Tellurium isotopic composition of the early solar system—A search for effects resulting from stellar nucleosynthesis, 126Sn decay, and mass-independent fractionation

Geochimica et Cosmochimica Acta, Vol. 69, No. 21, pp. 5099 –5112, 2005 Copyright © 2005 Elsevier Ltd Printed in the USA. All rights reserved 0016-7037/05 $30.00 ⫹ .00

doi:10.1016/j.gca.2005.04.020

Tellurium isotopic composition of the early solar system—A search for effects resulting from stellar nucleosynthesis, 126Sn decay, and mass-independent fractionation MANUELA A. FEHR,1,2,* MARK REHKÄMPER,1,3 ALEX N. HALLIDAY,1,4 UWE WIECHERT,1 BODO HATTENDORF,5 DETLEF GÜNTHER,5 SHUHEI ONO,6 JENNIFER L. EIGENBRODE,6,7 and DOUGLAS RUMBLE, III6 1

ETH Zürich, Institute of Isotope Geology and Mineral Resources, 8092 Zürich, Switzerland Swedish Museum for Natural History, Laboratory for Isotope Geology, 104 05 Stockholm, Sweden 3 Department of Earth Science and Engineering, Imperial College, London SW7 2AZ, United Kingdom 4 Department of Earth Sciences, Oxford OX1 3PR, United Kingdom 5 ETH Zürich, Laboratory of Inorganic Chemistry, 8093 Zürich, Switzerland 6 Carnegie Institution of Washington, Geophysical Laboratory, Washington, DC 20015, USA 7 The Pennsylvania State University, Astrobiology Research Center, University Park, PA 16802, USA 2

(Received August 16, 2004; accepted in revised form April 11, 2005)

Abstract—New precise Te isotope data acquired by multiple collector inductively coupled plasma mass spectrometry (MC-ICPMS) are presented for selected extraterrestrial and terrestrial materials. Bulk samples of carbonaceous, ordinary and enstatite chondrites as well as the metal and sulfide phases of iron meteorites were analyzed to search for nucleosynthetic isotope anomalies and to find evidence of formerly live 126Sn, which decays to 126Te with a half-life of 234,500 yr. None of the meteorites show evidence of mass dependent Te isotope fractionations larger than 2‰ for ␦126/128Te. Following internal normalization of the data to 125 Te/128Te, the Te isotope ratios of all analyzed meteorites were found to be identical to a terrestrial standard, within uncertainties. This provides evidence that the regions of the solar disk that were sampled during accretion of the meteorite parent bodies were well mixed and homogeneous on a large scale, with respect to Te isotopes. The data acquired for bulk carbonaceous chondrites indicate that the initial 126Sn/118Sn ratio of the solar system was ⬍4 ⫻ 10⫺5, but this is dependent on the assumption that no redistribution of Sn and Te occurred since the start of the solar system. Five Archean sedimentary sulfides that display both mass dependent and mass-independent isotope effects for S yield internally normalized Te isotope data, which indicate that mass-independent Te isotope effects are absent. The mass dependent fractionations in these samples are constrained to be less than ⬃1‰ for ␦126/128Te. Copyright © 2005 Elsevier Ltd gations of the 126Sn-126Te decay system are important in this regard because 126Sn has a very short half-life and it cannot be formed by spallation. The only event that would be expected to produce sufficient amounts of 126Sn to generate resolvable anomalies of its daughter 126Te in at least some meteorite parent bodies, is a supernova explosion that occurred just before the formation of the solar system. The discovery of 126Te excesses that correlate with Sn/Te ratios in meteorites would therefore provide powerful confirmation of the theory that a supernova injected freshly synthesized nuclides into the molecular cloud from which our solar system formed, providing evidence of a trigger (Cameron and Truran, 1977). (2) Tellurium isotopes offer a broad spectrum of nucleosynthetic sites to test models for the stars that contributed nuclides to the molecular cloud from which the Sun and planets were formed (Smith and De Laeter, 1986). A number of studies have concluded that the solar nebula was, in general, well mixed (e.g., Suess, 1965; Reynolds, 1967) and it is now known that the isotope compositions of many elements are identical to within ⬃0.01% or less in bulk samples of the Earth, Moon, Mars (as sampled by SNC meteorites) and various meteorite parent bodies (e.g., Reynolds, 1967; Schönbächler et al., 2003). Large-scale isotopic variations, however, have been proposed for a number of elements, including O (Clayton, 2004), Mn or Cr (Lugmair and Shukolyukov, 1998), Mo (Dauphas et al., 2002), and Ru (Chen et al., 2003; Papanastassiou et al., 2004). On a smaller scale, nucleosynthetic anomalies of

1. INTRODUCTION

The tellurium (Te) isotope compositions of early solar system objects have been difficult to determine precisely but are of great current interest for the following reasons: (1) The short-lived nuclide 126Sn decays to 126Te through the intermediate 126Sb with a half-life of 234,500 ⫾ 7100 yr (Oberli et al., 1999). As 126Sn cannot be produced in significant amounts by the s-process because of the exceedingly short half-life of 125Sn (9.6 d), it is predominantly an r-process nuclide that is probably formed in supernova environments (Qian et al., 1998). The initial solar system abundances of most short-lived radionuclides with halflives of ⬃10 to 100 Myr (e.g., 129I, 146Sm, 244Pu) are commonly assumed to reflect steady state concentrations that were established through continuous galactic nucleosynthesis and free decay in the interstellar medium (Wasserburg et al., 1996; Meyer and Clayton, 2000). The discovery of radiogenic effects from the decay of short-lived nuclides with half-lives of less than 1 Myr (e.g., 26Al, 41Ca) in meteorites, however, requires that these isotopes were produced either within the nascent solar system by spallation or in a late stellar nucleosynthetic event that took place just before the collapse of the protosolar cloud (Lee et al., 1998; Meyer and Clayton, 2000). Cosmochemical investi* Author to whom correspondence should be addressed, at Laboratory for Isotope Geology, Swedish Museum for Natural History, 50007 Box, 104 05 Stockholm, Sweden ([email protected]). 5099

␦33S, ␦34S, and ⌬33S data are from Ono et al. (2003). d

Number of individual Te isotope measurements for mean values.

From Fehr et al. (2004). c

b

a

6 SV1 55.2 SV1 55.2 SV1 110.5 SV1 110.5 RHDH2a 161.7d RHDH2a 161.7d RHDH2a 161.7d RHDH2a 205.2d RHDH2a 205.2d RHDH2a 299.2d RHDH2a 299.2d

6 20 Te standard Pyrrhotitec Pyritec

␦33S No.b Sample

Five Archean sedimentary pyrites as well as two more recent terrestrial sulfides, a pyrite and a pyrrhotite, were analyzed in the present study (Table 1). For the latter samples, which were previously analyzed by Fehr et al. (2004), only the mean results of several individual analyses are reported. The Precambrian sulfides are from the Fortescue Group in the Hamersly Basin, Western Australia, and they have ages of ⬃2.7 Ga (Arndt et al., 1991). The S isotope compositions of these samples were previously determined by Ono et al. (2003), who detected significant variations in both ␦34S and ⌬33S (Table 1). The latter parameter defines the deviation of the S isotope composition from a purely mass dependent relationship (Thiemens et al., 2001). These variations were interpreted to reflect the combined effects of MIF in the atmosphere and biologically mediated mass dependent fractionations in the oceans (Farquhar et al., 2000).

1.21 0.86 0.97 0.57 0.12 0.37 ⫺0.81 6.13 7.26 8.27

2.1. Terrestrial Sulfides

The quoted analytical uncertainties reflect the external reproducibilities of the standard obtained on a particular measurement session. For the mean values, the uncertainties reflect the combined error of the individual measurements. Multiple digestions were performed only for the pyrrhotite (5) and pyrite (3). ␦xS ⫽ [(xS/32S)sample/(xS/32S)CDT ⫺ 1] ⫻ 103, where x is either 33 or 34. ⌬33S ⫽ 103 ⫻ ln(1 ⫹ ␦33S/103) ⫺0.515 ⫻ 103 ln(1 ⫹ ␦34S/103). The precision (2s) for sulfur isotope data is ⫾ 0.2‰ for ␦33S, ⫾ 0.4‰ for ␦34S, and better than 0.1‰ for ⌬33S. ␧xTeN ⫽ {[(xTe/128Te)N sample ⫺ (xTe/128Te)N std]/(xTe/128Te)N std} ⫻ 104, where x is 120, 122, 124, 126, or 130; N: normalized to 125Te/128Te ⫽ 0.22204 (Lee and Halliday, 1995) with the exponential law. ␦126/128TeM ⫽ {[(126Te/128Te)M sample ⫺ (126Te/128Te)M std mean]/(126Te/128Te)M std mean} ⫻ 103; M: measured.

0.7 ⫾ 0.9 ⫺1.0 ⫾ 0.9 ⫺0.2 ⫾ 0.7 ⫺0.1 ⫾ 0.5 0.1 ⫾ 0.4 0.6 ⫾ 0.6 ⫺1.0 ⫾ 2.1 2.1 ⫾ 2.2 ⫺43 ⫾ 59 71 ⫾ 62

⫺0.8 ⫾ 1.6 ⫺0.1 ⫾ 1.8

1.2 ⫾ 1.4 5.6 ⫾ 29.1 13.5 ⫾ 22.3 ⫺50.6 ⫾ 132.1 ⫺2246 ⫾ 2754

⫺8.4 ⫾ 59.4

⫺0.9 ⫾ 0.9 ⫺0.2 ⫾ 0.9 0.5 ⫾ 0.9

⫺0.5 ⫾ 0.4 0.3 ⫾ 0.2

0.1 ⫾ 0.6 1.1 ⫾ 0.7 6.8 ⫾ 9.3 ⫺0.2 ⫾ 0.3 0.0 ⫾ 0.8 ⫺8.2 ⫾ 9.5

⫺0.1 ⫾ 0.1

0.3 ⫾ 1.4 ⫺0.8 ⫾ 2.2 ⫺50.2 ⫾ 19.2

0.0 ⫾ 0.3 0.0 ⫾ 0.6 0 ⫾ 18

15 ⫾ 44 ⫺48 ⫾ 64 ⫺940 ⫾ 1683

0.77 0.75 0.68 0.57 ⫺2.49 ⫺2.44 ⫺2.45 ⫺1.97 ⫺1.05 6.20

Precambrian sulfides

0.1 ⫾ 1.0 ⫺0.9 ⫾ 2.6 ⫺13.3 ⫾ 33.1

0.0 ⫾ 0.9 0.3 ⫾ 0.4 0.2 ⫾ 0.2 0.0 ⫾ 0.6 0.2 ⫾ 0.2 0.5 ⫾ 0.1 0.0 ⫾ 0.3 0.2 ⫾ 0.1 0.0 ⫾ 0.1 0.0 ⫾ 1.4 ⫺0.8 ⫾ 0.5 ⫺0.2 ⫾ 0.3 0 ⫾ 45 ⫺31 ⫾ 17 ⫺14 ⫾ 10

0.0 ⫾ 1.0 ⫺0.3 ⫾ 0.4 ⫺0.2 ⫾ 0.2

␧130TeN ⫾ 2␴ ␧126TeN ⫾ 2␴ ␧124TeN ⫾ 2␴ ␧122TeN ⫾ 2␴ ␧120TeN ⫾ 2␴

0.85 0.22 0.57 ⫺0.01 5.08 5.48 3.20 15.80 16.21 3.96

2. SAMPLES

⌬33S

This study attempts to address these four issues using multiple collector inductively coupled plasma mass spectrometry (MC-ICPMS). In the past, Te isotopic compositions have been determined by neutron activation analysis (Ballad et al., 1979; Oliver et al., 1981), positive ion thermal ionization mass spectrometry (Smith et al., 1978; De Laeter and Rosman, 1984; Smith and De Laeter, 1986), negative ion thermal ionization mass spectrometry (Wachsmann and Heumann, 1992; Richter et al., 1998), and MC-ICPMS (Lee and Halliday, 1995; Fehr et al., 2004). Of these, MC-ICPMS provides the best precision combined with excellent sensitivity. In the present study, the MC-ICPMS techniques of Fehr et al. (2004) were used to determine the Te isotope compositions of Archean sulfides, bulk samples of several carbonaceous, ordinary and enstatite chondrites as well as the metal and sulfide phases of iron meteorites.

␦34S

various elements are known in calcium-aluminum-rich inclusions (CAIs) (MacPherson, 2004) and presolar grains (Nittler, 2003; Zinner, 2004). Tellurium is particularly useful for studies of stellar nucleosynthesis because it has eight stable nuclides with a range of production mechanisms. 120Te is produced only by the p-process, 122–124Te only by the s-process, 128,130Te only by the r-process, whereas 125,126Te are produced by both the r- and sprocess. In addition, it has already been demonstrated that there exist large (‰–% level) nucleosynthetic Te isotope anomalies in presolar diamonds from Allende (Richter et al., 1998; Maas et al., 2001). (3) Cadmium isotope fractionations of almost 4‰ per amu mass difference have been reported for some chondritic meteorites that are though to reflect redistribution of this volatile element due to thermal metamorphism (Rosman and De Laeter, 1988; Wombacher et al., 2003). Tellurium is only slightly less volatile than Cd, such that it may also exhibit large stable isotope effects. (4) Some Archean sulfides have been found to display large mass-independent fractionations (MIF) for S isotopes (Ono et al., 2003). These isotope effects are thought to be due to photochemical reactions of gaseous S-species that were delivered to the atmosphere by volcanic activity (Farquhar et al., 2000). The chalcophile element Te has a (geo)chemical behavior akin to S, such that atmospheric reactions may also have generated MIF for Te isotopes.

␦126/128TeM ⫾ 2␴

M. A. Fehr et al.

Table 1. Te isotopic data of terrestrial sulfides.a

5100

Tellurium isotopic composition of the early solar system

2.2. Meteorites Tellurium isotope data were acquired for bulk samples of 7 carbonaceous chondrites (including Orgueil [CI], Murchison [CM], and Allende [CV]), three ordinary chondrites (Plainview, Mezö-Madaras, and Allan Hills (ALH) 84081), and the Abee enstatite chondrite (Table 2). Also analyzed were the metal phases of the iron meteorites North Chile (Coya Norte specimen) and Arispe, as well as metal and sulfide from Canyon Diablo and Toluca (Table 2). The detailed replicated results obtained for Allende and Canyon Diablo (metal and sulfide) were already published in Fehr et al. (2004), and only the averages from several repeat analyses are given here. 3. ANALYTICAL PROCEDURES The chemical and mass spectrometric techniques follow the methods described in Fehr et al. (2004). The most important aspects of the procedures and any modifications are outlined below. 3.1. Sample Preparation Chondrite samples weighing more than 1 g were crushed in a boron carbide or aluminum oxide mortar under a laminar flow of filtered air and a fraction of the powder, generally ⬃0.1–1 g (Table 3), was then dissolved without further treatment. For the iron meteorites, ⬃5–15 g and 1 g, respectively, of metal and sulfide were rinsed with ethanol and then leached twice with 50% aqua regia at room temperature. Following this, the samples were digested with aqua regia on a hot plate overnight. For the Archean sulfides, sample aliquots of ⬃50 mg were dissolved directly in aqua regia without any pretreatment. Tellurium was chemically isolated from the rock matrix using a twostage column chemistry procedure. For the iron meteorites, an additional solvent extraction step was used for the separation of Te from Fe before the ion-exchange chemistry. Out of the 40 blanks measured, 33 had between 2 and 30 pg Te, and only 4 displayed more than 100 pg Te. The occurrence of such high values was erratic and later investigations of this effect suggest that they were due to incomplete removal of Te from either the Teflon beakers that were used for sample handling or the ion-exchange columns. The typical blank contribution is 0.03%, which is calculated for a sample with 100 ng Te and a typical total chemistry blank of 30 pg. The maximum measured total Te chemistry blank was 400 pg, which is still negligible as most samples contained more than 30 ng Te (therefore ⬍1.3% blank contribution). Two of the Archean sulfides, however, contained only 1.5 ng (RHDH2a 205.2) and 3 ng (RHDH2a 161.7) of Te. For these samples, the blank Te contribution could conceivably be as high as 15 to 30% (assuming an extreme blank of 400 pg Te) but it is more likely ⬃1 to 2%, based on a typical blank of 30 pg Te. 3.2. Mass Spectrometry The Te isotope measurements were performed with a Nu Plasma MC-ICPMS at the ETH Zürich on Faraday cups. Sample solutions (normally 100 ng/g Te) were introduced in 0.1 M HNO3 by free aspiration using a Cetac MCN 6000 desolvating nebulizer. Before each analysis, the system was washed with 0.1 M HNO3 for 5–10 min and this was typically sufficient to achieve residual Te signal levels equivalent to less than 0.1% of those utilized during sample and standard analyses. Longer washout periods of 30 to 60 min were used only occasionally at the end of longer (⬎12 h) measurement sessions to reduce memory effects that had gradually built up in the preceding hours. Most measurements comprised the collection of 80 ratios (5 s integrations), but for samples with low Te concentrations only 20 – 60 ratios were collected. The measured Te isotope ratios were normalized (N) to 125Te/128Te ⫽ 0.22204 (Lee and Halliday, 1995) with the exponential law to correct for mass fractionation. Epsilon Te values for samples were calculated as the difference to the mean of the matching JMC Te standards measured on the same day using ␧1xxTeN ⫽ {[(1xxTe/128Te)N sample ⫺ (1xxTe/128Te)N std]/(1xxTe/128Te)N std} ⫻ 104. The reproducibility (2s) of the isotopic measurements for 100 –150 ng/g standard solutions of Te on 1 d is typically ⫾ 45 ␧ for 120Te/ 128 TeN, ⫾ 1.4 ␧ for 122Te/128TeN, ⫾ 1.0 ␧ for 124Te/128TeN, ⫾ 0.3 ␧ for

5101

Te/128TeN, and ⫾ 0.6 ␧ for 130Te/128TeN. As 17 analyses of the Allende chondrite have a reproducibility similar to that obtained for repeated analyses of the standard (Fehr et al., 2004), we use the latter data as a conservative estimate for the 2s uncertainty of sample measurements. For the calculation of mean values, the variances from duplicates were combined quadratically assuming a single population.

126

3.2.1. Tellurium stable isotope measurements Although the methods used in this study (Fehr et al., 2004) were not specifically developed or optimized to yield precise Te stable isotope data, some information about possible mass dependent fractionations can nonetheless be deduced. Consistent results were furthermore obtained for several Te isotope ratios. The data are reported as ␦126/128TeM values (M: measured), which were calculated using the standard-sample bracketing technique, where sample data are referenced to the mean result of the preceding and succeeding JMC Te measurements: ␦126/128TeM ⫽ {[(126Te/128Te)M sample ⫺ (126Te/128Te)M std mean]/(126Te/128Te)M std mean} ⫻ 103. The external reproducibility of the ␦-values obtained for the JMC standard varied significantly between individual measurement sessions and 2s uncertainties of ⫾ 0.2–1.7‰ for ␦126/128TeM were obtained. This variability most likely reflects variations in the stability of the instrumental mass bias due to either sample-induced matrix effects and/or small drifts in the operating conditions (e.g., gas flows, sample uptake rates). For most sessions, the reproducibility of the standard was ⫾ 0.9‰ or better and we use this value as a conservative estimate of the 2␴ uncertainty of the ␦126/128TeM data. Total chemistry yields were 70%– 80% for sulfides and chondrites, and only ⬃55% for the metal of iron meteorites (Fehr et al., 2004). Several experiments showed, however, that the incomplete yields did not generate apparent isotopic effects. A Te standard solution that was processed through the column chemistry displayed a ␦126/128TeM value of ⫹0.1 ⫾ 0.9‰. Several synthetic samples that were doped with Te (originally to demonstrate the accuracy of our ␧1xxTeN results; Fehr et al., 2004) furthermore revealed no Te stable isotope fractionations. This included a terrestrial diorite doped with various amounts of Te, a chondrite matrix (obtained by the collection of all Te-free matrix fractions that were eluted whilst solutions of Murchison and Allende were processed through the column) and an iron meteorite matrix (Fe, Ni and Te standard). These samples displayed ␦126/128TeM values of ⫹0.1 ⫾ 0.3‰, ⫺0.2 ⫾ 0.4‰, and ⫹0.2 ⫾ 0.4‰, respectively. This demonstrates that the separation procedure does not generate stable isotope fractionations, despite the low chemical yields. 3.2.2. Determination of Sn and Te concentrations Tin and Te concentrations were measured by quadrupole ICP-MS (Agilent 7500cs) at the ETH Zürich on sample solution aliquots that did not undergo chemical separation, to avoid any chemical fractionation of Sn from Te. The concentrations were determined after external calibration, using synthetic standard solutions. Samples and standards were prepared in 0.5 M HNO3 and 0.01 M HF to stabilize Sn and Te and prevent precipitation of SnO2. Rhodium and Cd were used as internal standards to compensate for nonspectral interferences. Each digest was diluted to yield a concentration below 1 g/L for the original sample. The analyses typically displayed a precision of better than 10% relative standard deviation (RSD, 2s) for concentrations above 0.5 ␮g/L Sn and Te and the limits of detection (LODs) were 0.01 ␮g/L for 118Sn and 0.025 ␮g/L for 125Te. This corresponds to abundances of ⬃0.01 ␮g/g Sn and 0.03 ␮g/g Te in the original samples, depending on the final dilution. The procedural blanks were below the detection limit for both elements. For the 118Sn/128Te concentration ratios, the precision was better than 10% RSD for Sn and Te concentrations of 10 times above the LOD and up to 30% RSD for concentrations close to the LOD. 4. RESULTS

4.1. Tin and Tellurium Concentrations The Sn and Te concentrations for chondrites, iron meteorites and terrestrial sulfides are shown in Table 3. For some samples, such as the troilite inclusions of Canyon Diablo and Toluca,

5102

M. A. Fehr et al. Table 2. Te isotopic data of meteorite samples.a

Sample

No.b

␧120TeN ⫾ 2␴ 0 ⫾ 45

Te standard

␧122TeN ⫾ 2␴

␧124TeN ⫾ 2␴

0.0 ⫾ 1.4

0.0 ⫾ 1.0

␧126TeN ⫾ 2␴ 0.0 ⫾ 0.3

␧130TeN ⫾ 2␴ 0.0 ⫾ 0.6

␦126/128TeM ⫾ 2␴ 0.0 ⫾ 0.9

Carbonaceous chondrites Mean Orgueil (CI1) ALH 83100 (CM1/2) ALH 83100 (CM1/2) ALH 83100 (CM1/2) Mean Murchison (CM2) (USNM5459) ALH 84028 (CV3) ALH 84028 (CV3) Mean Allende (CV3) (USNM-6159)c Allende (CV3) (USNM-6159), steel bomb Allende (CV3) (USNM-6159), steel bomb EET 92002 (CK4) ALH 83108 (CO3.5) ALH 83108 (CO3.5)

8

10 7 ⫺26 5

⫾ ⫾ ⫾ ⫾

19 49 34 30

⫺0.7 ⫺1.3 ⫺0.3 ⫺0.7

⫾ ⫾ ⫾ ⫾

0.5 1.0 0.7 0.9

⫺0.4 ⫺0.8 0.0 ⫺1.0

⫾ ⫾ ⫾ ⫾

0.3 0.9 0.6 0.6

⫺0.2 ⫺0.2 0.1 ⫺0.1

⫾ ⫾ ⫾ ⫾

0.1 0.2 0.2 0.2

0.5 0.7 ⫺0.1 0.2

⫾ ⫾ ⫾ ⫾

0.2 0.6 0.5 0.5

0.1 ⫺0.4 ⫺0.2 0.5

⫾ ⫾ ⫾ ⫾

0.3 0.9 0.9 0.9

8

⫺20 11 10 ⫺1

⫾ ⫾ ⫾ ⫾

20 49 34 12

⫺0.4 ⫺2.2 0.2 ⫺0.7

⫾ ⫾ ⫾ ⫾

0.6 1.0 0.7 0.4

⫺0.1 ⫺0.6 0.0 ⫺0.3

⫾ ⫾ ⫾ ⫾

0.5 0.9 0.6 0.3

0.0 ⫺0.1 0.1 0.0

⫾ ⫾ ⫾ ⫾

0.2 0.2 0.2 0.1

0.5 0.9 ⫺0.4 0.4

⫾ ⫾ ⫾ ⫾

0.3 0.6 0.5 0.2

0.0 0.1 ⫺0.1 0.4

⫾ ⫾ ⫾ ⫾

0.4 0.9 0.9 0.2

17

39 ⫾ 41 8 ⫺14 ⫺12 ⫺14

⫾ ⫾ ⫾ ⫾

41 34 34 30

⫺0.9 ⫾ 1.4 1.9 ⫺0.5 ⫺0.1 ⫺1.2

⫾ ⫾ ⫾ ⫾

2.0 0.7 0.7 0.9

⫺0.6 ⫾ 0.7 0.3 0.6 ⫺0.1 ⫺0.5

⫺0.2 ⫾ 0.3

⫺0.1 ⫾ 0.4

⫺0.1 ⫾ 0.9

⫾ ⫾ ⫾ ⫾

0.5 0.6 0.6 0.6

0.4 0.1 0.2 ⫺0.1

⫾ ⫾ ⫾ ⫾

0.4 0.2 0.2 0.2

⫺0.4 0.1 ⫺0.2 0.4

⫾ ⫾ ⫾ ⫾

0.7 0.5 0.5 0.5

⫺1.1 ⫺0.4 0.3 0.2

⫾ ⫾ ⫾ ⫾

0.9 0.9 0.9 0.9

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.6 0.6 1.0 0.6 1.0 0.6

⫺0.3 ⫺0.4 ⫺0.3 ⫺0.2 ⫺0.1 0.0

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.2 0.3 0.3 0.3 0.3 0.3

0.4 0.9 0.9 0.8 0.8 0.3

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.5 0.4 0.4 0.4 0.4 0.4

0.3 0.2 2.1 1.7 ⫺0.1 ⫺0.2

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.9 0.9 0.9 0.9 0.9 0.9

Ordinary chondrites Plainview (H5) (USNM-2285) Plainview (H5) (USNM-2285) Mezö-Madaras (L3.7) (USNM-4838) Mezö-Madaras (L3.7) (USNM-4838) ALH 84081 (LL6) ALH 84081 (LL6)

9 ⫺49 ⫺18 3 16 ⫺30

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

30 45 44 44 44 45

⫺1.5 ⫺2.3 ⫺1.0 ⫺1.7 ⫺0.5 ⫺1.1

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.9 1.6 1.4 1.6 1.4 1.6

⫺0.6 ⫺1.9 ⫺1.1 ⫺1.2 ⫺0.4 ⫺1.2

Enstatite chondrite 2 ⫾ 44

Abee (EH4) Abee (EH4)d

⫺1.4 ⫾ 1.4

⫺0.1 ⫾ 1.0

⫺0.3 ⫾ 0.3 0.0 ⫾ 0.3

0.5 ⫾ 0.4 0.9 ⫾ 0.4

1.0 ⫾ 0.9 0.8 ⫾ 0.9

Iron meteorites Mean Canyon Diablo (IA) metal and sulfide (USNM-676)c Mean Canyon Diablo (IA) sulfide (USNM-676) Canyon Diablo metal (IA) (USNM676) Canyon Diablo metal (IA) (USNM678) Canyon Diablo metal (IA) (USNM676) Canyon Diablo metal (IA) (USNM676) Mean Toluca (IA) sulfide (USNM75) Toluca metal (IA) (USNM-75) Toluca metal (IA) (USNM-75) Toluca metal (IA) (USNM-75) Arispe metal (IC) (ME2474) North Chile (IIA) metal

5

⫺4 ⫾ 17

⫺0.3 ⫾ 0.6

0.3 ⫾ 0.3

0.0 ⫾ 0.1

0.3 ⫾ 0.2

0.4 ⫾ 0.4

12

⫺8 ⫾ 13

⫺0.1 ⫾ 0.4

⫺0.2 ⫾ 0.2

0.0 ⫾ 0.1

0.5 ⫾ 0.1

0.2 ⫾ 0.3

⫺7 ⫾ 45

⫺0.5 ⫾ 0.8

0.5 ⫾ 0.6

⫺0.1 ⫾ 0.3

⫺0.5 ⫾ 0.4

0.2 ⫾ 0.9

⫺24 ⫾ 44

⫺0.5 ⫾ 1.2

0.6 ⫾ 0.9

0.3 ⫾ 0.3

0.1 ⫾ 0.4

0.4 ⫾ 0.9

⫺37 ⫾ 59

⫺1.2 ⫾ 1.5

⫺0.3 ⫾ 0.9

⫺0.3 ⫾ 0.3

⫺0.7 ⫾ 0.6

1.2 ⫾ 0.9

1 ⫾ 58

0.3 ⫾ 1.0

⫺0.4 ⫾ 0.8

⫺0.1 ⫾ 0.4

0.1 ⫾ 0.5

0.3 ⫾ 0.9

⫺0.5 0.8 0.7 0.7 1.8 1.0

⫺0.1 0.0 ⫺0.3 0.2 ⫺0.1 ⫺0.1

8

⫺1 49 11 ⫺39 ⫺1 26

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

15 44 59 58 62 62

⫺0.2 ⫺0.8 0.5 0.5 4.5 1.4

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.5 1.2 1.5 1.0 2.2 2.2

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.3 0.9 0.9 0.8 1.8 1.8

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.1 0.3 0.3 0.4 0.6 0.6

0.7 ⫺0.6 ⫺1.5 ⫺0.2 0.2 ⫺0.3

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.2 0.4 0.6 0.5 0.5 0.5

1.1 0.0 0.1 0.2 0.0 0.0

⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.3 0.9 0.9 0.9 0.9 0.9

a The quoted analytical uncertainties reflect the external reproducibilities of the standard obtained on a particular measurement session. For mean values, the uncertainties reflect the combined error of the individual measurements. Multiple digestions were performed only for Orgueil (n ⫽ 3), Murchison (n ⫽ 4) and Allende (n ⫽ 9). ␧xTeN ⫽ {[(xTe/128Te)N sample ⫺ (xTe/128Te)N std]/(xTe/128Te)N std} ⫻ 104, where x is 120, 122, 124, 126, or 130; N: normalized to 125Te/128Te ⫽ 0.22204 (Lee and Halliday, 1995) with the exponential law. ␦126/128TeM ⫽ {[(126Te/128Te)M sample ⫺ (126Te/128Te)M std mean]/(126Te/128Te)M std mean} ⫻ 103; M: measured. b Number of individual isotope measurements for mean values. c Te isotope data from Fehr et al. (2004). d No ␧120TeN, ␧122TeN, and ␧ 124TeN values are shown because of known Sn interference problems.

Tellurium isotopic composition of the early solar system

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Table 3. Tin and Te concentrations in meteorite samples and terrestrial sulfides. Sn (␮g/g) Sample

Weight (g)

This work

Te (␮g/g) Literature

This work

Literature

2.44 2.55 1.50 1.52 1.79 1.00 0.91 1.08 1.06 1.02 1.02 0.37 0.74 0.78 0.57 0.49 2.98

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.12 0.08 0.10 0.10 0.08 0.06 0.07 0.04 0.04 0.04 0.04 0.02 0.04 0.06 0.04 0.02 0.08

2.50d 2.26f 1.48d 1.60d 1.66f 1.04f 1.00d

12.2 6.7 0.176 0.080 0.100 0.081 0.061

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.4 0.4 0.020 0.018 0.006 0.017 0.008

72.4 4.99 0.072 0.047 2.30 1.03 1.21 1.19 0.93 13.8 10.4

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

1.6 0.12 0.020 0.008 0.10 0.06 0.04 0.09 0.04 0.2 0.2

118

Sn/128Tea

Chondrites Orgueil (CI1) Orgueil (CI1) ALH 83100 (CM1/2) Murchison (CM2) (USNM-5459) Murchison (CM2) (USNM-5459) ALH 84028 (CV3) Allende (CV3) (USNM-6159), steel bombb Allende (CV3) (USNM-6159), steel bombc Allende (CV3) (USNM-6159) Allende (CV3) (USNM-6159) Allende (CV3) (USNM-6159) EET 92002 (CK4) ALH 83108 (CO3.5) Plainview (H5) (USNM-2285) Mezö-Madaras (L3.7) (USNM-4838) ALH 84081 (LL6) Abee (EH4)

0.102 0.150 0.471 0.255 0.308 0.552 0.088 0.089 0.403 0.505 0.414 0.512 0.675 0.470 0.900 1.00 0.162

1.10 1.19 0.70 0.79 0.65 0.543 0.569 0.533 0.560 0.509 0.559 0.533 0.403 0.103 0.180 0.083 1.58

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

1.5d 1.64e 0.88d 0.89d 1e

0.02 0.02 0.01 0.02 0.02 0.015 0.024 0.015 0.010 0.012 0.010 0.020 0.017 0.006 0.006 0.006 0.02

0.58d 0.68e

0.56d 0.31e 0.75j–2.2k

0.96d 0.43g, 0.61h 0.875i 2.2g–2.75l

0.34 0.36 0.35 0.40 0.28 0.42 0.48 0.38 0.40 0.38 0.42 1.1 0.42 0.10 0.24 0.13 0.40

Iron meteorites Canyon Diablo sulfide (IA) (USNM-676) Toluca sulfide (IA) (USNM-75) Canyon Diablo metal (IA) (USNM-676) Arispe metal (IC) (ME2474) Arispe metal (IC) (ME2474) North Chile metal (IIA) North Chile metal (IIA)

0.732 0.630 6.12 11.3 5.51 5.39 8.30

0.018 0.014 2.91 0.451

⫾ ⫾ ⫾ ⫾

0.002 0.002 0.10 0.020

0.14 ⫾ 0.08 0.131 ⫾ 0.010

m

3.6 3.0n

⬍1n, 20.2o

5.0h 1.7h 0.09h

0.001 0.002 12.6 4.3 1.3 1.6

Terrestrial sulfides SV1 55.2 SV1 110.5 RHDH2a 161.7 RHDH2a 205.2 RHDH2a 299.2 pyrrhotite pyrrhotite pyrrhotite pyrrhotite pyrite pyrite

0.045 0.045 0.057 0.057 0.050 0.519 0.521 0.502 0.501 0.112 0.127

0.259 0.053 0.196 0.406 0.815 0.084 0.125 0.346 0.143 0.225 0.144

⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾ ⫾

0.008 0.006 0.008 0.006 0.016 0.006 0.002 0.004 0.012 0.008 0.006

0.003 0.008 2.08 6.60 0.27 0.063 0.079 0.22 0.12 0.012 0.011

a The 118Sn/128Te ratios of carbonaceous chondrites are estimated to have an uncertainty (2s) of ⫾20% from repeated analyses of Allende. All chondrites were digested using a high-pressure asher, except for two Allende dissolutions, which were performed using a Teflon bomb with steel jacketb,c. Multiple analyses on separate digestions were performed for Orgueil (n ⫽ 2), Murchison (n ⫽ 2), and Allende (n ⫽ 5), North Chile (n ⫽ 2), the pyrrhotite (n ⫽ 4), and the pyrite (n ⫽ 2). For Allende, 15 b and 4 c individual measurements were performed on the same sample solution. Literature data are from dFriedrich et al. (2002), eDe Laeter et al. (1974), fXiao and Lipschutz (1992), gSmith et al. (1977), hGoles and Anders (1962), i Keays et al. (1971), jDe Laeter and Jeffery (1967), kHamaguchi et al. (1969), lReed and Allen (1966), mD’Orazio and Folco (2003), nOnishi and Sandell (1957), oWinchester and Aten (1957).

these are the first published Sn and/or Te concentration data. Individual digestions of Allende obtained with either a Teflonlined steel bomb (n ⫽ 2) or a high-pressure asher (HPA-S, n ⫽ 3) show a reproducibility (2s) of 12% for the Te and of 10% for the Sn concentration. The 118Sn/128Te ratios determined for Allende vary by 20%, however, which suggests that Sn and Te may exhibit a heterogeneous distribution within this meteorite. For two dissolutions of both Orgueil and Murchison the variability of the Sn and Te concentrations is similar to and slightly larger than the Allende data, respectively. Based on this, we use the reproducibility obtained for Allende to estimate the uncertainty of the 118Sn/128Te ratios determined for all other carbonaceous chondrites (Table 3). The Te concentrations

determined in this study for carbonaceous chondrites agree very well with the results of Friedrich et al. (2002), whereas the Sn concentrations are generally lower by 6%–28% (Table 3). The cause of this systematic discrepancy is presently unclear. It is unlikely to reflect loss of Sn from the sample solutions by precipitation or adsorption, because Sn was stabilized by the addition of either HCl or HF. Incomplete digestion of resistant Sn phases (such as SnO2) is also improbable because similar Sn concentrations were determined for the HPA-S digests (which typically display some undissolved residue) and the steel bomb digests of Allende, which yield perfect sample solutions without any residue. For the ordinary chondrites, some of our results deviate

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significantly from the very scarce literature data (Table 3). It is possible that these discrepancies reflect either analytical errors (especially for the older literature results) or sample heterogeneity. The latter explanation is supported by the observation that many trace elements are more heterogeneously distributed in ordinary chondrites than in carbonaceous chondrites, which is partly thought to be caused by the inhomogeneous occurrence of metal phases (Wolf and Lipschutz, 1995; Friedrich et al., 2002, 2003). Good agreement between literature data and the present results is observed for the Abee enstatite chondrite (Table 3). The Sn and Te concentrations of metals from iron meteorites decrease in the order IA ⬎ IC ⬎ IIA, as is expected from the volatile element depletion patterns of these meteorites (Scott and Wasson, 1975). The Te concentrations of troilite inclusions from Canyon Diablo and Toluca are significantly higher compared to the data of Goles and Anders (1962) but this may simply reflect the variable trace element contents of individual sulfide inclusions. Several individual digestions of the terrestrial pyrrhotite display reproducibilities (2s) of 140 and 24% for the Sn and Te concentrations, respectively. This is in agreement with the observation of Yi et al. (1995) that the distribution of Sn in sulfides is highly heterogeneous and with the suggestion of Gaboury and Graham (2004) that the Sn of pyrites might be hosted in microinclusions. All terrestrial sulfides measured in this study have Sn concentrations of between 0.05 and 0.8 ␮g/g. This is at the lower end of the range of Sn abundances previously determined for terrestrial sulfides (0.1–250 ␮g/g; Yi et al., 1995). The Te contents of these samples are much more variable with concentrations that cover three orders of magnitude (0.05–72.4 ␮g/g), which is in accord with previously published data (Leutwein, 1972). The Archean sulfides with the lowest Te abundances (RHDH2a 161.7 and 205.2) have negative ⌬33S values, whereas sulfides with higher Te contents show positive ⌬33S (Table 1 and 2). Negative ⌬33S values are thought to be inherited from seawater sulfate, whereas positive ⌬33S is thought to be derived from sulfur (Ono et al., 2003). In addition, the sulfides with low Te are from the Carawine dolomite, whereas the ones with higher Te concentrations are from either the Lewin Shale (RHDH2a 299.2) or the Tumbiana Formation (SV1 55.2 and 110.5). However, given the small number of samples analyzed, it cannot be deduced whether the variations in Te contents determined for the sulfides are due to different formation mechanisms or distinct environments. 4.2. Isotope Compositions of Terrestrial Sulfides As sample size was limited, the Te isotopic compositions of two sulfides (RHDH2a 161.7 and 205.2) could not be analyzed to high precision due to the low Te concentrations (Table 3). For the data that are corrected for mass fractionation by internal normalization, any significant deviation of the isotope composition from ␧1XXTeN ⫽ 0 can be interpreted as a mass-independent isotope effect. All five Archean sulfides are, however, characterized by ␧1xxTeN values that are identical to the terrestrial standard, within error (Figs. 1, 2, Table 1). In particular, sample RHDH2a 299.2, which has the largest ⌬33S anomaly of ⫹6.2 displays ␧126TeN ⫽ ⫹0.4 ⫾ 0.5, and sample RHDH2a 161.7, which has the most negative ⌬33S value of ⫺2.46 is characterized by ␧126TeN of ⫺8.2 ⫾ 13.5. Furthermore, none of the sulfides display any clearly resolvable mass dependent

Fig. 1. Results obtained for ␧126TeN for chondrites, iron meteorites and terrestrial sulfides. Shown are the combined mean values of several measurements. The error bars (2s) denote the combined uncertainty of the individual sample measurements. m ⫽ metal; s ⫽ sulfide; HPA ⫽ high-pressure asher.

Te isotope fractionations, given the analytical precision of about ⫾ 0.9‰, with ␦126/128TeM values varying from ⫺1.0‰ to ⫹1.2‰ (Fig. 3, Table 1). 4.3. Meteorites The Te isotope compositions acquired with internal normalization for the bulk chondrite samples and the metal and sulfide phases of iron meteorites are summarized in Figures 1 and 2 and Table 2. All results are identical to the terrestrial standard within the 2s analytical uncertainties. This is in agreement with previous, albeit less precise, Te isotope data for bulk meteorite samples, which were obtained by TIMS (Smith et al., 1978; De Laeter and Rosman, 1984; Smith and De Laeter, 1986). Small isotopic variations are apparent for some samples, for example the ␧122TeN and ␧124TeN results of Arispe or the ␧130TeN values of the ordinary chondrites and Abee, but these are not resolvable given the present uncertainties. The ␧130TeN data show the largest overall variations, but it is not clear if these reflect true isotopic differences or small residual analytical artifacts. Similar variations are also observed for synthetic rock samples, which consist of a Te-free matrix that was doped with Te standard (Fehr et al., 2004), such that the latter explanation is considered to be more likely. The Te stable isotope compositions of the meteorites range between ⫺1.1‰ to ⫹2.1‰ in ␦126/128TeM (Fig. 3, Table 2), with an analytical uncertainty of about ⫾ 0.9‰ (2s). The only

Tellurium isotopic composition of the early solar system

Fig. 2. Results obtained for ␧120TeN, ␧122TeN, ␧124TeN, and ␧130TeN for chondrites, iron meteorites and terrestrial sulfides. m ⫽ metal; s ⫽ sulfide; HPA ⫽ high-pressure asher.

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Fig. 3. Te stable isotope data, as ␦126/128TeM, for different chondrites, iron meteorites and terrestrial sulfides. m ⫽ metal; s ⫽ sulfide; HPA ⫽ high-pressure asher.

sample with a clearly resolvable stable isotope effect is MezöMadaras (L3.7) but the measured fractionation is too small to completely exclude that the result reflects analytical artifacts from matrix effects. 5. DISCUSSION

5.1. Terrestrial Sulfides 5.1.1. Stable isotope fractionation of tellurium The general chemical and geochemical behavior of Te is similar to that of S and Se, and the latter elements display significant stable isotope fractionations in nature. Biologic processes typically generate large mass-dependent S isotope effects of up to 15% (Nielsen, 1979; Schidlowski et al., 1983) and Se isotopes have been observed to vary by as much as 1.5% in low temperature environments, mainly due to the inorganic and biologic reduction of various Se species (Johnson et al., 1999; Rouxel et al., 2002; Ellis et al., 2003; Johnson and Bullen, 2003). In contrast, high temperature isotope fractionations of S and Se in ore deposits and igneous rocks are small (Ohmoto and Rye, 1979; Ohmoto, 1986; Rouxel et al., 2002). There is only a single previous investigation, which addressed potential natural variations in the stable isotope composition of Te (Smithers and Krouse, 1968). Based on laboratory experiments, these workers claimed that the inorganic and biologically mediated reduction of tellurite to elemental tellurium can generate kinetic isotope effects of up to 6‰ that were thought to be due to the cleavage of Te-O bonds. In addition, Smithers and Krouse (1968) observed a 4‰ variation in 130Te/ 122 Te for six natural telluride samples that were analyzed relative to a laboratory standard.

The five Archean sulfides that were analyzed in this study are pyrites, which are thought to have formed under the influence of microorganisms by bacterial reduction of seawater sulfate and elemental sulfur (Ono et al., 2003). Such reactions are known to generate mass-dependent sulfur isotope effects. Present-day sedimentary sulfides are commonly depleted in 34S by 40 to 70‰ compared to seawater sulfate, but Archean sulfides typically show less depletion, probably because the oceans had a lower sulfate concentration at that time (Canfield, 1998; Canfield et al., 2000). The present samples are in accord with this trend, as they exhibit ␦34S values of ⬃0 to 16‰ (Table 1). The Archean sulfide samples have the potential to show Te isotope fractionations from biologic processing. The methods used in this study were not primarily developed for precise stable Te isotope measurements, however, and only fractionations larger than ⬃0.9‰ in ␦126/128TeM are resolvable. At this level of precision, none of the Archean sulfides display mass dependent stable isotope effects for Te (Fig. 3, Table 1). Several interpretations can account for the lack of Te isotope effects, assuming that the samples were not significantly altered following deposition. It is possible, that the fractionation factors were too small to generate resolvable variations. In this case, the small fractionation factors may reflect the lack of biologic processing of Te and/or the smaller relative mass difference between the most abundant Te isotopes, which can be measured to high precision (126Te, 128Te, 130Te), as compared to 32S and 34S. Alternatively, no fractionation may have been recorded because the sulfides quantitatively incorporated the Te from a reservoir of limited size. Seawater or pore fluids with low Te contents may constitute such a reservoir. Nevertheless, the present results do not invalidate the possibility of significant Te isotope fractionation on Earth and a more systematic study of a larger sample suite will be necessary to resolve this issue. 5.1.2. Mass-independent isotope fractionation of tellurium Mass-independent fractionations (MIF) of S isotopes are found only in Archean sulfides older than ⬃2450 Ma, whereas younger sulfides show only limited ⌬33S variations of less than ⫾ 0.1‰. The MIF of S isotopes is thought to be due to the UV-induced photolysis of gaseous SO2 that was expelled into the Archean atmosphere by volcanic activity. As such, these isotope anomalies provide compelling evidence that the Earth’s atmosphere was not oxygenated before 2450 Ma (Farquhar et al., 2000). Given the chemical similarity of S and Te, it is possible that certain Te species may also show mass-independent isotope effects. Significant outgassing of Te is known to occur during volcanic activity but in the current atmosphere, Te appears to be present mainly in the particulate fraction (Greenland and Aruscavage, 1986). A photolytic reaction analogous to that inferred for SO2 is, furthermore, unlikely for Te because TeO2 is nonvolatile with a boiling point (Tb) of 1245°C. Tellurium, however, forms the volatile hydride TeH2 (Tb ⫽ ⫺1°C) and various volatile halides (e.g., TeF4, TeCl2, TeCl4, TeBr2, TeBr4 with Tb ⬇ 190 to 410°C), which may have had longer residence times in the anoxic Archean atmosphere as compared to the present. These volatile compounds can also be decomposed by UV radiation (Samuel, 1946). If such photochemical reactions

Tellurium isotopic composition of the early solar system

indeed occurred in the early Earth, they may have generated MIF for Te, which could have been preserved in Archean sulfides. As MIF can be identified with isotope data that are acquired with internal normalization, such effects can be resolved much more precisely than mass dependent fractionations. Despite of this analytical advantage, none of the Archean pyrites (or the more recent sulfides) display any resolvable Te isotope anomalies (Figs. 1, 2, Table 1). A number of interpretations can account for the lack of MIF for Te isotopes: (1) the atmospheric content of suitable gaseous Te species was insufficient or the atmospheric residence time too short; (2) there are no (gas phase) reactions that can generate significant MIF for Te isotopes; (3) any MIF signature that was produced in the atmosphere was diluted or overprinted in the pyrites by isotopically normal Te. The present results unfortunately do not provide constraints on which interpretation is most likely to be correct. 5.2. Meteorites 5.2.1. Stable isotopes No previous attempts have been made to detect variations in the stable isotope composition of Te in meteorites, but some results are available for Se and Cd, which are broadly similar in volatility. Data for Se are limited to iron meteorites, where no isotopic fractionation was observed (Rouxel et al., 2002). In contrast, large Cd isotope fractionations were reported for unequilibrated ordinary chondrites (Rosman and De Laeter, 1988; Wombacher et al., 2003). Terrestrial rock and mineral samples and bulk carbonaceous chondrites, however, were found to display only very limited variations in Cd isotope compositions (from ⫺0.10 to ⫹0.12‰/amu) that are barely resolvable given an uncertainty of about ⫾ 0.12‰/amu (Wombacher et al., 2003). Cadmium is a highly volatile element with a 50% condensation temperature (at 10⫺4 bar) of ⬃652 K (Lodders, 2003) and the isotopic variations detected in meteorites are thought to reflect partial evaporation and/or condensation during redistribution of Cd on the parent bodies. Such fractionations are, in principle, also possible for Te, which is moderately volatile with a 50% condensation temperature of 709 K (Lodders, 2003). Heating experiments have furthermore shown that both Cd and Te can be redistributed and partially lost from chondrites by thermal metamorphism (Ngo and Lipschutz, 1980) and Ikramuddin et al. (1977) found that ⬃95%–99% of the Te was lost from the L3 chondrite Krymka in the temperature range of 400 to 1000°C. Mobilization of Cd, however, already begins at ⬃400°C, whereas Te requires temperatures of at least 700°C (Ngo and Lipschutz, 1980). Fractionation effects from thermal redistribution may therefore be smaller for Te than for Cd. Given the published Cd and Se stable isotope data, it is not surprising that the carbonaceous chondrites and iron meteorites analyzed in the present study do not show stable isotope effects for Te (Fig. 3, Table 2). The single unequilibrated ordinary chondrite that was analyzed, Mezö-Madaras (L3.7), yielded the most fractionated isotope composition with a mean ␦126/128TeM of 1.9 ⫾ 0.6‰. This result appears to be in accord with the large Cd isotope fractionations of up to 3.6‰/amu that have been determined for unequilibrated ordinary chondrites (Ros-

5107

man and De Laeter, 1988; Wombacher et al., 2003) but further verification of the anomaly with a method developed specifically for Te stable isotope ratio measurements is desirable. 5.2.2. Chondrites and the initial system

126

Sn/118Sn of the solar

Carbonaceous chondrites are, in principle, ideally suited for detecting anomalies in the abundance of 126Te that are due to decay of 126Sn, because they are very early objects that are likely to have retained the original Te isotope signature that they acquired in the nascent solar system. An inspection of the Sn/Te ratios reveals, however, that these meteorites exhibit only a very limited fractionation of these two elements. The moderately volatile elements are generally progressively depleted in carbonaceous chondrites with increasing volatility in the order CI ⬎ CM ⬎ CO ⬎ CV (Larimer and Wasson, 1988; Palme et al., 1988). This depletion pattern is thought to reflect incomplete condensation and loss of the more volatile elements from the inner solar system during planetesimal formation. Older studies have, furthermore, inferred that Te (680 K) has a slightly lower half mass condensation temperature than Sn (720 K) (Wai and Wasson, 1977; Wasson, 1985). A recent investigation has, however, calculated nearly identical condensation temperatures for Te and Sn, with values of 705 K and 703 K, respectively (Lodders, 2003). This result suggests that Sn and Te should not be significantly fractionated from one another during condensation of the nebular, which is in accord with the observation that the 118Sn/128Te ratios of bulk carbonaceous chondrites vary only slightly between ⬃0.35 and 1.1 (Table 3). The ␧126TeN data acquired for all carbonaceous chondrites is identical, within error, to the terrestrial standard (Fig. 1, Table 2). Assuming that the Sn/Te ratios of these meteorites reflect variable early volatile element depletion during condensation of the portion of the nebula that represents their parent bodies, the isotope data can be plotted in an isochron diagram, which yields an upper limit for the initial 126Sn/118Sn of the solar system of ⬍4 ⫻ 10⫺5 (Fig. 4). For this to accurately reflect the initial composition of the solar system the Sn/Te fractionation must be a very early (nebular) feature and there should have been no subsequent Sn/Te redistribution on a bulk scale. The CK4 carbonaceous chondrite Elephant Moraine (EET) 92002 has the largest 118Sn/128Te ratio of 1.1. Given that all other carbonaceous chondrites have a very limited variation in their 118 Sn/128Te of 0.34 – 0.42, the possibility that the 118Sn/128Te ratio of the CK chondrite was fractionated at a late stage needs to be considered. An elevated Sn/Te ratio could be due to loss of Te during metamorphism (Wang and Lipschutz, 1998), which is not unreasonable given that the CK4 meteorite has the highest metamorphic grade of the analyzed carbonaceous chondrites. As EET 92002 is an Antarctic meteorite, there is also a potential for Te loss during terrestrial alteration. The observation that other Antarctic carbonaceous chondrites analyzed in this study have Sn/Te ratios that are similar to the non-Antarctic samples argues against this scenario (Table 3 and Xiao and Lipschutz, 1992). The effect of Antarctic weathering on trace element abundances can be very heterogeneous, however (Crozaz et al., 2003). Given that CK chondrites have not been previously studied for their homogeneity regarding trace element contents it is possible that the high Sn/Te determined for

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Fig. 4. Sn-Te isochron diagram for bulk carbonaceous chondrites. The 118Sn/128Te data are from Table 3. The 126Te/128TeN ratios are combined mean values for each sample and are recalculated from the data of Table 2 assuming that ␧126TeN ⫽ 0 corresponds to 126Te/128TeN ⫽ 0.592260 (Fehr et al., 2004). Error bars (2s) represent the combined uncertainties of the sample and standard measurements: sstd⫹sample 2 2 ⫽ 公(sstd ⫹ ssample ); errors for 118Sn/128Te are 2s. The initial 126Sn/ 118 Sn for all carbonaceous chondrite samples is ⬍4 ⫻ 10⫺5 and ⬍3.4 ⫻ 10⫺4 without the CK chondrite EET 92002.

EET 92002 reflects sample heterogeneity and/or terrestrial alteration. Exclusion of this sample from the isochron plot (Fig. 4) results in a less well-constrained initial 126Sn/118Sn of ⬍3.4 ⫻ 10⫺4. The ␧126TeN results for the ordinary and enstatite chondrites, and the iron meteorites are identical to the carbonaceous chondrites (and the terrestrial standard), despite of more fractionated 118 Sn/128Te ratios (Table 2 and 3, Fig. 1). The parent bodies of these meteorites, however, experienced heating and/or melting episodes, during which Sn and Te were probably mobilized and the Te isotope compositions homogenized. The Te isotope data of these samples are therefore unsuitable for deducing the initial abundance of 126Sn. The same is true for the silicate Earth, which is characterized by a Te isotope composition that is unrelated to the bulk Earth’s Sn/Te ratio. This is because most of the proto-Earth’s Te budget was probably partitioned into the core, such that the Te isotope composition of the silicate Earth is dominated by the late (chondritic) veneer (Yi et al., 2000). The lack of a detectable ␧126TeN anomaly for any of the meteorites limits the amount of 126Sn that may have been present in the early solar system, but it does not demonstrate that 126Sn was completely absent or provides evidence against a supernova trigger of solar system formation. A type-II supernova would generate a freshly synthesized supply of short-lived isotopes such as 26Al, 41Ca, 53Mn, 60Fe, and 107Pd, even though very special conditions may be required to obtain initial abundances that are in accord with the published analytical data (Russell et al., 2001; Busso et al., 2003). If the 41Ca that has been detected in meteorites was indeed mainly derived from a supernova source, the short half-life of this isotope (⬃0.103 Myr) would also indicate that the free decay interval between nucleosynthetic production and the condensation of the first solids was short. The absence of Te isotope anomalies is nonetheless in accord with a supernova scenario for the following reasons. (1) The cosmochemical evidence for the former existence of many extinct isotopes is provided by anal-

yses of specific materials (e.g., CAIs for 26Al, 41Ca and highly volatile depleted iron meteorites for 107Pd) that feature highly fractionated parent-daughter ratios and which are therefore particularly suitable for the detection of isotope anomalies from radiogenic decay (Russell et al., 2001). The carbonaceous chondrites that were analyzed in the present study for Te isotopes, however, show only a very limited variation in Sn/Te ratio. This severely restricts our ability to detect the former presence of small amounts of live 126Sn, despite the extremely precise isotopic measurement techniques. (2) It is possible that anomalies in 126Te/128Te were not observed in this study because too little 126Sn was synthesized compared to other shortlived isotopes. To our knowledge there are no current supernova models that predict the nucleosynthetic production yields of various isotopes relative to 126Sn. In the absence of such models, the constraint of 126Sn/118Sn ⬍4 ⫻ 10⫺5 cannot be used to argue against a supernova trigger. (3) It is also conceivable that the lack of radiogenic Te isotope anomalies in bulk meteorites reflects the redistribution of Te (and/or Sn) by alteration processes. 5.2.3. Investigation of nucleosynthetic Te isotope anomalies A number of recent studies, many of which utilized MCICPMS, identified nucleosynthetic isotope anomalies for the intermediate mass elements Zr (Sanloup et al., 2000; Yin et al., 2001), Mo (Masuda and Lu, 1998; Dauphas et al., 2002, 2004; Yin et al., 2002; Chen et al., 2004), Ba (Harper et al., 1992; Hidaka et al., 2003), and Ru (Chen et al., 2003; Papanastassiou et al., 2004) in bulk samples of chondritic and iron meteorites. A few of these results have been contested (Lee and Halliday, 2002; Becker and Walker, 2003a,b; Schönbächler et al., 2003), but it is likely that some provide “real” evidence for the existence of isotopic heterogeneities in the early solar disk. Several investigations have furthermore recognized that various types of presolar grains display large isotopic anomalies for Ba, Mo, Ru, and Zr (Nicolussi et al., 1997, 1998a,b; Savina et al., 2003, 2004). These observations support the interpretation that the isotopic heterogeneities found for the meteorites could be related to the presence (or absence) of anomalous presolar grains or material derived from such grains (Dauphas et al., 2002; Yin et al., 2002). The presolar diamonds of the Allende chondrite display large (‰ to % level) nucleosynthetic Te isotope anomalies (Richter et al., 1998; Maas et al., 2001). If such exotic grains were heterogeneously distributed in the early solar system, then bulk meteorite samples may display evidence of this heterogeneity. The internally normalized Te isotope compositions acquired in the present study, however, are devoid of any significant anomalies (Fig. 2, Table 2). A few very small isotopic variations are nonetheless apparent in Figure 2. In the following, possible causes for these variations are discussed, based on the assumption that they do not reflect analytical artifacts. Some of the iron meteorites as well as the ordinary chondrites and Abee display slightly elevated or reduced abundances of the s-process isotopes 122Te and 124Te. This could reflect variations in the s-process contribution, as was previously inferred for Ru and Mo (Dauphas et al., 2002, 2004; Chen et al., 2003, 2004; Papanastassiou et al., 2004) but the most anomalous ␧122,124TeN data has large uncertainties. In addition, there is a hint of a variation in ␧130TeN, particularly between the metal of

f

Molar fraction of an element derived from a presolar component that can be admixed to a reservoir of terrestrial isotope composition without producing a resolvable isotopic difference. This is calculated for the mean (fmean) and the maximum (fmax) presolar isotope composition, based on molar isotope mixing.

Lodders (2003).

x represents the mass fraction of an element, which is derived from the presolar component; x ⫽ abundance in presolar phase ⫻ abundance of presolar component in bulk meteorite/abundance in bulk meteorite. e

d

⬍300 ⫻ 10

For Te from Maas et al. (2001); for Ba recalculated from Amari et al. (1995), assuming Zr and Si concentrations as in SiC of Kashiv et al. (2001); for Mo, Ru, and Zr from Kashiv et al. (2001).

For Allende from Jarosewich et al. (1987); for Murchison from Friedrich et al. (2002). c

Huss et al. (2003). a

b

17 ⫻ 10⫺6

49 ⫻ 10 3.96

3.96 6.9 Graphite in Murchison Zr

4.7

SiC in Murchison Zr

13.5

25

6.9

⬍130 ⫻ 10⫺6

⬍15 ⫻ 10⫺6

⬍320 ⫻ 10⫺6 ⫺6 ⫺6

85 ⫻ 10⫺6 0.69 1.4 8.8 SiC in Murchison Ru

13.5

⬍240 ⫻ 10⫺6 SiC in Murchison Mo

13.5

8.7

1.8

1.02

65 ⫻ 10⫺6

⬍20 ⫻ 10⫺6

Maas et al. (2001), Richter et al. (1998) Zinner et al. (1991), Ott and Begemann (1990), Hidaka et al. (2001) Becker and Walker (2003a), Nicolussi et al. (1998a) Becker and Walker (2003b), Savina et al. (2004) Schönbächler et al. (2003), Nicolussi et al. (1997) Schönbächler et al. (2003), Nicolussi et al. (1998b) ⬍20 ⫻ 10 ⬍40 ⫻ 10⫺6 Diamond in Allende SiC in Murchison Te Ba

340 13.5

0.64 7.3

1.2 2.9

2.33 2.31

181 ⫻ 10 34 ⫻ 10⫺6

⬍0.01 ⬍100 ⫻ 10⫺6

⫺6

Presolar phase Element

5109

⫺6

fmaxf fmeanf xe Abundance in Cl chondrites (␮g/g)d Abundance in bulk meteorite (␮g/g)c Abundance in presolar phase (␮g/g)b

The possible effect of a presolar component was estimated by calculating (1) the mass fraction x of an element which is contained in the presolar component relative to the bulk concentration of the element in a specific meteorite and (2) the molar fraction f of an element carried by a presolar phase that can be admixed to a reservoir of terrestrial isotope composition without producing a resolvable isotopic effect (Table 4). The molar fraction f is calculated both for the most extreme measured presolar isotope composition of an element (fmax) and the mean isotope composition determined for several presolar grains (fmean). For Te, fmean is less than 0.01 (1%) for a mean presolar isotope composition of ␧128Te ⫽ 40 ⫾ 15 and ␧130Te ⫽ 93 ⫾ 28, as measured by Maas et al. (2001), whereas fmax is smaller than 20 ⫻ 10⫺6 using the extreme isotope composition (␧120Te ⫽ ␧122Te ⫽ ␧123Te ⫽ ␧124Te ⫽ ␧125Te ⫽ ␧126Te ⫽ ⫺10,000, and ␧130Te ⫽ ⫺800) determined by Richter et al. (1998). The Te isotope composition of bulk Allende (without acid insoluble presolar diamonds) is within error identical to that of the Earth (as defined by terrestrial sulfides and the Te standard) and other meteorites (Figs. 1, 2). Assuming a homogeneous distribution of presolar diamonds over the disk, this implies that the fraction of Te derived from presolar diamonds in Allende with a composition as determined by Maas et al. (2001) has to be ⬍0.01. Otherwise, the Te isotopic composition

Abundance of presolar phase in bulk meteoritea (␮g/g)

5.2.4. Modeling

Table 4. Estimates of potential isotopic effects from presolar grains in bulk meteorites for Te, Ba, Mo, Ru, and Zr.

iron meteorites and the accompanying sulfides. As these phases are generally thought to be in equilibrium, these variations cannot be due to distinct presolar components. In terrestrial uraninites, fission produced 130Te from U has been observed by Hidaka and Masuda (1993) and Goles and Anders (1962) have reported U concentrations of 10 and 3.5 ng/g U for sulfides of Toluca and Canyon Diablo, respectively. A uranium content of ⬃3%–7% would be needed, however, to account for a 130Te excess of 1␧ unit, assuming that the sample is 4.6 Gyr old. Fission-produced Te therefore cannot be the cause of the observed small variations in ␧130TeN (⬃1 ␧ unit). The data of this study are in accord with a solar system that was well mixed on a large scale before the accretion of the Earth and the meteorite parent bodies. Rouxel et al. (2002) suggested that Te (and Se) isotope studies would be useful for investigating the origin of the extraterrestrial material that bombarded the Earth during the late accretionary stages. The uniformity of Te isotope compositions for the Earth and the analyzed meteorites indicates, however, that such a study would not provide any new constraints. The observed Te isotope homogeneity also contrasts with the large Te isotope anomalies of presolar diamonds and the isotopic heterogeneities identified in bulk meteorite samples for other elements, such as Ba, Mo, Ru, and Zr. There are three plausible interpretations for these observations. (1) Presolar grains with exotic Te isotope compositions are less common than the presolar carriers of the nucleosynthetic isotope signatures of other elements. (2) Such grains have Te isotope signatures that are significantly less extreme. (3) Alternatively, the presolar grains with anomalous Te may typically have low Te contents, such that they exhibit low Te/Mo and Te/Ru ratios. These interpretations are evaluated in the following, based on simple mass balance constraints.

Reference

Tellurium isotopic composition of the early solar system

5110

M. A. Fehr et al.

of the acid soluble components of Allende would be different from other meteorites and the Earth. This conclusion is in agreement with the calculated mass fraction x of Te in the presolar diamonds of Allende, which is ⬃200 ⫻ 10⫺6 (Table 4). Therefore, nanodiamonds could be more abundant by a factor of 50 compared to their measured content in Allende without producing a resolvable Te isotope effect. Presolar silicon carbide grains have anomalous isotopic compositions for many elements, including Ba, Mo, Ru, and Zr. The concentrations of these elements in the SiC grains of Murchison are about one order of magnitude larger than the Te contents of Allende presolar diamonds (Table 4). However, the abundance of presolar diamonds in Allende is also about an order of magnitude larger than the concentration of SiC grains in Murchison (Table 4). For Zr, the molar fraction f is 130 ⫻ 10⫺6 based on the average composition of the r-process component determined for presolar graphite grains and ⬍15 ⫻ 10⫺6 for one of the most extreme graphite grain compositions (Schönbächler et al., 2003). For the Zr s-process contribution (in SiC grains), fmean has to be ⬍300 ⫻ 10⫺6 (Schönbächler et al., 2003). The limits on the molar fraction f of admixed Ba and Mo from presolar material are similar to those of Zr (Table 4). Ruthenium in presolar SiC grains is enriched in 99Ru due to the decay of 99Tc and it was inferred that the maximum expected Ru isotope anomaly from the admixed presolar component would be about as large as the present uncertainties of bulk meteorite analyses (Savina et al., 2004). The fraction fmax of admixed Ru with an extreme presolar isotope composition that cannot be resolved in bulk samples is ⬍320 ⫻ 10⫺6, which is about an order of magnitude higher compared to the elements Te, Ba, Mo, and Zr. The maximum amounts of Ba, Mo, Zr, and Te that can be derived from presolar grains with extreme isotope compositions (fmax) is very similar, but fmean is about two orders of magnitude larger for Te than for the other elements (Table 4). The presently available data therefore suggest that nucleosynthetic isotope anomalies in bulk meteorite samples are much more likely for Ba, Mo, and Zr than for Te. This is supported by the mean Te isotope composition of presolar diamonds (Maas et al., 2001), which is much less extreme than the Ba, Mo, and Zr isotope data for presolar SiC and graphite grains (Ott and Begemann, 1990; Nicolussi et al., 1997, 1998a,b). It was suggested that at least some nanodiamonds have a solar rather than a presolar origin (Dai et al., 2002; Nittler, 2003), and this might explain the less extreme mean isotope composition of Te. 6. CONCLUSIONS

New Te isotopic data shows that chondrites, sulfide and metal separates of iron meteorites, and the Earth have identical Te isotope compositions. The results provide no evidence of stable Te isotope fractionations of larger than 2‰ for ␦126/128Te. Precambrian sulfides, which exhibit large mass-independent fractionations for S isotopes, furthermore display no such effects for Te. The meteorite data suggests that solar system materials were well mixed on a large scale in terms of Te isotope compositions, even though some presolar grains are known to have large Te isotope anomalies. Any ␧126TeN anomalies from the decay of 126Sn are either too small to be resolvable or absent.

Data obtained for bulk carbonaceous chondrites indicate that the initial 126Sn/118Sn of the solar system was less than 4 ⫻ 10⫺5. However, this assumes that the Sn/Te ratios of the samples reflect a very early nebular fractionation with no subsequent bulk-scale redistribution of Sn and Te. Future work needs to focus on more suitable samples for the detection of possible 126Te excesses from the decay of 126Sn. Such samples are early formed phases with high Sn/Te ratios, which are undisturbed and primary, as Te may be redistributed easily during alteration and metamorphism. Early formed metal from unequilibrated chondrites may be a suitable material, as Sn is enriched relative to Te in metal. Alternatively, the sequential leaching of carbonaceous chondrites may dissolve phases with different Sn/Te ratios that could reveal Te isotope differences, in the same manner as it is thought that 135Ba effects reflect the decay of 135Cs (Hidaka et al., 2001; Hidaka et al., 2003). Other promising samples are CAIs, which are the host of many anomalies from short-lived radionuclides. Some CAIs in Allende have high Sn concentrations (Mason and Martin, 1977) and may therefore also have high Sn/Te ratios. The petrography of potentially suitable inclusions should be carefully studied before isotopic analysis because some CAIs of oxidized CV3 carbonaceous chondrites (including Allende) appear to have experienced severe alteration (Krot et al., 1995). Acknowledgements—We thank Der-Chuen Lee, Sarah Woodland, Maria Schönbächler, and Don Porcelli for many helpful discussions. We are grateful to Glenn MacPherson, Brigitte Zanda, Marilyn Lindstrom, and Meenakshi Wadhwa for providing samples from the meteorite collections of the Smithsonian Institution of Washington, the Natural History Museum in Paris, NASA and the Field Museum of Chicago, respectively. We would like to thank Kate Freeman for making the Precambrian sulfide mineral separates available to us, Rio Tinto Exploration for providing access to the Archean cores, and the National Science Foundation (grant EAR-00-73831) for funding sample collection. Financial support was provided by the ETH Forschungskomission and the Schweizerische Nationalfonds (SNF). The very helpful reviews of Meenakshi Wadhwa and two anonymous referees were much appreciated and we acknowledge the prompt editorial handling of Rich Walker. Associate editor: R. Walker REFERENCES Amari S., Hoppe P., Zinner E., and Lewis R. S. (1995) Trace-element concentrations in single circumstellar silicon carbide grains from the Murchison meteorite. Meteoritics 30, 679 – 693. Arndt N. T., Nelson D. R., Compston W., Trendall A. F., and Thorne A. M. (1991) The age of the Fortescue Group, Hamersley Basin, Western Australia, from ion microprobe zircon U-Pb results. Austral. J. Earth Sci. 38, 261–281. Ballad R. V., Oliver L. L., Downing R. G., and Manuel O. K. (1979) Isotopes of tellurium, xenon and krypton in Allende meteorite retain record of nucleosynthesis. Nature 277, 615– 620. Becker H. and Walker R. J. (2003a) Efficient mixing of the solar nebula from uniform Mo isotopic composition of meteorites. Nature 425, 152–155. Becker H. and Walker R. J. (2003b) In search of extant Tc in the early solar system: 98Ru and 99Ru abundances in iron meteorites and chondrites. Chem. Geol. 196, 43–56. Busso M., Gallino R., and Wasserburg G. J. (2003) Short-lived nuclei in the early solar system: A low mass stellar source? Publ. Astron. Soc. Aust. 20, 356 –370. Cameron A. G. W. and Truran J. W. (1977) The supernova trigger for formation of the solar system. Icarus 30, 447– 461. Canfield D. E. (1998) A new model for Proterozoic ocean chemistry. Nature 396, 450 – 453.

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