Isotopic fractionation of germanium in iron meteorites: Significance for nebular condensation, core formation and impact processes

Earth and Planetary Science Letters 262 (2007) 21 – 36 www.elsevier.com/locate/epsl

Isotopic fractionation of germanium in iron meteorites: Significance for nebular condensation, core formation and impact processes Béatrice Luais ⁎ Centre de Recherches Pétrographiques et Géochimiques (CRPG) CNRS-UPR 2300, 15, rue Notre Dame des Pauvres, BP 20, 54501 Vandoeuvre-lès-Nancy Cedex, France Received 8 February 2007; received in revised form 29 June 2007; accepted 29 June 2007 Available online 12 July 2007 Editor: R.W. Carlson

Abstract The siderophile and volatile nature of germanium allow Ge isotopes to be used to investigate the early history of planetesimals, by tracing processes of core formation and impact through the study of magmatic and non-magmatic iron meteorites. Germanium isotopic compositions were measured using a hexapole-collision cell MC-ICPMS, with an external reproducibility of 0.06‰/amu. Iron meteorites display heavy Ge isotopic compositions (δ74Ge =δ74Ge/70Ge = −0.27 to +1.92‰) with respect to a JMC Ge standard, which contrast with the light isotopic compositions of the studied terrestrial materials (Aldrich Ge standard: δ74Ge = −1.68 − −1.72 ± 0.22‰; sphalerite: δ74Ge = −0.69‰). All the data fall on the theoretical mass fractionation line, indicating no isotopic anomalies. The samples from magmatic irons (IIA–IIB, IIIAB, IIC) have relatively homogeneous Ge isotopic compositions of δ74Ge = 1.77 ± 0.22‰ (2σ) for a large range of Ge contents (36–189 ppm). In contrast, the samples of the non-magmatic groups (IAB, IIE) have lower and more variable Ge isotopic compositions, with δ74Ge = 1.15 ± 0.20‰ for IAB Group (Ge = 254–493 ppm), δ74Ge = −0.27 to +0.43‰ for IIE Old group (Ge = 63–69 ppm), and δ74Ge = +1.40 ± 0.22‰ for the Watson sample (Ge = 50 ppm) from the IIE Young group. Except for IIE samples, no within-group correlation of δ74Ge with Ir or Ge contents is observed, demonstrating that fractional crystallization or crystal segregation do not fractionate Ge isotopes. The lack of inter-group correlation between the Ge isotopic composition of magmatic irons and Ni content, an indicator of redox processes, suggests that if Ge isotopes are fractionated by redox-induced diffusion during metal-silicate segregation, isotopic equilibration towards the initial composition must also occur during isotopic exchange at high temperature. Thus the average δ74Ge value of 1.77 ± 0.22‰ should be representative of their parent body precursors. The lack of correlation between the Ge isotopic composition and Ge/Ni ratios, and the high Ge/Ni ratio of IIA samples (NGe/Ni in CI chondrites) may reflect the high proportion of metal condensates present in the parent body, before metal-silicate differentiation. These metal phases would carry the Ge isotopic signature of a Ge-enriched gas phase of the solar nebula. For the IAB non-magmatic group, the light and homogeneous Ge isotopic composition may result from condensation from a 70 Ge-enriched vapour subsequent to partial melting within a porous parent body. The origin of the heating event, 26Al-decay or impact on the surface of the chondritic body, cannot be distinguished. In contrast, it can be demonstrated that impact events are involved in the formation of IIE irons: the negative δ74Ge–Ge correlation with increasing δ74Ge values from the Old group to the Watson sample of the Young group would result from non-Rayleigh evaporation of 70Ge during impact, with gas–melt back-reaction. The high δ74Ge of the Watson sample (Young group) is consistent with successive impacts on the surface of the IIE parent body. The very light

⁎ Tel.: +33 3 83 59 42 24; fax: +33 3 83 51 17 98. E-mail address: [email protected]. 0012-821X/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2007.06.031

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B. Luais / Earth and Planetary Science Letters 262 (2007) 21–36

Ge isotopic composition of the Old, less impacted IIE irons compared to other studied iron meteorites suggests a light Ge isotopic composition for the IIE parent body. © 2007 Elsevier B.V. All rights reserved. Keywords: germanium isotopes; MC-ICPMS; iron meteorites; core formation; impact; evaporation–condensation

1. Introduction The partitioning behaviour of siderophile elements has been commonly used to study the thermodynamic conditions of formation of iron meteorites. For example, the behaviour of Ge, Ni, in the initial Ge–Ni classification of iron meteorites (Wasson, 1974; Scott and Wasson, 1975), as well as Ir and Au variations, has constrained the origin and processes of formation of magmatic iron meteorites by metal-silicate segregation of undifferentiated precursor material. It is well-established that they represent the crystallized cores of asteroids (Scott and Wasson, 1975), thought nature of these precursors, earlier thought to be chondritic material (e.g. IIA group and enstatite chondrites, Kelly and Larimer, 1977; Cook et al., 2004), has recently become controversial because of the possible contemporary formation of iron meteorites with chondrules and CAIs (Kleine et al., 2005; Markowski et al., 2006; Scherstén et al., 2006). The formation of the nonmagmatic iron meteorites is more controversial. The occurrence of large proportions of chondritic silicate inclusions, which contrasts with their near absence in magmatic irons with the exception of IVA irons (Wasson et al., 2006), led Wasson et al. (1980), Wasson and Wang (1986), Olsen et al. (1994) and recently Wasson and Kallemeyn (2002) to propose mixing between unmelted silicate and impact-induced metallic melt at the nearsurface of a porous chondritic parent body. Petrological studies of various lithologies, on the other hand, favour internal heating, either by metamorphism (Bogard et al., 2000) or by 26Al decay (Benedix et al., 2000) as the source of melting. This would produce silicate and metallic melts which would then be mixed, either by migration of melt at depth or by catastrophic impact and reassembly. In both models, impact occurs at one stage of the formation of nonmagmatic irons. Small variations in Ge concentrations can identify the processes controlling the within-group evolution of metallic liquid as described above, whereas the large variability in Ge concentration of four orders of magnitude (0.01–600 ppm) between magmatic iron groups remains problematic. Germanium is a moderately siderophile element, but is also volatile, with 50% condensation temperature of 825 K at 10− 4 atm and 702 K at 10− 6 atm

(Wai and Wasson, 1979). These authors proposed that the range in Ge contents reflects loss of volatiles in the solar nebula. On the other hand, variations in Ge contents could also reflect Ge metal-silicate partitioning under distinct redox conditions, Ge partitioning in the Fe–Ni core being a direct function of fO2 (Schmitt et al., 1989; Capobianco et al., 1999; Holzheid and Palme, 2001). Ge isotopes can help constrain the origin of elemental fractionation. Kinetic isotopic fractionation is expected because of the following chemical characteristics of Ge. Because of the volatility of Ge, isotopic fractionation can occur by evaporation with preferential loss of the light isotopes in the gas phase at high temperature (Davis et al., 1990; Richter et al., 1999), such as during a high temperature impact event. Elemental partitioning of siderophile Ge in the metal phase during metal-silicate segregation can also be associated with isotopic fractionation, which depends on the relative mass-dependant diffusivities of the isotopes and on the elemental concentration contrasts between the two phases (Richter et al., 1999, 2003). For example, experiments by Roskosz et al. (2006) and Cohen et al. (2006) have demonstrated Fe isotopic fractionation between metal and silicate phases under reducing conditions. In addition to iron meteorites, study of pallasites is of primary importance for understanding processes at the core–mantle boundary. Pallasites, which are composed of olivine–metal associations, are thought to represent the core–mantle interface of planetesimals (Scott and Taylor, 1990). The close relationship of the Pallasite Main Group (PMG) samples with IIIAB irons (Scott and Wasson, 1976) suggest that they are formed by mixing of mantle olivine and a IIIAB-like core, during a highenergy event such as an impact (Wasson and Choi, 2003). These samples can provide information on the resulting core and mantle compositions subsequent to core formation. The present study examines whether germanium isotopes in iron meteorites can help decipher the proposed models. More specifically, this study investigates whether Ge isotopes can be used to: (1) identify distinct parent bodies; (2) trace core formation and provide information on the processes and thermodynamic conditions that prevailed during core segregation; (3) constrain condensation

B. Luais / Earth and Planetary Science Letters 262 (2007) 21–36

and/or evaporation processes at different stages of planetary formation. For this study, we have selected iron meteorite samples from magmatic groups (IIAB, IIC, IIIAB), which represent true planetary cores, as well as from non-magmatic groups (IAB, IIE), for determination of germanium isotopic composition. A pallasite sample was also analyzed for comparison with the IIIAB sample. 2. Analytical techniques Analytical developments of Ge isotopic measurements have been summarized in Luais et al. (2000) and Luais (2003). A detailed description of the technique is given below. As Ge forms volatile species with halogens (e.g. Cl) which could induce loss of Ge and isotopic fractionation, HCl cannot be used during the analytical procedure. Iron meteorite samples were cut in mm size chips with a diamond coated saw to remove 1–2 mm of the potentially oxidized surface corresponding either to terrestrial oxidation or to fusion crust, and then quickly cleaned in acetone. Chips were closely inspected under a binocular microscope to visually verify the absence of sulphide and/or silicate inclusions often present in non-magmatic irons and in some specific magmatic irons (Cape York). The chips were then leached in 2 M HNO3 in an ultrasonic bath for 5 min to remove any contamination during sawing, rapidly rinsed in 18.2 MΩ H2O and in ultra-pure acetone, then dried under a laminar flow hood. Fe and Ni are the main components of iron meteorites, at around 90–95% and 5–10% respectively, Zn being a trace element present in variable concentrations (2 to 25 ppm) (Luck et al., 2005). Ge must be separated perfectly from the matrix to eliminate potential isobaric interference species, such as FeO, FeOH and NiO at masses 72, 73, and 74, 76, respectively, and Zn interferences at mass 70. Whereas 70Zn interferences can be corrected (see below), isobaric interferences of Fe and Ni oxides are more difficult to correct precisely, as they can vary from one session to another depending on focussing parameters. Ge from the analyzed fraction must also be Ga-free, because a Ga isotopic standard is used for the external mass bias correction. Sample chips of 100–200 mg to ensure a representativity of the sample were digested in screw-top Teflon vials using diluted 2 M SeaStar HNO3 on hot plate at 55–60 °C, overnight or more until complete dissolution. After evaporation to dryness, the samples were redissolved in diluted 0.5 M SeaStar HNO3. Depending on the Ge concentration of the samples, 1 to 5 μg of Ge in 0.5 cc of 0.5 M HNO3 are loaded on a Bio-Rad type column filled with 2 cc of AG50x8 cation-exchange resin (200–400 mesh), pre-cleaned with 6 M HCl and H2O until neutral, then preconditioned in 0.5 M HNO3. Given the

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extremely low partition coefficient of Ge for this resin in dilute acid (DeCarlo et al., 1981), Ge passes totally through the resin without adsorption whereas other metal elements are strongly absorbed onto the resin. Ge is collected in a 2 cc fraction of 0.5 M HNO3, free of any matrix elements. Thus, only one ion-exchange column procedure is needed, which results in a full yield. 100% recovery is needed in order to avoid isotopic mass fractionation during column processing. Achievement of full yield is demonstrated by the Ge isotopic composition of the chemistry-processed JMC Ge standard and Fe-meteorite synthetic solution (a synthetic mixture of JMC Ge standard and Fe, Zn, Ni, Co, Ga element standards in the proportions of Fe-meteorites), for which the δ74Ge values are −0.22‰ and −0.15‰ respectively, which are indistinguishable within uncertainty. After Ge elution, matrix elements are removed from the column by rinsing with alternate 6 M HCl and H2O. The Ge fraction is evaporated at 55–60 °C on a hot plate. After complete evaporation, the sample is taken up in 0.01 M SeaStar HNO3 in preparation for MC-ICPMS analysis. Total procedural blank levels for Ge and Ga are better than 30 pg and 40 pg respectively, around 300 pg for Ni, and at the ng level for Fe and Zn (b 6 ng for Fe and b13 ng for Zn). Ge isotopic measurements were performed on a MultiCollector Plasma-Source Mass Spectrometer (Isoprobe GV Instruments, CRPG-Nancy). Standard/sample solutions are introduced in a free-aspiration mode into the MC-ICPMS via a PFA nebuliser at a flow rate of 50 μl/min, and then vapourized into a chilled cyclonic spray chamber and a quartz torch. The specificity of the Isoprobe is the RF-only hexapole collision cell. Whereas Ar in the collision cell reduces the energy spread of ions entering the simple focussing magnetic sector, appropriate H gas flux (0.2– 0.4 ml/min) removes Ar2 interferences on Ge masses (38Ar36Ar on mass 74, 38Ar38Ar and 40Ar36Ar on mass 76), permitting precise measurements of 74Ge. However, 76Ge cannot be analyzed because the ion beam cannot be correctly aligned in the right Faraday collector. External mass fractionation correction of Ge isotopic ratios with the international Ga isotopic standard SRM 994 (69Ga/71Ga= 1.50676; Machlan et al., 1986) is done using an exponential law. The choice of Ga as an external element (Hirata, 1997) is based on the similarity of its masses (m =69, 71) with those of Ge (m =70–76). Isotopic analyses of Ge are carried out by adding 100 ppb of SRM Ga standard to a 1 ppm dilution of the JMC commercial solution of 1000 ppm Ge standard, or to meteorite samples, in 0.01 N HNO3. Care is taken to insure that Ge total beam intensities and Ga/Ge intensity ratios agree within 7% between samples and standards. 72Ge/70Ge, 73Ge/70Ge, 74Ge/70Ge ratios are expressed in δ units with respect to a JMC Ge standard (‰ δxGe = ((xGe/70Ge)sa/(xGe/70Ge)JMC Std − 1) ⁎ 1000)

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Table 1 Germanium isotopic results for terrestrial sphalerite (ZnS sample) and Aldrich Ge standard, and iron meteorites Sample

Class a

Ref number b

Terrestrial materials Aldrich standard (f Ga corrected) f Aldrich standard (simple standard– sample bracketing method) Sphalérite (St Salvy mine, France)

Ni (%)

Ir (ppm)

Ge (ppm)

84 84

δ72Ge (‰)

± 2σ d

δ73Ge (‰)

± 2σ d

δ74Ge (‰)

± 2σ d

−0.88 −0.86

0.16 0.14

−1.35 −1.32

0.20 0.21

−1.72 −1.68

0.26 0.22

−0.58

Refs. Ni, Ir, Ge contents e

453

−0.30

−0.69

0.92 0.94 0.80 0.90 0.90 0.89 0.69

0.14 0.13 0.09 0.03 0.11 0.11 0.02

1.32 1.38 1.16 1.23 1.25 1.27 0.98

0.18 0.11 0.12 0.11 0.14 0.17 0.08

1.83 1.86 1.64 1.71 1.78 1.77 1.35

0.24 0.27 0.17 0.21 0.06 0.18 0.09

[1] [2] [1] [3] [1]

IIA IIA IIA IIA IIA

MNHN 675 BM 1985 M.204 MNHN MNHN MNHN

5 2 5 3 4

5.49 5.46 5.31 5.35 5.49

12.00 3.00 49.00 42.20 12.90

183 189 172 182 178

IIB

MNHN 3053

2

6.1

0.01

107

Wiley Perryville Kumerina Salt River IIC Mean

IIC IIC IIC IIC

BM 1959.914 BM 1959.979 BM 1938.22 BM 1985 M202

2 5 5 1

11.5 9.27 9.69 10.02

6.20 7.18 6.87 6.30

114 88 93.4 100

0.79 0.93 0.87 0.86 0.86

0.20 0.24 0.10 – 0.12

1.14 1.41 1.27 1.27 1.27

0.05 0.28 0.15 – 0.22

1.60 1.92 1.68 1.77 1.74

0.14 0.45 0.18 – 0.28

[1] [1] [1] [1]

Cape York (Agpalilik)

IIIAB

Copenhagen Geol Museum

2

7.58

2.95

37

0.95

0.09

1.41

0.28

1.91

0.39

[6]

Magura Paris #1 g Magura Paris #2 Magura London #3 Mean Magura Cosby's Creek Sarepta Landes Canyon Diablo Odessa Seeläsgen Toluca IAB Mean

IAB MG IAB MG IAB MG

MNHN 986 MNHN 986 BM 33925

3 2 8

6.48

3.65

429

MNHN 84 MNHN 732 MNHN MPI-HEID USNM 2261 USNM 639 BM USNM 6684

6 2 2 5 5 2 6

6.49 6.69 6.58 7.01 7.19 6.65 7.86

2.72 4.38 4.26 2.42 2.38 1.14 2.47

431 457 383 327 283 493 254

0.06 0.06 0.10 0.03 0.02 0.02 0.14 0.04 0.12 0.25 0.10 0.10

0.93 0.95 0.86 0.89 0.77 0.71 0.85 0.90 0.96 0.85 0.79 0.84

0.17 0.22 0.13 0.05 0.06 0.04 0.24 0.18 0.20 0.41 0.12 0.16

1.27 1.26 1.20 1.22 1.06 1.02 1.20 1.24 1.27 1.16 1.05 1.15

0.16 0.14 0.18 0.05 0.05 0.10 0.20 0.19 0.17 0.10 0.12 0.20

[7], [8]

IAB IAB IAB IAB IAB IAB IAB

0.65 0.65 0.63 0.64 0.55 0.52 0.60 0.63 0.65 0.57 0.55 0.59

MNHN n°3447 MNHN BM 83393 BM 19529.196 MPI 1034/3

2 4 2 2 5

8.21 7.55 8.42 7.51 7.96

4.90 6.50 2.10 1.12

49.96 63 64.9 67 68.77

0.71 0.27 0.07 −0.17 −0.10

0.01 0.18 0.09 0.24 0.11

0.96 0.33 0.09 −0.18 −0.20

0.06 0.25 0.05 0.03 0.15

1.40 0.43 0.17 −0.27 −0.22

0.22 0.35 0.02 0.20 0.16

[9] [10] [11] [11] [12]

3

11.16

0.11

60.7

0.94

0.06

1.36

0.16

1.83

0.21

[13]

MG MG MG MG MG MG (IIICD) sLL

Watson MontDieu Arlington Weekeroo Station) Miles

IIE IIE IIE IIE IIE

Brahin

MG Pallasite

Notes to table:

(Young group) (Old Group) (Old Group) (Old Group) (Old Group)

[1], [4], [5]

[7], [7], [7], [7], [7], [7], [7],

[8] [8] [8] [8] [8] [8] [8]

B. Luais / Earth and Planetary Science Letters 262 (2007) 21–36

Iron meteorites Braunau Walker County Scottsville Guadalupe Y Calvo Coahuila IIA Mean Sao Juliao de Moreira

nc

B. Luais / Earth and Planetary Science Letters 262 (2007) 21–36

Fig. 1. Long-term reproducibility of Ge isotopic measurements, and accuracy of mass bias correction for the “secondary” Ge Aldrich standard and the Magura iron meteorite. N = 84 analyses of Aldrich standard and n = 13 analyses of Magura were obtained during 9 sessions from 2001 to 2007. δ74Ge values are calculated with respect to the JMC Ge reference standard (“primary standard”). The two methods of mass bias correction are shown : (1) Solid symbols, heavy lines (δ74Ge average) and dark grey fields (2σ reproducibility) represent data with mass bias correction based that of the added SRM 994 Ga isotopic standard; (2) Open symbols, dotted lines (δ74Ge average) and light grey fields (2σ reproducibility) correspond to δ74Ge values calculated by standard bracketing from measured 74Ge /70Ge ratios. Note that no difference in δ74Ge values within 2σ analytical uncertainties can be identified between the three samples of Magura.

(Xue et al., 1997), in analogy with light stable isotope notation. The standard isotopic values correspond to the mean of the JMC standards analyzed before and after the sample. To test for possible biases in the fractionation correction technique, values obtained using Ga isotopic ratios for this correction are compared to those obtained by simple sample–standard bracketing of measured values. This latter technique implicitly assumes identical mass fractionation between standards and samples measured under the same analytical conditions. The only δxGe values that are retained are those showing a perfect match within error between these two methods of correction (Table 1). Zn interferences on mass 70 are corrected using 68Zn, with a

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mass bias correction applied on the 70Zn/68Zn ratio (70Zn/68Zn = 0.0329; Rosman, 1972). The Zn interference correction remains accurate even for a mixed Ge +Zn solution with a ratio of Zn/Ge= 0.75. Analyses of diluted JMC and Aldrich Ge commercial standards give a total beam intensity of about 10–12 V for 1 μg of Ge. The washing sequence between measurements includes 240 s washout in 0.6 M HNO3, followed by 240 s in 0.01 M HNO3 and 120 s in 0.01 M HNO3 blank solution. After washing, measurements of 0.01 M HNO3 blank solution (180 s on-peak integration time) were taken before each sample and standard measurement for procedural blank correction. Individual measurements of standard/sample correspond to 60 ratios with 10 s integration time, giving an internal precision on both the Ge standards and meteorite samples better than 0.05‰/ amu (2σ). Long-term total external reproducibility determined both on the secondary Aldrich Ge standard, and on the Magura iron meteorite (IAB) (three complete sample treatments including sample dissolution and chemical separation) is generally better than 0.06‰/amu (2σ) for 1 μg Ge (Table 1, Fig. 1). 3. Selected samples Ge isotopic compositions were determined on Femeteorites from different classes according to the initial Ge–Ni classification of Scott and Wasson (1975). They include samples from IIAB, IIC, IIIAB magmatic and IIE, IAB MG and sLL sub-group non-magmatic groups, and cover a large range of within-group Ni contents. Samples from former sub-groups (such as IIA and IIB, IAB and IIICD, IIE Old group and Young group: Watson) and endmembers of groups have been analyzed in order (1) to examine if their Ge isotopic composition matches the revised elemental classification of iron meteorites (Wasson, 1969; Wasson et al., 1989; Olsen et al., 1994; Choi et al., 1995; Wasson and Kallemeyn, 2002), (2) to determine precisely the distinct processes described for some sub-

Notes to table: a Class abbreviations as in Fig. 3. b Reference numbers of samples as: MNHN (Muséum National d'Histoire Naturelle, Paris), BM (Bristish Museum, Natural History Museum, London), USNM (Smithsonian Institution, Washington DC), MPI (Max-Planck-Institute für Chemie). c “n” is the number of replicates of a given sample. Replicate measurements were performed on the same sample dissolution, but distinct ion exchange chemistry, and different MC-ICPMS sessions to ensure a true and long-term external reproducibility, except for Magura samples. d 2σ uncertainty corresponds to 2σ external reproducibility calculated using the Student “t” test (Platzner et al., 1997). e References for Ni, Ir and Ge contents : [1] Wasson (1969); [2] Wasson (1974); [3] Wasson et al. (1998); [4] Malvin et al. (1984); [5] Ryan et al. (1990); [6] Wasson (1999); [7] Choi et al. (1995); [8] Wasson and Kallemeyn (2002); [9] Olsen et al. (1994); [10] Grossman (1997); [11] Wasson and Wang (1986); [12] Wasson (1994); [13] Wasson and Choi (2003). f f Ga correction following the exponential law as xGe/70Ge corr = xGe/70Gemeas * (mx/m70)f; m refers to the molar masses of xGe and 70Ge isotopes; f is the mass bias correction calculated using SRM 994 international Ga isotopic standard. g Magura isotopic measurements were performed on three distinct hand samples from the Museums of Paris (#1 and #2) and London (#3). δ74Ge values for these three samples agree within 2σ reproducibility, attesting the robustness of the analytical procedure.

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B. Luais / Earth and Planetary Science Letters 262 (2007) 21–36

Fig. 2. δ72Ge vs δ74Ge ‰ for iron meteorite samples, secondary Aldrich standard and a terrestrial sphalerite (ZnS, St Salvy Mine, France). Error bars are 2σ external reproducibility, based on replicate samples. All samples fall on the theoretical mass fractionation line (TMFL), for which the equation is: δ72Ge = (m72 − m70)/(m74 − m70)δ74Ge, where m refers to molar mass of the considered isotope. Magmatic group samples (IIAB, IIIAB, IIC) exhibit heavier isotopic compositions than non-magmatic samples (IAB, IIICD, IIE). Error bars for each sample are 2σ external reproducibility on sample replicates. The same results are obtained in a δ73Ge vs δ74Ge ‰ diagram (not shown here).

groups (e.g. IIE Old group and Young group; Wasson and Wang, 1986). In detail, the studied samples of the large IAB group cover the main range of Ni and Ge contents with Ge N 200 ppm and Ni b 9%, this range representing more than 84% of the IAB samples. These samples mainly belong to the IAB Main Group, even for the sample Seeläsgen formerly identified as IIICD, and to the IAB sLL sub-group (s for sub-group, L for low Au, and L for low Ni) for the Toluca sample (Wasson and Kallemeyn, 2002). The IIE irons define continuous trends on some trace element diagrams (Wasson and Wang, 1986; Olsen et al., 1994), and can be justified as a single group. Isotopic data however reveal more complex distinctions: Niemeyer (1980) initially subdivided the IIEs into two Ar–Ar age sub-groups, later confirmed by additional radiometric and exposure ages summarized in Snyder et al. (2001): a young (3.7± 0.2 Ga) group with 5–10 Ma cosmic-ray exposure (CRE) age, an old (4.5 ± 0.1 Ga) group with N 50 Ma CRE age (Bogard et al., 2000; Olsen et al., 1994). They also have distinct nitrogen isotopic signatures, with δ15N of −7.1‰ for the old group and −2.3‰ for the young group (Mathew et al., 2000). Ge isotopic measurements are given for a sample of Watson (Young group), and samples of Weekeroo Station, Miles, Arlington (Old group). A MontDieu sample (Grossman, 1997), for which no radiometric data are yet available, has been also analyzed.

The metal phase of the Brahin pallasite (Main Group Pallasite) have been analyzed for comparison with the IIIAB sample. In addition to meteorite samples, a terrestrial ZnS ore (sphalerite from St Salvy Mine, Massif Central, France) has also been analyzed because of its high Ge content (453 ppm of Ge) in the range of the IAB iron meteorites. 4. Results Germanium isotopic variations, expressed in δxGe units (‰ δxGe =δxGe/70Ge) for whole rock iron meteorite samples, are presented and discussed in a manner analogous to that normally used for light stable isotopes (e.g.; δ17O vs δ18O) (Table 1, Figs. 2 and 3). In Fig. 2, data are reported in a δ72Ge vs δ74Ge diagram. All iron meteorites fall within error of the theoretical mass fractionation line (TMFL), on which also plot a few terrestrial samples including the two Ge standards of distinct compositions, and the sphalerite sample (ZnS ore). This means that: (1) all isobaric interferences have been efficiently removed; (2) there are no resolvable nucleosynthetic isotopic anomalies at the delta unit scale, thus indicating a common isotopic source. Ge isotopic ratios of the studied iron meteorites span a total range of 0.55‰ per amu, and these variations are thus ascribed to mass fractionation processes. The iron meteorites are enriched in

B. Luais / Earth and Planetary Science Letters 262 (2007) 21–36

Fig. 3. Detail of δ74Ge distribution in iron meteorites with respect to their classification. Solid lines and areas for each group (except IIE) indicate the mean and range (2σ) of δ74Ge values. Magmatic and non-magmatic groups hosting large sulphide nodules such as Cape York (IIIAB) and Toluca, Odessa and Canyon Diablo (IAB) are underlined. Notations: MG Pall for Main Group pallasite, IAB MG for main group of the IAB group, and IAB sLL for sLL sub-group of the IAB group (Wasson and Kallemeyn, 2002). Error bars as in Fig. 2. Same symbols as in Fig. 2.

heavy isotopes (δ74Ge = −0.27 to +1.92‰) in contrast with the depletion of terrestrial samples (up to −1.72‰ for δ74Ge in sphalerite). The studied groups display resolvable, distinct Ge isotopic compositions (Fig. 3), with the non-magmatic groups IIE and IAB having lower delta values than the magmatic groups. In detail, group IIE has the lowest δ74Ge values and also the larger spread, ranging from −0.27 to +1.40‰, while the IAB group has δ74Ge from +1.02 to +1.27‰. The IIAB, IIC, and IIIAB magmatic groups have a similar range of δ74Ge values, from +1.35 to +1.92‰. There is no correlation between the isotopic ratio and the kamacite/taenite ratio of each sample, i.e. no difference in δ74Ge values between hexahedrites (IIA) and octahedrites (IIC, IIIAB). Ge isotopic compositions of sub-groups are investigated (Fig. 3). For the IIE samples, the Watson sample from the Young group has high δ74Ge (δ74Ge= +1.40‰) compared to the Old group (δ74Ge = −0.27 to +0.43‰). Sao Juliao de

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Moreira sample from the IIB sub-group has significantly lower values (δ74Ge = 1.35‰) than the average IIA group that is homogeneous (δ74Ge =1.77 ±0.18‰). Unlike the IIA samples, Sao Juliao de Moreira contains an unusually large amount of trapped melt (70%) (Wasson et al., 2007), which can play a role in the Ge isotopic budget. The former IAB Group and IIICD Seeläsgen sample cannot be distinguished (δ74Ge =1.02 to 1.27‰; and 1.16‰; respectively), confirming the recent classification of Wasson and Kallemeyn (2002), which assigns all of the studied samples, except Toluca, to the IAB Main Group. Toluca from the IAB sLL sub-group has δ74Ge values of 1.05 ± 0.13‰, identical to those of the Main Group. This is however consistent with appreciable overlap of trace element contents between of IAB sLL and the IAB Main Group. IAB Main Group, IAB sLL and Seeläsgen samples will therefore be treated as a single group and discussed under the IAB label further in this study. Some of these features are in agreement with the Ge isotopic results of Hirata (1997) in that his only analyzed IIB magmatic iron (Sikhote-Alin) has slightly heavier Ge isotopic compositions than his IAB non-magmatic irons, but at the limit of analytical error. However, some discrepancies exist between this study and his data for IAB samples. Taking into account the larger analytical error (up to 0.6–1‰ for some samples), the Hirata (1997) data show distinct Ge isotopic values for Odessa/Landes and Canyon Diablo IAB samples whereas this study reports constant Ge isotopic ratios for these IAB samples. Such discrepancies could potentially be explained by (1) sample inhomogeneity, such as sulphide inclusions in Canyon Diablo, or more likely (2) isotopic fractionation during sample dissolution and Ge purification. Hirata's method includes the use of HCl medium and a CCl4 organic phase. As noted in the analytical section, Ge forms a volatile GeCl4 complex with HCl. Possible incomplete back-extraction of Ge with pure water, leading to traces of CCl4 in the Ge fraction, which is later heated at 55 °C, could result in isotopic fractionation. For magmatic and non-magmatic groups hosting large sulphide nodules, such as Cape York (IIIAB), and Toluca, Odessa and Canyon Diablo (IAB), the Ge isotopic ratios of the pure metal phase analyzed here do not differ from those of sulphide-free irons of their respective groups. We could expect some kinetic isotopic fractionation of Ge associated with mass transport between sulphide and metal phases, because Ge is more siderophile than chalcophile. Nielsen et al. (2006) and Williams et al. (2006) also present for Tl and Fe isotopes respectively in iron meteorites, the lack of measurable isotopic fractionation in the metal phase whereas the sulphide phase is isotopically fractionated. Modelling by Nielsen et al. (2006) demonstrates the

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Fig. 4. δ74Ge vs Ni contents for magmatic irons (IIAB, IIC, IIIAB). Data for non-magmatic irons (IAB, IIE) are indicated for comparison. Ni is an indicator of redox processes. No correlations between the groups can be observed. References for Ni contents in Table 1. Error bars and symbols as in Fig. 2, and notations as in Fig. 3.

dependence of the diffusion-controlled isotopic fractionation on the relative modal abundances of the two phases (high metal/sulphide ratio). Finally, the metal phase of the Brahin pallasite has a δ74Ge of 1.83‰, which agrees within error with that of the Cape York IIIAB sample (δ74Ge = 1.92‰) (Fig. 3). 5. Discussion 5.1. Core formation and nebular condensation: oxidoreduction versus condensation processes for magmatic irons Variations of Ge isotopic compositions of magmatic iron samples are small compared with those of IIE nonmagmatic irons (Fig. 3), and there is no marked difference between magmatic groups. If the magmatic irons represent planetesimal cores, as is generally believed, they would have recorded fingerprints of processes linked to metalsilicate segregation, which occurs under reduced conditions during heating and melting of undifferentiated chondritic bodies during accretion. However metal-silicate segregation is not the only process that produced the metal found in planetesimal cores. Partially molten metal droplets, that agglomerated to form large diapirs gravitationally descending to the growing core, have several possible origins. They can originate in the chondrite body as metal grains, either (1) as direct condensates from the solar nebula for zoned grains

(Meibom et al., 1999) or (2) from processes related to chondrule formation such as evaporation/recondensation of vapour (Kong et al., 1999; Yu et al., 2003; Zanda, 2004) or (3) from reduction of FeO within chondrule melt (Connolly et al., 2001; Cohen and Hewins, 2004). They also can be formed (4) by reduction of Fe-silicates of the chondritic parent body (Leroux et al., 2003), or (5) by “mantle self-oxidation”, a mechanism in which perovskite precipitated at depth reacts with Al-peridotite, leaving an oxidized lower mantle and iron droplets (Wade and Wood, 2005). In the following sections we examine the effect of reduction and condensation processes on Ge isotopic compositions. 5.1.1. Reduction processes The potential effect of reduction on Ge isotopes is evaluated in a δ74Ge vs Ni diagram (Fig. 4). During metalsilicate equilibration, reduction results in a progressive decrease in Ni contents in the metal phase because of preferential enrichment of Fe in the metal, making Ni (or Ni/Fe) a potential tracer of redox process. Different degrees of oxidation among the IIA, IIB, IIC, IIIAB studied groups can be defined based on their Ni contents and the occurrence of oxygen-bearing phosphates. IIA samples with the lowest Ni contents and no phosphates reflect strongly reduced conditions while the IIIAB samples with higher Ni contents and the presence of phosphate minerals provide evidence of oxidized conditions (Olsen et al., 1999). IIC samples with the highest Ni

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Fig. 5. δ74Ge vs Ge/Ni ratios for magmatic irons (IIAB, IIC, IIIAB). Data for non-magmatic irons (IAB, IIE) are indicated for comparison. During condensation of a cosmic gas (Kelly and Larimer, 1977), Ni is enriched in the first condensates and volatile elements are depleted. With increasing condensation, the condensates are depleted in Ni whereas concentrations of volatiles increase. At each stage, redox processes alter the formed metal, resulting in an increase in Ni because of Fe oxidation and loss from the metal, and a small increase in volatile element content. The final stage of condensation is characterized by the lowest Ni and highest volatile contents, which are present in CI proportions. This leads to a Ge/Ni ratio (CI)=0.00295 (Anders and Grevesse, 1989) (dashed line). Low Ge/Ni ratios represent early condensates, and high Ge/Ni ratios late condensates. The crucial point here is the high Ge/Ni ratios of IIA samples. Error bars and symbols as in Fig. 2, and notations as in Fig. 3. References for Ge and Ni contents in Table 1.

contents are the most oxidized group among the studied samples. However variations in Ni contents from 11.5 ppm in IIC samples to 5.5 ppm in IIA samples are not correlated with Ge isotopic values. Small variations in germanium isotopes for all of the investigated magmatic groups can be quantified with a mean and standard deviation in δ74Ge of 1.77 ± 0.22 (2σ). This standard deviation is nevertheless similar to the analytical 0.06‰/amu (2σ) uncertainty. Two alternative interpretations can be proposed. The first is that isotopic fractionation during core segregation is small, approaching the limit of analytical error. It can be considered as negligible. The second is that reduction during metal segregation from an undifferentiated parent body can indeed fractionate Ge isotopes, but these effects are largely erased at high temperature by isotopic equilibration with time, as has been experimentally demonstrated for Fe isotopes (Cohen et al., 2006; Roskosz et al., 2006). Two concomitant processes, isotopic diffusion and isotopic exchange occur (Cohen et al., 2006). During reduction, chemical and isotopic diffusion is dominant and results in preferential migration of the lighter isotopes of Ge in the newly formed Fe–Ni metal, because light isotopes have higher diffusion rates than heavy isotopes (Richter et al., 1999). This massdependant diffusion leads to kinetic isotopic fractionation,

and would result in low δ74Ge in the metal phase with respect to the silicate phase. With time and at constant f O2, when little metal reduction can occur, isotopic exchange of Ge between metal (Ge-rich) and silicate (Ge-poor) phases becomes the dominant process. This process will tend toward equilibrium isotopic fractionation leading to isotopic compositions evolving toward the initial value of the parent body. The 2σ standard deviation of 0.22‰ can reflect either the state of isotopic equilibrium, consistent with the value of 0.20 ± 0.15‰/amu for Fe isotopes (Roskosz et al., 2006) or the slight range in δ74Ge of the parent bodies if complete equilibration is achieved. Large variations in Ge contents associated with small isotopic variability suggest that Ge isotopic re-equilibration is faster than elemental re-equilibration. The restricted range in δ74Ge thus probably represents the limited Ge isotopic variation of the parent bodies. Possible lack of metal-silicate isotopic fractionation, or Ge isotopic equilibration after core formation is further demonstrated by the similarity in Ge isotopic composition between the Brahin pallasite metal phase and the Cape York IIIAB sample. If the metal phase of the pallasite represents the IIIAB core–mantle interface (Wasson and Choi, 2003), its isotopic composition has not been changed by interactions, such as diffusion, with silicate mantle.

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Fig. 6. (a) δ74Ge vs Ir contents, (b) δ74Ge vs Ge contents, for magmatic (IIAB, IIC, IIIAB) and non-magmatic (IAB, IIE) irons. Ir is an indicator of fractional crystallization, and of mixing of liquid and solid metallic phases. There are no variations in Ir and Ge contents with δ74Ge values, except for IIE groups. References for Ir and Ge data as in Table 1. Error bars and symbols as in Fig. 2, and notations as in Fig. 3.

5.1.2. Condensation processes Kelly and Larimer (1977) have modelled the two processes of condensation and oxidation–reduction of Fe– Ni metal from a cosmic gas by comparing concentration variations of Ni, which is less volatile than Fe, with those of either more refractory or more volatile elements, such as Ga (and Ge). As Ni is more sensitive than Ge to late redox conditions, we chose to investigate the co-variation of δ74Ge and Ge/Ni ratio for the studied groups (Fig. 5), the latter as an indicator of nebular condensation at a given

oxidation and fractionation state. The IIIAB and IIC samples with low Ge/Ni ratios represent the initial stages of condensation, thus providing the evidence that the range of δ74Ge values of 1.77 ± 0.22‰ is established early. This isotopic signature is not modified in the later stage as seen in the IIA samples with high Ge/Ni ratios (Fig. 5). Samples with low Ni content (high Ge/Ni), such as the hexaedrite IIA samples, are preferred for studying primitive processes related to the origin of metal, as samples with high Ni content (low Ge/Ni) may have experienced secondary

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modifications due to metamorphism or redox processes (Meibom et al., 1999). The key point is that IIA samples have Ge/Ni ratios even higher than that of primitive material (Ge/Ni CI = 0.00295; Anders and Grevesse, 1989) but nevertheless have Ge isotopic ratios similar to those of the other irons. Even if the parent body had variable concentrations in Ge, this cannot explain these values higher than that of the CI primordial material. The occurrence of high Ge/Ni values in meteorites could indicate that CI composition does not represent primitive material; the latter would be less depleted in volatiles. However, as CI have solar Ge concentrations (Ringwood, 1979), the high Ge/Ni more likely reflects enrichment in Ge in the IIA precursor. The lack of significant isotopic fractionation with large Ge/Ni variations suggests that Ge has been incorporated by condensation processes, but at high levels in the IIA meteorites producing an anomalous enrichment. Higher quantities of metal condensates in the parent body of IIA samples than in others are necessary to explain their high Ge/Ni ratios. As Ge is moderately volatile, the Ge carrier phase may be a gas component, which would favour the transport of Ge in the form of the volatile species GeO or GeS. Early or late Ge component condensates have similar Ge isotopic compositions, implying that metal condensates in the parent body impose the Ge isotopic composition rather than metal produced by reduction of silicates. This can also suggest a unique Ge isotopic composition of the metal phase (primitive metal condensates), and thus of the parent body during accretion prior to core formation. The homogeneous Ge isotopic signature of magmatic irons is established during early processes and can result from large-scale isotopic homogenization. As stated above, primitive metal condensates are present in chondrule material. There is thus a genetic link between magmatic iron meteorites and chondrules. This link is also time-related as Hf–W studies have revealed early formation of iron meteorites, coeval with the first chondrules and possibly with CAIs (Kleine et al., 2005; Markowski et al., 2006; Scherstén et al., 2006). There are also certain geochemical and isotopic similarities between iron meteorites and chondrules notably the lack of, or limited isotopic fractionation of moderately volatile elements such as Si, K (Alexander et al., 2000; Clayton et al., 1991; Cohen et al., 2004; Humayun and Clayton, 1995; Yu et al., 2003), or Fe (Mullane et al., 2005) in chondrules. Isotopic uniformity in magmatic iron meteorites has also been reported for Fe isotopes, for which a fairly large number of isotopic data are available (Poitrasson et al., 2005; Weyer et al., 2005). These authors also argue for a homogeneous inner solar system, with homo-

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genization occurring perhaps even before chondrule formation. 5.1.3. No variations in δ74Ge values during the evolution of metallic liquids by fractional crystallization Once the metal liquid has been segregated, cooling produces crystallization. Wasson (1974) and Scott and Wasson (1975) have examined pure fractional crystallization of these metallic liquids using Ir content variations, Ir being strongly compatible in the solid metal. In a δ74Ge vs Ir content plot (Fig. 6a), there is no within-group correlation for the IIA and IIC groups, which shows that fractional crystallization of a metallic core does not fractionate Ge isotopes. This is strongly demonstrated in the IIA group which has large variations in Ir contents (3–49 ppm). The exception is our sample of Sao Juliao de Moreira belonging to the IIB group, which displays very low Ir contents (0.012 ppm) and a slightly lower δ74Ge value than the IIA irons, despite the fact that the IIA and IIB irons are thought to derive from a common parent body (Wasson, 1969; Cook et al., 2004). A recent study of the IIAB group (Wasson et al., 2007) provides a detailed thin-section description of Sao Juliao de Moreira which shows that this sample displays well-distinguished rounded troilite (FeS) and skeletal schreibersite (FeNi, P) crystals. These sulphide and phosphide inclusions may have been formed during trapping of a large amount (70%) of enriched (P, S) residual metallic liquid after 48% fractional crystallization (Wasson et al., 2007), these non-metallic elements being almost insoluble in metal (very low DS, P solid/liquid; Wasson and Richardson, 2001). It is therefore possible that the analyzed sample is an intimate mixture of metal and schreibersite, as troilite nodules are easier to avoid during sample preparation. The low δ74Ge value would then result from the low δ74Ge of FeNi phosphate phases. Further data are required to resolve this issue. 5.2. Impact processes for non-magmatic irons? The IAB and IIE non-magmatic samples have lower Ge delta values than magmatic irons, with the notable exception of the IIE Watson meteorite which overlaps with the magmatic irons. The IAB and IIE groups display different δ74Ge vs (Ir,Ge) trends (Fig. 6a and b), reflecting the distinct behaviour of refractory non-volatile Ir and volatile Ge during processes involved in the formation of non-magmatic irons. 5.2.1. The IAB group The IAB group trend resembles those of the magmatic irons. No δ74Ge–Ir or δ74Ge–Ge correlations exist,

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Fig. 7. Modelling the δ74Ge vs Ge variations for evaporation processes with comparison to IIE samples. The low δ74Ge sample (Miles) is taken as the initial composition. Solid curves (1, 2, 3) represent pure Rayleigh fractionation for evaporation of Ge as Ge metal, GeO and GeS, respectively, with gas–melt isotopic fractionation factors α of 0.97258, 0.97752 and 0.980945, respectively (α is defined as (m70/m74) β, β=0.5). Dotted curves (1, 2, 3) are defined with α of 0.99079, 0.99410 and 0.99501 for Ge, GeO and GeS species, respectively, corresponding to a β value of 0.13. Curves (not shown) for a β value of 0.1 (Tsuchiyama et al., 1994) would fit the data for GeO and GeS species only. Error bars as in Fig. 2.

because of the constant δ74Ge (δ74Ge = 1.15± 0.20‰), coupled with a very limited Ir range (2–4 ppm) and large Ge range (180–560 ppm) compared to that of individual magmatic groups (Fig. 6a and b). This suggests that the constant Ge isotopic ratios may result from magmatic processes. However these processes must be distinct from pure fractional crystallization to account for the contrasting range in Ir and Ge concentrations, and the small range of Ir (Wasson, 1974; Scott and Wasson, 1975; Choi et al., 1995). According to Wasson and Kallemeyn (2002), the combined effect of a limited degree of fractional crystallization of a metallic melt, followed by separation of crystals and melt by density within a porous chondritic parent body with similar volumes of solid and liquid metal, would produce the observed range in elemental variations. Wasson and Kallemeyn (2002) modelled elemental variations in terms of equilibrium mixing between solid and liquid metal phases in various proportions. High Au– low Ge samples correspond to a liquid-enriched pole whereas low Au–high Ge samples correspond to a mixture of 30% melt and 70% solid. The important point is that because of the negative Ge–Ni correlation reflecting partitioning for Ge and Ni between liquid metal and solid phases during crystallization (Wasson and Kallemeyn, 2002), this process would produce very high Ge/Ni ratios in the solid-enriched pole (Fig. 5), strongly above the primitive value (Ge/Ni CI = 0.00295; Anders and Grevesse, 1989). In Fig. 6b, it can then be inferred that as various proportions of melt and solid do not change the Ge delta values, melt and solid must have the same Ge isotopic

composition. Crystal segregation would not induce Ge isotopic fractionation, as concluded for pure fractional crystallization in magmatic irons. Benedix et al. (2000) and Takeda et al. (2000) propose distinct petrological models based on the comparison of recrystallized textures in silicate inclusions of IAB irons and in the silicates of genetically related winonaite achondrites, that they attribute to partial melting of chondritic precursors. In this scenario, IAB irons represent Fe–Ni–S cotectic melts. Benedix et al. (2000) suggest that catastrophic impact would then mix silicates and metallic melt together, while they are still molten. Subsequent slow cooling of large volumes of metallic melts, as indicated by the cooling rate in the taenite phase (Herpfer et al., 1994) would induce fractional crystallization. Takeda et al. (2000) propose that coexistence of silicate and metal result from inhomogeneous segregation of partial melts. These authors allow the possibility of a small impact event after crystallization to account for shock textures in silicate inclusions. Reconciling elemental and isotopic variations of Ge with petrological interpretations is somewhat difficult. First, magmatic processes, which can be partial melting and/or crystal segregation, would not produce Ge isotopic fractionation: the constant δ74Ge values of 1.15 ± 0.20‰ over a large range of Ge concentrations would be that of the initial metallic melt. All the studied IAB samples would therefore have been formed by similar magmatic processes affecting identical precursor material. Secondly, it is necessary to understand the low delta Ge values of IAB samples compared to magmatic samples. Non-magmatic processes, such as impacts, have been proposed by Wasson et al. (1980), Benedix et al. (2000) and Takeda et al. (2000), as contributing factors in the evolution of these IAB irons, but there is disagreement concerning the timing of these processes. According to Wasson and Kallemeyn (2002), impact is the major process involved in the initial melting event. According to Benedix et al. (2000), impact is involved in the intermediate stage permitting the reassembly of silicate and metallic melt. On the other hand, Takeda et al. (2000) suggest that the shock event occurred after the crystallization process. They all note petrographic evidence of minor shock events, such as very weak to weak shock effects (S1–S2 stages, Stoffler et al., 1991) in silicate-rich inclusions (Benedix et al., 2000), and in gabbroic inclusions (Takeda et al., 2000). Such mild shock events would not be able to induce melting, unless the impacted material is porous (Stoffler et al., 1991). However if impact is the process producing initial melting in the IAB irons, why is there no firm Ge isotopic evidence of this process? During impact, volatile

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elements such as Ge would be partly vapourized, inducing loss of Ge and isotopic fractionation with loss of light isotopes in the residue. No such correlation is seen in Fig. 6b. The Ge element-isotope systematics are similar to those of the volatile element nitrogen for IAB samples, e.g. relatively constant N isotopic ratios are associated with a large range in N contents (Franchi et al., 1993). The isotopically light Ge composition (δ74Ge = 1.15 ± 0.20‰) of these metallic melts can have two explanations: 1. Enrichment in 70Ge related to the condensation of a vapour phase. This scenario is similar to that described by Wombacher et al. (2003) for ordinary chondrites. During melting at high temperature producing metallic pools within the parent body, selective volatilization of Ge and other volatile elements can occur in the hot portion of the parent body. Transport of melt and 70Gerich vapour phase along cracks and fissures would be followed by melt crystallization and vapour condensation in pores of the cooler regions. These would thus be preferentially enriched in volatile elements and would have a light Ge isotopic composition. The latter stage implies that the parent body is highly porous. The constant Ge isotopic ratio implies a single stage-melting event. This model provides a framework that could explain the formation of IAB irons. The present Ge isotopic data cannot distinguish the origin of the melting event: impact (Wasson and Kallemeyn, 2002) or 26Al radioactive decay and thermal metamorphism up to 1200 °C (Benedix et al., 2000; Takeda et al., 2000). 2. The measured light Ge isotopic composition reflects that of the precursor material. This scenario implies that internal heating and mild shock events did not significantly fractionate the isotopes of the volatile germanium. 5.2.2. The IIE irons The IIE group shows a negative δ74Ge–Ge correlation with large variations in δ74Ge (Fig. 6b). Such a trend is expected for evaporation, with preferential loss of volatile Ge and light Ge isotopes. By comparison, there is no loss of refractory Ir as shown by the positive δ74Ge–Ir correlation (Fig. 6a). This implies that the increase in δ74Ge reflects the progressively heavier Ge isotopic composition of the residue as evaporation proceeds, with each IIE sample representing a residue of evaporation. In contrast with IAB non-magmatic irons, strong mineralogical and petrological evidence of impact have been described in metal and silicate phases of IIE irons, suggesting a link between the evaporation process and impact events. Shock stage S4 textures are described in the Old group (Bogard et al., 2000), and shock stage S3 textures in the Young Group

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(Olsen et al., 1994; Bogard et al., 2000). The distinct high Ge isotopic value of the Watson sample (δ74Ge = +1.40‰) of the Young Group, compared to the low Ge isotopic value of the Old group (δ74Ge = −0.27 to +0.43‰) would reflect the higher degree of Ge isotopic fractionation through evaporation. The well-defined δ74Ge–Ge negative correlation (Fig. 6b) supports the origin of the Young and Old groups from a single parent body, contrary to the suggestion of Mathew et al. (2000) who propose two distinct reservoirs for these two IIE sub-groups. The textural relationship between metal and melted silicates in the Young group samples are interpreted as successive impact heating events on the IIE parent body (Olsen et al., 1994), as supported by radiometric data (Bogard et al., 2000). The high δ74Ge values for the Watson sample can be explained as the result of progressive heavier isotopic compositions of the residue by Ge evaporation due to repeated impacts subsequent to those producing the Old group. The evaporation process for IIE samples is modelled in Fig. 7, with the initial isotopic composition represented by the Miles sample, which has the least fractionated Ge isotopic composition and the highest Ge content within the IIE group. It thus represents the residue that has been least re-processed by impacts. The large isotopic fractionation (range of δ74Ge values from −0.27‰ to +1.40‰) precludes volatility-controlled isotopic fractionation under equilibrium fractionation at high temperature (Richter et al., 2003). Modelling of kinetic isotopic fractionation shows that the IIE samples lie below the ideal kinetic Rayleigh fractionation curves for the evaporation of germanium as Ge metal or as GeO or GeS species (Fig. 7), which are the most stable gaseous compounds of Ge (Kelly and Larimer, 1977; Wai and Wasson, 1979). This implies that Ge isotopic mass fractionation is much lower than predicted for pure kinetic Rayleigh fractionation. A good fit for all IIE samples (Young and Old groups) can be achieved only by using higher values of the gas–melt isotopic fractionation factor α (α =0.9941–0.9950 for GeO and GeS species, respectively), corresponding to a β value of 0.13 (see explanations in Fig. 7 caption). Such lower β values are similar to those estimated by (Tsuchiyama et al., 1994) (β = 0.1) for diffusion between condensed phases. This β value of 0.13 implies additional gas–solid exchange during evaporation, such as might occur during back-reaction (partial re-condensation) between the residue and the ambient gas (Ozawa and Nagahara, 2001), partially suppressing the large isotopic fractionation expected for evaporation (Cohen et al., 2004). Deciphering the origin of the low initial Ge isotopic composition is somewhat speculative at present. We can perhaps infer an isotopically light precursor for IIE

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irons. However, it is premature to propose a light Ge isotopic composition for the potential H chondritic parent body, as its relationship with IIE irons is established only for IIE silicate inclusions on the basis of O isotopic data (Wasson and Wang, 1986; Olsen et al., 1994). An alternative explanation of the Ge isotopic systematic of the non-magmatic irons is that their low δ74Ge compositions may be influenced by silicate phases present as unmelted inclusions of olivine, plagioclase and pyroxene. In particular, the IIE trend could represent a mixing trend between a high δ74Ge Fe–Ni component and a low δ74Ge silicate pole. However, according to the negative δ74Ge–Ge correlation (Fig. 6b), this would imply a silicate phase with higher Ge contents than the metal phase, which is completely inconsistent with the siderophile character of Ge. Similarly, the low and constant δ74Ge compositions of IAB samples preclude mixing with silicates. 6. Conclusions Germanium isotopic compositions of iron meteorites and a terrestrial sphalerite have been measured by using a hexapole collision cell MC-ICPMS technique with a 2σ reproducibility of 0.06‰/amu. The Ge isotopic fractionation of iron meteorites follows the theoretical mass fractionation law, as do the terrestrial samples. There are no nucleosynthetic isotopic anomalies. Iron meteorites are characterized by heavier isotopic compositions than terrestrial samples. Within the analytical reproducibility, the germanium isotopic composition of magmatic and nonmagmatic irons can be distinguished, with higher δ74Ge values for magmatic irons (1.77 ± 0.22‰) than for nonmagmatic irons (−0.27 to 1.40‰). The lack of correlation between δ74Ge values and Ir contents for both magmatic and IAB non-magmatic irons shows that Ge isotopes are not fractionated during crystallization, partial melting and crystal segregation. Magmatic groups that are representative of planetesimal cores (IIA, IIIAB, IIC) have homogeneous δ74Ge of 1.77 ± 0.22‰, for variable Ge contents (36–189 ppm). If core formation (metal-silicate segregation) induces isotopic fractionation by diffusion between silicate and metal phases, then isotopic exchange over time at high temperature towards the isotopic composition of the parent bodies must have occurred. In this case, the 2σ deviation would represent the restricted range of Ge isotopic composition of the undifferentiated precursors. Group IIA samples, which have superchondritic Ge/Ni ratios, have similar Ge isotopic compositions to low Ge/Ni groups (IIC, IIIAB), demonstrating that Ge isotopic compositions of irons can also be controlled by metal condensates present in the parent body,

before metal-silicate segregation. This strongly suggests that the homogeneous Ge isotopic composition of magmatic irons was established early in the history of the solar nebula, at high temperature encouraging isotopic equilibration. The low and homogeneous isotopic compositions of IAB samples (δ74Ge = 1.15 ± 0.20‰) can be explained by limited vapourization, transport and condensation of the volatile Ge subsequent to partial melting within a porous parent body. The negative δ74Ge–Ge correlation of the IIE samples (δ74Ge = − 0.27 to + 0.43‰ for the Old group; δ74Ge = +1.40‰ for the Watson sample from the Young group) can be produced by evaporation during impact events. Evaporation processes, unlike pure Rayleigh fractionation, may reflect gas–melt exchange in which back reaction during recondensation partially suppresses an extensive isotopic fractionation. The high δ74Ge value of the Watson sample is explained by successive impact events, as is supported by its young formation age. Acknowledgments An early version of the manuscript benefitted from constructive comments of J. Wasson and M. Chaussidon. Thanks to the MNHN Paris, NHM London, MPI Mainz, Smithsonian Institution of Washington, Geological Institute of Copenhagen University, for meteorite sample donation, to F. Kuntz for donation of the Brahin pallasite sample, to O. Legendre (BRGM Orléans) for the sphalerite sample from St-Salvy mine (France). L. Reisberg carefully edited the English. C. Fournier is thanked for management of the clean lab and MCICPMS laboratory. Funding from CNRS programs (Interieur Terre, DyeTI, and PNP) are acknowledged. The editor R. Carlson and two anonymous reviewers are strongly acknowledged for considerable constructive comments. This is CRPG contribution no. 1863. References Alexander, C.M.O.D., Grossman, J.N., Wang, J., Zanda, B., BourotDenise, M., Hewins, R.H., 2000. The lack of potassium isotopic fractionation in Bishunpur chondrules. Meteorit. Planet. Sci. 35, 859–868. Anders, E., Grevesse, N., 1989. Abundances of the elements: meteoritic and solar. Geochim. Cosmochim. Acta 53, 197–214. Benedix, G.K., McCoy, T.J., Keil, L., Love, S.G., 2000. A petrological study of the IAB iron meteorites: constraints on the formation of the IAB-Winonaite parent body. Meteorit. Planet. Sci. 35, 1127–1141. Bogard, D.D., Garrison, D.H., McCoy, T.J., 2000. Chronology and petrology of silicates from IIE iron meteorites: evidence of a complex parent body evolution. Geochim. Cosmochim. Acta 64, 2133–2154.

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