Chronology of the angrite parent body and implications for core formation in protoplanets

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1 Available online at Geochimica et Cosmochimica Acta 84 (212) Chronology of the angrite parent body and implications for core formation in protoplanets Thorsten Kleine a,b,, Ulrik Hans b, Anthony J. Irving c, Bernard Bourdon b,d a Institut für Planetologie, Westfälische Wilhelms-Universität Münster, Wilhelm-Klemm-Str. 1, Münster, Germany b Institute of Geochemistry and Petrology, Department of Earth Sciences, ETH Zurich, Clausiusstrasse 25, CH-892 Zurich, Switzerland c Department of Earth and Space Sciences, University of Washington, Seattle, WA 98195, USA d Laboratoire de Géologie des Lyon, Ecole Normale Supérieure de Lyon, CNRS and UCBL, 46, Allée d Italie, Lyon cedex 7, France Received 1 February 211; accepted in revised form 23 January 212; available online 2 February 212 Abstract Angrites formed by some of the earliest igneous activity in the solar system and provide insights into the early stages of planetary melting and differentiation. Moreover, they are pivotal reference points for early solar system chronology. In order to study the processes and timescales of metal segregation in early protoplanets and to assess the distribution of short-lived radionuclides in the early solar system, the 182 Hf 182 W system was applied to a comprehensive suite of angrites. 182 Hf 182 W isochron ages for angrites are in excellent agreement with previously reported 27 Pb 26 Pb and 53 Mn 53 Cr results but are 1 Myr older than ages obtained from 26 Al 26 Mg chronometry. These inconsistencies probably reflect a disturbance of the Al Mg system in the angrite feldspars, but could alternatively be explained by a heterogeneous distribution of 26 Al in the early solar system. Based on the Hf W results four texturally and temporally resolved groups of angrites can be identified that were derived from at least two distinct mantle sources. These mantle sources are the result of separate events of core formation, both of which took place within 2 Myr of CAI formation. Thus, core formation in the angrite parent body did not occur as a single event of metal segregation from a global magma ocean but rather took place under varying conditions by several more local events. The disparate Hf W systematics of the two distinct angrite source regions indicate that convection in the magma ocean was inefficient in homogenizing the composition of the mantle, possibly as a result of a continuous bombardment with small planetesimals during ongoing core formation. Such impacts could have constantly removed primordial and earlier formed crust, facilitating rapid cooling of the magma ocean, which solidified as early as 3.6 ±.7 Myr after CAI formation. Ó 212 Elsevier Ltd. All rights reserved. 1. INTRODUCTION Angrites are a small but diverse group of refractory mafic to ultramafic meteorites that formed by melting in the mantle of a differentiated early solar system body. They are characterized by predominantly igneous textures (reflecting variable magmatic cooling rates) and unlike the Corresponding author at: Institut für Planetologie, Westfälische Wilhelms-Universität Münster, Wilhelm-Klemm-Str. 1, Münster, Germany. Tel.: (direct), (secretary). address: thorsten.kleine@uni-muenster.de (T. Kleine). more common eucrites show little evidence for impact and shock metamorphism (e.g., Mittlefehldt et al., 1998). Angrites have preserved a record of some of the earliest igneous activity in the solar system, and retain paleomagnetic evidence for an early dynamo in the metallic core of their parent body (Weiss et al., 28). Therefore they can be used to gain insights into the processes of accretion, differentiation and melting in some of the earliest protoplanets (e.g., Wasserburg et al., 1977; Lugmair and Galer, 1992; Nyquist et al., 1994; Markowski et al., 27; Weiss et al., 28; Schiller et al., 21). Moreover, angrites are pivotal reference points for early solar system chronology and ideally suited for assessing the distribution of short-lived /$ - see front matter Ó 212 Elsevier Ltd. All rights reserved. doi:1.116/j.gca

2 Chronology of the angrite parent body 187 radioisotopes in the early solar system (Lugmair and Galer, 1992; Markowski et al., 27; Spivak-Birndorf et al., 29; Nyquist et al., 29). Several long- and short-lived isotope systems have been applied to angrites. The results of these studies are generally in agreement with each other and indicate that magmatism on the angrite parent body occurred between 4 and 1 Myr after the beginning of the solar system (see summary in Nyquist et al., 29). However, the timescales of parent body accretion and differentiation are less well constrained. Based on 87 Rb 87 Sr systematics it was concluded that the angrite parent body accreted late, no earlier than 2 Myr after CAI formation (Lugmair and Galer 1992; Nyquist et al., 1994; Halliday and Porcelli, 21). However, based on 26 Al 26 Mg model ages for silicate differentiation combined with thermal modeling it was argued that accretion of the angrite parent body occurred within the first 2 Myr after CAI formation (e.g., Schiller et al., 21), inconsistent with the Rb Sr results. The short-lived 182 Hf 182 W system has proven well suited to determining the timescales of planetary accretion and differentiation. The fact that both Hf and W are refractory and have very different geochemical behavior during metal silicate separation renders this chronometer uniquely useful to study the timescales of core formation (Lee and Halliday, 1995; Harper and Jacobsen, 1996; Kleine et al., 29). As evident from strong 182 W deficits in most magmatic iron meteorites, core formation in their parent bodies must have occurred within 1 Myr after formation of Ca,Al-rich inclusions (CAI) (Kleine et al., 25a; Markowski et al., 26; Scherstén et al., 26; Burkhardt et al., 28; Qin et al., 28; Kleine et al., 29), 1 2 Myr earlier than formation of the ordinary chondrites (Kleine et al., 28; Kleine and Rudge, 211). The Hf W systematics of igneous achondrites are more difficult to interpret, as they reflect not only Hf/W fractionation during core formation but also additional fractionations during later melting and metamorphic processes (Quitté et al., 2; Kleine et al., 25b; Markowski et al. 27; Kleine et al., 29). Hafnium tungsten data for three angrites (D Orbigny, Sahara 99555, NWA 2999) reported by Markowski et al. (27) indicate that core formation in the angrite parent body took place within the first 4 Myr of the solar system. This age constraint is too imprecise, however, to distinguish between an early (<2 Myr after CAI formation) and late formation (>2 Myr after CAI formation) of the angrite parent body. Here we present internal Hf W mineral isochrons for eight angrites spanning the compositional, textural and temporal range of specimens known to date, and including new specimens identified since our initial studies (Markowski et al., 27; Kleine et al., 29). The sample suite includes finer grained angrites (Northwest Africa 1296, Sahara 99555, D Orbigny), coarser grained angrites (Angra dos Reis, Northwest Africa 459, Northwest Africa 481, Lewis Cliff 861), and a more metal-rich metaclastic angrite (paired specimens Northwest Africa 2999 and 4931). The new Hf W results are compared to age constraints from other chronometers and are used to assess the distribution of short-lived 26 Al, 53 Mn and 182 Hf in the early solar system. Such a comparison provides an important consistency test for the use of these short-lived radioisotopes as chronometers for early solar system processes. The Hf W results are then used to assess the chronology of core formation and silicate melting in the angrite parent body, and the process of metal segregation in the angrite parent body is evaluated. 2. PETROLOGY OF ANGRITE SPECIMENS The twelve known angrites are remarkable among achondrites not only because of their very ancient formation ages, but also because of their highly refractory bulk compositions, distinctive mineralogy, and wide variety of igneous textures. The angrites are tied together as a group particularly by their oxygen isotope compositions [plotting just below but parallel to the terrestrial fractionation line (e.g.,greenwood et al., 25)]. Yet for such a small group of related specimens, the textures range from dendritic/plumose (with or without macrocrysts) to subophitic (and vesicular) to plutonic cumulate (partially annealed) to metaclastic (with an exotic metal component). Detailed petrologic information on the specimens analyzed in this study is summarized here to provide context for organizing and understanding the chronologic results. Photomicrographs of seven of the studied specimens (all at the same scale) are presented in Fig. 1. The mineralogy of all angrites reflects their very refractory bulk compositions in that they are composed predominantly of Al Ti-rich clinopyroxene, pure anorthite, calcic olivine and usually kirschsteinite, with accessory phosphates (merrillite, Ca silicophosphate), troilite, kamacitic metal, and in some examples Fe Ti oxides (titanomagnetite or ulvöspinel), chromian spinel, rhönite, celsian, baddeleyite or glass. NWA 1296 is a very fine-grained, aphyric specimen with plumose to dendritic textures (Fig. 1a) suggestive of very rapid quenching from a high temperature melt. Sahara and D Orbigny (Fig. 1b, c) are both somewhat coarser grained, diabasic rocks with subophitic textures and prominent vesicles (indicative of a magmatic vapor phase of unknown composition). NWA 459 has a much coarser grained, plutonic igneous texture (Fig. 1e), and is characterized by strongly zoned pyroxenes and accessory rhönite and grain boundary glasses (Kuehner and Irving, 27a, b). Angra dos Reis (the only angrite witnessed to fall, in 1869) and NWA 481 (Fig. 1f) both have annealed plutonic cumulate textures, but Angra dos Reis is composed predominantly of Al Tirich clinopyroxene and is devoid of plagioclase, whereas NWA 481 has patches (possibly former xenolithic clasts) of recrystallized anorthite in addition to abundant clinopyroxene and contains calcic olivine (but no kirschsteinite). NWA 2999/4931 is very different from the other angrites in that it is an annealed breccia composed of metaplutonic igneous clasts (containing chromian spinel) and sparse angular anorthite grains, plus 8 volume% of exotic impactor metal (Humayun et al., 27); the metamorphism of this material was non-isochemical, and resulted in formation of complex disequilibrium corona and symplectite textures (see Kuehner et al., 26).

3 188 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (212) Fig. 1. Optical thin section images of angrite specimens. All images are in plane polarized light at the same scale (width of field = 9 mm). (a) Northwest Africa Very fine grained assemblage of skeletal Al Ti-rich clinopyroxene (tan), olivine + kirschsteinite (gray), interstitial anorthite (white) and oxide + sulfide + rare metal (black). (b) Sahara Medium grained subophitic texture. Compositionally zoned Al Ti-rich clinopyroxene (brown), anorthite (white laths), olivine + kirschsteinite (pale green, partly skeletal), and oxide + sulfide + rare metal (black). Note the small vesicles (center and lower center). (c) D Orbigny. Medium grained subophitic texture. Mineral phases are the same as for Sahara 99555, but some anorthite grains are more stubby. Note the vesicles at upper left and upper right. (d) Lewis Cliff 861. Coarse granular assemblage of compositionally zoned Al Ti-rich clinopyroxene (pink to light purplish brown), anorthite (white), olivine (gray, with thin linear kirschsteinite lamellae) and oxides + sulfide + rare metal (black). (e) Northwest Africa 459. Coarse granular assemblage of compositionally zoned Al Ti-rich clinopyroxene (pale brown to purple), olivine + kirschsteinite (pale green), anorthite (white) and oxide + sulfide + rare metal (black). (f) Northwest Africa 481. Coarse granular assemblage of Al Ti-rich clinopyroxene (tan), olivine (pale green), anorthite (white) and oxides + sulfide + rare metal (black). Note the large polygranular anorthite clast at upper right. (g) Northwest Africa Fragmental breccia crosscut by secondary (terrestrial) iron hydroxide veinlets. Primary silicate phases (olivine, Al Ti-rich clinopyroxene and anorthite) are all pale colored and difficult to distinguish. Black grains are metal (kamacite) and purplish brown grains are Cr Al-rich spinel.

4 Chronology of the angrite parent body ANALYTICAL METHODS 3.1. Sample preparation and chemical separation Samples were carefully cleaned with abrasive paper and washed with ethanol in an ultrasonic bath. Whole-rock powders were prepared from 25 to 7 mg chips when sufficient sample material was available. The remaining material was gently crushed in an agate mortar and separated in 4 2 lm and <4 lm fractions using nylon sieves. Mineral separates were prepared from the former fraction using heavy liquids, a Frantz magnetic separator and handpicking under a binocular microscope. Due to the limited amount of available material, not all mineral separates are of high purity. However, for each sample (except LEW 861) high-purity pyroxene separates were obtained. Samples were dissolved in pre-cleaned Savillex beakers using concentrated HF:HNO 3 (3:1) at 12 C. All the high-purity pyroxene separates were washed in cold 1 M HCl prior to dissolution to remove W-rich phosphates located at the rims of pyroxene grains. However, only the 1 M HCl leachate of pyroxenes from NWA 1296 was analyzed for its Hf/W ratio and W isotopic composition. After digestion, samples were dried and re-dissolved several times in HNO 3 H 2 O 2 to remove organic compounds. The samples were then completely dissolved in 6 M HCl.6 M HF and a 1% aliquot was spiked with a mixed 18 Hf 183 W tracer for Hf and W concentration measurements by isotope dilution (Kleine et al., 24). The chemical separation of Hf and W from the spiked aliquots was performed using ion exchange techniques described in Kleine et al. (24). The separation of W from the unspiked aliquots initially followed our previously established techniques [labeled method A in Table 2 (Kleine et al., 22, 24; Burkhardt et al., 28; Touboul et al., 29)] but during the later course of this study a newly developed technique was used (labeled method B in Table 2). An outline of the new separation scheme is summarized in Table 1. After aliquoting, the unspiked aliquot was dried, re-dissolved in 1 M HCl.1 M HF, and loaded onto cation exchange columns filled with 5 ml (wet volume) Bio- Rad AG5W8 resin (2 4 mesh). From this column, W (together with other high field strength elements) was collected with one resin volume of 1 M HCl.1 M HF (following Patchett and Tatsumoto, 198). This cut was dried, re-dissolved in.6 M HF-.4% H 2 O 2 and loaded on precleaned BioRad Ò PolyPrep columns filled with 1 ml (wet volume) AG 18 (2 4 mesh) resin. The matrix containing mostly Ti, Zr and Hf was eluted using 1 ml 1 M HCl 2% H 2 O 2 (following Quitté et al., 2; Scherstén et al., 26). Tungsten was then collected in 7 ml 6 M HCl 1 M HF. Methods A and B have been applied to different fractions from D Orbigny and Sahara (see Table 2) and results obtained using these two different methods agree very well, i.e., they plot on a single well-defined isochron (see Section 4.2). Data obtained for D Orbigny and Sahara by Markowski et al. (27) using yet another chemical separation method also plot on this isochron, demonstrating that all three methods provide accurate and reproducible Hf W data. Table 1 Two-column separation procedure for W isotope analyses. Step Volume (ml) Acid Column A: 5 ml BioRad 5W8 Clean 3 5 Alternating 6 M HCl and 2MHF Clean 5 1 M HCl.1 M HF Load sample + elute 2 1 M HCl.1 M HF HFSE Elute HFSE 5 1 M HCl.1 M HF Column B: 1 ml BioRad AG 18 Clean 1 3 M HNO 3 Clean 1 6 M HNO 3.2 M HF Clean 1 6 M HCl 1 M HF Equilibrate M HF Load HFSE cut 6.6 M HF.4% H 2 O 2 Rinse (Ti, Zr, Hf) 1 1 M HCl 2% H 2 O 2 Rinse 2 1 H 2 O Rinse 2 6 M HCl 1 M HF Elute W 8 6 M HCl 1 M HF The newly developed technique permits chemical separation of W with lower blanks (typically <1 pg) compared to those achieved using the old technique (typically 3 4 pg, sometimes >1 pg). The improved blanks are important for accurate W isotope ratio measurement of the W-poor and radiogenic pyroxenes because they minimize the need for blank correction. Throughout this study blanks were sufficiently low such that blank corrections were small or insignificant. In general, only corrections larger than.1 e (1 e =.1%) were made and this was necessary for only four samples (see Table 2). For all other samples blank corrections would have been smaller than 2 ppm and, hence, insignificant. For these samples no blank correction was made. For the four samples that were corrected for W blank, corrections range from.12 to.65 e and are smaller than the analytical uncertainties of the W isotope measurements. For instance, the largest blank correction of.65 e was necessary for D Orbigny px-2 but the analytical uncertainty on its 182 W/ 184 W ratio prior to blank correction was already 1.7 e. Usually two or three blanks were processed together with each set of samples and the average of the blanks for each session were used for the blank correction. An uncertainty of 5% was assumed for the blank correction because blanks in one analytical session can vary by up to 5%. The blank was assumed to have a terrestrial W isotopic composition because the main sources of the blank were the acids and the Savillex Teflon Isotope measurements All isotope measurements were performed using the Nu Plasma MC-ICPMS at ETH Zurich, equipped with a Cetac Aridus desolvating nebulizer. Prior to measurements samples were re-dissolved and dried several times in HNO 3 H 2 O 2 and then taken up in.3 1 ml.56 M

5 19 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (212) Table 2 Hf W data for angrites. Sample Weight (mg) Method a Blank b (pg) Hf (ppb) W (ppb) 18 Hf/ 184 W c ±2SD e 182 W d ±2SD e 183 W d ±2SD D Orbigny Fines 6.17 A ± ±.24.2 ±.15 wr B ± ±.5.1 ±.39 ol A ± 2.98 ±.5.32 ±.37 px A ± ± ±.2 px-2 (h.p.) 15.2 B ± ± 2. e 1.77 ±.99 px-3 (h.p.) 1.54 B ± ± 4.4 e 2.85 ± 2.25 Sahara wr B ± ±.41.7 ±.24 ol A ± ±.47.1 ±.25 px A ± ±.54 e.39 ±.69 px-2 (h.p.) B ± ± ±.3 NWA 1296 Fines B ± ±.3. ±.27 wr B ± ± ±.24 lm 28.3 B ± ± ±.71 px-1 (h.p.) B ± ± ±.3 px-2 (h.p) B ± ±.42.1 ±.26 px-w n/a B ± ± ±.26 NWA 459 Fines A ± 3.65 ±.43.6 ±.21 wr A ± ±.34.6 ±.29 Plag A ± ± ±.27 ol A ± ± ±.3 px (h.p.) A ± ±.46.7 ±.28 NWA 481 Fines A ± ±.18.2 ±.22 wr B ± 2.89 ±.19.3 ±.15 ol A ± ± ± ± ±.24 ol A ± ±.66.2 ±.5 px A ± ±.35.8 ± ±.29.8 ±.21 px A ± ± ±.39 px B ± ±.8.35 ±.35 px-4 (h.p.) B ± ± ±.59 LEW 861 ol A ± 3.74 ± ±.28 px A ± ±.3.11 ±.21 NWA 4931 wr B ± ±.66.7 ±.51 m ± ±.57.8 ±.34 nm B ± ± ± 1. Angra dos Reis Mix B ± ± ±.63 px (h.p.) B ± ± 1.8 e 2.28 ± 1.59 Notations: fines = 64 lm fraction, wr = whole-rock, ol = olivine, px = pyroxene, lm = light minerals, h.p. = high-purity, m = metal, nm = non-metal, px-w = wash fraction of pyroxenes. a Chemical separation procedure used for W isotope analyses; method A is described in Kleine et al. (24), method B is described in Table 1. b Total procedure W blanks for measurement of W isotopic compositions. c Uncertainties of the 18 Hf/ 184 W ratios refer to the last significant digits and are better than ± 1% (2r) in most cases. d e 182 W and e 183 W are the relative deviations from the terrestrial W isotope composition: e 18i W=[( 18i W/ 184 W) sample / ( 18i W/ 184 W) std. 1] 1 4. The uncertainties of the e 182 W and e 183 W values are 1.5 times the uncertainty of the mass spectrometric run (see text). e W isotopic composition corrected for blank. Corrections are.2 e for Sahara px-1;.65 e for D Orbigny px-2;.51 e for D Orbigny px-3; and.12 e for Angra dos Reis px.

6 Chronology of the angrite parent body 191 HNO 3.24 M HF. When sufficient amounts of W were available, isotope measurements were performed with an ion beam of Aon 182 W, which was obtained for 2 ppb W at an uptake rate of 1 ll/min. For many of the mineral separates, however, smaller quantities of W were available and the W isotope measurements were then performed with smaller ion beam intensities between A and Aon 182 W. Each measurement consisted of 12 s baseline integrations and up to 4 W isotope ratio measurements of 5 s each. Instrumental mass bias was corrected relative to 186 W/ 183 W = using the exponential law. Potential isobaric interferences of Os on masses 184 and 186 were corrected by monitoring 188 Os and were negligible for all samples. The 182 W/ 184 W and 183 W/ 184 W ratios of all samples were determined relative to two runs of an Alfa Aesar standard metal (Kleine et al. 24) bracketing the sample run and are reported in e 18i W units as the deviation of the 18i W/ 184 W ratio from the terrestrial standard value in parts per 1,.The reproducibility of the Alfa Aesar standard over one measurement session was always equal to or better than.3,.5, and 1 e units (2r) for the 2, 1, and 5 ppb W standards, respectively. The reproducibilities of the standards are equal to 1.5 times the uncertainty of an individual mass spectrometric run. For samples, the uncertainties of the e 182 W and e 183 W values, therefore, were calculated as 1.5 times the uncertainty of the mass spectrometric run. The accuracy of the 182 W/ 184 W measurements is monitored by also measuring the stable 183 W/ 184 W ratio. For almost all samples, the 183 W/ 184 W ratios agree with the terrestrial value (Table 2) but for four samples (D Orbigny px-2; NWA 459 plag; NWA 481 ol; Angra dos Reis px; see Table 2) higher 183 W/ 184 W are observed. Such spuriously high 183 W/ 184 W have been reported in many previous Hf W studies and are attributed to an organic interference on mass 183 that for all other samples was successfully removed by treatment with HNO 3 H 2 O 2. For samples with elevated 183 W/ 184 W, e 182 W values are calculated using the 182 W/ 184 W ratio normalized to 186 W/ 184 W = Previous studies on carbonaceous chondrites and eucrites have demonstrated that this approach provides accurate e 182 W values (e.g., Kleine et al. 24). For one pyroxene fraction from D Orbigny a lower-than-terrestrial 183 W/ 184 W ratio is observed. For this sample less than 1 ng W was available for W isotope measurement, resulting in an ion beam intensity of as little as Aon 182 W. At such low intensities we sometimes observed unidentified interferences on mass 184, leading to offsets in all isotope ratios involving 184 W. For these samples the 182 W/ 183 W ratio normalized to 186 W/ 183 W seems to be unaffected, so that for D Orbigny px-2 the e 182 W value was calculated using the 182 W/ 183 W ratio. Note that the e 182 W thus obtained for D Orbigny px-2 is consistent with the Hf W data for the other D Orbigny samples. For the Hf and W isotope dilution measurements, the uncertainties result from those in the measured 18 Hf/ 177 Hf and 183 W/ 184 W ratios of ±.2% (2r) and the correction for W blank (2 1 pg). Hf blanks were 61 pg and insignificant. The resulting uncertainty of the Hf/W ratio is better than ±1% (2r) in most cases Overview of Hf W data 4. RESULTS The Hf W data for mineral separates, whole-rocks and fines fractions of angrites are summarized in Table 2 and plotted in Fig. 2. Tungsten concentrations in the analyzed mineral fractions range from 56 ppb in a pyroxene from Angra dos Reis to 4 ppb in some fractions from NWA 481 and LEW 861. The D Orbigny and Sahara whole-rocks and the D Orbigny fines fraction display W contents within this range. The NWA 459 and NWA 481 whole rocks and fines have much higher W concentrations between 586 and 73 ppb and the NWA 481 fines stand out by having by far the highest W content of 3641 ppb. The origin of this enrichment in W is unclear but is probably related to the presence of W-rich phosphates in the fines fraction (Shirai and Humayun, 21). In the metal-bearing angrite NWA 4931 most of the W resides in metal (666 ppb), while a non-magnetic fraction of this meteorite displays the lowest W concentration of all samples analyzed for this study. The NWA 4931 wholerock contains 116 ppb W, about 6 times less than the NWA 481 whole-rock. The Hf concentrations are highly variable among angrite minerals and almost the entire Hf resides in high-ca pyroxenes. The Hf contents of the analyzed pyroxenes range from 2.4 to 5.1 ppm. Olivine separates and the NWA 4931 metal have much lower Hf concentrations ranging from 28 to 213 ppb. Wholerocks and fines have intermediate Hf contents ranging from 1.2 to 3.3 ppm. As for W, only the NWA 4931 wholerock displays a lower Hf content. The 18 Hf/ 184 W ratios of angrite whole-rocks are only modestly elevated compared to chondrites and range from 4.6 to 8, and a corresponding range in 182 W/ 184 W ratios from.9 to 5 e 182 W. However, the variability of Hf and W concentrations among the different angrite minerals led to a large range in 18 Hf/ 184 W ratios from.15 to 71, and a corresponding range in 182 W/ 184 W ratios from -1 to 56 e 182 W. This large range in 18 Hf/ 184 W and e 182 W makes it possible to calculate precise Hf W isochrons Isochron regressions All isochron regressions were calculated using the model 1 fit of IsoPlot. For D Orbigny and Sahara previously published Hf W data (Markowski et al., 27) were included in the regressions, as long as the samples were not leached. Data for leached samples were excluded from the regressions because the leaching may have induced a fractionation of Hf from W. The two most radiogenic pyroxene fractions from D Orbigny obtained in the present study have 183 W/ 184 W ratios different from the terrestrial value, which are attributed to small interferences on masses 183 or 184 (see Section 3 above). To assess whether the calculated e 182 W values of these two fractions from D Orbigny are accurate, the regression of the D Orbigny data were repeated without the two pyroxene fractions. Excluding only px-3 from the regression results in a 182 Hf/ 18 Hf of (7.17 ±.17) 1 5, identical to (7.15 ±.17) 1 5 ob-

7 192 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (212) tained by regressing all D Orbigny data. Excluding both px- 2 and px-3 from the regression results in a 182 Hf/ 18 Hf of (7.2 ±.2) 1 5, indistinguishable and only slightly less precise than the value obtained by regressing all the D Orbigny data. Consequently, in spite of the observed 183 W/ 184 W variations, the e 182 W values for the px-2 and px-3 fractions are accurate and fully consistent with the Hf W data for other fractions of D Orbigny. As shown in Fig. 2, the 18 Hf/ 184 W and 182 W/ 184 W ratios are correlated for all the analyzed angrites, such that precise isochrons could be obtained (MSWD in most cases). However, for NWA 459 and NWA 2999/4931 there is significant scatter on the isochron and several fractions including the whole-rocks and fines plot off the isochron. In these angrites, the Hf W system is disturbed, which may at least in part be due to terrestrial weathering. For these two angrites, only washed mineral separates were included in the isochron calculations. For NWA 2999/4931 previously published data for a metal and a pyroxene fraction (Markowski et al., 27) were included in the regression. The initial 182 Hf/ 18 Hf ratios of the analyzed angrites range from (7.15 ±.17) 1 5 to (4.2 ±.24) 1 5, corresponding to time intervals, Dt CAI, between 3.9 and 11.3 Myr after CAI formation using the initial 182 Hf/ 18 Hf of CAIs of (9.72 ±.44) 1 5 (Burkhardt et al., 28).These relative Hf W ages can be transformed to an absolute timescale using the Pb Pb ages of angrites. For reasons explained below, absolute ages, t D Orb, were calculated using a ±.3 Ma Pb Pb age of D Orbigny (Amelin, 28a; Bouvier and Wadhwa 21; Brennecka et al., 21b). The initial 182 Hf/ 18 Hf ratios of angrites along with their relative and absolute Hf W ages are summarized in Table Angrite groups Based on their Hf W systematics, the investigated angrites can be subdivided into four groups that also differ in their petrology (Table 4): (i) the more rapidly crystallized quenched angrites NWA 1296, D Orbigny and Sahara 99555, which have a fine-grained plumose/dendritic or subophitic texture, Hf W ages of Dt CAI 4 Myr and initial e 182 W values of about 2.4; (ii) the coarser grained plutonic angrites LEW 861, NWA 459 and NWA 481 with Hf W ages of Dt CAI 1 Myr and initial e 182 W values of about 1.5; (iii) the more metal-rich metaclastic angrite NWA 2999/4931 with a Hf W age of Dt CAI 7.5 Myr and a more elevated initial e 182 W of about +1.2; and (iv) Angra dos Reis which has a texture and Hf W age similar to those of the other plutonic angrites but a much higher initial e 182 W value of about Relative to the other angrites, NWA 2999/4931 and Angra dos Reis stand out by having much higher initial e 182 W values at a given 182 Hf/ 18 Hf. Both meteorites also share an unusual chemical composition that seems to require addition of a refractory component similar in composition to CAI (Longhi, 1999; Gellissen et al., 27). As such these two specimens may belong to a distinct group of angrites, i.e., they may derive from the same source (Table 4). 5. CHRONOMETRY OF ANGRITES Angrites are pivotal reference points in the chronology of the early solar system and are ideally suited for an intercalibration of different chronometers because they cooled so rapidly that potential differences in closure temperatures could not result in resolvable age differences among the various chronometers (Lugmair and Galer, 1992; Lugmair and Shukolyukov 1998). Furthermore, due to their high U/Pb ratios, precise Pb Pb ages are available for angrites, such that the relative ages obtained from short-lived systems can be linked to an absolute timescale. Pb Pb ages for angrites along with their relative ages obtained from the shortlived Al Mg, Mn Cr, and Hf W chronometers are summarized in Table Comparison of Hf W and Pb Pb ages Previous studies emphasized the good agreement between Hf W and Pb Pb ages for angrites but also noted that the absolute Hf W age of CAI, calculated relative to an Pb Pb age of ±.12 Ma for the D Orbigny angrite (Amelin, 28a), are Myr older than those obtained from Pb Pb chronometry (Burkhardt et al., 28; Kleine et al., 29; Nyquist et al., 29). However, these Pb Pb ages were calculated assuming that their 238 U/ 235 U ratios are invariant and identical to the terrestrial U isotope composition, for which a 238 U/ 235 U of was adopted (Amelin et al., 22; Connelly et al., 28; Amelin, 28a, b). However, the recent discovery of U isotope variations in CAI (Amelin et al., 21; Brennecka et al., 21a) and angrites (Brennecka et al., 21b; Amelin et al., 211; Brennecka and Wadhwa 211; Kaltenbach et al. 211) demonstrates that this assumption is not valid and that precise Pb Pb ages can only be obtained in concert with U isotope measurements. Such combined U and Pb isotope data are currently available for the forsterite-bearing Allende CAISJ11 (Amelin et al., 21) and most of the angrites. Importantly, all the angrites investigated so far (including most of the angrites examined in this study) have indistinguishable 238 U/ 235 U ratios, indicating that angrites are characterized by a uniform U isotopic composition with a 238 U/ 235 U ratio of ±.2 (Brennecka et al., 21b; Amelin et al., 211; Brennecka and Wadhwa, 211; Kaltenbach et al., 211; Larsen et al., 211). Published Pb Pb ages for angrites, which were calculated assuming a 238 U/ 235 U ratio of , can thus be re-calculated using the subsequently measured U isotopic composition of the angrites (Table 3). Nyquist et al. (29) demonstrated that in a plot of ln( 182 Hf/ 18 Hf) vs. Pb Pb age samples with concordant Hf W and Pb Pb ages plot along a line whose slope is given by the 182 Hf decay constant. Fig. 3 shows that for angrites the initial 182 Hf/ 18 Hf ratios correlate with their Pb Pb ages as predicted for decay of 182 Hf; the slope of the regression line of.76 ±.12 is consistent with the value of.78 ±.2 of the 182 Hf decay constant. An important observation from Fig. 3 is that CAI also plot on the correlation defined by the angrites. Owing to U isotopic variations among CAI there is currently uncertainty in the true

8 Chronology of the angrite parent body D ORBIGNY px 2 px 3 3 SAHARA px 2 2 wr px 1 fines ol m = (7.15 ±.17) x 1 5 i = 2.4 ±.2 MSWD = 1.1 Δt CAI = 3.9 ±.7 Myr t D Orb ±.3 Ma px 1 1 m = (6.83 ±.14) x 1 5 i = 2.1 ±.2 MSWD = 1.2 wr Δt CAI = 4.5 ±.7 Myr t D Orb = ±.5 Ma ol Hf/ 184 W 18 Hf/ 184 W NWA 1296 px 1 NWA 459 px 8 px fines m = (7.1 ±.28) x 1 lm 5 i = 2. ±.3 MSWD = 1.9 Δt CAI = 4.2 ±.8 Myr px w t D Orb = ±.7 Ma wr m = (4.63 ±.17) x 1 5 i = 1.2 ±.2 MSWD = 2. fines Δt CAI = 9.5 ±.8 Myr t D Orb = ±.7 Ma ol, plag Hf/ 184 W 18 Hf/ 184 W 15 6 NWA 481 LEW px 4 4 px 5 wr ol 1, ol 2, fines px 1 px 2 px 3 m = (4.52 ±.16) x 1 5 i = 1.5 ±.1 MSWD = 1.2 Δt CAI = 9.8 ±.8 Myr t D Orb = ±.7 Ma ol m = (4.8 ±.42) x 1 5 i = 1. ±.7 Δt CAI = 9. ± 1.3 Myr t D Orb = ± 1.2 Ma Hf/ 184 W 18 Hf/ 184 W 2 15 NWA 2999/4931 px nm 3 ANGRA DOS REIS px 1 2 mix 5 m wr Hf/ 184 W m = (5.43 ±.34) x 1 5 i = 1.2 ±.3 MSWD =.7 Δt CAI = 7.5 ± 1. Myr t D Orb = ±.9 Ma m = (4.2 ±.24) x i = 2.51 ±.45 MSWD =. Δt CAI = 11.3 ± 1. Myr wr t D Orb = ±.9 Ma Hf/ 184 W Fig. 2. Hf W isochrons for angrites. Previously reported data for angrites are shown with gray symbols and are from Markowski et al. (27). The Angra dos Reis whole-rock is from Quitté et al (2). All regressions were calculated using the model 1 fit of IsoPlot. m = initial 182 Hf/ 18 Hf, i = initial e 182 W. Time intervals relative to the formation of CAI, Dt CAI, were calculated using ( 182 Hf/ 18 Hf) i = (9.72 ±.44) 1 5 for CAI (Burkhardt et al., 28). Absolute ages, t D Orb, were calculated using the ±.3 Ma Pb Pb age for D Orbigny.

9 194 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (212) Table 3 Radiometric ages for angrites. Sample 182 Hf/ 18 Hf 1 5 (±2r) t D Orb (Ma) a (Hf W) t (Ma) b (Pb Pb) Dt CAI (Myr) (Hf W) Dt CAI (Myr) c (Pb Pb) Dt CAI (Myr) d (Al Mg) Dt D Orb (Myr) (Hf W) Dt D Orb (Myr) e (Mn Cr) D Orbigny 7.15 ± ± ± ± ±.6 5. ±.2 Sahara ± ± ± ± ±.6 5. ±.2.5 ±.4 NWA ± ± ±.8.3 ±.6 NWA ± ± ± ± ± ±.6 NWA ± ± ± ± ± ± ±.7 LEW ± ± ±.3 9. ± ± ± ±.6 NWA 2999/ ± ± ± ± ± ±.9 5. ± 1.1 Angra dos Reis 4.2 ± ± ± ± ± ±.8 a Absolute ages, t D Orb, are calculated relative to an Pb Pb age of ±.3 Ma. b Pb Pb ages were re-calculated using their measured U isotope composition (see text). c Calculated using the ±.5 Ma Pb Pb age of CAI SJ11 (Amelin et al., 21). d Al Mg data from Spivak-Birndorf et al. (29) and Schiller et al. (21). e Mn Cr data from Lugmair and Shukolyukov (1998), Shukolyukov and Lugmair (28), Shukolyukov et al. (29). Table 4 Characteristics of angrite groups. Group Samples Texture Dt CAI (Myr) (e 182 W) i Quenched angrites D Orbigny; Sah 99555; NWA 1296 Fine-grained, plumose/dendritic to subophitic Plutonic angrites NWA 459; NWA 481; LEW 861 Plutonic AdoR group Angra dos Reis Plutonic NWA 2999 /4931 Annealed breccia, meta-plutonic clasts, metal-rich ln( 182 Hf/ 18 Hf) slope = (λ 182 Hf) calc. =.76 ±.12 [(λ 182 Hf) meas. =.78 ±.2] Sahara D Orbigny NWA 459 Pb Pb age (Ma) NWA 481 LEW 861 NWA 2999/4931 Angra dos Reis CAI (SJ11) CAI (2364 B 1) Fig. 3. Measured 182 Hf/ 18 Hf and Pb Pb ages for angrites and CAI. The slope of the regression line corresponds to the 182 Hf decay constant and is in excellent agreement with the measured value of the 182 Hf decay constant, indicating that the Hf W and Pb Pb systems provide concordant ages for angrites and CAI. References for Pb Pb ages are given in the text and in Table 3. Pb Pb age of CAI (Brennecka et al., 21a). Therefore, ages for two CAI are shown in Fig. 3: (i) CAI SJ11 with a Pb Pb age of ±.5 Ma, calculated using its measured U isotopic composition (Amelin et al., 21); and (ii) CAI 2364 B-1 from the CV chondrite NWA 2364, for which Bouvier and Wadhwa (21) reported a Pb Pb age of ±.17 Ma. For this CAI the U isotopic composition was not measured, however, so that its Pb Pb age was calculated assuming that its 238 U/ 235 U ratio is identical to that of the SRM95a and 96 U isotopic standards (Bouvier and Wadhwa, 21). This assumption was based on the observations that (i) the Th/U ratio of CAI 2364 B-1 is very low and (ii) that other CAI with low Th/U show no 235 U excesses from the decay of 247 Cm (Brennecka et al. 21a). Note, however, that a later study identified a CAI with low Th/U but 238 U/ 235 U distinct from the value determined for the SRM95a and 96 U isotopic standards (Amelin et al., 21). Fig. 3 reveals that the Hf W data are consistent with both these Pb Pb ages for CAI because both ages plot within uncertainty of the 182 Hf/ 18 Hf vs. Pb Pb age correlation line defined by the angrites. This is further illustrated by calculating absolute Hf W ages for CAI using the relative Hf W and absolute Pb Pb ages of the angrites and the initial 182 Hf/ 18 Hf of CAI. These absolute Hf W ages for CAI average at ±.8 Ma, consistent with the Pb Pb ages for both CAI SJ11 and CAI 2364 B-1. Thus, the Hf W ages currently have insufficient precision to distinguish which of these two Pb Pb dates for CAI better represent their absolute age. However, the comparison of the relative ages, Dt CAI,of the individual angrites shows that the Hf W data are more consistent with an absolute age of ±.5 Ma for CAI, as determined for CAI SJ11. All the Pb Pb formation intervals of the angrites calculated relative to the Pb Pb age of this CAI are consistent with their Hf W formation intervals (Table 3). In contrast, the Pb Pb formation intervals calculated relative to a Pb Pb age of ±.17 Ma for CAI 2364 B-1 tend to be slightly longer than those obtained from the Hf W system. For

10 Chronology of the angrite parent body 195 instance, the Hf W age difference between D Orbigny and CAI of 3.9 ±.7 Myr is in excellent agreement with the Pb Pb age difference between D Orbigny and CAI SJ11 of 3.8 ±.6 Ma but only marginally overlaps with the Pb Pb age difference between D Orbigny and CAI 2364 B-1 of 4.9 ±.4 Myr. Likewise, for NWA 481 and LEW 861 the Hf W ages of 9.8 ±.8 and 9. ± 1.3 Myr are consistent with their Pb Pb age differences calculated relative to CAI SJ11 but are inconsistent with the Pb Pb differences of 11.2 ±.3 and 1.7 ±.3 Myr calculated for these samples relative to CAI 2364 B-1. These results suggest that the reported Pb Pb age of CAI 2364 B-1 might be slightly too old, perhaps due to unaccounted U isotopic variations in this CAI. Clearly, U isotopic data for this and other CAI are needed to more precisely define the absolute age of CAI and test the consistency between Pb Pb and Hf W ages of CAI Comparison of Hf W and Al Mg ages: crystallization age of D Orbigny and Sahara The comparison between Al Mg and Hf W ages can only be made for CAI and the oldest of the angrites (D Orbigny and Sahara 99555) because at the time the younger angrites crystallized, 26 Al was already extinct. The Al Mg systematics in D Orbigny and Sahara were investigated in detail by Spivak-Birndorf et al. (29) and Schiller et al. (21). A linear regression (calculated using IsoPlot) of the D Orbigny data from both studies yields an isochron (MSWD = 4.8) with an initial 26 Al/ 27 Al of (4.4 ±.55) 1 7. Regression of the Al Mg data for Sahara from Spivak-Birndorf et al. (29) and Schiller et al. (21) also yields an isochron (MSWD = 2.) whose initial 26 Al/ 27 Al of (4.5 ±.57) 1 7 is indistinguishable from that of D Orbigny. The initial 26 Al/ 27 Al of D Orbigny and Sahara correspond to formation intervals, Dt CAI, of 5. ±.2 Myr after CAI formation (Table 3). The uncertainty of this age is larger than those reported in earlier studies (e.g., Spivak-Birndorf et al., 29; Schiller et al., 21) because we included the uncertainty in the 26 Al half-life (t 1/2 =.73 ±.3; see Nyquist et al., 29). Spivak-Birndorf et al. (29) interpreted the Al Mg isochron ages to reflect the timing of crystallization of D Orbigny and Sahara However, Schiller et al. (21) argued that the Al Mg systematics in these angrites might be disturbed and that the crystallization of D Orbigny and Sahara is best dated by their Al Mg model ages of 3.9 ±.3 Myr and 4.1 ±.3 Myr, calculated based on small 26 Mg excesses in the D Orbigny and Sahara whole-rocks compared to chondrites. Based on the observation that the Pb Pb age of Sahara is 3.6 ±.5 Myr younger than that for CAI SJ11, Larsen et al. (211) argued that the 5 Myr Al Mg isochron age of Sahara reflects a lower-than-canonical initial 26 Al/ 27 Al ratio of the angrite parent body. However, others have interpreted the good agreement between the 5. ±.2 Myr Al Mg isochron ages for D Orbigny and Sahara and the difference between their Pb Pb ages and CAI 2364 B- 1 of 4.9 ±.4 Myr (D Orbigny) and 4.6 ±.4 Myr (Sahara 99555) as evidence that 26 Al was homogeneously distributed in the early solar system (Bouvier and Wadhwa, 21; Bouvier et al., 211). However, the validity of the Pb Pb age for CAI 2364 B-1 is uncertain because the U isotopic composition has not been measured for this CAI (see above), so that the significance of the apparent agreement between the Al Mg and Pb Pb ages (relative to CAI 2364 B-1) remains unclear. Overall, there is considerable uncertainty regarding the crystallization age of D Orbigny and Sahara based on their Al Mg systematics, and on the agreement (or disagreement) of the Al Mg and Pb Pb ages. Fig. 4 compares the aforementioned Al Mg and Pb Pb ages for D Orbigny and Sahara to their Hf W ages obtained in the present study. For D Orbigny, the Hf W interval of Dt CAI = 3.9 ±.7 Myr is shorter than the 5. ±.2 Myr interval obtained from the Al Mg isochron. However, the Hf W age agrees well with the 3.8 ±.6 Myr Pb Pb age difference between CAI SJ11 and D Orbigny, and with an Al Mg model age of 3.9 ±.3 Myr for the D Orbigny whole-rock (Schiller et al., 21). For Sahara the Hf W age overlaps with both the Al Mg isochron age and the Pb Pb formation interval calculated rel- Sahara & D Orbigny Sahara D Orbigny Δt CAI (Myr) (Al Mg isochron) (Al Mg mafic min. isochr.) (Hf W) (Al Mg isochron) (Al Mg model) (Pb Pb rel. CAI 2364 B-1) (Pb Pb rel. CAI SJ11) (Hf W) (Al Mg isochron) (Al Mg model) (Pb Pb rel. CAI 2364 B-1) (Pb Pb rel. CAI SJ11) (Hf W) Fig. 4. Comparison of relative Al Mg, Pb Pb and Hf W ages for angrites Sahara and D Orbigny. References for the Al Mg and Pb Pb ages are given in the text and in Table 3. Al Mg model ages are from Schiller et al. (21) and were calculated by assuming that the Al Mg systematics of the angrite whole-rocks reflect a single event of Al/Mg fractionation from a chondritic source. Al Mg isochron ages were recalculated from data in Spivak-Birndorf et al. (29) and Schiller et al. (21) using IsoPlot. Relative Pb Pb ages were calculated using two different CAI ages: ±.5 Ma for CAI SJ11 (Amelin et al., 21) and ±.2 Ma for CAI 2364 B-1 (Bouvier and Wadhwa, 21). The Hf W ages are in good agreement with Pb Pb ages calculated relative to CAI SJ11, with the Al Mg model ages, and with the Al Mg age obtained for an isochron regressed through the data for mafic minerals and wholerock samples. In contrast, Al Mg ages obtained from feldsparcontrolled isochrons are 1 Myr younger, indicating either a disturbance of the Al Mg system in the angrite feldspars or a heterogeneous distribution of 26 Al in the early solar system.

11 196 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (212) ative to CAI SJ11. Also plotted in Fig. 4 is the Hf W age obtained for a combined regression of the Hf W data for D Orbigny and Sahara Fig. 5 shows the results of this regression and reveals that all analyzed fractions from these two angrites plot on a single, well defined isochron (MSWD = 1.3). This indicates that both angrites crystallized contemporaneously from similar or identical magmas, consistent with the Al Mg data. The combined isochron yields a Hf W age of 4.2 ±.6 Myr after CAI formation, which we consider the best estimate for the crystallization age of D Orbigny and Sahara This Hf W age marginally overlaps the 5. ±.2 Myr Al Mg isochron age for these angrites but it is highly unlikely that these two ages are identical. There are three different scenarios that could account for the mismatch between the Hf W and Al Mg isochron ages for D Orbigny and Sahara 99555: (i) the Hf W formation interval is too short because the Hf W isochron for CAI does not reflect the 182 Hf/ 18 Hf at the time of CAI formation; (ii) at the time of CAI formation the precursor material of the angrites had a lower-than-canonical 26 Al/ 27 Al; (iii) the Al Mg systematics in Sahara and D Orbigny are disturbed so that the Al Mg isochrons do not reflect the time of crystallization. The first of these possibilities can be excluded because the Hf W data for CAI show no evidence for disturbance (Burkhardt et al., 28). There are small nucleosynthetic W isotope anomalies in some CAI and it is currently unclear how this affects the Hf W systematics. However, it is unlikely that nucleosynthetic anomalies will change the slope of the CAI isochron significantly. The second of the aforementioned possibilities is that the angrite parent body accreted from precursor material that initially had a lower 26 Al/ 27 Al than CAI. Using the Hf W formation interval of D Orbigny and Sahara of Dt CAI = 4.2 ±.6 Myr (see Fig. 5) and the initial 26 Al/ 27 Al obtained from a combined D Orbigny and Sahara isochron [ 26 Al/ 27 Al = (4.44 ±.33) 1 7 ; recalculated Sahara D Orbigny px 1 px Hf/ 184 W px 2 px 3 m = (6.99 ±.11) x 1 5 i = 2.24 ±.13 MSWD = 1.3 Δt CAI = 4.2 ±.6 Myr Fig. 5. Hf W isochron for D Orbigny and Sahara Regression (including data from Markowski et al., 27) calculated using the model 1 fit of IsoPlot. m = initial 182 Hf/ 18 Hf, i = initial e 182 W. For definition of Dt CAI see Fig. 2. All fractions from D Orbigny and Sahara plot on a single isochron, indicating that these two angrites formed contemporaneously from identical or similar magmas. from data in Spivak-Birndorf et al. (29) and Schiller et al. (21)] an initial 26 Al/ 27 Al of ( / 1.1) 1 5 at the time of CAI formation is calculated, lower than the initial 26 Al/ 27 Al of (5.23 ±.13) 1 5 that is characteristic for CAI (Jacobsen et al., 28). Such a low initial 26 Al/ 27 Al of the angrite precursor material is consistent with the value inferred by Larsen et al. (211) based on a correlation between 26 Mg excesses and nucleosynthetic 54 Cr anomalies. The third possibility that could account for the disparity in the calculated Al Mg and Hf W formation interval between CAI and the angrites D Orbigny and Sahara has been advanced by Schiller et al. (21), based on the observation that the Al Mg model ages of the D Orbigny and Sahara whole-rocks and an Al Mg isochron based on mafic minerals from these two angrites provide a better match to results from other chronometers as the feldspar-controlled Al Mg isochrons. However, Schiller et al. (21) also noted that there is no textural evidence for a disturbance of the Al Mg systematics in the angrites. Moreover, it is unclear as to whether the Al Mg wholerock model ages date crystallization of the angrites or rather are related to Al/Mg fractionation during magma ocean solidification, which predated the extrusion and crystallization of the angrite melts (Schiller et al., 21). Nevertheless, Tonui et al. (23) observed that Sm Nd data for a plagioclase separate from D Orbigny plot significantly above a 4.56 Ga isochron, indicating disturbed isotope systematics at least in the plagioclase. This and the observation that the Hf W isochron ages for D Orbigny and Sahara are in excellent agreement with both the Al Mg model and mafic mineral isochron ages (Fig. 4) suggests that the Al Mg system in the angrite feldspars is disturbed and that the feldspar-controlled Al Mg isochrons may not provide a reliable estimate of the 26 Al/ 27 Al at the time of angrite crystallization. In summary, the Hf W data strongly suggest that the crystallization age of the oldest angrites, D Orbigny and Sahara 99555, is 4.2 ±.6 Myr after CAI formation, i.e., older than the age inferred based on Al Mg chronometry (Spivak-Birndorf et al., 29; Bouvier and Wadhwa, 21). The mismatch between the Hf W and Al Mg ages for these two angrites may reflect either disturbed Al Mg systematics in angrite feldspars or a heterogeneous distribution of 26 Al in the early solar system. Distinguishing between these two possibilities is difficult and will require more precise constraints on the initial 182 Hf/ 18 Hf and the Pb Pb age of CAI, and a better understanding of any disturbance in the Al Mg systematics of the angrites. The disturbed Sm Nd systematics of D Orbigny plagioclase reveal in any case that this angrite may not be suitable as a common anchor for short-lived systems, and hence cannot be used to argue for or against 26 Al homogeneity in the early solar system Comparison of Hf W and Mn Cr ages In Fig. 6, initial 182 Hf/ 18 Hf ratios for angrites are plotted against their initial 53 Mn/ 55 Mn ratios. In this plot, samples having concordant Hf W and Mn Cr ages should plot on a single straight line, whose slope can be calculated from