A nebula setting as the origin for bulk chondrule Fe isotope variations in CV chondrites

Earth and Planetary Science Letters 296 (2010) 423–433

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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l

A nebula setting as the origin for bulk chondrule Fe isotope variations in CV chondrites Dominik C. Hezel a,⁎, Andrew W. Needham b, Ros Armytage b, Bastian Georg b, Richard L. Abel a, Erika Kurahashi a, Barry J. Coles c, Mark Rehkämper c, Sara S. Russell a a b c

IARC, Department of Mineralogy, Natural History Museum, Cromwell Road, London, SW7 5BD, UK University of Oxford, Department of Earth Sciences, Parks Road, OX1 3PR, Oxford, UK Department of Earth Science and Engineering, Imperial College, London, SW7 2AZ, UK

a r t i c l e

i n f o

Article history: Received 29 January 2010 Received in revised form 19 May 2010 Accepted 20 May 2010 Available online 19 June 2010 Editor: R.W. Carlson Keywords: Fe isotopes Si isotopes micro-CT chondrules Solar nebula

a b s t r a c t We combined micro computer tomography with Fe and Si isotope measurements of Mokoia, Allende and Grosnaja chondrules. Ten Mokoia chondrules contain 0.9 to 11.8 vol.% opaque phases (metal + sulfide), and 6 Allende chondrules contain 0.0 to 6.6 vol.% opaque phases. Hence, the Fe isotope composition of many chondrules is dominated by the Fe isotope composition of their opaque phases. We studied Fe isotopes of 35 bulk chondrules. The range is different for each of the three meteorites studied and largest for Allende with δ56Fe ranging from − 0.82 to + 0.37‰. Six out of seven chondrules analysed for their Si isotope composition in Mokoia and Grosnaja have similar δ29Si of around − 0.12‰. One anomalous chondrule in Mokoia has a δ29Si of + 0.58‰. We exclude isotopically heterogeneous chondrule precursors and different isotopic chondrule reservoirs as the source of the observed Fe isotope variation among bulk chondrules. We conclude that the observed bulk chondrule Fe isotope variation is the result of evaporation and re-condensation processes in a nebula setting with high dust densities, required to explain the comparatively low isotope fractionations. Subsequent parent body alteration slightly overprinted this pre-accretionary Fe isotope variation. © 2010 Elsevier B.V. All rights reserved.

1. Introduction CV chondrites are a diverse group of primitive meteorites and one of the best studied. They contain about 45 vol.% chondrules, 40 vol.% matrix, and the highest abundance of CAIs (3 vol.%, Hezel et al., 2008). They are divided into a reduced and an oxidised group. Allende, Mokoia and Grosnaja belong to the oxidised group. Oxidised CVs might have developed from reduced CVs in an oxidation event (e.g. Krot et al., 1995). The petrology and bulk isotope composition of chondrules reveal processes and conditions that prevailed in the solar nebula and on the meteorite parent body. All isotope studies of bulk chondrules show only minor fractionations of a few per mil or less, except for O. For example, the typical range for the volatile element K is −3 to 5‰, with only a few chondrules going down to − 15.5‰ and up to +17.8‰ of δ41/39K (Humayun and Clayton, 1995; Alexander et al., 2000; Alexander and Grossman, 2005). Typical fractionations for major elements range from −1.33 to +1.21‰ for δ56/54Fe (Alexander and Wang, 2001; Zhu et al., 2001; Kehm et al., 2003; Mullane et al., 2005; Poitrasson et al., 2005; Needham et al., 2009), and from −0.69 to +0.47‰ for δ29/28Si

⁎ Corresponding author. E-mail address: [email protected] (D.C. Hezel). 0012-821X/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2010.05.029

(Molini-Velsko et al., 1986; Clayton et al., 1991; Georg et al., 2007; Fitoussi et al., 2009). Data for Mg isotopes are inconsistent, as different workers used different standards. The reported maximum range of fractionation is 1.7‰ (Galy et al., 2000; Nguyen et al., 2001; Young et al., 2002; Bouvier et al., 2009). Other meteoritic components, such as CAIs, can have larger isotope fractionations (MacPherson, 2004). For example, Hezel et al. (2009) report Fe and Si isotope data for bulk CAIs from Allende and Mokoia. These have δ56Fe as low as −1.29‰ and δ29Si between −0.78 and +0.83‰. These fractionations are usually attributed to parent body alteration. Individual CAI minerals have larger fractionations. Melilite analyses from Allende type B CAIs have δ29Si of up to around +12‰ and δ25Mg of up to around + 8.5‰ (Young et al., 2002; Richter et al., 2007; Knight et al., 2009). These fractionations are explained by high temperature evaporation processes in the nebula. Cosmic spherules, which formed during atmospheric entry, are small (up to a few 100 µm at maximum), commonly glassy objects. These objects can have large isotopic fractionations of up to + 51.1‰ in δ57Fe, + 8.8‰ in δ29Si, +8.0‰ in δ25Mg and + 183‰ in δ41K, which is close to what is expected from Rayleigh fractionation (Alexander et al., 2002; Engrand et al., 2005; Taylor et al., 2005). However, tektites, which are solidified melts that formed in the aftermath of giant impacts on Earth, have no fractionations of K (∼ 0‰ δ41/39K, Humayun and Koeberl, 2004; Herzog et al., 2008) and minor to

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significant fractionations in Cd (− 0.7 to + 7.6 εCd/amu, Wombacher et al., 2003), Zn (0.17 to 2.49‰, δ66/64Zn Moynier et al., 2009) and Cu (+1.99 to +6.98‰ δ65/63Cu, Moynier et al., 2010). A simple Rayleigh fractionation explanation does not apply to the tektite data. Chondrule formation and evolution is characterised by three major steps: (i) aggregation of the chondrule precursors, (ii) formation of chondrules during flash heating and (iii) secondary processing in the parent bodies. Chondrules from a single meteorite might have formed in single or in separate nebula regions. The observed isotope fractionations can, hence, potentially be produced in each of the three steps of chondrule formation and evolution or in different reservoirs. Each possibility has been advocated in the past as the source of bulk chondrule isotope fractionation. Identifying the correct origin of the isotope fractionation is crucial to characterise the formation conditions of chondrules and, ultimately, understand chondrule formation. Here we combine Fe and Si bulk chondrule isotope measurements, petrologic studies of the same chondrules and 3 dimensional measurements of opaque phase modal abundances in chondrules using micro computer tomography (micro-CT). The combination of these methods allows a detailed discussion of the different origins for isotope fractionation mentioned above. 2. Methods 2.1. Chondrule separation The chondrites were crushed with a hammer to release the chondrules from the matrix. The meteorite was wrapped in weighing paper to prevent contamination from the hammer. Chondrules were then handpicked from the rock dust under a binocular microscope. In a few cases, a ceramic knife was used to remove what could be adhering matrix material. More than a hundred chondrules have been separated, however, only a small fraction was significantly above 1 mg, the minimum mass required for the combined isotope and petrologic analyses. Individual chondrules were placed between two glass plates and a small force applied on the top glass plate broke the chondrule in several pieces, which were used for Fe and Si isotope measurements and petrologic studies. Weighing paper between the chondrule and the glass plates prevented contamination from the glass. 2.2. Sample preparation and mass spectrometry The results of the stable isotope measurements are given in the δnotation, where "
x x

δ el =

 el= y el sample

ðx el= y elstd Þ

# −1 × 1000

For Fe, xel (element) denotes 56Fe and 57Fe, respectively, yel is 54Fe and the standard reference material (std) is IRMM-14. For the Si isotope measurements xel is 29Si and 30Si, respectively, yel is 28Si and the standard reference material is NBS-28. Mixtures of HCl, HNO3 and HF were used to digest the samples, which were heated in closed Savillex Teflon beakers on hotplates. The silicate rich samples were generally refluxed overnight in sealed beakers with HF–HCl and HF–HNO3 mixtures, repeating until dissolution was complete. Chromatographic separation of Fe was performed using methods reported by Dauphas et al. (2004), using AG1-X8 anion exchange resin, and complete recovery of Fe was achieved with this technique. The procedural blank of b50 ng Fe produces a sample to blank ratio of at least 5000:1, introducing a negligible error. The Fe isotope compositions were measured at the University of Oxford and the MAGIC Laboratories of Imperial College London

(Table 1), using Nu Plasma MC–ICP–MS instruments. At Oxford, the analyses were carried out at low mass resolution, using techniques previously described by Belshaw et al. (1998, 2000) and Zhu et al. (2002). These analyses generally utilised sample solutions with Fe concentrations of N25 ppm to overcome isobaric interferences from Ar species. The typical reproducibility of Fe isotope data during the course of this study was ±0.04‰ 2 s.e. (standard error) for δ56Fe and ±0.06‰ 2 s.e. for δ57Fe, as determined from multiple analyses of several standard solutions (in-house rock standards and international, pure Fe, standards) over a period of three years. The uncertainties (2 s.e.) of the Fe results reported in Tables 1 and 2 are slightly larger at ±0.1‰ for δ56Fe and ± 0.2‰ for δ57Fe, due to lower Fe concentrations, smaller sample sizes, fewer repeat analyses, and/or variations in instrument stability. The Fe isotope measurements at Imperial College utilised a Nu Plasma HR instrument in conjunction with a Nu Instruments DSN desolvating nebuliser for sample introduction. The analyses were carried out in pseudo high-resolution mode, with a 0.03 mm source slit. This configuration provided a mass resolving power of ∼ 7500 and sufficient transmission to allow routine analyses of 2 ppm Fe sample solutions. The external reproducibility of the data was typically ±0.07‰ 2 s.e. for δ56Fe and ±0.15‰ 2 s.e. for δ57Fe, whilst the internal reproducibility was generally about a factor of two better. All samples measured in Oxford and in London were analysed between three and six times, non-consecutively, during long analytical sessions of around 8 h duration. Each sample measurement was bracketed by two or more analyses of an IRMM-14 Fe solution that was made up to closely match the Fe concentration of the sample. Only the average results obtained from these multiple analyses are reported here. The Si isotopic composition of the chondrules was analysed using the procedure previously described in Georg et al. (2006). Approximately 1–2 mg of chondrule material was crushed in an alumina mortar and subsequently fused with ∼100 mg solid NaOH at 720 °C for 10 min in a furnace. The resulting fusion cake was transferred into a weakly acidified (HCl) solution. The final solutions (40 ml to 120 ml) were stored in pre-cleaned PPP bottles with Si concentrations, typically between 1.3 and 7 ppm. The ion-chromatography to separate Si from the other cations used BioRad polyprep columns (Hercules, CA, USA) filled with 1.8 ml of strong acidic cation exchange DOWEX 50W-X12 (200–400 mesh) resin in the H+ form. The resin was pre-cleaned with several rinses of MQ-e, HNO3 and HCl to ensure the complete removal of matrix elements and the efficiency of the exchange process. The solutions were run on a MC–ICP–MS (NuPlasma HR) at the Department of Earth Sciences, University of Oxford. The machine was run in medium resolution mode which allows for pseudo highresolution where all three silicon isotope beams can be determined free of interferences. A self-aspirating microflow PFA nebuliser with an uptake rate of 75 μl/min was used to aspirate the samples into an Aridus II (CETAC, Omaha, NE, USA) desolvator. The injector within the torch was made of sapphire (rather than quartz) so as to lower the internal Si blanks. The typical sensitivity for a 500 ppb Si solution was ∼8 V total with 1011 Ω resistors on the collectors. To correct for instrumental mass drift, sample-standard bracketing was used with the international Si standard NBS-28 as the bracketing standard. Each measurement represents 20 cycles with an integration time of 10 s for every cycle. The accuracy of the measurements in each run was checked using various calibrated standards such as IRMM018, Diatomite and BHVO-1. Samples were repeated in different sessions, and the reported data represent the average delta values of multiple runs. The Diatomite values for a 15 month period of δ30Si = 1.22 ± 0.16‰ and δ29Si = 0.62 ± 0.11‰ (±2 s.e., n = 299) are consistent with the calibrated values of δ30Si = 1.26 ± 0.20‰ and δ29Si = 0.64 ± 0.14‰ (Reynolds et al., 2007). Similarly the long term averages of silicon standard IRMM018 (δ30Si = − 1.67 ± 0.20‰, δ29Si = − 0.86 ± 0.13‰) and the rock standard BHVO-1 (δ30Si =

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Table 1 Fe and Si isotope compositions of bulk chondrules.

n

δ56Fe

2 s.e.

δ57Fe

2 s.e.

Grosnaja—BM 35217 #1 ? #2 – #3 IA #4 IA

2 3 3 2

−0.21 − 0.20 0.21 0.09

0.33 0.20 0.09 0.28

− 0.76 − 0.35 0.24 0.08

0.11 0.48 0.29 0.27

Grosnaja—BM 63624 #1 IA

2

− 0.01

0.04

0.05

0.06

Sample

#3 Mokoia—BM #2 #3 #4 #5 #20 #25 #27 #30 #31 #36 #38 #39 #40 #41 #48 #55 #64 Allende #1 #9 #29 #35Lo #37 #40Lo #42 #50 #75Lo #77Lo #82Lo #85Lo

Type

– 1910, 729 IAB IA IA IA IIAB IA BO IA IA IA IAB IA IA IA IA BO IA

BO IAB IA IA I/IIB IA IA IAB I/IIA IA IA IA

1

−0.25

3 3 3 5 2 3 2 3 3 3 4 2 4 3

0.09 0.00 0.32 − 0.08 − 0.18 − 0.33 − 0.47 − 0.41 −0.41 − 0.20 0.40 − 0.10 −0.11 0.12

0.33 0.02 0.21 0.08 0.01 0.07 0.01 0.04 0.09 0.04 0.07 0.08 0.09 0.02

0.22 0.05 0.41 − 0.01 − 0.04 − 0.51 − 0.68 − 0.56 − 0.58 − 0.19 0.58 − 0.24 − 0.18 0.14

0.91 0.18 0.58 0.15 0.10 0.08 0.14 0.06 0.10 0.11 0.28 0.03 0.23 0.13

3 3

− 0.05 0.07

0.06 0.13

− 0.07 0.16

0.18 0.22

4 5 4 6 6 6 4 5 6 6 6 6

− 0.51 0.18 − 0.39 − 0.13 − 0.82 − 0.15 0.37 − 0.28 − 0.06 0.21 − 0.17 − 0.54

0.09 0.11 0.05 0.07 0.09 0.06 0.13 0.03 0.05 0.08 0.08 0.10

−0.71 0.37 − 0.52 −0.16 − 1.24 −0.17 0.56 − 0.42 −0.03 0.37 −0.23 −0.75

0.12 0.25 0.15 0.11 0.24 0.12 0.07 0.11 0.11 0.12 0.17 0.14

n

δ29Si

2 s.e.

δ30Si

Mass

Radius1

FeO ol

2 s.e.

(mg)

(μm)

(wt.%)

0.92 0.7 3.32 10.9

622 568 954 1418

– –

7.4

1247

15 15

− 0.10 − 0.16

0.06 0.12

− 0.21 − 0.27

0.10 0.22

15 15

− 0.06 − 0.12

0.11 0.10

− 0.05 − 0.19

0.16 0.21

− 0.37

15

0.58

0.17

1.19

0.33

15 15 15

− 0.17 − 0.13 − 0.11

0.05 0.06 0.07

− 0.26 − 0.23 − 0.18

0.09 0.13 0.15

3.65 1.35

0.73

2

0.4

479

6.3 4.6 2.6 1.0 3.02 1.82 0.9 1.3 3.42 1.7 1.5 3.7 3.9 8.3 3.42 1.62 1.62

1182 1064 880 640 925 781 618 698 964 763 732 989 1007 1295 964 751 751

0.18 3.08 2.82 0.80 25.98 0.82 2.99 2.78 2.77 6.47 0.80 1.28 8.68 1.36 1.55 3.75 8.50

4.1 3.42 1.8 0.5 1.7 0.42 1.8 2.6

1024 964 778 508 763 479 778 880 – – 508 540

2.09 3.64 3.30 4.81 10.50 7.23 4.71 5.73 11.62 8.27 5.93 1.52

– – 0.5 0.6

–

1 Model chondrule radii are calculated from chondrule masses assuming a chondrule density of ρ = 3 g/cm− 3; 2chondrule fragments, i.e. true mass of the chondrule is larger; ol: chondrule olivine; LoFe measurements done in London. All others were done in Oxford.

− 0.30 ± 0.15‰, δ29Si = − 0.16 ± 0.08‰) are in agreement with the previously published values of δ30Si = −1.65 ± 0.22‰, δ29Si = −0.85 ± 0.14‰ and δ30Si = − 0.33 ± 0.14‰, δ29Si = −0.16 ± 0.09‰, respectively (Reynolds et al., 2007; Abraham et al., 2008). 2.3. Electron microprobe and scanning electron microscopy Chemical analyses of the chondrule silicate phases were obtained by means of a CAMECA SX100 electron microprobe, equipped with five wavelength dispersive spectrometers. The accelerating voltage was set to 15 kV and the beam current to 20 nA. The spot size was set to 1 μm. The built in PAP-algorithm (e.g. Pouchou and Pichoir, 1991) was used for correction. Back scattered electron images of all samples were taken with a LEO 1455 VP scanning electron microscope.

180 µA; 0.5 µm focal spot; 3142 projections with 250 ms second exposure and a voxel (the three dimensional equivalent of pixels) size of 8–10 µm. In order to reduce the effects of beam hardening the X-rays were filtered with a 0.1 mm thick copper plate. The long axis of the object was oriented vertically within the beam, thus ensuring maximum resolution (Kothari et al., 1998) whilst minimising streak artefacts (Zou et al., 2004). The micro-CT data was reconstructed using CT-PRO software version 2.0 (Metris X-Tek). Each voxel is assigned a CT (grey) value derived from a linear attenuation coefficient (i.e. density) of the material being scanned. Hence, a micro-CT scan is not a true image, unlike a radiograph, but rather a mathematical representation of an object. Phases can be separated based on material density, measured by X-ray transmission (Zonneveld, 1987; Spoor et al., 1993). As a result it is easy to collect measurements, such as volume fraction of a material, from within an object.

2.4. Micro-computed tomography 3. Results Pieces of meteorite samples (∼10 cm3 each) were scanned using a HMX-ST CT 225 System (Metris X-Tek, Tring, UK). The instrument utilises a cone beam projection system (see Johnson et al., 2007) with a four megapixel Perkin Elmer XRD 1621 AN3 HS detector panel. The Xray and scan parameters were as follows: tungsten target; 180 kV;

3.1. Petrology of the chondrules We measured a total of 35 chondrules: 17 from Mokoia (BM 1910, 729), 12 from Allende (BM 1981, M5) and 6 from Grosnaja (BM 35217,

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Table 2 Fe isotope composition of bulk chondrites and matrix.

Lo

Sample

n

δ56Fe

2 s.e.

δ57Fe

2 s.e.

Bulk Grosnaja GrosnajaLo GrosnajaLo Mokoia MokoiaLo MokoiaLo AllendeLo AllendeLo AllendeLo Orgueil Orgueil

2 6 6 3 6 6 6 6 6 3 4

− 0.02 0.09 0.13 − 0.07 0.06 0.05 0.02 −0.01 0.01 − 0.03 −0.04

0.10 0.10 0.07 0.09 0.08 0.13 0.07 0.14 0.18 0.13 0.08

0.00 0.15 0.19 − 0.08 0.10 0.07 0.04 0.00 0.12 − 0.02 − 0.08

0.06 0.25 0.26 0.30 0.15 0.18 0.17 0.16 0.32 0.24 0.11

Matrix Grosnaja Grosnaja GrosnajaLo GrosnajaLo GrosnajaLo Mokoia MokoiaLo MokoiaLo AllendeLo AllendeLo AllendeLo

3 1 6 6 6 1 6 6 6 6 6

0.09 −0.04 −0.01 0.15 0.12 − 0.18 0.04 0.09 0.02 0.13 0.15

0.14 0.01 0.05 0.10 0.14

0.23 0.01 0.01 0.21 0.24 − 0.16 0.09 0.08 − 0.03 0.25 0.21

0.26

0.15 0.07 0.08 0.18 0.06

0.17 0.31 0.12 0.14 0.13 0.16 0.29 0.15

Fe measurements done in London. All others were done in Oxford.

BM 63624). Of these 23 are of type IA (1 of these is I/IIA), 4 of IAB, 1 of IIAB, 1 of I/IIB and 3 are BO chondrules (Supplementary Fig. S1). One object in Grosnaja consists nearly entirely of Al-rich clinopyroxene (#1 of BM 35217 in Supplementary Fig. S1c). The splits for SEM of two Grosnaja chondrules were lost during section preparation. None of the chondrule borders visible in SEM images had any matrix attached to it. Model chondrule diameters are calculated from initial chondrule masses assuming a chondrule density of ρ = 3 g/cm− 3. We use the measured chondrule olivine FeO-concentration as an approximation for the silicate FeO-concentration of the bulk chondrule. We do not have 3dimensional chemical information of the chondrules, but two-dimensional bulk compositions obtained from sections can have large errors, especially for Fe and if opaque phases are included in the bulk chondrule composition (Hezel, 2007; Hezel and Kieswetter, in press). In addition, we only have small splits of the chondrules, which make 2D bulk analyses even less representative. The FeO-concentrations of olivine should be the same throughout the chondrule, and as olivine is about in equilibrium with the pyroxenes and the mesostasis during chondrule crystallisation, the olivine FeO-concentration is a reasonable first approximation for the FeO-concentration of the silicate portion of the chondrule. The opaque phases metal, sulfide and magnetite occur either as large blebs inside the chondrule (e.g. Mokoia chondrules #3, #4 in Supplementary Fig. S1), as smaller grains scattered throughout the chondrule (e.g. Mokoia chondrule #31, Allende chondrule #42 in Supplementary Fig. S1) or preferentially located at the border of the chondrule (e.g. Mokoia chondrules #41, #55 in Supplementary Fig. S1). In a single chondrule the three opaque phases occur separate from, and also tightly intergrown with each other. In general, chondrules have no FeO enrichment at their border (e.g. Mokoia chondrules #3, #5 and #27, Allende chondrules #3 and #42 and Grosnaja chondrules #4 and #1-BM 63624 in Supplementary Fig. S1) and, hence, there is no evidence for a widespread exchange of Fe between chondrule and matrix silicate. Opaques in some chondrules have a non-uniform texture and a few others are rimmed, indicating that these experienced in-situ alteration after solidification. In Mokoia chondrules interaction between opaques and surrounding chondrule silicate is limited. One exception is chondrule #4, where the opaque phases are surrounded by a couple of µm thin layer of FeO-rich silicate. In this case, some Fe from the

opaque phase was oxidised and reacted with the surrounding silicate. In Allende, a significant amount of opaques in chondrules were oxidised and replaced by fayalitic olivine and Ca,Fe-rich pyroxene (e.g. chondrules #35, #40, #75, #77, #82 in Supplementary Fig. S1), as described by e.g. Krot et al. (1998). Grosnaja chondrule #3 contains sulfide in Fe-rich veins. The veins are probably oxidised sulfide or metal, similar to what is observed in Allende chondrules. Grosnaja is a hard and strongly cemented rock. 3.2. Micro-CT of chondrules Modal abundances of opaque phases obtained from chondrule sections can be highly unreliable, as opaques tend to be heterogeneously distributed within a chondrule, as previously shown for CR chondrite chondrules by Ebel et al. (2008) and CV chondrite chondrules by Hezel and Kiesswetter (in press). We determined modal abundances of opaque phase in 10 Mokoia and 7 Allende chondrules using micro-CT. The chondrules were randomly chosen from the CT scan and are not the same chondrules as used for the isotope studies. Opaque phase modal abundances range from 0.9 to 11.8 vol.% in the Mokoia chondrules and from 0.0 to 6.6 vol.% in the Allende chondrules, i.e., it appears that Mokoia chondrules have higher opaque phase modal abundances than Allende chondrules (Fig. 1). However, a significant amount of sulfides in Allende chondrules were replaced, and their initial modal abundances were probably higher. The majority of chondrules for which we measured isotope compositions are of type I. The Fe concentration of their olivines typically range from b1 to 3 wt.% (Table 1). In contrast, sulfides contain around 64 wt.% and metal (kamacite) around 90 wt.% Fe. If more than ∼0.65 vol.% metal or more than ∼1.85 vol.% sulfide is present in a chondrule, the opaques will dominate their Fe budget (Fig. 2). Fig. 1 illustrates that the Fe budget of most chondrules in Mokoia and many chondrules in Allende are dominated by their opaque phases. Hence, the Fe isotope composition of many chondrules is dominated by the Fe isotope composition of their opaque phases. 3.3. Fe isotope compositions All chondrules, matrices and bulk chondrites fall on a single massdependent fractionation line with a slope of δ56Fe= 0.674·δ57Fe−0.018 in the Fe 3-isotope plot (Supplementary Fig. S2). Allende chondrules show the largest variation in δ56Fe, between−0.82 and +0.37‰ (Table 1, Fig. 3A). Mullane et al. (2005) reported a similar range for Allende, although with a few chondrules that were even more fractionated resulting in a total range of between−1.33 and +0.65‰. The Mokoia chondrules we measured have a slightly smaller range, between−0.47 and +0.40‰, and Grosnaja chondrules have the smallest range, between −0.21 and +0.21‰ in δ56Fe. The range of δ56Fe for all chondrules decrease with increasing chondrule olivine FeO-concentrations (Fig. 4A) and with an increase in chondrule mass (Fig. 4B). Fig. 5 displays all chondrules that show high opaque phase abundances in the BSE-image. These chondrules do not have a distinct Fe isotope composition, and scatter nearly across the whole range, from−0.41 to + 0.37‰ in δ56Fe. There is also no correlation between the isotope composition of the chondrules and their petrographic type either (Fig. 5). We measured one aliquot of the CI chondrite Orgueil (BM 1960, 331) as a reference, which is unfractionated relative to the IRMM-14 standard, in agreement with previous results (Kehm et al., 2003; Poitrasson et al., 2004). Matrix and bulk compositions of all three chondrites are on average unfractionated relative to Orgueil (Table 2, Fig. 3b). 3.4. Si isotope compositions We measured the Si isotopes of 4 chondrules from Mokoia and 3 from Grosnaja (marked with ‘*’ in Supplementary Fig. S1), for which

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427

Fig. 1. Modal abundances of opaque phases in chondrules from Mokoia and Allende. The dark grey field at the bottom indicates the range in which the silicate dominates the Fe budget of the bulk chondrule, assuming the silicate contains 2.5 wt.% FeO. The light grey field indicates when the opaque phases start to dominate the Fe budget of the chondrule. Starting at the basis of the hatched area the opaques start to dominate the Fe budget, if all opaques are metal. Starting at the top of the hatched area the opaques start to dominate the Fe budget, if all opaques are sulfide (cf. Fig. 2). Opaques dominate the Fe budget in most Mokoia and many Allende chondrules.

we had previously measured their Fe isotopes (Table 1). All data plot on a single mass fractionation line with a slope of δ29Si = 0.5041 · δ30Si −0.0207 (Supplementary Fig. S3). Six of the 7 chondrules are of type IA, 1 is of type IAB. All 6 type IA chondrules are indistinguishable from each other with an average δ29Si of−0.12‰ (Fig. 6). This uniformity and slightly light isotopic composition is similar to a previous report of 5 chondrules from the LL3.4 chondrite Chainpur (average δ29Si of −0.26 ± 0.22‰ 2 s.d.; Georg et al., 2007). It is also similar to the values of two Bjurböle chondrules with δ29Si of−0.10 and−0.26‰, reported by Molini-Velsko et al. (1986). Molini-Velsko et al. (1986) reported a range of−0.56 to + 0.02‰, with an average of 0.28‰ for OC and CC bulk chondrites, Georg et al. (2007) reported a δ29Si for CC of−0.30 ± 0.06‰ 2 s.d. (average of 5 chondrites) and for OC of−0.30 ± 0.06‰ 2 s.d.‰ (average of 6 chondrites), and Fitoussi et al. (2009) reported a δ29Si for CC of−0.19 ± 0.01‰ 2 s.d. (average of 3 chondrites) and for OC of−0.22 ± 0.04‰ 2 s.d.‰ (average of 3 chondrites). Our average chondrule value is close to, but slightly more heavy, than the bulk chondrite value reported by Fitoussi et al. (2009). We measured a δ29Si of +0.58‰ in the type IAB chondrule #2 from Mokoia. This value is significantly different from any other result

reported so far. The only petrologic or chemical difference with the other chondrules is a dominance of pyroxene, i.e. an enrichment in Si. It has been demonstrated that Si-rich chondrules can be the result of gas–chondrule interaction (Hezel et al. 2003; Tissandier et al. 2002; Libourel et al., 2006). The large Si isotope fractionation in this particular chondrule might be the result of such a gas–chondrule interaction. However, more measurements of similar chondrules are required to confirm this hypothesis. In this study, we will not further discuss this anomalous chondrule, and assume that in general chondrules in CV chondrites have a uniform Si isotope composition with δ29Si around −0.12‰. 4. Discussion There are 4 possible mechanisms capable of producing the observed bulk chondrule isotope fractionations: (i) aggregation of isotopically heterogeneous chondrule precursor grains, (ii) formation of chondrules in isotopically distinct nebula reservoirs and subsequent mixing of the chondrules, (iii) elemental redistribution between the solid chondrule and the matrix during a metasomatic

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Fig. 2. When the (mol Fe in opaque)/(mol Fe in silicate) ratio is N 1, the opaque phase dominate the bulk chondrule Fe budget. When the ratio is b1, the silicate phase dominate the bulk chondrule budget. For example, metal dominates the Fe budget of the chondrule, when the chondrule contains more than 0.65 vol.% metal and the chondrule silicate has an FeO-concentration of 2.5 wt.%. Sulfide dominates the Fe budget of a chondrule, when the chondrule contains more than 1.85 vol.% metal and the chondrule silicate has an FeO-concentration of 2.5 wt.%.

event on the parent body, and (iv) elemental redistribution between the molten chondrule and the surrounding gas during chondrule formation. 4.1. Chondrule precursors were not responsible for bulk chondrule Fe isotope variations Not much is known about chondrule precursor material and so far nothing has been unequivocally identified as such. Most workers agree that chondrule precursor aggregates agglomerated from smaller components. The term ‘dust-ball’ has been coined for these aggregates. Although the nature of chondrule precursor grains is unknown, these were probably individual minerals, and hence had different chemical compositions. It has been suggested that precursor heterogeneities are the reason for the bulk chondrule compositional variations in a single chondrite (e.g. Hezel et al., 2006; Jones and Schilk, 2009). In the same way, it has been hypothesised that chondrule precursors had different isotopic compositions and that this explains the bulk chondrule Fe isotope heterogeneities (e.g. Mullane et al., 2005). The designation ‘dust-ball’ implies a chondrule precursor aggregate consisting of hundreds, thousands or even millions of individual grains, each having the size of only a couple of µm or less. However, after agglomerating little more than 10–15 chondrule precursor grains, the resulting aggregate will have the average composition of the precursor grain population already. Hence, a chondrule must not have consisted of more than a dozen or so precursor grains to retain any precursor grain heterogeneities. For CV chondrite chondrules this means that precursor grain sizes must have been larger than 300 µm (Hezel and Palme, 2007). Currently no process is known that could have produced such large precursor grains. These precursor grains must also have had larger Fe isotope fractionations than the chondrules, as any mixing will reduce the initial range. Presolar material, defined by their extreme isotopic anomalies, might be regarded as a source to produce significant isotopic heterogeneity among precursor grains. However, true stardust was rare in the parental cloud of the solar system and around 95% of the material in this cloud was isotopically homogeneous (Zhukovska et al., 2008). Isotopically heterogeneous chondrule precursors are therefore not a likely source for the bulk chondrule isotope variations. In addition, the Si isotope compositions of 6 chondrules we measured are identical and similar to reports in the literature. If chondrule precursors were

Fig. 3. (A) Variation of bulk chondrule Fe isotope compositions. (B) Bulk chondrites and matrices of all samples are unfractionated relative to Orgueil.

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bution contradicts the existence of different Fe isotope reservoirs and sustains the supposition that chondrules of a single chondrite formed in the same chemical and isotopic reservoir. Furthermore, it could be expected that different Fe isotope reservoirs were also different in their Si isotope compositions. Yet, all except one chondrule have similar δ29Si (Fig. 6). 4.3. Influence on bulk chondrule isotope compositions during elemental redistribution on the meteorite parent body

Fig. 4. The extend of the bulk chondrule isotope variation decreases with increasing FeO in chondrule olivine (A) and chondrule mass (B). Some of the studied chondrules were fragments (grey symbols), i.e. their masses are underestimated relative to their true initial masses.

isotopically heterogeneous, this should as well be expected for Si isotopes. 4.2. All chondrules of a single chondrite formed in a single chemical and isotopic reservoir It is controversial whether individual components of a single meteorite formed in a single (e.g. Klerner and Palme, 1999; Hezel and Palme, 2008; Hezel and Palme, 2010) or in different chemical reservoirs and were then mixed together (Anders, 1964; Larimer and Anders, 1967,1970; Grossman and Wasson, 1982,1983a,1983b; Shu et al., 1996; Alexander, 2005; Zanda et al., 2006). Different chemical reservoirs could have had different Fe isotope compositions and then Fe isotope variations among chondrules in a single chondrite could reflect different parental reservoirs. In this case a multimodal distribution of bulk chondrule δ56Fe would be expected with each peak representing a different isotope reservoir. A histogram of our Fe isotope data displays a continuous distribution (Fig. 7). Such a distribution is expected when various parameters, such as different chondrule sizes, FeO-concentrations, heating times, etc. act on a single population of chondrules during an element/isotope exchange between chondrules and their surroundings. This continuous distri-

Krot et al. (1995) and Krot et al. (1998) describe a series of alteration processes that affected CVox chondrite chondrules: (i) Fealkali metasomatism, (ii) replacement of low-Ca pyroxene by fayalitic olivine, (iii) fayalitic olivine veins, (v) fayalitic rims around olivine phenocrysts, and (v) replacement of magnetite–sulfide nodules by fayalitic olivine and Ca–Fe pyroxene. Ash et al. (1999) and Ash and Young (2000) showed that this alteration was accompanied by massdependent O-isotope fractionations. The mechanism and location (nebular or parent body) of these alteration processes are controversial (e.g. Ash et al., 1999; Brearley, 2003). The extent of alteration is in the range of 0–50%, i.e. 50–100% of the chondrite is unaltered, and most chondrules show various degrees of alteration (Brearley, 2006). Bonal et al. (2006) classified Allende as petrologic type N3.6 and Mokoia and Grosnaja as ∼ 3.6, based on studies of organic matter. This has been confirmed by Busemann et al. (2007) and Cody et al. (2008). Although it is obvious that some alteration took place on the parent body, we will argue that Fe exchange between chondrules and matrix on the parent body was limited and had only a minor influence on the bulk chondrule Fe and Fe isotope composition. Chondrules and matrix are in strong thermodynamic disequilibrium. Matrices are typically FeO-rich, containing up to ∼30 wt.% FeO, whereas the majority of chondrules we studied are FeO-poor (type I, cf. Table 1). The Fe-enrichment in fayalitic rims around chondrule olivine phenocrysts contribute only minor to the bulk chondrule Fe. Only if the phenocryst is very Fe-poor, these fayalitic rims can nearly double the bulk chondrule FeO (i.e. from e.g. 2 wt.% to 4 wt.% bulk chondrule FeO). It has been shown that the Fe-rich rims around phenocrysts and the Ferich veins in chondrules are best explained by parent body alteration (e.g. Krot et al., 1995). However, the bulk chondrule Fe budget of many chondrules in Mokoia and Allende is dominated by opaque phases and not silicates (Fig. 1). In these cases, Fe needs to be re-distributed between the opaque phases of the chondrule and the matrix to change the bulk chondrule Fe isotope compositions. Many opaques in Allende chondrules are replaced by fayalitic olivine and Ca,Fe-rich pyroxene, whereas opaques in Mokoia chondrules show less such alteration. Matrix olivine surrounding chondrules are neither depleted in Fe due to loss of Fe from matrix to chondrules, nor enriched in Mg, due to Fe– Mg exchange between chondrules and matrix. The Fe-enrichment of chondrule silicates is restricted to the regions of the sulfide replacement. No evidence can be found that this replacement involved any elemental exchange between chondrule and matrix silicates. We argue that the sulfide replacement was an isochemical oxidation event that occurred inside the chondrule, without Fe exchange between chondrule and matrix. As cited above, it has been suggested that alteration of chondrules occurred in the nebula (e.g. Brearley, 2003 and references therein), and the oxidation of sulfides might be best explained in a nebula setting. We conclude that parent body alteration did not significantly change the bulk Fe and Fe isotope composition of most chondrules. CAIs always have light Fe isotope compositions and are more fractionated than chondrules (e.g. Hezel et al., 2009). Their Fe isotope composition is easily altered, as they contain only minor amounts of Fe. It is likely that their light Fe isotope compositions are the result of kinetic fractionation during a metasomatic event on the parent body. It has for example been suggested by Williams et al. (2005) that metasomatism produced a light Fe isotope signature in mantle rocks.

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Fig. 5. All but one chondrules displayed have large amounts of opaque phases and, hence, their bulk Fe concentrations are dominated by the opaque phases (cf. Figs. 1 and 2). The exception is the BO chondrule at the bottom. The chondrules span nearly the entire range of measured δ56Fe values, i.e. chondrules with high opaque phase abundances do not have a particular isotope composition. Likewise, there is no correlation between petrographic type and isotope composition. This is especially obvious for the two BO chondrules: Mokoia #55 is unfractionated, whereas Allende #1 has a very light composition. op: opaque phase modal abundance.

Parent body alteration probably produced the fayalitic rims around chondrule olivine phenocrysts and fayalitic veins in chondrules (e.g. Krot et al., 1998). This event might have introduced a certain amount of light Fe from the matrix into the chondrules. Bulk chondrule Fe isotope compositions do not appear to scatter around the bulk chondrite value, but all are slightly shifted to lighter compositions (Fig. 3A). If parent body alteration shifted CAIs to light Fe isotope compositions, the slight shift in the scatter of bulk chondrule Fe isotope compositions to lighter values might also be explained by parent body alteration: initially, bulk chondrules scattered around the bulk chondrite value and were then shifted to lighter compositions during the metasomatic event.

There exist two other arguments against large amounts of Fe exchange between chondrules and matrix on the parent body: (i) The matrices of the CV chondrites have unfractionated δ56Fe relative to Orgueil (Fig. 3B). The introduction of kinetically fractionated light Fe from the matrix into chondrules would shift the isotope compositions of all chondrules in the same direction, i.e. towards isotopically light composition. However, a number of the studied chondrules have heavy isotope compositions. (ii) Chondrules display the largest Fe isotope variations compared to other non-biogenic solar system rocks (Fig. 8). Large isotope fractionations in magmatic rocks are often related to melt separation, which does not occur in chondrules (e.g. Williams et al., 2005; Teng et al., 2008). An exception are metasomatic

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Fig. 6. All but one chondrule have the same isotopic composition of δ29Si−0.12‰ within error.

events as discussed above. A parent body setting for the observed isotopic scatter is therefore unlikely. 4.4. Fractionation of bulk chondrule Fe isotopes in a complex system during evaporation and re-condensation in the nebula The extent of isotope fractionation during evaporation and recondensation depends on several independent parameters, such as the bulk chondrule Fe concentration, the chondrule size, the peak temperature of chondrule formation, the duration of the chondrule heating event, the number of repetitive chondrule heating events, the dust densities and the total pressure. The latter is in particular important for the extent of isotope fractionation, as high vapour pressures above the chondrule surface prohibits large evaporation and, hence, extensive isotope fractionation. This large number of

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Fig. 8. Compilation of Fe isotope compositions of various solar system materials. 1Data considered for this range are from Zhu et al. (2001), Kehm et al. (2003), Mullane et al. (2005) and Needham et al. (2009) and our data. 2Teng et al. (2008). All other data are taken from Beard and Johnson (2004).

parameters defines a complex system, and the Fe isotope composition of each chondrule is determined by its individual set of parameters in this system. For example, if a certain amount of Fe is lost or gained during evaporation/re-condensation, the Fe isotope composition of a chondrule with a high Fe concentration will be less affected than of a chondrule with a low Fe concentration. If the chondrules have similar Fe concentrations, but different sizes, the larger chondrule with the lower surface to volume ratio will lose less Fe than the smaller chondrule. Hence, isotope fractionation will be larger in the smaller chondrule and smaller in the larger chondrule. A higher peak temperature and longer heating event experienced by one chondrule will lead to a more extensive fractionation than in a chondrule that experienced a lower peak temperature and shorter heating event. Variable combinations of these parameters result in complex correlations of the sort δ56Fe ∝ Fe concentration of the chondrule, or δ56Fe ∝ chondrule size/mass. Such complex relationships are observed in the Fe isotope compositions of the studied chondrules. The δ56Fe of chondrules is highly variable and decreases with increasing chondrule mass and FeO in chondrule olivine (Fig. 4). This kind of relationship is expected for an evaporation/re-condensation setting and supports this process as source for the bulk chondrule Fe isotope compositions. It is, however, noted that the x-axis in Fig. 4A would ideally display the bulk chondrule Fe content of combined opaques and silicates. This concentration is unknown due to problems outlined in the Method section. If the bulk chondrule Fe content of combined opaques and silicates were displayed on the x-axis, part of the chondrules would shift to the right. We assume that this shift would not change the observed pattern. 5. Conclusions

Fig. 7. Bulk chondrule Fe isotope compositions are continuously distributed. They should display a multimodal distribution if chondrules originated from isotopically different reservoirs.

Chondrules in CV chondrites have variable Fe isotope, but similar Si isotope compositions. The absence of variable Si isotope compositions is most certainly because of the high Si-concentration in chondrules, and because Si diffusion is much slower than Fe diffusion, due to the tetrahedral coordination of Si. The Fe budget of many chondrules is dominated by their opaque phases. Hence, the bulk chondrule Fe isotope composition of these chondrules is dominated by the Fe isotope composition of their opaque phases. The variable Fe isotope compositions of chondrules cannot be explained by

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aggregation of isotopically heterogeneous chondrule precursor grains, isotopically different nebula regions or isotopic exchange between chondrules and matrix in the chondrite parent body. We conclude that the variable Fe isotope compositions are the result of elemental exchange between the still molten or at least hot chondrule and the surrounding nebula gas. This variation was subsequently slightly overprinted and shifted by a metasomatic event on the parent body. A nebula setting excludes simple Rayleigh fractionation into vacuum, as this would produce fractionations up to several tens of per mil (e.g. Alexander, 2001; Davis et al., 2005). Small fractionations can be achieved if the partial pressure of the evaporating species in the surrounding gas is high. It was recently proposed by Cuzzi and Alexander (2006) and Alexander et al. (2008) that high dust/gas ratios can explain small isotope fractionations in chondrules. We consider our results as evidence for this suggested scenario. There is, however, currently no known setting for such a high dust/gas ratio. Johansen et al. (2007) proposed that planetesimals formed in gravitational instabilities. During such a collapse, the dust/gas ratio constantly increases until the final planetesimal is formed. We suggest it might be possible that chondrule formation and the isotopic exchange occurred during such a collapse phase.

Acknowledgements We are grateful for constructive reviews by S. Weyer and C. M. O'D. Alexander which improved the quality and clarity of this paper. Many thanks to Catherine Unsworth, Tony Wighton, Lauren Howard and John Spratt of the EMMA labs at the NHM for all their help. This study was supported by PPARC/STFC.

Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2010.05.029.

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