planetesimal differentiation in the early Solar System?

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1 Available online at Geochimica et Cosmochimica Acta 77 (2012) Al 26 Mg deficit dating ultramafic meteorites and silicate planetesimal differentiation in the early Solar System? Joel A. Baker a,, Martin Schiller a,b, Martin Bizzarro b a School of Geography, Environment and Earth Sciences, Victoria University of Wellington, P.O. Box 600, Wellington 6014, New Zealand b Centre for Star and Planet Formation, Natural History Museum of Denmark, University of Copenhagen, Øster Voldgade 5-7, Copenhagen DK-1350, Denmark Received 10 June 2011; accepted in revised form 17 October 2011; available online 21 October 2011 Abstract Meteorites with significantly sub-chondritic Al/Mg that formed in the first 2 million years of the Solar System should be characterised by deficits in the abundance of 26 Mg (d 26 Mg * ) due to the absence of in-growth of 26 Mg from the decay of shortlived 26 Al (t 1/2 = 0.73 Myr). However, these 26 Mg deficits will be small (d 26 Mg * > 0.037&) even for material that formed at the same time as the Solar System s oldest solids calcium aluminium-rich inclusions and thus measurement of these deficits is analytically challenging. Here, we report on a search for 26 Mg deficits in three types of ultramafic meteorites (pallasites, ureilites and aubrites) by multiple-collector inductively coupled plasma mass spectrometry. A range of analytical tests were carried out including analysis of: (1) a range of synthetic Mg solution standards; (2) Mg gravimetrically doped with a high purity 26 Mg spike; (3) Mg cuts collected sequentially from cation exchange separation columns with fractionated stable Mg isotope compositions; (4) Mg separated from samples that was bracketed by analyses of both DSM-3 and Mg separated from a natural olivine sample subjected to the same chemical processing as the samples. These tests confirm it is possible to resolve differences in d 26 Mg * from the terrestrial materials that are &. However, if Mg yields from chemical separation are low or an inappropriate equilibrium-isotopically fractionated standard is used this will generate analytical artefacts on d 26 Mg * when this is calculated with the kinetic/exponential mass fractionation law as is the case when correcting for instrumental mass bias during mass spectrometric analysis. Olivine from four different main group pallasites and four bulk ureilites have small deficits in the abundance of 26 Mg with d 26 Mg DSM-3 ¼ 0:0120 0:0018& and d26 Mg DSM-3 ¼ 0:0062 0:0023&, respectively, relative to terrestrial olivine (d 26 Mg DSM-3 ¼þ0:0029 0:0028&). Six aubrites have d26 Mg DSM-3 ¼þ0:0015 0:0020&, which is identical to terrestrial olivine. Model ages from these deficits can be calculated by assuming that 26 Al was homogeneously distributed in the planetesimalforming regions of the proto-planetary disc at the same level as calcium aluminium-rich inclusions (CAIs). The absence of 26 Mg deficits in aubrites, means these can only be constrained to have formed relatively late >2.9 Myr after CAI formation. Model ages calculated from pallasite olivine deficits would suggest that pallasite olivine crystallised and was diffusively isolated on its parent body 1:24 þ0:40 0:28 Myr after formation of CAIs. Similarly, ureilites would have experienced silicate partial melting and lowering of their bulk Al/Mg ratios 1:9 þ2:2 0:7 Myr after CAI formation. The model ages for silicate differentiation on the main group pallasite parent body are intermediate between those for metal-silicate fractionation for core formation obtained from magmatic iron meteorites and those for asteroidal silicate magmatism obtained from basaltic meteorites. Ó 2011 Elsevier Ltd. All rights reserved. Corresponding author. Tel.: ; fax: address: joel.baker@vuw.ac.nz (J.A. Baker) /$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi: /j.gca

2 416 J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) INTRODUCTION A variety of long- (absolute) and short-lived (relative) chronometers are used to date meteorites and their components as a result of the processes of solid formation and planetary accretion and differentiation in the early Solar System. In particular, over the past several decades, application of absolute Pb Pb and relative 53 Mn 53 Cr, 182 Hf 182 W and 26 Al 26 Mg dating techniques have led to an increasingly clearer picture of these timescales (e.g., Lugmair and Galer, 1992; Lugmair and Shukolyukov, 1998; Srinivasan et al., 1999; Amelin et al., 2002; Kleine et al., 2002; Yin et al., 2002; Bizzarro et al., 2005; Spivak- Birndorf et al., 2009; Wadhwa et al., 2009). However, a number of types of meteorites and their components are not easily dated with conventional application of these isotopic systems. Examples of meteorites that are difficult to date include those where: (1) the meteoritic material may contain very low lithophile trace element abundances characterised by a low ratio of the parent (P) radioactive isotope to the daughter (D) radiogenic isotope. (2) Particularly in combination with (1) some isotopic systems are highly sensitive to the affects of terrestrial contamination, such as the Pb Pb chronometer (Torigoye-Kita et al., 1995a; Amelin, 2006). (3) Co-existing phases in a meteorite with high and low P/D ratios have re-equilibrated due to slow cooling of the meteorite parent body and/or due to later thermal or shock events, especially where the high P/D phase has high diffusivity for the element of the daughter isotope. Numerous examples of meteorite dating studies with the 53 Mn 53 Cr, 182 Hf 182 W and 26 Al 26 Mg chronometers have interpreted the obtained ages to reflect secondary events rather than the primary crystallization age of a meteorite (or its components) (e.g., Wadhwa et al., 2003; Shukolyukov and Lugmair, 2004; Kleine et al., 2005a; Touboul et al., 2009). 26 Al 26 Mg dating has been used as a relative dating tool for meteorites since the first demonstration that live 26 Al was present in the Solar System s oldest dated solids (calcium aluminium-rich inclusions; CAIs) when they formed (Lee et al., 1977; Gray and Compston, 2004). Since then both bulk analytical methods applied to whole rock samples and mineral separates (thermal ionisation and plasma source mass spectrometry) and in situ (secondary ionisation and laser ablation plasma source mass spectrometry) analytical methods have utilised the 26 Al 26 Mg chronometer to date solid formation (CAIs and chondrules) and planetesimal magmatism (e.g., Russell et al., 1996; Kita et al., 2000; Bizzarro et al., 2005; Simon et al., 2005; Young et al., 2005; Jacobsen et al., 2008; Spivak-Birndorf et al., 2009). In these cases, the 26 Al 26 Mg chronometer has been primarily used to date meteorites or their components that have high P/D (i.e., Al/Mg) ratios. The initial abundance of 26 Al in the Solar System is sufficiently high that meteorites which formed very early in the Solar System with markedly sub-chondritic Al/Mg (i.e., 27 Al/ 24 Mg 0.101; Lodders, 2003) might have less radiogenic 26 Mg as compared to the bulk Solar System (Fig. 1). Fig. 1 shows a Mg isotopic evolution curve for Fig. 1. d 26 Mg * isotopic evolution of the Solar System in the first 5 million years after CAI formation. Meteorites or meteoritic material characterised by sub-chondritic Al/Mg ( 27 Al/ 24 Mg <0.101) will be characterised by resolvable d 26 Mg * deficits if they formed within 2 million years of CAIs. The isotopic evolution curve shown is calculated using the initial 26 Al value for the Solar System and its proto-planetary disc based on the assumption that the planetesimal-forming region of the proto-planetary disc had the same initial 26 Al/ 27 Al as indicated by a regression through Mg isotope data for chondrites and CAIs (Jacobsen et al., 2008; Schiller et al., 2010a). the Solar System based on an initial 26 Al/ 27 Al value of derived from a regression through high precision Mg isotope data for CAIs and bulk chondrites (Jacobsen et al., 2008; Schiller et al., 2010a) and the chondritic Al/ Mg ratio, which shows it is possible to hypothesize that the early Solar System evolved from an initial d 26 Mg * deficit of 0.037& to essentially its present value within 5 million years. While these potential d 26 Mg * deficits are small, improved analytical methods, and evidence that planetesimals accreted and differentiated very early in the Solar System (e.g., Kleine et al., 2005b) make it plausible that it will be possible to detect such deficits in appropriate meteoritic material. This would provide a potential dating methodology for high-mg meteoritic material in a manner that is analogous to 182 Hf 182 W dating of metal phases in meteorites with Hf/W 0 (e.g., Kleine et al., 2005b; Markowski et al., 2006a,b; Schersten et al., 2006). This type of approach has recently been utilised by Villeneuve et al. (2010) to date chondrule formation, and Mg-rich olivines in chondrites and the Eagle Station pallasite by ion probe methods. Herein we describe the methods and results of a search for deficits in the abundance of 26 Mg in three classes of meteorites (pallasites, ureilites and aubrites) with subchondritic Al/Mg ratios. We show that it is possible to resolve small variations in d 26 Mg * in meteorites compared to terrestrial material and chondrites and potentially place some new age constraints on the formation of pallasites, ureilites and aubrites, and silicate planetary differentiation of their parent bodies in the early Solar System.

3 26 Al 26 Mg deficit dating ultramafic meteorites STANDARDS, SAMPLES AND ANALYTICAL TECHNIQUES 2.1. Standards and samples A range of synthetic Mg solution standards and Mg separated from an in-house olivine standard (J11) taken from a spinel peridotite were analysed in this study. The primary Mg solution standard used for bracketing analyses of other Mg solution standards as well as Mg separated from terrestrial olivine and meteorite samples was DSM-3 (Galy et al., 2003). Various ICP-MS Mg solution standards (Aristar, Alfa Aesar 1, and Alfa Aesar 2) and SRM980 were analysed versus DSM-3. Aristar Mg gravimetrically spiked with a high purity (>99.5%) 26 Mg spike and Aristar Mg subjected to various chemical separation techniques was also analysed versus unspiked/unprocessed Aristar Mg (Section 2.5). Mg separated from olivine minerals in samples J11, JK3 and JB281 was analysed versus the DSM-3 Mg standard and, in a number of cases, Mg separated from J11 mantle olivine was also used as the bracketing standard for Mg isotope analysis of Mg separated from these terrestrial olivines and the meteorite samples. Where Mg separated from J11 mantle olivine was used as the bracketing standard, Mg was separated from J11 olivine in exactly the same fashion as the terrestrial olivines or meteorite samples being analysed. J11 is an anhydrous spinel peridotite collected from a Plio-Quaternary intraplate volcanic cone located in Jordan (Shaw et al., 2007), whereas JK3 is a hydrous amphibolebearing spinel peridotite collected from Plio-Quaternary intraplate volcanism located in Ataq (southern Yemen) (Baker et al., 1998) and JB281 are olivine phenocrysts from a near-primary continental flood basalt sourced from the Afar mantle plume erupted in Yemen during the Oligocene (Baker et al., 1996). Three different types of meteorites with sub-chondritic Al/Mg were analysed in this study olivine from main group pallasites, and bulk ureilite and aubrite samples. Main group pallasites are stony-iron meteorites containing large mm-sized olivine crystals set in evolved iron nickel metal that are thought to represent the core-mantle boundary of differentiated planetesimals (Wasson and Choi, 2003). Ureilites are coarse-grained meteorites composed primarily of olivine and pyroxene, but also containing Fe Ni metal, sulphide phases and various forms of carbon (e.g., Goodrich, 1992). Ureilites are highly depleted in incompatible lithophile elements and variably depleted in siderophile elements (e.g., Boynton et al., 1976; Goodrich, 1992; Mittlefehldt, 2007, reference therein). While ureilites exhibit some features typical of both primitive and differentiated meteorites and their precise origin remains enigmatic, they are generally considered to represent partial (s)melting residues of a differentiated planetesimal or planetesimals (Warren et al., 2006). Aubrites are coarse-grained meteorites primarily composed of variably brecciated Fe-free orthopyroxene with associated small and varying amounts of plagioclase, high-ca pyroxene and forsterite as well as a suite of accessory metal and sulphide phases (Mittlefehldt, 2007, reference therein). Aubrites are highly depleted in both lithophile and siderophile trace elements and, although their origins are also unclear, are interpreted as being coarse-grained igneous cumulates from a highly reduced, differentiated planetesimal. Magnesium separated from olivine from four main group pallasites (Admire, Brenham, Esquel and Molong) was analysed in this study. These pallasites contain olivine with a restricted range of Fo contents (87 89; Wasson and Choi, 2003). Small sub-mm-sized fragments of olivine devoid of chromite inclusions were hand-picked from the pallasite meteorites (and also terrestrial mantle and basalt samples) under a binocular microscope. Prior to digestion, the olivine samples were washed with ultra-clean water (>18.2 MX) and gently acid washed with cold 2 M HCl for 5 10 min to remove any secondary material resulting from oxidation of iron nickel metal and/or the olivine. Bulk samples of four ureilites were analysed in this study. SAH98505 is a coarse-grained ureilite dominated by olivine (Fo 81 ) and pigeonite (Fs 12.9 )(Grossman, 1999). El Gouanem is a ureilite find from Morocco and has a typical ureilitic texture and is primarily composed of olivine with Fo (Grossman and Zipfel, 2001). NWA2234 is a crystalline ureilite composed of coarse, shocked dusty olivine (cores Fo 82 92, rims Fo ) and pigeonite (Fs 17 )(Russell et al., 2004). NWA766 is a ureilite containing 80% olivine (Fo 76 ) and 20% pigeonitic pyroxene (Fs 18.7 ), but marked by the presence of a high-si Al glass (Skirdji and Warren, 2001). Bulk samples of six aubrites were analysed Norton County, Pena Blanca Spring, Mt. Egerton, Shallowater, Cumberland Falls and Bishopville. All of these aubrites are dominated by unbrecciated (Mt. Egerton and Shallowater) to variably brecciated Fe-free enstatite pyroxene crystals, although Cumberland Falls contains some unequilibrated chondritic material (Neal and Lipschutz, 1981) and Bishopville contains a much higher modal abundance of plagioclase than the other aubrites (Watters and Prinz, 1979). Representative fragments of ca mg of the ureilites and aubrites were hand-picked under a binocular microscope. These fragments were then crushed to a powder with an agate mortar and pestle Sample digestion and chemical separation of Mg Samples were digested in a 3:1 mixture of concentrated HF and HNO 3 acid in savillex Teflon capsules on a hotplate at 130 C. Approximately 5 10 mg of olivine from the terrestrial samples and pallasites was digested, whereas ca. 50 mg aliquots of powdered ureilite or aubrite material was digested. After evaporation of the HF HNO 3 acid, samples were sequentially refluxed and evaporated with concentrated HNO 3, 7 M HCl and aqua regia to bring them fully into solution. Samples were finally converted to chloride form by evaporation of 7 M HCl prior to dissolution in concentrated HCl for the first Mg chemical separation step. All acids used in this study were high purity Seastar acids, where necessary, diluted with >18.2 MX ultra-clean water. Chemical separation techniques for purification of Mg utilised in this study have been previously described in

4 418 J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) detail by Handler et al. (2009) and Schiller et al. (2010a,b) and are only briefly described here. An amount of sample equivalent to ca. 1 mg of Mg was subjected to up to five chemical separation steps comprising: (1) anion exchange separation (0.5 ml Bio-Rad AG1-X4 resin) of Fe in concentrated HCl whereby Fe is retained on the resin; (2) separation of Ca in 3 M HNO 3 on 0.25 ml Eichrom DGA resin whereby Ca is retained on the resin; (3) cation exchange separation of other major and trace elements with the exception of Mn and Ni in 1 M HNO M HF on 1 ml of Bio-Rad AG50W-X mesh resin; (4) cation exchange separation of Mn in 0.5 M HCl 95% acetone on 1 ml of Bio-Rad AG50W-X mesh resin and (5) separation of Ni using 1 ml of Eichrom Ni-spec resin. Different types of samples were subjected to different chemical separation steps. The aubrites and relatively Nirich ureilites were subjected to the most extensive chemical separation steps (aubrites steps 1 4 above; ureilites steps 1 5 above) to ensure the complete removal of matrix elements. However, the terrestrial and pallasite olivines have relatively simple compositions compared to the aubrites and ureilites. Thus, the olivines were analysed after just step (1), steps (1 3) and steps (1 4) to assess the extent to which different chemical separation procedures affected the Mg isotope results. When Mg separated from J11 was used as a bracketing standard, it was treated to exactly the same Mg chemical separation procedure as the same being analysed to minimise any potential analytical artefacts. All Mg separates were screened for the presence of matrix elements prior to Mg isotope analysis. Mg extracted from olivines after the first anion exchange step was >98.5% (terrestrial olivine) and >99.0% (pallasite olivine) pure. Mn and Ni were the most abundant matrix elements remaining ( %) after this first chemical separation step, along with very minor amounts of Al, Ca and Cr (<0.1% total), although no Ni was present in the pallasite olivine Mg due to the very low Ni contents of pallasite olivine. After complete processing, Mg purity was always >99.5% in all the analysed terrestrial and meteoritic samples, with only small amounts of Ni (<0.5%) remaining in the Mg separated from the terrestrial olivines. Mg procedural blanks were always <0.001% of the processed Mg for each sample. All of the chemical separation steps were checked to ensure that they resulted in 100% Mg yields Al/Mg ratio measurements by ICP-MS Al/Mg ratios were measured on aliquots of dissolved samples taken before chemical separation of Mg with an Agilent 7500CS ICP-MS in Victoria University of Wellington s Geochemistry Laboratory using He in a collision cell to minimise interferences on Al and Mg isotopes ( 27 Al, 24 Mg and 25 Mg). Sample analyses were bracketed with analyses of a gravimetrically prepared Al/Mg = 0.1 solution made from Aristar single element ICP-MS solutions. The error assigned to the Al/Mg ratio is ±2% (2 sd) based on repeated measurements of USGS basaltic rock standards BCR-2 and BHVO-2 (Schiller et al., 2010b). Al/Mg ratios of the pallasite olivines were not measured precisely as these were always < Mg isotope analyses by MC-ICP-MS Mg isotope ratios were measured in pseudo-high-resolution mode with a Nu Plasma MC-ICP-MS in Victoria University of Wellington s Geochemistry Laboratory. Mg solutions containing ca. 1 3 ppm Mg were introduced into the plasma with a DSN-100 desolvating nebuliser system. The mass spectrometer was operated at a resolution of ca , which enables resolution of polyatomic interferences on the high mass side (e.g., 12 C + 2, 12 C 14 N + ) of Mg. 24 Mg, 25 Mg and 26 Mg were either monitored by the L5, Ax and H6 Faraday collectors equipped with X resistors or by L4, Ax and H5 collectors where the L4 collector was equipped with a X resistor allowing larger ion beams of ca , 3 6 and 3 6 V to be measured on masses 24, 25 and 26 respectively. Results using the latter collector configuration do not differ significantly from the former configuration except that, in some analytical sessions, internal errors were improved by about a factor of 1.5 when measuring the larger ion beams. A single Mg isotopic analysis comprises a total of 480 s of baseline measurements and 1600 s of data acquisition in four blocks (four blocks of 80 5 s integrations). Sample analyses were either bracketed by analyses of the DSM-3 standard or by Mg separated from the J11 in-house olivine standard. J11 olivine has a stable Mg isotopic composition that is slightly lighter than DSM-3 and more representative of terrestrial Mg than DSM-3 (Handler et al., 2009). We used Mg separated from J11 as the bracketing standard for some of our analyses for two reasons. Firstly, it is possible (but not known) that the slight isotopic difference between DSM-3 and Earth could be the result of equilibrium isotopic fractionation processes (Young and Galy, 2004). If so, pure equilibrium fractionation would produce very marginally erroneous mass-bias-corrected 26 Mg * values for the mass-independent abundance of 26 Mg when calculated using the kinetic (=exponential) fractionation law by 0.004& per 0.1& difference in the stable isotopic difference (d 25 Mg) between the sample and bracketing standard. Secondly, we used Mg separated from J11 as the bracketing standard for some analyses as this meant that the sample and standard had both experienced exactly the chemical separation procedures, providing an additional test of our analytical methodology. The mass-independent abundance of 26 Mg (d 26 Mg * ) was calculated by internally normalising the 26 Mg/ 24 Mg to 25 Mg/ 24 Mg = (Catanzaro et al., 1966) using the exponential mass fractionation law (b = 0.511) and calculating the difference between this value for the sample and the average value of the bracketing standards in the per mil notation. Stable Mg isotope data (d 25 Mg) are reported in the per mil notation as the difference between the sample and the average value of the bracketing standards. Single Mg isotope analyses have uncertainties (2 se) on d 26 Mg * that are ±0.021& to ±0.012& when the uncertainties on the bracketing standards are quadratically incorporated into the error on the sample. Each Mg isotope analysis presented in Tables 1 4 represent the weighted mean of 2 38 such measurements resulting in final 2 se

5 Table 1 Mg isotope data for standard solutions (Aristar, Alfa Aesar, SRM980) and terrestrial olivine separated from mantle (J11, JK3) and basalt (JB281) samples. Sample d 26 Mg * ±2 se d 26 Mg * ±2 se d 25 Mg ±2 se d 26 Mg ±2 se n Bracketing standard Chemistry D Solution standards Average & 2 se Weighted mean & 2 se Aristar J11 olivine None Alfa Aesar J11 olivine None Alfa Aesar J11 olivine None SRM980 (NZ) J11 olivine None Aristar (d 26 Mg * = &) A Aristar None Aristar (d 26 Mg * = &) Aristar None Aristar (d 26 Mg * = &) Aristar None Aristar (d 26 Mg * = &) Aristar None Aristar (column processed) B Aristar a Aristar (column processed) Aristar a, dga, c Aristar (column processed) Aristar a, dga, c Terrestrial mantle and basalt olivine C J11 digestion DSM-3 a J11 digestion DSM-3 a J11 digestion DSM-3 a, dga, c J11 digestion DSM-3 a, dga, c, cmn J11 digestion DSM-3 a, dga, c, cmn J11 digestion J11 olivine a, dga, c, cmn JK J11 olivine a, dga, c JB J11 olivine a, dga, c J11 olivine mean versus DSM J11/JB281/JK3 olivine mean versus J A Aristar solutions with gravimetrically prepared artificial d 26 Mg * excesses. B About 1000 lg of Aristar Mg standard solution chemically processed through the Mg chemical separation procedure listed. C J11 = mantle olivine from an anhydrous spinel peridotite (Jordan); JK3 = mantle olivine from a hydrous spinel peridotite (Yemen); JB281 = olivine phenocrysts from an Oligocene continental flood basalt (Yemen). D a = Anion chemical separation (Fe); dga = TODGA Eichrom separation (Ca); c = cation exchange separation (most elements except Mn and Ni); cmn = cation exchange separation in HCl/ acetone (Mn). 26 Al 26 Mg deficit dating ultramafic meteorites 419

6 420 J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) Table 2 Mg isotope data for Aristar Mg standard solution isotopically fractionated on cation exchange columns. Sample d 26 Mg * ±2 se d 25 Mg ±2 se d 26 Mg ±2 se n Weighted mean & 2 se Aristar (column cut; ml) A Aristar (column cut; ml) Aristar (column cut; ml) ml B ml ml ml ml ml A About 20 ml aliquots of Aristar Mg sequentially collected from 5000 lg of Aristar Mg standard passed through a cation exchange column with a resin bed of 4.5 ml in 1 M HNO 3. B The 1 2 ml aliquots of Aristar Mg sequentially collected from 10,000 lg of Aristar Mg standard passed through a cation exchange column with a AG50W-X8 resin bed of 4.5 ml in 1 M HCl. analytical uncertainties of ±0.013& to ±0.004&. While the errors for the data presented here are internal errors, the external reproducibility on d 26 Mg * was typically found to be 50% greater than internal errors based on repeated measurements of samples, terrestrial standards and standards with gravimetrically prepared excesses in 26 Mg (herein; Schiller et al., 2010b). Thus we estimate that the external reproducibility of sample analyses is a factor of 1.5 greater than internal errors. A typical analysis of a meteorite sample comprises the mean of 6 12 individual measurements carried out over an 8 14 h period consuming about lg of Mg. In a number of cases, this type of analysis would then be repeated on either remaining Mg left over from the first 6 12 analyses on another day in another analytical session, or on Mg separated in a new chemical separation chemistry from remaining digested material of the sample, or of Mg separated from the same sample in a new chemical separation chemistry where new material was digested a second or third time. In the data presented in Tables 1 4, n refers to the number of repeat measurements carried out on each solution. When a sample was analysed in a number of different ways and results pooled into a final mean, n is expressed as, for example 35/4/3 (Table 1; Admire pallasite olivine measured versus DSM-3). In this case, three different separates of olivine from the Admire pallasite were digested, subjected to chemical separation on Mg three occasions utilising three different types of chemical separation, and analysed in total 35 times in four analytical sessions Analytical tests Six analytical tests were carried out on standards, terrestrial samples and meteorite samples to assess potential analytical artefacts and demonstrate the precision and accuracy of the presented Mg isotope data: (1) a range of synthetic Mg solution standards were analysed versus DSM-3. (2) Aristar Mg doped with a high purity 26 Mg spike to produce 26 Mg * anomalies in the range of & with an accuracy of <5% was analysed against undoped Aristar Mg. (3) Aristar Mg subjected to various parts of the Mg chemical separation procedures was analysed versus unprocessed Aristar Mg. (4) Aristar Mg cuts collected sequentially from cation exchange separation columns with fractionated stable Mg isotope compositions were analysed to examine the effects of incomplete Mg recovery on measurements of the mass-independent abundance of 26 Mg. In this experiment, a larger volume of resin (4.5 ml) was utilised than in the processing of samples, in order to ensure that isotopically fractionated cuts of Mg would be obtained. (5) Mg separated from three different terrestrial olivine samples and some meteorite samples was analysed versus both DSM-3 and Mg separated from the in-house J11 mantle olivine standard. (6) The Mg isotopic composition of both the mantle olivine (J11) and pallasite olivine samples were measured after different stages (anion exchange separation of Fe ± TODGA separation of Ca ± cation exchange separation of most major and trace elements ± cation exchange separation of Mn in 0.5 M HCl 95% acetone) of chemical separation of Mg to assess the extent to which the progressive removal of matrix elements affected the results. 3. RESULTS 3.1. Analytical tests and analyses of solution and mineral standards Analysis of Mg solution standards versus DSM-3 Four Mg solution standards were analysed versus DSM- 3 and yielded light stable Mg isotopic compositions of d 25 Mg = 0.41& to 2.33& (Table 1). With the exception of the most fractionated standard (SRM980) all the other Mg solution standards show small apparent excesses in the abundance of 26 Mg with respect to DSM-3 of d 26 Mg * = & to & (Table 1 and Fig. 2) Analysis of Aristar Mg doped with a high purity 26 Mg spike Four Aristar Mg standard solutions were gravimetrically doped with >99.5% pure 26 Mg to create solutions with

7 Table 3 Mg isotope data for olivine from pallasite meteorites. Sample Type 27 Al/ 24 Mg d 26 Mg * ±2 se d 26 Mg * ±2 se d 25 Mg ±2 se d 26 Mg ±2 se n Bracketing standard Chemistry A Pallasites Average & 2 se Weighted mean & 2 se Admire digestion 1 Fo DSM-3 a Admire digestion J11 olivine B a Admire digestion J11 olivine a, dga, c Admire digestion DSM-3 a, dga, c Admire digestion J11 olivine a, dga, c Admire digestion J11 olivine a, dga, c, cmn Admire digestion DSM-3 a, dga, c, cmn Admire digestion DSM-3 a, dga, c, cmn Admire mean versus DSM /4/3 Admire mean versus J /4/3 Brenham digestion 1 Fo J11 olivine a Brenham digestion DSM-3 a Brenham digestion J11 olivine a, dga, c Brenham digestion DSM-3 a, dga, c, cmn Brenham mean versus DSM /2/2 Brenham mean versus J /2/2 Esquel digestion 1 Fo J11 olivine a Esquel digestion DSM-3 a Esquel digestion J11 olivine a, dga, c Esquel digestion DSM-3 a, dga, c, cmn Esquel mean versus DSM /2/2 Esquel mean versus J /2/2 Molong digestion 1 Fo DSM-3 a, dga, c Molong digestion J11 olivine a, dga, c Molong digestion DSM-3 a, dga, c Molong digestion J11 olivine a, dga, c Molong digestion DSM-3 a, dga, c, cmn Molong digestion J11 olivine a, dga, c, cmn Molong mean versus DSM /3/3 Molong mean versus J /3/3 A a = Anion chemical separation (Fe); dga = TODGA Eichrom separation (Ca); c = cation exchange separation (most elements except Mn and Ni); cmn = cation exchange separation in HCl/ acetone (Mn). B Mg separated from J11 a mantle olivine from an anhydrous spinel peridotite (Jordan). 26 Al 26 Mg deficit dating ultramafic meteorites 421

8 Table 4 Mg isotope and Al/Mg data for ureilite and aubrite meteorites. Sample Type 27 Al/ 24 Mg d 26 Mg * ±2 se d 26 Mg * ±2 se d 25 Mg ±2 se d 26 Mg ±2 se n Bracketing standard Chemistry A Ureilites Average & 2 se Weighted mean & 2 se SAH98505 Fo /2/1 DSM-3 a, dga, c, cmn, dmg /3/1 J11 olivine B El Gouanem Fo /2/1 DSM-3 a, dga, c, cmn, dmg /1/1 J11 olivine NWA2234 Fo /2/1 DSM-3 a, dga, c, cmn, dmg /2/1 J11 olivine NWA766 Fo /2/1 DSM-3 a, dga, c, cmn, dmg /1/1 J11 olivine Aubrites Norton County /2/1 DSM-3 a, dga, c, cmn Pena Blanca Spring /1/1 DSM-3 a, dga, c, cmn Mt. Egerton /1/1 DSM-3 a, dga, c, cmn Shallowater /1/1 DSM-3 a, dga, c, cmn Cumberland Falls /2/1 DSM-3 a, dga, c, cmn Bishopville /1/1 DSM-3 a, dga, c, cmn Bishopville (pyroxene) /1/1 DSM-3 a, dga, c, cmn A a = Anion chemical separation (Fe); dga = TODGA Eichrom separation (Ca); c = cation exchange separation (most elements except Mn and Ni); cmn = cation exchange separation in HCl/ acetone (Mn); dmg = dimethylgloxamine Ni-specific chemistry. B Mg separated from J11 a mantle olivine from an anhydrous spinel peridotite (Jordan). 422 J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012)

9 26 Al 26 Mg deficit dating ultramafic meteorites Mg * excesses of 0.010&, 0.020&, 0.030& and 0.200&. These solutions were then analysed a number of times versus undoped Aristar Mg. While the doped standards were analysed a variable number of times with differences in the resultant 2 se on the final d 26 Mg * values, all solutions yielded d 26 Mg * values within 2 se analytical uncertainties of the expected value (Table 1 and Fig. 2). For example, the solution with a 0.010& 26 Mg excess produced a mean d 26 Mg * = ± & after eight analyses Aristar Mg subjected to chemical separation of Mg Approximately 1 mg of Aristar Mg was processed through two different types of chemistries anion exchange separation of Fe and (on two occasions) anion exchange separation of Fe + TODGA separation of Ca + cation exchange separation of most major and trace elements. Both the abundance of 26 Mg (d 26 Mg * ) and stable Mg isotopic composition of the column-processed standards produced values within 2 se analytical uncertainties of zero (Table 1 and Fig. 2) Aristar Mg collected sequentially from a cation exchange separation column Aristar Mg collected sequentially from a cation exchange column in both 1 M HNO 3 and 1 M HCl shows that heavier isotopes of Mg preferentially pass faster through the column as compared to lighter isotopes of Mg (Table 2). These different Mg cuts do not generally yield calculated d 26 Mg * values that are within 2 se analytical uncertainties of zero. In particular, heavy (d 25 Mg = +1.38&) and light (d 25 Mg = 0.76&) Mg in these cuts is characterised by apparent deficits (d 26 Mg * = &) and excesses (d 26 Mg * = &) in the abundance of 26 Mg relative to unprocessed Aristar Mg (Table 2 and Figs. 2 and 3) Analysis of terrestrial olivine standards Mg separated from olivine crystals taken from three terrestrial materials was analysed versus DSM-3 and Mg separated from J11 mantle olivine. The mean d 26 Mg * obtained on Mg from J11 mantle olivine analysed versus DSM-3 was ± & (Table 1 and Fig. 2). The stable Mg isotopic composition of d 25 Mg DSM-3 ¼ 0:10 0:06& is within error of the value published by Handler et al. (2009). Mg separated from mantle olivine samples (J11 and JK3) as well as basaltic olivine (JB281) yields d 26 Mg * ( ± &) and d 25 Mg (0.029 ± 0.062&) values, as would be expected, within 2 se uncertainty of zero when measured versus Mg separated from J11 olivine in a previous digestion and chemical separation pass. Fig. 2. Summary of the abundance of 26 Mg (d 26 Mg * ) of pallasite olivine, ureilite and aubrite meteorites, terrestrial standards, and chondrite meteorites (Schiller et al., 2010a). The grey field represents the range of terrestrial mantle and basalt olivine Analysis of the Mg isotopic composition of J11 mantle and pallasite olivine samples after different stages of chemical separation of Mg Mg separated from both the J11 mantle and pallasite olivines was measured after different stages (anion exchange separation of Fe ± TODGA separation of Ca ± cation exchange separation of most major and trace elements ± cation exchange separation of Mn in 0.5 M HCl 95% acetone) of chemical separation of Mg to assess the extent to which the progressive removal of matrix elements affected the results. The results show that consistent results for the d 26 Mg * of both J11 and pallasite olivine are produced irrespective of the extent of the chemical separation procedures used to purify Mg, provided that Si (HF digestion) and Fe (anion exchange separation) have been removed from the olivine (Tables 1 and 3 and Fig. 4). When all data are combined, the average d 26 Mg * of both the terrestrial and pallasite olivines are marginally more positive ( ± &) when DSM-3 is used as the bracketing standard rather than Mg separated from J11 mantle olivine (Tables 1 and 3).

10 424 J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) Fig. 3. Variations in the abundance of 26 Mg (d 26 Mg * ) in isotopically fractionated (d 25 Mg) cuts of Mg produced by sequentially collecting Mg from cation exchange columns (AG50W-X8 resin). The line on the graph represents the expected (erroneous) d 26 Mg * values that would result from applying a kinetic/exponential mass bias (b = 0.511) correction to the isotopic analyses when, in fact, Mg has been fractionated on the cation exchange columns by an equilibrium (b = 0.521) fractionation process. Fig Al 26 Mg isochron diagram for pallasite olivine and ureilite meteorites. Isochrons are anchored with the mean values determined for non-cai bearing chondrites (Schiller et al., 2010a; 27 Al/ 24 Mg = and d 26 Mg * = ± &; 2 se). Initial 26 Al values are calculated using our conservative estimate of data reproducibility i.e., 1.5 times the 2 se uncertainty and only using data obtained versus the DSM-3 standard. Also shown is the regression through data for chondrite meteorites and CAIs (Schiller et al., 2010a). Fig. 4. d 26 Mg * values for terrestrial olivine (J11, JK3, JB281) and pallasite olivine (Admire) analysed after different chemical separation procedures and utilising different bracketing standards (i.e., DSM-3 [filled square symbols] and Mg separated from J11 mantle olivine [open square symbols]). a = anion exchange separation of Fe; c = anion exchange separation of Fe + TODGA separation of Ca + cation exchange separation of most major and trace elements (except Mn and Ni); Mn = anion exchange separation of Fe + TODGA separation of Ca + cation exchange separation of most major and trace elements + cation exchange separation of Mn in 0.5 M HCl 95% acetone Pallasites All of the olivine separates from the main group pallasites have 27 Al/ 24 Mg ratios that are effectively zero (<0.001). Analyses of the abundance of 26 Mg of the pallasite olivines repeatedly show resolvable deficits with respect to the terrestrial standard, whether bracketed by analyses of DSM-3 or Mg separated from J11 mantle olivine (Table 3 and Fig. 2), and irrespective of the chemical separation methodology used to purify Mg (Fig. 4). The mean d 26 Mg DSM-3 of all the pallasite olivine analyses is ± &, which is slightly less negative than the value obtained when pallasite olivines were measured against Mg separated from J11 mantle olivine i.e., d 26 Mg J11 ¼ 0:0162 0:0016&. There is no resolvable difference in d 26 Mg * values between olivines from the different main group pallasites. We use a model approach to calculate initial 26 Al/ 27 Al ratios of pallasites (and other meteorites) studied here by assuming that each meteorite group originated from a parent body that accreted from material with Al/Mg ratios and a present-day Mg isotope composition represented by non-cai bearing chondrite meteorites. A regression through the 27 Al/ 24 Mg d 26 Mg * pallasite olivine data including the average composition of non-cai-bearing chondrites (Schiller et al., 2010a) defines a line with a slope and initial ( 26 Al/ 27 Al) = 1.6 ± (Fig. 5). The stable Mg isotopic composition of olivine from the four main group pallasites show very limited variations from d 25 Mg DSM-3 = 0.05 ± 0.14& to 0.15 ± 0.11&, which are within error of the data previously reported by Handler et al. (2009) for pallasite olivine, and these overlap the values for olivine from Earth s upper mantle. These values are also identical to those reported by Teng et al. (2010) for a wide range of oceanic basalts, peridotite xenoliths and chondrite meteorites. Our d 25 Mg DSM-3 values for pallasite and mantle olivine are, however, inconsistent with those published by Chakrabarti and Jacobsen (2010) whose d 25 Mg DSM-3 values for pallasite and mantle olivine are systematically lighter due to as yet unknown analytical artefacts that have likely comprised the accuracy of Mg stable isotopic data presented by Chakrabarti and Jacobsen (2010) Ureilites Three of the ureilites yielded markedly sub-chondritic 27 Al/ 24 Mg ratios from to (El Gouanem, SAH98505, NWA2234) (Table 4). However, NWA766

11 26 Al 26 Mg deficit dating ultramafic meteorites 425 has a higher 27 Al/ 24 Mg (0.070) that may be consistent with the presence of high Si Al glass in this ureilite. All d 26 Mg * values measured in the ureilites whether bracketed by DSM-3 or Mg separated from J11 olivine yield slight deficits in the abundance of 26 Mg relative to the terrestrial standards, although in some cases (SAH98505 and NWA766 bracketed by DSM-3) these are just within 2 se analytical uncertainty of the terrestrial standard (Table 4 and Fig. 2). The mean d 26 Mg DSM-3 of all the ureilite analyses is ± & which is slightly less negative than when these samples were measured against Mg separated from J11 mantle olivine i.e., d 26 Mg J11 ¼ 0:0080 0:0023&. The d 26 Mg DSM-3 measured on ureilites in this study are within analytical uncertainty of those measured by Larsen et al. (2011) on two ureilites, including SAH A regression through the 27 Al/ 24 Mg d 26 Mg * ureilite data including the average composition of non- CAI-bearing chondrites (Schiller et al., 2010a) defines a line with a slope and initial ( 26 Al/ 27 Al) = 8.8 ± (Fig. 5). Stable Mg isotope data range from d 25 Mg DSM- 3 = 0.15 ± 0.13& to 0.34 ± 0.13& and overlap values measured for samples of Earth s mantle and basalts as well chondrites (Handler et al., 2009; Yang et al., 2009; Schiller et al., 2010a; Teng et al., 2010) Aubrites With the exception of Bishopville, the six bulk aubrite samples all have markedly sub-chondritic 27 Al/ 24 Mg ratios that range from (Norton County) to (Cumberland Falls) (Table 4). Bishopville has a 27 Al/ 24 Mg ratio (0.370) that is considerably higher than that measured on a larger bulk sample of this meteorite (0.039; Easton, 1985) suggesting that a feldspar-rich region of Bishopville was initially sampled in this study. Therefore, a second pyroxene-rich mineral separate of this meteorite was subsequently prepared, digested and analysed and yielded a lower 27 Al/ 24 Mg = The abundance of 26 Mg for all the aubrite samples, including the high 27 Al/ 24 Mg sample of Bishopville, are all within 2 se analytical uncertainties of the terrestrial standard used to bracket the analyses (DSM-3) (Table 4 and Fig. 2). The mean d 26 Mg DSM-3 of all the aubrite analyses is ± &. A regression through the 27 Al/ 24 Mg d 26 Mg * data including the average composition of non- CAI-bearing chondrites (Schiller et al., 2010a) defines a line with a slope of zero and maximum possible slope and initial ( 26 Al/ 27 Al) = (given the error on the regression). Stable Mg isotope data range from d 25 Mg DSM-3 = ± 0.11& to 0.18 ± 0.05& and overlap values measured for samples of Earth s mantle and basalts (Handler et al., 2009; Yang et al., 2009; Teng et al., 2010; Schiller et al., 2010a). 4. DISCUSSION 4.1. Precision and accuracy of d 26 Mg * data Resolving small deficits in 26 Mg abundances in meteorites for dating purposes requires a careful assessment as to whether it is possible to accurately and precisely measure d 26 Mg * values to <±0.005&. Multiple Mg isotope analyses (n = 10 40) of single samples, by pooling of analyses, which is a common practice in application of all short-lived chronometers to early Solar System chronometry (e.g., Lugmair and Shukolyukov, 1998; Kleine et al., 2005a,b; Markowski et al., 2006a,b; Wadhwa et al., 2009; Villeneuve et al., 2010; Bizzarro et al., 2011), can produce weighted mean d 26 Mg * values with internal 2 standard errors (se) as low as ±0.006& to 0.002& (Tables 1 4). The 2 se values quoted in this study include the uncertainties incorporated from the bracketing standards as well as that on the sample. Previous high precision Mg isotope studies of meteoritic material (e.g., Baker et al., 2005; Bizzarro et al., 2005) did not calculate d 26 Mg * as weighted means or incorporate uncertainties from the bracketing standard (or the sample) into the final 2 se and simply calculated the average and 2 se from the d 26 Mg * values as the mean of n measurements and 2 se = 2 sd/n. The difference in doing this ( average & 2 se ) as compared to the approach adopted here ( weighted mean & 2 se ) and in Schiller et al. (2010a,b) is evident from Tables 1, 3 and 4. While the two different approaches do not result in significant differences in the mean or average d 26 Mg * values of more than &, it is common for the quoted 2 se to vary significantly using the two different approaches. In about 55% of cases the 2 se calculated without incorporating the uncertainties from the sample and standard analyses ( average and 2 se ) is lower than that calculated using the weighted mean approach. In 30% of cases the calculated 2 se values are essentially the same (within 0.001&) irrespective of the method used to calculate them. We consider that the approach adopted herein provides a more realistic and conservative estimate of the analytical uncertainties on Mg isotope analysis by MC-ICPMS as the average and 2 se approach can yield overly optimistic estimates of 2 se, particularly where the number of repeat analyses (n) is small and, fortuitously, a small spread in individual d 26 Mg * values is obtained. The accuracy of the presented Mg isotope data can be assessed by the analyses of column-processed standards, gravimetrically 26 Mg spiked standards, terrestrial olivines and terrestrial and pallasite olivines subject to different amounts of chemical purification of Mg. In all cases, these data yield d 26 Mg * values that are, at worst (J11 digestion 3; Table 1), within 1.4 times the 2 se of the expected value and, apart from this example, all within 2 se uncertainties of the expected values. This demonstrates that the data are accurate to quoted uncertainties i.e., <1.5 times the final 2 se obtained on any particular sample. While it is possible to measure d 26 Mg * both precisely and accurately to < ± 0.005&, our analytical tests do reveal some artefacts that have potential to produce inaccurate data, although these are not important for the meteorite data obtained in this study. In particular, some isotopically fractionated and light ICP-MS Mg solution standards (Alfa Aesar and Aristar) have apparent excesses in d 26 Mg *. This reflects an artefact of these standards containing Mg that has, in part, experienced equilibrium stable isotopic fractionation. When an exponential/kinetic mass fractionation