.

William G. Melson, Gary R. Byerly, Joseph A. Nelen, and Timothy O’Hearn, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560. Thomas L. Wright, U.S. Geological Survey, Reston, Virginia 22092. Tracy Vallier, Scripps Institution of Oceanography, La Jolla, California 92037.

Introduction

The importance of understanding the chemical variation of abyssal volcanic rocks has been emphasized in a number of recent reports (Engel, Engel, and Havens, 1965; Kay, Hubbard, and Gast, 1970; Cann, 1971; and Melson, et al., 1976). Many attempts have been made in the last dozen years to relate this chemical variation to regional differences in the upper mantle, differences in such tectonic parameters as spreading rates, or variation due to age of the sample and thus length of exposure to weathering and alteration. We present here a compilation of data on location, tectonic setting, depth, age, and major element chemistry for glassy abyssal volcanic rocks analyzed by electron microprobe in this laboratory. The analytical technique along with a discussion on the homogeneity of abyssal volcanic glasses is presented herein by Byerly, et al. Melson, et al. (1976) also discuss a portion of this data set and have tabulated averages based on a clustering procedure. We do not include at this time any interpretation of the data set, but will use it for a series of papers on regional variations and statistical evaluation of global patterns in abyssal volcanic rock chemistry. It is further hoped that other investigators will be able to use this data set for comparison to other data sets or statistical tests of newer petrogenetic models.

The majority of the samples included in this set are from dredges along the mid-oceanic ridge. A smaller number of samples are dredged from fracture zones, aseismic ridges, and seamounts. We have tried to include all glasses that were recovered in DSDP cores which penetrated sills or pillow basalts.

At present the geographical distribution of samples is heavily biased towards the Juan de Fuca and Galapagos regions, which are well represented in the collections of the National Museum of Natural History, Smithsonian Institution. The Indian Ocean spreading system is poorly represented. Long segments of many spreading systems are not covered. The data set continues to grow as we receive samples from other investigators and samples from new areas are recovered through dredging and the new International Program for Ocean Drilling.

The catalog[*] is arranged by ocean, and within ocean by latitude of the sample location. For DSDP cores the arrangement is stratigraphic from top to bottom of the hole. In dredges the samples are arranged in order of increasing FeO* (= total Fe calculated as FeO) content to aid in rapid scanning of the data set.

ACKNOWLEDGMENTS.—We are deeply indebted to the numerous institutions and individuals who contributed samples for this study and who are indicated in the “Key to the Catalog.” Mr. Eugene Jarosewich contributed substantially by providing classical analyses of basaltic glasses that were used as standards.

Literature Cited

Cann, J. R.

1971.    Major Element Variations in Ocean-Floor Basalts. Philosophical Transactions of the Royal Society of London, A268:495–505.

Engel, A. E., C. G. Engel, and R. G. Havens

1965.    Chemical Characteristics of Oceanic Basalts and Upper Mantle. Bulletin of the Geological Society of America, 76:719–734.

Kay, R., N. J. Hubbard, and P. W. Gast

1970.    Chemical Characteristics and Origins of Oceanic Ridge Volcanic Rocks. Journal of Geophysical Research, 75:1585–1613.

Melson, W. G., T. L. Vallier, T. L. Wright, G. R. Byerly, and J. Nelen

1976.    Chemical Diversity of Abyssal Volcanic Glass Erupted Along Pacific, Atlantic, and Indian Ocean Sea-Floor Spreading Centers. In G. H. Sutton, M. H. Manghnani, and R. Moberly, editors, Geophysics of the Pacific Ocean Basin and Its Margin. Washington, D.C.: American Geophysical Union.

Key to the Catalog

LAT.

Latitude of sample

LONG.

Longitude of sample

DEPTH

Water depth in meters; in the case of dredges, the greatest depth

AGE

Probable age (millions of years) of sample: M = magnetic anomaly age, S = overlying sediment age, <10 = samples less than 10 million years old, that is, younger than magnetic anomaly five

S

Probable site of eruption of intrusion: F = fracture zone, M = marginal basin, N = aseismic ridge, R = spreading rise or ridge

T

Type of sample: digits = Deep Sea Drilling Project site number, D = dredge, C = core, F = field

E

Probable origin: E = eruptive, I = instrusive

FeO*

Total iron calculated as ferrous oxide

VG–#

Sample number in data file of Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution

SO

Source of sample:

AM

=  Dr. Fabrizio Aumento, Dalhousie University

AN

=  Dr. Tj. van Andel, Oregon State University

BR

=  Dr. Wilfred Bryan, Woods Hole Ocranographie Institution

BW

=  Dr. V. T. Bowen, Woods Hole Oceanographic Institution

CL

=  Dr. David Clague, Scripps Institution of Oceanography

CM

=  Dr. John Campsie, Ocean Study Group, Copenhagen

CN

=  Dr. J. R. Cann, University of East Anglia

CR

=  Dr. Jack Corliss, Oregon State University

DP

=  Deep Sea Drilling Project of the Scripps Institution of Oceanography

JH

=  Dr. Leonard Johnson, U.S. Naval Oceanographic Office

LM

=  Lamont-Doherty Geological Observatory

MY

=  Dr. Akiho Miyashiro, SUNY Albany

SC

=  Scripps Institution of Oceanography (obtained by Dr. Tracy Vallier and Mr. Richard Nishimori)

SH

=  Dr. Anatol Sharaskin, Vernadsky Institution of Geochemistry and Analytical Chemistry

TH

=  Dr. Geoffrey Thompson, Woods Hole Oceanographic Institution

US

=  National Museum of Natural History, Smithsonian Institution (formerly of the United States National Museum)

VG

=  Dr. Peter Vogt, U. S. Naval Oceanographic Office

The Harleton, Texas, Chondrite

Roy S. Clarke, Jr., Edward P. Henderson, and Brian Mason

ABSTRACT

The Harleton meteorite fell near Harleton, Texas (32°40.5’N, 94°30.7’W) on 30 May 1961. A single stone, weighing 8.36 kg, was recovered. Chemical analyses and microscopic examination show that it is an olivine-hypersthene chondrite (L6). Principal minerals are olivine and orthopyroxene, with minor amounts of plagioclase, nickel-iron, and troilite, and accessory chromite.

Roy S. Clarke, Jr., Edward P. Henderson, and Brian Mason, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560.

Introduction

The Harleton meteorite fell four miles [6 km] east of Harleton, Harrison County, Texas, at 2230 (Central Time) on 30 May 1961, on the property of J. J. and Bernice Craver; the coordinates of the place of fall are 32°40.5’N, 94°30.7’W. A single stone, weighing 8.36 kg, was recovered from a shallow crater in sandy ground within ten minutes of the time of fall. An examination of that specimen shows that other pieces were shed from it during high velocity flight, but none was recovered. The fireball accompanying the fall was seen over a considerable area of east Texas; the meteorite was traveling approximately west to east.

The Harleton meteorite was purchased from the property owners by E. P. Henderson on 12 June 1961. Dr. Henderson represented both the Smithsonian Institution and an informal consortium of interested institutions in this difficult negotiation. The meteorite arrived at the National Museum of Natural History in Washington on 13 June and a mold for subsequent preparation of models was made during the day of 14 June. Late that same day the Harleton meteorite was broken in half in a jaw crusher. One of the halves was broken down into smaller pieces using the same technique. Most of these fragments were distributed to scientific colleagues for detailed studies on 15 June, just 16 days after the fall. Details of the various transactions involved are in Smithsonian accession file 236677, and the meteorite has been assigned catalog number USNM 2576.

The prompt recovery and distribution of the Harleton meteorite made it a particularly interesting specimen for the study of phenomena associated with freshly fallen meteorites. Published reports that discuss these results are listed here: Bainbridge, Suess, and Wänke, 1962; Davis, Stoenner, and Schaeffer, 1963; Eberhardt, et al., 1966; Fireman, 1967; Fireman, DeFelice, and Tilles, 1963; Goel and Kohman, 1962; Honda and Arnold, 1964; Kohman and Goel, 1963; Mabuchi, et al., 1968; Osborn, 1972; Rowe, Van Dilla, and Anderson, 1963; Shedlovsky, Cressy, and Kohman, 1967; and Suess and Wänke, 1962.

ACKNOWLEDGMENTS.—We are indebted to a group of scientific colleagues for providing funding to help in the purchase of the Harleton meteorite.

Morphology

The stone is roughly cuboidal in form, 19 × 14 × 13 cm in maximum dimensions. The morphology shows several significant features resulting from ablation and spallation during its fall through the atmosphere. These are shown in a series of pictures taken of its various sides.

1. One side is quite flat and has a square outline (Figure 1).

2. The fusion crust that covers the flat side has different markings than the crust on the other surfaces (Figure 2), yet its texture and thickness are essentially the same as on the other sides.

3. Around the edges of this rear square face (Figure 1) there is a slight ridge, which is due to the accumulation of overturned stringers of surface material. Molten material was pushed backwards on the side faces by the strong lateral flow, and part of this was deposited as delicate stringers of glass by the doughnut-shaped vortex trailing the rear face. This vortex carried material outward from the edges with the side faces, depositing it on the rear face and moving it toward the edges. Thus at the edges there was a mingling of two forces, the stronger lateral flow off the sides of the meteorite

FIGURE 1.—Rear surface of the Harleton meteorite; maximum width of specimen, 14 cm.

FIGURE 2.—Side view of Harleton meteorite, showing flow lines on fusion crust; maximum length of specimen, 21 cm.

and the slower-moving vortex pushing the molten material from the rear face to the edge. Where these forces met, a ridge formed around the rear face of a falling orientated body.

4. The square face (Figure 1) is to some degree ablated but less so than the other faces, and it has totally different flight markings.

5. Since this side has a square outline it represents

FIGURE 3.—Side view of Harleton meteorite, showing cuboidal form; maximum length of specimen, 19 cm.

a break in the main mass late in the high-velocity flight, but while the meteorite was still falling fast enough to form fusion crust. The sharp edges around this face have been only slightly rounded, indicating a brief period of ablation.

6. This stone not only has an unusual flat square face but also has other surfaces which are almost at right angles to it (Figure 3). Thus there are three rectilinear directions represented in the surfaces of the meteorite.

7. The shape of this mass, therefore, shows that the meteorite broke up during high-velocity flight. After each break the object adjusted its position so that its center of gravity and the stagnation point on its front face were on the line of fall of the meteorite. The weight of the object decreased by spallation and ablation, and its velocity was decreased to free-fall velocity by atmospheric friction.

The morphology of this meteorite shows that other fragments must have fallen as individual stones along its trajectory. None was recovered, however, probably because the brushy terrain made search and sighting difficult.

Structure and Mineralogy

When viewed with the naked eye, a broken surface of the meteorite appears remarkably homogeneous, the material being compact and pale gray. With a hand lens brightly reflecting particles of nickel-iron and troilite are seen dispersed

FIGURE 4.—Photomicrograph (reflected light) of a polished surface of the Harleton meteorite; metal particles are white, troilite (and chromite) pale gray, silicates dark gray. Greatest dimension, 20 mm.

throughout, and chondrules are few and poorly defined. These features are demonstrated by a low-power photomicrograph of a polished surface (Figure 4). An unusual feature of broken surfaces of this meteorite is the complete absence of any rusting around the metal particles, despite exposure to the atmosphere for many years; this implies the absence of lawrencite. Only two thin sulfide veins were observed on a 170 cm^2 broken surface. One crossed the specimen for 14 cm, while the other was only 3 cm long.

Thin sections under the microscope show a granular aggregate of olivine, orthopyroxene, and opaque minerals (Figure 5). The original chondritic texture has almost disappeared. Chondrules are recognizable by their vaguely circular form and their usually coarser grain size than the surrounding matrix. The fusion crust shows three well-developed zones. The outermost zone is very thin, averaging about 0.05 mm, and consists of black opaque glass, the opacity appearing to be largely due to finely dispersed magnitite or magnesioferrite. The next zone averages 0.1 mm, is transparent, and consists largely of granular olivine and pyroxene like that in the body of the meteorite; it also contains occasional pools of isotropic, fused plagioclase. The innermost zone is considerably thicker, averaging 0.3 mm; it is black and opaque in transmitted light, but in reflected light it is seen to consist of the same minerals as the body of the meteorite, but impregnated with a network of minute troilite veinlets following grain boundaries, cleavages, and fractures.

Reflected light examination of polished surfaces revealed the presence of troilite, kamacite, taenite, chromite, and trace amounts of copper. Troilite is mainly present in isolated coarse-grained patches, but also occurs in association with kamacite and/or taenite. A variety of metal structures is seen on etched surfaces. Large clear kamacite areas are present, as are polycrystalline kamacite masses. Some kamacite areas contain silicate inclusions and several contain small troilites. Taenite inclusions generally are smaller than kamacite and have complex structures. They appear to be pearlitic and occasionally include small grains of troilite. Small areas of copper were observed twice within kamacite and in association with small grains of troilite, and once at a grain boundary within a large area of troilite. Several areas that suggest shock melting of troilite or kamacite were observed, but these are comparatively rare. The overall impression is one of a compact rock with clean and well-defined crystal boundaries.

As mentioned above, the principal minerals are olivine and pyroxene. Other characteristic minerals are nickel-iron, plagioclase, feldspar, and troilite. Minor minerals include chromite and a phosphate mineral (apatite and/or whitlockite) and copper is present as a trace mineral. Ramdohr (1973:89) has recorded the following opaque minerals: kamacite, taenite, plessite, copper, troilite, chalcopyrrhotite (trace), mackinawite, and chromite. Our notes on some of the minerals follow.

Olivine: The refractive indices are α = 1.682, γ = 1.720, indicating a content of 25 mole percent of the Fe_2SiO_4 component, according to the determinative curve of Poldervaart (1950). Microprobe analyses show 24.6 mole percent Fe_2SiO_4 in the olivine. The chemical analysis of the acid-soluble nonmagnetic fraction (Table 1), essentially olivine with troilite and a little phosphate, indicates 26 mole percent Fe_2SiO_4.

Orthopyroxene: The refractive indices are α = 1.678, γ = 1.689, indicating a content of 21 mole percent of the FeSiO_3 component, according to the determinative curve of Kuno (1954). Microprobe analyses show the Fe/Fe + Mg mole percentage of the orthopyroxene to be 22.4, with a content of 0.85% CaO in this material. The chemical analysis of the acid-insoluble nonmagnetic fraction (Table 1), essentially orthopyroxene and plagioclase, shows an Fe/Fe+Mg mole percentage of 23. In terms of the conventional subdivision of meteoritic pyroxene, this falls in the compositional range of hypersthene.

Diopside: An x-ray diffractogram of the acid-insoluble fraction shows a low-intensity peak corresponding to diopside. This mineral was not certainly identified microscopically or with the microprobe, although some turbid and fine-grained material is probably clinopyroxene.

Plagioclase: The plagioclase feldspar is finegrained, granular, untwinned, and interstitial to the other minerals. Microprobe analyses give a composition of An_9, with a K_2O content of 1.11%.

Chromite: Bunch, Keil, and Snetsinger (1967, table 2) give the following microprobe analysis of Harleton chromite: Cr_2O_3, 56.0; Al_2O_3, 5.1; V_2O_3, 0.73; TiO_2, 2.88; FeO, 33.4; MgO, 2.11; MnO, 0.57.

FIGURE 5.—Photomicrograph (transmitted light) of a thin section of the Harleton meteorite, showing fusion crust on left; width of section, 1.8 mm.

Chemical Composition

Fragments of the meteorite free of fusion crust were selected for chemical analysis (Table 1). They were crushed to a fine powder containing a few residual larger particles of metal. This material was separated into magnetic and nonmagnetic fractions, which were individually leached with 20% HCl in a nitrogen atmosphere. Sample solutions containing the HCl-soluble magnetic fraction (I) and the HCl-soluble nonmagnetic fraction (II) re-

TABLE 1.—Chemical analysis of the Harleton meteorite

Constituent

I

II

III

Composite

Fe (metal)

72.4

-

-

7.28

Ni

11.3

<0.1

<0.1

1.14

Co

0.61

<0.01

-

0.06

FeS

2.99

11.84

-

6.58*

SiO_2

4.57

31.84

57.51

39.85

TiO_2

-

0.02

0.21

0.10

Al_2O_3

<0.1

<0.1

5.10

2.14

Cr_2O_3

-

-

-

0.52*

FeO

3.2

21.3

10.5

14.3 **

MnO

<0.1

0.39

0.34

0.33

MgO

5.00

33.35

19.82

24.77

CaO

<0.2

0.59

3.65

1.81

Na_2O

-

-

2.2

0.92

K_2O

<0.1

0.3

0.1

P_2O_5

<0.1

0.48

<0.1

0.23

H_2O(-)

-

0.17

0.07

_____

_____

_____

_____

Total

100.06

99.8

99.8

100.2

Total Fe

76.8

24.1

8.1

22.6

I = HCl-soluble magnetic fraction (10.05%of sample.)

II = HCl-soluble nonmagnetic fraction (47.86% of sample.)

III = Total acid-insoluble fractions (41.95% of sample.)

*S and Cr were determined on powder prepared from total meteorite.

**Based on composite total Fe, direct S, and metallic Fe.

suited, along with acid-insoluble fractions that were combined to form a separate sample (III). These three samples were analyzed separately, the results along with a composite analysis being given in Table 1. The analytical procedures used were mainly the well-known methods of Hillebrand, et al. (1953) and Peck (1964).

The major elements were determined on aliquots of solutions I and II, and on a solution resulting from the fusion of a sample of the acid-insoluble fraction (III) with Na_2CO_3. SiO_2 was determined gravimetrically by dehydration from perchloric acid and volatilization with HF. The R_2O_3 group was precipitated with NH_4OH, ignited and weighed, and dissolved and taken to volume in 10% H_2SO_4 after pyrosulfate fusion. No correction of SiO_2 values was made for SiO_2 present in the R_2O_3 precipitate. This was followed by precipitation of the ammonium sulfide group. CaO was determined gravimetrically by double precipitation of calcium oxalate and ignition to CaO. MgO was determined gravimetrically by double precipitation as Mg-(NH_4)PO_4 and ignition to Mg_2P_2O_7. Mn was determined in the Mg_2P_2O_7 and found to be negligible.

Total Fe and Al_2O_3 were determined on aliquots of the R_2O_3 group solution. Fe was titrated with K_2Cr_2O_7 after reduction with silver. Al_2O_3 was determined gravimetrically with 8-hydroxyquinoline from a tartrate-cyanide solution. Loveland, Schmitt, and Fisher (1969), by neutron activation analysis, found 1.21% Al in Harleton; this is equivalent to 2.19% Al_2O_3, in excellent agreement with the gravimetric determination.

The following determinations were made on separate aliquots of I and II, and samples of III. TiO_2 was determined spectrophotometrically with tiron using the method of Rigg and Wagenbauer (1961). Ni was determined gravimetrically with dimethylglyoxime. MnO was determined spectrophotometrically as MnO_4−1 using KIO_4 as the oxidant. P_2O_5 was determined spectrophotometrically as the phosphomolybdate and Co with nitroso-R, both as described as Moss, Hey, and Bothwell (1961). The alkalis were determined flame-photometrically (Grimaldi, 1960).

Cr and S were determined on a powder sample prepared from fragments of the total meteorite, fused with Na_2CO_3 and KNO_3. Cr was determined spectrophotometrically as Cr_2O_7−2 on the leachate from the fusion. The same solution was used for the gravimetric determination of S as BaSO_4.

FeS in I and II was calculated from the S value determined by passing the gas evolved on solution of the sample into lead acetate. The evolved H_2S was precipitated as PbS and converted to PbSO_4 for a gravimetric determination of S. The FeS figure used in the composite analysis was based on the direct determination of S in the total meteorite.

FeO was calculated in III by assuming that all of the iron was present as FeO. In II it was assumed that all of the Fe present was in the form of FeO with the exception of that combined with S. The SiO_2/MgO ratio in I was very close to that observed in II. On this basis, it was assumed that the FeO/SiO_2 ratio in I would be the same as that ob-

TABLE 2.—Normative composition of the Harleton meteorite in weight percent

Constituent

Weight percent

Constituent

Weight percent

Olivine

43.2

Apatite

0.5

Hypersthene

25.5

Chromite

0.8

Diopside

4.9

Ilmenite

0.2

Albite

7.8

Troilite

6.6

Anorthite

1.4

Nickel-iron

8.5

Orthoclase

0.6

served in II. This ratio was used to calculate the FeO value for I. The FeO value for the composite analysis was based on the total Fe, the direct S, and the metallic Fe values. The metallic Fe value in I was determined by correcting the total Fe value in I for Fe combined as FeS and FeO.

The normative composition of the Harleton meteorite, calculated from the analysis, is given in Table 2. The figures are consistent with the observed mineralogical composition. From the analysis acid-soluble silicate, which is olivine, amounts to approximately 42%, compared to 43.2% in the norm; acid-insoluble silicate (pyroxene and plagioclase) is 42%, and the norm shows 9.8% plagioclase and 30.4% pyroxene.

The chemical and mineralogical composition of the Harleton meteorite shows that it is a typical olivine-hypersthene chondrite, according to the classification of Prior (1920). The total iron content, 22.6%, is that of an L (low-iron) group chondrite in the Urey-Craig classification (1953). Van Schmus and Wood (1967:764) classify Harleton as an L6 chondrite.

Literature Cited

Bainbridge, A. E., H. E. Suess, and H. Wänke

1962.    The Tritium Content of Three Stony Meteorites and One Iron Meteorite. Geochimica et Cosmochimica Acta, 26:471–473.

Bunch, T. E., Klaus Keil, and K. G. Snetsinger

1967.    Chromite Composition in Relation to Chemistry and Texture of Ordinary Chondrites. Geochimica et Cosmochimica Acta, 31:1569–1582.

Davis, R., Jr., R. W. Stoenner, and O. A. Schaeffer

1963.    Cosmic-Ray Produced Ar^37 and Ar^39 Activities in Recently Fallen Meteorites. Pages 355–365 in Radioactive Dating. Vienna: International Atomic Energy Agency.

Eberhardt, P., O. Eugster, J. Geiss, and K. Marti

1966.    Rare Gas Measurements in 30 Stone Meteorites. Zeitschrift für Naturforschung, 21a:414–426.

Fireman, E. L.

1967.    Radioactivities in Meteorites and Cosmic-Ray Variations. Geochimica et Cosmochimica Acta, 31:1691–1700.

Fireman, E. L., J. DeFelice, and D. Tilles

1963.    Tritium and Radioactive Argon and Xenon in Meteorites and Satellites. Pages 323–334 in Radioactive Dating. Vienna: International Atomic Energy Agency.

Goel, Parmatma S., and Truman P. Kohman

1962.    Cosmogenic Carbon–14 in Meteorites and Terrestrial Ages of “Finds” and Craters. Science, 136:875–876.

Grimaldi, F. S.

1960.    Dilution-Addition Method for Flame Spectrophotometry. U.S. Geological Survey Professional Paper, 400B:494–495.

Hillebrand, W. F., G. E. F. Lundell, H. A. Bright, and J. I. Hoffman

1953.    Applied Inorganic Analysis. 2nd edition, 1034 pages. New York: John Wiley and Sons.

Honda, Masatake, and James R. Arnold

1964.    Effects of Cosmic Rays on Meteorites. Science, 143:203–212.

Kohman, T. P., and P. S. Goel

1963.    Terrestrial Ages of Meteorites from Cosmogenic C^14. Pages 395–411 in Radioactive Dating. Vienna: International Atomic Energy Agency.

Kuno, H.

1954.    Study of Orthopyroxene from Volcanic Rocks. American Mineralogist, 39:30–46.

Loveland, W., R. A. Schmitt, and D. E. Fisher

1969.    Aluminum Abundances in Stony Meteorites. Geochimica et Cosmochimica Acta, 33:375–385.

Mabuchi. H., J. Tobailem, C. Leger, and R. Bibron

1968.    Radioactivite induite par le rayonnement cosmique dans la meteorite Granès. Geochimica et Cosmochimica Acta, 32:949–963.

Moss, A. A., M. H. Hey, and D. I. Bothwell

1961.    Methods for the Chemical Analysis of Meteorites, I: Siderites, Mineralogical Magazine, 32:802–816.

Osborn, T. W.

1972.    Evidence for Association between Ir and Al in L Chondrites. Nature Physical Science, 239:10–11.

Peck, L.

1964.    Systematic Analysis of Silicates. U.S. Geological Survey Bulletin, 1170:1–89.

Poldervaart, A.

1950.    Correlation of Physical Properties and Chemical Composition in Plagioclase, Olivine, and Orthopyroxene Series. American Mineralogist, 35:1067–1079.

Prior, G. T.

1920.    The Classification of Meteorites. Mineralogical Magazine, 19:51–63.

Ramdohr, Paul

1973.    The Opaque Minerals in Stony Meteorites. 244 Pages. Amsterdam: Elsevier Publishing Company.

Rigg, Tryson, and Heinrich A. Wagenbaur

1961.    Spectrophotometric Determination of Titanium in Silicate Rocks. Analytical Chemistry, 33:1347–1349.

Rowe, M. W., M. A. Van Dilla, and E. C. Anderson

1963.    On the Radioactivity of Stone Meteorites. Geochimica et Cosmochimica Acta, 27:983–1001.

Shedlovsky, Julian P., Philip J. Cressy, Jr., and Truman P. Kohman

1967.    Cosmogenic Radioactivities in the Peace River and Harleton Chondrites. Journal of Geophysical Research, 72:5051–5058.

Suess, H. E., and H. Wänke

1962.    Radiocarbon Contents and Terrestrial Age of Twelve Stony Meteorites and One Iron Meteorite. Geochimica et Cosmochimica Acta, 26:475–480.

Urey, H. C, and H. Craig

1953.    The Classification of the Stone Meteorites and the Origin of the Meteorites. Geochimica et Cosmochimica Acta, 4:36–82.

Van Schmus. W. R., and J. A. Wood

1967.    A Chemical and Petrologic Classification for the Chondritic Meteorites. Geochimica et Cosmochimica Acta, 31:747–765.

Eight L-Group Chondrites: A Comparative Study

Ananda Dube, Becky J. Fredriksson, Eugene Jarosewich, Joseph A. Nelen,
Albert F. Noonan, Jane O’Keefe, and Kurt Fredriksson

ABSTRACT

The bulk chemical compositions of eight L-group chondrites, Atarra, Cranganore, Karkh, Kuttippuram, Mabwe-Khoywa, Mirzapur, Segowlie, and Sultanpur, all falls and mostly from the collection of the Geological Survey of India, Calcutta, have been determined by classical wet chemical analysis. In addition, the compositions and abundances of individual mineral phases were determined with electron-probe techniques.

Seven of the eight stones represent the petrographie types 4, 5, and 6, while one, Mabwe-Khoywa, cannot be classified because it is a polymict breccia. Still, the practically identical bulk and phase compositions imply a similar (chemically identical) parent; thus any metamorphic recrystallization must have occurred before the various fragments accreted. Therefore, we interpret these L-group chondrites, in addition to four of the H-group previously reported, as products of multiple impacts of various intensities on similar materials if not on the same parent body.

Ananda Dube, Geological Survey of India, Calcutta 13, India; Becky J. Fredriksson, Eugene Jarosewich, Joseph A. Nelen, Albert F. Noonan, and Kurt Fredriksson, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560; Jane O’Keefe, Geology Department, Rice University, Houston, Texas 77001.

Introduction

The eight hypersthene (L-group) chondrites, all observed falls, described here were selected for detailed studies because they represent wide ranges in macroscopic texture, hardness, color, and visibility of chondrules. This is reflected to some degree in their classifications by Hey (1966), Mason (1975), and Wasson (1974). Two are black chondrites, L4 and L6 respectively; two are gray chondrites, both L6; one is a brecciated white chondrite, L6; one is a brecciated, veined intermediate chondrite, L5; and one is a crystalline chondrite, L6. Mabwe-Khoywa looks like a polymict breccia and has so far only been classified as “L.” All eight meteorites were analyzed by classical wet chemical techniques (Jarosewich, 1966). Major phase compositions were determined by standard electron-probe techniques; in addition, the abundances of the major phases (modes) were determined by our previously described microprobe technique (Fredriksson, et al., 1975). Bulk chemical compositions are given in Table 1; the compositions and abundances of the major silicates appear in Table 2. The mode and the norm (mineralogical composition calculated from the chemical analysis) are presented in Table 3.

These chondrites are listed, classified and/or described by Cotter (1912), Hey (1966), Mason (1975), Silberrad (1932); their olivine compositions were given by Mason (1963, 1967). Detailed information on the morphology and the distribution of samples is given by Murthy, Srivastava, and Dube (1969) for the Indian falls (all but Karkh and Mabwe-Khoywa). In the following descriptions we do not repeat such information, but give a few more references as required. Rather, we present our observations on a number of polished thin sections (some ultrathin prepared by Mr. G. Moreland of the Smithsonian Institution) and some additional microprobe analyses not included in Tables 2 and 3. After the name of each chondrite we give pertinent references followed by the composition of

TABLE 1.—Chemical composition in weight percent

TABLE 2.—Abundance and average composition of major silicate phases

TABLE 3.—Normative and modal compositions in weight percent

olivine in mole percent fayalite (Fa), pyroxene in mole percent wollastonite (Wo), enstatite (En), and ferrosilite (Fs), plagioclase in mole percent orthoclase (Or), albite (Ab), and anorthite (An). These data are followed by observations on texture, analyses of phases of special interest or significance, etc.

ACKNOWLEDGMENTS.—We are grateful to the former Director General of the Geological Survey of India, Dr. C. Karunakaran, for encouraging this study. Mr. R. K. Sundaram and Drs. S. V. P. Iyengar and M. V. N. Murthy also gave much assistance. We are indebted to the Smithsonian Systems Analysis Division for data reductions. The work has been supported in part by grants from the Smithsonian Research Foundation (Sg 3352213 and 3352363) and Foreign Currency Program.

Description of Samples

ATARRA

Olivine: Fa24.9

Pyroxene: Wo_2 En_16 Fs_22

Plagioclase: Or_5 Ab_78 An_17

The stone has been classified as a black chondrite and L4. We would rather call it a gray (or intermediate) veined chondrite, while the L4 designation seems appropriate (Figure 1). The olivine composition is remarkably constant: 30 randomly selected olivines varied between 16.4% and 16.8% Fe. In the dominant pyroxene (Pyroxene 1, Table 2), the iron varies between 10.0% and 12.4% Fe. Abundant shock effects are visible: twisted pyroxene lamellae in radiating pyroxene chondrules, and fracturing and wavy extinction in olivine crystals.

CRANGANORE

Walker, 1924

Olivine: Fa24.7

Pyroxene: Wo1.6 En_77 Fs_21

Plagioclase: Or_6 Ab_77 An_17

Cranganore has been classified as a gray chondrite and L6, which generally agrees with our observations, although the hand specimen is rather rusty and more crumbly than most L6 chondrites. Most thin sections show a structure similar to that in Figure 7 (Segowlie), but some parts show fairly well-delineated chondrules, and approach an L5 texture. The matrix is mostly fragmental to hypidiomorphic-granular, but the grain size varies considerably within millimeter or centimeter distances; this indicates that the last recorded event was brecciation not recrystallization.

FIGURE 1.—Photomicrograph of thin section of the Atarra L4 chondrite. The stone is fragmentary with relatively few chondrules, while the olivine composition is constant throughout. Length of section, 4.5 mm.

KARKH

Fermor, 1907

Olivine: Fa23.3

Pyroxene. Wo0.2 En_80 Fs_20

Plagioclase: Or_4 Ab_78 An_18

This stone has been classified as a gray chondrite and as L6, which seems appropriate. Our hand specimens, although small, shows some brecciation or veining; shock effects are also evident in the thin sections. The clinopyroxene (Pyroxene 2, Table 2) in Karkh is more inhomogeneous and also has distinctly lower Ca and higher Fe than the almost pure diopsides in the other seven chondrites, in-

FIGURE 2.—Photomicrograph of thin section of the Karkh L6 chondrite. Quartz-corundum normative glass, white, between olivine bars, light gray, in a “ghost” chondrule. Opaques are nickel-iron and troilite. Length of section, 0.6 mm.

cluding Cranganore, described above, which also is an L6 gray chondrite. Another anomaly is the presence of small amounts of silica-alumina-rich glass found mostly in ghost chondrules, i.e., barely distinguishable in thin sections (Figure 2), and also in the matrix as completely integrated fragments of chondrules. This clear, perfectly isotropic glass is rather inhomogeneous as exemplified below. All values are in weight percent.

Analyses 2 and 7 were made in the indistinct chondrule shown in Figure 2; analyses 20 and 23 in an integrated chondrule fragment. In all cases the glass is quartz (~10% to ~50%) and corundum (~4% to ~14%) normative and is thus not simply vitrified plagioclase. Although shock melting may have occurred, the glass, and probably its parent material, is not in equilibrium with the major olivine-pyroxene phases.

KUTTIPPURAM

Keil and Fredriksson, 1964

Olivine: Fa24.3

Pyroxene: Wo1.6 En_78 Fs_20

Plagioclase: Or_5 Ab_76 An_19

This meteorite has been described as a veined white crystalline hypersthene chondrite and L6, appropriate for the main part. Our small sample and also one thin section (Figure 3), however, contain fragments consisting mainly of olivine, accessory pyroxene, and interstitial Si-Al-rich glass, with practically no metal or troilite. It is remarkable that olivine is of the same composition in both parts: 22.5 ± 0.5% FeO. The glass is fairly constant in composition with ~67 ± 1% SiO_2 21.5 ± 1% Al_2O_3, 2.1% CaO. Sodium and potassium are somewhat variable (also the Na/K ratio); our best estimate is ~10% Na_2O and 0.8% K_2O. Thus the glass is consistently quartz (~10%–20%) and corundum (~6%) normative and resembles the average glass in the Karkh chondrite described above.

MABWE-KHOYWA

Coulson, 1940

Olivine: Variable, dominant Fa23.25; minor Fa~10, others present

Pyroxene: Variable, dominant Wo~1 En~78 Fs~21 (extreme Wo_2 En_68 Fs_30); minor strongly variable diopsides; enstatite, rare ~ Wo_2 En_90 Fs_8

Though it has been classified as an L or hyper-

FIGURE 3.—Photomicrograph of thin section of the Kuttippuram L6 veined white chondrite. Note the “achondritic” fragment, lower half, consisting mostly of olivine and glass, and almost devoid of metal and troilite. The olivine-pyroxene composition in the fragment is the same as in matrix and chondrules. Length of section, 3 mm.

sthene chondrite, the variable olivine-pyroxene composition and especially Figures 4 and 5 illustrate why the classification is exceedingly difficult. Our sample of Mabwe-Khoywa is a complex rock, a breccia or conglomerate. In addition to a great variety of chondrules it contains lithic fragments, rounded or angular; some of these were pre-existing chondrites, while others appear to be achondritic. Most olivines and pyroxenes, 90%–95% in these parts, with the exceptions described below, still are in the dominant range, as are numerous individual larger mineral fragments. The main part of the matrix also consists of comminuted chondrules and chondrite fragments. In general, but not always, olivines in the matrix seem to be richer in iron than olivines in fragments, while the lowest iron content is found in some rare chondrules (Figure 5). Also, the larger olivines in the “achondritic” fragments are more iron rich, similar to the dominant olivine, than the smaller grains. At least three types of pyroxene were found. The dominant orthopyroxene is found in almost all types of chondrules and fragments and is essentially homogeneous, i.e., equilibrated. The diopsidic pyroxene varies at least from Wo_20 to Wo_40 with corresponding variations in iron and magnesium. The third pyroxene, approximately enstatite, is rare and was only found in random point analyses (several hundreds) and we could not establish in which kind of fragment it occurs. A few pyroxenes in the range between the orthopyroxene and the diopside could represent overlap between host crystals and exsolution lamellae.

Some of the chondrules and fragments in Figure 4 deserve brief individual descriptions. In the olivine-pyroxene chondrule (Figure 5), the olivine has the lowest iron content encountered, and most crystals are zoned from Fa_8 in the center to Fa_10 at the edge. The pyroxene appears to be monoclinic and has exsolution lamellae, but nearly constant composition (~Wo_1 En_87 Fs_12); thus the nonparallel extinction and the exsolution lamellae are probably shock produced. Further indication of shock effects is the opaque spherule consisting of apparently quenched eutectic troilite plus nickel-iron similar to those in the H-chondrite Pulsora (Fredriksson, et al., 1975).

Adjacent to the chondrule is a chondritic fragment containing a fragment of a barred olivine and glass chondrule. The olivine bars (Figure 5) seem intergrown with the first chondrule but belong to the equilibrated main type, Fa~24. Also other olivines and pyroxenes in this chondrite fragment belong to the main type. The glass in the barred chondrule is almost albitic but has slight normative excesses of quartz and corundum.

Figure 4 shows a large “achondritic” fragment consisting largely of olivine and glass with metal, troilite, and chondrules conspicuously lacking. The smaller olivine crystals have low iron content, Fa~12, while the larger crystals are somewhat zoned but close in composition (Fa_20_22) to the “equilibrated” main olivine. This might indicate that the larger crystals are relatively late xenoliths. The interstitial glass is enriched in silica (~62%) and alumina (~8%), but has low soda (~0.8%) and substantial amounts of iron (16% FeO) and magnesia (~5%).

FIGURE 4.—Photomicrograph of thin section of the hypersthene chondrite. Mabwe-Khoywa. The chondritic main mass is L3 or L4, olivines in most chondrules and fragments have approximately Fa_24; one chondrule indicated by arrow and shown in Figure 5 has Fa8–11. In the achondritic fragment (right), the small olivines are in the range Fa11–14, some of the larger Fa20–23, while the interstitial glass is albitic with excess normative quartz and corundum. Length of section, 12.5 mm.

The olivine with highest iron content was found in a lithic fragment (possibly a chondrule). The crystal has 25.5% FeO (Fa_28) in the center and 24% at the edge. Adjacent glass is again mostly albitic but has a substantial excess of normative corundum along with some iron and magnesium. It is remarkable that Mabwe-Khoywa in spite of all this diversity of components has a bulk composition almost identical to the other L-chondrites. Indeed, even the abundances and average compositions of the phases are close. Mabwe-Khoywa is apparently the product of multiple impacts, and after the first, or first few, impacts its chemistry was not significantly altered.

MIRZAPUR

Olivine: Fs24.8

Pyroxene: Wo1.8 En_78 Fs_21

Plagioclase: Or3.4 Ab_80 An_17

It has been classified as a veined intermediate chondrite (Cia) and L6; we prefer L5 (towards L4) because our specimen is relatively unconsolidated and the thin sections show well-delineated chon-

FIGURE 5.—An Mg-rich olivine and pyroxene chondrule (shown by arrow in Figure 4), is lined with nickel-iron and troilite, black. It is intergrown with matrix (lower right) and with a fragment of a barred “equilibrated” olivine (Fa_25) chondrule (lower left), which is part of a fragment of preexisting chondrite (shown under arrow in Figure 4). Length of section, 0.6 mm.

drules (Figure 6), some with Si-Al enriched and only partly devitrified glass. This glass has approximately 66% SiO_2, 21% Al_2O_3, and 8% soda; the normative composition is again almost albitic with some quartz and minor corundum and pyroxene. Shock alterations, including fracturing wavy extinction and veining (Fredriksson, et al., 1964), are abundant and the glass may be impact-produced; however, it would hardly be in equilibrium with olivine-pyroxene at metamorphic temperatures of 800° to 1000° C.

SEGOWLIE

Olivine: Fa25.0

Pyroxene: Wo1.8 En_77 Fs_21

Plagioclase: Or_5 Ab_78 An_17

Segowlie perfectly fits the definitions for a crystalline chondrite and almost as well for the L6 group. The photomicrograph in Figure 7 shows individual patches with equigranular, recrystallized structure, identical to that in the H-chondrite Andura (Fredriksson, et al., 1975). However, the grain size varies grossly between the different parts, i.e., the stone is still a breccia. Thus it seems unlikely even in this case that the structure is due to simple metamorphic recrystallization. Some parts (Figure 8) also show rather well-delineated chondrules, i.e., a structure more like L5.

SULTANPUR

Walker, 1924; Coulson, 1940

Olivine: Fa26.2

Pyroxene: Wo_2 En_76 Fs_22; Wo_44 En_48 Fs_8

Plagioclase: Or_5 Ab_77 An_18

Sultanpur has been classified as a black chondrite and an L6, but both designations are somewhat questionable. It is more a black chondrite than Atarra, but not in the same sense as shock products like Farmington and McKinney (Fredriksson, et al., 1964; Heymann, 1967). Sultanpur seems rather to be close to a brecciated (and veined) gray chondrite. The grain size is much finer and the chondrules better defined than in the L6 Segowlie; thus Sultanpur could probably also be an L5, particularly because the plagioclase does not occur as clear, well-defined grains. Evidence of shock effects, e.g., fracturing and undulatory extinction, in both olivine and pyroxene are common, however.

Discussion

Bulk and phase compositions, as well as phase abundances (Tables 1, 2), clearly show the close relation among these eight meteorites in spite of their different appearances and classifications. Essentially only the contents of nickel-iron and troilite vary significantly (Table 3), while the silicate fractions are chemically almost identical even in the obviously fragmentary Mabwe-Khoywa. The silicate fraction compositions in Table 4 were obtained by deducting metal, troilite, Fe for ilmenite and chromite, and Ca for apatite from the bulk analysis. While normative and modal amounts of pyroxene are close, both olivine and plagioclase are higher in the norm, which must be compensated for by

FIGURE 6.—Photomicrograph of thin section of the Mirzapur, L5 veined intermediate chondrite (see text). Note the black, opaque, shock veins and shock altered areas in the center. The chondrules are quite distinct, although often deformed in some parts. Length of section, ~20 mm.

the modal residue. Because the residue in the modal analyses (Table 3) is substantial, an attempt has been made to establish its chemical composition by deducting the components of the modal minerals and normative metal, sulfide, chromite, ilmenite, and apatite from the bulk analyses. The compositions of these residues are compared to the bulk silicates in Table 4. Although this modal residue is largely due to overlap between grains in the probe analysis, it appears that in all cases except for Atarra, the residue has 1–3% normative corundum (Al_2O_3) and more iron-rich, including Atarra, normative olivine than the bulk silicates. Thus all these L-chondrites appear to have small amounts of a matrix component enriched in Al and Fe, similar to the H-chondrites previously described by Fredriksson, et al. (1975). Moreover, the comparison in Table 5 of the bulk silicate

FIGURE 7.—Photomicrograph of thin section of “normal” part of the Segowlie, L6 crystalline chondrite. Note the extreme integration of the chondrule (right), which is, however, finegrained (as is part of the matrix above it) compared to the adjacent matrix. Length of section, 4.5 mm.

compositions of our eight L-chondrites with the four H-chondrites described by Fredriksson, et al. (1975, table 4) shows that they differ essentially only in the FeO content; if normalized to the same FeO content, the silicates seem identical within analytical uncertainties.

Conclusion

Eight L-chondrites with different textures and classifications are chemically and mineralogically almost identical. This indicates a similar origin from chemically the same parent material and by the same processes, probably impact events.

The silicate compositions in these L-chondrites, except for the FeO contents, are close to those of four previously described H-chondrites (Fredriksson, et al., 1975), and if normalized to the same iron content, the silicates are chemically identical. Thus, we imply a common parent material for H- and L-(also LL-) chondrites. The essential difference between the groups only lies in the amounts of iron reduced and the portion of such produced metal remaining. Chemically the parent material might well have been close to the main semi-amorphous silicates in the carbonaceous chondrites such as Ivuna or Orgueil (Boström and Fredriksson, 1966). Early impacts may have caused the chemical alterations,

TABLE 4.—Chemical composition of bulk silicates from chemical analyses and of unidentified fractions “residue” in the modal analyses

FIGURE 8.—Photomicrograph of thin section of chondritis part of the Segowlie chondrite (cf., Figure 7). Note the patchy variation in grain size. Length of section, 3.6 mm.

TABLE 4.—Silicate composition of eight L- and four H-group chondrites

e.g., loss of volatiles, and extensive melting and drop (chondrule) formation because of low density and high compressibility of the target material. Later impacts on a parent body caused more brecciation, destruction of preexisting chondrules and chondrites, and “solid state” recrystallization rather than melting and formation of liquid droplets, i.e., new chondrules, although this must also have occurred as it did on the moon.

Literature Cited

Boström, K., and K. Fredriksson

1966.    Surface Conditions of the Orgueil Meteorite Parent Body as Indicated by Mineral Associations. Smithsonian Miscellaneous Collections, 151(3): 1–39.

Cotter, G. de P.

1912.    Notes on Indian Aerolites Recorded since 1906. Records of the Geological Survey of India, 42:265–277.

Coulson, A. L.

1940.    Catalogue of Indian Meteorites. Memoirs, Geological Survey of India, 75:1–346.

Femor, L. L.

1907.    Notes on Some Indian Aerolites. Records of the Geological Survey of India, 35:90–02.

Fredriksson, K., P. DeCarli, R. O. Pepin, J. H. Reynolds, and G. Turner

1964.    Shock Emplaced Argon in a Stony Meteorite. Journal of Geophysical Research, 69:1403–1411.

Fredriksson, K., A. Dube, E. Jarosewich, J. Nelen, and A. Noonan

1975.    The Pulsora Anomaly: A Case against Metamorphic Equilibration in Chondrites. In George S. Switzer, editor, Mineral Sciences Investigations 1972–1973. Smithsonian Contributions to the Earth Sciences, 14:41–53.

Hey, M. H.

1966.    Catalogue of Meteorites. 3rd edition, 637 pages. London: British Museum of Natural History.

Heymann, D.

1967.    On the Origin of Hypersthene Chondrites: Ages and Shock Effects of Black Chondrites. Icarus, 6:189–221.

Hobson, G. V.

1927.    Six Recent Indian Aerolites. Records of the Geological Survey of India, 60:128–152.

Jarosewich, E.

1966.    Chemical Analyses of Ten Stony Meteorites. Geochimica et Cosmochimica Acta, 30:1261–1265.

Keil, K., and K. Fredriksson

1964.    The Iron, Magnesium, and Calcium Distribution in Coexisting Olivines and Rhombic Pyroxenes of Chondrites. Journal of Geophysical Research, 69:3487–3515.

Mason, B.

1963.    Olivine Composition in Chondrites. Geochimica et Cosmochimica Acta, 27:1011–1023.

1967.    Olivine Composition in Chondrites—A Supplement. Geochimica et Cosmochimica Acta, 31:1100–1103.

1975.    List of Meteorites in the National Museum of Natural History, Smithsonian Institution. In George S. Switzer, editor, Mineral Sciences Investigations 1972–1973. Smithsonian Contributions to the Earth Sciences, 14:71–83.

Murthy, M. V. N., S. N. P. Srivastava, and A. Dube

1969.    Indian Meteorites. Memoirs, Geological Survey of India, 99:1–192.

Silberrad, C. A.

1932.    List of Indian Meteorites. Mineralogical Magazine, 23:290–304.

Walker, H.

1924.    Recent Falls of Aerolites in India. Records of the Geological Survey of India, 55:133–142.

Wasson, J. T.

1974.    Meteorites. 316 pages. New York: Springer-Verlag.

The Ras Tanura, Saudi Arabia, Chondrite

Brian Mason

ABSTRACT

The Ras Tanura meteorite, which fell near Dhahran, Saudi Arabia, on 23 February 1961, is an olivine-bronzite (H6) chondrite; principal minerals are olivine and orthopyroxene, with minor amounts of plagioclase, nickel-iron, and troilite, and accessory chromite and whitlockite.

The Ras Tanura meteorite was an observed fall on 23 February 1961 at approximately 1142 GCT. An eyewitness at Dhahran, about 40 km south of the point of fall (26°40’N, 50°09’E) described the event as follows:

Brian Mason, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D. C. 20560.

A distinct boom similar to a sonic boom was heard, followed by a rolling rumble, like thunder. In an otherwise perfectly clear sky there appeared a dense white cloud with a less dense zig-zag tail trailing away to the south. The tail indicated a track from south to north. The tail disappeared after about 10 minutes. The dense white cloud persisted for about 30 minutes. During this time it drifted slowly to the east about 15°. The cloud diameter subtended an angle of about 2°. The elevation angle was about 60°. The object was observed on an azimuth of approximately 10°.

A single small crusted stone, weight approximately 6 grams, is all that was recovered. It was picked up on the South Pier of the Aramco refinery at Ras Tanura. One observer believed he saw a waterspout offshore as if caused by a falling body “shortly after the explosion.” It is thus probable that the main mass of the meteorite fell in the sea.

The stone was turned in to Aramco officials, who forwarded it to the Air Force Intelligence Center, Wright-Patterson Air Force Base, Ohio, as an unidentified flying object. It was eventually transferred to the Smithsonian Astrophysical Observatory and thence to the National Meteorite Collection, where it is cataloged as USNM 2171.

Ras Tanura is a well-recrystallized olivine-bronzite chondrite, H6 in the Van Schmus-Wood classification. A cut surface shows abundant metal particles in a pale gray matrix; no chondrules are visible. A slight amount of rusting has taken place in association with the metal particles. In thin section the meteorite shows an allotriomorphic-granular texture. Average grain size is approximately 0.1 mm, but with scattered larger olivine grains, up to 0.5 mm. Sparse chondrules are recognizable by textural differences, principally a coarser grain size and interstitial fine-grained material (probably devitrified glass), and by a vaguely circular outline.

The principal minerals are olivine and orthopyroxene. Microprobe analyses give a composition of Fa19.0 for the olivine and Fs17.4 for the orthopyroxene; the orthopyroxene contains 0.75% CaO. A small amount of diopside, composition Wo_44 En_50-Fs_6, is associated with the orthopyroxene; the diopside contains 0.53% Na_2O, 0.31% TiO_2, and 0.73% Cr_2O_3, whereas the orthopyroxene contains <0.1% Na_2O, 0.10% TiO_2, and 0.14% Cr_2O_3. Plagioclase, composition An_10 with a K_2O content of 0.95%, is also present. Accessory whitlockite was identified with the microprobe. The opaque minerals are nickel-iron, troilite, and chromite.

Geochemical Differences among Components of the Allende Meteorite

Brian Mason and Philip M. Martin

ABSTRACT

Four discrete groups of inclusions have been distinguished in the Allende meteorite. Group I are melilite-rich chondrules; Groups II and III are aggregates made up largely of spinel and fassaite; Group IV are olivine-rich chondrules and aggregates. Groups I, II, and III have been previously classed together as Ca/Al-rich inclusions, but each group has a distinctive trace-element pattern, most clearly shown by the lanthanide distribution. Group I has a relatively unfractionated pattern (with a small positive Eu anomaly) at about 10–15 times chondrites; Group II a highly fractionated pattern with depletion of the heavier lanthanides and negative Eu and positive Tm and Yb anomalies; Group III an unfractionated pattern at about 20 times chondrites, except for negative Eu and Yb anomalies; Group IV a relatively unfractionated pattern at 2–4 times chondrites. The Yb anomalies can be ascribed to highly reducing conditions producing Yb 2+, or to fractional condensation in a cooling nebula; the Tm anomaly may be due to fractional condensation or some unrecognized factor. The complex pattern of trace element distribution in these Allende inclusions indicates a complex history of formation of this meteorite from the ancestral nebula.

Brian Mason, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D. C. 20560. Philip M. Martin, Geology Department, University of Leicester, Leicester LE1 7RH, England.

Introduction

The fall of the Allende meteorite on 8 February 1969, provided a large amount of a Type III carbonaceous chondrite for research. A detailed description of the fall and recovery, and of the chemical and mineralogical composition of the meteorite has been published by Clarke, et al. (1971). The following principal components were distinguished and their volume percentages estimated by point-counting: matrix, largely of Fe-rich olivine (~60%), Mg-rich chondrules (~30%), Ca/Al-rich chondrules (~’5%), and Ca/Al-rich aggregates (~5%). The chondrules and aggregates have been the subject of intensive research and speculation as to their origin (Marvin, Wood, and Dickey, 1970; Grossman, 1972; Osborn, et al., 1974; Wänke, et al., 1974). The purpose of this paper is to show that they comprise at least four groups of distinctive chemical and mineralogical composition. The differences between the groups are indicative of extensive chemical fractionations prior to the accumulation of the components into the parent body of the Allende meteorite.

ACKNOWLEDGMENTS.—We would like to express our appreciation to Mr. E. Jarosewich, Mr. J. Nelen, Dr. A. L. Graham, Mrs. P. Muir, and Mr. W. Nance for assistance in the analytical work. The spark source mass spectrometric analyses were made in the Research School of Earth Sciences, Australian National University, and we are indebted to Dr. S. R. Taylor for the use of these facilities over several weeks in 1971 and 1973. The Smithsonian Research Foundation and the National Aeronautics and Space Administration (Grant NGR–09–015–170) contributed part of the costs of this research, and their contribution is gratefully recorded.

Experimental Methods

A large number of Allende specimens in the collection of the National Museum of Natural History (NMNH) were inspected, and those showing prominent and readily extractable chondrules and aggregates were selected for detailed examination. From these, over 20 chondrules and aggregates were extracted and subjected to complete chemical and mineralogical analysis, and many others investigated to a lesser degree.

As far as possible the chondrules and aggregates were extracted from the matrix by simple mechanical tools and hand-picking, to avoid introducing contaminants; however, in some instances methylene iodide was used to separate the denser matrix from the less dense chondrules and aggregates. Mineral identification and analysis were carried out using microscope, x-ray diffraction, and electron-microprobe techniques. Compositional data for major elements and some minor elements were obtained by microprobe analyses of glasses prepared by fusion with a lithium tetraborate flux. Some of the analyses of Groups II, III, and IV components have low summations (90–95%) because of the non-determination of some elements (Ni, Cr, S) and possible loss of alkalies during fusion. The trace element data were obtained by spark source mass spectrometry, using the procedures described by Taylor (1965, 1971). The precision is about ±5% and the accuracy about ± 10%, but they vary somewhat, being poorer for elements below mass number 105 than for those of higher mass number. Some trace elements could not be determined because their concentrations were below the sensitivity of the technique. Among the lanthanides, Lu was not determined because of its use as an internal standard. The single mass number (169) for Tm is subject to interference from a multiple carbon ion, and this element cannot be accurately determined at low concentrations (<0.3 ppm). The 169 line, however, was unexpectedly strong in Group II samples, indicating the presence of a positive Tm anomaly. The platinum metals were observed in most Groups I and II samples, but accurate determination was not possible because of the lack of suitable standards.

Results

The results of the chemical analyses are summarized in Tables 1 and 2 and Figures 1–3.

The Allende samples can be grouped in various ways using chemical, petrological, and textural parameters. From such considerations four distinct groups can be distinguished: Group I: Ca/Al-rich chondrules with relatively unfractionated lantha-nide abundance patterns; Group II: Ca/Al-rich aggregates with strongly fractionated lanthanide abundance patterns: Group III: Ca/Al-rich aggregates with relatively unfractionated lanthanide abundance patterns, except for Eu and Yb anomalies; and Group IV: Mg-rich chondrules and aggregates, with low and relatively unfractionated lanthanide patterns.

Group I, the largest group studied, can be characterized as melilite chondrules, rounded bodies up to 25 mm diameter with melilite (40%–75%), Al/Ti-rich pyroxene (10%–25%), spinel (10%–25%), and anorthite (5%—15%) as major phases. Samples falling in this group are 26, 29, 30, 31, 32, 33, 3529G, 3658, and 3898. They have a coarsely crystalline igneous texture and are characterized chemically by high Al_2O_3 (22%–23%), high CaO (23%–33%), low SiO_2 (25%–35%), and a relatively unfractionated lanthanide abundance pattern averaging 10–15 times chondrites but with a small positive Eu anomaly.

Sample 5171 probably belongs to Group I but has a fine-grained texture. It has low Al_2O_3 and Ba and high Rb compared to other Group I samples. Mineralogically it is characterized by the presence of grossular (about 10%), as well as the major phases melilite, pyroxene, spinel, and anorthite.

Another well-defined group, Group II, comprises samples 37, 3598, 3803, 4691, and 4692, which are fine-grained irregular aggregates (except for 4691, which is texturally similar to the chondrules of Group I), with spinel, pyroxene, melilite and/or grossular as important phases, and minor contents of nepheline and sodalite. Their chondrite-normalized lanthanide abundance patterns are very interesting (Figure 1). The light lanthanides (La-Sm) are relatively high, with rapidly diminishing abundances of the the heavier lanthanides. Superimposed upon these patterns are marked negative Eu and positive Tm and Yb anomalies.

The two remaining fine-grained Ca/Al-rich aggregates, 3593 and 4698, classified as Group III,

TABLE 1.—Chemical abundances in components of the Allende meteorite (oxides in weight numbers are individual chondrules and inclusions from NMNH 3529; the figures for the

possess chondrite-normalized lanthanide abundance patterns quite distinct from those of Group II. They have a relatively unfractionated lanthanide pattern at about 20 times chondrite abundances, on which are superimposed marked negative Eu and Yb anomalies. Group III also have lower average SiO_2 and higher Al_2O_3, Y, Zr, Mo, and possibly Nb than Group II. Mineralogically, Group III aggregates are characterized by the presence of spinel, pyroxene, grossular, sodalite, and nepheline; a high spinel content and the absence of melilite distinguish Group III from Group II.

The three samples classified as Group IV, 24, 3529P, and 4060, are characterized by high SiO_2 and MgO and low TiO_2 contents, and consist largely of olivine with lesser amounts of clinoenstatite and glass. They have a relatively low and unfractionated lanthanide distribution at 2–4 times percent, elements in ppm; 4-digit sample numbers are NMNH catalog numbers, 2-digit sample bulk composition are from a compilation by Mason, in press)

chondrite abundances. Texturally 3529P and 4060 are fine-grained aggregates whereas 24 is a large chondrule with a coarsely crystalline igneous texture.

Among the samples examined are a few that cannot be readily accommodated in the fourfold grouping here proposed. One such chondrule (3509 in Table 1) is that described by Clarke, et al. (1971: 39–41, type c chondrule), consisting of nepheline, sodalite, pyroxene, and olivine; the lanthanide concentrations are low (at about average chondritic abundances) with a positive Eu anomaly. One finegrained aggregate (NMNH 3529–36) has trace element abundances similar to those for Group I chondrules, except for extremely low Sr (5.1 ppm) and Ba (3.4 ppm) and high Rb (5.4 ppm) and Cs (0.32 ppm); mineralogically it is notable for a considerable content of grossular. Another fine-grained

TABLE 2.—Range and average chemical composition for each of the Allende groups (oxides in weight percent, elements in ppm; the numbers in parentheses after the group are the number of specimens analyzed)

aggregate (NMNH 3643), also with a considerable grossular content, has trace element concentrations generally similar to those of Group II, but its lanthanide distribution pattern lacks the characteristic negative europium anomaly. Further search will probably turn up additional unique chondrules and aggregates in the Allende meteorite.

The Ca/Al relationship between the four groups deserves special attention, since it has been shown that this ratio (by weight) is practically constant at 1.09 in chondrites (Ahrens and von Michaelis, 1969). Figure 2 shows that Groups I, II, and III materials deviate widely from this ratio, whereas Group IV materials are close to it. The figure also illustrates a significant compositional difference between Group I and Groups II and III, Group I having a relatively high Ca/Al ratio, the latter two, relatively low ones. Mineralogically this is expressed by the dominance of melilite in Group I samples, and its absence or presence in only minor amounts in other samples (except 4691).

Some other features of major element chemistry deserve notice (Table 2). Group IV is consistently higher in SiO_2 and MgO and lower in TiO_2 than the other groups. In Group I, Na_2O and FeO are notably low; in fact, some of these melilite chondrules are essentially iron- and alkali-free, and the ferrous iron and alkali in others may have been a later introduction, since microprobe analysis shows that these elements are concentrated at the chondrule margins.

Returning now to the trace element data, we find characteristic differences between the different groups. Group I samples have notably high Sr and low Rb, giving unusually high Sr/Rb ratios. Gray, Papanastassiou, and Wasserburg (1973) and Weth-erill, Mark, and Lee-Hu (1973) have analyzed similar samples with concordant results, and they have also found that the 87Sr/86Sr ratios are the lowest yet recorded, i.e., the Sr is “primordial” and has had essentially no radiogenic 87Sr added to it. Other Allende samples analyzed by them have much higher Rb, lower Sr, and higher 87Sr/86Sr ratios, and probably belong to Group II or Group III as these are defined here.

The neighboring elements Y, Zr, and Nb show rather coherent geochemical behavior, with a distinctive pattern in each of the four groups. They are most highly enriched in Group III (however, only two samples were analyzed, so the statistical basis is weak), notably enriched (at about 10 times chondritic level) in Group I, and relatively low (at about average chondritic level) in Groups II and IV. On examining the data for the individual samples (Table 1), an interesting relationship is apparent; the ratio Zr/Nb varies considerably from 1.1 to 53, whereas the ratio Zr/Y is remarkably constant, ranging only from 2.1 to 2.9. The Zr/Y coherence

FIGURE 1.—Chondrite-normalized lanthanide abundance patterns of the average of each group and of a bulk sample of the Allende meteorite: Group I: Ca, Al-rich chondrules; Groups II and III: Ca, Al-rich aggregates; Group IV: Mg-rich chondrules and aggregates. Note the Eu and Yb anomalies in Groups II and III and the Tm anomaly in Group II.

is illustrated in Figure 3. The wide range of Zr/Nb ratios is somewhat surprising, since Graham and Mason (1972) found a limited range (10–32) for this ratio over a variety of meteoritic, lunar, and terrestrial materials. A geochemical coherence between Zr and Y has not previously been remarked upon to our knowledge, yet Figure 3 shows that this coherence for the Allende samples is comparable to that between Zr and Hf. The cause may lie in similar condensation behavior for Zr and Y. Another possibility is that under the highly reducing conditions under which the Allende materials were formed (evidenced by the occurrence of Eu and possibly also Yb in the divalent state), Zr was trivalent, in which case its ionic radius is very close to that of Y and the two elements would tend to be closely associated geochemically. Taylor, et al. (1972) have provided evidence for the presence of trivalent zirconium in lunar rocks.

Quantitative data on the platinum-group elements (Ru, Rh, Pd, Os, Ir, Pt) could not be obtained from the spark source mass spectrometer records, because of the lack of suitable standards. However, these elements were clearly present in Group I and Group III samples and were not seen in Group II and Group IV samples. Estimates based on relative intensities would give the following averages for Group I and Group III samples (in ppm): Pt 10; Ir 7; Os 7; Ru 8; Rh 1; Pd 1; the latter element appears to be depleted relative to the other platinum-group metals, a feature probably related to its having a considerably lower boiling point and hence being less readily condensed. Grossman (1973) has analyzed 16 Allende inclusions for Ir; all except one (which belongs to Group II from its Eu content) have Ir concentrations at 5–11 ppm. Molybdenum, a comparable siderophile element with a very high boiling point, is present in all Group I and Group III samples at a level of 4–14 ppm, is seen in two Group II samples near the lower limit of detection (0.6 ppm), and is present in Group IV samples at about 1 ppm.

In most samples (Table 1) uranium was close to or below the limit of detectability (0.05 ppm); however, at that level it is clearly considerably enriched over the chondritic mean (0.01 ppm). Uranium

FIGURE 2.—The relationship between CaO and Al_2O_3 (weight percent) in individual samples.

is enriched in Groups I, II, and III, but not in Group IV. Thorium shows a similar pattern.

Discussion

Several other investigators have published data on minor and trace elements in materials from the Allende meteorite, and it is of interest to correlate their data with those presented here. Unfortunately, in most instances no mineralogical or petrological description accompanies the chemical data, which makes the assignment of type (chondrule or aggregate) uncertain. However, the large melilite chondrules are the most prominent and most easily extracted inclusions in the Allende meteorite, and it is evident that sampling has been biased towards these (as is true for the present research also).

Gast, Hubbard, and Wiesmann (1970) provided K, Rb, Sr, Ba, and lanthanide data for what they described as a Ca-rich inclusion from Allende. The lanthanide distribution matches that for Group I, as does the high Sr and Ba contents, and their material was probably a melilite chondrule. Tanaka and Masuda (1973) have published Ca, Al, Sr, Ba, and lanthanide data for the Allende meteorite and some selected components. Their Ca/Al-rich chondrule was a melilite chondrule, and its lanthanide distribution pattern closely parallels that from Group I. They analyzed two samples of a finegrained aggregate; the patterns are closely similar and agree with that for Group II. An olivine chondrule gave an unfractionated lanthanide distribution pattern at approximately twice chon-dritic abundance, consistent with Group IV. Only one of their analyzed samples does not fit readily within the grouping proposed here; their inclusion O has a lanthanide distribution pattern rather similar to Group III, but the negative Eu and Yb anomalies are much less pronounced.

The most extensive set of data so far published

FIGURE 3.—The relationship between Zr and Y in individual samples.

is that of Grossman (1973) on 16 specimens of Ca/Al-rich inclusions selected by Dr. E. A. King, Jr., of the University of Houston. Grossman determined Fe, Mn, Na, Sc, La, Sm, Eu, Yb, Co, Ir, and Au in these specimens and on two samples of the bulk meteorite. Through the kindness of Dr. King, we received splits of specimens 1, 4, and 8 analyzed by Dr. Grossman. These had been selected as covering a wide compositional range, and our analyses show that sample 1 belongs to Group I, sample 4 to Group III, and sample 8 to Group II. Grossman’s data for the remaining samples indicate that they all belong to Group I.

Perhaps the most intriguing feature revealed by the present research is the occurrence of at least three different lanthanide distribution patterns in the materials comprised under the description of Ca/Al-rich inclusions. Although Groups I, II, and III do show characteristic differences in major element composition, these differences are relatively much less than those for the trace elements. The presence of marked Yb anomalies in Groups II and III, and of a Tm anomaly in Group II appears to be unique to these Allende materials; to our knowledge Yb and Tm are closely coherent with the neighboring lanthanides both in terrestrial and in lunar materials.

It is remarkable at first glance that in Group II a negative Eu anomaly is associated with a positive Yb anomaly, whereas in Group III a negative Eu anomaly is associated with a negative Yb anomaly. However, when the Eu and Yb abundances in individual samples in these two groups are compared, one finds that the Eu/Yb ratio is relatively constant (Table 3). Figure 1 shows that the Yb anomaly is positive in Group II because it is superimposed on a fractionated lanthanide pattern with rapidly diminishing abundances of the heavier elements, whereas it is negative in Group III because it is superimposed on an unfractionated lanthanide pattern in which all the other elements (except Eu) are strongly enriched.

Thus in Groups II and III Eu and Yb show a geochemical coherence unrelated to the other lanthanides. After Eu, Yb is the lanthanide most readily reduced to the divalent state, and this provides a possible explanation for this coherence. However, this explanation requires highly reducing conditions at some stage in the origin of Group II and HI materials, more reducing, for example, than for lunar rocks, which show Eu anomalies but do not have Yb anomalies. In this connection, too, it is puzzling to find no Yb anomalies either in Group I materials or their constituent minerals (Mason and Martin, 1974). Boynton (1975) comments that it seems unlikely that divalent Yb can explain the Yb anomalies, since Sm is nearly as easily reduced as Yb, and there is no evidence of an Sm anomaly. He points out that the condensation of lanthanides from the solar nebula may be controlled by thermodynamic equilibrium between gas and condensed solids, and that highly fractionated lanthanide patterns may result if condensates are removed from the gas before condensation is complete. Both Yb and Eu are predicted to be extremely depleted in the early condensate without the requirement of condensation in the divalent state. According to Boynton’s model, the Group II inclusions may be a condensate from a previously fractionated gas rather than from a gas of solar composition. Thus the cosmochemical properties of Eu and Yb (determined by gas-solid equilibria) may be quite different from crystallochemical properties (determined by liquid-solid and solid-solid equilibria), and may allow an unambiguous determination of which process is yielding a specific lanthanide pattern.

The most remarkable and thought-provoking anomaly is the positive Tm anomaly in Group II samples. Group II samples 37 and 3598 were analyzed in 1971, when no attempt was made to measure Tm because of possible interference on the 169 mass number. In 1973 it was realized that the 169 mass number line in many Allende samples was much too strong to be due to the multiple carbon interference, and it was found that in Group I samples Tm values were obtained that fell on the smooth chondrite-normalized curve linking Er and Yb. However, the Tm values for Group II samples are anomalously high in comparison to Er and Yb (Figure 1). A similar Tm anomaly in some Allende aggregates has recently been reported by Conrad, Schmitt, and Boynton (1975). The cause of this Tm anomaly remains to be elucidated. Boynton (1975) has discussed rare earth anomalies in terms of fractional condensation in the solar nebula, and this may be invoked as a possible cause. Another possibility, suggested by the search for evidence for the original presence of superheavy elements in the Allende meteorite, is that the Tm excess may be a previously unrecognized fission product.

Marvin, Wood, and Dickey (1970), who first described the Ca/Al-rich inclusions in the Allende meteorite, noted the general similarity between their chemistry and mineralogy and the sequence of compounds calculated to be among the highest temperature condensates to form in a solar nebula. Grossman (1972, 1973) has greatly extended this work; he concluded (1973:1137) that the Ca/Al-rich inclusions in the Allende meteorite are enriched in Sc, Ir, and the lanthanides by a factor of 22.8 ± 2.2 relative to Type I carbonaceous chondrites, and that the large enrichments of these refractory trace elements provided strong evidence that the inclusions are high-temperature condensates from the solar nebula. All but two of the samples he analyzed belong to Group I, and his conclusion appears well-founded for this group. Can this conclusion be extended to Groups II and III, or do they have a different origin? The trace element distribution in the Group III samples is not greatly different from Group I, except for the negative Eu and Yb anomalies in the lanthanide distribution, and these two groups may well be genetically related. However, the strongly fractionated lanthanide distribution in Group II samples, the Eu, Tm, and Yb anomalies, and their relative depletion in other refractory elements such as Zr, Mo, and the platinum metals, strongly suggest extensive chemical differentiation of their parent material.

One aspect still insufficiently explored is the possible correlations between mineralogy, texture, and chemistry of the Allende inclusions. Group I so far comprises only melilite-rich chondrules; Group II comprises four aggregates and one chondrule; the two representatives of Group III are both aggregates; Group IV comprises the common olivine-rich chondrules, but also includes some aggregates. The term aggregate covers a variety of fine-grained white, pink, or pale gray inclusions with irregular form; if an aggregate approaches spherical form, the distinction between aggregate and chondrule may be arbitrary. The form of the aggregates and their contacts with the enclosing matrix vary greatly; some are angular fragments with sharp contacts, many are irregular in form, with intricate mossy-like margins against the matrix, others are elongated shards or stringers. Subparallel elongation of the latter type in a hand specimen often indicates a vaguely oriented fabric, suggesting an original flattening during deposition (as plastic glass fragments?) or a secondary flattening produced by later deformation. The overall impression is that some aggregates may be devitrified glass fragments, whereas others represent agglutinated dust particles.

The formation of chondrules, specifically the melilite-rich chondrules of Group I, presents an additional problem. Grossman’s condensation curves (1972) are for gas-solid equilibria, and if the Group I material originated in this way, it should have separated as a dust; the formation of chondrules involved segregation of this dust and subsequent melting, by such processes as lightning or shock effects. Alternatively, these chondrules may have originated in some nonequilibrium process, such as that suggested by Blander and Katz (1967): metastable nucleation of liquid droplets, formed by super-saturation of the nebular gas, component by component, as subcooling took place below the temperatures where equilibrium solids should have crystallized out.

TABLE 3.—Abundances (ppm) of Eu and Yb in Group II and Group III samples, and the Eu/Yb ratios

The mineralogy and chemistry of the Group I materials certainly support the concept that they represent an early high-temperature condensate. The condensation sequence established by Grossman shows that this would be followed by forsteritic olivine and enstatite, essentially the material of the Group IV chondrules and aggregates. The position of the Group II and Group III materials in this sequence is more enigmatic. In minor and traceelement concentrations, Group III materials are not very different from Group I, the most marked feature being the negative Eu and Yb anomalies. In view of the small number of Group III specimens, further speculation on their origin may be deferred. Group II materials are in marked contrast to Groups I and III in minor and trace elements, with relatively high alkalies, low refractories (Zr, Hf, Mo, and platinum metals), and especially the strongly fractionated lanthanide pattern. Evidently the medium from which this material condensed had undergone or was undergoing considerable fractionation of these elements. So far a complementary condensate, enriched in the heavier lanthanides, has not been identified; however, since Group II material forms only a small amount of the total meteorite (probably about 5% or less), the complementary enriched material may be too dispersed to be recognizable. Possibly Group II materials represent a late condensate, subsequent to the segregation of Group I and Group IV. The Group IV material is almost entirely magnesium-iron silicates; separation of this material from the parent medium would enrich the latter in other major elements, among which would be Ca, Al, and the alkalies. Since olivine is a poor host for most trace elements, these too would remain in the parent medium and increase in concentration. Pyroxene is a mineral which preferentially concentrates the heavier lanthanides, and its separation along with olivine would lead to a depletion of these elements in the residual material. In this connection the composition of interstitial glass in the olivine-pyroxene chondrules reported by Clarke, et al. (1971) is revealing; this glass composition is (weight percent) SiO_2, 47.8; Al_2O_3, 24.7; FeO, 0.7; MgO, 6.3; CaO, 17.5; Na_2O; 2.0; K_2O, 0.2, a composition similar in many respects to the bulk composition of Group II aggregates.

This discussion thus suggests at least two periods of segregation for the Ca/Al-rich inclusions, one before the separation of olivine and one following it. Much additional work needs to be done to verify this, and to fill out the details of the process. Clearly, however, the Allende meteorite preserves a remarkable record of the origin of chondrules and chondrites, but one which requires extensive and intensive research to decipher.

Literature Cited

Ahrens, L. H., and H. von Michaelis

1969.    The Composition of Stony Meteorites, III: Some Inter-element Relationships. Earth and Planetary Science Letters, 5:395–400.

Blander, M., and J. L. Katz

1967.    Condensation of Primordial Dust. Geochimica et Cosmochimica Acta, 31:1025–1034.

Boynton, W. V.

1975.    Fractionation of the Solar Nebula: Condensation of Yttrium and the Rare Earth Elements. Geochimica et Cosmochimica Acta, 39:569–584.

Clarke, R. S., Jr., E. Jarosewich, B. Mason, J. Nelen, M. Gómez, and J. R. Hyde

1971.    The Allende, Mexico, Meteorite Shower. Smithsonian Contributions to the Earth Sciences, 5:1–53.

Conrad, R. L., R. A. Schmitt, and W. V. Boynton

1975.    Rare-Earth and Other Elemental Abundances in Allende Inclusions. Meteoritics, 10:384–387.

Gast, P. W., N. J. Hubbard, and H. Wiesmann

1970.    Chemical Composition and Petrogenesis of Basalts from Tranquillity Base. Proceedings of the Apollo 11 Lunar Science Conference, 2:1143–1163.

Graham, A. L., and B. Mason

1972.    Niobium in Meteorites. Geochimica et Cosmochimica Acta, 36:917–922.

Gray, C. M., D. A. Papanastassiou, and C. J. Wasserburg

1973.    Primitive 87Sr/86Sr in the Allende Carbonaceous Chondrite. Icarus, 20:213–239.

Grossman, L.

1972.    Condensation in the Primitive Solar Nebula. Geochimica et Cosmochimica Acta, 36:597–619.

1973.    Refractory Trace Elements in Ca-Al-rich Inclusions in the Allende Meteorite. Geochimica et Cosmochimica Acta, 37:1119–1140.

Marvin, U. B., J. A. Wood, and J. S. Dickey, Jr.

1970.    Ca-Al Rich Phases in the Allende Meteorite. Earth and Planetary Science Letters, 7:346–350.

Mason, B.

In press. Meteorites. Chapter B–1 in The Data of Geochemistry, sixth edition. U.S. Geological Survey Professional Paper, 440–B, Part 1.

Mason, B., and P. M. Martin

1974.    Minor and Trace Element Distribution in Mellilite and Pyroxene from the Allende Meteorite. Earth and Planetary Science Letters, 22:141–144.

Osborn, T. W., R. G. Warren, R. H. Smith, H. Wakita, D. L. Zellmer, and R. A. Schmitt

1974.    Elemental Composition of Individual Chondrules from Carbonaceous Chondrites, Including Allende. Geochimica et Cosmochimica Acta, 38:1359–1378.

Tanaka, T., and A. Masuda

1973.    Rare-Earth Elements in Matrix, Inclusions, and Chondrules of the Allende Meteorite. Icarus, 19:523–530.

Taylor, S. R.

1965.    Geochemical Analysis by Spark Source Mass Spectrography. Geochimica et Cosmochimica Acta, 29:1243–1261.

1971.    Geochemical Application of Spark Source Mass Spectrography, II: Photoplate Data Processing. Geochimica et Cosmochimica Acta, 35:1187–1196.

Taylor S. R., M. Kaye, P. Muir, W. Nance, R. Rudowsky, and N. Ware

1972.    Composition of the Lunar Uplands: Chemistry of Apollo 14 Samples from Fra Mauro. Proceedings of the Third Lunar Science Conference, 2:1231–1249.

Wänke, H., H. Baddenhausen, H. Palme, and B. Spettel

1974.    On the Chemistry of the Allende Inclusions and Their Origin as High Temperature Condensates. Earth and Planetary Science Letters, 23:1–7.

Wetherill, G. W., R. Mark, and C. Lee-Hu

1973.    Chondrites: Initial Strontium-87/Strontium-86 Ratios and the Early History of the Solar System. Science, 182:281–283.

The St. Mary’s County, Maryland, Chondrite

Albert F. Noonan, Eugene Jarosewich, and Roy S. Clarke, Jr.

ABSTRACT

The St. Mary’s County, Maryland, meteorite fell on an afternoon in June of 1919. One individual stone of perhaps 500 grams is known to have been recovered, but only a 23.4 gram fragment was preserved. A 2.9 gram piece was made available for laboratory investigations.

Chemical parameters determined from the bulk chemical analysis, compositions of the major silicates, and the overall texture of the meteorite show it to be an LL3 chondrite. The average compositions of olivine, 14.4 (Fa, 0.2–76.6), and low-Ca pyroxene, 9.9 (Fs, 0.5–37.5), with % MD’s of 51.9 and 67.3, respectively, show that this meteorite is among the least equilibrated of the ordinary chondrites. Calcium and manganese concentrations in olivine plotted against fayalite content show mixing of an Fe-rich matrix fraction with one and possibly two chondrule fractions. Oxide plots of minor elements in low-Ca pyroxene (<1% CaO) suggest a general increase in chromium, manganese, and possibly calcium and aluminum concentration with increasing iron content.

Albert F. Noonan, Eugene Jarosewich, and Roy S. Clarke, Jr., Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D. C. 20560.

Introduction

A bright fireball accompanied by a lingering smoke trail and a series of violent detonations that shook buildings and windows over a wide area announced the fall of the Saint Mary’s County, Maryland, meteorite to hundreds of residents of tidewater Virginia and Maryland on a June afternoon in 1919. It approached from the southwest at about 4 PM, attracting the attention of retired sea-captain John Forrest, who was working in a tobacco field adjoining his home along St. Jerome Creek, near the village of Ridge, St. Mary’s County, Maryland. An object that later proved to be a meteorite struck the ground approximately 20 feet [6 m] from where he stood, producing a cloud of dust. Captain Forrest removed the stone from the ground, and believing it to be hot carried it on his hoe to the house. This meteorite specimen, most of which has since been lost, was the only one recovered. The recovery site (approximately 38°10’N, 76°23’W) is on a peninsula between the Potomac River and the Chesapeake Bay, and it was felt at the time that other specimens may have landed in the adjoining waters, particularly in the Chesapeake Bay. The exact date of fall is not known, but it was within a few days of 20 June 1919.

The main mass of the meteorite, probably some 500 grams or more, remained at Captain Forrest’s home serving as a doorstop, and his widow and son reported seeing it as late as 1933. Since that time, however, it has disappeared without a trace, and all of the people involved with the recovery have left the area. The surviving piece was broken off the main mass by Captain Forrest within a few days of the fall and was given to William Matthews of Great Mills, Maryland. This 24.25 gram specimen came into the possession of Matthews’ grandson, Francis D. Cecil, who later loaned it to the Maryland Academy of Sciences, Baltimore, Maryland. The fragment was shown to Dr. William F. Foshag of our staff and identified as a meteorite in 1937. The recovery of the meteorite and a description of the surviving specimen have been recorded by Mather (1938, 1954) and Cecil (1944).

Early in 1969 the pyramidal specimen, fusion-crusted on three of its four sides, was borrowed from the Maryland Academy of Sciences through the courtesy of Paul S. Watson. The specimen weighed 23.4 grams when it was received at the Smithsonian Institution. It was photographed, a model was made, and a 2.9 gram piece was removed for analysis. A polished section and two polished thin sections were prepared for optical and elecron-microprobe examination (NMNH 5423), and the remaining material was used for a bulk chemical analysis. A 19.7 gram piece was returned to the Maryland Academy of Sciences.

Texture

The St. Mary’s County meteorite is a black, matrix-rich chondrite of low petrographic type (Figure 1). In thin section (0.03 mm), the matrix is opaque and contains small angular silicate crystals and finely dispersed metal and sulfide particles. The predominant inclusions are broken and intact olivine and/or low-Ca pyroxene chondrules, which range up to 0.6 mm in diameter. These frequently contain primary albitic glass and metal droplets. In addition, several metal-sulfide chondrules were observed, but these phases usually occur in chondrule rims and as small inclusions in chondrules and in the matrix.

FIGURE 1.—Overall chondritic texture of the St. Mary’s County (LL3) meteorite; length of specimen, 19 mm.

Chemistry

The bulk chemical analysis of the St. Mary’s County meteorite is presented in Table 1. It was performed on a 1.5 gram sample that appeared to be reasonably representative of the 23.4 gram specimen. The procedures of Jarosewich (1966) were used with the modification that the portion taken for major constituent analysis was completely dissolved, rather than analyzing an acid-soluble and acid-insoluble fraction. The carbon content is high compared to other unequilibrated ordinary chondrites; only Sharps is reported to have a higher carbon concentration of 0.95% (Fredriksson, Jarosewich, and Nelen, 1969). The SiO_2/MgO ratio of 1.64 is typical of ordinary chondrites and the Fe/SiO_2 and Fe°/Fe ratios of 0.50 and 0.14 are within the range recognized for LL-group chondrites (Van Schmus and Wood, 1967).

TABLE 1.—Chemical analyses of the St. Mary’s County chondrite

Mineralogy

The analyses of the St. Mary’s County silicates were made with the ARL microprobe model SEMQ, utilizing an accelerating potential of 15 KV and a beam current of 0.15 μA. Results have been corrected using procedures outlined by Bence and Albee (1968) and the correction factors of Albee and Ray (1970). Two spots separated by a 1–2 micron spacing were counted for 10 seconds and averaged for each analyzed area. An advantage of our microprobe is that nine elements can be determined simultaneously, thus providing minor element data from spots that are analyzed for major elements.

OLIVINE

Olivine was analyzed for Si, Al, Fe, Mg, Ca, K, Na, Cr, and Mn. The primary mineral standard used for major element determinations was Springwater olivine with 13.26% Fe. The minor elements calcium and manganese were determined using An 80 and Rockport fayalite with 11.64% Ca and 1.30% Mn, respectively. Springwater olivine was analyzed routinely every twenty samples to provide an internal check on instrument stability. Iron concentrations of those analyses with oxide summations between 98%–101% were applied to a calculated calibration curve to establish the fayalite content of each well-analyzed grain.

A total of 269 olivine analyses were made, which includes analyses of crystals from 137 separate chondrules and 82 matrix crystals. For all olivine analyses, the average iron content () is 10.4%, the average fayalite content () is 14.4 (range, Fa, 0.2–76.6) and the percent mean deviation (%MD) is 51.9. (The % MD is the parameter used by Dodd, Van Schmus, and Koffman, 1967, as an index of olivine and pyroxene heterogeneity.) These results compare well with the interpretation of an x-ray diffractometer trace of the meteorite in which the

FIGURE 2.—Plot of CaO concentration vs fayalite content of olivine (lines join compositions in zoned crystals). Note particularly the high to low CaO values of the Mg-rich olivine in metal-rich chondrules, possibly representing a separate chondrule fraction, and the relatively Fe-rich fragments from the matrix. Only one olivine fragment from the matrix with the composition Fa 76.6 and containing 0.14% CaO is not represented in the plot.

olivine peak (130) indicates a broad range in composition from Fa 0 to Fa 80 with a small peak at Fa 12 (B. Mason, pers. comm.). The shape of the peak shows that most of the olivine is Mg-rich, in the range Fa 0–15, with diminishing amounts of more Fe-rich material. Compared with olivine compositions from other unequilibrated chondrites (Dodd, Van Schmus, and Koffman, 1967), St. Mary’s County is clearly the least equilibrated of the ordinary chondrites. Although the major element analyses and x-ray diffraction data are useful for classification purposes, they do not provide clues as to the number or origin of olivine fractions that may exist in a particularly unequilibrated chondrite like St. Mary’s County, information which may become apparent with minor element data.

When CaO and MnO concentrations in olivine from relatively metal-rich and metal-poor chondrules and from silicate fragments in the matrix are plotted against fayalite content (Figures 2 and 3), at least two important points become apparent. Firstly, silicate inclusions in the matrix are not simply fragments of crushed chondrules. The olivine has a mean fayalite content of 21.2, which, if the silicates are of primary composition, cannot

FIGURE 3.—Plot of MnO concentration vs fayalite content of olivine. Only one olivine fragment from the matrix with the composition Fa 76.6 containing 0.46% MnO is not represented in the plot.

be derived solely from the metal-rich and metal-poor chondrule olivine with mean fayalite contents of 3.0 and 14.1, respectively. Indeed, several abraded chondrules with forsteritic olivine adjacent to matrix material suggest that all silicates do reflect primary compositions. Therefore, although silicates in the matrix may include olivine fragments derived from chondrules similar in composition to those studied, it must contain a separate Fe-rich olivine fraction. Secondly, olivine in metal-rich chondrules is enriched in Ca and depleted in Mn relative to olivine in metal-poor chondrules, possibly suggesting two separate chondrule fractions. The high-Ca olivine is Mg-rich (Fa, 0–10), occurs predominantly in metal-rich chondrules, and trends from a high of 0.67% CaO for nearly pure forsterite to below 0.10% CaO at about Fa 10 (Figure 2). This grouping corresponds to the proposed high-temperature metal-rich chondrule fraction in Sharps (Dodd, 1971). This fraction is depleted in MnO but shows a positive rather than a negative correlation with fayalite content, also to about Fa 10 (Figure 3). Olivine with greater than about 8% Fa (Fa 10), which includes most metal-poor chondrules and silicate fragments from the matrix, shows no correlation of either CaO or MnO concentration with fayalite content.

Most of the compositional zonation displayed in

FIGURE 4.—Plot of MnO vs ferrosilite content of low-Ca pyroxene (< 1% CaO) suggests an increase in manganese with increasing iron (lines join compositions in zoned crystals).

these olivines is normal, with Mg-rich cores and relatively Fe-rich rims. One single crystal metal-poor chondrule, however, was observed with a core of Fa 6.2 and a rim of Fa 0.2.

PYROXENE

Pyroxene was analyzed for Si, Al, Fe, Mg, Ca, K, Mn, Ti, and Cr. The primary mineral standards used for pyroxene analyses were Johnstown hypersthene (Si, 25.06%; Fe, 12.14%; Mg, 16.40%) and an omphacite standard (Al, 4.71%; Ca, 9.83%). The more abundant minor elements, manganese and chromium, were determined using Rockport fayalite (Mn, 1.30%) and a chromite standard (Cr, 40.00%). Johnstown hypersthene was analyzed every twenty samples to provide a check on instrumental drift, which was negligible. Iron concentrations of the low-Ca (<1% CaO) pyroxenes with oxide summations between 97%–101% were applied to a calculated calibration curve to establish the ferrosilite content of each grain.

A total of 274 separate analyses were made of pyroxene in 23 metal-rich and 67 relatively metal-poor chondrules and from 101 silicate fragments in the matrix. Of these, 222 analyses are of the more abundant low-Ca (<1% CaO) pyroxene, mostly clinoenstatite and clinobronzite, and 52 analyses

FIGURE 5.—Binary oxide plot of Cr_2O_3 vs MnO suggests that chromium increases as does manganese with increasing iron content of pyroxene (cf. Figure 4).

of more calcic pyroxene (1–18% CaO), which occurs both as fragments in the matrix and as rims on chondritic low-Ca pyroxene. The average iron content () of the low-Ca pyroxene is 5.2%, the average ferrosilite content () is 9.9 (range, Fs 0.5–37.5) and the percent mean deviation (%MD) is 67.3, which is a significantly higher value than has been reported for other unequilibrated ordinary chondrites (Dodd, Van Schmus, and Koffman, 1967). As with the olivine analyses, iron determinations are useful for classification purposes and comparative studies, but minor element data may help distinguish possible pyroxene fractions.

Binary oxide plots were prepared of Al_2O_3, CaO, MnO, and Cr_2O_3 values against ferrosilite content of low-Ca pyroxene. When MnO concentrations are plotted against ferrosilite content (Figure 4), it becomes evident that pyroxene from metal-rich chondrules is Mg-rich and low in MnO relative to both fragments in the matrix and metal-poor chondrule pyroxene. But unlike the olivine, both the Mg-rich fraction and the pyroxene in metal-poor chondrules and the matrix overlap to a large measure and show a continuous positive correlation of MnO with ferrosilite. Therefore, the pyroxene data provides no compelling evidence to suggest that metal-rich chondrules and metal-poor chondrules represent two distinct chondrule fractions although that possibility may be supported by the olivine data. Also of interest in the MnO vs plot is the extensive mixing of points representing matrix analyses ( = 9.8) with pyroxene analyses from both metal-rich ( = 3.0) and metal-poor ( = 12.6) chondrules. The pyroxene fragments in the matrix clearly could be fragments of pyroxene derived from chondrules similar to the compositions studied.

Of the other minor elements determined, a Cr_2O_3 vs MnO binary plot shows that a positive correlation exists (Figure 5), both elements apparrently increasing in concentration with Fe. No clear trends are apparent for Al or Cr against ferrosilite content. However, Cr_2O_3 against CaO and Al_2O_3 plots show a subtle positive correlation suggesting that calcium and aluminum also increase in concentration to some degree with increasing iron.

LESS ABUNDANT MINERALS

The less abundant minerals identified in the St. Mary’s County meteorite include small amounts of taenite, kamacite, troilite, and chromite, and minor amounts of plagioclase, pentlandite, copper, spinel, and a phosphate mineral, probably whitlockite.

The feldspathic fraction occurs both as primary albitic glass, which is commonly found in ordinary chondrites, and as crystalline calcic plagioclase. Two abraded fragments were found to contain twinned plagioclase with compositions in the range An 80–83, which occur with low-Ca, Fe-Mg silicates. Both fragments appear equally medium-grained; however, one occurs with equilibrated olivine (Fa 13) and pyroxene (Fs 12) while the other contains less equilibrated and higher iron silicates with compositions in the range Fa 23–25 and Fs 15–18 for olivine and pyroxene, respectively. One plagioclase matrix fragment was found with a composition An 26.

The most commonly occurring opaque minerals are troilite, taenite, and kamacite. Typical of most unequilibrated ordinary chondrites and some carbonaceous chondrites, globular metal grains or chondrules rimmed with troilite were observed. The taenite in these features contains from 30%–50% Ni and frequently displays the familiar M-shaped Ni-concentration profile when in contact with kamacite. Two pentlandite grains with 16% Ni were found in troilite with taenite and kamacite and presumably represent disequilibrium assemblages, which are not often found in unequilibrated ordinary chondrites (Dodd, Van Schmus, and Koffman, 1967). In addition to the large inclusions, metal also occurs in small, occasionally sulfide-rimmed droplets in glassy chondrules. Scanning-electron pictures of Na, Ca, and P distribution in these 10 micron and less diameter features suggests that phosphate nucleation, probably whitlockite, has occurred at the metal-silicate interface, or, if rimmed with sulfide, at the metal-sulfide interface.

Literature Cited

Albee, Arden L., and L. Ray

1970.    Correction Factors for Electron Probe Microanalysis of Silicates, Oxides, Carbonates, Phosphates, and Sulfates. Analytical Chemistry, 42:1408–1414.

Bence, A. E., and A. L. Albee

1968.    Empirical Correction Factors for the Electron Microanalysis of Silicates and Oxides. Journal of Geology, 76:382–403.

Cecil, Francis D.

1944.    The St. Mary’s Meteorite of 1919. Sky and Telescope, 3(12):9.

Dodd, R. T., Jr.

1971.    The Petrology of Chondrules in the Sharps Meteorite. Contributions to Mineralogy and Petrology, 31:201–227.

Dodd, R. T., Jr., W. R. Van Schmus, and D. M. Koffman

1967.    A Survey of the Unequilibrated Ordinary Chondrites. Geochimica et Cosmochimica Acta, 31(6):921–951.

Fredriksson, K., E. Jarosewich, and J. Nelen

1969.    The Sharps Chondrite—New Evidence on the Origin of Chondrules and Chondrites. Pages 155–165 in P. Millman, Meteorite Research. Dordrecht: D. Reidel Publishing Company.

Jarosewich, E.

1966.    Chemical Analysis of Ten Stony Meteorites. Geochimica et Cosmochimica Acta, 30:1261–1265.

Mather L. Bryant, Jr.

1938.    A New Meteorite from Maryland. The Natural History Society of Maryland Bulletin, 8(8):63–68.

1954.    Maryland Meteorites. Maryland Naturalist, 24(1–4):2–18.

Van Schmus, W. R., and J. A. Wood

1967.    A Chemical-petrologic Classification for the Chondritic Meteorites. Geochimica et Cosmochimica Acta, 31:747–765.

Mineralogical and Chemical Compositions of Silicate Inclusions in the El Taco, Campo del Cielo, Iron Meteorite

Frank Wlotzka and Eugene Jarosewich

ABSTRACT

The silicate minerals in these inclusions (olivine, orthopyroxene, diopside and a sodic plagioclase) have the same composition as silicates in chondrites except for a lower FeO-content of the mafic minerals and a lower K-content of the feldspar. Small grains of K-feldspar occur with the plagioclase. Olivine contains 3.5% to 4.5% FeO, orthopyroxene 3.8% to 4.6% FeO. The range of compositions within one inclusion is smaller, however, and distinct differences in mean FeO-content occur between different inclusions; the fayalite to ferrosilite ratio varies from inclusion to inclusion. The outer zone of olivine and pyroxene grains is depleted in iron oxide compared to the central part. Thus three types of disequilibria occur: (1) disequilibrium between different inclusions, (2) disequilibrium between olivine and pyroxene in a given inclusion, and (3) disequilibrium within single olivine and pyroxene grains. Both the mineral composition and the bulk chemistry of the silicate part of the inclusions are close to normal chondrites, if compared on an iron-free basis. Variations that do not seem to be caused by nonrepresentative sampling are found between different inclusions. The variations affect mainly the elements Ca, Al, and the alkalies. They can be explained by the removal of diopside and/or feldspar from a chondritic material, perhaps through partial melting. Other features suggest that the lower iron content of the mafic minerals was produced by reduction of chondritic silicates.

Frank Wlotzka, Max-Planck-Institut für Chemie, Mainz, Germany. Eugene Jarosewich, Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560.

Introduction

The El Taco meteorite is one of the Campo del Cielo, Argentina, irons. It was recovered by an expedition described by Cassidy, et al. (1965). The El Taco specimen, which weighed 2000 kg, was cut at the Max-Planck-Institut für Chemie at Mainz, Germany (Hintenberger, Schultz, and Weber, 1969). Three cuts were made parallel to the longest axis of the meteorite, so that two slices of about 2 cm thickness and two end pieces resulted. One of the slices (slice A, Figure 1) is on exhibit in the Meteorite Hall of the National Museum of Natural History, and the other (slice B) was made available for research by the Max-Planck-Institut für Chemie, Mainz.

Slice A revealed a large number of silicate inclusions (Figure 1). One inclusion of this kind had already been found in the El Toba specimen of Campo del Cielo and described by Park, Bunch, and Massalski (1966). Silicate inclusions are known from other iron meteorites, for instance, Linwood (Henderson and Perry, 1949), Woodbine (Mason, 1967), Netschaevo (Olsen and Jarosewich, 1971), and Odessa (Marshall and Keil, 1965). A thorough description of the mineral compositions of the silicate inclusions from most of the iron meteorites known to contain them was given by Bunch, Keil, and Olsen (1970). All these iron meteorites (except Netchaevo) belong to the same chemical group IAB of Wasson(1970).

In continuation of earlier work (Wlotzka and Jarosewich, 1969), this paper describes the mineralogy and the bulk chemical composition of a

FIGURE 1.—El Taco slice A, etched. Silicate inclusions appear black. Long axis of section, 111 cm.

number of the El Taco silicate inclusions in more detail, with special emphasis on possible variations between different inclusions.

ACKNOWLEDGMENTS.—We wish to thank Professor Hintenberger for providing a facility for an ingenious cutting of this large specimen of the El Taco meteorite, Dr. H. Wänke and Dr. K. Fredriksson for their stimulating discussions, Mr. R. S. Clarke for selecting and providing some samples, and Dr. B. Mason for his suggestions in the preparation of the manuscript. One of us (F. W.) wishes to thank the members of the Department of Mineral Sciences of the Smithsonian Institution for their hospitality and help during his stay there, especially Mrs. P. R. Brenner, Mrs. C. Trice, Mr. C. Obermeyer, Mr. J. Nelen, and Mr. A. F. Noonan. We also thank Mr. G. Moreland for the preparation of excellent polished thin sections. The completion of this study was made possible through a special fund provided by the Secretary of the Smithsonian Institution.

Macroscopic Features

Campo del Cielo is a coarse octahedrite. Figure 1 shows the Widmanstätten pattern of the etched slice. The silicates are scattered over the whole area of the sectioned slice. Most of them are located at grain boundaries of the primary taenite crystals, but several appear also in the centers of these

FIGURE 2.—Silicate inclusions of slice A.

crystals. About nine different primary taenite crystals can be recognized in slice A (Figure 1). There is no disturbance of the Widmanstätten structure, i.e., flow or shock features, visible either around the silicates or in other parts of the metal areas, as has been reported for smaller Campo del Cielo specimens by Bunch and Cassidy (1968). This indicates clearly that these inclusions were incorporated into the metal before the taenite-kamacite transformation took place, or at temperatures above about 750° C.

The outline of most of the silicate inclusions is angular with usually straight margins. This is suggestive of a fragmental nature of these inclusions. It can be noted, however, that several of the straight boundaries are parallel to kamacite laths in the Widmanstätten pattern and may be secondary features. This point is difficult to prove, and most observers fail to see it. It seems, however, that some of the silicate-metal boundaries and indentations in the silicate inclusions are controlled by the Widmanstätten pattern or the growth directions of the kamacite laths. Examples are shown in Figure 2. The microscopic texture of the silicates shows replacement of silicates by opaques (mostly by graphite, which also seems to replace troilite and metal) in many places, which could explain such a reshaping of the silicate inclusions. In contrast to other iron meteorites with silicate inclusions, for instance, Woodbine (Mason, 1967), no swathing kamacite is found around the inclusions of El Taco. Macroscopically the inclusions differ in the relative amounts of graphite and opaques and in the grain size of the silicates.

FIGURE 3.—Line drawing of part of El Taco slice B. Silicate inclusions are hatched. Open bars in inclusions 2 and 3 give locations of samples taken for chemical analysis, lines 1, 3, 4, 7, and 9 give locations of the corresponding polished thin sections. Inclusion 6 was a veinlike inclusion which was exposed along the indicated cutting line and from which section 6 was made. Width of cut, 6 cm.

Sampling

Inclusion 1 was taken from the surface of El Taco. A piece of 3.5 grams was used for chemical analysis; from another piece, polished thin section (PTS) 2 was made. Inclusions 2 and 3 were obtained from the research slice B, and their locations and the sampling are shown in Figure 3. PTS 6 was made from a small inclusion (No. 6, Figure 3). Additional thin sections were made from inclusions that were exposed during the cutting of slices A and B. These are: location A I–5/6: inclusion 4, PTS 10; location A–15/18: inclusion 5, PTS 5; location A I–4/7: inclusion 8, PTS 8; location B III-II/7: inclusion 12, PTS 12; and location not known: inclusion 11, PTS 11.

Microscopic Texture

Between the metal host and the silicate inclusion, a border zone consisting of schreibersite, troilite, and graphite is found. Figure 4 shows an example: schreibersite usually forms a thin band against the metal, followed and partly replaced by troilite and graphite. In some places a very fine-grained eutectoid intergrowth of metal and troilite and/or sphalerite is found. The same type of border zone was described by Park, Bunch, and Massalski (1966) for the El Toba silicate inclusion.

The silicate portions of the inclusions differ considerably in grain size and distribution of opaques and each inclusion has its own texture. Two main types of textures can be distinguished: (1) large silicate crystals set in a matrix of graphite or sometimes troilite and metal (Figure 5), and (2) allotriomorphic intergrowths of smaller, equigranular grains (Figure 6).

The textural relationships of the opaque minerals graphite, troilite, and metal with one another and with the silicates are more complex and will be described in a separate paper (Wlotzka, in prep.). The conclusions of this study are that graphite is a late addition in these inclusions and replaces metal, troilite, and silicates in the texture of type 1 mentioned above. In type 2, graphite (if present) occurs as small flakes dispersed throughout the silicates and does not show tendencies to form veins like in type 1 (Figure 7).

In contrast to mesosiderites and many chondrites and achondrites no brecciation structure is found. The silicates have an igneous texture, and the equigranular

FIGURE 4.—Border zone between metal host and silicate inclusion 8, reflected light. (M = metal, S = schreibersite, T = troilite, G = graphite, with some rounded silicate grains in the lower part, width of image = 1.5 mm.)

and allotriomorphic intergrowth suggests a metamorphic recrystallization. Graphite and sometimes metal veins fill cracks in the silicate fragment in some places. Only in one thin section a straight boundary between two different kinds of structures or “fragments” was found; all other sections studied appear to be homogeneous throughout, except for the development of a coarser texture towards the metal host, with euhedral silicate crystals in a graphite or troilite matrix.

Table 1 gives the main minerals found in three different inclusions and their relative amounts as determined by point counting. Noteworthy are the large variations in the relative amounts of the opaque minerals graphite, troilite, and metal. Where graphite is high, troilite is low, and vice versa. Thus the sum of the opaques and consequently

FIGURE 5.—Coarse silicate texture, inclusion 8: a, transmitted light (black areas are graphite), note graphite “nibbling” at silicate grains; b, reflected light, crossed polarizers.

of the silicates is rather constant. It is interesting that the amount of silicates in these three inclusions of 75.9%, 77.3%, and 74.5% is quite close to the average silicate portion of H-chondrites, which is 76% (Table 7). The silicate portion of two of the three inclusions analyzed chemically is higher, but this does not give the true silicate amount, as coarse metal particles were removed prior to analysis.

In addition to the main minerals listed in Tables 1 and 2, the following accessory minerals were found: schreibersite and cohenite (at the border to the metal host), apatite, chromite, and sphalerite. The apatite is a chlorapatite; the microprobe analysis of 11 grains in different sections gave an average (in weight percent) of CaO, 52.6; P_2O_5; 41.1; Cl, 6.00; FeO, 0.30; MgO, 0.11; Na_2O, 0.46; F is not determined, sum 99.2 (corrected for Cl = 0). This is smiliar to the composition of the Campo del Cielo apatite published by Bunch, Keil, and Olsen (1970). The silicates often contain small inclusions, mostly troilite or metal. Veining with iron oxides formed by terrestrial weathering is found in most specimens.

Mineral Composition

In Table 2 the orthopyroxene, olivine, diopside, plagioclase, and K-feldspar compositions are given for one inclusion as determined by standard procedures with an ARL SEMQ microprobe. The results were corrected by the method of Bence and Albee (1968). During the study of these inclusions it was found that small differences in the Fe/(Fe + Mg)

FIGURE 6.—Fine silicate texture, inclusion 5: a, transmitted light (black vein in upper left is graphite); b, reflected light (white grains in graphite vein are metal, in silicate area troilite).

ratio occur in olivines and pyroxenes of different inclusions. Special care was taken, by repeated measurements, to establish these differences beyond doubt. Table 3 shows the results of repeated measurements of the same grains in two inclusions. Table 4 gives the mean FeO contents for olivines and FeO and CaO contents for orthopyroxenes. The spread in FeO values for a given inclusion is shown in Figure 8.

Two results are evident from Table 3 and Figure 8: (1) The spread of FeO content found within one and the same inclusion is small, but real, as the values are reproducible by repeated measurements. (2) The difference of FeO content in olivines in different inclusions is real and reproducible. The clearest cases are sections 5 and 10, with an average of 4.5 ± 0.2% FeO and 3.6 ± 0.2% FeO respectively. Figure 8 shows that the ranges of FeO content in olivine in these sections do not overlap. Most olivine compositions of the other inclusions center around values of 4% FeO. The pyroxene average values are also distinctly different for different

TABLE 1.—Mineral composition of silicate inclusions 1, 4, and 5, determined by point-counting in weight percent

Mineral

Inclusion 1

Inclusion 4

Inclusion 5

Olivine

33

21

43

Orthopyroxene

35

44

21

Clinopyroxene

2.5

0.8

2.7

Plagioclase

5.4

11.5

7.8

Sum of silicates

75.9

77.3

74.5

Graphite

22

14

0.8

Troilite

0.1

3.4

16.5

Nickel-iron

0.6

5.5

4.9

Iron oxide*

1.7

—

3.2

Total

100.3

100.2

99.9

Points counted

1300

1000

1000

*Formed by terrestrial weathering.

FIGURE 7.—Graphite flakes in silicate area of inclusion 3, reflected light. (Width of image = 280 microns.)

inclusions, and the spread within any one inclusion is larger than for the olivines. Compared to the mafic minerals, the plagioclase is very constant in composition at about 3.0% CaO, except for inclusion 6 with 3.4% CaO and inclusion 8, which shows the largest spread in An content of the feldspar (Table 5).

The outer rim of the mafic minerals is lower in FeO content than the central part. Table 6 gives FeO values for several olivine centers and rims of inclusion 8. In Figure 9 the results of microprobe step-scans across several olivine and pyroxene grains are plotted. The large amounts of Fe just outside the grains are due either to a thin rim of iron oxide (formed by terrestrial weathering) or to troilite adjacent to the silicate grain. The drop in iron measured at the grain boundary, despite the steep rise just next to it, shows that this is a real effect, since any overlapping effect of the electron beam at the grain boundary should have raised the iron reading, not lowered it. Mg is plotted together with iron to show that the measured points are still in the silicate grain. The rise in MgO content compensating the drop in FeO (Table 6) is not visible here, because it is too small relative to the Mg level. In orthopyroxene Ca shows higher values in the grain centers and lower ones towards the rim. There is also a general tendency in the orthopyroxene for high Fe grains to also have high Ca contents, as shown in the CaO/FeO plot in

TABLE 2.—Composition of silicate minerals from inclusion 8, determined by electron microprobe (numbers in parentheses give number of grains analyzed)

Constituent

Orthopyroxene (65)

Olivine (63)

Diopside (33)

Plagioclase (53)

K-Feldspar (2)

SiO_2

58.9

42.00

55.2

64.9

64.0

Al_2O_3

0.25

—

0.74

22.4

19.8

MgO

35.2

53.8

17.9

—

—

FeO

4.37

4.00

1.81

0.15

0.52

CaO

0.79

0.04

21.6

3.24

1.17

TiO_2

0.20

—

0.60

—

—

Cr_2O_3

0.37

0.02

1.24

—

—

MnO

0.39

0.33

0.34

—

—

Na_2O

—

—

0.83

10.0

1.0

K_2O

—

—

—

0.57

14.7

Total

100.4

100.2

100.3

101.2

101.2

FIGURE 8.—Histograms of microprobe FeO values of different inclusions. (The suffix “.1” or “.2” following a section number means that grains in two different areas were measured; results from 1 are hatched except for inclusion 12.)

Figure 10. The same is true for the average Fe and Ca contents of the orthopyroxenes in different inclusions, higher FeO being accompanied by higher CaO values (Figure 11).

TABLE 3.—FeO content of olivine grains in two thin sections as measured on two different days by electron microprobe

SECTION 1

SECTION 10

Grain No.

Day 1

Day 2

Grain No.

Day 1

Day 2

18

4.48

4.39

34

3.39

3.31

23

4.07

4.10

36

3.60

3.77

25

4.24

4.30

37

3.79

3.49

89

4.04

4.00

40

3.61

3.57

95

4.31

4.30

41

3.72

3.58

80

4.39

4.39

45

3.86

3.73

Mean

4.26

4.25

56

3.57

3.36

62

3.40

3.38

59

2.98

2.99

71

3.83

3.73

Mean

3.56

3.49

Chemical Composition

Samples of three silicate inclusions, weighing 3.5 (no. 1), 6.7 (no. 2), and 9 (no. 3) grams, respectively, were taken for chemical analysis. Number 1 was a surface inclusion, and the location of the other two is shown in Figure 3. These samples contained varying amounts of metal as grains dispersed throughout the silicate matrix and veinlets or protrusions extending from the main mass of meteorite into the silicates. Each sample was ground in a tungsten carbide mill and passed through a 100-mesh sieve to separate coarse metallic particles from finely ground silicates. Since for this study analyses of the silicates are of prime interest, only the fraction less than 100 mesh was used for analysis. The bulk chemical analysis was performed in a manner similar to that described earlier (Jarosewich, 1966). The results are shown in Table

TABLE 4.—Average FeO content of olivine and FeO and CaO content (weight percent) of orthopyroxene of all thin sections studied

TABLE 5.—Average CaO, Na_2O, and K_2O content of plagioclase in different inclusions as measured by electron microprobe

7. Because the carbon, troilite, and metal content of these inclusions varies, the analyses were recalculated on a carbon, troilite, and metal-free basis in Table 8 and compared with bulk analyses of other silicates from iron meteorites and with chondrites. The chondrite values for Tables 7 and 8 were calculated from an average of 19 H- and 27 L-group chondrite analyses by one of us (Jarosewich). Only meteorites showing no signs of alteration by weathering were chosen for the calculation of the averages.

Discussion

COMPARISON TO CHONDRITES

Mason (1967) first noted that the minerals of the silicate inclusions in iron meteorites are essentially those of chondrites: olivine, orthopyroxene, diop-side, and an albitic plagioclase. This distinguishes them clearly from achondrites and mesosiderite silicates, which contain anorthitic plagioclase. The main difference from normal chondrites is the FeO content of olivine and pyroxene, which is intermediate between enstatite and bronzite chondrites (Table 2).

The bulk chemical composition of whole inclusions does not match that of chondrites. As the amounts of metal, troilite, and graphite are highly variable (Table 1), only the bulk compositions of the silicates are compared here. Some of the metal and troilite and most of the graphite are probably secondary in origin, i.e., they were introduced later into the silicate fragments.

Table 8 gives the bulk chemical composition of the chemically analyzed inclusions 1, 2, and 3, recalculated for the silicate portion only. Included in Table 8 are calculated bulk chemical values for inclusion 1 and for the additional inclusions 4 and 5. These values were calculated from the modal composition of these inclusions (Table 1) and the chemical composition of the minerals as determined by the electron microprobe. The agreement between the chemical analysis and that calculated from the modal composition of inclusion 1 shows that these two samples from the same inclusion are similar.

FIGURE 9.—Microprobe step-scan for Fe and Mg across an orthopyroxene grain in section 8. Grain center at left, boundary towards feldspar at right. Fe increase between the grains is caused by thin vein of iron oxide formed by terrestrial weathering.

FIGURE 10.—CaO vs FeO in orthopyroxene from inclusion 6.

FIGURE 11.—CaO vs FeO, average values of orthopyroxenes from different sections.

This is an important test to determine whether or not the chemical differences between different inclusions, as shown in the values of Table 8, are due to sample inhomogeneities or are real. These variations are already evident from the variations in modal composition of inclusions 4 and 5 (Table 1). In Table 8 an average composition for the inclusions 1 to 5 has been calculated, although it seems doubtful that such an average gives meaningful values.

A consistent feature of the inclusions is the lower FeO-content compared to chondrites, which is caused by the lower FeO-content of the mafic minerals. In Table 9, therefore, the silicate composition

TABLE 6.—FeO content (weight percent) of olivines in inclusion 8

Grain No.

Center

Rim

38

4.18

3.71

43

3.83

3.45

44

4.09

3.80

45

4.28

3.79

51

4.35

3.98

55

4.39

3.62

79

4.39

3.91

81

4.34

4.00

Mean

4.23

3.78

TABLE 7.—Chemical composition of three El Taco silicate inclusions and average of 19 H. and 26 L-type chondrites (some unpublished results are included in averages; analyst: E. Jarosewich)

Chondrite

Constituent

No. 1 Surface

Location No.2 11B-25/2-5

No.3 11B-24/19-22

H-type

L-type

Fe

1.75

0.80

0.25

15.71

6.47

Ni

0.18

0.27

0.10

1.72

1.25

Co

0.02

0.02

<0.01

0.09

0.06

FeS

0.33

7.29

1.32

5.56

6.00

SiO_2

37.82

45.86

40.52

36.41

39.81

TiO_2

0.13

0.12

0.13

0.12

0.13

Al_2O_3

1.36

3.41

1.64

2.11

2.20

Cr_2O_3

0.10

0.33

0.30

0.52

0.52

FeO

3.95

4.94

3.33

10.54

14.68

MnO

0.35

0.32

0.31

0.31

0.33

MgO

28.63

28.27

31.94

23.23

24.69

CaO

1.09

1.09

2.48

1.74

1.87

Na_2O

0.52

1.23

0.47

0.84

0.96

K_2O

0.04

0.11

0.06

0.09

0.11

P_2O_5

0.39

0.26

0.03

0.27

0.22

H_2O+

1.13

0.85

0.76

0.46

0.53

H_2O-

0.07

0.17

0.12

0.13

0.14

C

23.00

5.35

16.41

0.12

0.19

Total

100.86

100.69

100.17

99.97

100.16

Total Fe

5.03

9.27

3.68

27.49

21.69

has been recalculated without FeO, to give a better comparison to chondrites. The variations between different inclusions and the deviations from the chondritic average are most pronounced for the elements Al, Cr, Ca, Na, K, and P. Also the mean MgO content is higher, although the values of inclusions 2 and 4 are close to the chondrite average. The chemical variability and the deviations from chondritic values are also shown by the element ratios included in Table 9. The mean Mg/Si atom ratio (× 100) of 110 is higher than in normal chondrites (94), and closer to the value for carbonaceous chondrites of 104 (Ahrens and von Michaelis, 1968); on the other hand the Ca/Si ratio of 3.5 is lower than the chondritic average of 5.0 and more similar to the ratio of enstatite chondrites, which is 3.0 (Ahrens and von Michaelis, 1968). These agreements seem to be fortuitous, however, in view of the large variations between the inclusions. It is also evident from Table 9 that inclusions which show agreement in one of the element ratios with the chondritic value do not do so in other ratios. The best illustration of the chemical variations is given by the Ca/Al ratio. It varies here by a factor of 5 between 28 and 140, whereas this ratio is very constant at 74±6 (atom %) for all stony meteorite groups including the achondrites (Ahrens and von Michaelis, 1969).

TABLE 8.—Composition of the silicate portion of inclusions 1-5, compared to chondrites and mesosiderites

TABLE 9.—Composition of silicate inclusions 1–5 and chondrite average from Table 8, recalculated also without FeO (element ratios are calculated from atom percentages)

FIGURE 12.—Border zone of silicate inclusion 3, transmitted light. Coarse silicate grains occur next to the graphite zone (black, top) at the metal to silicate inclusion boundary, smaller silicate grains with graphite flakes (black) further inside (bottom). (Width of image = 1.5 mm.)

As Ca and Al are located in diopside and feldspar in these inclusions, we tried to match the composition of chondrites more closely by the addition or subtraction of one or both of these minerals. A fit for Ca and Al should also result in correct values for Na and K, as these elements are found in feldspar only. Because the composition of the minerals in these inclusions is so close to that of chondrites, it is self-evident that a mixture of these minerals in the right proportions should also produce a good match to the bulk chemistry of chondrites. Nevertheless, it is interesting to see how much diopside or feldspar of the analyzed composition has to be added or subtracted to come close to a chondritic silicate composition. For this comparison FeO has again been omitted from the calculations, as its content is much lower in El Taco olivines and pyroxenes than in ordinary chondrites. The results are shown in Figure 13: chondritic compositions can be matched quite well by addition or subtraction of small amounts of diopside and/or plagioclase for all inclusions. The correction is not sufficient, however, to match the Mg/Si ratio of normal chondrites completely. To achieve this, also the pyroxene/olivine ratio has to be increased: a ratio of 2:1 as in inclusion 4 would be appropriate.

Thus it seems feasible to derive these inclusions from chondritic material (neglecting the FeO content, which is discussed further below), from which mainly feldspar and/or diopside was lost (the grains in inclusion 2 and 3 are so small that they may be neglected). A mechanism for this derivation would be partial melting and redistribution of the melt, as albite and diopside would be the first minerals to melt, at 1120° and 1390° C, respectively. It is important to note, however, that different inclusions need different amounts of diopside and feldspar and that these minerals are not lost in proportional amounts, as already evident from the highly variable Ca/Alf ratio of these inclusions. It can be seen from the values of Table 1 that the relative amounts of olivine and pyroxene also vary by a factor of 2. These variations are probably not due to sampling errors, as the good agreement of chemical and calculated analysis of different areas of inclusion 1 (Table 8) shows. It may be argued, however, whether the lower melting silicates were entirely lost from the inclusion or may still be present in a small border zone towards the metal. A simple calculation shows that a surface layer 0.5 mm thick of an inclusion (Figure 12) with the dimensions 3 × 3 × 6 cm (like inclusion 3) would contain 8% of the material of this inclusion by volume. This border zone would then have to consist entirely of feldspar and/or diopside to account for the material which was lost (Figure 13). This is not observed in thin sections that are available from the metal-silicate interface of inclusions 2 and 3. Feldspar is somewhat enriched in this zone but it is still a minor phase in relation to pyroxene and olivine. Thus, it seems safe to conclude that the variations in chemical and mineralogical composition of these inclusions are due neither to inhomogeneities nor to nonrepresentative sampling.

FIGURE 13.—Open bars: Main element content of silicate portions of inclusions 1 to 5, normalized to average chondrite composition (from Table 9). Hatched bars: Composition after addition or subtraction of the indicated amounts of diopside and plagioclase feldspar of the composition measured in inclusions 1–5.

The match with chondrites for the main elements extends also to the minor ones, such as Ti and Mn. Notable exceptions, however, are Cr and P, which are highly variable, and K, which is usually too low.

In chondrites, Cr is mainly contained in chromite. If FeO is reduced to iron metal, the next element to be reduced would be Cr, according to its redox potential. If chromite is consumed by reduction, first Fe is reduced, so that the residual chromite should become very high in Cr. This is found in the El Taco chromites (Bunch, Keil, and Olsen, 1970, table 7). Upon further reduction, Cr also leaves this mineral and enters troilite. The troilite of the El Taco inclusions contains 0.3% Cr (Bunch, Keil, and Olsen, 1970, table 8). Thus, Cr can be lost from these inclusions in variable amounts according to its degree of reduction and the fate of the troilite of a specific inclusion. Inclusion 1, which is lowest in FeS (Table 1), is also lowest in Cr (Table 9).

In the case of P, both reduction and partial melting may be responsible for its variations. Apatite may be reduced to schreibersite, as their coexistence close to one another shows. It could also be removed or redistributed by partial melting. In the thin sections small apatite grains are found throughout the silicates, usually in association with troilite grains, but larger ones occur only at the border to the metal host or to metal/graphite veins.

VARIATIONS IN MINERAL COMPOSITION

Besides the variations in bulk composition, slight but distinct differences exist in mineral composition, as shown in Table 2 and Figure 8. In several of the FeO-histograms of Figure 8 measurements of different areas of the same inclusion are plotted. These show that in a given inclusion the distribution of FeO values for pyroxenes and olivines and the average are constant, but that different inclusions can have differences outside the analytical error. Also, the ratio of Fe/Mg in olivine to this ratio in pyroxene varies, as low FeO contents in olivine may coexist with high FeO content of pyroxene, and vice versa.

In general, in the El Taco inclusions and in the silicate inclusions of other IAB irons (Bunch, Keil, and Olsen, 1970), the Fe/(Fe + Mg) atom ratios are higher in pyroxene than in olivine, whereas in normal chondrites and terrestrial rocks, olivine in equilibrium with orthpyroxene is more iron-rich (Keil and Fredriksson, 1964; Medaris, 1969). This led Olsen and Fuchs (1967) to the conclusion that equilibrium was not reached in these silicate inclusions. Later, however, Olsen and Bunch (1970) argued that equilibrium should have been reached during the long cooling times of iron meteorites (1° to 10° C per 10^6 years, according to Goldstein and Short, 1967:1733). As they did not find any signs of disequilibruim, such as variable Fe contents or zoning, they concluded that an equilibrium with a reverse trend occurs at low FeO/(FeO + MgO) percentages, so that the pyroxene becomes more iron-rich than the coexisting olivine. The measurements on El Taco show that there are clear indications of disequilibrium between different inclusions. The most striking indication is the variation in the ratio of Fe/Mg in olivine over Fe/Mg in pyroxene (Table 4). The other is the slight zoning of the FeO content in the mafic minerals. Table 6 and Figure 9 show that the rims of olivine, pyroxene, and diopside grains have lower iron contents than the central part. The crystallization zoning should be just the opposite, as Fe-rich olivines and pyroxenes melt at lower temperatures than their Mg-rich counterparts.

It should be emphasized that we are speaking of three kinds of disequilibrium: (1) disequilibrium between different inclusions, expressed as different Fe/Mg ratios in olivine compared to Fe/Mg in pyroxene; (2) disequilibrium within single olivine and pyroxene grains of one inclusion, the rim being depleted in iron compared to the center; and (3) “equilibrium” within one inclusion, that is a narrow range of Fe/Mg ratios in olivines and pyroxene. In a given inclusion, the large olivine and pyroxene grains which are close to the metal border usually have the highest FeO contents. On the other hand, there is no simple relationship between the average grain size of an inclusion and the average FeO content of its mafic minerals. Contrary to the observation on single grains mentioned above, inclusion 5 with the highest average FeO content in the olivine has one of the smallest grain-size textures.

The composition of the main minerals is very close to that of chondrites except for the lower FeO content of the mafic minerals. This implies a slightly more reduced condition of this material, but not nearly as reduced as that of enstatite chondrites. Enstatite in enstatite chondrites contains 0.7% FeO (type I) or 0.2% FeO (type I), according to Keil (1968). Their more reduced state is also reflected in the presence of 1% to 3% Si in the kamacite metal (Ringwood, 1961: Keil, 1968). No Si was detected in the metal of the El Taco silicate inclusions. It has also been shown by Olsen and Fuchs (1967) that the coexistence of schreibersite and phosphates in these inclusions allows for an oxygen fugacity during their last high-temperature transformation of 10−20 atm, which is in the same range as that of normal chondrites. The minor-element content of the orthopyroxene is close to the values for this mineral from normal chondrites but usually higher than that of enstatite from enstatite chondrites. (Table 2).

FELDSPAR COMPOSITION

The anorthite content of 15 mole percent is close to that of normal chondrites, but is even closer to the feldspar of enstatite chondrites (Table 10). Thus the El Taco oligoclase fits into this general sequence of increasing An content with decreasing Fe content of olivine and pyroxene. The Or content of 3 mole percent is notably low in El Taco. This is the reason that the addition of this feldspar to the bulk silicate composition of the inclusions gave a good match to chondritic values for Na, but not for K (Figure 13). For the same reason the analysis calculated from the mode of inclusion 1 is too low in potassium, whereas it agrees well with the chemical analysis for sodium (Table 8). The Na/K atomic ratio of the feldspar is 26.6, whereas

TABLE 10.—Feldspar composition (mole percentages) in El Taco inclusions compared to the average compositions in L-, H-, and enstatite chondrites

Sample

Anorthite

Qrthoclase

El Taco, range

14.3–15.9

2.7–3.0

average

15.0

2.8

L-chondrites*

10

5

H-chordrites*

12

6

Enstatite chondrites, type 2†

15

4

*Van Schraus and Ribbe, 1968, table 1,

†Keil, 1968, table 4

in the bulk analyses this ratio varies between 12 and 20 and is thus closer to the value of 13.5 for the average chondrite (Table 7). This means that an additional phase containing potassium must be present. This was found as small K-feldspar grains, which are only a few microns across and occur usually at plagioclase-graphite grain boundaries. The average composition of two grains large enough for analysis is given in Table 2. Small K-feldspar grains associated with albite in the Odessa iron meteorite were reported by El Goresy (1967). Bunch, Keil, and Olsen (1970) observed K-enriched zones in the feldspar of Odessa and Toluca and thought them due to K-leaching by terrestrial weathering. This explanation can be excluded for El Taco by the microprobe analysis of these grains. They can be explained by exsolution from the albitic plagioclase because of the expanding solvus at low temperatures. The temperature of exsolution of the albite according to the solvus determined by Goldsmith and Newton (1974) should be below 600° C and lower than the last equilibration temperature of chondritic feldspars. Recently, a phase containing 50 mole percent Or was observed as exsolution lamellae in the albitic plagioclase of a silicate nodule in the San Cristobal iron (Scott and Bild, 1974).

CLINOPYROXENE-ORTHOPYROXENE THERMOMETRY

The Wo content of the diopsides in El Taco of 44 mole percent would correspond to a temperature of about 1000° C on the diopside solvus of Davis and Boyd (1966). The KD value for the Fe/Mg equilibrium between diopside and orthopyroxene as used by Kretz (1963) and McCallum (1968) gives a much higher temperature. For the pyroxenes of inclusion PTS 8 KD = (Fe/Mg) orthopyroxene/(Fe/Mg) clinopyroxene gives a value of 1.11. On McCallum’s equilibrium curve this would correspond to a temperature of about 1700° C. This temperature is much too high, and if reduction of FeO from the silicates took place the temperature derived in this way is probably meaningless.

ORIGIN OF THE INCLUSIONS

The similarity between chondritic and the El Taco silicate minerals favors a close relation between the two. Whether this is a genetic one, so that El Taco silicate inclusions were derived by transformation of chondrite material, cannot be answered with certainty. Furthermore, there are many ways in which a chondrite could be transformed to give the El Taco silicates. Three features have to be explained: the lower FeO content of the mafic minerals, the apparent differences in bulk composition, and the disequilibrium between different inclusions.

One simple model that would explain most of the El Taco silicate inclusion features would be the incorporation of chondritic matter into a hot iron metal, at or near to the nickel-iron melting point of about 1500° C. Heating of a chondrite with H-type olivine and pyroxene in an atmosphere of CO and CO_2 in equilibrium with one another will lead to a reduction of iron silicates above temperatures of about 1200° C. It has been shown by Mueller (1963) and Speidel and Nafziger (1968) that an olivine-silica-metal assemblage will change towards lower fayalite contents in the olivine when heated from 1100° to 1300° C. A reduction of the silicates to about 6 mole percent FeO, which is found in most of the silicates of iron meteorites with silicate inclusions (Bunch, Neil, and Olsen, 1970), may then reflect the equilibrium at the temperature of molten nickel-iron. At this temperature albite and diopside will melt also, but not olivine and enstatite, a condition which could explain the depletion of the former minerals discussed above. The apparent disequilibrium between different inclusions was probably not established at these high temperatures, but later during cooling, and may be the last change produced by a continuing reduction of iron oxide from the silicates. The lower FeO mole percent of olivine may then simply reflect the fact that olivine is easier to reduce than pyroxene. That this is the case was shown in the ureilite Haverö, where olivine next to carbon-rich veins is reduced to lower FeO contents, whereas pyroxene is not (Wlotzka, 1972).

The question remains, why did the high-temperature reduced assemblage not change back to its original, more oxidized form upon slow cooling? This difficulty can be overcome by changing from a closed to an open system, which would be the result in a change from rapid heating to slow cooling. A rapid heating process, as by an introduction of silicates into a molten metal pool or invasion of a chondritic rock by a metal melt, would essentially result in a closed system with respect to C/CO/CO_2 because the time for escape of the gases was too short. Thus the equilibrium between CO and CO_2 would establish the low Fa and Fs content of the mafic minerals at this high temperature. Slow cooling, which took place during the γ−α transformation between 850° and 600° C over some 100 million years (Goldstein and Short, 1967), will result in an open system, where oxygen removed by the reduction of Fe-silicates by carbon may escape and provide for the continuing reduction. The differences from inclusion to inclusion can then be explained by local differences in carbon available per inclusion or weight-unit of silicates. These differences are mainly due to a different FeO content of the olivines, since the pyroxene compositions are more uniform. As olivine is more easily reduced than pyroxene, its composition will follow these changes better than that of pyroxene. Thus the low FeO content of olivines from the small, veinlike inclusion 6 may be caused by the higher amounts of carbon available from the metal in this area.

The proportionality of Fe and Ca in orthopyroxene (Figure 10) is not in contradiction to a loss of FeO by reduction. Both FeO and CaO will be leaving the pyroxene grains by diffusion—FeO because it is reduced to Fe metal, and CaO because at lower temperatures the solubility of CaSiO_3 in the enstatite is lowered (Davis and Boyd, 1966).

Another possibility is a genetic relationship of the material of the El Taco inclusions to chondrites like Kakangari (Graham and Hutchison, 1974), which already contain olivines with a low fayalite and pyroxene with a low ferrosilite content, so that a reduction is not necessary. Only the late reduction leading to the disequilibrium between olivine and pyroxene and between different inclusions has to be operative here too.

Summary and Conclusions

It has been shown that (1) the silicate minerals of the El Taco inclusions have a composition close to the minerals of chondrites with the exception of a lower iron content of olivine and pyroxene; (2) the iron content of olivine and pyroxene shows small but distinct variations from inclusion to inclusion; and (3) the bulk chemical composition is close to chondrites with deviations mainly in the alkalies, Ca, and Al.

We suggest an explanation of these features by an in-situ reduction of chondritic material at the temperatures of molten nickel-iron together with partial melting and redistribution of the more fusible minerals, feldspar and diopside. The varying Fe/Mg ratios of olivine and pyroxene in different inclusions are explained as a minor disequilibrium due to continuing reduction at lower temperatures. The following observations appear to support this explanation:

Partial melting is supported by the rare-earth data on Woodbine by Masuda (1969) and Hintenberger and Berghof on El Taco (1970). The rare earth pattern shows a depletion in Eu and other deviations from the chondritic pattern, which can be explained by loss of a feldspar component.

Evidence for reduction of the mafic minerals is the FeO-depletion of the marginal zones of olivines and pyroxenes. This is supported by the observed reduction of Cr, which moves from chromite into troilite. Another hint in this direction is the lower Ni-content of the metal particles in the silicate inclusions compared to the metal host, which has been reported by Bunch, Keil, and Olsen (1970). This could be a dilution effect caused by the newly formed metallic iron.

Not much evidence is available for an in-situ process. A simple relation between size or location of the inclusions and their composition has not yet been found, except for the fact that the small veinlike inclusion 6 has one of the lowest and most variable FeO-contents in its mafic minerals. More inclusions of this type should be studied to answer this question.

On the other hand, the variations in composition and texture between large inclusions, together with a homogeneous composition of individual inclusions, are easier to understand if the differentiations took place before incorporation into the metal host. In any case, however, these differentiations are small compared to the differentiations which separate achondrites or mesosiderites from chondritic material. These inclusions are still close to primitive material and not what one would expect at the boundary between iron core and mantle of a planetary body. The “raisin-bread” theory, which sees the iron meteorites formed in small pools dispersed through silicates (Urey, 1959), fits the observations on El Taco better. It does not seem possible, however, to go one step further, as Wasson (1970) does, and assume that group I irons and their silicate inclusions are primitive condensates and were never molten. These inclusions vary in bulk and mineral composition, which means they were either formed under different conditions in different regions (if one wants to retain their primitive nature) or they were differentiated by melting processes. The latter seems much more probable.

Literature Cited

Ahrens, L. H., and H. von Michaelis

1968.   Fractionation of Some Abundant Lithophile Element Ratios in Chondrites.    Pages 257–272 in L. H. Ahrens, Origin and Distribution of the Elements. Oxford: Pergamon Press.

1969.    The Composition of Stony Meteorites, III: Some Inter-Element Relationships. Earth and Planetary Science Letters, 5:395–403.

Bence, A. E., and A. L. Albee

1968.    Empirical Correction Factors for the Electron Microanalysis of Silicates and Oxides. Journal of Geology, 76:382–403.

Bunch, T. E., and W. Cassidy

1968.    Impact-induced Deformation in the Campo del Cielo Meteorite. Pages 601–612 in B. French and N. Short, Shock Metamorphism of Natural Materials. Baltimore: Mono Book Corp.

Bunch, T. E., K. Keil, and E. Olsen

1970.    Mineralogy and Petrology of Silicate Inclusions in Iron Meteorites. Contributions to Mineralogy and Petrology, 25:297–340.

Cassidy, W. A., L. M. Villar, T. E. Bunch, T. P. Kohman, and D. J. Milton

1965.    Meteorites and Craters of Campo del Cielo, Argentina. Science, 149:1055–1064.

Davis, B. T. C, and F. R. Boyd

1966.    The Join Mg_2Si_6O–CaMgSi_2O_6 at 30 kb Pressure and Its Application to Pyroxenes from Kimberlites. Journal of Geophysical Research, 71:3567–3576.

El Goresy, A.

1967.    Quantitative Electron Microprobe Analyses of K-Feldspar Grains from the Odessa Iron Meteorite [Abstract]. 30th Meeting of the Meteoritical Society, October 25–27, 1967. Moffett Field, California: Ames Research Center.

Goldsmith, J. R., and R. C. Newton

1974.    An Experimental Determination of the Alkali Feldspar Solvus. Pages 337–359 in W. S. Mackenzie and J. Zussman, The Feldspars. Manchester: University Press.

Goldstein, J. I., and J. M. Short

1967.    The Iron Meteorites: Their Thermal History and Parent Bodies. Geochimica et Cosmochimica Acta, 31:1733–1770.

Graham, A. L., and R. Hutchison

1974.    Is Kakangari a Unique Chondrite? Nature, 251:128–129.

Henderson, E. P., and S. H. Perry

1949.    The Linwood (Nebraska) Meteorite. Proceedings of the United States National Museum, 99:357–360.

Hintenberger, H., and W. Berghof

1970.    Mass Spectrographic Determination of Trace Elements in Meteorites. Pages 657–665 in K. Ogata and T. Hayakawa, Recent Developments in Mass Spectroscopy. Tokyo: University of Tokyo Press.

Hintenberger, H., L. Schultz, and H. Weber

1969.    Rare Gases in the Iron and the Inclusions of the Campo del Cielo Meteorite, El Taco. Pages 895–900 in P. M. Millman, Meteorite Research. Dordrecht, Holland: D. Reidel Publishing Company.

Jarosewich, E.

1966.    Chemical Analyses of Ten Stony Meteorites. Geochimica et Cosmochimica Acta, 30:1261–1265.

Keil, K.

1968.    Mineralogical and Chemical Relationships among Enstatite Chondrites. Journal of Geophysical Research, 73:6945–6976.

Keil, K., and K. Fredriksson

1964.    The Iron, Magnesium, and Calcium Distribution in Coexisting Olivines and Rhombic Pyroxenes of Chondrites. Journal of Geophysical Research, 69:3487–3515.

Kretz, R.

1963.    Distribution of Magnesium and Iron between Orthopyroxene and Calcic Pyroxene in Natural Mineral Assemblages. Journal of Geology, 71:773–785.

Marshall, R. R., and K. Keil

1965.    Polymineral Inclusions in the Odessa Iron Meteorite. Icarus, 4:461–479.

Mason, B.

1967.    The Woodbine Meteorite, with Notes on Silicates in Iron Meteorites. Mineralogical Magazine, 36:120–126.

Masuda, A.

1969.    Lanthanides in the Silicate Inclusions of the Woodbine Meteorite. Nature, 224:164–165.

McCallum, I. S

1968.    Equilibrium Relationships among the Coexisting Minerals in the Stillwater Complex, Montana. PhD dissertation, University of Chicago.

Medaris, L. G.

1969.    Partitioning of Fe2+ and Mg2+ between Coexisting Synthetic Olivine and Orthopyroxene. American Journal of Science, 267:945–968.

Mueller, R. F.

1963.    A Comparison of Oxidative Equilibria in Meteorites and Terrestrial Rocks. Geochimica et Cosmochimica Acta, 27:273–278.

Olsen, E., and T. E. Bunch

1970.    Empirical Derivation of Activity Coefficients for the Magnesium-Rich Portion of the Olivine Solid Solution. American Mineralogist, 55:1829–1842.

Olsen, E., and L. Fuchs

1967.    The State of Oxidation of Some Iron Meteorites. Icarus, 6:242–253.

Olsen, E., and E. Jarosewich

1971.    Chondrules: First Occurrence in an Iron Meteorite. Science, 174:583–585.

Park, F. R., T. E. Bunch, and T. B. Massalski

1966.    A Study of the Silicate Inclusions and Other Phases in the Campo del Cielo Meteorite. Geochimica etCosmochimica Acta, 30:399–414.

Powell, Benjamin N.

1971.    Petrology and Chemistry of Mesosiderites, II: Silicate Textures and Compositions and Metal-Silicate Relationships. Geochimica et Cosmochimica Acta, 35(1):5–34.

Ringwood, A. E.

1961.    Silicon in the Metal Phase of Enstatite Chondrites and Some Geochemical Implications. Geochimica et Cosmoschimica Acta, 25:1–13.

Scott, E. R. D., and R. W. Bild

1974.    Structure and Formation of the San Cristobal Meteorite, Other IB Irons and Group IIICD. Geochimica et Cosmochimica Acta, 38:1379–1391.

Speidel, D. H., and R. H. Nafziger

1968.    P–T–f_02 Relations in the System Fe–O–MgO–SiO_2. American Journal of Science, 266:361–379.

Urey, H. C.

1959.    Primary and Secondary Objects. Journal of Geophysical Research, 64:1721–1737.

Van Schmus, W. R., and P. H. Ribbe

1968.    The Composition and Structural State of Feldspar from Chondritic Meteorites. Geochimica et Cosmochimica Acta, 32:1327–1342.

Wasson, J. T.

1970.    The Chemical Classification of Iron Meteorites, IV: Irons with Ge Concentrations Greater than 190 ppm and Other Meteorites Associated with Group I. Icarus, 12:407–423.

Wlotzka, F.

1972.    Haverö Ureilite: Evidence for Recrystallization and Partial Reduction. Meteoritics, 7:591–600.

Wlotzka, F., and E. Jarosewich

1969.    The Mineralogical and Chemical Composition of Silicate Inclusions in the El Taco (Campo del Cielo) Iron Meteorite. Meteoritics, 4:298–299.

* This catalog is available on IBM cards or on magnetic tape and may be purchased from Department of Mineral Sciences, National Museum of Natural History, Smithsonian Institution, Washington, D.C. 20560.