Oxygen and magnesium isotope compositions of calcium-aluminum-rich inclusions from Rumuruti (R) chondrites

Transcription

1 1 Oxygen and magnesium isotope compositions of calcium-aluminum-rich inclusions from Rumuruti (R) chondrites S. S. Rout 1*, A. Bischoff 1, K. Nagashima 2, A. N. Krot 2, G. R. Huss 2, and K. Keil 2 1 Institut für Planetologie, Westfälische Wilhelms Universität Münster, Wilhelm Klemm Strasse 10, 48149, Münster, Germany 10 2 School of Ocean, Earth Sciences and Technology, Hawai i Institute of Geophysics and Planetology, 1680 East West Road, Honolulu, HI 96822, USA To be submitted to: Geochimica et Cosmochimica Acta * address of the corresponding author: suryarout@uni-muenster.de ABSTRACT. We report oxygen and magnesium isotope compositions of Ca,Al-rich inclusions (CAIs) from several Rumuruti (R) chondrites measured in situ using a Cameca ims-1280 ion microprobe. On a three isotope oxygen diagram, compositions of individual minerals in most R CAIs analyzed fall along a slope 1 line. Based on the variations of Δ 17 O values within an inclusion, the R CAIs are divided into (i) 16 O-rich (Δ 17 O ~ 23 to 26 ), (ii) uniformly 16 O- depleted (Δ 17 O ~ 2 ), and (iii) isotopically heterogeneous (Δ 17 O ranges from 25 to +5 ). One of the hibonite-rich CAIs, H030/L, has an intermediate Δ 17 O value of 12 and a highlyfractionated composition (δ 18 O ~ +47 ). We infer that like most CAIs in other chondrite groups, the R CAIs formed in an 16 O-rich gaseous reservoir. The uniformly 16 O-depleted and isotopically heterogeneous CAIs subsequently experienced oxygen isotope exchange during remelting in an 16 O-depleted nebular gas, possibly during chondrule formation, and/or during fluid-assisted thermal metamorphism on the R chondrite parent asteroid. Three hibonite-bearing CAIs and one spinel plagioclase-rich inclusion were analyzed for magnesium isotope compositions. The CAI with the highly-fractionated oxygen isotopes, H030/L, shows a resolvable excess of 26 Mg ( 26 Mg*) corresponding to the initial 26 Al/ 27 Al ratio of ~ Three other CAIs show no resolvable excess of 26 Mg. The absence of 26 Mg* in the spinel plagioclaserich CAI found in a metamorphosed R chondrite could have resulted from metamorphic resetting. Two other hibonite-bearing CAIs occur in the R chondrites which appear to have experienced only minor degree of thermal metamorphism. These inclusions could have formed from the precursors with the lower than the canonical 26 Al/ 27 Al ratio. -1-

2 INTRODUCTION The Rumuriti-type (R) chondrites are mineralogically, chemically and isotopically distinct from carbonaceous, ordinary, and enstatite chondrites (e.g., Schulze et al., 1994; Bischoff et al., 1994; Rubin and Kallemeyn, 1994; Kallemeyn et al., 1996). The major chemical, mineralogical and petrographic characteristics of R chondrites are the following: (i) similar refractory and moderately volatile lithophile element abundances as in ordinary chondrites, except for S, Se, and Zn; (ii) high degree of oxidation, which can be inferred from high fayalite content in olivine (38 40 mol%), low modal abundance of metallic iron, and high modal abundance of olivine; (iii) presence of Ti Fe 3+ -rich chromite and platinum group element-rich phases; (iv) matrix/chondrule modal abundance similar to carbonaceous (CO, CR, and CV) chondrites; (v) intermediate abundance of chondrules with non-porphyritic texture (barred-olivine, radialpyroxene and cryptocrystalline) chondrules between those in carbonaceous and ordinary chondrites, and (vi) high values of Δ 17 O = δ 17 O 0.52 δ 18 O (e.g., Schulze et al., 1994; Bischoff et al., 1994; Kallemeyn et al., 1996; Schulze, 1998). On a three isotope oxygen diagram, compositions of most R chondrites plot along mass-fractionation line with a slope of ~0.52 and Δ 17 O value of ~2.9. Some R chondrites deviate from this line and have a lower Δ 17 O value, probably due to the terrestrial alteration in cold and hot deserts (e.g., Bischoff et al., 1994; Schulze et al., 1994; Kallemeyn et al., 1996; Weisberg et al., 1991; Greenwood et al., 2000; Pack et al., 2004). More than half of the R chondrites are monomict regolith breccias containing different types of light and dark clasts, of different petrologic types of R chondrites, (Bischoff, 2000) and solar wind implanted gases (e.g., Bischoff and Schultz, 2004; Bischoff et al., 1994, 2006). Only a few detailed isotopic studies have been done so far on chondrules, CAIs, and individual minerals in R chondrites (e.g., Weisberg et al., 1991; Greenwood et al., 2000; Bischoff and Srinivasan, 2003; Pack et al., 2004). Weisberg et al., (1991) was the first to analyze bulk oxygen isotope compositions of chondrules separated from the R chondrites Allan Hills (ALH) 85151, Yamato (Y) , and Carlisle Lake. Most chondrules from R-chondrites plot along a mass-fractionation line with Δ 17 O value of ~2.7. The only exception is a chondrule from Carlisle Lake that has a lower Δ 17 O value (0.68 ) and plots close to the ordinary chondrite -2-

3 field. This, along with mineralogical observations, led Weisberg et al. (1991) to suggest a possible link between the R and ordinary chondrites. Greenwood et al. (2000) reported oxygen isotope compositions of magnetite grains in the R chondrite Pecora Escarpment (PCA) On a three isotope oxygen diagram, these magnetites and the magnetites from the ordinary chondrite Semarkona (LL3.0) plot along the same line with a slope of ~0.7, suggesting that magnetites in both meteorites resulted from oxidation of metal by water of a similar oxygen isotope composition. Greenwood et al. (2000) also measured oxygen isotope compositions of pyroxene and olivine in a porphyritic olivine pyroxene chondrule and several olivine clasts with variable fayalite content from PCA The chondrule silicates have Δ 17 O values only slightly lower than that of the bulk host meteorite, whereas the olivine clasts are more 16 O-enriched (Δ 17 O 2 ). Pack et al. (2004) reported a systematic oxygen isotope study of the refractory forsterite grains in several carbonaceous and ordinary chondrites, as well as in the R chondrite Dar al Gani 013. On a three isotope oxygen diagram, the refractory forsterite grains, the outer non-refractory fayalitic rim, the bulk composition of Dar al Gani 013, and chondrule silicates from PCA reported by Greenwood et al. (2000) plot along a single line with a slope of 0.8, named the Rumuruti Mixing Line (RML). Recently, Rout and Bischoff (2007, 2008a,b) described the mineralogy and petrography of the Al-rich objects, chondrules and CAIs, in several R chondrites. They showed that (i) the abundance of R CAIs is intermediate between those in ordinary and carbonaceous chondrites. (ii) The spinel-bearing CAIs are a dominant type of refractory inclusions in R chondrites. Hibonite is a rare phase; corundum, melilite, and grossite are not observed. (iii) Most CAIs in R chondrites contain abundant secondary minerals resulting from the extensive parent body and/or nebular alteration processes. The mineralogically pristine CAIs are exceptionally rare. In this paper, we report oxygen and magnesium isotope compositions of eleven CAIs described by Rout and Bischoff (2007, 2008a,b). SAMPLES AND ANALYTICAL TECHNIQUES 140 Eleven CAIs in 9 polished thin and thick sections of 6 R chondrites (Table 1) were studied using optical and scanning electron microscopy (SEM), electron microprobe, and secondary ion mass-spectrometry (SIMS). Backscattered electron images of the CAIs were obtained with a -3-

4 JEOL LV5900 SEM at the University of Hawai i, with an accelerating voltage of 15 kv. Quantitative mineral compositions were obtained with a JEOL 8900 electron microprobe at the University of Münster, Interdisciplinary Center for Electron Microscopy and Microanalysis (ICEM; Table 2). The electron microprobe was operated at 15 kv; a probe current of 15 na was used to analyse the minerals with a beam diameter of ~1 µm. Natural and synthetic minerals of well-known compositions were used as standards: jadeite (Na), sanidine (K), diopside (Ca), apatite (P), V-metal (V), disthen (Al), fayalite (Fe), chromite (Cr), flourite (F), NiO (Ni), hypersthene (Si), forsterite (Mg), rhodonite (Mn), rutile (Ti), tugtupite (Cl), scapolite (S), and willemite (Zn). The matrix corrections were made by the Фρ(z) procedure of Armstrong (1991). Oxygen isotope analyses were carried out with the Cameca ims-1280 ion microprobe at the University of Hawai i. CAIs and the constituent minerals analyzed for oxygen isotopes are listed in Table 1. A focused Cs + primary beam of 200 pa was used to raster regions 7 7 µ m for presputtering. Oxygen isotopes were measured in a raster mode from central regions ~5 5 µm in size. Three oxygen isotopes were measured by a combination of multicollection mode and peak jumping. 16 O and 17 O were measured simultaneously using multi-collection Faraday Cup (FC) and mono-collection electron multiplier (EM), respectively. 18 O was measured with monocollection EM by peak jumping. Each measurement consisted of 30 cycles and took ~8 min. Mass resolving power (MRP) for 17 O was set to ~5600, sufficient to separate 17 O from the interference 16 OH signal. The 18 O secondary ion intensity was between and counts per second. Deadtime corrections were made with deadtime of 30.5 ns estimated using titanium isotopes (Fahey et al., 1987). The instrumental mass fractionation (IMF) was corrected using the terrestrial standards: Burma spinel for spinel and hibonites; Cr-augite for Al-rich diopside, diopside, and fassaite; and San Carlos olivine for olivine. The reproducibility of measurements of the standards were ~ ±1.9 for Burma spinel, ±2.1 for Cr-augite and ±1.8 for San Carlos olivine. Magnesium and aluminum isotope compositions were measured in situ with the University of Hawai i Cameca ims-1280 ion microprobe. The mass spectrometer was operated at +10 kev with a 50 ev energy window. Minerals with high 27 Al/ 24 Mg ratios (anorthite and hibonite) were analyzed with a focused 5 7 µm 16 O primary beam in monocollection mode using EM and FC detectors for magnesium and aluminum isotopes, respectively. A primary current of 150 or 300 pa was used depending on concentration of magnesium in a mineral analyzed. 24 Mg +, 25 Mg + and -4-

5 26 Mg +, and 27 Al + were measured in 120 cycles on a monocollector EM and FC in a peak jumping mode for 4, 10, 10, and 2 sec, respectively. The MRP was set to ~3800, sufficient to separate interfering hydrides and doubly charged 48 Ca ++. Automated centering of the secondary beam in the field aperture of the mass spectrometer, high-voltage offset control, and mass-peak centering were applied before each measurement. Offset control and peak centering were also applied at cycle 60. Magnesium isotope compositions are reported as per mil deviations (Δ 25 Mg and Δ 26 Mg) from the terrestrial magnesium isotope ratios [ 25 Mg/ 24 Mg = ; 26 Mg/ 24 Mg = (Catanzaro et al., 1966)] and as an excess or a deficit of 26 Mg (δ 26 Mg*). Instrumental mass fractionation was corrected by standard-sample bracketing by comparing each measurement with the Δ 25 Mg and Δ 26 Mg measured in the terrestrial Madagascar hibonite and Miyakejima anorthite. Some of the igneous CAIs have an intrinsic isotopic mass fractionation that probably resulted from evaporation during melting. The intrinsic mass fractionation (F Mg ) was calculated by subtracting the mean Δ 25 Mg for the appropriate standard mineral from Δ 25 Mg measured for each sample. This intrinsic mass fractionation was propagated to Δ 26 Mg using an exponential law with an experimentally defined mass fractionation exponent of (Davis et al., 2005). Excesses or deficits of 26 Mg for CAIs and chondrule were calculated by comparing the measured Δ 26 Mg with value calculated from this law. The choice of the mass fractionation law is not critical to the data set obtained in this study because the degree of intrinsic mass fractionation is relatively small for most samples, and because the Al/Mg ratios of the minerals that control the 26 Al 26 Mg isochrons are sufficiently high that the uncertainty introduced by the choice of law is insignificant. The reported uncertainties include both the internal precision of an individual analysis and the external reproducibility for standard measurements during a given analytical session. The relative sensitivity factors for aluminum and magnesium were determined from the 27 Al + / 24 Mg + measured by SIMS and the Al/Mg ratios determined previously for each standard mineral using electron microprobe

6 RESULTS MINERALOGY 240 A detailed description of a mineralogy of CAIs in R chondrites has been reported by Rout and Bischoff (2007, 2008a,b). Here, we briefly summarize only the mineralogy of the CAIs analyzed for oxygen and magnesium isotope compositions. Spinel-rich CAIs CAI D013/44 has a spinel fassaite core surrounded by a monomineralic diopside rim (Figs. 1a c). Spinel grains near the CAI periphery are enriched in FeO (~2.3 wt%) compared to those in the core. CAI D013B/118 has a spinel fassaite perovskite core surrounded by a double-layered rim of fassaite and diopside (Figs. 1d f). The inner, fassaite layer is intergrown with spinel of the CAI core. Some islands of fassaite are also present in the spinel core. The presence of unaltered perovskite and low FeO content in spinel in the CAI core (~1 wt%; Table 2) indicate that the inclusion escaped significant alteration. However, spinel in the outer portion of the CAI have high FeO content (~7 wt%). CAI 2446D/33B is an irregularly-shaped CAI fragment composed of FeO-rich spinel and minor fassaite and secondary Na,Al-rich minerals (Fig. 2). The CAI is surrounded by a discontinuous rim of diopside. CAI Dfr/53 is the largest ( µm) inclusion found in R chondrites (Fig. 3). A thin rim of diopside overlies a spinel fassaite core that contains abundant secondary Na,Al-rich minerals and rare inclusions of perovskite. Spinel has relatively high FeO content (~12 15 wt%), whereas fassaite is almost devoid of it (~ wt% FeO; Table 2). Spinel plagioclase-rich CAIs CAI 753/6L consists mainly of spinel and plagioclase; it is surrounded by a thick rim of diopside (Figs. 4a,b). A complex mixture of fine-grained spinel and fassaite is present below the diopside rim. Minor Al-rich alteration products are also found within the inclusion. Most spinel grains surrounded by anorthite are euhedral, suggesting crystallization from a melt. Anorthite -6-

7 (An ) contains low Na 2 O (<0.2 wt%) and FeO contents (~0.85 wt%), whereas spinel is enriched in FeO (~20 wt%; Table 2). CAI 753C/51 consists of euhedral spinel grains in a plagioclase groundmass (Figs. 4c,d). Plagioclase in the CAI core has low Na 2 O content; plagioclase near the CAI periphery is enriched in Na 2 O and FeO and almost Mg-free (Table 2). The spinel grains are hercynitic in composition (Table 2). Hibonite-rich CAIs CAI H030/L consists of large euhedral crystals of hibonite and small FeO-rich (17 wt%) spinel grains; it is rimmed by a thin layer of diopside (Figs. 5a,b). Hibonite has low MgO (<0.02 wt%) and TiO 2 (~ wt%) contents (Table 2). CAI 2446D/2L is a concentrically zoned inclusion (Figs. 5c f). Its hibonite core is overlain by fine-grained spinel with many ilmenite grains scattered within it along with alteration products. The CAI is rimmed by diopside. The hibonites have low TiO 2 (1.4 wt%) and FeO (0.75 wt%) contents (Table 2). The spinel grains are too fine-grained to be analyzed quantitatively by electron microprobe. Fassaite-rich spherules CAI Dfr/23 is an unrimmed fassaite-rich spherule containing euhedral spinel grains and skeletal olivine grains (Figs. 6a e). The fassaite groundmass shows a chemical zoning in Al 2 O 3 and TiO 2. Although the spinel and olivine grains are compositionally uniform and have high FeO contents (Table 2), there are significant differences in FeO contents between the individual olivine grains, which range from ~25 to ~34 wt%. CAI 753B/14 is an unrimmed spherule composed of a large spinel grain surrounded by fassaite (Fig. 6f). The fassaite has low TiO 2 (~1.5 wt%) and high Al 2 O 3 (~13 wt%) contents; the spinel grain has high FeO content (~19 wt%; Table 2). CAI 1476/124 is an unrimmed fassaite hibonite spherule (Fig. 7). The fassaite is chemically zoned in Al 2 O 3 (16.3 to 27 wt%) and TiO 2 (up to 2.7 wt%). The hibonite contains low MgO content (0.96 wt%) (Table 2). -7-

8 OXYGEN ISOTOPES Oxygen isotope compositions of 10 CAIs from R chondrites measured are listed in Table 3 and plotted in Figure 8. Based on the oxygen isotope data of individual minerals in an inclusion, the R CAIs can be divided into four groups: (i) 16 O-rich; (ii) 16 O-depleted, (iii) isotopically heterogeneous, and (iv) a CAI with a highly fractionated composition. On a three isotope oxygen diagram, CAIs of the first three groups plot slightly off the Allende CAI line and follow line with a slope of ~1, named by Young and Russell (1998) as a primary nebular line (Fig. 8a). The 16 O-rich CAIs include the unrimmed fassaite hibonite spherule 1476/124 (Fig. 7), the spinel-rich CAIs rimmed by diopside, 2446D/33B (Fig. 2) and D013/44 (Figs. 1a c) rimmed by diopside. Within 2σ uncertainty of our measurements, these CAIs have similar Δ 17 O values ranging from 22.5 to 25.5 (Table 3; Fig. 8). It is unknown whether the CAI 2446D/33B is isotopically uniform, because only spinel grains were analyzed. The 16 O-depleted CAIs include two unrimmed fassaite spinel±olivine spherules, Dfr/23 (Figs. 6a e) and 753B/14 (Fig. 6f). Within 2σ uncertainty of our measurements, these CAIs have similar Δ 17 O values ranging from 1.6 to 3.3 (Table 3; Fig. 8). These compositions are similar to those of chondrules in carbonaceous chondrites (e.g., Krot et al., 2006; Yurimoto et al., 2008 and references therein). The isotopically heterogeneous CAIs are the most common type of inclusions in R chondrites. These include the heavily-altered, spinel-rich CAI Dfr/53 rimmed by diopside (Fig. 3), the heavily-altered, hibonite-rich CAI 2446D/2L rimmed by diopside (Figs. 5c f), the spinelrich CAI D013B/118 rimmed by diopside (Figs. 1d f), and the spinel plagioclase-rich CAI 753/6L rimmed by diopside (Figs. 4a,b). The spinel and hibonite grains in these CAIs have 16 O- rich compositions, with Δ 17 O values ranging from 20.5 to The fassaite grains and diopside rims are 16 O-depleted to varying degrees: Δ 17 O value ranges from 13.8 to +5.1 (Table 3; Fig. 8). Fassaite of the CAI Dfr/53 has the heaviest oxygen isotope composition reported in CAIs (Δ 17 O ~ +5.0±2.5 ). Within 2σ uncertainty, this value is similar to the bulk oxygen isotope compositions of the R chondrites and chondrules (Weisberg et al., 1991; Bischoff et al., 1994; Schulze et al., 1994; Kallemeyn et al., 1996; Greenwood et al., 2000; Pack et al., 2004) (Fig. 9). -8-

9 The hibonite spinel-rich CAI H030/L rimmed by diopside (Figs. 5a,b) is the only inclusion with a highly fractionated oxygen isotope composition: Δ 17 O ~ 12, δ 17 O ~ +12, and δ 18 O ~ +47 (Table 3; Fig. 8). Only hibonite grains in this inclusion were analyzed. Oxygen isotope compositions of these hibonites are similar to the HAL-like hibonite grains in other chondrite groups reported by Lee et al. (1980 a,b), Ireland et al. (1992), and Ushikubo et al. (2007). 380 ALUMINUM-MAGNESIUM ISOTOPE SYSTEMATICS Magnesium isotopes were analyzed in four CAIs containing minerals with high 27 Al/ 24 Mg ratio, hibonite or plagioclase, 2446D/2L, 1476/124, H030/L, and 753B/51. The data obtained are listed in Table 4 and shown in Figure 10. The resolvable excess of 26 Mg ( 26 Mg*) due to decay of 26 Al is detected only in one of these CAIs, H030/L. The 26 Mg* correlated with Al/Mg ratio and define an internal isochron corresponding to the initial 26 Al/ 27 Al ratio [( 26 Al/ 27 Al) 0 ] of (7.1±1.0) 10 7, which is much lower than the canonical ratio of (4 5) 10 5 typical for most CAIs from primitive (unmetamorphosed) chondrites. The inferred ( 26 Al/ 27 Al) 0 in this CAI is slightly lower than the value reported by Bischoff and Srinivasan (2003) in the same inclusion [( 26 Al/ 27 Al) 0 = (1.4±0.3) 10 6 ]. The upper limits for the ( 26 Al/ 27 Al) 0 in the CAIs 2446D/2L, 1476/124, and 753B/51 are <3 10 6, <5 10 6, and <5 10 7, respectively (Table 4; Fig. 10). Some of the analyses of hibonite grains in the CAIs 2446D/2L, 1476/124, and H030/L are characterized by the isotopically light magnesium (resolvable from the terrestrial value within 2σ uncertainties): F Mg are up to 4.7, 3.2, and 8.4 respectively (Table 4). Magnesium isotope composition of anorthite in the CAI 753B/51 is normal and no excess of 26 Mg* is found (Table 4). DISCUSSION OXYGEN ISOTOPE COMPOSITIONS OF THE R CHONDRITE CAIS: IMPLICATION FOR THE INITIAL 410 ISOTOPE COMPOSITION OF CAIS AND THE NATURE OF THE SUBSEQUENT ISOTOPE EXCHANGE The oxygen isotope compositions of the R chondrite CAIs are generally similar to those in other chondrite groups (e.g., Yurimoto et al., 2008 and references therein; Makide et al., 2008). (i) On a three isotope oxygen diagram, all CAIs studied, except H030/L, plot along massindependent fractionation line with a slope of ~1. (ii) Spinel, hibonite, and fassaite in most of -9-

10 them are 16 O-rich (Δ 17 O < 20 ). (iii) There are only two igneous CAIs uniformly 16 O-depleted (Δ 17 O ~ 2 ). (iv) CAIs with highly-fractionated isotope compositions are rare. However, compared to CAIs in primitive (unmetamorphosed) chondrites (CO3.0, CR2, CM2), those in R chondrites often show oxygen isotope heterogeneity. The number of the isotopicallyheterogeneous CAIs would possibly be even higher if fine-grained pyroxenes and plagioclase were measured. In addition, none of the studied CAIs contains melilite (probably as a result of alteration and thermal metamorphism) a mineral typically showing 16 O-depletion in isotopically heterogeneous CAIs from CV chondrites (Yurimoto et al., 2008 and references therein). The fact that 7 out of 10 R CAIs measured contain minerals having 16 O-rich compositions (Table 3; Fig. 8) indicates that the R CAIs originated in an 16 O-rich gaseous reservoir (Krot et al., 2002), like most CAIs in ordinary, enstatite and carbonaceous chondrite groups (McKeegan et al., 1998; Sakai and Yurimoto, 1999; Guan et al., 2000a,b; Fagan et al., 2001; Itoh et al., 2004; Makide et al., 2008). It has been previously shown that most CAIs in primitive carbonaceous chondrites, e.g., CM2 (Sakai and Yurimoto, 1999), CO3.0 (Itoh et al., 2004), CH3.0 (Krot et al., 2008a), and CR2 (Makide et al., 2008), are uniformly 16 O-rich. In contrast, many R CAIs studied are isotopically heterogeneous, suggesting subsequent isotope exchange in an 16 O-poor external reservoir. This exchange could have occurred in the solar nebula, e.g., during CAI remelting (Yurimoto et al., 1998; Aléon et al., 2002; Itoh et al., 2004; Krot et al., 2005, 2008a), and/or during fluid-assisted thermal metamorphism on the R chondrite parent body, as was proposed for the majority of CAIs from metamorphosed CV chondrites (Krot et al., 2007). The isotopically heterogeneous CAIs, Dfr/53, 2446D/2L, D013B/118, and 753/6L, contain 16 O-rich spinel and hibonite (Δ 17 O < 20 ) and 16 O-depleted Al,Ti-diopside (Δ 17 O ranges from 13.8 to +5.1 ). Since the pyroxenes, 16 O-depleted to varying degrees, are found in the outer portions of the CAIs 2446D/2L, D013B/118, and 753/6L, which show clear evidence for postcrystallization alteration and thermal metamorphism, this depletion is most likely due to fluidassisted thermal metamorphism on the R chondrite parent asteroid, rather than thermal processing in the solar nebula, although the latter cannot be entirely excluded. For example, the similar 16 O-depletion of diopside from the Wark-Lovering rims has been previously reported in an isotopically-heterogeneous CAI from Efremovka (CV ) (Krot et al., 2002). Ito and Messenger (2008) found three zones with oscillating oxygen isotope compositions ( 16 O-poor -10-

11 16 O-rich 16 O-poor) within the Wark-Lovering rim sequence around an Allende (CV>3.6) coarse-grained CAI. These variations have been attributed to fluctuations in oxygen isotopic composition of the nebular gas in the CAI-forming region (e.g., Itoh and Yurimoto, 2003). Recently, Fagan et al. (2007) described an 16 O-depleted diopside rim around an 16 O-rich hibonite-spinel-rich CAI in the most primitive carbonaceous chondrite, Acfer 094. They suggested that the rim either condensed from an 16 O-depleted nebular gas or experienced preferential melting and isotope exchange during late-stage reheating. Fassaite in Dfr/53 has the highest degree of oxygen isotope heterogeneity: spinel in the CAI core is 16 O-rich (Δ 17 O ~ 25 ), whereas the coexisting fassaite is 16 O-poor (Δ 17 O ~ +5.0±2.5 ). The latter composition is similar to those of bulk chondrules in R chondrites (Fig. 9). Although this CAI is rather extensively altered, it seems likely that it experienced partial melting and isotope exchange during formation of the R chondrite chondrules. The spinel grains must have preserved the original 16 O-rich composition, consistent with a slow dissolution rate of spinel in the CAI melts (Stolper, 1982; Stolper and Paque, 1986; Beckett and Stolper, 1994) and the slow rate of oxygen self-diffusion in spinel (Ryerson and McKeegan, 1994). Two igneous CAIs, the unrimmed fassaite-rich spherules Dfr/23 and 753B/14, are uniformly 16 O-depleted (Fig. 8) to a level commonly observed in chondrules from carbonaceous chondrites (e.g., Krot et al., 2006 and references therein). Similarly 16 O-depleted CAIs have been previously reported in enstatite (Guan et al., 2000a,b), CB (Krot et al., 2001a), CH (Krot et al., 2008a), and CR (Krot et al., 2005) chondrites. We infer that these CAIs experienced complete melting and isotope exchange in an 16 O-poor nebular gas, most likely during chondrule formation. This interpretation is consistent with the lack of rims around these CAIs (most CAIs are surrounded by the Wark-Lovering rims) and with the presence of skeletal olivine crystals, indicative of fast cooling, in one of them (Fig. 6a). It is not clear whether these spherules experienced melting and incomplete isotope exchange during formation of R chondrules, which have much heavier oxygen isotope compositions (Fig. 9), or the melting and isotope exchange occurred prior to formation of R chondrules. In summary, CAIs in R chondrites originated in an 16 O-rich gaseous reservoir and subsequently experienced isotope exchange in an 16 O-poor reservoir. The isotope exchange in some CAIs clearly occurred during melting, possibly during chondrule formation, whereas isotope exchange in several other CAIs could have occurred either during fluid-assisted thermal -11-

12 510 metamorphism on the R chondrite parent body or during brief reheating events in the solar nebula. CAI WITH HIGHLY-FRACTIONATED OXYGEN COMPOSITION CAI H030/L is moderately 16 O-depleted (Δ 17 O ~ 13 ) and characterized by the highly fractionated oxygen isotope composition (δ 18 O ~ +47 ), which is similar to those observed in the HAL-like inclusions. The highest fractionation of oxygen isotopes (δ 18 O up to +73 ; Δ 17 O ~ 0.5 ) was found in hibonite of a relict CAI inside a porphyritic Type I chondrule in the CR chondrite Acfer 209 (Nagashima et al., 2008). Based on the oxygen isotope compositions of the Allende FUN inclusion HAL, Lee et al. (1980 a,b) proposed a back-reaction model in which the FUN and non-fun CAIs formed from a single 16 O-rich reservoir; the FUN CAI melts experienced distillation and mass-dependent fractionation of oxygen isotopes followed by gas-melt isotope exchange to varying degrees with an 16 O-poor gas. According to this model, H030/L formed by melting of 16 O-rich precursors, similar in composition to typical CAIs (δ 17,18 O ~ 50 ); it experienced melt evaporation and fractionation of oxygen isotopes (along line AB in Fig. 8a). Because hibonite is 16 O-depleted compared to typical CAIs (Δ 17 O ~ 13 and ~ 25 ), it must have subsequently experienced isotope exchange with an 16 O-poor nebular gas. Because of the fine-grained nature of the CAI, neither spinel no diopside in this CAI were analyzed. As a result, it is not possible to infer a path of isotope exchange of H030/L. If we consider the most 16 O-depleted phase among the R CAIs studied fassaite in the CAI Dfr/53 as a sample of the 16 O-poor nebular reservoir, the isotope exchange in the hibonite H030/L may have occurred along line BC as illustrated in Figure 8a. This melt distillation isotope exchange hypothesis, however, appears to be inconsistent with the isotopically normal magnesium in the H030/L hibonites (Table 4), unless subsequent magnesium isotope exchange is invoked (see below). An alternative model to describe the oxygen isotopic fractionation within H030/L can be proposed. The H030/L precursors may be 16 O-rich initially and later underwent oxygen isotopic exchange along the Allende CAI mixing line (line AD in Fig. 8a) or was initially 16 O-depleted and plotted at the point D in Fig. 8a. Later distillation process further fractionated the oxygen isotopic composition from the point D to F along a mass-dependent fractionation line with Δ 17 O ~ Similar oxygen isotopic exchange processes might be also involved in Allende F CAI -12-

13 TE (El Goresy et al., 1991; Krot et al., 2008b) and the FUN CAI, #3 discovered in CR carbonaceous chondrite, Gao Guenie (b) (Makide et al., 2008). OXYGEN ISOTOPIC EVOLUTION OF THE R CHONDRITE-FORMING REGION Bulk R chondrites and their chondrules are 16 O-poor (e.g., Greenwood et al., 2000; Weisberg et al., 1991; Schulze et al., 1994; Bischoff et al., 1994; Kallemeyn et al., 1996) and record a completely different oxygen isotope composition compared to the R chondrite CAIs, indicating formation in isotopically distinct nebular regions (Figs. 9a,b). The oxygen isotopic composition of the R-chondrite forming region or the nebula certainly changed with time as also inferred from other chondritic groups (e.g., Krot et al., 2005; Yurimoto et al., 2008; Krot et al., 2006). It is generally believed that CAIs formed in an 16 O-rich nebular region with ambient temperature (> 1350 K), possibly near the protosun, <1 AU (e.g., Shu et al., 1996; Krot et al., 2001b, 2008a). Such high temperatures can only be attained during the early stages (class 0 I) of the Sun evolution, when the mass accretion rate was high, ~ solar mass per year (e.g., Morfill, 1988; Cassen, 1994; Boss, 1998, 2003). Early formation of CAIs within a short period of time is supported by the recent high precision 207 Pb 206 Pb and 26 Al 26 Mg isotope measurements (Amelin et al., 2002; Jacobsen et al., 2008). We suggest that CAIs in R chondrites formed near the protosun in the presence of 16 O-rich nebular gas and were subsequently transported to a region depleted in 16 O, where R chondrules formed and R chondrites accreted. Some of the R CAIs were remelted in this region, which resulted in oxygen isotope exchange and formation of 16 O-depleted CAIs and, possibly, isotopically heterogeneous CAIs. The outward transport of the CAIs can be due to the outward radial diffusion in a turbulent disk (e.g., Cuzzi and Davis, 2003; Cuzzi et al., 2003), x wind (Shu et al., 1996), or photophoresis (e.g., Krauss and Wurm, 2005; Wurm and Krauss, 2006; Krauss et al., 2007; Teiser et al., 2008; Wurm et al., 2008; Haack and Wurm, 2008; Wurm and Haack, 2008). Photophoresis has been successfully shown to be an efficient mechanism to transport CAIs radially away from the Sun in an optically thin, gas-rich disk (Wurm and Krauss, 2006). The high-temperature CAI-forming region could have been dust depleted due to its sublimation. The photophoresis may have prevented CAIs from moving towards the Sun and transported them outward to dust-enhanced nebular regions. Alternatively, Wurm and Haack (2008) suggested that -13-

14 610 CAIs could have been transported by photophoresis over the disk surface during FU Orionis outbursts while the disk is still optically thick. The origin of 16 O-depleted nebular region has been recently discussed in the light of the CO self-shielding model (e.g., Yurimoto and Kuramoto, 2004; Lyons and Young, 2005; Sakamoto et al., 2007) and numerical models of the evolution of a viscous protoplanetary disk (Cuzzi and Zahnle, 2004; Ciesla and Cuzzi, 2006; Ciesla, 2007, 2008). According to these models, a significant enrichment in 17,18 O-enriched water can be achieved near the snow line during the early stages of protoplanetary disk evolution due to inward migration of icy particles. The 17,18 O-enriched compositions of R chondrules and high oxidation state of R chondrules and matrices may indicate that they formed near the snow line, in a region with enhanced abundance of 17,18 O-rich water vapor. ALUMINUM-MAGNESIUM ISOTOPE SYSTEMATICS OF THE R CHONDRITE CAIS Among the four CAIs analyzed for magnesium isotopes, only one hibonite-rich CAI, H030/L, shows a resolvable excess of 26 Mg ( 26 Mg*) corresponding to the initial 26 Al/ 27 Al ratio of (7±1) 10 7, which is significantly lower than the canonical ratio of (4 5) Two other hibonite-bearing inclusions, 2446D/2L and 1476/124, and a spinel plagioclase-rich inclusion, 753B/51, show no resolvable 26 Mg*. The 26 Al-poor, hibonite-bearing CAIs in R chondrites share some similarities with hibonite-rich inclusions in CM chondrites (Ireland, 1988,1990) and pyroxene hibonite spherules in several chondrite groups (MacPherson et al., 1989; Ireland et al., 1991; Russell et al., 1998; Simon et al., 1998). Below we compare the mineralogy and isotope compositions of the hibonite-rich CAIs from R chondrites with those in other chondrite groups and discuss their possible origin. Hibonite-rich CAI 2446D/2L 640 CAI 2446D/2L has a hibonite-rich core composed of platy hibonite crystals and surrounded by a spinel-rich mantle with numerous inclusions of ilmenite and a diopside rim. This CAIs is texturally similar to some of the spinel hibonite±perovskite (SHIBs) inclusions in CM chondrites (e.g., Ireland, 1988, 1990). However, the low TiO 2 content (< 2 wt%), the absence of the resolvable 26 Mg*, and the high 27 Al/ 24 Mg ratio in the 2446D/2L hibonite grains make this inclusion more similar to the platy hibonite crystals (PLACs) than to SHIBs (Ireland 1988, 1990; -14-

15 Hinton et al., 1988). On the contrary, the magnesium mass fractionation towards the lighter isotopes is not seen in the PLACs, rather the SHIBs show large variation in intrinsic isotopic fractionation. A metamorphic redistribution of magnesium isotopes in 2446D/2L could have resulted in the low δ 26 Mg values and increase of the 27 Al/ 24 Mg ratio. This metamorphic event was probably responsible for the high FeO contents in the spinels, transformation of perovskites to ilmenites, and relatively high FeO contents in the hibonites (Table 2). Based on our observations, we cannot conclude unambiguously whether the lack of resolvable excess of 26 Mg in 2446D/2L is due to metamorphic resetting or this CAI originated without live 26 Al. Hibonite-rich CAI H030/L CAI H030/L shows the highest excess of 26 Mg among the R CAIs measured. The highlyfractionated oxygen isotope composition of this inclusion and the high 27 Al/ 24 Mg ratio of its hibonite (up to ~57,000; Table 4) are similar to those in F (fractionated)- and HAL-like CAIs (e.g., Lee et al., 1979, 1980a,b; Allen et al., 1980; Davis et al., 1982; Hinton and Bischoff, 1984; Ireland and Compston, 1987; Ireland et al., 1988, 1992; Russell et al., 1998; Ushikubo et al., 2007). The HAL-like inclusions are characterized by the large mass-dependent fractionation of oxygen, calcium, and titanium isotopes, depletion in cerium, ytterbium, and europium, normal magnesium isotope compositions, and resolvable 26 Mg* corresponding to a much lower ( 26 Al/ 27 Al) 0 compared to the canonical value (e.g., Lee et al., 1976; MacPherson et al., 1995). A comparison of the isotopic data published for six HAL-like inclusions analyzed so far is given in Table 5. H030/L has the second highest 27 Al/ 24 Mg ratio among all the HAL-like inclusions. It shows no large degree of intrinsic fractionation of magnesium isotopes. Ushikubo et al. (2007) reported the canonical ( 26 Al/ 27 Al) 0 value of (5.3±0.2) 10 5 in a HAL-like CAI from Kainsaz (CO3), but the ( 26 Al/ 27 Al) 0 values in all other HAL-like inclusions, including H030/L, are much lower (Table 5). Assuming a uniform distribution of 26 Al in the early Solar System (e.g., Thrane et al., 2006; Jacobsen et al., 2008), the inferred ( 26 Al/ 27 Al) 0 of H030/L represents a time difference of about ~4.5 Myr between the closure of its Al Mg isotope systematics and the formation of CAIs with the canonical ( 26 Al/ 27 Al) 0. The similar value of ( 26 Al/ 27 Al) 0 observed in a HAL-like CAI DH-H1 from Dhajala (Ireland et al., 1992) was attributed to a late redistribution of radiogenic magnesium. Since both CAIs contain ferrous spinel, indicative of Fe Mg interdiffusion during -15-

16 thermal metamorphism, the same interpretation can be used to explain the Al Mg isotope systematics of H030/L. Recent experiments have shown that temperature, pressure and oxygen fugacity of the surrounding gas play an important factor during the evaporation of silicate liquids (e.g., Richter et al., 2002, 2007, 2008; Davis and Richter 2003). The evaporation kinetics are significantly high at vacuum or at a finite hydrogen pressure with the rate being two orders of magnitude larger in the latter compared to the former (Richter et al., 2007, 2008). This evaporation under a finite hydrogen pressure leads to significant evaporation of magnesium and silicon from the melt because of increase in the saturation vapour pressure of Mg and SiO. Several workers have shown based on vacuum evaporation experiments of synthetic forsterite (e.g., Davis et al., 1990; Wang et al., 1993) and bulk Allende (Floss et al., 1996) that large isotopic fractionation of magnesium, silicon, and oxygen can be produced through evaporation from the liquid state, and that oxidizing conditions are intrinsic to an evaporation process. Bulk Allende (Floss et al., 1996), which lost more than 97% of its mass during heating above 2000ºC contained hibonite in the residue. These hibonites show cerium and europium anomalies, mass-dependent fractionation in calcium and titanium, enrichment in heavy oxygen isotopes, and almost complete loss of magnesium. Davis et al. (1990) also showed that a non-equilibrium evaporation in molten state can produce the isotopic fractionation seen in FUN CAIs. Wang et al. (2001) further evaporated synthetic materials of solar elemental ratios of iron, magnesium, silicon, titanium, calcium, and aluminum oxides doped with REE at 1800 and 2000ºC for different duration. They found that the residue after heating the sample to 2000 ºC for about 25 minutes (mass loss of ~95%) has Ca,Aloxides with almost complete loss of FeO, MgO, and SiO 2 (MgO and SiO 2 evaporates in fixed proportion after 50% mass loss), minor titanium loss, and low vaporization of CaO and Al 2 O 3. On the basis of evaporation experiments on Type B CAI-like material, Richter et al. (2002) suggested that significant fractionation of magnesium and silicon isotopes can be achieved by slow cooling at low ambient pressure and in temperature ranges of ºC. Based on the experimental and isotopic data discussed above, the following model for the origin H030/L can be proposed: (i) Condensation of the solid precursor of H030/L or aggregation of a dust ball composed of presolar grains. -16-

17 (ii) Heating of the precursors above 2000ºC and the molten liquid evaporated (by a nonequilibrium process) almost all magnesium and silicon; oxygen was significantly mass fractionated due to evaporative loss. From the experimental data of oxygen isotope massfractionation of artificially evaporated forsterite and bulk Allende the mass loss involved during the non-equilibrium evaporation of the precursor of H030/L is estimated to be 97 98%. (iii) The observed correlation of 26 Mg* with the 27 Al/ 24 Mg ratio suggests that live 26 Al was present during the distillation process. As the hibonite crystallized it had the imprint of the 26 Al of the liquid. If the precursors had the canonical 26 Al/ 27 Al ratio, the distillation event occurred 4.5 Myr later. (iv) Incorporation of minor magnesium due to reaction with the surrounding nebular gas. Iron must also have been incorporated during the post accretion metamorphism, which led to the formation of the Fe-rich spinel as stated above. Pyroxene hibonite spherule 1476/ Pyroxene hibonite spherules are a unique type of refractory inclusions found in several carbonaceous chondrite groups (e.g., Kurat 1975; Grossman et al., 1988; MacPherson et al., 1989; Ireland et al., 1991). They are characterized by the large excesses in 48 Ca and 50 Ti, the diverse rare earth element (REE) patterns, and the absence or deficits of 26 Mg* (e.g., MacPherson et al., 1989; Ireland et al., 1991; Russell et al., 1998; Simon et al., 1998). Ireland et al. (1991) found large 48 Ca and 50 Ti excesses, deficit of 26 Mg*, and enrichment of heavy isotopes of magnesium in the pyroxene glass in heavy isotopes of magnesium compared to the hibonites in the pyroxene hibonite spherules from Lanće (CO3) and Murchison (CM) (Table 6). Russell et al. (1998) reported similar 26 Mg* deficits and isotopically anomalous calicum and titanium in the pyroxene hibonite spherules from Colony (CO3.1) and ALH (CO3.3) (Table 6). Simon et al., (1998) described a pyroxene hibonite spherule from Yamato CO3 chondrite with a large excess of 26 Mg (Table 6) in the pyroxene and normal magnesium isotope composition in the hibonites, and suggested that the hibonite grains are relicts. Based on the analysis of the CMAS phase diagram, Beckett and Stolper (1994) concluded that the precursors of the hibonite-pyroxene spherules consisted of hibonite, spinel, melilite, and perovskite, which were heated up to ~ ºC. The vacuum evaporation experiments on the synthetic and meteorite samples have shown that no significant isotopic fractionation occurs -17-

18 at <50% weight loss during distillation (Floss et al., 1996; Wang et al., 2001). Alternatively, fast cooling at a moderately high total pressure (> 10 2 bar) of the CAI melts could have prevented significant fractionation of magnesium isotopes (Richter et al., 2002). The hibonite grains of the fassaite-rich spherule 1476/124 show neither fractionated magnesium isotopes nor 26 Mg*; magnesium isotopes in fassaite were not analyzed. The fassaite and hibonite grains have similar 16 O-rich compositions (Δ 17 O ~ 25 ). The bulk chemical composition of 1476/124 projected from spinel onto the gehlenite anorthite forsterite plane of the CMAS (CaO MgO Al 2 O 3 SiO 2 ) system plots above the spinel saturation surface, in the anorthite field, indicating that hibonite is not the liquidus phase from a melt of such composition, and, hence, must be relict. Based on these observations, we infer that the 16 O-rich precursors of 1476/124 were melted at ~ ºC under nearly equilibrium conditions and cooled rapidly which resulted in the lack of mass-dependent fractionation of oxygen and magnesium isotopes. These precursors must have suffered multiple events of heating if the hibonite grains are relict. Calcium and titanium isotope compositions need to be measured to understand the nature of the precursor material of 1476/124. If it shows excesses of 48 Ca and 50 Ti, as many other pyroxene hibonite spherules, then the precursor may have included presolar grains. The lack of 26 Mg* in hibonite grains of 1476/124 requires an explanation. The late-stage formation of 1476/124, after decay of 26 Al, seems unlikely considering its 16 O-rich isotopic composition, similar to those of most CAIs with the canonical 26 Al/ 27 Al ratio. Considering wellpreserved igneous zoning of pyroxenes in 1476/124 and one of the lowest petrologic types of its host meteorite, NWA 1476, the metamorphic resetting of its Al-Mg systematics of 1476/124 seems unlikely. One possible explanation is heterogeneous distribution of 26 Al in the early Solar System. We note, however, that recent high-precision magnesium isotope measurements of bulk CAIs (Thrane et al., 2006; Jacobsen et al., 2008), bulk chondrites, Mars, Moon, and the Earth (Thrane et al., 2006) suggest a uniform distribution of 26 Al, at least in the inner Solar System, the region where probably most meteorites originated. Alternatively, 1476/124 formed in a localized region of the solar nebula lacking canonical abundance of 26 Al, the region where most 26 Al-poor CAIs (FUN, PLACs) may have formed (e.g., Papanastassiou and Brigham, 1989; MacPherson et al., 1995; Ireland, 1988, 1990). Finally, Wood (1996) suggested that the 26 Al-poor inclusions are the residues of partial evaporation of the presolar precursors consisting of a volatile phase rich in 26 Al and refractory grains devoid of it. Such partial evaporation should produce an isotopically -18-

19 anomalous inclusion with less than the canonical abundance of 26 Al. To test this hypothesis, calcium and titanium isotope compositions need to be measured. Spinel plagioclase-rich CAI 753B/ The spinel plagioclase-rich CAI 753B/51 shows no evidence for 26 Mg* (the negative 26 Mg* values appear to be an instrumental artifact and do not reflect true deficits in 26 Mg). The CAI is texturally similar to the plagioclase-rich, igneous (Type C) inclusions (Wark, 1987; Beckett and Grossman, 1988). The Type C CAIs often show low initial 27 Al/ 26 Al ratios, similar to those in most plagioclase-rich chondrules (Wark, 1987 and references therein; MacPherson et al., 1995; Sheng et al., 1991). About 50% of Type C CAIs in CV chondrites appear to have experienced late-stage (~2 3 Myr after formation of CAIs with the canonical 26 Al/ 27 Al ratio) melting during chondrule formation (Krot et al., 2007). 753B/51 could be a Type C CAI melted during such late-stage event, after nearly complete decay of 26 Al. Alternatively, the CAI experienced redistribution of magnesium isotopes during thermal metamorphism on the R chondrite parent asteroid, which is consistent with the presence of abundant secondary oligoclase in this inclusion (Figs. 4c,d). CONCLUSIONS Oxygen and magnesium isotope compositions were measured in situ in 10 and 4 R CAIs, respectively, using a Cameca ims On a three isotope oxygen diagram, compositions of individual minerals in most R CAIs analyzed fall along slope-1 line. Based on the variations of Δ 17 O values within an individual inclusion, the R CAIs are divided into (i) 16 O-rich (Δ 17 O ~ 23 to 26 ), (ii) uniformly 16 O- poor (Δ 17 O ~ 2 ), and (iii) isotopically heterogeneous (Δ 17 O range 25 to +5 ). One of the hibonite-rich CAIs, H030/L, has an intermediate Δ 17 O value of 12 and a highly-fractionated oxygen (δ 18 O ~ +47 ). 3. We conclude that the R CAIs formed in an 16 O-rich gaseous reservoir, like the majority of CAIs in other chondrite groups. The isotopically heterogeneous CAIs subsequently experienced oxygen isotope exchange during remelting in an 16 O-depleted nebular gas, possibly during chondrule formation, and/or during fluid-assisted thermal metamorphism on the R chondrite parent asteroid. -19-

20 Three hibonite-bearing CAIs and one anorthite-bearing inclusion were analyzed for magnesium isotope compositions. Only the hibonite-rich CAI, H030/L, shows a resolvable excess of 26 Mg corresponding to the initial 26 Al/ 27 Al ratio of ~ The absence of 26 Mg* in the anorthite-bearing CAI from a metamorphosed R chondrite could be related to metamorphic resetting of its Al Mg systematics. Two other 26 Al-poor, hibonite-bearing CAIs occur in R chondrites, which appear to have experienced only minor degree of thermal metamorphism. These inclusions could have formed from the precursors with the lower 26 Al/ 27 Al ratio than the canonical value. Acknowledgements: We thank Ursula Heitmann and Thorsten Grund (Münster) for sample preparation and technical support. We would also like to thank R. N. Clayton for helpful discussion regarding the oxygen isotopes and Gerhard Wurm for sharing his ideas on photophoresis. This study is part of the dissertation of S. S. R. at the Westfälische Wilhelms- Universität Münster. S. S. R. thanks for the great support during his visit at the Hawai i Institute of Geophysics and Planetology. This work was supported by NASA grants NNX07AI81G (A. N. Krot, PI), and NAG (K. Keil, PI). 900 REFERENCES Aléon J., Krot A. N. and McKeegan K. D. (2002) Calcium-aluminum-rich inclusions and amoeboid olivine aggregates from the CR carbonaceous chondrites. Meteorit. Planet. Sci. 37, Allen J. M., Grossman L., Lee T., and Wasserburg G. J. (1980) Mineralogy and petrography of HAL, an isotopically-unusual Allende inclusion. Geochim. Cosmochim. Acta 44, Amelin Y., Krot A. N., Hutcheon I. D. and Ulyanov A. A. (2002) Lead isotopic ages of chondrules and calcium-aluminum-rich inclusions. Science 297, Armstrong J. T. (1991). Quantitative elemental analysis of individual microparticles with electron beam instruments. In Electron Probe Quantitation, (eds. K. F. J. Heinrich and D. E. Newbury) New York, Plenum Press. pp Beckett J. R. and Grossman L. (1988) The origin of type C inclusions from carbonaceous chondrites. Earth Planet. Sci. Lett. 89, Beckett J. R. and Stolper E. (1994) The stability of hibonite, melilite and other aluminous phases in silicate melts: Implications for the origin of hibonite-bearing inclusions from carbonaceous chondrites. Meteoritics 29, Bischoff A. (2000) Mineralogical characterization of primitive, type 3 lithologies in Rumuruti chondrites. Meteorit. Planet. Sci. 35,

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27 Table 1. Description of the R chondrite CAIs analyzed for oxygen and magnesium isotopes. Minerals analyzed for Petrologic Size Meteorite type ref CAI # Mineralogy Oxygen Magnesium (µm) isotopes isotopes Dar al Gani 013* R3-6 1,2 D013/ Sp + Fas + Di Sp, Di Dar al Gani 013* R3-6 1,2 D013B/ Sp + Fas + Di + Pv Sp, Fas Dhofar 1223 R3 3 Dfr/ Fas + Ol + Sp Fas, Sp Dhofar 1223 R3 3 Dfr/ Sp + Fas + Sod + Alt + Pv Fas, Sp Hughes 030* R3-6 4 H030/L 170 Hib + Sp Hib Hib NWA 753* R /6L 200 Sp + An + Fas + Di Di, Sp NWA 753* R B/ Sp + Fas Sp, Fas NWA 753* R C/ Sp + An + Olg An NWA 2446* R D/2L 150 Sp + Hib + Di Hib, Di Hib NWA 2446* R D/33B 280 Sp + Di + Alt Sp NWA 1476* R / Hib + Al-Di + Fas Fas, Al-Di, Hib Hib *breccias. Alt = Na,Al-rich secondary minerals; An = anorthite; Di = diopside; Al-Di = Al-rich diopside; Fas = fassaite; Hib = hibonite; Ol = olivine; Olg = oligoclase; Pv = perovskite; Sod = sodalite; Sp = spinel. References: 1 = Jäckel et al. (1996); 2 = Grossman (1996); 3 = Russell et al. (2005); 4 = Grossman (1998); 5 = Grossman and Zipfel (2001); 6 = Russell et al. (2003). -27-

28 Table 2. Electron microprobe analyses of minerals in CAIs from R chondrites. CAI type / # / Mineral SiO 2 TiO 2 Al 2 O 3 Cr 2 O 3 FeO MnO MgO CaO Na 2 O NiO ZnO Total Sp-rich D013/44 Sp < <0.02 <0.01 <0.03 < Di < <0.03 < D013B/118 Sp < < < <0.01 <0.03 < Fas < < D/33B Sp < Dfr/53 Sp < < <0.01 < Fas < <0.01 <0.03 n.d Sp-Plag-rich 753/6L Sp Di < C/51 Sp < < An 43.0 < < <0.03 < <0.03 < Hib-rich H030/L Hib < < <0.01 < <0.01 <0.03 < D/2L Hib < < < < Di < < < Fas-rich spherules Dfr/23 Sp 0.06 < Fas < <0.01 <0.06 < B/14 Sp Fas < <0.01 <0.06 < /124 Hib < < <0.06 < Al- Di 45.4 < < < <0.01 < Fas <0.01 < Al-Di = Al-rich diopside; An = anorthite; Di = diopside; Fas = fassaite; Hib = hibonite; Sp = spinel. -28-

29 Table 3. Oxygen isotope compositions of individual minerals in the CAIs from R chondrites. CAI # / Mineral Spot # δ 18 O, 2σ δ 17 O, 2σ Δ 17 O, 2σ 16 O-rich 1476/124 (Fas-rich spherule) Fas Al-Di Hib D/33B (Hib-rich) Sp Sp D013/44 (Sp-rich) Di Sp Di Sp Uniformly 16 O-poor Dfr/23 (Fas-rich spherule) Fas Sp Fas B/14 (Fas-rich spherule) Fas Sp Fas Isotopically heterogeneous Dfr/53 (Sp-rich) Sp Sp Fas Fas D/2L (Sp-Hib-rich) Hib Di D013B/118 (Sp-rich) Sp Fas /6L (Sp-Plag-rich) Sp Di Di Highly-fractionated H030/L (Hib-rich) Hib Hib Al-Di = Al-rich diopside; An = anorthite; Di = diopside; Fas = fassaite; Hib = hibonite; Plag = plagioclase; Sp = spinel. -29-

30 Table 4. Magnesium isotope compositions of CAIs from R chondrites. CAI # / Mineral Spot # F Mg 26 Mg/ 24 Mg δ 26 Mg 27 Al/ 24 Mg ( 26 Al/ 27 Al) D/2L (Hib-rich) No Hib ± ± ± 6 Hib ± ± ± 7 < Hib ± ± ± 6 Hib ± ± ± /124 (Fas-rich spherule) Hib ± ± ± 4 Hib ± ± ± 4 < Hib ± ± ± 4 H030/L (Hib-rich) Hib ± ± ± 4860 Hib ± ± ± 2471 (7.1±1.0) 10 7 Hib ± ± ± 389 Hib ± ± ± B/51 (Sp-Plag-rich) An ± ± ± 185 < An ± ± ± 62 An = anorthite; Hib = hibonite; Plag = plagioclase. -30-

31 Table 5. Comparison of the isotopic compositions of the H030/L and six HAL-type CAIs. F O F Ca F Ti CAI ref ( per amu) ( per amu) ( per amu) 27 Al/ 24 Mg* ( per amu) ( 26 Al/ 27 Al) 0 H030/L # ~57, ± 11.9 (7.1 ± 1.0) 10 7 HAL 1,2,6 ~ ± ± 2.2 ~42,000 ** (5.2±1.7) 10 8 DH-H1 3, ± ± 2.2 ~17,000 ~19 (8.4 ± 0.5) ,6 ~ ± ± 3.8 ~60,000 (28 ± 13) (92 ± 10) ~ ± ± 3.8 ~1, ± 10.6 (1.4 ± 0.6) 10 6 SP ± 2.0 ~2, ± 21.4 (2.4 ± 0.3) 10 5 Kz # 16.4 ± ± 1.3 ~4, ± 29.2 (5.3 ± 0.2) 10 5 # The range of oxygen isotope fractionation values are based on two different paths of oxygen isotope fractionation as described in the text. The lower and upper limit corresponds to the path ADE and ABC in the Fig. 8a respectively. For the path ADE, F O = [δ 18 O(E) - δ 18 O(D)]/2 and for path ABC, F O = (δ 18 O(E) + 50)/2, where δ 18 O(E) and δ 18 O(D) are the δ 18 O values at the point E and D in Fig 8a, respectively. A similar method was also applied for calculating the F O value of Kz1-2. Mean and 2σ standard deviation for measured spots (when <3 points were measured, the error given is propagated measurement error). *Highest 27 Al/ 24 Mg measured. **Fractionation is below the error limit. Data are from two heterogeneous areas within the same inclusion, which also shows heterogeneity in 26 Mg*. The ( 26 Al/ 27 Al) 0 ratio corresponds to the highest δ 26 Mg value measured within the core of the inclusion. References: 1 = Lee et al. (1979); 2 = Fahey et al. (1987); 3 = Hinton and Bischoff (1984); 4 = Hinton et al. (1988); 5 = Ireland and Compston (1987); 6 = Ireland et al. (1992); 7 = Russell et al. (1998); 8 = Ushikubo et al. (2007). F Mg -31-

32 Table 6. Calcium, titanium, and magnesium isotope compositions of the hibonite pyroxene spherules from different chondrites. CAI ref / Mineral δ 48 Ca, δ 50 Ti, δ 26 F Mg, Mg ( per amu) 1476/124 Hib 0.9 ± ± 3.7 LA3413-1/31 1 Hib 24.5 ± ± ± ± 1.4 Px glass 10.8 ± ± ± ± Hib 41.9 ± ± ± ± 0.9 Px glass 35.5 ± ± ± ± Hib 32.5 ± ± ± ± 1.0 Px glass 33.8 ± ± ± ± 1.2 ALH SP15 2 Hib 23.8 ± ± ± ± 2.1 Px 25.8 ± ± ± 1.7 Colony Sp1 2 Hib 1.5 ± ± ± ± 2.3 Px 4.0 ± ± ± 1.2 MYSM3 3 Hib 31.4 ± ± ± ± 2.6 Px 31.4 ± ± ± ± 3.4 Y Hib 44.9 ± ± ± ± 2.2 Px 44.9 ± ± ± ± 3.0 Mean and 2σ standard deviation for measured spots (when <3 points were measure, the error given is propagated measurement error). Both, hibonite and pyroxene have similar calcium and titanium isotope compositions; the reported value represents the mean weighed average of three spots. *Isotopic mass fractionation insignificant. **Hibonites have F Mg ~ 0. References: 1 = Ireland et al. (1991); 2 = Russell et al. (1998); 3 = Simon et al. (1998). Hib = hibonite; Px = pyroxene. -32-

33 FIGURE CAPTIONS Fig. 1. Backscattered electron (BSE) images of the CAIs D013/44 (a c) and D013B/118 (d f). Regions outlined in a and d are shown in detail in b c and e f respectively. The CAIs are composed of spinel and fassaite and rimmed by diopside. Here and in Figures 2 7, the ion microprobe spots for oxygen isotopes are outlined and numbered; the numbers correspond to the analysis numbers listed in Tables 3 and 4. Di = diopside; Fas = fassaite; Sp = spinel; Pv = perovskite. Fig. 2. BSE images of a spinel-rich CAI fragment 2446D/33B surrounded by a discontinuous rim of diopside. Region outlined in a is shown in detail in b. Alt = alteration products; Di = diopside; Sp = spinel; Fas = fassaite. Fig. 3. BSE images (a c) and elemental maps in Al (d), Ca (e) and Ti Kα (f) x-rays of a CAI Dfr/53. Regions outlined in a are shown in detail in b c. The CAI core consists of spinel, fassaite, perovskite, and abundant alteration minerals, including sodalite; the core is rimmed by diopside. Alt = alteration products; Di = diopside; Fas = fassaite; Pv = perovskite; Sod = sodalite; Sp = spinel. Fig. 4. BSE images of the spinel plagioclase-rich CAI 753/6L (a, b) and 753C/51 (c, d). Regions outlined in a and c are shown in detail in b and d, respectively. An = anorthite; Di = diopside; Olg = oligoclase; Fas = fassaite; Sp = spinel. Fig. 5. BSE images of the CAIs H030/L (a b) and 2446D/2L (c f). Regions outlined in a and c are shown in detail in b and d f, respectively. Alt = alteration products; Di = diopside; Hib = hibonite; Sp = spinel; Ilm = ilmenite. Fig. 6. BSE images (a, f) and elemental maps in Al (b), Ca (c), Fe (d), and Mg Kα (e) x-rays of the fassaite-rich spherules Dfr/23 (a e) and 753B/14 (f). Di = diopside; Fas = fassaite; Ol = olivine; Pv = perovskite; Sp = spinel. Fig. 7. BSE image (a) and elemental maps in Al (b), Ca (c), Mg (d), Si (e), and Ti Kα (f) x-rays of a CAI 1476/124 (a) consist of euhedral laths of hibonites within a fassaite groundmass. The ion microprobe spots for oxygen (rectangular box) and magnesium (circles) isotope measurements are outlined and numbered; the numbers correspond to the analysis numbers listed in Tables 3 and 4, respectively. Al-Di = Al-diopside; Fas = fassaite; Hib = hibonite. Fig. 8. Oxygen isotope compositions of CAIs from the R chondrites. In a, the data are plotted on a three-isotope oxygen diagram, δ 17 O vs. δ 18 O. In b, the same data are plotted as a deviation from the terrestrial fractionation line, Δ 17 O. In a the paths of oxygen isotopic evolution of the inclusion H030/L is also illustrated. Error bars are 2σ. Fig. 9. Oxygen isotope compositions of the R chondrite CAIs (this study), chondrules, refractory forsterite grains, and magnetite, and bulk oxygen isotope compositions of R chondrites (data from Weisberg et al., 1991; Bischoff et al., 1994; Schulze et al., 1994; Kallemeyn et al., 1996; Greenwood et al., 2000; Pack et al., 2004). In a the data are plotted in a threeisotope oxygen diagram, δ 17 O vs. δ 18 O. In b, the same data are plotted as a deviation from the terrestrial fractionation line, Δ 17 O. Error bars are 2σ. -33-

34 Fig. 10. The aluminum magnesium evolutionary diagrams for the CAIs from R chondrites. Error bars are 2σ. -34-

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