Chronology of the Angrite Parent Body and Implications for Core Formation in Protoplanets

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Geochimica et Cosmochimica Acta 84 (2012) 186–203 www.elsevier.com/locate/gca

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 fu¨r Planetologie, Westfa¨lische Wilhelms-Universita¨tMu¨nster, Wilhelm-Klemm-Str. 10, 48149 Mu¨nster, Germany b Institute of Geochemistry and Petrology, Department of Earth Sciences, ETH Zurich, Clausiusstrasse 25, CH-8092 Zurich, Switzerland c Department of Earth and Space Sciences, University of Washington, Seattle, WA 98195, USA d Laboratoire de Ge´ologie des Lyon, Ecole Normale Supe´rieure de Lyon, CNRS and UCBL, 46, Alle´e d’Italie, 69364 Lyon cedex 7, France

Received 10 February 2011; accepted in revised form 23 January 2012; available online 2 February 2012

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 182Hf–182W system was applied to a comprehensive suite of angrites. 182Hf–182W isochron ages for angrites are in excellent agreement with previously reported 207Pb–206Pb and 53Mn–53Cr results but are 1 Myr older than ages obtained from 26Al–26Mg 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 26Al 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 bom- bardment 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 ± 0.7 Myr after CAI formation. Ó 2012 Elsevier Ltd. All rights reserved.

1. INTRODUCTION more common eucrites show little evidence for impact and shock metamorphism (e.g., Mittlefehldt et al., 1998). Angrites are a small but diverse group of refractory ma- Angrites have preserved a record of some of the earliest fic to ultramafic meteorites that formed by melting in the igneous activity in the solar system, and retain paleomag- mantle of a differentiated early solar system body. They netic evidence for an early dynamo in the metallic core of are characterized by predominantly igneous textures their parent body (Weiss et al., 2008). Therefore they can (reflecting variable magmatic cooling rates) and unlike the 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; Corresponding author at: Institut fu¨r Planetologie, Westfa¨lische Nyquist et al., 1994; Markowski et al., 2007; Weiss et al., Wilhelms-Universita¨tMu¨nster, Wilhelm-Klemm-Str. 10, 48149 2008; Schiller et al., 2010). Moreover, angrites are pivotal Mu¨nster, Germany. Tel.: +49 251 83 33406 (direct), +49 251 83 33496 (secretary). reference points for early solar system chronology and ide- E-mail address: [email protected] (T. Kleine). ally suited for assessing the distribution of short-lived

0016-7037/$ - see front matter Ó 2012 Elsevier Ltd. All rights reserved. doi:10.1016/j.gca.2012.01.032 Chronology of the angrite parent body 187 radioisotopes in the early solar system (Lugmair and Galer, bution of short-lived 26Al, 53Mn and 182Hf in the early solar 1992; Markowski et al., 2007; Spivak-Birndorf et al., 2009; system. Such a comparison provides an important consis- Nyquist et al., 2009). tency test for the use of these short-lived radioisotopes as Several long- and short-lived isotope systems have been chronometers for early solar system processes. The Hf–W applied to angrites. The results of these studies are generally results are then used to assess the chronology of core for- in agreement with each other and indicate that magmatism mation and silicate melting in the angrite parent body, on the angrite parent body occurred between 4 and and the process of metal segregation in the angrite parent 10 Myr after the beginning of the solar system (see sum- body is evaluated. mary in Nyquist et al., 2009). However, the timescales of parent body accretion and differentiation are less well con- 2. PETROLOGY OF ANGRITE SPECIMENS strained. Based on 87Rb–87Sr systematics it was concluded that the angrite parent body accreted late, no earlier than The twelve known angrites are remarkable among 2 Myr after CAI formation (Lugmair and Galer 1992; Ny- achondrites not only because of their very ancient forma- quist et al., 1994; Halliday and Porcelli, 2001). However, tion ages, but also because of their highly refractory bulk based on 26Al–26Mg model ages for silicate differentiation compositions, distinctive mineralogy, and wide variety of combined with thermal modeling it was argued that accre- igneous textures. The angrites are tied together as a group tion of the angrite parent body occurred within the first particularly by their oxygen isotope compositions [plotting 2 Myr after CAI formation (e.g., Schiller et al., 2010), just below but parallel to the terrestrial fractionation line inconsistent with the Rb–Sr results. (e.g.,Greenwood et al., 2005)]. Yet for such a small group The short-lived 182Hf–182W system has proven well sui- of related specimens, the textures range from dendritic/plu- ted to determining the timescales of planetary accretion mose (with or without macrocrysts) to subophitic (and and differentiation. The fact that both Hf and W are vesicular) to plutonic cumulate (partially annealed) to refractory and have very different geochemical behavior metaclastic (with an exotic metal component). Detailed pet- during metal–silicate separation renders this chronometer rologic information on the specimens analyzed in this study uniquely useful to study the timescales of core formation is summarized here to provide context for organizing and (Lee and Halliday, 1995; Harper and Jacobsen, 1996; understanding the chronologic results. Kleine et al., 2009). As evident from strong 182W deficits Photomicrographs of seven of the studied specimens (all in most magmatic iron meteorites, core formation in their at the same scale) are presented in Fig. 1. The mineralogy of parent bodies must have occurred within 1 Myr after all angrites reflects their very refractory bulk compositions formation of Ca,Al-rich inclusions (CAI) (Kleine et al., in that they are composed predominantly of Al–Ti-rich 2005a; Markowski et al., 2006; Schersteń et al., 2006; clinopyroxene, pure anorthite, calcic olivine and usually Burkhardt et al., 2008; Qin et al., 2008; Kleine et al., kirschsteinite, with accessory phosphates (merrillite, Ca sil- 2009), 1–2 Myr earlier than formation of the ordinary icophosphate), troilite, kamacitic metal, and in some exam- chondrites (Kleine et al., 2008; Kleine and Rudge, ples Fe–Ti oxides (titanomagnetite or ulvo¨spinel), chromian 2011). The Hf–W systematics of igneous achondrites are spinel, rho¨nite, celsian, baddeleyite or glass. NWA 1296 is a more difficult to interpret, as they reflect not only Hf/W very fine-grained, aphyric specimen with plumose to den- fractionation during core formation but also additional dritic textures (Fig. 1a) suggestive of very rapid quenching fractionations during later melting and metamorphic pro- from a high temperature melt. Sahara 99555 and D’Orbi- cesses (Quitte´ et al., 2000; Kleine et al., 2005b; Markow- gny (Fig. 1b, c) are both somewhat coarser grained, diaba- ski et al. 2007; Kleine et al., 2009). Hafnium–tungsten sic rocks with subophitic textures and prominent vesicles data for three angrites (D’Orbigny, Sahara 99555, NWA (indicative of a magmatic vapor phase of unknown compo- 2999) reported by Markowski et al. (2007) indicate that sition). NWA 4590 has a much coarser grained, plutonic core formation in the angrite parent body took place igneous texture (Fig. 1e), and is characterized by strongly within the first 4 Myr of the solar system. This age con- zoned pyroxenes and accessory rho¨nite and grain boundary straint is too imprecise, however, to distinguish between glasses (Kuehner and Irving, 2007a, b). Angra dos Reis (the an early (<2 Myr after CAI formation) and late forma- only angrite witnessed to fall, in 1869) and NWA 4801 tion (>2 Myr after CAI formation) of the angrite parent (Fig. 1f) both have annealed plutonic cumulate textures, body. but Angra dos Reis is composed predominantly of Al–Ti- Here we present internal Hf–W mineral isochrons for rich clinopyroxene and is devoid of plagioclase, whereas eight angrites spanning the compositional, textural and NWA 4801 has patches (possibly former xenolithic clasts) temporal range of specimens known to date, and including of recrystallized anorthite in addition to abundant clinopy- new specimens identified since our initial studies (Markow- roxene and contains calcic olivine (but no kirschsteinite). ski et al., 2007; Kleine et al., 2009). The sample suite in- NWA 2999/4931 is very different from the other angrites cludes finer grained angrites (Northwest Africa 1296, in that it is an annealed breccia composed of metaplutonic Sahara 99555, D’Orbigny), coarser grained angrites (Angra igneous clasts (containing chromian spinel) and sparse dos Reis, Northwest Africa 4590, Northwest Africa 4801, angular anorthite grains, plus 8 volume% of exotic impac- Lewis Cliff 86010), and a more metal-rich metaclastic ang- tor metal (Humayun et al., 2007); the metamorphism of this rite (paired specimens Northwest Africa 2999 and 4931). material was non-isochemical, and resulted in formation of The new Hf–W results are compared to age constraints complex disequilibrium corona and symplectite textures from other chronometers and are used to assess the distri- (see Kuehner et al., 2006). 188 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203

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 1296. 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 99555. 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 86010. 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 4590. 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 4801. 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 2999. 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. Chronology of the angrite parent body 189

3. ANALYTICAL METHODS Table 1 Two-column separation procedure for W isotope analyses. 3.1. Sample preparation and chemical separation Step Volume Acid (ml) Samples were carefully cleaned with abrasive paper and Column A: 5 ml BioRad 50WÂ8 washed with ethanol in an ultrasonic bath. Whole-rock Clean 3 Â 5 Alternating 6 M HCl and powders were prepared from 250 to 700 mg chips when suf- 2MHF ficient sample material was available. The remaining mate- Clean 5 1 M HCl – 0.1 M HF rial was gently crushed in an agate mortar and separated in Load sample + elute 2 1 M HCl – 0.1 M HF 40–200 lm and <40 lm fractions using nylon sieves. Min- HFSE eral separates were prepared from the former fraction using Elute HFSE 5 1 M HCl – 0.1 M HF heavy liquids, a Frantz magnetic separator and handpick- Column B: 1 ml BioRad AG 1Â8 ing under a binocular microscope. Due to the limited Clean 10 3 M HNO3 amount of available material, not all mineral separates Clean 10 6 M HNO3 – 0.2 M HF are of high purity. However, for each sample (except Clean 10 6 M HCl – 1 M HF LEW 86010) high-purity pyroxene separates were obtained. Equilibrate 2 Â 51MHF Samples were dissolved in pre-cleaned Savillex beakers Load HFSE cut 6 0.6 M HF – 0.4% H2O2 using concentrated HF:HNO (3:1) at 120 °C. All the Rinse (Ti, Zr, Hf) 10 1 M HCl – 2% H2O2 3 Rinse 2 Â 1HO high-purity pyroxene separates were washed in cold 1 M 2 Rinse 2 6 M HCl – 1 M HF HCl prior to dissolution to remove W-rich phosphates lo- Elute W 8 6 M HCl – 1 M HF cated 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 The newly developed technique permits chemical sepa- in HNO3–H2O2 to remove organic compounds. The ration of W with lower blanks (typically <10 pg) com- samples were then completely dissolved in 6 M HCl– pared to those achieved using the old technique 0.06 M HF and a 10% aliquot was spiked with a mixed 180 183 (typically 30–40 pg, sometimes >100 pg). The improved Hf– W tracer for Hf and W concentration measure- blanks are important for accurate W isotope ratio mea- ments by isotope dilution (Kleine et al., 2004). surement of the W-poor and radiogenic pyroxenes because The chemical separation of Hf and W from the spiked they minimize the need for blank correction. Throughout aliquots was performed using ion exchange techniques de- this study blanks were sufficiently low such that blank cor- scribed in Kleine et al. (2004). The separation of W from rections were small or insignificant. In general, only cor- the unspiked aliquots initially followed our previously rections larger than 0.1 e (1 e = 0.01%) were made and established techniques [labeled method A in Table 2 (Kleine this was necessary for only four samples (see Table 2). et al., 2002, 2004; Burkhardt et al., 2008; Touboul et al., For all other samples blank corrections would have been 2009)] but during the later course of this study a newly smaller than 2 ppm and, hence, insignificant. For these developed technique was used (labeled method B in Ta- samples no blank correction was made. For the four sam- ble 2). An outline of the new separation scheme is summa- ples that were corrected for W blank, corrections range rized in Table 1. After aliquoting, the unspiked aliquot was from 0.12 to 0.65 e and are smaller than the analytical dried, re-dissolved in 1 M HCl–0.1 M HF, and loaded onto uncertainties of the W isotope measurements. For in- cation exchange columns filled with 5 ml (wet volume) Bio- stance, the largest blank correction of 0.65 e was necessary Rad AG50WÂ8 resin (200–400 mesh). From this column, for D‘Orbigny px-2 but the analytical uncertainty on its W (together with other high field strength elements) was 182W/184W ratio prior to blank correction was already collected with one resin volume of 1 M HCl–0.1 M HF (fol- 1.7 e. Usually two or three blanks were processed together lowing Patchett and Tatsumoto, 1980). This cut was dried, with each set of samples and the average of the blanks for re-dissolved in 0.6 M HF-0.4% H2O2 and loaded on pre- Ò each session were used for the blank correction. An uncer- cleaned BioRad PolyPrep columns filled with 1 ml (wet tainty of 50% was assumed for the blank correction be- volume) AG 1Â8 (200–400 mesh) resin. The matrix con- cause blanks in one analytical session can vary by up to taining mostly Ti, Zr and Hf was eluted using 10 ml 50%. The blank was assumed to have a terrestrial W iso- 1 M HCl–2% H2O2 (following Quitte´ et al., 2000; Schersteń topic composition because the main sources of the blank et al., 2006). Tungsten was then collected in 7 ml 6 M HCl– were the acids and the Savillex Teflon. 1 M HF. Methods A and B have been applied to different fractions from D’Orbigny and Sahara 99555 (see Table 2) and results obtained using these two different methods 3.2. Isotope measurements agree very well, i.e., they plot on a single well-defined isochron (see Section 4.2). Data obtained for D’Orbigny and All isotope measurements were performed using the Nu Sahara 99555 by Markowski et al. (2007) using yet another Plasma MC-ICPMS at ETH Zurich, equipped with a Cetac chemical separation method also plot on this isochron, Aridus desolvating nebulizer. Prior to measurements demonstrating that all three methods provide accurate samples were re-dissolved and dried several times in and reproducible Hf–W data. HNO3–H2O2 and then taken up in 0.3–1 ml 0.56 M 190 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203

Table 2 Hf–W data for angrites. Sample Weight (mg) Methoda Blankb (pg) Hf (ppb) W (ppb) 180Hf/184Wc ±2SD e182Wd ±2SD e183Wd ±2SD D’Orbigny Fines 60.17 A 33 1231 282.5 5.14 ± 2 1.77 ± 0.24 À0.02 ± 0.15 wr 47.61 B 24 1580 238.7 7.81 ± 6 4.36 ± 0.50 À0.10 ± 0.39 ol 48.76 A 33 174.0 104.9 1.96 ± 2 À0.98 ± 0.50 À0.32 ± 0.37 px-1 71.24 A 33 2402 271.3 10.44 ± 5 6.07 ± 0.26 0.14 ± 0.20 px-2 (h.p.) 15.20 B 11 4070 67.47 71.2 ± 1.9 56.0 ± 2.0e 1.77 ± 0.99 px-3 (h.p.) 10.54 B 5 3725 64.23 68.4 ± 1.5 52.1 ± 4.4e À2.85 ± 2.25 Sahara 99555 wr 58.24 B 24 1503 294.0 6.03 ± 3 2.85 ± 0.41 0.07 ± 0.24 ol 104.85 A 175 239.7 191.5 1.48 ± 1 -0.89 ± 0.47 0.10 ± 0.25 px-1 48.74 A 175 4087 247.5 19.5 ± 3 13.08 ± 0.54e 0.39 ± 0.69 px-2 (h.p.) 32.83 B 11 3974 154.5 30.3 ± 2 21.80 ± 0.68 0.26 ± 0.30 NWA 1296 Fines 84.65 B 7 1739 407.6 5.04 ± 2 1.89 ± 0.30 0.00 ± 0.27 wr 117.77 B 7 1666 304.1 6.47 ± 3 2.09 ± 0.46 À0.18 ± 0.24 lm 28.30 B 7 1091 251.5 5.12 ± 2 2.15 ± 0.86 À0.54 ± 0.71 px-1 (h.p.) 106.36 B 7 1936 146.8 15.56 ± 6 10.92 ± 0.46 À0.22 ± 0.30 px-2 (h.p) 106.39 B 7 1957 184.9 12.49 ± 5 7.93 ± 0.42 À0.01 ± 0.26 px-w n/a B 7 230.4 1709 0.159 ± 1 À1.65 ± 0.37 À0.15 ± 0.26 NWA 4590 Fines 28.35 A 41 3351 730.0 5.42 ± 3 0.65 ± 0.43 0.60 ± 0.21 wr 244.42 A 41 2955 586.3 5.95 ± 2 3.29 ± 0.34 À0.06 ± 0.29 Plag 69.63 A 41 86.33 134.2 0.759 ± 4 À1.02 ± 0.37 1.39 ± 0.27 ol 46.26 A 41 27.97 226.9 0.145 ± 1 À1.03 ± 0.23 À0.18 ± 0.30 px (h.p.) 41.41 A 41 5136 235.0 25.8 ± 1 12.61 ± 0.46 À0.07 ± 0.28 NWA 4801 Fines 130.54 A 41 2290 3641 0.742 ± 3 À1.05 ± 0.18 0.02 ± 0.22 wr 236.53 B 24 2572 661.5 4.59 ± 2 0.89 ± 0.19 0.03 ± 0.15 ol 89.39 A 41 213.1 390.9 0.643 ± 3 À1.01 ± 0.36 1.16 ± 0.19 À0.96 ± 0.40 1.07 ± 0.24 ol-2 15.88 A 41 165.1 322.4 0.604 ± 4 À1.22 ± 0.66 À0.20 ± 0.50 px-1 91.24 A 41 3725 379.4 11.59 ± 5 4.45 ± 0.35 0.08 ± 0.22 4.67 ± 0.29 0.08 ± 0.21 px-2 23.00 A 41 3756 386.4 11.47 ± 7 4.22 ± 0.37 À0.17 ± 0.39 px-3 28.73 B 11 3449 232.3 17.5 ± 1 7.76 ± 0.80 0.35 ± 0.35 px-4 (h.p.) 47.98 B 5 3125 168.3 21.9 ± 1 10.71 ± 0.69 0.16 ± 0.59 LEW 86010 ol 25.88 A 25 149.8 404.2 0.437 ± 3 À0.74 ± 0.65 0.55 ± 0.28 px 48.92 A 25 3880 394.8 11.6 ± 6 4.89 ± 0.30 0.11 ± 0.21 NWA 4931 wr 51.26 B 24 790.9 116.2 8.03 ± 9 5.07 ± 0.66 À0.07 ± 0.51 m 24.09 20 176.0 665.7 0.312 ± 1 1.13 ± 0.57 À0.08 ± 0.34 nm 44.64 B 11 838.6 37.98 26.1 ± 3 17.4 ± 1.4 0.42 ± 1.00 Angra dos Reis Mix 44.25 B 5 2504 93.27 31.7 ± 2 17.2 ± 1.1 À0.01 ± 0.63 px (h.p.) 29.36 B 5 2628 56.11 55.3 ± 4 28.2 ± 1.8e 2.28 ± 1.59 Notations: ﬁnes = 640 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. (2004), method B is described in Table 1. b Total procedure W blanks for measurement of W isotopic compositions. c Uncertainties of the 180Hf/184W ratios refer to the last signiﬁcant digits and are better than ± 1% (2r) in most cases. d 182 183 18i 18i 184 e W and e W are the relative deviations from the terrestrial W isotope composition: e W=[( W/ W)sample/ 18i 184 4 182 183 ( W/ W)std. À 1] Â 10 . The uncertainties of the e W and e W values are 1.5 times the uncertainty of the mass spectrometric run (see text). e W isotopic composition corrected for blank. Corrections are 0.20 e for Sahara 99555 px-1; 0.65 e for D’Orbigny px-2; 0.51 e for D’Orbigny px-3; and 0.12 e for Angra dos Reis px. Chronology of the angrite parent body 191

HNO3–0.24 M HF. When sufficient amounts of W were 4. RESULTS available, isotope measurements were performed with an ion beam of 2 Â 10À11 Aon182W, which was obtained 4.1. Overview of Hf–W data for 20 ppb W at an uptake rate of 100 ll/min. For many of the mineral separates, however, smaller quantities of W The Hf–W data for mineral separates, whole-rocks and were available and the W isotope measurements were then fines fractions of angrites are summarized in Table 2 and performed with smaller ion beam intensities between plotted in Fig. 2. Tungsten concentrations in the analyzed 2 Â 10À12 A and 1 Â 10À11 Aon182W. Each measure- mineral fractions range from 56 ppb in a pyroxene from ment consisted of 120 s baseline integrations and up to Angra dos Reis to 400 ppb in some fractions from 40 W isotope ratio measurements of 5 s each. Instrumental NWA 4801 and LEW 86010. The D‘Orbigny and Sahara mass bias was corrected relative to 186W/183W = 1.9859 99555 whole-rocks and the D‘Orbigny fines fraction display using the exponential law. Potential isobaric interferences W contents within this range. The NWA 4590 and NWA of Os on masses 184 and 186 were corrected by monitoring 4801 whole rocks and fines have much higher W concentra- 188Os and were negligible for all samples. The 182W/184W tions between 586 and 730 ppb and the NWA 4801 fines and 183W/184W ratios of all samples were determined rela- stand out by having by far the highest W content of tive to two runs of an Alfa Aesar standard metal (Kleine 3641 ppb. The origin of this enrichment in W is unclear et al. 2004) bracketing the sample run and are reported in but is probably related to the presence of W-rich phos- e18iW units as the deviation of the 18iW/184W ratio from phates in the fines fraction (Shirai and Humayun, 2010). the terrestrial standard value in parts per 10,000.The repro- In the metal-bearing angrite NWA 4931 most of the W re- ducibility of the Alfa Aesar standard over one measurement sides in metal (666 ppb), while a non-magnetic fraction of session was always equal to or better than 0.3, 0.5, and this meteorite displays the lowest W concentration of all 1 e units (2r) for the 20, 10, and 5 ppb W standards, samples analyzed for this study. The NWA 4931 whole- respectively. The reproducibilities of the standards are rock contains 116 ppb W, about 6 times less than the equal to 1.5 times the uncertainty of an individual mass NWA 4801 whole-rock. The Hf concentrations are highly spectrometric run. For samples, the uncertainties of the variable among angrite minerals and almost the entire Hf e182W and e183W values, therefore, were calculated as 1.5 resides in high-Ca pyroxenes. The Hf contents of the ana- times the uncertainty of the mass spectrometric run. lyzed pyroxenes range from 2.4 to 5.1 ppm. Olivine sep- The accuracy of the 182W/184W measurements is moni- arates and the NWA 4931 metal have much lower Hf tored by also measuring the stable 183W/184W ratio. For al- concentrations ranging from 28 to 213 ppb. Whole- most all samples, the 183W/184W ratios agree with the rocks and fines have intermediate Hf contents ranging from terrestrial value (Table 2) but for four samples (D’Orbigny 1.2 to 3.3 ppm. As for W, only the NWA 4931 whole- px-2; NWA 4590 plag; NWA 4801 ol; Angra dos Reis px; rock displays a lower Hf content. see Table 2) higher 183W/184W are observed. Such spuri- The 180Hf/184W ratios of angrite whole-rocks are only ously high 183W/184W have been reported in many previous modestly elevated compared to chondrites and range from Hf–W studies and are attributed to an organic interference 4.6 to 8, and a corresponding range in 182W/184W ratios on mass 183 that for all other samples was successfully re- from 0.9 to 5 e182W. However, the variability of Hf and moved by treatment with HNO3–H2O2. For samples with W concentrations among the different angrite minerals led elevated 183W/184W, e182W values are calculated using the to a large range in 180Hf/184W ratios from 0.15 to 71, 182W/184W ratio normalized to 186W/184W = 0.92767. Pre- and a corresponding range in 182W/184W ratios from -1 vious studies on carbonaceous chondrites and eucrites have to 56 e182W. This large range in 180Hf/184W and e182W demonstrated that this approach provides accurate e182W makes it possible to calculate precise Hf–W isochrons. values (e.g., Kleine et al. 2004). For one pyroxene fraction from D’Orbigny a lower-than-terrestrial 183W/184W ratio is 4.2. Isochron regressions observed. For this sample less than 1 ng W was available for W isotope measurement, resulting in an ion beam inten- All isochron regressions were calculated using the model sity of as little as 1 Â 10À12 Aon182W. At such low inten- 1 fit of IsoPlot. For D’Orbigny and Sahara 99555 previ- sities we sometimes observed unidentified interferences on ously published Hf–W data (Markowski et al., 2007) were mass 184, leading to offsets in all isotope ratios involving included in the regressions, as long as the samples were 184W. For these samples the 182W/183W ratio normalized not leached. Data for leached samples were excluded from to 186W/183W seems to be unaffected, so that for D’Orbigny the regressions because the leaching may have induced a px-2 the e182W value was calculated using the 182W/183W fractionation of Hf from W. The two most radiogenic ratio. Note that the e182W thus obtained for D’Orbigny pyroxene fractions from D’Orbigny obtained in the present px-2 is consistent with the Hf–W data for the other study have 183W/184W ratios different from the terrestrial D’Orbigny samples. value, which are attributed to small interferences on masses For the Hf and W isotope dilution measurements, the 183 or 184 (see Section 3 above). To assess whether the cal- uncertainties result from those in the measured 180Hf/177Hf culated e182W values of these two fractions from D’Orbigny and 183W/184W ratios of ±0.2% (2r) and the correction for are accurate, the regression of the D’Orbigny data were re- W blank (2–10 pg). Hf blanks were 61 pg and insignificant. peated without the two pyroxene fractions. Excluding only The resulting uncertainty of the Hf/W ratio is better than px-3 from the regression results in a182Hf/180Hf of ±1% (2r) in most cases. (7.17 ± 0.17) Â 10À5, identical to (7.15 ± 0.17) Â 10À5 ob- 192 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203 tained by regressing all D’Orbigny data. Excluding both px- 5. CHRONOMETRY OF ANGRITES 2 and px-3 from the regression results in a182Hf/180Hf of (7.20 ± 0.20) Â 10À5, indistinguishable and only slightly Angrites are pivotal reference points in the chronology less precise than the value obtained by regressing all the of the early solar system and are ideally suited for an inter- D’Orbigny data. Consequently, in spite of the observed calibration of different chronometers because they cooled 183W/184W variations, the e182W values for the px-2 and so rapidly that potential differences in closure temperatures px-3 fractions are accurate and fully consistent with the could not result in resolvable age differences among the var- Hf–W data for other fractions of D’Orbigny. ious chronometers (Lugmair and Galer, 1992; Lugmair and As shown in Fig. 2, the 180Hf/184W and 182W/184W ra- Shukolyukov 1998). Furthermore, due to their high U/Pb tios are correlated for all the analyzed angrites, such that ratios, precise Pb–Pb ages are available for angrites, such precise isochrons could be obtained (MSWD 6 1.2 in most that the relative ages obtained from short-lived systems cases). However, for NWA 4590 and NWA 2999/4931 there can be linked to an absolute timescale. Pb–Pb ages for ang- is significant scatter on the isochron and several fractions rites along with their relative ages obtained from the short- including the whole-rocks and fines plot off the isochron. lived Al–Mg, Mn–Cr, and Hf–W chronometers are summa- In these angrites, the Hf–W system is disturbed, which rized in Table 3. may at least in part be due to terrestrial weathering. For these two angrites, only washed mineral separates were in- 5.1. Comparison of Hf–W and Pb–Pb ages cluded in the isochron calculations. For NWA 2999/4931 previously published data for a metal and a pyroxene frac- Previous studies emphasized the good agreement be- tion (Markowski et al., 2007) were included in the tween Hf–W and Pb–Pb ages for angrites but also noted regression. that the absolute Hf–W age of CAI, calculated relative to The initial 182Hf/180Hf ratios of the analyzed angrites an Pb–Pb age of 4564.42 ± 0.12 Ma for the D’Orbigny ang- range from (7.15 ± 0.17) Â 10À5 to (4.02 ± 0.24) Â 10À5, rite (Amelin, 2008a), are 1–1.5 Myr older than those ob- corresponding to time intervals, DtCAI, between 3.9 and tained from Pb–Pb chronometry (Burkhardt et al., 2008; 11.3 Myr after CAI formation using the initial Kleine et al., 2009; Nyquist et al., 2009). However, these 182Hf/180Hf of CAIs of (9.72 ± 0.44) Â 10À5(Burkhardt Pb–Pb ages were calculated assuming that their 238U/235U et al., 2008).These relative Hf–W ages can be transformed ratios are invariant and identical to the terrestrial U isotope to an absolute timescale using the Pb–Pb ages of angrites. composition, for which a 238U/235U of 137.88 was adopted For reasons explained below, absolute ages, tD’Orb, were (Amelin et al., 2002; Connelly et al., 2008; Amelin, 2008a, calculated using a 4563.4 ± 0.3 Ma Pb–Pb age of D’Orbi- b). However, the recent discovery of U isotope variations gny (Amelin, 2008a; Bouvier and Wadhwa 2010; Brennecka in CAI (Amelin et al., 2010; Brennecka et al., 2010a) and et al., 2010b). The initial 182Hf/180Hf ratios of angrites angrites (Brennecka et al., 2010b; Amelin et al., 2011; along with their relative and absolute Hf–W ages are sum- Brennecka and Wadhwa 2011; Kaltenbach et al. 2011) dem- marized in Table 3. onstrates that this assumption is not valid and that precise Pb–Pb ages can only be obtained in concert with U isotope 4.3. Angrite groups measurements. Such combined U and Pb isotope data are currently available for the forsterite-bearing Allende Based on their Hf–W systematics, the investigated ang- CAISJ101 (Amelin et al., 2010) and most of the angrites. rites can be subdivided into four groups that also differ in Importantly, all the angrites investigated so far (including their petrology (Table 4): (i) the more rapidly crystallized most of the angrites examined in this study) have indistin- “quenched” angrites NWA 1296, D‘Orbigny and Sahara guishable 238U/235U ratios, indicating that angrites are 99555, which have a fine-grained plumose/dendritic or characterized by a uniform U isotopic composition with a 238 235 subophitic texture, Hf–W ages of DtCAI 4 Myr and initial U/ U ratio of 137.78 ± 0.02 (Brennecka et al., 2010b; e182W values of about À2.4; (ii) the coarser grained “plu- Amelin et al., 2011; Brennecka and Wadhwa, 2011; Kalten- tonic” angrites LEW 86010, NWA 4590 and NWA 4801 bach et al., 2011; Larsen et al., 2011). Published Pb–Pb ages 182 238 235 with Hf–W ages of DtCAI 10 Myr and initial e W values for angrites, which were calculated assuming a U/ U of about À1.5; (iii) the more metal-rich metaclastic angrite ratio of 137.88, can thus be re-calculated using the subse- NWA 2999/4931 with a Hf–W age of DtCAI 7.5 Myr and a quently measured U isotopic composition of the angrites more elevated initial e182W of about +1.2; and (iv) Angra (Table 3). dos Reis which has a texture and Hf–W age similar to those Nyquist et al. (2009) demonstrated that in a plot of of the other plutonic angrites but a much higher initial ln(182Hf/180Hf) vs. Pb–Pb age samples with concordant e182W value of about +2.5. Relative to the other angrites, Hf–W and Pb–Pb ages plot along a line whose slope is given NWA 2999/4931 and Angra dos Reis stand out by having by the 182Hf decay constant. Fig. 3 shows that for angrites much higher initial e182W values at a given 182Hf/180Hf. the initial 182Hf/180Hf ratios correlate with their Pb–Pb ages Both meteorites also share an unusual chemical composi- as predicted for decay of 182Hf; the slope of the regression tion that seems to require addition of a refractory compo- line of 0.076 ± 0.012 is consistent with the value of nent similar in composition to CAI (Longhi, 1999; 0.078 ± 0.002 of the 182Hf decay constant. An important Gellissen et al., 2007). As such these two specimens may be- observation from Fig. 3 is that CAI also plot on the corre- long to a distinct group of angrites, i.e., they may derive lation defined by the angrites. Owing to U isotopic varia- from the same source (Table 4). tions among CAI there is currently uncertainty in the true Chronology of the angrite parent body 193

60 30 D’ORBIGNY px−2 SAHARA 99555 px−3

40 20 px−2 W W 182 182 px−1 ε 20 ε 10 m = (7.15 ± 0.17) x 10−5 m = (6.83 ± 0.14) x 10−5 i = −2.4 ± 0.2 i = −2.1 ± 0.2 wr MSWD = 1.1 MSWD = 1.2 px−1 ΔtCAI = 3.9 ± 0.7 Myr wr ΔtCAI = 4.5 ± 0.7 Myr 0 fines 0 ol tD’Orb ≡ 4563.4 ± 0.3 Ma tD’Orb = 4562.8 ± 0.5 Ma ol 0 204060800 10203040 180Hf/184W 180Hf/184W 15

12 NWA 1296 NWA 4590 px−1 px

10 8 px−2 W W 182 4 182 ε ε 5 fines m = (7.01 ± 0.28) x 10−5 wr m = (4.63 ± 0.17) x 10−5 lm i = −2.0 ± 0.3 i = −1.2 ± 0.2 0 MSWD = 1.9 MSWD = 2.0 ΔtCAI = 4.2 ± 0.8 Myr 0 fines ΔtCAI = 9.5 ± 0.8 Myr px−w tD’Orb = 4563.1 ± 0.7 Ma tD’Orb = 4557.8 ± 0.7 Ma ol, plag –4 0 5 10 15 20 0 102030 180Hf/184W 180Hf/184W 15 6 NWA 4801 LEW 86010

px px−4 4 10

px−3 W W 2

182 px−1 182

ε 5 ε px−2 m = (4.52 ± 0.16) x 10−5 m = (4.80 ± 0.42) x 10−5 i = −1.5 ± 0.1 0 i = −1.0 ± 0.7 MSWD = 1.2 ΔtCAI = 9.0 ± 1.3 Myr wr Δt t 0 CAI = 9.8 ± 0.8 Myr ol D’Orb = 4558.3 ± 1.2 Ma t = 4557.5 ± 0.7 Ma ol−1, ol−2, fines D’Orb –2 0 5 10 15 20 25 04812 180Hf/184W 180Hf/184W 20

NWA 2999/4931 px−nm ANGRA DOS REIS 30 px 15

10 20 W W mix 182 182 ε ε −5 −5 5 m = (5.43 ± 0.34) x 10 m = (4.02 ± 0.24) x 10 wr i = 1.2 ± 0.3 10 i = 2.51 ± 0.45 MSWD = 0.7 MSWD = 0.0 ΔtCAI = 7.5 ± 1.0 Myr ΔtCAI = 11.3 ± 1.0 Myr m wr 0 tD’Orb = 4559.8 ± 0.9 Ma tD’Orb = 4556.0 ± 0.9 Ma 0 01020300 204060 180Hf/184W 180Hf/184W

Fig. 2. Hf–W isochrons for angrites. Previously reported data for angrites are shown with gray symbols and are from Markowski et al. (2007). The Angra dos Reis whole-rock is from Quitte´ et al (2000). All regressions were calculated using the model 1 ﬁt of IsoPlot. m = initial 182 180 182 Hf/ Hf, i = initial e W. Time intervals relative to the formation of CAI, DtCAI, were calculated using 182 180 À5 ( Hf/ Hf)i = (9.72 ± 0.44) Â 10 for CAI (Burkhardt et al., 2008). Absolute ages, tD’Orb, were calculated using the 4563.4 ± 0.3 Ma Pb–Pb age for D’Orbigny. 194 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203

Table 3 Radiometric ages for angrites. 182 180 a b c d e Sample Hf/ Hf tD’Orb (Ma) t (Ma) DtCAI (Myr) DtCAI (Myr) DtCAI (Myr) DtD’Orb (Myr) DtD’Orb (Myr) Â 105 (±2r) (Hf–W) (Pb–Pb) (Hf–W) (Pb–Pb) (Al–Mg) (Hf–W) (Mn–Cr) D’Orbigny 7.15 ± 0.17 4563.4 ± 0.3 4563.4 ± 0.3 3.9 ± 0.7 3.8 ± 0.6 5.0 ± 0.2 Sahara 99555 6.87 ± 0.15 4562.8 ± 0.5 4563.8 ± 0.3 4.5 ± 0.7 3.4 ± 0.6 5.0 ± 0.2 0.5 ± 0.4 NWA 1296 7.01 ± 0.28 4563.1 ± 0.7 4.2 ± 0.8 0.3 ± 0.6 NWA4590 4.63 ± 0.17 4557.8 ± 0.7 4557.8 ± 0.4 9.5 ± 0.8 9.4 ± 0.6 5.6 ± 0.6 NWA4801 4.52 ± 0.16 4557.5 ± 0.7 4557.0 ± 0.3 9.8 ± 0.8 10.2 ± 0.6 5.9 ± 0.6 6.5 ± 0.7 LEW86010 4.80 ± 0.42 4558.3 ± 1.2 4557.5 ± 0.3 9.0 ± 1.3 9.7 ± 0.6 5.1 ± 1.2 5.1 ± 0.6 NWA 2999/4931 5.43 ± 0.34 4559.8 ± 0.9 4560.7 ± 0.5 7.5 ± 1.0 6.4 ± 0.7 3.5 ± 0.9 5.0 ± 1.1 Angra dos Reis 4.02 ± 0.24 4556.0 ± 0.9 4556.6 ± 0.2 11.3 ± 1.0 10.6 ± 0.6 7.4 ± 0.8 a Absolute ages, tD’Orb, are calculated relative to an Pb–Pb age of 4563.4 ± 0.3 Ma. b Pb–Pb ages were re-calculated using their measured U isotope composition (see text). c Calculated using the 4567.2 ± 0.5 Ma Pb–Pb age of CAI SJ101 (Amelin et al., 2010). d Al–Mg data from Spivak-Birndorf et al. (2009) and Schiller et al. (2010). e Mn–Cr data from Lugmair and Shukolyukov (1998), Shukolyukov and Lugmair (2008), Shukolyukov et al. (2009).

Table 4 Characteristics of angrite groups. 182 Group Samples Texture DtCAI (Myr) (e W)i Quenched angrites D’Orbigny; Sah 99555; NWA 1296 Fine-grained, plumose/dendritic to subophitic 4 À2.4 Plutonic angrites NWA 4590; NWA 4801; LEW 86010 Plutonic 10 À1.5 AdoR group Angra dos Reis Plutonic 11 +2.5 NWA 2999 /4931 Annealed breccia, meta-plutonic clasts, metal-rich 7.5 +1.2

238 235 182 was calculated assuming that its U/ U ratio is identical –9.2 slope = (λ Hf)calc. = 0.076 ± 0.012 182 [(λ Hf)meas. = 0.078 ± 0.002] to that of the SRM950a and 960 U isotopic standards (Bou- vier and Wadhwa, 2010). This assumption was based on the

Hf) observations that (i) the Th/U ratio of CAI 2364 B-1 is very –9.6 235 180 low and (ii) that other CAI with low Th/U show no U

Hf/ excesses from the decay of 247Cm (Brennecka et al. 182 2010a). Note, however, that a later study identified a CAI ln( –10.0 with low Th/U but 238U/235U distinct from the value determined for the SRM950a and 960 U isotopic standards (Amelin et al., 2010). –10.4 4556 4560 4564 4568 Fig. 3 reveals that the Hf–W data are consistent with both these Pb–Pb ages for CAI because both ages plot with- Pb−Pb age (Ma) in uncertainty of the 182Hf/180Hf vs. Pb–Pb age correlation Sahara 99555 NWA 4801 Angra dos Reis line defined by the angrites. This is further illustrated by D’Orbigny LEW 86010 CAI (SJ101) calculating ‘absolute’ Hf–W ages for CAI using the relative NWA 4590 NWA 2999/4931 CAI (2364 B−1) Hf–W and absolute Pb–Pb ages of the angrites and the initial 182Hf/180Hf of CAI. These absolute Hf–W ages for CAI Fig. 3. Measured 182Hf/180Hf and Pb–Pb ages for angrites and average at 4567.5 ± 0.8 Ma, consistent with the Pb–Pb ages CAI. The slope of the regression line corresponds to the 182Hf decay constant and is in excellent agreement with the measured for both CAI SJ101 and CAI 2364 B-1. Thus, the Hf–W value of the 182Hf decay constant, indicating that the Hf–W and ages currently have insufficient precision to distinguish Pb–Pb systems provide concordant ages for angrites and CAI. which of these two Pb–Pb dates for CAI better represent References for Pb–Pb ages are given in the text and in Table 3. their absolute age. However, the comparison of the relative ages, DtCAI,of the individual angrites shows that the Hf–W data are more Pb–Pb age of CAI (Brennecka et al., 2010a). Therefore, consistent with an absolute age of 4567.2 ± 0.5 Ma for ages for two CAI are shown in Fig. 3: (i) CAI SJ101 with CAI, as determined for CAI SJ101. All the Pb–Pb forma- a Pb–Pb age of 4567.2 ± 0.5 Ma, calculated using its mea- tion intervals of the angrites calculated relative to the Pb– sured U isotopic composition (Amelin et al., 2010); and Pb age of this CAI are consistent with their Hf–W forma- (ii) CAI 2364 B-1 from the CV chondrite NWA 2364, for tion intervals (Table 3). In contrast, the Pb–Pb formation which Bouvier and Wadhwa (2010) reported a Pb–Pb age intervals calculated relative to a Pb–Pb age of of 4568.22 ± 0.17 Ma. For this CAI the U isotopic compo- 4568.22 ± 0.17 Ma for CAI 2364 B-1 tend to be slightly sition was not measured, however, so that its Pb–Pb age longer than those obtained from the Hf–W system. For Chronology of the angrite parent body 195 instance, the Hf–W age difference between D’Orbigny and 99555) as evidence that 26Al was homogeneously distrib- CAI of 3.9 ± 0.7 Myr is in excellent agreement with the uted in the early solar system (Bouvier and Wadhwa, Pb–Pb age difference between D’Orbigny and CAI SJ101 2010; Bouvier et al., 2011). However, the validity of the of 3.8 ± 0.6 Ma but only marginally overlaps with the Pb–Pb age for CAI 2364 B-1 is uncertain because the U iso- Pb–Pb age difference between D’Orbigny and CAI 2364 topic composition has not been measured for this CAI (see B-1 of 4.9 ± 0.4 Myr. Likewise, for NWA 4801 and LEW above), so that the significance of the apparent agreement 86010 the Hf–W ages of 9.8 ± 0.8 and 9.0 ± 1.3 Myr are between the Al–Mg and Pb–Pb ages (relative to CAI 2364 consistent with their Pb–Pb age differences calculated rela- B-1) remains unclear. Overall, there is considerable uncer- tive to CAI SJ101 but are inconsistent with the Pb–Pb dif- tainty regarding the crystallization age of D’Orbigny and ferences of 11.2 ± 0.3 and 10.7 ± 0.3 Myr calculated for Sahara 99555 based on their Al–Mg systematics, and on these samples relative to CAI 2364 B-1. These results sug- the agreement (or disagreement) of the Al–Mg and Pb–Pb gest that the reported Pb–Pb age of CAI 2364 B-1 might ages. be slightly too old, perhaps due to unaccounted U isotopic Fig. 4 compares the aforementioned Al–Mg and Pb–Pb variations in this CAI. Clearly, U isotopic data for this and ages for D’Orbigny and Sahara 99555 to their Hf–W ages other CAI are needed to more precisely define the absolute obtained in the present study. For D’Orbigny, the Hf–W age of CAI and test the consistency between Pb–Pb and Hf– interval of DtCAI = 3.9 ± 0.7 Myr is shorter than the W ages of CAI. 5.0 ± 0.2 Myr interval obtained from the Al–Mg isochron. However, the Hf–W age agrees well with the 3.8 ± 0.6 Myr 5.2. Comparison of Hf–W and Al–Mg ages: crystallization Pb–Pb age difference between CAI SJ101 and D’Orbigny, age of D’Orbigny and Sahara 99555 and with an Al–Mg model age of 3.9 ± 0.3 Myr for the D’Orbigny whole-rock (Schiller et al., 2010). For Sahara The comparison between Al–Mg and Hf–W ages can 99555 the Hf–W age overlaps with both the Al–Mg iso- only be made for CAI and the oldest of the angrites chron age and the Pb–Pb formation interval calculated rel- (D’Orbigny and Sahara 99555) because at the time the younger angrites crystallized, 26Al was already extinct. The Al–Mg systematics in D’Orbigny and Sahara 99555 Sahara 99555 & D’Orbigny were investigated in detail by Spivak-Birndorf et al. (Al–Mg isochron) (2009) and Schiller et al. (2010). A linear regression (calcu- (Al–Mg mafic min. isochr.) lated using IsoPlot) of the D’Orbigny data from both stud- (Hf–W) ies yields an isochron (MSWD = 4.8) with an initial 26Al/27Al of (4.40 ± 0.55) Â 10À7. Regression of the Al– Sahara 99555 Mg data for Sahara 99555 from Spivak-Birndorf et al. (Al–Mg isochron) (Al–Mg model) (2009) and Schiller et al. (2010) also yields an isochron (Pb–Pb rel. CAI 2364 B-1) (MSWD = 2.0) whose initial 26Al/27Al of (Pb–Pb rel. CAI SJ101) (Hf–W) (4.50 ± 0.57) Â 10À7 is indistinguishable from that of D’Orbigny. The initial 26Al/27Al of D’Orbigny and Sahara D’Orbigny 99555 correspond to formation intervals, DtCAI,of (Al–Mg isochron) 5.0 ± 0.2 Myr after CAI formation (Table 3). The uncer- (Al–Mg model) tainty of this age is larger than those reported in earlier (Pb–Pb rel. CAI 2364 B-1) (Pb–Pb rel. CAI SJ101) studies (e.g., Spivak-Birndorf et al., 2009; Schiller et al., (Hf–W) 2010) because we included the uncertainty in the 26Al half-life (t1/2 = 0.73 ± 0.30; see Nyquist et al., 2009). 3456 Spivak-Birndorf et al. (2009) interpreted the Al–Mg iso- Δ t CAI (Myr) chron ages to reflect the timing of crystallization of D’Orbi- gny and Sahara 99555. However, Schiller et al. (2010) Fig. 4. Comparison of relative Al–Mg, Pb–Pb and Hf–W ages for argued that the Al–Mg systematics in these angrites might angrites Sahara 99555 and D’Orbigny. References for the Al–Mg be disturbed and that the crystallization of D’Orbigny and Pb–Pb ages are given in the text and in Table 3. Al–Mg model and Sahara 99555 is best dated by their Al–Mg model ages ages are from Schiller et al. (2010) and were calculated by assuming of 3.9 ± 0.3 Myr and 4.1 ± 0.3 Myr, calculated based on that the Al–Mg systematics of the angrite whole-rocks reflect a single event of Al/Mg fractionation from a chondritic source. Al– small 26Mg excesses in the D’Orbigny and Sahara 99555 Mg isochron ages were recalculated from data in Spivak-Birndorf whole-rocks compared to chondrites. Based on the observa- et al. (2009) and Schiller et al. (2010) using IsoPlot. Relative Pb–Pb tion that the Pb–Pb age of Sahara 99555 is 3.6 ± 0.5 Myr ages were calculated using two different CAI ages: 4567.2 ± 0.5 Ma younger than that for CAI SJ101, Larsen et al. (2011) ar- for CAI SJ101 (Amelin et al., 2010) and 4568.2 ± 0.2 Ma for CAI gued that the 5 Myr Al–Mg isochron age of Sahara 2364 B-1 (Bouvier and Wadhwa, 2010). The Hf–W ages are in good 99555 reflects a lower-than-canonical initial 26Al/27Al ratio agreement with Pb–Pb ages calculated relative to CAI SJ101, with of the angrite parent body. However, others have inter- the Al–Mg model ages, and with the Al–Mg age obtained for an preted the good agreement between the 5.0 ± 0.2 Myr Al– isochron regressed through the data for mafic minerals and whole- Mg isochron ages for D’Orbigny and Sahara 99555 and rock samples. In contrast, Al–Mg ages obtained from feldspar- the difference between their Pb–Pb ages and CAI 2364 B- controlled isochrons are 1 Myr younger, indicating either a disturbance of the Al–Mg system in the angrite feldspars or a 1 of 4.9 ± 0.4 Myr (D’Orbigny) and 4.6 ± 0.4 Myr (Sahara heterogeneous distribution of 26Al in the early solar system. 196 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203 ative to CAI SJ101. Also plotted in Fig. 4 is the Hf–W age from data in Spivak-Birndorf et al. (2009) and Schiller obtained for a combined regression of the Hf–W data for et al. (2010)] an initial 26Al/27Al of (2.5 + 2.0/ D’Orbigny and Sahara 99555. Fig. 5 shows the results of À1.1) Â 10À5at the time of CAI formation is calculated, this regression and reveals that all analyzed fractions from lower than the initial 26Al/27Al of (5.23 ± 0.13) Â 10À5 that these two angrites plot on a single, well defined isochron is characteristic for CAI (Jacobsen et al., 2008). Such a low (MSWD = 1.3). This indicates that both angrites crystal- initial 26Al/27Al of the angrite precursor material is consis- lized contemporaneously from similar or identical magmas, tent with the value inferred by Larsen et al. (2011) based on consistent with the Al–Mg data. The combined isochron a correlation between 26Mg excesses and nucleosynthetic yields a Hf–W age of 4.2 ± 0.6 Myr after CAI formation, 54Cr anomalies. which we consider the best estimate for the crystallization The third possibility that could account for the disparity age of D’Orbigny and Sahara 99555. This Hf–W age mar- in the calculated Al–Mg and Hf–W formation interval be- ginally overlaps the 5.0 ± 0.2 Myr Al–Mg isochron age tween CAI and the angrites D’Orbigny and Sahara 99555 for these angrites but it is highly unlikely that these two has been advanced by Schiller et al. (2010), based on the ages are identical. observation that the Al–Mg model ages of the D’Orbigny There are three different scenarios that could account for and Sahara 99555 whole-rocks and an Al–Mg isochron the mismatch between the Hf–W and Al–Mg isochron ages based on mafic minerals from these two angrites provide for D’Orbigny and Sahara 99555: (i) the Hf–W formation a better match to results from other chronometers as the interval is too short because the Hf–W isochron for CAI feldspar-controlled Al–Mg isochrons. However, Schiller does not reflect the 182Hf/180Hf at the time of CAI forma- et al. (2010) also noted that there is no textural evidence tion; (ii) at the time of CAI formation the precursor mate- for a disturbance of the Al–Mg systematics in the angrites. rial of the angrites had a lower-than-canonical 26Al/27Al; Moreover, it is unclear as to whether the Al–Mg whole- (iii) the Al–Mg systematics in Sahara 99555 and D’Orbigny rock model ages date crystallization of the angrites or are disturbed so that the Al–Mg isochrons do not reflect the rather are related to Al/Mg fractionation during magma time of crystallization. The first of these possibilities can be ocean solidification, which predated the extrusion and crys- excluded because the Hf–W data for CAI show no evidence tallization of the angrite melts (Schiller et al., 2010). Never- for disturbance (Burkhardt et al., 2008). There are small theless, Tonui et al. (2003) observed that Sm–Nd data for a nucleosynthetic W isotope anomalies in some CAI and it plagioclase separate from D’Orbigny plot significantly is currently unclear how this affects the Hf–W systematics. above a 4.56 Ga isochron, indicating disturbed isotope sys- However, it is unlikely that nucleosynthetic anomalies will tematics at least in the plagioclase. This and the observation change the slope of the CAI isochron significantly. that the Hf–W isochron ages for D’Orbigny and Sahara The second of the aforementioned possibilities is that the 99555 are in excellent agreement with both the Al–Mg mod- angrite parent body accreted from precursor material that el and mafic mineral isochron ages (Fig. 4) suggests that the initially had a lower 26Al/27Al than CAI. Using the Hf–W Al–Mg system in the angrite feldspars is disturbed and that formation interval of D’Orbigny and Sahara 99555 of the feldspar-controlled Al–Mg isochrons may not provide a 26 27 26 27 DtCAI = 4.2 ± 0.6 Myr (see Fig. 5) and the initial Al/ Al reliable estimate of the Al/ Al at the time of angrite obtained from a combined D’Orbigny and Sahara 99555 crystallization. isochron [26Al/27Al = (4.44 ± 0.33) Â 10À7; recalculated In summary, the Hf–W data strongly suggest that the crystallization age of the oldest angrites, D’Orbigny and Sa- hara 99555, is 4.2 ± 0.6 Myr after CAI formation, i.e., older than the age inferred based on Al–Mg chronometry (Spi- 60 Sahara 99555 vak-Birndorf et al., 2009; Bouvier and Wadhwa, 2010). D’Orbigny px−2 The mismatch between the Hf–W and Al–Mg ages for these px−3 two angrites may reflect either disturbed Al–Mg systematics 40 in angrite feldspars or a heterogeneous distribution of 26Al in the early solar system. Distinguishing between these two W

182 possibilities is diﬃcult and will require more precise con- ε straints on the initial 182Hf/180Hf and the Pb–Pb age of 20 px−2 m = (6.99 ± 0.11) x 10−5 CAI, and a better understanding of any disturbance in px−1 i = −2.24 ± 0.13 the Al–Mg systematics of the angrites. The disturbed Sm– MSWD = 1.3 Nd systematics of D’Orbigny plagioclase reveal in any case Δt CAI = 4.2 ± 0.6 Myr 0 that this angrite may not be suitable as a common anchor 020406080for short-lived systems, and hence cannot be used to argue 26 180 Hf/ 184W for or against Al homogeneity in the early solar system.

Fig. 5. Hf–W isochron for D’Orbigny and Sahara 99555. Regres- 5.3. Comparison of Hf–W and Mn–Cr ages sion (including data from Markowski et al., 2007) calculated using 182 180 182 the model 1 ﬁt of IsoPlot. m = initial Hf/ Hf, i = initial e W. In Fig. 6, initial 182Hf/180Hf ratios for angrites are plot- For deﬁnition of DtCAI see Fig. 2. All fractions from D’Orbigny ted against their initial 53Mn/55Mn ratios. In this plot, sam- and Sahara 99555 plot on a single isochron, indicating that these ples having concordant Hf–W and Mn–Cr ages should plot two angrites formed contemporaneously from identical or similar on a single straight line, whose slope can be calculated from magmas. Chronology of the angrite parent body 197

since Hf and W can be strongly fractionated by large-scale –9.2 differentiation processes such as core formation and mantle differentiation. Therefore, the Hf–W data can shed light on the timescales of these important differentiation processes.

Hf) Markowski et al. (2007) observed that Sahara 99555 and

180 –9.6 D’Orbigny evolved with diﬀerent time-integrated

Hf/ 180Hf/184W than NWA 2999, suggesting that at least two 182 distinct mantle sources had formed during the early differ- ln( entiation of the angrite parent body. Fig. 8 shows the W λ182 λ53 –10.0 slope = ( Hf/ Mn)calc. = 0.38 ± 0.04 isotope evolution of all eight angrites examined in the pres- λ182 λ53 [( Hf/ Mn)meas. = 0.42 ± 0.04] ent study and confirms the presence of at least two mantle sources in the angrite parent body. The Hf–W systematics –14.0 –13.0 –12.0 of the quenched and plutonic angrites are consistent with ln(53Mn/55Mn) an evolution in a reservoir with a modestly elevated Hf/ W ratio compared to chondrites. These two groups of ang- D’Orbigny LEW 86010 Solar system initial (CI chondrites) rites may thus derive from a common source. In contrast, NWA 4801 NWA 2999/4931 Solar system initial (26Al–53Mn correlation) NWA 2999/4931 and Angra dos Reis have much more radiogenic initial W isotope compositions, such that their 182 180 53 55 Fig. 6. Comparison of measured Hf/ Hf and Mn/ Mn source(s) must have evolved with much higher Hf/W than ratios in angrites. A linear regression through these data has a slope 182 53 the other angrites. that is consistent with the ratio of the Hf and Mn decay The presence of at least two sources with distinct Hf/W constants, indicating that the angrites have concordant Hf–W and ratios in the mantle of the angrite parent body raises the Mn–Cr ages. Also shown are two different estimates for the solar system initial 53Mn/55Mn: (9.1 ± 1.7) Â 10À5 as inferred from a question as to whether this is the sole result of core forma- 26Al/27Al–53Mn/55Mn correlation line for meteorites (Nyquist tion or if it reflects additional Hf/W fractionation during la- et al., 2009) and (6.3 ± 0.7) Â 10À5 as determined from a Mn–Cr ter mantle differentiation. Fractionation of Hf from W in isochron for Orgueil (Trinquier et al., 2008). Mn–Cr data for silicate systems may occur in the presence of ilmenite and angrites are from Lugmair and Shukolyukov (1998), Glavin et al. clinopyroxene both of which are characterized by D(Hf)/ (2004), Shukolyukov and Lugmair (2008), and Shukolyukov et al. D(W) > 1 (Righter and Shearer, 2003). However, as dis- (2009). cussed by Markowski et al. (2007), the near-chondritic REE patterns of most angrites (Mittlefehldt et al., 2002; Gellissen et al., 2007) argue against significant amounts of 182 53 the ratio of the Hf and Mn half-lives (Nyquist et al., residual clinopyroxene in the angrite source regions. Fur- 2009). Fig. 6 reveals that while NWA 4801, LEW 86010 thermore, the chondritic Hf/Sm of most angrites (see Mar- and D’Orbigny have concordant Hf–W and Mn–Cr ages, kowski et al., 2007) indicates that no ilmenite is present in NWA 2999 plots slightly above but within uncertainty of the angrite sources (Blichert-Toft et al., 2002). Thus, the the line of concordant ages. Also plotted in Fig. 6 are two distinct Hf/W ratios of the two angrite sources must solely 53 55 different estimates for the solar system initial Mn/ Mn: result from metal–silicate fractionation during core forma- À6 (6.3 ± 0.7) Â 10 as inferred from Mn–Cr data of inner tion. Consequently, core formation in the angrite parent solar system objects (Trinquier et al., 2008); and body was not a single event of metal segregation in a mantle 6 (9.1 ± 1.7) Â 10À as derived from a correlation of initial characterized by a uniform Hf/W ratio. Instead, metal seg- 26 27 53 55 Al/ Al and Mn/ Mn ratios in meteorites (Nyquist regation in the angrite parent body must have occurred 182 180 et al. 2009). Both these estimates plot on the Hf/ Hf more locally and under varying conditions that led to differ- 53 55 vs. Mn/ Mn correlation line defined by the angrites, indi- ent degrees of W depletion in the separate angrite source re- cating that the Hf–W and Mn–Cr systematics of the ang- gions. The lower Hf/W ratio in the source of the quenched rites are consistent with current estimates for the solar and plutonic angrites compared to the source of NWA 182 180 53 55 system initial Hf/ Hf and Mn/ Mn ratios. However, 2999/4931 and Angra dos Reis may reflect a higher propor- the Hf–W and Mn–Cr data cannot yet distinguish which of tion of (stranded) metal in the former. Alternatively it may the two aforementioned estimates for the solar system reflect more oxidizing conditions and hence less siderophile 53 55 initial Mn/ Mn is correct, although a value of behavior of W during core formation (Palme and Rammen- 6 (9.1 ± 1.7) Â 10À seems to be somewhat too high. see, 1981) in the source of the quenched and plutonic angrites. 6. DIFFERENTIATION OF THE ANGRITE PARENT The presence of the two separate angrite sources could BODY alternatively be explained if NWA 2999/4931 and Angra dos Reis derive from a different parent body from the other A key observation from the Hf–W data is that all the angrites. However, this possibility is rejected here in light of 182 184 angrites have initial W/ W ratios that are elevated the very distinctive mineralogy, elemental ratios, D17O val- compared to the chondritic value at the time of angrite crys- ues and 54Cr anomalies shared by all angrites (e.g., Mit- tallization. Thus, prior to crystallization, the angrite tlefehldt et al., 1998; Greenwood et al., 2005; Trinquier sources must have evolved in a reservoir with higher- et al., 2008; Shukolyukov et al., 2009), and the fact that than-chondritic Hf/W. This observation is not surprising NWA 2999 plots on the Mn–Cr bulk isochron defined by 198 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203 the two subophitic angrites and LEW 86010 (Shukolyukov 4.0 and Lugmair, 2007). We thus interpret these two sources to 2 Myr result from an early differentiation of a single angrite parent 1 Myr 3 Myr body. 2.0 180Hf/184W = 8 ± 4

6.1. Chronology of diﬀerentiation W 0.0

182 180Hf/184W = 3.1 ± 1.4 6.1.1. Timing of core formation ε The simplest conceptual framework for using the Hf–W data to constrain the timescale of core formation is a two- –2.0 stage model, which uses the present-day 180Hf/184W and CHUR (180Hf/184W = 1.23±0.15) 182W/184W ratios of the angrite bulk rocks and calculates –4.0 the time at which they separated from a chondritic Hf–W 0.8 0.6 0.4 reservoir (Lee and Halliday, 1995; Harper and Jacobsen, 182Hf/180Hf × 104 1996; Kleine et al., 2009). Most angrites have model ages D’Orbigny NWA 4590 NWA 2999/4931 that range from 0to4 Myr after CAI formation, with Sahara 99555 NWA 4801 Angra dos Reis an average age of 1.9 ± 0.8 Myr (2 SE) (Fig. 7). NWA 2999 and Angra dos Reis, however, plot outside the two- NWA 1296 LEW 86010 stage evolutionary ﬁeld in Fig. 7 and have negative model Fig. 8. Hf–W evolution diagram for angrites as deﬁned by their ages. This indicates either that the source(s) of these two initial e182W and 182Hf/180Hf. The Hf–W evolution of chondrites angrites had a more complex history involving more than (CHUR) is constrained by the initial W isotopic composition of one Hf/W fractionation event or that the analyzed whole- CAI (Burkhardt et al., 2008) and the present-day W isotopic rocks are not representative bulk rocks. It is noteworthy composition of chondrites (Kleine et al., 2002; Kleine et al., 2004) that the NWA 2999 and NWA 4931 whole-rocks have dif- and is shown for reference. Solid lines are W isotope evolution ferent 180Hf/184W and 182W/184W, despite that fact that curves for quenched and plutonic angrites as well as NWA 2999/ these two meteorites are almost certainly paired. The two- 4931 and Angra dos Reis. Regressions and error envelopes were calculated using IsoPlot. stage model age of NWA 4931 is positive and consistent

6 with those of the other angrites, which may indicate that t the Hf–W data for the NWA 2999 whole-rock are not rep- CHUR < 0 2 Myr resentative of a true bulk rock (perhaps due to the high metal content of this angrite). Whether a similar argument can 0 Myr 4 Myr 4 be made for Angra dos Reis is less clear but the coarse- grained, nearly mono-mineralic nature of this rock renders

W it diﬃcult to obtain a representative bulk rock sample be- 182

ε cause of a heterogeneous distribution of intercumulus 2 phases. The two-stage model ages provide a reliable estimate for the timing of core formation only if the Hf–W systematics t CHUR > 0 of the whole-rocks solely record Hf/W fractionation during 0 core formation. This requires that during solidification of 468the mantle and later re-melting of the angrite sources no 180 184 Hf/ W significant Hf/W fractionation took place. The extent of any Hf/W fractionation after core formation can be as- D’Orbigny NWA 1296 NWA 4931 sessed by comparing the Hf/W ratio of the angrite bulk Sahara 99555 NWA 4801 NWA 2999 rocks to that of their sources. There are two approaches Angra dos Reis that may be used for this task. The first approach uses 180 184 Fig. 7. Present-day e182W vs. 180Hf/184W for angrite whole-rocks. the time-integrated Hf/ W, which can be calculated Data shown with gray symbols are from Markowski et al. (2007). from the difference in the initial W isotope compositions The Hf–W systematics of samples that plot in the white area are of the solar system (as determined based on CAI) and the consistent with a two-stage model and one event of Hf/W angrites themselves. This approach assumes that the Hf/ fractionation. For these samples two-stage model ages, tCHUR, W fractionation resulting from the early differentiation of are positiv and may have chronological significance. Samples that the angrite parent body occurred at the time of CAI forma- plot in the hatched area had a more complex Hf–W evolution tion. The quenched and plutonic angrites both have time- involving more than one event of Hf–W fractionation. For these integrated 180Hf/184W ratios of 3–3.5, whereas NWA samples t is negative and, hence, meaningless. The petrogen- CHUR 2999/4931 and Angra dos Reis have much higher time-inte- esis of angrites probably involved more than one Hf/W fraction- 180 184 ation event, and so the two-stage model ages may not provide grated Hf/ Wof9. 182 accurate ages of core formation. Instead, metal segregation in the The second approach uses the initial e W and 182 180 angrite parent body must have occurred earlier than given by the Hf/ Hf of the angrites as determined from their inter- two-stage model ages. See text for details. nal isochrons to reconstruct the Hf–W isotope evolution Chronology of the angrite parent body 199 of their sources. This approach involves no assumption 0 regarding the timing of Hf/W fractionation. Fig. 8 reveals that a common source of quenched and plutonic angrites Whole−rocks evolved with a180Hf/184W ratio of 3.1 ± 1.4, whereas –1 NWA 2999/4931 and Angra dos Reis must derive from a Partial melting & source (or sources) with a 180Hf/184W ratio of 8 ± 4. These extrusion of angrites Source source 180Hf/184W ratios are similar to the time-integrated W –2 182 values, suggesting that these estimates are reliable and that ε (A) the primordial differentiation of the angrite parent body oc- CHUR curred at about the time of CAI formation. –3 All analyzed whole-rocks of the quenched and plutonic (B) apparent model age of core formation 180 184 angrites have Hf/ W ratios that are higher (4.6–7.8; (C) true age of core formation 180 184 –4 see Table 2) than that of their source ( Hf/ W= 0246810 3.1 ± 1.4; see Fig. 8). Thus, the melting that produced these Time (Myr after CAI formation) angrites involved some Hf/W fractionation but surprisingly Fig. 9. Schematic W isotope evolution diagram for the source of the angrite specimens themselves have higher 180Hf/184W ra- quenched and plutonic angrites. Shown are W isotope evolution tios than their source. This is unexpected because during curves for (i) chondrites (CHUR), (ii) the source of the quenched partial melting in silicate systems W is thought to be more and plutonic angrites (using 180Hf/184W = 3.5 and assuming a time incompatible than Hf, so the angrite melts should have low- of core formation of 0.38 Myr after CAI formation), and (iii) a er Hf/W ratios than their source, opposite to what is ob- whole-rock sample of the quenched angrites (D’Orbigny, Sahara served. However, residual metal in the angrite source 99555, NWA 1296). Since the whole-rock (A) has a higher could retain W and cause elevated Hf/W ratios in the angrite 180Hf/184W than its source, the two-stage model ages for core 180 184 melts compared to their source. Thus, the Hf–W systematics formation calculated using the present day Hf/ W and 182 184 seem to require the presence of metal in the source region of W/ W of the whole-rocks (B) are younger than the true age the quenched and plutonic angrites. It is noteworthy that of core formation (C). Consequently, core formation in the angrite parent body must have occurred earlier than given by the two-stage metal in the source of these angrites could also account model ages of 2 Myr. for its lower Hf/W ratio compared to the source of NWA 2999/4931 and Angra dos Reis. Evidently the Hf–W systematics of the angrite whole- slightly higher than the initial 53Mn/55Mn of D‘Orbigny rocks do not result from just one Hf/W fractionation event of (3.24 ± 0.04) Â 10À6 (Glavin et al., 2004). The Mn–Cr during core formation but also reflect subsequent Hf/W bulk rock isochron yields the timing of the last Mn–Cr iso- fractionation during the melting that produced the angrites. tope equilibration in the mantle of the angrite parent body, Consequently, the two-stage model ages may not accurately which most probably is related to the solidification of the provide the time of core formation. As the angrite whole- mantle after core formation. Petrologic evidence suggests rocks have higher Hf/W than their sources, core formation that olivine and spinel are important minerals in the angrite must have occurred earlier than given by the two-stage source regions and that many compositional characteristics model ages (Fig. 9), so that metal segregation in the angrite of angrites can be accounted for by varying proportions of parent body probably occurred within the first 2 Myr these two minerals (Mittlefehldt et al., 2002). Importantly, after CAI formation. Alternatively the age of core forma- olivine and spinel have largely different Mn/Cr ratios, such tion can be estimated by determining the time at which that various proportions of these minerals could be respon- the Hf–W isotope evolution of the angrite sources deviated sible for Mn/Cr fractionations within the mantle. The dif- from that of chondrites. Fig. 8 reveals that both the sources ference in the initial 53Mn/55Mn ratios of the bulk rock of the quenched and plutonic angrites as well as that of isochron and D‘Orbigny corresponds to an age difference NWA 2999/4931 and Angra dos Reis separated from the of 0.3 ± 0.2 Myr, indicating that the Mn/Cr fractionation chondritic evolution within the first 2 Myr after CAI for- in the angrite parent body mantle occurred about contem- mation. Thus, core formation in the angrite parent body poraneously to the extrusion of the oldest angrite. Combin- must have occurred within the first 2 Myr after solar sys- ing this age difference with the Hf–W age of D‘Orbigny of tem formation, consistent with the estimate from the two- DtCAI=3.9 ± 0.7 Myr reveals that mantle differentiation stage model ages derived above. took place at 3.6 ± 0.7 Myr after CAI formation. That mantle differentiation in the angrite parent body 6.1.2. Timing of mantle solidification occurred 3.6 ± 0.7 Myr after CAI formation provides fur- The Hf–W data provide little information on the timing ther support for our conclusion that the distinct Hf/W ra- of mantle differentiation in the angrite parent body, as the tios of the two angrite sources result solely from core dominant process fractionating Hf and W is core formation formation, which occurred within the first 2 Myr after (see above). However, information regarding the timescales CAI formation (see above). If mantle differentiation would of silicate differentiation may be obtained from the Mn–Cr have been responsible for the Hf/W fractionation between systematics of the angrites. A bulk rock Mn–Cr isochron the source of NWA 2999/4931 and Angra dos Reis and that for four angrites (LEW 86010, NWA 2999, Sahara 99555, of the other angrites, the former source would have to have D’Orbigny) corresponds to an initial 53Mn/55Mn of evolved with a 180Hf/184W ratio of greater than 12 (3.40 ± 0.14) Â 10À6 (Shukolyukov and Lugmair, 2007), (Fig. 8). The NWA 2999/4931 and Angra dos Reis whole- 200 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203 rocks have much lower 180Hf/184W ratios, so substantial ins-Tanton et al., 2011). The Hf–W evidence for accretion Hf/W fractionation would need to have occurred during of the angrite parent body within the first 1.5 Myr after the melting that produced these angrites. However, there CAI formation, therefore, is consistent with 26Al heating is no evidence for such large trace element fractionations being the dominant heat source responsible for differentia- during angrite melting (see above), so mantle differentiation tion, and is also consistent with the past presence of a core as dated by the Mn–Cr bulk rock isochron does not seem to dynamo in the angrite parent body (Weiss et al., 2008). have induced any significant Hf/W fractionation. Although the angrite parent body accreted sufficiently early to allow for global melting by heating from 26Al de- 6.1.3. Chronology of the angrite parent body cay, core formation was not a single event of metal segrega- The combined Hf–W and Mn–Cr constraints permit tion in a global magma ocean characterized by a uniform reconstructing the chronology of the early differentiation Hf/W ratio. Metal–silicate fractionation rather must have of the angrite parent body and reveal that accretion and occurred more locally and under varying conditions that core formation terminated no later than 2 Myr after led to different degrees of W depletion in the separate ang- CAI formation and was followed by differentiation during rite source regions, as is evident from the presence of at solidification of the mantle that occurred at 3.6 ± 0.7 Myr least two mantle reservoirs that acquired their distinct Hf/ after CAI formation. This in turn was followed by the W ratios as a result of core formation. Furthermore, the extrusion of the first angrite lavas at 4 Myr after CAI for- preservation of these distinct mantle reservoirs indicates mation, which continued for 7 Myr, as constrained by the some persistent heterogeneity in the mantle that has not crystallization age of the youngest angrite, Angra dos Reis, been erased by convection in a magma ocean. This is sur- of DtCAI 11 Myr. prising because convection is expected to have resulted in Using thermal modeling of planetesimals heated by decay efficient homogenization of the mantle. The accretion and of 26Al, Qin et al. (2008) observed that there is a time gap be- differentiation of the angrite parent body, therefore, cannot tween accretion and core formation. This gap reflects the per- have occurred instantaneously as a single event, but must iod between the time at which a planetesimal became large have been a more protracted, multi-stage process. enough to trap the heat produced by 26Al decay in its interior Hevey and Sanders (2006) argued that in bodies heated and the time at which the temperature became high enough by 26Al decay convection in the magma ocean began at 50% for melting and differentiation to occur. According to this silicate melting, which was reached about 0.8 Myr after thermal modeling, a planetesimal that underwent core for- (instantaneous) accretion. On the angrite parent body, mation within the first 2 Myr after CAI formation must therefore, convection would have started between 0.8 and have accreted within in the first 1.5 Myr after CAI forma- 2.3 Myr after CAI formation, given that accretion occurred tion (Qin et al., 2008). Thus, the angrite parent body must within the first 1.5 Myr after CAI formation. The end of have accreted within the first 1.5 Myr of the solar system. convection and solidification of the magma ocean probably The early accretion of the angrite parent body is consis- occurred around 3.6 ± 0.7 Myr after CAI formation, as gi- tent with Hf–W evidence for accretion and differentiation of ven by the Mn–Cr age of the bulk rock isochron for ang- the parent bodies of magmatic iron meteorites within rites (see above).The Hf–W and Mn–Cr chronological 1 Myr after CAI formation, but is difficult to reconcile constraints for core formation and magma ocean solidifica- with Rb–Sr evidence for a late accretion of the parent tion, therefore, would suggest that convection in the mag- bodies of eucrites and angrites. The elevated initial 87Sr/86Sr ma ocean continued for 1–4 Myr after the core had of the angrite parent body compared to the initial 87Sr/86Sr fully formed (based on the Mn–Cr bulk rock isochron age of CAI has been taken as evidence for a relatively late of 3.6 ± 0.7 Myr and the Hf–W age for core formation of accretion of the angrite parent body, more than 2 Myr <2 Myr). However, the presence of at least two mantle res- after CAI formation (Lugmair and Galer, 1992; Nyquist ervoirs that acquired their distinct Hf/W ratios as a result et al., 1994; Halliday and Porcelli, 2001). However, once of core formation within the first 2 Myr after CAI forma- the presence of recently discovered nucleosynthetic Sr iso- tion indicates that convection in the magma ocean cannot tope anomalies in CAI is taken into account, the Rb–Sr have been efficient in isotopically and chemically homoge- model time between CAI formation and angrite accretion nizing the mantle of the angrite parent body. This observa- is significantly reduced and consistent with the Hf–W age tion seems inconsistent with the absence of mass- constraints (Hans et al., 2011). independent O isotopic variations in the angrites, which was used to argue for efficient homogenization of the man- 6.2. Implications for accretion and differentiation of the tle by convection (Greenwood et al., 2005). However, equil- angrite parent body ibrated ordinary chondrites also do not show significant mass-independent O isotope variations (Clayton et al., Thermal modeling shows that the decay of 26Al provides 1991), although global magma oceans never existed on sufficient energy for large-scale melting in planetary bodies the parent bodies of these meteorites. This strongly suggests with radii larger than 20 km that accreted within the first that the homogenization of mass-independent O isotope 1.5 Myr of the solar system (Hevey and Sanders, 2006; variations among components of primitive meteorites does Sahijpal et al., 2007). Such bodies formed global magma not require global melting and, consequently, the lack of oceans, in which efficient metal–silicate separation and me- mass independent O isotope variations among the angrites tal segregation could take place. These planetesimals also is not evidence for efficient homogenization in a global had a sufficient heat flux to produce core dynamos (Elk- magma ocean on their parent body. Chronology of the angrite parent body 201

Mantle solidification on the angrite parent body at stantly removed newly formed, insulating crust, thereby 3.6 ± 0.7 Myr after CAI formation (as dated by the Mn– facilitating rapid cooling of the magma ocean, as required Cr bulk rock isochron) indicates a surprisingly rapid cooling by the Mn–Cr systematics of bulk angrites. of the magma ocean. Detailed numerical modeling of the differentiation history of planetesimals shows that such ra- ACKNOWLEDGMENTS pid cooling can only be achieved if persistent insulating crust (either unmelted chondritic crust or freshly formed igneous We thank the NASA Meteorite Working Group for generously crust) is constantly being removed by impacts. Otherwise providing a sample of LEW 86010, Ben Weiss and Museu Nacional, Rio de Janeiro, Brazil who made a sample of Angra dos Reis avail- cooling of the magma ocean would have taken at least 6– able for us, and Ted Bunch for providing a sample of NWA 1296. 7 Myr (Gupta and Sahijpal, 2010). Thus, a continuous bom- Other samples were portions of official type material or were pur- bardment of the angrite parent body by small planetesimals chased from private sources. We are grateful to John Kashuba seems to be required to account for the rapid solidification and Kevin Righter for loaning several thin sections, and to Ted indicated by the Mn–Cr data. Such planetesimal impacts Bunch for the photomicrographs. Liping Qin, Richard Walker would also account for the lack of complete homogenization and an anonymous referee provided thoughtful reviews that led to in the magma ocean, and the presence of at least two distinct significant improvement of this manuscript. This work was sup- mantle reservoirs. The impacts would have added cold ported by the Swiss National Science Foundation (PP00P2_ material to the magma ocean and might have resulted in 123470). more localized melting and metal–silicate separation events under varying conditions, thereby accounting for the different degrees of W depletion in the two separate angrite source REFERENCES regions. Moreover, more localized metal segregation events Amelin Y. (2008a) U–Pb ages of angrites. Geochim. Cosmochim. may have been inefficient, possibly leaving behind some me- Acta 72, 221–232. tal in the angrite source regions, as observed at least for the Amelin Y. (2008b) The U–Pb systematics of angrite Sahara 99555. source of the quenched and plutonic angrites. Geochim. Cosmochim. Acta 72, 4874–4885. Amelin Y., Krot A. N., Hutcheon I. D. and Ulyanov A. A. (2002) 7. CONCLUSIONS Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science 297, 1678–1683. The Hf–W isochron ages for angrites are in excellent Amelin Y., Kaltenbach A., Iizuka T., Stirling C. H., Ireland T. R., agreement with their Pb–Pb and Mn–Cr ages, indicating Petaev M. and Jacobsen S. B. (2010) U–Pb chronology of the 238 235 that these three systems provide concordant ages for the Solar System’s oldest solids with variable U/ U. Earth angrites. Furthermore, the absolute Hf–W age of CAI as Planet. Sci. Lett. 300, 343–350. Amelin Y., Kaltenbach A. and Stirling C. H. (2011) The U–Pb calculated relative to the angrites is 4567.5 Ma, consistent systematics and cooling rate of plutonic angrite NWA 4590. with a recently reported Pb–Pb age for the Allende CAI Lunar Planet. Sci. Conf. XLII. #1682 (abstr.). SJ101 (Amelin et al., 2010). However, the Hf–W ages for Blichert-Toft J., Boyet M., Te´louk P. and Albare`de F. (2002) D’Orbigny and Sahara 99555 are 1 Myr older than ages 147Sm–143Nd and 176Lu–176Hf in eucrites and the differentiation obtained from the Al–Mg system, indicating either a distur- of the HED parent body. Earth Planet. Sci. Lett. 204, 167–181. bance of the Al–Mg system in angrite feldspars, or a heter- Bouvier A. and Wadhwa M. (2010) The age of the Solar System ogeneous distribution of 26Al in the early solar system. redefined by the oldest Pb–Pb age of a meteoritic inclusion. Nat. Given the evidence for disturbed Sm–Nd systematics in Geosci. 3, 637–641. feldspars from D’Orbigny, the former interpretation seems Bouvier A., Spivak-Birndorf L. J., Brennecka G. A. and Wadhwa more likely. We note, however, that currently available age M. (2011) New constraints on early Solar System chronology from Al–Mg and U–Pb isotope systematics in the unique data do not provide strong evidence either for or against a 26 basaltic achondrite Northwest Africa 2976. Geochim. Cosmo- homogeneous distribution of Al in the early solar system. chim. Acta 75, 5310–5323. The Hf–W data require that core formation in the ang- Brennecka G. A. and Wadhwa M. (2011) 238U/235U ratios of rite parent body occurred within 2 Myr of CAI forma- angrites: adjusting absolute ages of anchors. Mineral. Mag. 75, 26 tion, early enough that heating by Al decay caused the 579. formation of a global magma ocean in which efficient me- Brennecka G. A., Weyer S., Wadhwa M., Janney P. E., Zipfel J. and tal–silicate separation could take place. Nevertheless, the Anbar A. D. (2010a) 238U/235U variations in meteorites: extant 247 Hf–W data indicate that core formation in the angrite par- Cm and implications for Pb–Pb dating. Science 327, 449–451. ent body was not a single event of metal segregation but Brennecka G. A., Wadhwa M., Janney P. E. and Anbar A. D. rather occurred more locally and under varying conditions (2010b) Towards reconciling early solar system chronometers: the 238U/235U ratios of chondrites and D’Orbigny pyroxenes. in partially or fully molten regions. Metal segregation in the Lunar Planet. Sci. XLI. #2117 (abstr.). angrite parent body thus was fundamentally different from Burkhardt C., Kleine T., Palme H., Bourdon B., Zipfel J., Friedrich core formation processes in the large terrestrial planets. The J. and Ebel D. (2008) Hf–W mineral isochron for Ca, Al-rich preservation of distinct Hf–W signatures from separate core inclusions: age of the solar system and the timing of core forming events indicates that the angrite parent body es- formation in planetesimals. Geochim. Cosmochim. Acta 72, caped efficient homogenization by convection in a magma 6177–6197. ocean. This most plausibly was caused by the addition of Clayton R. B., Mayeda T. K., Goswami J. N. and Olsen E. J. cold material by impacts of small planetesimals during (1991) Oxygen isotope studies of ordinary chondrites. Geochim. ongoing core formation. These impacts may have con- Cosmochim. Acta 55, 2317–2337. 202 T. Kleine et al. / Geochimica et Cosmochimica Acta 84 (2012) 186–203

Connelly J., Bizzarro M., Thrane K. and Baker J. A. (2008) The cooling history of the H chondrite parent body. Earth Planet. Pb–Pb age of angrite SAH99555 revisited. Geochim. Cosmo- Sci. Lett. 270, 106–118. chim. Acta 72, 4813–4824. Kleine T., Touboul M., Bourdon B., Nimmo F., Mezger K., Palme Elkins-Tanton L. T., Weiss B. P. and Zuber M. T. (2011) H., Yin Q. Z., Jacobsen S. B. and Halliday A. N. (2009) Hf–W Chondrites as samples of differentiated planetesimals. Earth chronology of the accretion and early evolution of asteroids Planet Sci. Lett. 305, 1–10. and terrestrial planets. Geochim. Cosmochim. Acta 73, 5150– Gellissen M., Palme H., Korotov R.L. and Irving A.J. (2007) NWA 5188. 2999, a unique angrite with a large chondritic component. Kuehner S. M. and Irving A.J. (2007a) Grain boundary glasses in Lunar Planet. Sci. XXXVIII. #1612 (abstr.). plutonic angrite NWA 4590: evidence for rapid decompressive Glavin D. P., Kubny A., Jagoutz E. and Lugmair G. W. (2004) partial melting and cooling on Mercury? Lunar Planet Sci. Mn–Cr isotope systematics of the D’Orbigny angrite. Meteorit. XXXVIII. #1522 (abstr.). Planet. Sci. 39, 693–700. Kuehner S.M. and Irving A.J. (2007b) Primary ferric iron-bearing Greenwood R. C., Franchi I. A., Jambon A. and Buchanan P. C. rho¨nite in plutonic igneous angrite NWA 4590: implications for (2005) Widespread magma oceans on asteroidal bodies in the redox conditions on the angrite parent body. American early Solar System. Nature 435, 916–918. Geophysical Union, Fall Meeting 2007. #P41A-0219 (abstr.). Gupta G. and Sahijpal S. (2010) Differentiation of Vesta and the Kuehner S.M., Irving A.J., Bunch T.E., Wittke J.H., Hupe´ G. M. parent bodies of other achondrites. J. Geophys. Res. 115, and Hupe´ A.C. (2006) Coronas and symplectites in plutonic E08001. angrite NWA 2999 and implications for Mercury as the angrite Halliday A. N. and Porcelli D. (2001) In search of lost planets - the parent body. Lunar Planet. Sci. XXXVII. #1344 (abstr.). paleocosmochemistry of the inner solar system. Earth Planet. Larsen K. K., Trinquier A., Paton C., Schiller M., Wielandt D., Sci. Lett. 192, 545–559. Ivanova M. A., Connelly J. A., Nordlund A., Krot A. N. and Hans U., Kleine T. and Bourdon B. (2011) Strontium isotope Bizzarro M. (2011) Evidence for magnesium isotope heteroge- anomalies in Ca–Al-rich inclusions and the Rb–Sr chronology neity in the solar protoplanetary disk. Astrophys. J. 735, L37– of volatile depletion revisited. Lunar Planet. Sci. XLII. #2672 L50. (abstr.). Lee D. C. and Halliday A. N. (1995) Hafnium–tungsten chronom- Harper C. L. and Jacobsen S. B. (1996) Evidence for 182Hf in the etry and the timing of terrestrial core formation. Nature 378, early solar system and constraints on the timescale of terrestrial 771–774. accretion and core formation. Geochim. Cosmochim. Acta 60, Longhi J. (1999) Phase equilibrium constraints on angrite petro- 1131–1153. genesis. Geochim. Cosmochim. Acta 63, 573–585. Hevey P. J. and Sanders I. S. (2006) A model for planetesimal Lugmair G. W. and Galer S. J. G. (1992) Age and isotopic meltdown by 26Al and its implications for meteorite parent relationships among the angrites Lewis Cliff 86010 and Angra bodies. Meteorit. Planet. Sci. 41, 95–106. dos Reis. Geochim. Cosmochim. Acta 56, 1673–1694. Humayun M., Irving A. J. and Kuehner S. M. (2007) Siderophile Lugmair G. W. and Shukolyukov A. (1998) Early solar system elements in metal from metal-rich angrite NWA 2999. Lunar timescales according to 53Mn–53Cr systematics. Geochim. Cos- Planet. Sci. XXXVIII. #1338 (abstr.). mochim. Acta 62, 2863–2886. Jacobsen B., Yin Q.-Z., Moynier F., Amelin Y., Krot A. N., Markowski A., Quitte´ G., Halliday A. N. and Kleine T. (2006) Nagashima K., Hutcheon I. D. and Palme H. (2008) 26Al–26Mg Tungsten isotopic compositions of iron meteorites: chronolog- and 207Pb–206Pb systematics of Allende CAIs: canonical solar ical constraints vs. cosmogenic effects. Earth Planet. Sci. Lett. initial 26Al/27Al ratio reinstated. Earth Planet. Sci. Lett. 272, 242, 1–15. 353–364. Markowski A., Quitte´ G., Kleine T., Halliday A., Bizzarro M. and Kaltenbach A., Stirling C. H. and Amelin Y. (2011) Revised ages of Irving A. J. (2007) Hf–W chronometry of angrites and the angrites. Mineral. Mag. 75, 1137. earliest evolution of planetary bodies. Earth Planet. Sci. Lett. Kleine T. and Rudge J. F. (2011) Chronometry of meteorites and 262, 214–229. the formation of the Earth and Moon. Elements 7, 41–46. Mittlefehldt D.W., McCoy T.J., Goodrich C.A. and Kracher A. Kleine T., Mu¨nker C., Mezger K. and Palme H. (2002) Rapid (1998) Non-chondritic meteorites from asteroidal bodies.In accretion and early core formation on asteroids and the Planetary Materials, (ed. J. J. Papike), Rev. Mineral.. 36, pp. 4- terrestrial planets from Hf–W chronometry. Nature 418, 952– 131–4-142. 955. Mittlefehldt D. W., Killgore M. and Lee M. T. (2002) Petrology Kleine T., Mezger K., Mu¨nker C., Palme H. and Bischoff A. (2004) and geochemistry of D’Orbigny, geochemistry of Sahara 99555, 182Hf–182W isotope systematics of chondrites, eucrites, and and the origin of angrites. Meteorit. Planet. Sci. 37, 345. Martian meteorites: chronology of core formation and mantle Nyquist L. E., Bansal B., Wiesmann H. and Shih C. Y. (1994) differentiation in Vesta and Mars. Geochim. Cosmochim. Acta Neodymium, Strontium and Chromium Isotopic Studies of the 68, 2935–2946. Lew86010 and Angra-Dos-Reis Meteorites and the Chronology Kleine T., Mezger K., Palme H., Scherer E. and Mu¨nker C. (2005a) of the Angrite Parent Body. Meteoritics 29, 872–885. Early core formation in asteroids and late accretion of Nyquist L. E., Kleine T., Shih C. Y. and Reese Y. (2009) The chondrite parent bodies: evidence from 182Hf–182W in CAIs, distribution of short-lived radioisotopes in the early solar metal-rich chondrites and iron meteorites. Geochim. Cosmo- system and the chronology of asteroid accretion, differentiation, chim. Acta 69, 5805–5818. and secondary alteration. Geochim. Cosmochim. Acta 73, 5115– Kleine T., Mezger K., Palme H., Scherer E. and Mu¨nker C. (2005b) 5136. The W isotope composition of eucrite metals: constraints on the Palme H. and Rammensee W. (1981) The significance of W in timing and cause of the thermal metamorphism of basaltic planetary differentiation processes: evidence from new data on eucrites. Earth Planet. Sci. Lett. 231, 41–52. eucrites. Proc. 12th Lunar Planet. Sci. Conf., 949–964. Kleine T., Touboul M., Van Orman J. A., Bourdon B., Maden C., Patchett P. J. and Tatsumoto M. (1980) A routine high-precision Mezger K. and Halliday A. (2008) Hf–W thermochronometry: method for Lu–Hf isotope geochemistry and chronology. closure temperature and constraints on the accretion and Contrib. Mineral. Petrol. 75, 263–267. Chronology of the angrite parent body 203

Qin L., Dauphas N., Wadhwa M., Masarik J. and Janney P. E. Shukolyukov A., Lugmair G. W. and Irving A. J. (2009) Mn–Cr (2008) Rapid accretion and differentiation of iron meteorite isotope systematics of angrite Northwest Africa 4801. Lunar parent bodies inferred from 182Hf–182W chronometry and Planet. Sci. XL. #1381 (abstr.). thermal modeling. Earth Planet. Sci. Lett. 273, 94–104. Spivak-Birndorf L., Wadhwa M. and Janney P. E. (2009) Quitte` G., Birck J.-L. and Alle`gre C. J. (2000) 182Hf–182W 26Al–26Mg systematics in D’Orbigny and Sahara 99555 ang- systematics in eucrites: the puzzle of iron segregation in the rites: implications for high-resolution chronology using extinct early solar system. Earth Planet. Sci. Lett. 184, 83–94. chronometers. Geochim. Cosmochim. Acta 73, 5202–5211. Righter K. and Shearer C. K. (2003) Magmatic fractionation of Hf Tonui, E. K., Ngo, H. H., and Papanastassiou, D. A. (2003) Rb–Sr and W: constraints on the timing of core formation and and Sm–Nd study of the D’Orbigny angrite. Lunar Planet. Sci. differentiation in the Moon and Mars. Geochim. Cosmochim. XXXIV. #1812 (abstr.). Acta 67, 2497–2507. Touboul M., Kleine T., Bourdon B., Van Orman J. A., Maden C. Sahijpal S., Soni P. and Gupta G. (2007) Numerical simulations of and Zipfel J. (2009) Hf–W thermochronometry: II. Accretion the differentiation of accreting planetesimals with 26Al and 60Fe and thermal history of the acapulcoite-lodranite parent body. as the heat sources. Meteorit. Planet. Sci. 42, 1529–1548. Earth Planet. Sci. Lett. 284, 168–178. Schersteń A., Elliott T., Hawkesworth C., Russell S. S. and Trinquier A., Birck J. L., Alle`gre C. J., Go¨pel C. and Ulfbeck D. Masarik J. (2006) Hf–W evidence for rapid differentiation of (2008) 53Mn–53Cr systematics of the early Solar System iron meteorite parent bodies. Earth Planet. Sci. Lett. 241, 530– revisited. Geochim. Cosmochim. Acta 72, 5146–5163. 542. Wasserburg G. J., Tera F., Papanastassiou D. A. and Huneke J. C. Schiller M., Baker J. A. and Bizzarro M. (2010) 26Al–26Mg dating (1977) Isotopic and chemical investigations on Angra dos Reis. of asteroidal magmatism in the young Solar System. Geochim. Earth Planet. Sci. Lett. 35, 294–316. Cosmochim. Acta 74, 4844–4864. Weiss B. P., Berdahl J. S., Elkins-Tanton L., Stanley S., Lima E. A. Shirai N. and Humayun M. (2010) Initial tungsten isotopic and Carporzen L. (2008) Magnetism on the angrite parent body compositions for angrites obtained from phosphates. Lunar and the early differentiation of planetesimals. Science 322, 713– Planet. Sci. XLI. #2642 (abstr.). 716. Shukolyukov A. and Lugmair G. W. (2007) The Mn–Cr isotope systematics of bulk angrites. Lunar Planet. Sci. XXXVIII. Associate editor: Richard J. Walker #1432 (abstr.). Shukolyukov A. and Lugmair G. W. (2008) Mn–Cr chronology of eucrite CMS 04049 and angrite NWA 2999. Lunar Planet. Sci. XXXIX. #2094 (abstr.).