Oxygen isotopic and chemical compositions of cosmic spherules collected from the Antarctic ice sheet: Implications for their precursor materials

Geochimica et Cosmochimica Acta, Vol. 69, No. 24, pp. 5789 –5804, 2005 Copyright © 2005 Elsevier Ltd Printed in the USA. All rights reserved 0016-7037/05 $30.00 ⫹ .00

doi:10.1016/j.gca.2005.08.002

Oxygen isotopic and chemical compositions of cosmic spherules collected from the Antarctic ice sheet: Implications for their precursor materials TORU YADA,1,2,*,† TOMOKI NAKAMURA,2 TAKAAKI NOGUCHI,3 NORIKO MATSUMOTO,3 MINORU KUSAKABE,4 HAJIME HIYAGON,1 TAKAYUKI USHIKUBO,1,‡ NAOJI SUGIURA,1 HIDEYASU KOJIMA,5 and NOBUO TAKAOKA2 1

Department of Earth and Planetary Science, Graduate School of Science, University of Tokyo, 7-3-1 Hongo, Bunkyo-ku, Tokyo 113-0033, Japan 2 Department of Earth and Planetary Sciences, Graduate School of Science, Kyushu University, 6-10-1 Hakozaki, Fukuoka 812-8581, Japan 3 Department of Minerals and Biological Sciences, Ibaraki University, Bunkyo 2-1-1, Mito, Ibaraki 310-8512, Japan 4 Institute for Study of the Earth’s Interior, Okayama University, 827 Yamada, Misasa, Tottori 682-0193, Japan 5 National Institute of Polar Research, Department of Polar Science, Graduate University for Advanced Studies, 1-9-10 Kaga, Itabashi, Tokyo 173-8515, Japan (Received October 14, 2004; accepted in revised form August 1, 2005)

Abstract—Bulk chemical compositions and oxygen isotopic compositions were analyzed for 48 stony cosmic spherules (melted micrometeorites) collected from the Antarctic ice sheet using electron- and ion-microprobes. No clear correlation was found between their isotopic compositions and textures. The oxygen isotopic compositions showed an extremely wide range from ⫺28‰ to ⫹93‰ in ␦18O and from ⫺21‰ to ⫹13‰ in ⌬17O. In ␦18O-␦17O space, most samples (38 out of 48) plot close to the terrestrial fractionation line, but 7 samples plot along the carbonaceous chondrite anhydrous mineral (CCAM) line. Three samples plot well above the terrestrial fractionation line. One of these has a ⌬17O of ⫹13‰, the largest value ever found in solar system materials. One possible precursor for this spherule could be 16O-poor planetary material that is still unknown as a meteorite. The majority of the remaining spherules are thought to be related to carbonaceous chondrites. Copyright © 2005 Elsevier Ltd which are supposed to have resulted from ion-molecular reactions in the early solar nebular or in interstellar clouds (Messenger, 2000; Floss et al., 2004), their pristine mineralogy which is similar to the hypothetical accretional dust in the early solar nebula (Bradley et al., 1983), and their amorphous silicate component (GEMS: glass embedded with metals and sulfides) whose spectrum fits that of interstellar dust (Bradley, 1994). Cosmic spherules are considered to be extraterrestrial samples based on various lines of evidence: high Ni or Ir concentrations (Millard and Finkelman, 1970; Yamakoshi, 1984; Brownlee et al., 1984), the presence of cosmogenic nuclides such as 10Be, 26Al, and 53Mn (Nishiizumi, 1983; Raisbeck et al., 1985; Raisbeck and Yiou 1987; Nishiizumi et al., 1991, 1995), oxygen isotopic compositions (Clayton et al., 1986; Engrand et al., 1998; Taylor et al., 2005), and solar noble gas signatures (Osawa et al., 2003). Nishiizumi et al. (1995) indicated that only 6 of 42 stony cosmic spherules had 26Al/10Be ratios consistent with those of meteorite-sized bodies; the remainder were originally dust grains in interplanetary space. Thus, most of the spherules were originally small bodies in interplanetary space, not ablation spheres from meteorites. Determination of their precursor materials has been difficult because melting during entry into the atmosphere changed their original compositions, mineralogies and textures. Bulk chemistry provides essential information for determining the precursor materials, as indicated by Brownlee et al. (1997) who showed similarities between stony (S) type spherules and CM chondrites in Mn/Si ratio. However, other evidence is necessary to support this view. Since oxygen isotopic anomalies were first discovered in refractory inclusions in Allende (CV3) (Clayton et al., 1973), both oxygen isotopic compositions of chondrites and achondrites, as well as their minerals and constituent components have been found to vary widely (Clayton,

1. INTRODUCTION

One of the major interests for studies of micrometeorites (MMs) and interplanetary dust particles (IDPs), the two major types of extraterrestrial dust collected on Earth, is the determination of their precursor bodies. Interplanetary dust captured by the Earth should be a broader collection of bodies orbiting at heliocentric distances greater than that of the Earth, such as asteroids, comets, and Edgeworth-Kuiper belt objects (Grün et al., 1985; Liou et al., 1996; Kortenkamp and Dermott, 1998; Landgraf et al., 2002). Among the possible parent bodies above, the primary precursors must be asteroids and comets, although the abundance ratio of the two sources is under debate (i.e., Gustafson 1994; Kortenkamp and Dermott, 1998; Ishimoto, 2000). Hydrous IDPs consist of phyllosilicates (Tomeoka and Buseck, 1985), and also, most of fine-grained MMs likely consisted of phyllosilicates originally (Kurat et al., 1994; Genge et al., 1997; Nakamura et al., 2001; Noguchi et al., 2002). Thus, they resemble the matrices of hydrous carbonaceous (C) chondrites in mineralogy and chemistry. Parent bodies of C chondrites are supposed to be C, D, and T type asteroids (Hiroi et al., 2001; Burbine et al., 2002). On the other hand, most anhydrous IDPs are thought to originate from comets, because of their porous structures (Brownlee, 1985), their high carbon contents comparable to those of comets (Thomas et al., 1993), their H, C and N isotopic anomalies

* Author to whom correspondence should be addressed ([email protected]). † Present address: Materials Science Division, Argonne National Laboratory, 9700 S. Cass Ave., Argonne, IL 60439, USA ‡ Present address: Department of Geological Sciences, Arizona State University, P.O. Box 871404, Tempe, AZ 85287-1404, USA 5789

5790

T. Yada et al.

Fig. 1. SEM images of typical textures of S-type spherules. All the images except for (d) are secondary electron images, and (d) is a backscattered electron image. (a) M240u02, a barred olivine (BO) type spherule. Bar-shaped olivines and tiny magnetites are embedded in interstitial glass. The observed dents are sputtering spots made by the ion microprobe, and white flakes are remnants of the gold coating. (b) K02KS300, the largest spherule of the homogenous glass (Glass) type. The observed dents are ion microprobe sputtering spots. The spots arrowed with numbers correspond to analysis points shown in Table 3. (c) M03KS063, a cryptocrystalline glass (CC) type spherule. This type has textural heterogeneity, but no apparent mineral grains are observed except for very fine-grained minerals. The observed dents are sputtering spots by the ion microprobe. (d) M03KS019, a porphyritic olivine (Po) type spherule. This type is composed of porphyritic olivines, magnetites and mesostasis. (e) KW540260, a relict grain-bearing (RGB) type spherule. This spherule contains irregularshaped, forsteritic olivines, which resembles the amoeboid olivine aggregates found in C chondrites. Two dents observed in it are sputtering holes made by the ion microprobe and white flakes around and inside it are remnants of the gold coating. (f) MY240184, the only Ca, Al and Ti-rich (CAT) spherule in this study. It consists of homogeneous glass. The dent in its center is a sputtering hole by the ion microprobe.

1993). In this sense, oxygen isotopic compositions of planetary materials enable us not only to distinguish the origin of the materials, but also inform us of conditions and processes in the early solar system. Thus, oxygen isotopic compositions of S-type spherules should help reveal their precursors. MMs are the dominant extraterrestrial materials accreting to the Earth’s surface today. The influx of micrometeoroids to Earth is estimated to be 40,000 ⫾ 20,000 ton/yr (Love and Brownlee, 1993), and some parts of the incoming micrometeoroids are observed as MMs on the Earth’s surface, whose accretion rate is estimated to range from 2700 ⫾ 1400 to 16,000 ⫾ 9100 ton/yr (Taylor et al., 1998; Duprat et al., 2001; Yada et al., 2004), whereas that of meteorites is estimated to be 2.9 to 7.3 ton/yr (Bland et al., 1996), or ⬎ 510 ton/yr (Halliday et al., 1989). Cosmic spherules comprise ⬃80% of Antarctic MMs (AMMs) in the ⬎ 100-␮m size fraction (Maurette et al., 1991; Taylor, private communication), but despite their abun-

dance, their oxygen isotopic compositions have not been extensively studied. Clayton et al. (1986) analyzed the oxygen isotopic ratios of pooled samples of cosmic spherules from deep-sea sediments (up to 600 iron [I] type and up to 315 stony [S] type samples). The S-type spherules ranged from 22‰ to 25‰ in ␦18OSMOW and 9‰ to 11‰ in ␦17OSMOW and plotted below the terrestrial fractionation line (TFL), and the I-type spherules were enriched in ␦18O by 40‰ to 50‰ on the TFL. Clayton et al. (1986) argued that the original isotopic compositions of S-type spherules were similar to those of C chondrites and that the I-type spherules acquired their unusual high isotopic signature through O-exchange with the upper atmosphere. However, Davis et al. (1991) disputed the existence of heavy oxygen isotopic components in the upper atmosphere as proposed by Clayton et al. (1986), because the large isotopic fractionations of iron observed in I-type spherules suggested that the heavy isotopic compositions of both iron and oxygen

Oxygen isotope compositions of stony cosmic spherules

resulted from isotopic fractionation which the spherules experienced during atmospheric entry heating and subsequent evaporation. Since individual spherules may have different origins, the analysis of the oxygen isotopic composition of individual spherules is essential to determine their precursors. Engrand et al. (1998) analyzed the oxygen isotopes of five Antarctic cosmic spherules and suggested that their precursors are related to C chondrites, based on the fact that they plot below the TFL. Taylor et al. (2005) also analyzed oxygen isotopes in several Antarctic spherules and reached a similar conclusion. In this study, we analyzed both bulk chemical compositions and oxygen isotopic compositions of 48 individual Antarctic spherules to constrain their precursors and origins. 2. SAMPLES AND METHODS 2.1. Samples Most of the samples used in the present study were collected from the blue ice around the Yamato Mountains (Yada and Kojima, 2000; Terada et al., 2001; Yada et al., 2004). One spherule, however, is from sediment at a water tank in the Dome Fuji station, which is situated at the top of the ice shield of east Antarctica (Nakamura et al., 1999; Noguchi et al., 2000). The candidate spherules were handpicked from the glacial sands and their chemical compositions were determined by scanning electron microscopy, using the JEOL JSM5800LV at the National Institute of Polar Research (NIPR), Japan, and the JEOL JSM5600LV at Ibaraki University, Japan, in a low vacuum mode without carbon coating. The EDS spectra (energy dispersive X-ray spectroscopy) showed high peaks in Mg, Si and Fe and low peaks in Al, Ca and S, suggesting that the spherules were extraterrestrial. Fortyeight S-type spherules were used for the present study. Each spherule was glued to a duralumin cylinder (6 mm in diameter and 9 mm in height) with epoxy resin, and was later polished to expose its cross section. 2.2. Electron Microprobe Analysis The carbon coated surfaces of the sectioned S-type spherules were photographed to document their internal textures using the SEM-EDS, followed by bulk chemical analyses using electron microprobes (JEOL JXA8800M, NIPR, Japan and JEOL JXA733, Kyushu University, Japan) with an electron acceleration voltage of 15 kV and a beam current of 10nA. Standard minerals and samples were analyzed with a defocused electron beam of 5 ␮m in diameter under the same analytical conditions. ZAF corrections were applied for quantitative analyses. The detection limit is ⬃0.01 wt.% for elements analyzed, and the reproducibility is within ⫾2% for the major elements as determined by repeated measurements of natural mineral standards (olivine and chromite). Five to 19 spots were analyzed for each spherule and the average value was assumed to represent the bulk chemical composition. 2.3. Ion Microprobe Analysis Oxygen isotopic ratios are expressed in the conventional delta (␦) notation as defined by the following equation:

␦⫽

冉

RX RSMOW

冊

⫺ 1 ⫻ 1000

(‰).

where RX and RSMOW stand for the 18O/16O or 17O/16O ratio of sample X and Standard Mean Ocean Water, respectively. ⌬17O is calculated using the equation ⌬17O ⫽ ␦17OSMOW ⫺ 0.52 ⫻ ␦18OSMOW (Clayton, 1993), which expresses the degree of isotopic deviation from the mass-dependent terrestrial fractionation trend. The oxygen isotopic compositions of the spherules were analyzed with a secondary ion mass spectrometer (SIMS), CAMECA IMS6F at Kyushu University. Each sample was gold-coated. Primary Cs⫹ ions were accelerated at 19.5 kV to sputter the surface of the sample. The negatively charged secondary ions were accelerated with 9.5 kV and introduced into a double focus-

5791

ing mass spectrometer. The width of the energy slit was set to 110 eV, and the mass resolution was set at ⬃5000, sufficient to separate the 16 OH⫺ peak from the 17O⫺ peak. The tail of the 16OH⫺ peak under the 17 O peak center was calculated based on the 18O⫺ tail and 16OH⫺ intensity and kept below 0.02% of the 17O count rate throughout the analyses. The 16O⫺ ions, normally adjusted to ⬃3 ⫻ 107 counts per second (cps) using San Carlos olivine, were measured with a Faraday cup, and 17O⫺ and 18O⫺ ions were measured using pulse counting with an electron multiplier. The measurement time of the mass peaks were set 0.5, 4, and 2 s for 16O, 17O and 18O, respectively. Each analysis consisted of 60 to 80 cycles through the peaks. Approximately 0.8 mA of electron beam, accelerated at 9.5 kV, was sprayed over the analytical point with a normal-incidence electron flood gun to compensate for the positive charge on the sample surface (Slodzian, 1987). San Carlos olivine (SCO, Fo89.2), which has a bulk oxygen isotopic ratio of ␦18OSMOW ⫽ 5.4‰ and ␦17OSMOW ⫽ 2.9‰ (Kusakabe et al., 2004), was used as our standard mineral. It was mounted on the duralumin cylinder and polished in the same way as the samples. The cylinder was set in the same holder with the samples and analyzed repeatedly before and after the sample analyses. Typical statistical uncertainties in the oxygen isotopic analysis of SCO are ⫾1.2‰ for ␦18OSMOW and ⫾1.4‰ for ␦17OSMOW at 2␴, and the standard deviation calculated from the repeated SCO analyses within a day is typically ⬍ ⫾2.5‰ for both 18O/16O and 17O/16O ratios. All the samples’ data were normalized to the average of the SCO measurement on the same day to compensate for instrumental mass fractionation, which typically ranged from ⫺35‰ to ⫺15‰ in ␦18OSMOW relative to the SMOW ratio, 2.0052 ⫻ 10⫺3 in 18O/16O (Baertschi, 1976). The calculation error of each analysis for a sample includes the statistical uncertainty of the sample analysis and the standard deviation associated with the fractionation correction derived from the repeated analyses of the standard mineral (SCO) on the same day. In general, errors in the ⌬17O measurement of samples are smaller than those of ␦18OSMOW, because the errors in ⌬17O were calculated from statistical uncertainties of the samples’ analyses and the standard deviation of ⌬17O from the repeated analyses of the mineral standard, which is typically smaller than that of ␦18OSMOW. Based on the descriptions in McKeegan (1987), Leshin et al. (1997), and Yurimoto et al. (1998), matrix effects between silicates and oxide minerals are ⬍ 5‰ in ␦18OSMOW under low-energy filtering conditions. In our study, no matrix effect correction was performed on the ion microprobe data. 3. RESULTS

Electron micrographs of six S-type spherules are shown in Figure 1. We classified the 48 S-type spherules into six types based on the classification of Taylor et al. (2000). The types are: barred olivine (BO, 19 samples), glass (Glass, 10 samples), cryptocrystalline (CC, 10 samples), porphyritic olivine (Po, 5 samples), relict grain-bearing (RGB, 3 samples), and Ca, Al, and Ti-rich (CAT, 1 sample). Results of bulk major element analyses for the 48 S-type spherules by EPMA are shown in Table 1, and the average chemical compositions for each type of spherule are summarized in Table 2. Also, the average elemental abundances normalized to Si and CI in atomic ratios are displayed in Figure 2. M240168 (BO type), F96g343, and M03KS115 (Po type) are excluded from Figure 2 for the sake of good resolution of the groups, because they have some elements largely different from CI, for example, M240168 is ⫻5.4 enriched in Al relative to Si and CI, Ca (⫻5.6 CI) and Ti (⫻12 CI), F96g343 is enriched in Al (⫻5.1 CI), and M03KS115 is enriched in Al (⫻18 CI), Cr (⫻7.7 CI), Ca (⫻5.2 CI) and Ti (⫻4.6 CI). In general, S-type spherules are extremely depleted in volatile elements such as Na, S and P and moderately depleted in Fe and Ni relative to CI. These features have been observed in previous studies (Blanchard et al., 1980; Brownlee et al., 1997; Taylor et al., 2000; Alexander et al., 2002; Taylor et al., 2005).

Table 1. Bulk chemical compositions of 48 S-type spherules. Sample no.

F96g343a

M240168

M240u02

MY240159

MY240184

M03KS012

M03KS019

M03KS021

M03KS039

M03KS051

M03KS057

M03KS063

M03KS066

M03KS071

M03KS115

K02KS086

K02KS135

Type Size (␮m) Analysis point SiO2 (wt.%) TiO2 Al2O3 FeO MnO MgO CaO Na2O K 2O Cr2O3 NiO P2 O 5 S Total Mg/Al (atom) Si/Al (atom)

Po 75 ⫻ 70 5 31.8 0.16 11.7 26.9 0.23 26.5 3.24 N.D. N.D. 0.47 0.18 N.A. 0.01 101.20 2.9 2.3

BO 70 ⫻ 70 10 37.0 1.42 14.4 8.51 0.12 26.4 11.8 N.D. 0.01 0.17 0.35 0.03 N.D. 100.28 2.3 2.2

BO 135 ⫻ 100 13 41.0 0.08 1.40 14.2 0.35 43.4 0.69 N.D. N.D. 0.12 0.13 N.A. N.D. 101.33 39.1 24.8

Glass 95 ⫻ 95 8 45.3 0.05 1.84 3.75 0.05 49.7 0.15 0.01 N.D. 0.01 N.D. 0.01 N.D. 100.92 34.1 20.9

CAT 70 ⫻ 70 8 35.0 0.56 14.8 0.01 0.01 39.4 9.53 0.01 0.02 0.01 0.03 N.D. N.D. 99.32 3.36 2.01

CC 180 ⫻ 130 14 46.7 0.20 3.98 11.3 0.08 34.3 3.87 0.01 N.D. 0.03 0.13 N.D. N.D. 100.54 10.9 10.0

Po 165 ⫻ 150 15 34.6 0.18 1.79 32.4 0.16 25.5 3.10 0.03 N.D. 0.43 2.55 0.02 0.01 100.81 18.0 16.4

CC 170 ⫻ 145 14 40.9 0.13 2.23 22.5 0.33 29.9 4.77 0.02 N.D. 0.06 0.14 N.D. 0.01 100.97 16.9 15.6

RGB 80 ⫻ 75 11 41.1 0.04 0.62 9.18 0.18 48.8 0.61 0.01 N.D. 0.42 0.02 0.01 0.01 100.90 99.5 56.2

Glass 55 ⫻ 55 7 46.6 0.12 3.03 14.4 0.34 32.5 3.21 0.02 0.01 0.02 N.D. 0.01 N.D. 100.32 13.6 13.1

BO 50 ⫻ 45 6 35.1 0.14 3.03 24.1 0.05 32.7 2.68 0.03 N.D. 0.27 1.77 0.04 N.D. 99.97 13.7 9.8

CC 60 ⫻ 45 5 39.8 0.19 3.92 17.3 0.08 35.5 3.34 N.D. N.D. 0.07 0.31 N.A. 0.01 100.61 11.5 8.6

BO 70 ⫻ 60 10 42.8 0.15 3.40 25.4 0.50 23.2 2.90 0.09 N.D. 0.39 1.00 0.09 N.D. 99.95 8.6 10.7

Glass 90 ⫻ 90 13 56.5 0.14 2.88 1.22 0.34 33.6 2.50 0.01 N.D. N.D. 0.01 0.01 N.D. 97.17 14.8 16.7

Po 140 ⫻ 140 14 28.7 0.42 37.2 9.26 0.63 10.2 8.42 0.04 N.D. 3.77 0.05 N.D. N.D. 98.71 0.35 0.65

CC 120 ⫻ 90 12 45.7 0.15 3.13 11.3 0.52 36.8 2.67 0.02 N.D. 0.38 0.15 N.D. N.D. 100.87 14.9 12.4

BO 125 ⫻ 105 10 34.9 0.13 2.96 35.4 0.17 24.9 2.08 0.02 N.D. 0.18 0.15 0.02 N.D. 100.88 10.6 10.0

a

All sample names except for F96g343 should be prefaced with “Y98”. N.D. ⫽ not detected. N.A. ⫽ not analyzed.

Table 1. (Continued) Sample No.

K02KS189

K02KS209

K02KS269

K02KS281

K02KS300

KW54064

KW54087

KW54094

KW540100

KW540204

KW540216

KW540219

KW540235

KW540241

KW540260

KW540265

KW540273

Type Size (␮m) Analysis point SiO2 (wt.%) TiO2 Al2O3 FeO MnO MgO CaO Na2O K2 O Cr2O3 NiO P2O 5 S Total Mg/Al (atom) Si/Al (atom)

BO 95 ⫻ 90 8 41.0 0.09 2.71 25.3 0.44 27.2 2.02 0.02 N.D. 0.24 0.01 0.18 N.D. 99.16 12.7 12.9

CC 120 ⫻ 105 10 41.0 0.13 4.08 19.4 0.45 31.9 2.70 0.01 N.D. 0.08 0.96 0.02 N.D. 100.66 9.9 8.5

BO 185 ⫻ 175 5 48.2 0.17 2.88 19.4 0.50 26.4 2.52 0.04 N.D. 0.47 0.14 N.D. N.D. 100.66 11.6 14.2

BO 130 ⫻ 130 8 51.2 0.19 3.52 13.7 0.18 26.9 3.45 0.02 N.D. 0.88 0.05 0.07 0.01 100.22 9.6 12.3

Glass 300 ⫻ 300 5 42.0 0.22 4.84 10.7 0.15 38.4 3.35 0.01 N.D. 0.01 N.D. N.D. N.D. 99.65 10.0 7.4

BO 100 ⫻ 70 8 47.9 0.03 0.48 17.8 0.62 32.1 0.37 N.D. N.D. 0.33 0.07 0.04 N.D. 99.78 85.2 85.2

BO 55 ⫻ 50 11 39.0 0.02 0.06 29.9 0.42 30.4 0.05 0.01 N.D. 0.01 0.01 0.01 N.D. 99.94 633 546

BO 70 ⫻ 55 9 39.4 0.15 3.24 17.9 0.12 35.6 2.81 0.02 N.D. 0.07 1.32 N.D. N.D. 100.64 13.9 10.3

CC 110 ⫻ 80 12 43.6 0.15 3.65 23.8 0.44 25.4 2.37 0.02 N.D. 0.22 1.09 N.D. N.D. 100.83 8.8 10.1

BO 65 ⫻ 65 8 34.1 0.10 3.21 33.5 0.27 15.3 12.4 0.01 N.D. 0.17 0.96 0.15 0.01 100.09 6.0 9.0

Glass 55 ⫻ 55 8 56.6 0.14 0.23 11.2 0.47 30.7 0.63 0.03 N.D. 0.02 0.01 N.D. N.D. 100.10 168 207

BO 75 ⫻ 60 11 38.3 0.03 0.64 25.6 0.64 34.1 0.55 0.01 N.D. 0.22 0.02 0.01 N.D. 100.13 67.5 50.8

Glass 85 ⫻ 85 7 49.0 0.10 3.25 18.3 0.32 28.9 1.42 0.01 N.D. 0.03 0.02 0.01 N.D. 101.35 11.2 12.8

BO 50 ⫻ 50 7 35.1 0.10 2.01 33.0 0.30 27.6 1.28 0.01 N.D. 0.79 0.71 0.03 N.D. 100.92 17.4 14.8

RGB 100 ⫻ 90 12 35.7 0.10 1.96 24.0 0.16 36.1 2.57 0.02 N.D. 0.19 0.26 0.12 0.01 101.12 23.3 15.4

RGB 75 ⫻ 75 8 37.8 0.08 1.73 20.0 0.17 39.2 2.01 0.01 N.D. 0.22 0.17 0.14 0.14 101.66 28.7 18.5

BO 70 ⫻ 50 7 38.4 0.14 3.27 18.4 0.07 36.2 2.85 0.01 N.D. 0.26 1.53 0.07 N.D. 101.25 14.0 10.0

41.7 0.17 4.11 18.6 0.27 31.9 2.85 0.01 0.01 0.31 0.36 0.01 100.37 35.7 31.6 CC 135 ⫻ 130 12 56.2 0.06 0.60 4.19 0.32 37.6 0.54 N.D. 0.01 0.50 0.08 N.A. N.D. 100.11 79.0 79.4 Glass 155 ⫻ 150 18 49.9 0.13 3.19 12.0 0.37 32.3 2.01 N.D. N.D. 0.03 0.02 N.A. N.D. 100.00 12.8 13.3 BO 180 ⫻ 165 8 40.3 0.07 3.01 26.7 0.15 30.4 0.08 N.D. 0.01 0.28 0.25 N.A. 0.01 101.27 12.8 11.4 BO 210 ⫻ 200 17 38.4 0.11 1.92 28.2 0.27 29.5 0.96 N.D. N.D. 1.18 0.54 N.A. 0.01 101.03 19.4 16.9 Po 130 ⫻ 125 12 45.0 0.13 3.73 20.1 0.32 28.5 1.00 N.D. 0.22 0.47 0.08 N.A. 0.16 99.71 9.7 10.2 BO 145 ⫻ 140 19 33.6 0.18 6.09 46.6 0.29 9.55 3.13 N.D. N.D. 0.44 0.27 N.A. 0.01 100.22 2.0 4.7 CC 80 ⫻ 55 9 41.4 0.16 4.04 16.0 0.02 35.1 3.14 0.02 N.D. 0.02 1.08 0.02 N.D. 100.91 11.0 8.7 CC 60 ⫻ 45 8 41.2 0.16 3.69 15.8 0.17 36.8 3.06 N.D. N.D. 0.04 0.07 N.A. 0.01 101.00 12.6 9.5 Glass 65 ⫻ 60 7 43.6 0.05 1.67 0.25 0.02 54.7 0.94 N.D. 0.01 N.D. 0.01 0.01 N.D. 101.24 41.3 22.1 CC 75 ⫻ 45 8 40.6 0.18 3.86 12.2 0.17 40.2 2.37 0.02 N.D. 0.18 0.01 0.05 N.D. 99.86 13.2 8.9 Glass 60 ⫻ 60 7 49.8 0.15 3.33 7.89 0.39 35.0 2.49 N.D. N.D. 0.02 0.01 N.D. N.D. 99.10 13.3 12.7 Glass 75 ⫻ 75 8 45.6 0.13 2.97 12.7 0.22 32.6 4.96 0.01 N.D. 0.12 0.02 0.01 0.01 99.40 13.9 13.0 BO 65 ⫻ 60 10 38.0 0.14 2.65 31.8 0.17 25.7 1.46 0.01 N.D. 0.13 0.02 0.05 N.D. 100.16 12.3 12.2 Type Size (␮m) Analysis point SiO2 (wt.%) TiO2 Al2O3 FeO MnO MgO CaO Na2O K 2O Cr2O3 NiO P 2O 5 S Total Mg/Al (atom) Si/Al (atom)

Po 110 ⫻ 110 9 37.0 0.12 2.24 31.0 0.26 27.6 1.58 0.02 N.D. 0.52 0.52 0.15 0.01 100.97 15.6 14.0

Average K11KS248 K11KS246 K11KS239 K11KS236 K11KS234 K11KS233 K11KS147 K11KS142 KW540319 KW540308 KW540301 KW540299 KW540279 KW540277 Sample No.

Table 1. (Continued)

Na, S and P depletions were considered to represent preferential evaporative loss during atmospheric entry heating. Fe and Ni depletions are explained by preferential evaporation, or the extraction of Fe-Ni metals from silicate melts due to large inertial force during their atmospheric entry (Brownlee et al., 1984). Note that the depletions of Fe and Cr seem to be correlated with the types of spherules. Looking at individual types of spherules, the BO spherules consist mainly of barred olivine, magnetite and interstitial glass (Fig. 1a). They are less depleted in volatile elements among the six types, and in particular are the least depleted in Mn (Fig. 2). The Glass spherules are composed of homogeneous glass and depleted in Cr, Fe, Na and P relative to the other types of spherules except for the CAT spherule (Figs. 1b and Fig. 2). However, they are also depleted in refractory elements such as Al, Ca and Ti relative to CI. The CC spherules are either heterogeneous in chemical composition, or contain fine-grained minerals as observed in cross section (Fig. 1c). They display a flat pattern for the refractory elements and are depleted in Cr, Fe, Na, P and S after CAT and Glass type spherules (Fig. 2). The Po spherules consist of porphyritic olivine and magnetite ⬍ 5 ␮m in size and interstitial glass (Fig. 1d). They are the least depleted in volatile elements except for Mn and P (Fig. 2). The RGB spherules are composed mainly of relict anhydrous mineral grains 5 to 50 ␮m in size (Fig. 1e). They have less Al, Ti and Ca than CI and contain more Mg, because most of them are composed mainly of forsterite (Fig. 2). One of the RGB spherules seems to preserve amoeboid olivine aggregates (AOA) in the interior (Fig. 1e). The amoeboid olivines are MgO-rich (Fo97.4 –99.3), whereas the surrounding euhedral olivines are Fo75.2 and are thought to have grown secondarily from a eutectic melt. The CAT spherule is the most enriched in refractory elements such as Al, Ca, Ti and Mg and the most depleted in volatile and moderately volatile elements such as S, P, Na, Fe, Mn and Cr among all the types of spherules (Fig. 2). It consists of homogeneous glass (Fig. 1f), although the CAT spherules reported by Taylor et al. (2000) have barred textures. Of the various spherule types, the BO, Glass, CC, and CAT types are thought to have experienced total melting during atmospheric entry. Their Si-Mg-Fe compositions are shown in Figure 3. Tephra particles found in Antarctic blue ice and microtektites found around Australia are also plotted in the diagram (Katsushima et al., 1984; Glass et al., 2004). They are more silicic than S-type spherules, showing obviously different chemical compositions than the spherules. Most of the spherules plot between the olivine (Ol) and orthopyroxene (Opx) solid solution lines. The Fe contents generally decrease in the order: Po, BO ⬎ CC ⬎ Glass ⬎ CAT (see Table 2). The results of the oxygen isotopic analyses are shown in Table 3. Most of the data in Table 3 represent averages of two to six sets of replicate analyses of different spots. Some spherules were analyzed only once because they were too small for duplicate analyses. In some cases, multiple analyses of the same spherule gave oxygen isotopic ratios that differed by ⬎ 2␴ errors (in these cases, all the data for the spherules are shown in Table 3). Some of these differences can be explained by heterogeneity in the spherules because of the existence of relict grains (e.g., Fig. 1e), but some of them cannot be explained by the relict grains, because they have totally homogeneous internal textures (e.g., Fig. 1b). In the case that a spherule 5793

5794

T. Yada et al. Table 2. The averages and standard deviations of electron microprobe data for individual types of spherules. Barred olivine (x ⫾ sd), n ⫽ 19

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Cr2O3 NiO Total

39.7 0.2 3.2 25.0 0.3 28.3 2.8 0.3 0.5 100.4

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

4.9 0.3 3.0 9.1 0.2 7.4 3.4 0.3 0.6 0.6

Glass (x ⫾ sd), n ⫽ 10 48.5 0.1 2.7 9.2 0.3 36.8 2.2 0.0 0.0 99.9

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

5.0 0.0 1.2 5.9 0.1 8.5 1.5 0.0 0.0 1.2

Cryptocrystalline (x ⫾ sd), n ⫽ 10 43.7 0.2 3.3 15.4 0.3 34.4 2.9 0.2 0.4 100.6

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

was fairly large (⬎100 ␮m in diameter) and an intense evaporation occurred during atmospheric entry, its oxygen isotopic ratio might not be homogenized in the few seconds available (Love and Brownlee, 1991). Engrand et al. (1998) also noted heterogeneities in the oxygen isotopic ratios of the S-type spherules they measured. The S-type spherules show a wide range of oxygen isotopic ratios from ⫺28‰ to ⫹93‰ in ␦18OSMOW and from ⫺35‰ to

5.0 0.0 1.1 5.9 0.2 4.3 1.1 0.2 0.5 0.4

Porphylitic (x ⫾ sd), n⫽5 35.4 0.2 11.3 23.9 0.3 23.7 3.5 1.1 0.7 100.3

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

6.2 0.1 15.0 9.5 0.2 7.6 2.9 1.5 1.1 1.0

Relict grain-bearing (x ⫾ sd), n⫽3 38.2 0.1 1.4 17.7 0.2 41.3 1.7 0.3 0.2 101.3

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

2.7 0.0 0.7 7.7 0.0 6.6 1.0 0.1 0.1 0.5

CAT (n ⫽ 1) 35.0 0.6 14.8 0.0 0.0 39.4 9.5 0.0 0.0 99.3

⫹46‰ in ␦17OSMOW (Table 3). The oxygen isotopic compositions of our spherules are plotted on a three-isotope diagram in Figure 4a. The terrestrial fractionation line (TFL) represents the mass-dependent fractionation line for terrestrial materials and its slope is 0.52. In contrast, C chondrite anhydrous minerals generally plot on the CCAM line with a slope of ⬃1. No obvious correlation can be observed between oxygen isotope compositions and internal textural types, as shown in Figure 4b,

Fig. 2. Abundances of major and minor elements in Antarctic S-type spherules normalized to Si and CI (Lodders and Fegley, 1998). An average for each of the six groups of spherules is shown in the graph. Note that the statistics of the phosphorous data are poor because ten of spherules were not analyzed for the element.

Oxygen isotope compositions of stony cosmic spherules

5795

Fig. 3. Relative Si-Mg-Fe compositions of 48 S-type spherules. Glass, CC and CAT types of spherules are clearly depleted in Fe relative to spherules of other types. Data from tephra in Antarctic ice and Australian microtektites are also plotted (Katsushima et al., 1984; Glass et al., 2004).

although it is worthwhile to note that all three RGB spherules plot on or close to the CCAM line. Also, we did not observe any correlation between the size of the spherules and their oxygen isotopic ratios, which is consistent with the results of Clayton et al. (1986). Note that the average of all the S-type spherules (␦18OSMOW ⫽ 23.1‰, ␦17OSMOW ⫽ 9.8‰) is comparable to the literature values determined by conventional isotopic analysis (␦18OSMOW ⫽ 22.5‰–25.2‰, ␦17OSMOW ⫽ 9.5‰–11.0‰; Clayton et al., 1986). This indicates that our data do not suffer any systematic fractionation due to instrumental problems during the SIMS analysis. Reanalysis of four spherules (M240u02, M03KS063, KW540260 and KW540319) was performed with CAI minerals in Allende on the same mount of the spherules, to confirm the absence of systematic instrumental mass fractionation. The Allende sample used in the reanalysis was detailed in Akaki and Nakamura (2005). The reanalyzed data for the four spherules reproduced our previous data, and the CAI minerals plot on the CCAM line in the range of their 2␴ errors (see Table 3 and Fig. 4a). These results also prove the absence of any systematic fractionation due to instrumental problems in our SIMS analysis. 4. DISCUSSION

4.1. Factors Changing Oxygen Isotopic Ratios of Spherules—Evaporation and Atmospheric Oxygen Mixing The oxygen isotopic compositions of S-type spherules show dispersive distributions in the ␦18OSMOW-␦17OSMOW plot (Fig.

4a) that are more variable than the bulk oxygen isotopic compositions of known meteorites (Fig. 4c). The variation of their oxygen isotopic ratios includes the variation of the previous data by Engrand et al. (1998) and Taylor et al. (2005), but is much larger than theirs. The variations are considered to result from original features of precursor materials of the S-type spherules as well as secondary effects that might have occurred since their entry to the Earth’s atmosphere and storage in the glacier ice. First, we discuss possible secondary effects: massdependent fractionation due to evaporation and mixing of atmospheric oxygen with the cosmic spherules during their atmospheric entry. The spherules have experienced melting and subsequent evaporation more or less at the time of their atmospheric entry. Significant evaporation should induce both elemental and isotopic fractionation. The elemental fractionation causes depletions of relatively volatile elements and enrichment of refractory elements of the spherules. For example, Fe is the most volatile element among major elements such as Si, Mg, Fe, Al and Ca. Thus spherules that are depleted in Fe might have experienced more evaporation and oxygen isotopic fractionation. In this sense, CAT spherules, which have almost no FeO, should have experienced more severe evaporation and subsequent isotopic fractionation than other spherules. For example, the isotopic fractionation of oxygen was roughly estimated to be 15‰ to 26‰ in ␦18OSMOW based on the observed Mg, Si and Fe isotopic fractionation in one of CAT spherules (Alexander et al., 2002). Since CAT spherules should be one of the most heavily fractionated cases in oxygen isotopes, most spher-

5796

T. Yada et al. Table 3. Oxygen isotopic compositions of 48 S-type spherules and CAI minerals in Allende. Sample no.

F96g343a M240168 M240u02-1 M240u02-2 M240u02-3b M240u02-4b M240u02 AVG MY240159 MY240184 M03KS012 M03KS019 M03KS021 M03KS039 M03KS051 M03KS057 M03KS063-1 M03KS063-2b M03KS063 AVG M03KS066 M03KS071 M03KS115-1 M03KS115-2 M03KS115-3 M03KS115-4 M03KS115-5 M03KS115-6 M03KS115 AVG K02KS086 K02KS135 K02KS189 K02KS209 K02KS269 K02KS281-1 K02KS281-2 K02KS281-3 K02KS281 AVG K02KS300-1 K02KS300-2 K02KS300-3 K02KS300-4 K02KS300-5 K02KS300 AVG KW54064 KW54087 KW54094 KW540100-1 KW540100-2 KW540100-3 KW540100-4 KW540100 AVG KW540204 KW540216 KW540219 KW540235 KW540260-1 KW540260-2 KW540260-3b KW540260-4b KW540260-5b KW540260-6b KW540260 AVG KW540241 KW540265 KW540273 KW540277 KW540279-1 KW540279-2

Type

N

Oxygen isotopic group

␦18OSMOW (‰)

Error (2␴)

␦17OSMOW (‰)

Error (2␴)

⌬17O (‰)

Error (2␴)

Mg/Al (atom)

Si/Al (atom)

Po BO — — — — BO Glass CAT CC Po CC RGB Glass BO — — CC BO Glass — — — — — — Po CC BO BO CC BO — — — BO — — — — — Glass BO BO BO — — — — CC BO Glass BO Glass — — — — — — RGB BO RGB BO BO — —

1 2 — — — — 2 2 1 4 4 6 2 2 2 — — 1 2 2 — — — — — — 6 4 2 2 2 4 — — — 3 — — — — — 5 1 1 2 — — — — 4 1 2 1 2 — — — — — — 2 1 2 1 2 — —

TFL O-rich — — — — 16 O-rich TFL TFL TFL TFL 16 O-poor 16 O-rich TFL TFL — — 16 O-poor TFL 16 O-rich — — — — — — TFL TFL TFL TFL TFL TFL — — — TFL — — — — — TFL TFL TFL TFL — — — — TFL TFL TFL TFL TFL — — — — — — 16 O-rich TFL 16 O-rich TFL TFL — —

24.7 –16.7 –3.3 –3.3 –0.7 0.2 –1.8 26.2 93.0 19.6 7.9 56.5 3.3 36.9 24.0 51.5 52.1 51.8 9.2 2.3 22.3 22.3 11.9 13.3 12.0 6.7 14.7 10.7 40.8 30.1 30.3 15.3 12.6 9.4 5.8 9.3 44.9 48.5 29.1 49.6 54.6 45.3 20.7 38.5 8.1 21.6 24.7 27.3 27.3 25.2 41.9 8.3 29.6 31.2 –14.0 –18.1 –18.3 –11.8 –18.2 –15.8 –16.0 16.9 0.2 27.5 36.9 37.3 31.7

2.6 3.6 2.8 2.8 5.5 5.5 — 5.0 3.8 4.4 2.7 4.5 3.0 3.5 3.7 3.8 3.9 — 4.3 2.7 2.7 2.7 2.7 2.7 2.7 2.7 — 4.2 3.4 4.4 2.9 3.4 3.1 3.1 3.1 — 3.2 3.2 3.2 3.2 3.3 — 3.0 2.6 4.7 3.1 3.1 3.1 3.1 — 3.0 3.6 4.7 3.6 2.8 2.8 4.3 4.3 4.3 4.3 — 3.5 4.8 3.2 3.6 3.6 3.5

11.2 –29.2 –15.3 –15.9 –13.3 –12.9 –14.2 10.5 46.2 8.5 0.0 34.0 –1.6 19.9 9.9 40.0 39.9 40.0 6.0 –3.9 8.4 9.2 2.5 3.7 1.7 0.3 4.3 5.5 21.5 14.0 14.2 7.3 7.4 4.6 1.7 4.5 21.3 22.7 15.0 23.5 26.4 21.8 10.1 22.0 –0.9 13.1 13.6 16.6 17.9 15.3 19.1 4.9 13.8 17.2 –22.3 –25.1 –27.2 –24.0 –20.2 –20.6 –23.3 9.1 –8.4 11.6 19.2 17.0 13.7

3.1 2.7 2.0 1.7 2.9 2.9 — 2.6 2.2 3.3 2.1 3.4 2.5 2.6 2.6 2.6 2.6 — 3.2 2.2 2.3 2.2 2.2 2.2 2.2 2.2 — 2.8 1.6 2.9 2.3 2.6 1.7 1.6 1.7 — 1.6 1.6 1.6 1.6 1.6 — 2.3 3.2 2.5 2.5 2.5 2.5 2.5 — 2.4 2.4 2.8 2.6 1.8 1.8 3.1 3.1 3.1 3.1 — 2.0 3.4 1.6 2.0 2.0 2.0

–1.6 –20.5 –13.6 –14.2 –12.9 –13.0 –13.3 –3.2 –2.1 –1.7 –4.2 4.6 –3.3 0.7 –2.5 13.2 12.9 13.0 1.2 –5.1 –3.2 –2.4 –3.7 –3.3 –4.5 –3.2 –3.4 0.0 0.3 –1.6 –1.6 –0.6 0.8 –0.3 –1.4 –0.3 –2.1 –2.5 –0.1 –2.2 –2.0 –1.8 –0.6 1.9 –5.1 1.9 0.8 2.4 3.8 2.2 –2.7 0.6 –1.6 0.9 –15.0 –15.7 –17.7 –17.9 –10.8 –12.4 –14.9 0.3 –8.6 –2.7 0.0 –2.4 –2.8

2.6 3.0 2.3 2.1 1.7 1.7 — 3.0 2.0 2.4 2.0 2.4 2.8 2.3 2.7 2.3 1.9 — 3.4 2.2 2.2 2.1 2.1 2.1 2.1 2.1 — 2.2 2.1 3.3 1.9 2.2 2.1 2.1 2.1 — 2.1 2.1 2.1 2.1 2.1 — 2.5 2.6 2.2 2.7 2.7 2.7 2.7 — 2.0 2.2 2.1 2.5 2.2 2.2 2.1 2.1 2.1 2.1 — 1.8 3.0 2.1 2.5 1.8 1.8

2.9 2.3 — — — — 39.1 34.1 3.4 10.9 18.0 16.9 99.5 13.6 13.7 — — 11.5 8.6 14.8 — — — — — — 0.35 14.9 10.6 12.7 9.9 11.6 — — — 9.6 — — — — — 10.0 85.2 633 13.9 — — — — 8.8 6.0 168 67.5 11.2 — — — — — — 23.3 17.4 28.7 14.0 12.3 — —

2.3 2.2 — — — — 24.8 20.9 2.0 10.0 16.4 15.6 56.2 13.1 9.8 — — 8.6 10.7 16.7 — — — — — — 0.65 12.4 10.0 12.9 8.5 14.2 — — — 12.3 — — — — — 7.4 85.2 546 10.3 — — — — 10.1 9.0 207 50.8 12.8 — — — — — — 15.4 14.8 18.5 10.0 12.2 — —

16

Oxygen isotope compositions of stony cosmic spherules

5797

Table 3. (Continued)

␦18OSMOW (‰)

Error (2␴)

␦17OSMOW (‰)

Error (2␴)

⌬17O (‰)

Error (2␴)

Mg/Al (atom)

Si/Al (atom)

— TFL 16 O-poor TFL TFL — — 16 O-rich TFL TFL TFL TFL TFL TFL — — — — TFL — — — TFL

33.2 34.0 48.5 49.0 10.3 –26.2 –30.3 –28.2 16.8 19.4 47.5 37.1 10.7 25.7 21.7 23.6 14.2 22.6 20.5 16.2 14.7 11.1 14.0

3.5 — 3.1 3.2 2.9 2.6 3.6 — 4.2 4.4 3.0 3.1 3.6 2.9 3.7 3.4 3.4 3.4 — 2.9 2.9 2.9 —

16.0 15.6 29.4 25.8 1.5 –35.7 –33.6 –34.7 7.6 6.2 22.0 21.2 6.3 9.7 10.3 12.6 7.6 13.6 11.0 7.4 7.8 4.1 6.5

2.0 — 2.4 3.0 2.3 3.1 2.4 — 2.9 2.4 2.5 3.0 2.4 2.9 2.6 3.3 3.3 3.3 — 2.2 2.2 2.2 —

–1.3 –2.1 4.2 0.3 –3.9 –22.0 –17.9 –19.9 –1.1 –3.8 –2.7 1.9 0.7 –3.6 –1.0 0.3 0.2 1.8 0.3 –1.0 0.2 –1.6 –0.8

1.8 — 2.0 3.0 2.5 2.7 1.8 — 2.7 2.9 2.1 3.2 2.1 3.0 2.4 2.2 2.2 2.2 — 2.4 2.4 2.4 —

— 15.6 13.9 13.3 13.2 — — 41.3 12.6 11.0 2.0 9.7 19.4 12.8 — — — — 12.8 — — — 79.0

— 14.0 13.0 12.7 8.9 — — 22.1 9.5 8.7 4.7 10.2 16.9 11.4 — — — — 13.3 — — — 79.4

—

—

23.1

—

9.8

—

–2.2

—

35.7

31.6

— — — — —

— — — — —

–17.3 –13.8 –11.7 –11.2 –26.5

2.6 2.6 4.2 4.2 3.6

–22.0 –18.9 –16.9 –17.5 –26.7

1.8 1.8 3.3 3.3 2.5

–13.0 –11.7 –10.9 –11.7 –12.9

1.6 1.6 2.0 2.0 1.9

— — — — —

— — — — —

Type

N

Oxygen isotopic group

— Po Glass Glass CC — — Glass CC CC BO Po BO BO — — — — Glass — — — CC

— 3 2 1 1 — — 1 2 4 3 2 4 3 — — — — 4 — — — 3

48 spherules average

—

Al21-SP2-1 (Sp ⫹ Mel)c Al21-SP2-2 (Sp ⫹ Mel)c Al21-Mel4-1 (Mel ⫹ Sp)c Al21-Mel4-2 (Mel ⫹ Sp)c Al21-SP5-1 (Sp ⫹ Mel)c

— — — — —

Sample no. KW540279-3 KW540279 AVG KW540299 KW540301 KW540308 KW540319-1 KW540319-2b KW540319 AVG K11KS142 K11KS147 K11KS233 K11KS234 K11KS236 K11KS239 K11KS246-1 K11KS246-2 K11KS246-3 K11KS246-4 K11KS246 AVG K11KS248-1 K11KS248-2 K11KS248-3 K11KS248 AVG

a

All spherules’ names except for F96g343 should be prefaced with “Y98”. Reanalyzed data in 2004. c These CAI minerals in Allende were analyzed on the same mount with the reanalyzed spherules. b

ules other than CAT should have experienced less isotopic fractionations of oxygen than the CAT spherule have, such as ⬍ 15‰ to 26‰ in ␦18OSMOW. The mixing of atmospheric oxygen with the spherules is another factor that might change their isotopic compositions from those of the precursor materials. This would have taken place mainly during melting, because diffusion of oxygen in silicate melts is much faster than in solid silicates. O2 is more dominant than atomic O or CO2 in upper mesosphere, where a particle entering the Earth’s atmosphere experiences the frictional heating with atmospheric atoms and molecules. It is possible to calculate an amount of O2 hitting the melted spherule’s surface. However, it is difficult to estimate the degree of mixing of atmospheric O2 with the spherule by calculation, because the probability of a reactive collision when O2 hits the surface of silicate melt is poorly known (Yu et al., 1995). Thiemens et al. (1995) has confirmed that the oxygen isotopic ratio of atmospheric O2 was constant (⫹23.5‰ in ␦18OSMOW on the TFL) up to in an altitude of 60.9 km. Thus the oxygen isotopic ratio of spherules will tend to plot closer to the atmospheric value as the atmospheric oxygen mixes with the spherules (Fig. 4c). Actually, the average oxygen isotopic ratio of all 48 spherules, ⫹23‰ in ␦18OSMOW, is close to the atmospheric value, although it is off the TFL. This may indicate certain amounts of atmospheric oxygen could have mixed with spher-

ules during their atmospheric entry. The mixing of atmospheric O2, whose ⌬17O is zero by definition, with spherules decreases the absolute values of their ⌬17O. For example, if the original ⌬17O value of M240168 (⌬17O ⫽ ⫺21‰) were ⫺24‰, which corresponds to the typical lowest value of ⌬17O of CAI minerals in meteorites, the ratio of mixing atmospheric oxygen is calculated to be ⬃13%. It can be said that this ratio should be a nominal upper limit of mixing atmospheric oxygen for this spherule. Although it is impossible to estimate exactly how much atmospheric oxygen mixed with spherules, it can be said that the quantity should not be significant, ⬍ 13%, because the ⌬17O values of the 48 S-type spherules are already more varied than those of meteorites. Based on the observed ⌬17O values, we divided the fortyeight S-type spherules into three groups as shown in Table 3 and Figure 5a: a TFL group (38 spherules) that plots close to the TFL, a 16O-rich group (7 spherules) that plots close to the CCAM, and a 16O-poor group (3 spherules) that has a ⌬17O more positive than 4‰. Any mass-dependent fractionation should shift the oxygen isotopic compositions of the spherules parallel to the TFL and should not change their ⌬17O values. However, the mixing of atmospheric oxygen may decrease the absolute values of ⌬17O compared with those of their precursors. Thus, high absolute values of ⌬17O observed in 16O-rich or -poor spherules must come from the precursors. In contrast,

5798

T. Yada et al.

Fig. 4. (a) An oxygen three isotope plot of 48 S-type spherules classified according to the six types (see text). A typical 2␴ error is indicated by the cross. The isotopic compositions of the spherules vary widely from ⫺26‰ to ⫹93‰ in ␦18OSMO W and from ⫺35‰ to ⫹46‰ in ␦17OSMOW. (b) The same plot as (a) with respect to classifications based on their internal textures. (c) An oxygen three isotope plot showing chondritic and atmospheric oxygen isotopic reservoirs that may be related to the spherules. An enlarged view of the area enclosed by dotted lines is shown to the right. Data for the matrices of C chondrites are from Clayton and Mayeda (1999) and those of mesospheric CO2 are from Thiemens et al. (1995).

some TFL spherules might originally have had more negative or positive values of ⌬17O but lost them due to the dilution by atmospheric oxygen, because there is a possibility that such spherules would come closer to the TFL due to relatively high degrees of mixing with atmospheric oxygen than others. The precursors of the spherules may be generally considered to be similar to C chondrites, because the average oxygen

isotopic compositions plot beneath the TFL (⌬17O ⫽ ⫺2.2‰). This is consistent with the results of Clayton et al. (1986) and Engrand et al. (1998, 1999), and the interpretations that the most probable precursors of unmelted MMs are hydrated C chondrites: CI chondrites, CM chondrites, and Tagish Lake (Kurat et al., 1994; Genge et al., 1997; Nakamura et al., 2001; Noguchi et al., 2002). However, the wide variation in ⌬17O,

Oxygen isotope compositions of stony cosmic spherules

Fig. 5. (a) A plot of bulk Mg/Al vs. Si/Al ratios of 48 S-type spherules. An enlarged view of the area enclosed by dotted lines is shown to the right. Most spherules (⬃40%) plot around the CI value. (b) The same plot as (a) with respect to oxygen isotopic classifications. (c) Mg/Al vs. Si/Al ratios of possible precursors are shown for comparison. Data for the matrices of C chondrites are from Zolensky et al. (1993), and those for F96 unmelted AMMs are from Imae and Noguchi (unpublished data).

5799

5800

T. Yada et al.

ranging from ⫺21‰ to ⫹13‰, indicates that the spherules originate from various kinds of planetary materials. This idea is consistent with the concept that interplanetary dust reaching the Earth must originate from a more varied collection of bodies orbiting heliocentrically outside the Earth than just meteorites. Below, we further discuss the precursors of the spherules classified in this study in view of their oxygen isotopic and chemical variations. 4.2. TFL Group of Spherules This group is dominant among the analyzed spherules, which is consistent with the results of Engrand et al. (1998) and Taylor et al. (2005). The spherules of this group show the greatest mass-dependent variations among the three groups and other meteoritic materials. Since various kinds of planetary materials, including terrestrial materials, plot along the TFL with ⌬17O of ⫾4‰, it is difficult to estimate the precursor materials of the spherules from the oxygen isotopic compositions alone. Fortunately, their bulk chemical compositions help constrain their precursors. Figure 5a shows a plot of Mg/Al vs. Si/Al for the S-type spherules. The CI value is given for comparison. The ratios of these three refractory lithophile elements are distinctively different from one chondrite group to another and can be used to understand the evolution of chondritic materials (Larimer and Wasson, 1988). Alexander et al. (2002) and Taylor et al. (2005) also applied this diagram to cosmic spherules for the discussion about evaporation they experienced and estimation of their precursors. Six of our TFL spherules have Mg/Al and Si/Al ratios ⬎ 20 (Fig. 5b). Spherules with Mg/Al and Si/Al ratios ⬎20 are depleted in Al relative to CI. Most olivines and low-Ca pyroxenes in chondrites have Mg/Al and Si/Al ratios ⬎ 20 (Brearley and Jones, 1998) and, thus, they are possible precursors of these spherules. This is consistent with the observation of unmelted, crystalline MMs, which typically consist of crystalline olivines or low-Ca pyroxenes (Kurat et al., 1994). Brownlee et al. (1997) suggested that these spherules originated from “coarse-grained precursors.” If these were originally olivines and low-Ca pyroxenes with ⌬17O within ⫾4‰, there are a lot of candidate precursors, such as olivines and pyroxenes in ordinary (O), enstatite (E) and Rumuruti (R) chondrites, isolated olivines or olivines and pyroxenes from chondrules in C chondrites and those which originated from achondrites. In addition, precursors sometimes cannot be explained by known meteoritic materials, by taking their FeO contents and ␦18OSMOW values into account. For example, the bulk chemistry of KW54087 is close to that of olivine with ⌬17O of 1.9‰. However, its high FeO content (29.9 wt.%) suggests that it did not experience severe evaporation, thus its ␦18OSMOW value of 38.5‰ which is much higher than that of O chondrite (4‰– 6‰) cannot be explained by only mass-dependent fractionation due to evaporation during its atmospheric entry. The 28 TFL spherules that plot around the CI value in the Mg/Al vs. Si/Al diagram might be fine-grained chondritic materials (Fig. 5b). Possible precursors are materials like the matrices of hydrated C chondrites. The oxygen isotopic compositions of hydrated matrices of C chondrites have ⌬17O values ranging from ⫺2‰ to ⬃0‰ (Fig. 4c, Clayton and

Mayeda, 1999). The Mg/Al and Si/Al ratios of C chondrite matrices almost overlap those of these TFL spherules, as well as those of unmelted MMs (Fig. 5c, Zolensky et al., 1993; Imae and Noguchi, unpublished data). Four TFL spherules have low Mg/Al and Si/Al ratios (⬍5) in their bulk chemical compositions. This can be interpreted in two ways: one is a result of evaporational elemental fractionation from materials whose bulk Mg/Al and Si/Al ratios were originally similar to those of CI, and another interpretation is that the precursors were originally Al-rich such as CAI minerals or mesostasis of chondrules. The former explanation fits MY240184 well. This is the only CAT spherule and it has the highest oxygen isotopic composition among the analyzed spherules (␦18OSMOW ⫽ 93‰). It is also largely depleted in moderately volatile elements such as Fe, Mn, and Cr and is enriched in refractory elements such as Al, Ca, and Ti, relative to CI. This refractory elemental composition and heavy oxygen isotopic ratio are also described in Taylor et al. (2000, 2005). This type of spherule also has fractionated isotopic ratios for Mg, Si and Fe (Alexander et al., 2002). It is thought to be the evaporation remnant of the material whose chemical composition is similar to CI. The spherule appears to have experienced strong heating and evaporation during atmospheric entry, resulting in extreme mass-dependent isotopic fractionation as well as elemental fractionation. The bulk chemical composition of MY240184 is comparable to that of evaporation residues after 83% mass lost starting from a material of solar composition, applying the evaporation experimented data of Wang et al. (2001). In their paper, the oxygen isotopic ratio of the residue would have been fractionated by 37‰ in ␦18OSMOW relative to CI starting material. If we apply this result of the evaporation experiment to MY240184, the oxygen isotopic composition of its precursor should be 56‰ in ␦18OSMOW, which is out of the range of those of CI chondrites, the chondrite group having the largest ␦18OSMOW value (Fig. 4c). This mismatch might be explained by the different evaporation conditions in the experiments of Wang et al. (2001) from that of MY240184. 4.3.

16

O-Rich Group of Spherules

There are seven spherules that are categorized into a 16O-rich group. They are distributed along the CCAM line in the oxygen three isotope plot (Fig. 4a). Five of the spherules have 2␴ errors that touch the CCAM line, but M240u02 and M240168 plot significantly to the right of the CCAM line. They partly overlap the distribution of FUN (Fractionated with Unknown Nuclear effects) inclusions (Lee and Papanastassiou, 1974; Davis et al., 2000); however, FUN inclusions are not likely the precursors of the spherules, because they are very rare in CAIs. If we consider that the deviation from the CCAM line is not a primary feature, but is due to secondary effects, three interpretations can be proposed; the first is that the observed oxygen isotopic deviations from the CCAM line are due to mass-dependent fractionation during atmospheric entry heating, the second is that mixing with atmospheric oxygen (23.5‰ on the TFL) is responsible, and the third is that the oxygen isotopic feature is an artifact caused by matrix effects during SIMS analysis. If we assume that the precursors of 16O-rich spherules plot exactly on the CCAM line, the latter two factors are considered to be minor for the following reasons. Mixing with atmospheric

Oxygen isotope compositions of stony cosmic spherules

oxygen alone fails to explain the shift, because the line between the two 16O-rich spherules and atmospheric O2 is almost parallel to the CCAM line on the oxygen three-isotope plot (Figs. 4a and 4c). This indicates that it is impossible to reproduce the isotopic compositions of the two spherules simply by the mixing of the atmospheric O2 with CCAM precursors. Matrix effects alone cannot explain the deviation because the massdependent fractionation required to produce the deviation from the CCAM line is ⬃20‰ in ␦18OSMOW, which is much larger than the effects (⬍5‰ in ␦18OSMOW) reported by Yurimoto et al. (1998). If we assume that the mass-dependent fractionation due to evaporation is solely responsible for the observed isotopic shifts of the 16O-rich spherules from the CCAM precursors, we can estimate the average evaporational mass fractionation of the spherules based on the seven 16O-rich spherules to be ⫹11.3‰ in ␦18OSMOW. If we apply this isotopic fractionation to the average of all 48 spherules, the average isotopic ratio of their precursors is estimated to be 11.8‰ in ␦18OSMOW and 3.9‰ in ␦17OSMOW. These values overlap those of the matrices of CM chondrites (Fig. 4c, Clayton and Mayeda, 1999), which is consistent with studies of unmelted micrometeorites (Kurat et al., 1994; Genge et al., 1997; Engrand et al., 1999; Nakamura et al., 2001; Noguchi et al., 2002). As is the case for the TFL spherules, Mg/Al and Si/Al ratios can help determine possible precursors for the 16O-rich spherules. Five of the 16O-rich spherules have Mg/Al and Si/Al ratios ⬎ 20 (Fig. 5b), suggesting that their precursors consist of olivines or low-Ca pyroxenes in C chondrites. In particular, KW540260 contains relict AOAs, as shown in Figure 1e. This is the first AOA found in MMs and IDPs. At present, AOAs are only found in C chondrites (Grossman and Steele, 1976; Hashimoto and Grossman, 1987; Komatsu et al., 2001; Imai and Yurimoto, 2003), therefore the precursor of this spherule may be closely related to C chondrites. Mg-rich olivines in AOAs have been reported to be as 16O-rich as CAIs, which is consistent with more negative ⌬17O values observed in the spherule than those of chondrules’ or isolated olivines in C chondrites (Leshin et al., 1997; Hiyagon and Hashimoto, 1999; Jones et al., 2000; Krot et al., 2002; Imai and Yurimoto, 2003). Imai and Yurimoto (2003) suggested that 16O-rich, Mg-rich olivines are pristine, whereas 16O-poor, Mg-poor olivines resulted from secondary alteration on the parent bodies of the C chondrites. This suggests that the precursor of the AOA-bearing spherule did not experience heavy aqueous and thermal alteration on the parent body. M240168 has a ␦18OSMOW value of ⫺16.7‰ with ⌬17O ⫽ ⫺20.5‰, and Mg/Al and Si/Al ratios ⬍ 5. Its refractory chemical composition and 16O-rich isotopic composition indicate that its precursor is a fragment of a CAI. This suggestion is consistent with the observation that CAI minerals in both unmelted AMMs and anhydrous IDPs have oxygen isotopic compositions close to those in C chondrites (McKeegan, 1987; Greshake et al., 1996; Engrand et al., 1999). The remaining 16O-rich spherule, M03KS071, has Mg/Al and Si/Al ratios close to the CI value. Its oxygen isotopic composition overlaps those of bulk CO chondrites (Clayton and Mayeda, 1999). This spherule may originate from the same parent body as that of CO chondrites.

4.4.

5801

16

O-Poor Group of Spherules: New Chondritic Materials?

Three spherules of this group plot clearly above the TFL and show 16O-poor isotopic compositions relative to the meteorites (Figs. 4a and 4c). M03KS063 has a ⌬17O value of 13.2‰, the highest value ever reported for solar system materials. The mixing of atmospheric oxygen with the spherules might possibly decrease the ⌬17O values of 16O-poor spherules, but here, for simplicity, we assume that the effect of mixing atmospheric oxygen into them is negligible. There is a possibility that this high ⌬17O value can be attributed to mixing with atmospheric CO2 at high altitude, which is characterized by a high ⌬17O (up to 12.0‰) (Fig. 4c, Thiemens et al., 1995). This explanation, however, seems unlikely, because CO2 is only a minor component (⬃350 ppm) in mesospheric air (Thiemens et al., 1995). The Mg/Al and Si/Al ratio of this spherule plot close to the CI value (Fig. 5b), suggesting that it is neither a mineral fragment, a CAI-related material, nor a severely evaporated remnant of CI-like material. R chondrites are known to have the largest ⌬17O value (⫹2.9‰) of all chondrite groups (Kallemeyn et al., 1996) and might be the precursor materials of the other two 16 O-poor spherules, M03KS021 and KW540299, because their ⌬17O values, 4.6 ⫾ 2.4‰ and 4.2 ⫾ 2.0‰, respectively, are comparable to those of the R chondrites within uncertainties. However, the difference in ␦18OSMOW between R chondrite (⬃5‰) and the two spherules (48.5‰–56.5‰) is too large to be explained by mass-dependent fractionation, because they are not largely depleted in FeO (12.7–22.5 wt.%). Therefore, the two 16O-poor spherules with lower ⌬17O values also cannot be explained simply by R chondrite precursors. A possible precursor for the 16O-poor spherules is an extremely 16O-poor chondritic material that is still unknown in our meteorite collections. The existence of 16O-rich solid materials and 16O-poor nebular gas in the early solar nebula is now widely accepted based on the heterogeneous distribution of oxygen isotopic compositions among meteoritic components (e.g., Clayton and Mayeda, 1984, 1999; Rowe et al., 1994; Clayton et al., 1991; Clayton, 1993). The assumed component of nebular gas has a 16O-poor composition of ␦18O ⫽ 30.0‰, ␦17O ⫽ 24.2‰, and ⌬17O ⫽ 8.6‰ in Clayton and Mayeda (1984). Using the model of Clayton and Mayeda (1984), Wiens et al. (1999) estimated that the nebular gas component rose up to ⫹33‰ in ␦18OSMOW and ⫹10‰ in ⌬17O, assuming 200% solid /gas enrichment. This ⌬17O value is high enough to account for the ⌬17O value of M03KS063, thus the precursor of M03KS063 could originate from a chondritic parent body that completely exchanged its oxygen with such a 16O-poor nebular gas. This possibility has the additional advantage that it requires less mass-dependent fractionation to achieve the ␦18OSMOW value of the 16O-poor spherules. If extremely 16O-poor materials existed in the early solar nebula, this would have important implications for the oxygen isotopic cosmochemistry of the solar system. According to Clayton (2003), there are three hypotheses to explain the oxygen isotope heterogeneity in the early solar system: two-component mixing (Clayton et al., 1973; Clayton and Mayeda, 1984), mass-independent chemical fractionation (Thiemens and Heidenreich, 1983; Thiemens, 1996), and isotope-selective photochemistry related to self-shielding of UV irradiation (Navon and Wasserburg, 1985; Clayton, 2002; Yu-

5802

T. Yada et al.

rimoto and Kuramoto, 2004; Lyons and Young, 2005). The assumed average solar oxygen isotopic ratio for the first hypothesis is between (␦18OSMOW, ␦17OSMOW) ⫽ (16.4, 11.4)‰ and (12.3, 7.5)‰, for the second one it is at the intersection of the TFL and the CCAM line, and for the third it is (␦18OSMOW, ␦17OSMOW) ⫽ (⬃⫺50, ⫺50)‰. Recently, Hashizume and Chaussidon (2005) analyzed oxygen isotopic compositions of solar wind component penetrated on the surface of the metallic grains in lunar regolith and found that it plotted below the TFL. Their result supports the third hypothesis. Even though all three models can explain the component that is 16O-poor relative to the TFL, the mechanisms in the latter two models are considered to be reasonable for explaining the 16O-poor planetary material, because 16O-poor products (O3, atomic O or H2O) relative to their starting materials will inevitably be generated from the solar oxygen isotopic ratio by the fractionation processes in these models. Actually, the 16O-poor spherule, M03KS063, plots close to the extension of the CCAM line with a slope of 1 in the ␦18OSMOW-␦17OSMOW space. Recently, an extremely 16O-rich glassy chondrule was discovered in Acfer 214, a CH chondrite, which also plots on the extension of the CCAM line at ␦18OSMOW ⫽ ⫺73.0‰ ⬃ ⫺75.8‰ and ␦17OSMOW ⫽ ⫺78.0‰ ⬃ ⫺73.9‰ (Kobayashi et al., 2003). Thus, there is evidence that both the lower and upper limits of oxygen isotopic variations in planetary materials have expanded along the CCAM line, which is the mass-independent fractionation line assumed in the latter two models. 5. CONCLUSIONS

We analyzed 48 S-type spherules collected from the Antarctic ice sheet for oxygen isotopic and bulk chemical compositions. On the basis of microscopic texture and chemical composition, they are classified into 6 types; barred olivine (BO), glass (Glass), cryptocrystalline (CC), porphyritic (Po), relict grain-bearing (RGB), and Ca, Al, Ti-rich (CAT). Their bulk chemistries indicate systematic depletions of FeO in CAT, Glass and CC spherules, reflecting the effect of vaporization during atmospheric entry. However, no correlation was observed between oxygen isotopic composition and texture type or bulk chemical composition. The ␦18OSMOW values range widely from ⫺28‰ to ⫹93‰. Mass-dependent fractionation due to evaporation and mixing with atmospheric oxygen might have changed their oxygen isotopic ratios from those of their precursor, but the large variations in oxygen isotopic compositions are believed to be an original feature of their precursors. Based on the ⌬17O values, the spherules are separated into 3 groups: those having isotopic signatures close to the TFL (38 spherules), those close to the CCAM line (7 spherules), and those with high ⌬17O values (3 spherules). By comparing the oxygen isotopic compositions with their Mg/Al and Si/Al ratios, we suggest possible precursors for the spherules; the majority of the analyzed spherules may be related to C chondrites. One spherule has the largest ⌬17O value (13‰) ever found in solar system materials. A possible precursor of this spherule is a planetary material that had exchanged all its oxygen with 16O-poor nebular gas. Acknowledgments—We are grateful to Mr. J. Mikada and Ms. A. Nakazawa for collecting some of the samples in this study and to Prof. Minoru Sekiya for discussion. We thank Dr. S. Tachibana for giving us

the artificial forsterite and Mr. T. Akaki for lending the thin sections of Allende meteorite. We give great thanks to Drs. Michael E. Zolensky, Susan Taylor, Cécile Engrand and an anonymous referee for constructive reviews for this paper. We are particularly grateful to Drs. Akira Yamaguchi and Naoya Imae for helping us with the operation of SEM-EDS and EPMA in National Institute of Polar Research. We wish to express special thanks to Drs. Christine Floss, Wallis F. Calaway and Michael R. Savina for improving this paper in both English and discussion. This study is partially supported by the Research Fellowships of the Japan Society for the Promotion of Science for Young Scientists, given to T.Y., by the research fund of NIPR, by the cooperative program of the Institute for Cosmic Ray Research, University of Tokyo, and by the Scientific Research Grant-in-Aid of the Ministry of Education, Culture, Sports, Science and Technology (No. 11440163 to H.K. and No.13740318 to T.N.). Associate editor: G. F. Herzog REFERENCES Akaki T. and Nakamura T. (2005) Formation processes of compound chondrules in CV3 carbonaceous chondrites: Constraints from oxygen isotopic ratios and major element concentrations. Geochim. Cosmochim. Acta 69, 2907–2929. Alexander C. M. O’D., Tayler S., Delaney J. S., Ma P., and Herzog G. F. (2002) Mass-dependent fractionation of Mg, Si and Fe isotopes in five stony cosmic spherules. Geochim. Cosmochim. Acta 66, 173–183. Baertschi P. (1976) Absolute 18O content of standard mean ocean water. Earth Planet. Sci. Lett. 31, 341–344. Blanchard M. B., Brownlee D. E., Bunch T. E., Hodge P. W., and Kyte F. T. (1980) Meteoroid ablation spheres from deep-sea sediments. Earth Planet. Sci. Lett. 46, 178 –190. Bland P. A., Smith T. B., Jull A. J. T., Berry F. J., Bevan A. W. R., Cloudt S., and Pillinger C. T. (1996) The flux of meteorites to the Earth over the last 50000 years. Mon. Not. R. Astron. Soc. 283, 551–565. Bradley J. P. (1994) Chemically anomalous, preaccretionally irradiated grains in interplanetary dust from comets. Science 265, 925–929. Bradley J. P., Brownlee D. E., and Veblen D. R. (1983) Pyroxene whiskers and platelets in interplanetary dust: Evidence of vapor phase growth. Nature 301, 473– 477. Brearley A. and Jones R. H. (1998) Chondritic meteorites. In Planetary Materials (ed. J. J. Papike), pp. 3-1–3-398. Mineralogical Society of America, Washington, DC. Brownlee D. E. (1985) Cosmic dust: Collection and research. Annu. Rev. Earth Planet. Sci. 13, 147–173. Brownlee D. E., Bates B. A., and Wheelock M. M. (1984) Extraterrestrial platinum group nugget in deep sea sediments. Nature 309, 693– 695. Brownlee D. E., Bates B., and Schramm L. (1997) The elemental composition of stony cosmic spherules. Meteorit. Planet. Sci. 32, 157–175. Burbine T. H., McCoy T. J., Meibom A., Gladman B. and Keil K. (2002) Meteoritic parent bodies: Their number and identification. In Asteroid III (eds. W. F. Bottke, A. Jr.Cellino, P. Paolicchi and R. P. Binzel), pp. 653– 667. University of Arizona Press, Tucson. Clayton R. N. (1993) Oxygen isotopes in meteorites. Annu. Rev. Earth Planet. Sci. 21, 115–149. Clayton R. N. (2002) Self-shielding in the solar nebula. Nature 415, 860 – 861. Clayton R. N. (2003) Oxygen isotopes in the solar system. Space Sci. Rev. 106, 19 –32. Clayton R. N. and Mayeda T. K. (1984) The oxygen isotope record in Murchison and other carbonaceous chondrites. Earth Planet. Sci. Lett. 67, 151–161. Clayton R. N. and Mayeda T. (1999) Oxygen isotopic studies of carbonaceous chondrites. Geochim. Cosmochim. Acta 63, 2089 – 2104. Clayton R. N., Grossman L., and Mayeda T. K. (1973) A component of primitive nuclear composition in carbonaceous meteorites. Science 182, 485– 488.

Oxygen isotope compositions of stony cosmic spherules Clayton R. N., Mayeda T. K., and Brownlee D. E. (1986) Oxygen isotopes in deep-sea spherules. Earth Planet. Sci. Lett. 79, 235–240. Clayton R. N., Mayeda T. K., Goswami J. N., and Olsen E. J. (1991) Oxygen isotope studies of ordinary chondrites. Geochim. Cosmochim. Acta 55, 2317–2337. Davis A. M., Clayton R. N., Mayeda T. K., and Brownlee D. E. (1991) Large mass fractionation of iron isotopes in cosmic spherules collected from deep-sea sediments. Lunar Planet. Sci. 22, 281–282. Davis A. M., McKeegan K. D., and MacPherson G. J. (2000) Oxygen isotopic compositions of individual minerals from the FUN inclusion Vigarano 1623–5. Meteorit. Planet. Sci. 35, A47. Duprat J., Maurette M., Engrand C., Matrajt G., Immel G., Hammer C., Gounelle M. and Kurat G. (2001) An estimation of the contemporary micrometeorite flux obtained from surface snow sample collected in central Antarctica. Meteorit. Planet. Sci. 36, Suppl., A52. Engrand C., McKeegan K. D., Leshin L. A., and Brownlee D. E. (1998) In-situ measurement of oxygen isotopic compositions of deep-sea and Antarctic cosmic spherules [abstract]. Lunar Planet. Sci. 29, #1473. Engrand C., McKeegan K. D., and Leshin L. A. (1999) Oxygen isotopic compositions of individual minerals in Antarctic micrometeorites: Further links to carbonaceous chondrites. Geochim. Cosmochim. Acta 63, 2623–2636. Floss C., Stadermann F. J., Bradley J. P., Dai Z. R., Bajt S., and Graham G. (2004) Carbon and nitrogen isotopic anomalies in an anhydrous interplanetary dust particle. Science 303, 1355–1358. Genge M. J., Grady M. M., and Hutchison R. (1997) The textures and compositions of fine-grained Antarctic micrometeorites: Implications for comparisons with meteorites. Geochim. Cosmochim. Acta 61, 5149 –5162. Glass B. P., Huber H., and Koeberl C. (2004) Geochemistry of Cenozoic microtektites and clinopyroxene-bearing spherules. Geochim. Cosmochim. Acta 68, 3971– 4005. Greshake A., Hoppe P., and Bischoff A. (1996) Mineralogy, chemistry and oxygen isotopes of refractory inclusions from interplanetary dust particles and micrometeorites. Meteorit. Planet. Sci. 31, 739 – 748. Grossman L. and Steele I. M. (1976) Amoeboid olivine aggregates in the Allende meteorite. Geochim. Cosmochim. Acta 40, 149 –155. Grün E., Zook H. A., Fechtig H., and Giese R. H. (1985) Collisional balance of the meteoritic complex. Icarus 62, 244 –272. Gustafson B. Å. S. (1994) Physics of zodiacal dust. Annu. Rev. Earth Planet. Sci. 22, 553–595. Halliday I., Blackwell A. T., and Griffin A. A. (1989) The flux of meteorites on the Earth’s surface. Meteoritics 24, 173–178. Hashimoto A. and Grossman L. (1987) Alteration of Al-rich inclusions inside amoeboid olivine aggregates in the Allende meteorite. Geochim. Cosmochim. Acta 51, 1685–1704. Hashizume K. and Chaussidon M. (2005) A non-terrestrial 16O-rich isotopic composition for the protosolar nebula. Nature 434, 619 – 622. Hiroi T., Zolensky M. E., and Pieters C. M. (2001) The Tagish Lake meteorite: A possible sample from a D-type asteroid. Science 293, 2234 –2236. Hiyagon H. and Hashimoto A. (1999) 16O excesses in olivine inclusions in Yamato-86009 and Murchison chondrites and their relation to CAIs. Science 283, 828 – 831. Imai H. and Yurimoto H. (2003) Oxygen isotopic distribution in an amoeboid olivine aggregate from the Allende CV chondrite: Primary and secondary processes. Geochim. Cosmochim. Acta 67, 765–772. Ishimoto H. (2000) An analytical model for the mass distribution of the interplanetary dust cloud. Adv. Space Res. 25, 281–284. Jones R. H., Saxton J. M., Lyon I. C., and Turner G. (2000) Oxygen isotopes in chondrule olivine and isolated olivine grains from the CO3 chondrite Allan Hills A77307. Meteorit. Planet. Sci. 35, 849 – 857. Kallemeyn G. W., Rubin A. E., and Wasson J. T. (1996) The compositional classification of chondrites: The R chondrite group. Geochim. Cosmochim. Acta 60, 2243–2256. Katsushima T., Nishio F., Ohmae H., Ichikawa M., and Takahashi S. (1984) Composition of dirt layers in the bare ice areas near the Yamato Mountains in Queen Maud Land and the Allan Hills in

5803

Victoria Land, Antarctica. Mem. Natl Inst. Polar Res. Sec. Issue 34, 174 –187. Kobayashi S., Imai H., and Yurimoto H. (2003) New extreme 16O-rich reservoir in the early solar system. Geochem. J. 37, 663– 669. Komatsu M., Krot A. N., Petaev M. I., Ulyanov A. A., Keil K., and Miyamoto M. (2001) Mineralogy and petrography of amoeboid olivine aggregates from the reduced CV3 chondrites Efremovka, Leoville and Vigarano: Products of nebular condensation, accretion and annealing. Meteorit. Planet. Sci. 36, 629 – 641. Kortenkamp S. J. and Dermott S. F. (1998) Accretion of interplanetary dust particles by the Earth. Icarus 135, 469 – 495. Krot A. N., McKeegan K. D., Leshin L. A., MacPherson G. J., and Scott E. R. D. (2002) Existence of an 16O-rich gaseous reservoir in the solar nebula. Science 295, 1051–1054. Kurat G., Koeberl C., Presper T., Brandstätter F., and Maurette M. (1994) Petrology and geochemistry of Antarctic micrometeorites. Geochim. Cosmochim. Acta 58, 3879 –3901. Kusakabe M., Maruyama S., Nakamura T., and Yada T. (2004) CO2 Laser-BrF5 fluorination technique for analysis of oxygen three isotopes of rocks and minerals. J. Mass Spectrom. Soc. Jpn. 52, 205–212. Landgraf M., Liou J.-C., Zook H. A., and Grün E. (2002) Origin of solar system dust beyond Jupiter. Astron. J. 123, 2857–2861. Larimer J. W. and Wasson J. T. (1988) Refractory lithophile elements. In Meteorites and the Early Solar System (eds. J. F. Kerridge and M. S. Matthews), pp. 394 – 415. University of Arizona Press, Tucson. Lee T. and Papanastassiou D. A. (1974) Mg isotopic anomalies and correlation with O and Si effects. Geophys. Res. Lett. 1, 335–228. Leshin L. A., Rubin A. E., and McKeegan K. D. (1997) The oxygen isotopic composition of olivine and pyroxene from CI chondrites. Geochim. Cosmochim. Acta 61, 835– 845. Liou J.-C., Zook H. A., and Dermott S. F. (1996) Kuiper belt dust grains as a source of interplanetary dust particles. Icarus 124, 429 – 440. Lodders K. and Fegley B., Jr. (1998) The Planetary Scientist’s Companion. Oxford University Press, New York. Love S. G. and Brownlee D. E. (1991) Heating and thermal transformation of micrometeoroids entering the Earth’s atmosphere. Icarus 89, 26 – 43. Love S. G. and Brownlee D. E. (1993) A direct measurement of the terrestrial mass accretion rate of cosmic dust. Science 262, 550 – 553. Lyons J. R. and Young E. D. (2005) CO-self shielding as the origin of oxygen isotopic anomalies in the early solar nebula. Nature 430, 317–320. Maurette M., Olinger C., Michel-Levy M. C., Kurat G., Pourchet M., Brandstätter F., and Bourot-Denise M. (1991) A collection of diverse micrometeorites recovered from 100 tonnes of Antarctic blue ice. Nature 351, 44 – 46. McKeegan K. D. (1987) Oxygen isotopes in refractory stratospheric dust particles: Proof of extraterrestrial origin. Science 237, 1468 – 1471. Messenger S. (2000) Identification of molecular-cloud material in interplanetary dust particles. Nature 404, 968 –971. Millard H. T. and Finkelman R. B. (1970) Chemical and mineralogical compositions of cosmic and terrestrial spherules from marine sediment. J. Geophys. Res. 75, 2125–2133. Nakamura T., Imae N., Nakai I., Noguchi T., Yano H., Terada K., Murakami T., Fukuoka T., Nogami K., Ohashi H., Nozaki W., Hashimoto M., Kondo N., Matsuzaki H., Ichikawa O., and Ohmori R. (1999) Antarctic micrometeorites collected at the Dome Fuji Station. Antarct. Meteorite Res. 12, 183–198. Nakamura T., Noguchi T., Yada T., Nakamuta Y., and Takaoka N. (2001) Bulk mineralogy of individual micrometeorites determined by X-ray diffraction analysis and transmission electron microscopy. Geochim. Cosmochim. Acta 65, 4385– 4397. Navon O. and Wasserburg G. J. (1985) Self-shielding in O2—A possible explanation for oxygen isotopic anomalies in meteorites? Earth Planet. Sci. Lett. 73, 1–16. Nishiizumi K. (1983) Measurement of 53Mn in deep-sea iron and stony spherules. Earth Planet. Sci. Lett. 63, 223–228.

5804

T. Yada et al.

Nishiizumi K., Arnold J. R., Fink D., Klein J., Middleton R., Brownlee D. E., and Maurette M. (1991) Exposure history of individual cosmic spherules. Earth Planet. Sci. Lett. 104, 315–324. Nishiizumi K., Arnold J. R., Brownlee D. E., Caffee M. W., Finkel R. C., and Harvey R. P. (1995) Beryllium-10 and aluminum-26 in individual cosmic spherules from Antarctica. Meteoritics 30, 728 – 732. Noguchi T., Nakamura T., and Nozaki W. (2002) Mineralogy of phyllosilicate-rich micrometeorites and comparison with Tagish Lake and Sayama meteorites. Earth Planet. Sci. Lett. 202, 229 – 246. Noguchi T., Imae N., Nakamura T., Nozaki W., Terada K., Mori T., Nakai I., Kondo N., Sasaki M., Murakami T., Fukuoka T., Nogami K., Ohmori R., and Ohashi H. (2000) A consortium study of Antarctic micrometeorites recovered from the Dome Fuji station. Antarct. Meteorite Res. 13, 270 –284. Osawa T., Nagao K., Noguchi T., Nakazawa A., and Mikada J. (2003) Remnant extraterrestrial noble gases in Antarctic cosmic spherules. Antarct. Meteorit. Res. 16, 196 –219. Raisbeck G. M. and Yiou F. (1987) 10Be and 26Al in individual cosmic spherules from Antarctica [abstract]. Meteoritics 24, 31819. Raisbeck G. M., Yiou F., Klein J., Middleton R. and Brownlee D. E. (1985) 26Al/10Be in deep sea spherules as evidence of cometary origin. In Properties and Interactions of Interplanetary Dust (eds. R. H. Giese and P. Lamy), pp. 169 –174. D. Reidel, Dordrecht, the Netherlands. Rowe M. W., Clayton R. N., and Mayeda T. K. (1994) Oxygen isotopes in separated components of CI and CM meteorites. Geochim. Cosmochim. Acta 58, 5341–5347. Slodzian G. (1987) An electron mirror configuration for regulating the potential on insulators in ion emission microscopy. Optik-Stuttgart 77, 148 –155. Taylor S., Lever J. H., and Harvey R. P. (1998) Accretion rate of cosmic spherules measured at the South Pole. Nature 392, 899 – 903. Taylor S., Lever J. H., and Harvey R. P. (2000) Numbers, types and compositions of an unbiased collection of cosmic spherules. Meteorit. Planet. Sci. 35, 651– 666. Taylor S., Alexander C. M. O’D., Delaney J., Ma P., Herzog G. F., and Engrand C. (2005) Isotopic fractionation of iron, potassium and oxygen in stony cosmic spherules: Implications for heating histories and sources. Geochim. Cosmochim. Acta 69, 2647–2662. Terada K., Yada T., Kojima H., Noguchi T., Nakamura T., Murakami T., Yano H., Nozaki W., Nakamuta Y., Matsumoto N., Kamata J., Mori T., Nakai I., Sasaki M., Itabashi M., Setoyanagi T., Nagao K., Osawa T., Hiyagon H., Mizutani S., Fukuoka T., Nogami K., Ohmori R., and Ohashi H. (2001) General characterization of Antarctic micrometeorites collected by the 39th Japanese Antarctic Research Expedition: Consortium studies of JARE AMMs (III). Antarct. Meteorit. Res. 14, 89 –107.

Thiemens M. H. (1996) Mass-independent isotopic effects in chondrites: The role of chemical processes. In Chondrules and the Protoplanetary disk (eds. R. H. Hewins, R. H. Jones and E. R. D. Scott), pp. 107–118. Cambridge University Press, Cambridge, UK. Thiemens M. H. and Heidenreich J. E., III (1983) The mass-independent fractionation of oxygen: A novel isotope effect and its possible cosmochemical implications. Science 219, 1073–1075. Thiemens M. H., Jackson T., Zipf E. C., Erdman P. W., and van Egmond C. (1995) Carbon dioxide and oxygen isotope anomalies in the mesosphere and stratosphere. Science 270, 969 –972. Thomas K. L., Blanford G. E., Keller L. P., Klöck W., and McKay D. S. (1993) Carbon abundance and silicate mineralogy of anhydrous interplanetary dust particles. Geochim. Cosmochim. Acta 57, 1551–1566. Tomeoka K. and Buseck P. R. (1985) Hydrated interplanetary dust particle linked with carbonaceous chondrites? Nature 314, 338 – 340. Wang J., Davis A. M., Clayton R. N., Mayeda T. K., and Hashimoto A. (2001) Chemical and isotopic fractionation during the evaporation of the FeO - MgO - SiO2 - CaO - Al2O3 - TiO2 rare earth element melt system. Geochim. Cosmochim. Acta 65, 479 – 494. Wiens R. C., Huss G. R., and Burnett D. S. (1999) The solar oxygenisotopic composition: Predictions and implications for solar nebular processes. Meteorit. Planet. Sci. 34, 99 –107. Yada T. and Kojima H. (2000) The collection of micrometeorites collected in the Yamato Meteorite Ice Field of Antarctica in 1998. Antarct. Meteorit. Res. 13, 9 –18. Yada T., Nakamura T., Takaoka N., Noguchi T., Terada K., Yano H., Nakazawa T., and Kojima H. (2004) The global accretion rate of extraterrestrial materials in the last glacial period estimated from the abundance of micrometeorites in Antarctic glacier ice. Earth Planets Space 56, 823– 835. Yamakoshi K. (1984) Chemical compositions of magnetic, stony spherules from deep sea sediments determined by instrumental neutron activation analysis. Geochem. J. 18, 147–152. Yu Y., Hewins R. H., Clayton R. N., and Mayeda T. K. (1995) Experimental study of high temperature oxygen isotope exchange during chondrule formation. Geochim. Cosmochim. Acta 59, 2095– 2104. Yurimoto H., Ito M., and Nagasawa H. (1998) Oxygen isotope exchange between refractory inclusion in Allende and solar nebula gas. Science 282, 1874 –1877. Yurimoto H. and Kuramoto K. (2004) Molecular cloud origin for the oxygen isotope heterogeneity in the solar system. Science 305, 1763–1766. Zolensky M. E., Barrett R., and Browning L. (1993) Mineralogy and composition of matrix and chondrule rims in carbonaceous chondrites. Geochim. Cosmochim. Acta 57, 3123–3148.