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

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Geochimica et Cosmochimica Acta 73 (2009) 5018–5050

Oxygen- and magnesium-isotope compositions of calcium–aluminum-rich inclusions from CR2 carbonaceous chondrites

Kentaro Makide a, Kazuhide Nagashima a, Alexander N. Krot a,*, Gary R. Huss a,Ian D. Hutcheon b, Addi Bischoff c

a School of Ocean and Earth Science and Technology, Hawai‘i Institute of Geophysics and Planetology, University of Hawai‘i at Manoa,

1680 East–West Road, Honolulu, HI 96822, USAb Glenn T. Seaborg Institute, Lawrence Livermore National Laboratory, Livermore, CA 94551, USA

c Institut fur Planetologie, Westfalische Wilhelms-Universitat Muenster, Wilhelm-Klemm-Str. 10, 48149 Muenster, Germany

Received 8 September 2008; accepted in revised form 5 January 2009; available online 29 May 2009

Abstract

We report both oxygen- and magnesium-isotope compositions measured in situ using a Cameca ims-1280 ion microprobein 20 of 166 CAIs identified in 47 polished sections of 15 CR2 (Renazzo-type) carbonaceous chondrites. Two additional CAIswere measured for oxygen isotopes only. Most CR2 CAIs are mineralogically pristine; only few contain secondary phyllosil-icates, sodalite, and carbonates – most likely products of aqueous alteration on the CR2 chondrite parent asteroid. Spinel,hibonite, grossite, anorthite, and melilite in 18 CAIs have 16O-rich (D17O = �23.3 ± 1.9&, 2r error) compositions and showno evidence for postcrystallization isotopic exchange commonly observed in CAIs from metamorphosed CV carbonaceouschondrites. The inferred initial 26Al/27Al ratios, (26Al/27Al)0, in 15 of 16 16O-rich CAIs measured are consistent with thecanonical value of (4.5–5) � 10�5 and a short duration (<0.5 My) of CAI formation. These data do not support the‘‘supra-canonical” values of (26Al/27Al)0 [(5.85–7) � 10�5] inferred from whole-rock and mineral isochrons of the CV CAIs.A hibonite–grossite-rich CAI El Djouf 001 MK #5 has uniformly 16O-rich (D17O = �23.0 ± 1.7&) composition, but shows adeficit of 26Mg and no evidence for 26Al. Because this inclusion is 16O-rich, like CAIs with the canonical (26Al/27Al)0, we inferthat it probably formed early, like typical CAIs, but from precursors with slightly nonsolar magnesium and lower-than-canon-ical 26Al abundance. Another 16O-enriched (D17O = �20.3 ± 1.2&) inclusion, a spinel–melilite CAI fragment Gao-Guenie (b)

#3, has highly-fractionated oxygen- and magnesium-isotope compositions (�11 and 23&/amu, respectively), a deficit of26Mg, and a relatively low (26Al/27Al)0 = (2.0 ± 1.7) � 10�5. This could be the first FUN (Fractionation and UnidentifiedNuclear effects) CAI found in CR2 chondrites. Because this inclusion is slightly 16O-depleted compared to most CR2 CAIsand has lower than the canonical (26Al/27Al)0, it may have experienced multistage formation from precursors with nonsolarmagnesium-isotope composition and recorded evolution of oxygen-isotope composition in the early solar nebula over0:9�0:7þ2:2 My. Eight of the 166 CR2 CAIs identified are associated with chondrule materials, indicating that they experienced

late-stage, incomplete melting during chondrule formation. Three of these CAIs show large variations in oxygen-isotope com-positions (D17O ranges from �23.5& to �1.7&), suggesting dilution by 16O-depleted chondrule material and possiblyexchange with an 16O-poor (D17O > �5&) nebular gas. The low inferred (26Al/27Al)0 ratios of these CAIs (<0.7 � 10�5) indi-cate melting >2 My after crystallization of CAIs with the canonical (26Al/27Al)0 and suggest evolution of the oxygen-isotopecomposition of the inner solar nebula on a similar or a shorter timescale. Because CAIs in CR2 and CV chondrites appear tohave originated in a similarly 16O-rich reservoir and only a small number of CR2 and CV CAIs were affected by chondrulemelting events in an 16O-poor gaseous reservoir, the commonly observed oxygen-isotope heterogeneity in CAIs from meta-morphosed CV chondrites is most likely due to fluid–solid isotope exchange on the CV asteroidal body rather than gas–meltexchange. This conclusion does not preclude that some CV CAIs experienced oxygen-isotope exchange during remelting,

0016-7037/$ - see front matter � 2009 Elsevier Ltd. All rights reserved.

doi:10.1016/j.gca.2009.01.042

* Corresponding author.E-mail address: [email protected] (A.N. Krot).

mailto:[email protected]

CAIs in CR chondrites 5019

instead it implies that such remelting is unlikely to be the dominant process responsible for oxygen-isotope heterogeneity in CVCAIs. The mineralogy, oxygen and magnesium-isotope compositions of CAIs in CR2 chondrites are different from those inthe metal-rich, CH and CB carbonaceous chondrites, providing no justification for grouping CR, CH and CB chondrites intothe CR clan.� 2009 Elsevier Ltd. All rights reserved.

1. INTRODUCTION

Calcium, aluminum-rich inclusions (CAIs) are the oldestsolids formed in the solar nebula with an age of �4.567 Gyr(Chen and Wasserburg, 1981; Amelin et al., 2002). Theduration of formation and subsequent thermal processing(melting and evaporation) of CAIs in the solar nebula,which can be inferred from 26Al–26Mg systematics of CAIsfrom primitive (unmetamorphosed) chondrites, providesimportant constraints on the mechanism of CAI formationand chronology of the early Solar System (e.g., MacPher-son et al., 1995; Thrane et al., 2006; Jacobsen et al.,2008). 26Al decays to 26Mg with t1/2 � 0.73 My; the initial26Al/27Al ratios, (26Al/27Al)0, inferred from in situ magne-sium isotopic measurements of minerals with high27Al/24Mg ratios (anorthite, gehlenitic melilite, hibonite,grossite) in most CAIs using secondary ion mass spectrom-etry (SIMS, ion microprobe) show a reasonably well-de-fined value of (4.5–5) � 10�5, called the ‘‘canonical” ratio(MacPherson et al., 1995).

This canonical value has been recently challenged byhigh-precision in situ measurements of magnesium isotopesin CAI minerals with low 27Al/24Mg ratios (spinel andfassaitic pyroxene) using multicollector, SIMS (MC-SIMS)and laser-ablation, multicollector, inductively-coupled plas-ma mass spectrometry (LA-MC-ICPMS), as well as by‘‘whole-rock” magnesium-isotope measurements of CAIsusing MC-ICPMS, from which much higher than thecanonical 26Al/27Al ratio (up to 7 � 10�5) was inferred (Biz-zarro et al., 2004, 2005; Young et al., 2005; Taylor et al.,2005; Thrane et al., 2006; Cosarinsky et al., 2007). In spiteof the disagreement over the actual value of the initial26Al/27Al ratio (c.f. Thrane et al., 2006 and Jacobsenet al., 2008), both studies reported very small ranges in ini-tial 26Al/27Al ratios about the mean which they interpretedas an evidence for a very brief duration of CAI formation(<0.02–0.1 My). In contrast, Young et al. (2005) and Cosa-rinsky et al. (2007) concluded that CAI formation lasted for>0.3 My.

Note that most of the recent high-precision isotope mea-surements have been made on coarse-grained igneous (com-pact Type A and Type B) CAIs from the CV chondrites,which experienced multistage thermal processing in the so-lar nebula and on the CV asteroidal body (e.g., Krot et al.,1998 and references therein) that could have resulted in dis-turbance or resetting of their 26Al–26Mg systematics. Theseprocesses could also have affected oxygen-isotope composi-tions of the CV CAIs, which are typically characterized bylarge variations: melilite and anorthite are systematically16O-depleted relative to spinel, hibonite, and fassaite (e.g.,Yurimoto et al., 2008 and references therein). Althoughthere is a general agreement that this heterogeneity is due

to isotopic exchange between initially 16O-rich CAIs and16O-poor external reservoir, the nature of the exchange(gas–melt vs. gas–solid) and its time and place remainhighly controversial (e.g., Yurimoto et al., 1994, 1998;Ryerson and McKeegan, 1994; Itoh and Yurimoto, 2003;Dyl et al., 2008; Krot et al., 2008a). It may have resultedfrom very early (nearly contemporaneously with formationof CAIs) gas–solid or gas–melt interactions in the solar neb-ula regions that experienced rapid fluctuations of oxygen-isotope compositions (Itoh and Yurimoto, 2003; Aleonet al., 2007; Dyl et al., 2008). Alternatively, it may have re-sulted from late-stage, fluid-assisted thermal metamor-phism on the CV asteroidal body (Krot et al., 2008a).

To constrain the duration of formation and subsequentthermal processing of CAIs and the evolution of oxygenisotopic composition of the solar system, we measured theoxygen- and magnesium-isotope compositions of a suiteof CAIs from the CR (Renazzo-type) carbonaceous chon-drites. The CR2 meteorites were chosen because they areamong the least metamorphosed chondrites; they did, how-ever, experience low-temperature aqueous alteration to dif-ferent degrees (Bischoff et al., 1993a; Weisberg et al., 1993,1995; Krot et al., 2002 and references therein). Most CR2CAIs are mineralogically pristine objects that escaped sec-ondary alteration processes, and thus probably preservedsolar nebula records virtually unchanged (Aleon et al.,2002; Weber et al., 1995; Weber and Bischoff, 1997; Krotet al., 2005a; Weisberg et al., 2007).

2. ANALYTICAL PROCEDURES

2.1. Identification and mineralogical characterization of

CAIs

The mineralogy and petrography of refractory inclu-sions [calcium–aluminum-rich inclusions (CAIs) and amoe-boid olivine aggregates (AOAs) with large CAIs] from 47polished thin sections of 15 (some are paired) CR2 carbona-ceous chondrites [Asuka 881828, Elephant Moraine (EET)92042, 96286, Graves Nunataks (GRA) 95229, Temple Bar,Acfer 097, Acfer 209, El Djouf 001, and Gao-Guenie (b)]were studied using a Cameca SX-50 electron microprobeand a JEOL JSM-5900LV scanning electron microscope(SEM) equipped with a Thermo Electron energy dispersiveX-ray microanalysis system. Refractory inclusions from theCR2 chondrites, Renazzo and Al Rais, were excluded fromour study, because many of them experienced extensiveaqueous alteration resulting in replacement of anorthiteand melilite by phyllosilicates.

The refractory inclusions were identified using X-rayelemental mapping in Mg, Ca, and Al Ka with a resolutionof 2–5 lm/pixel; the elemental maps were acquired using

5020 K. Makide et al. / Geochimica et Cosmochimica Acta 73 (2009) 5018–5050

five spectrometers of the Cameca SX-50 microprobe operat-ing at 15 kV accelerating voltage, 50–100 nA beam currentand �1–2 lm beam size. The Mg, Ca, and Al X-ray imageswere combined by using a RGB-color scheme and ENVIsoftware to obtain false color maps to identify all CAIsand AOAs larger than 5–10 lm in apparent diameter. Theidentified refractory inclusions were studied in secondaryelectron and backscattered electron (BSE) images beforeand after ion probe measurements to verify the mineralogyof the sputtered regions. Some of the sections had been pre-viously gold-coated and were recoated with carbon after re-moval of gold coating. We made no attempts to removetraces of gold in holes and cracks of CAIs; this resultedin appearances of bright white spots in BSE images of theseCAIs.

Electron probe microanalyses were performed with theCameca SX-50 electron microprobe using a 15 keV acceler-ating voltage, 10–20 nA beam current, beam size of �1–2 lm and wavelength dispersive X-ray spectroscopy. Foreach element, counting times on both peak and backgroundwere 30 s (10 s for Na and K). Matrix effects were correctedusing PAP procedures. The element detection limits withthe Cameca SX-50 were (in wt.%): 0.04 (K2O), 0.07(Cr2O3), 0.08 (Na2O), and 0.09 (FeO). The representativechemical compositions of minerals in the CR2 CAIs arelisted in Tables EA1–EA3.

The modal abundances of CAIs in CR2 chondrites are<1 vol.% (Weber and Bischoff, 1997; Aleon et al., 2002; He-zel et al., 2008). In general, the CAIs are small (10–300 lmin apparent diameter), fine-grained, and irregularly-shaped;most of them are fragments of larger inclusions. Twenty-two of 166 identified CAIs were selected for detailed isoto-pic measurements. These CAIs contain regions large en-ough (15–20 lm in size) for high-precision oxygen- andmagnesium-isotope measurements using ion microprobe,i.e., without cracks and terrestrial weathering products.The careful selection of samples is critical for the high-pre-cision oxygen- and magnesium-isotope measurements.

2.2. Oxygen-isotope measurements

Oxygen-isotope compositions were measured in situ withthe University of Hawai‘i Cameca ims-1280 ion microprobeusing primary Cs+ ion beam accelerated to 10 keV and im-pacted the sample with an energy of 20 keV. Negativelycharged secondary ions were accelerated to �10 keV andan energy window of 50 eV was used. A normal incidentelectron flood gun was used for charge compensation withhomogeneous electron density over region of �75 lm indiameter. Two analytical procedures were used to measureoxygen-isotope compositions.

In the first procedure, a �1.8 nA focused Cs+ primaryion beam was focused to a diameter of �7–10 lm and ras-tered over a 25 � 25 lm2 area for 250 s of presputtering toremove the carbon coating. Then the raster size was re-duced to 10 � 10 lm2 for automated centering of secondarybeam to the mass spectrometer, followed by data collection.Secondary ions of 16O�, 17O�, and 18O� were measuredsimultaneously in multicollection mode with the magneticfield controlled by a nuclear magnetic resonance (NMR)

probe. 16O� and 18O� were measured by multicollectorFaraday cups (FCs) with low mass resolving power(MRP �2000), while 17O� was measured using the axialmonocollector electron multiplier (EM) in pulse countingmode with MRP �5600, sufficient to separate the interfer-ing 16OH� signal. Typical count rates of 16O� and 17O�

ion beams were �4 � 108 and �1.5 � 105 cps, respectively.The relatively high count rate of the 17O� signal may pro-duce a loss in gain due to aging of the first dynode of theEM. To minimize this effect, the 17O� signal was measuredfor only 4 s in each cycle, after which the beam was de-flected into a monocollector FC for 10 s (procedure origi-nally proposed by Kita et al., 2007). Signals for 16O� and17O� collected during the 4-s interval were used to deter-mine the 17O/16O ratio and the 16O� and 18O� signals col-lected during the 10-s interval were used to determine the18O/16O ratio. Each measurement consisted of 30 cycles.As a further control on the gain of the EM, the pulse-heightdistribution was checked every five measurements and theEM high voltage was adjusted when necessary to maintaina constant pulse-height distribution.

In the second procedure, to reduce the beam size, a�200 pA Cs+ primary ion beam was focused to a diameterof �5 lm and rastered over �7 � 7 lm2 area for presput-tering (150 s). After presputtering, the raster size was re-duced to 5 � 5 lm2 for automated centering of thesecondary-ion beam followed by data collection. Three oxy-gen isotopes were measured by combination of multicollec-tion mode and peak-jumping. 16O� and 17O� weremeasured simultaneously using the multicollection FCand the monocollection EM, respectively. Subsequently,18O was measured with the monocollection EM by peak-jumping. Mass resolving power for 16O� and for 17O�

and 18O� were set to �2000 and �5600, respectively. The18O� secondary ion intensity was typically �1.5 � 105

counts per second. 18O� was collected for only 3 s, while16O� and 17O� were measured for 8 s. The measurementconsisted of 30 cycles, and again the EM gain was checkedand adjusted every five measurements.

Oxygen-isotope compositions are reported as d17O andd18O, & deviations from Standard Mean Ocean Water(SMOW):

d17;18OSMOW¼ ½ð17;18O=16OsampleÞ=ð17;18O=16OSMOWÞ�1��1000

ð1Þ

We also report 16O enrichment relative to the terrestrialfractionation (TF) line as:

D17O ¼ d17O� 0:52� d18O ð2Þ

The instrumental mass fractionation (IMF) was correctedfor each session using San Carlos olivine (Fo89), Eagle Sta-tion pallasite olivine (Fo79), Miyakejima anorthite, syn-thetic enstatite, augite, and Burma spinel standards.Melilite and grossite (for which standards were lacking)compositions were corrected by assuming their IMF aresimilar to those of San Carlos olivine. Kita et al. (2007) re-ported similar IMF factors for olivine and gehlenitic meli-lite. Because of the difficulty of precisely determiningrelative efficiencies among FC and EM detectors, measuredD17O values on standards showed systematic shifts (typi-


cally <±1&) from the terrestrial value (D17O = 0). Themean of D17O on standard measurements was used to cor-rect systematic shift from the terrestrial value. The reporteduncertainties include both the internal measurement preci-sion on an individual analysis and the external reproducibil-ity for standard measurements during a given analyticalsession. For the first procedure, the point-to-point repro-ducibility (external reproducibility) on the multiple analysesof the standards was 0.5–1& (2 times the standard devia-tion of the mean, 2r) for both d17O and d18O. For the sec-ond procedure, the external reproducibility on the multipleanalyses of the standards was 1–1.5& (2r) for both d17Oand d18O.

To avoid possible contamination on the measured spotsby primary 16O� ions used for magnesium-isotope measure-ments, oxygen isotopes were typically measured beforemagnesium isotopes. In some cases, however, CAIs selectedfor oxygen-isotope measurements have been previouslymeasured for magnesium isotopes. These sections wererepolished and recoated, and spots for oxygen-isotope anal-yses were selected to be as far as possible from the regionssputtered during magnesium-isotope measurements.

2.3. Magnesium-isotope measurements

Magnesium- and aluminum-isotope compositions weremeasured in situ with the UH Cameca ims-1280 ion micro-probe using primary 16O� ion beam accelerated to 13 keVwith an impact energy of 23 keV. Positively charged sec-ondary ions were accelerated to +10 keV and an energywindow of 50 eV was used. Two analytical procedures wereused to measure 26Al–26Mg systematics.

Minerals with high 27Al/24Mg ratios (anorthite, grossite,hibonite, and gehlenitic melilite) were analyzed with a fo-cused 5–7 lm 16O� primary beam in monocollection modeusing EM and FC detectors for magnesium isotopes and27Al, respectively. Depending on concentration of magne-sium, a primary current of 150 or 300 pA was used.24Mg+, 25Mg+, 26Mg+, and 27Al+ were measured in peak-jumping mode for 120 cycles using count times of 4, 10,10, and 2 s, respectively. The MRP was set to �3800, suffi-cient to separate interfering hydrides and doubly charged48Ca++. Automated centering of the secondary beam inthe field aperture of the mass spectrometer, high-voltageoffset control, and mass-peak centering were applied beforeeach measurement. Offset control and peak centering werealso applied at cycle 60.

In most cases, minerals with low 27Al/24Mg ratios (spineland low-Ca pyroxene) were measured with a �20 lm diam-eter 16O� primary beam in multicollector mode using fourFC detectors. The primary current was set to �5 nA, andthe 25Mg+ count rate was typically �1.2 � 107 cps and>2 � 107 cps for spinel and low-Ca pyroxene, respectively.The magnetic field was controlled by the NMR probe.The MRP was set to �4200. Each measurement consistedof 40 cycles with a 10 s integration time per cycle. Auto-mated centering of secondary beam in the field apertureand high voltage offset control were applied before eachmeasurement. In a few cases where the spinel grains weretoo small, monocollection mode was used.

Magnesium-isotope compositions are reported as devia-tions in parts per thousand (&) from the terrestrial magne-sium isotopic ratios [25Mg/24Mg = 0.12663, 26Mg/24Mg =0.13932 (Catanzaro et al., 1966)]:

D25Mg ¼ ½ðð25Mg=24MgÞmeas=0:12663Þ � 1� � 1000

D26Mg ¼ ½ðð26Mg=24MgÞmeas=0:13932Þ � 1� � 1000

Instrumental mass fractionation (IMF) was corrected bystandard-sample bracketing by comparing each measure-ment with the isotope ratios measured in the appropriateterrestrial standards, all of which were assumed to havethe Catanzaro et al. (1966) isotope ratios. Standards in-cluded Burma spinel, Madagascar hibonite, synthetic meli-lite glass, Miyakejima anorthite, and enstatite. Grossite, forwhich we have no standard, was corrected by assuming itsIMF is similar to that of hibonite. However, this assump-tion may not be appropriate (see Section 3). The CAI sam-ples themselves have an intrinsic isotopic massfractionation that probably resulted from evaporation dur-ing melting and crystallization. The intrinsic mass fraction-ation is thus calculated as the difference between the D25Mgmeasured for each sample and the mean D25Mg measuredfor the appropriate standard minerals. By our convention,positive mass fractionation indicates enrichment in the hea-vier isotopes.

In order to determine the contribution of radiogenic26Mg to the measured 26Mg/24Mg ratio, we propagatedthe intrinsic mass fractionation from 25Mg/24Mg to26Mg/24Mg using an exponential law with an experimen-tally defined mass fractionation exponent of 0.514 (Daviset al., 2005a). First the IMF-corrected ratios were con-verted to ‘‘phi” coordinates:

/25Mg ¼ 1000� lnðð25Mg=24MgÞIMFcorr=0:12663Þ/26Mg ¼ 1000� lnðð26Mg=24MgÞIMFcorr=0:13932Þ

The fractionation-corrected d26Mg0 was calculated asfollows:

d26Mg0 ¼ /26Mg� ð1=0:514Þ � /25Mg

Then d26Mg0 was converted to the fractionation-corrected26Mg/24Mg ratio:

ð26Mg=24MgÞFRACcorr ¼ 0:13932� expðd26Mg0=1000Þ

The fractionation-corrected 26Mg/24Mg ratio was then con-verted into standard delta notation:

d26Mg ¼ ½ðð26Mg=24MgÞFRACcorr=0:13932Þ � 1� � 1000

For situations where the Al/Mg ratio is low and the isotopeeffects are small, d26Mg0 and d26Mg can be almost indistin-guishable. But for minerals with high Al/Mg ratios, the dif-ference between the two can be a percent or more, and ind26Mg0 coordinates, the isochron is a curve. The choice ofmass fractionation law is not critical to this data set becausethe degree of intrinsic mass fractionation is relatively smallfor most samples and because the 27Al/26Mg ratios of theminerals that control the Al–Mg isochrons are sufficientlyhigh that the uncertainty introduced by the choice of lawis insignificant. The reported uncertainties include boththe internal precision of an individual analysis and the


external reproducibility for standard measurements duringa given analytical session.

The relative sensitivity factors for aluminum and magne-sium were determined from the 27Al+/24Mg+ measured bySIMS and the Al/Mg ratios measured previously by elec-tron microprobe for each standard mineral. The sensitivityfactor for grossite was assumed to be the same as for hibo-nite (Weber et al., 1995), but once again, this may not beappropriate (see Section 3).

3. RESULTS

3.1. Mineralogy and petrography

The mineralogy and chemical compositions of refractoryinclusions (CAIs and AOAs) in CR2 chondrites have beendescribed in detail by Weisberg and Prinz (1990), Weberand Bischoff (1997), Aleon et al. (2002), Krot et al.(2004), and Weisberg et al. (2007). Based on their mineral-ogy and petrography, the CR2 refractory inclusions can bedivided into melilite-rich, grossite-rich, hibonite-rich, andanorthite-rich CAIs, and AOAs (Table 1). Some of theCAIs are associated with chondrule-like materials com-posed of ferromagnesian olivine and pyroxene, Fe,Ni-me-tal, and glassy or fine-grained, crystalline mesostasis.Below we briefly describe mineralogy and petrography ofthe refractory inclusions that we measured for oxygen-and magnesium-isotope compositions. Regions sputteredfor oxygen-isotope analyses, and, occasionally, for magne-sium-isotope analyses, are indicated in BSE images andX-ray elemental maps of the CAIs (Figs. 1–10). Represen-tative electron microprobe analyses of minerals from theCR CAIs measured for oxygen- and magnesium-isotopecompositions can be found in the electronic annex Tables1–3.

3.1.1. Grossite- and hibonite-bearing/rich CAIs

El Djouf 001 MK #5 (Figs. 1a and b and EA1) is a hib-onite–grossite-rich CAI composed of three fragments sepa-rated by matrix material. Two of the fragments consist oflath-shaped hibonite surrounded by a grossite mantle withnumerous inclusions of perovskite. Another fragment con-sists of grossite with abundant inclusions of perovskite. Allfragments are rimmed by melilite and diopside.

Gao-Guenie (b) #10 (Figs. 1c and d) is a grossite-richCAI consisting of two fragments separated by a crackand matrix material. Grossite contains rare inclusions of

Table 1Abundances of different mineralogical types of CAIs in CR2carbonaceous chondrites.

Type Number Abundance(%)

Number of CAI–CHDcompound objects

Mel–Sp-rich 118 71 1Grs- and Hib-bearing/rich

19 12 1

An-rich 22 13 6Other 7 4 0Total 166 100 8

perovskite, hibonite and Ca-aluminate (CaAl2O4). Bothfragments are rimmed by melilite and Al-diopside.

GRA 95229-17 #7 (Fig. 2a) is a melilite-rich CAI frag-ment surrounded by a grossite + perovskite + spinel + hib-onite mantle and a Wark–Lovering rim sequence composedof spinel + hibonite, melilite, and Al-diopside. Secondaryalteration minerals occur at the boundary between the meli-lite core and the grossite-rich mantle.

GRA 95229-17 #8 (Fig. 2b) is a melilite–spinel–perov-skite CAI fragment surrounded by a grossite + perov-skite + spinel + hibonite mantle and a Wark–Loveringrim sequence of melilite and diopside. Secondary alterationminerals occur at the boundary between the melilite coreand grossite-rich mantle. The similar mineralogy of GRA

95229-17 #7 and GRA 95229-17 #8 and the presence inthe same thin section may indicate that these CAIs repre-sent fragments of the same inclusion. However, in contrastto GRA 95229-17 #7, melilite in GRA 95229-17 #8 enclosesabundant inclusions of euhedral spinel grains.

GRA 95229-31 #3 (Fig. 2c) is composed largely of gros-site; melilite and perovskite are accessory. The CAI is sur-rounded by a Wark–Lovering rim sequence of spinel andmelilite.

3.1.2. Melilite–rich CAIs

Acfer 097 PL92521 #1 (Fig. 3a) is a CAI fragment com-posed of melilite, abundant anhedral grains of spinel, andtiny inclusions of perovskite. The CAI is rimmed by amonomineralic layer of Al-diopside with aluminum contentincreasing towards the CAI core. Small spinel grains areconcentrated near the diopside rim.

Acfer 097 PL92521 #2 (Fig. 3b) is a melilite-rich CAIfragment with rare inclusions of euhedral spinel grains.Melilite is extensively replaced by a fine-grained mixtureof secondary minerals.

Acfer 209 PL91165 #2 (Fig. 3c) is a CAI fragment com-posed of melilite with euhedral inclusions of spinel grainsand minor perovskite. The unbroken surface of the CAIis rimmed by spinel with abundant inclusions of perovskite,and melilite.

Asuka 881828-61-4 #1 (Figs. 4a and b) is a melilite-richCAI fragment with abundant inclusions of tiny spinelgrains surrounded by Al,Ti-diopside, and minor anorthite.The anorthite replaces melilite. Several regions, 20–30 lmin size, composed largely of spinel occur near the edge ofthe fragment.

EET 92042-22 #4 (Fig. 4c) is a melilite-rich CAI fragmentcontaining anhedral grains of Al,Ti-diopside intergrownwith perovskite. The CAI is surrounded by a rim of diopside;small spinel grains are concentrated near the rim. The CAI islocated near a fusion crust and appears to have experiencedsome melting during the atmospheric entry.

EET 96286-6 #7 (Fig. 4d) is a fragmented melilite-richCAI with abundant anhedral inclusions of spinel; perov-skite is minor. Spinel is heterogeneously distributed in theCAI and is mainly concentrated near the CAI edge rimmedby diopside. Melilite near the rim is corroded by anorthite.

El Djouf 001 PL91172 #1 (Fig. 5a) a melilite-rich CAIfragment containing abundant inclusions of perovskiteand Al,Ti-diopside.

Fig. 1. Backscattered electron (BSE) images of the grossite ± hibonite-rich CAIs El Djouf 001 MK #5 (a and b) and Gao-Guenie (b) #10 (cand d). El Djouf 001 MK #5 consists of three fragments. Regions outlined in ‘‘a” are shown in detail in ‘‘b” and Fig. EA1. Gao-Guenie (b) #10

consists of two fragments separated by matrix material. Bright white veins are FeO-rich terrestrial weathering products. Bright grains in (aand b) represent remnants of gold coating used during previous magnesium-isotope measurements. Here and in Figs. 2–10 ion microprobespots produced during oxygen-isotope measurements are outlined by squares; black numbers in white circles correspond to analysis numberslisted in Table 2. Ion probe spots for magnesium-isotope compositions are indicated by ellipses and crosses; white numbers in black circlesnear crosses correspond to analysis numbers listed in Table 3. Al-Di, Al-diopside; CA, calcium aluminate (CaAl2O4); Di, diopside; FeS, ironsulfide; Grs, grossite; Hib, hibonite; Mel, melilite; Pv, perovskite; Sec, secondary alteration minerals; Sp, spinel.


Gao-Guenie (b) #1 (Fig. 5b) is a melilite + spinel CAIfragment rimmed by Al-diopside. Subhedral spinel grainsare evenly dispersed throughout the inclusion; several re-gions, 20–30 lm in size, composed largely of spinel are con-centrated near the Al-diopside rim.

Gao-Guenie (b) #2 (Fig. 5c) is a melilite + spinel CAIrimmed by Al-diopside. Spinel is concentrated in irregularregions near the rim. Melilite in the outer portion of theinclusion is extensively replaced by anorthite.

Gao-Guenie (b) #3 (Fig. 6a) is a CAI fragment com-posed of melilite with coarse euhedral inclusions of spinel.

Gao-Guenie (b) #6 (Figs. 6b and c) is an amoeboid olivineaggregate composed of forsteritic olivine, Al,Ti-diopside,Fe,Ni-metal, and several spinel + perovskite ± meliliteCAIs rimmed by Al,Ti-diopside. Perovskite occurs as tinyinclusions in spinel and as larger grains surrounded byAl,Ti-diopside in melilite. Melilite is partly replaced byanorthite.

GRA 95229-18 #2 (Fig. 6d) is a melilite-rich CAI frag-ment containing an anhedral grain of spinel and a grainof sodalite; the latter replaces melilite.

Temple Bar MK #4 (Fig. 6e) is a fragment of a melilite-rich CAI with rare inclusions of spinel and perovskite.

3.1.3. Anorthite-rich CAI

GRA 95229-18 #22 (Fig. 7) is a CAI fragment composedof massive anorthite and anhedral grains of Al,Ti-diopside,both poikilitically enclosing euhedral spinel grains.

3.1.4. CAI–chondrule compound objects

Acfer 209 PL91165 #1 (Fig. 8) is a compound objectcomposed of an anorthite-diopside-spinel CAI surroundedby a thick mantle of chondrule-like material. The chondruleportion consists of ferromagnesian olivine, low-Ca pyrox-ene, high-Ca pyroxene, Fe,Ni-metal, troilite, and anorthiticplagioclase. There are textural differences between plagio-clase in the core and in the mantle of the object, clearly vis-ible in Al Ka X-ray elemental map (Fig. 8b): the coreplagioclase is granular (5–10 lm in size), whereas plagio-clase in the mantle is lath-shaped.

Acfer 209 PL91165 #3 (Fig. 9) is a compound objectcomposed of a relict, anorthite–spinel–Al,Ti-diopside CAIsurrounded by a mantle of chondrule-like material. The rel-ict CAI consists of lath-shaped anorthite, Al-Ti-diopsidewith minor perovskite grains, and abundant euhedralgrains of spinel. Spinel is poikilitically enclosed by anorthiteand Al,Ti-diopside. The spinel grain sizes decrease towards

Fig. 2. BSE images of the grossite- and/or hibonite-bearing CAIsGRA 95229-17 #7 (a), GRA 95229-17 #8 (b) and GRA 95229-31 #3

(c). Al-Di, Al-diopside; An, anorthite; Di, diopside; Grs, grossite;Hib, hibonite; Mel, melilite; Pv, perovskite; Sec, secondaryalteration minerals; Sp, spinel. Bright white veins are FeO-richproducts of terrestrial weathering.

Fig. 3. BSE images of the melilite–spinel CAIs Acfer 097 PL92521

#1 (a), Acfer 097 PL92521 #2 (b), and Acfer 209 PL91165 #2 (c).Al,Ti-Di, Al,Ti-diopside; Di, diopside; Mel, melilite; Pv, perov-skite; Sec, secondary alteration minerals; Sp, spinel.


the chondrule mantle. The latter is composed of ferromag-nesian olivine, low-Ca pyroxene, high-Ca pyroxene, Fe,Ni-metal, and FeS; some of the metal grains are replaced bythe terrestrial weathering products.

Gao-Guenie (b) #4 (Fig. 10) is a compound object com-posed of a spinel-anorthite core surrounded by a mantle oflow-Ca pyroxene poikilitically enclosing small olivinegrains, and minor Al,Ti-diopside. The spinel grains appear

to be corroded by the surrounding plagioclase; the spinelgrain sizes decrease towards the chondrule mantle.

3.2. Oxygen-isotope compositions

Oxygen-isotope compositions of individual minerals inthe 22 CR2 CAIs measured are shown in Fig. 11 and listedin Table 2. In Fig. 11a–c, the data are plotted on three-isotope oxygen diagrams (d17O vs. d18O). In Fig. 11d, thesame data are plotted as deviations from the terrestrial

Fig. 4. BSE images of the melilite-rich CAIs Asuka 881828-61-4 #1 (a and b), EET 92042-22 #4 (c), and EET 96286-6 #7 (d). Region outlinedin ‘‘a” is shown in detail in ‘‘b”. Al,Ti-Di, Al,Ti-diopside; An, anorthite; Di, diopside; Fus, fusion crust; Mel, melilite; Pv, perovskite; Sp,spinel.


fractionation (TF) line, D17O; each column represents datafor an individual inclusion. Based on the oxygen-isotopecompositions, four groups of CAIs can be identified.

(i) Data for melilite, anorthite, spinel, hibonite, andgrossite of most CAIs are uniformly 16O-rich, withan average D17O value of �23.3 ± 1.9& (2 standarddeviations). A slight spread in the data along theCCAM is seen in melilite, grossite, and spinel (theminerals with enough analyses to see the trend) andimplies a slight variation in the initial oxygen compo-sitions of the CAIs (see also Fig. 11d). The data clus-ter slightly to the right of the carbonaceous chondriteanhydrous mineral (CCAM) line, suggesting a smalldegree of mass-dependent isotope fractionation(Fig. 11a, b, and d), possibly due to melt evaporation(Davis et al., 2005b).

(ii) Data for the melilite–spinel CAI Gao-Guenie (b) #3

plot along a mass-dependent fractionation line paral-lel to the TF line with D17O value of �20.3 ± 1.2&

(Fig. 11b and d). Spinel grains are more 18O-rich thanmelilite.

(iii) Two anorthite-rich, igneous CAIs associated withchondrule materials, Acfer 209 PL91165 #1 andAcfer 209 PL91165 #3, are isotopically heteroge-neous and plot along the CCAM line (Fig. 11c). Rel-ict spinel in Acfer 209 PL91165 #3 and core anorthitein Acfer 209 PL91165 #1 are 16O-rich and plot near

the typical CR2 CAIs (Fig. 11a and d). Anorthitegrains near the CAI–chondrule boundaries andferromagnesian silicates (olivine, low-Ca pyroxene,high-Ca pyroxene) in the chondrule portions are16O-depleted to varying degrees (D17O ranges from�18.5& to �4.8&).

(iv) Spinel grains in the compound object Gao-Guenie (b)

#4 have highly-fractionated oxygen-isotope composi-tions and are 16O-depleted (D17O � �9&) relative totypical CR2 CAIs (Fig. 11c and d). Anorthite andlow-Ca pyroxene are uniformly 16O-depleted(D17O � �2&) and plot along the CCAM line.

3.3. Magnesium-isotope compositions

Twenty of the twenty-two CR2 CAIs were also analyzedfor magnesium isotopes. The results are listed in Tables 3and 4 and illustrated in Figs. 12–19.

3.3.1. Intrinsic mass fractionation

In Fig. 12, we plot the values of intrinsic mass-depen-dent fractionation of individual CAI minerals. The datafor grossite and sodalite are not plotted because appropri-ate standards were not available to correct magnesium-isotope compositions of these minerals for IMF.

Most of the 16O-rich CAIs (shown with white back-ground on the left of Fig. 12) have uniform magnesium

Fig. 5. BSE images of the melilite ± spinel-rich CAIs El Djouf 001

PL91172 #1 (a), Gao-Guenie (b) #1 (b), and Gao-Guenie (b) #2

(c). Al,Ti-Di, Al,Ti-diopside; An, anorthite; Di, diopside; Mel,melilite; Pv, perovskite; Sp, spinel.


isotopic compositions enriched in 25Mg/24Mg. Typicalenrichments are 2–5&/amu and are probably due to meltevaporation (e.g., Grossman et al., 2002; Richter et al.,2006). Two melilite-rich inclusions, Acfer 097 PL92521 #2

(Fig. 3c) and GRA95229-17 #7 (Fig. 2a), show significantvariations in D25Mg, suggesting either multistage formationhistory or more prolonged evaporation during meltcrystallization compared to CAIs with uniform D25Mg.However, no clear relationship was observed between massfractionation and position of the measured spots within the

inclusions. The melilite-rich CAI, GRA 95229-17 #8

(Fig. 2b) has a much larger enrichment of heavy magnesiumisotopes (D25Mg � 15&) compared to other 16O-rich inclu-sions. The melilite-rich core is surrounded by a more refrac-tory mantle rich in grossite, hibonite, and spinel, consistentwith extensive evaporative loss during crystallization of theCAI melt. The CAI, Asuka 881828-61-4 #1 (Fig. 4a and b)shows typical, slightly heavy values of D25Mg for melilite,but anorthite has D25Mg values ranging from zero to��4& (Fig. 12). Anorthite in this inclusion replaces meli-lite, suggesting formation by gas–solid reaction in the solarnebula (Fig. 4b). There is another kind of postcrystalliza-tion isotopic exchange that is often seen in CV chondriteCAIs (Simon et al., 2005). CV CAIs are typically moder-ately enriched in heavy Mg isotopes, with D25Mg � +5&,but it is quite common to find D25Mg values that approach0& near the outside of CAIs, apparently as a result of ex-change with an isotopically normal reservoir. This effect isusually seen in the outer 100–200 lm of the CAI; someCAIs do not show this effect, but most do. The CR2 CAIsstudied here also typically have D25Mg � +5&. We did notsee any evidence for exchange with an external reservoir for16O-rich CAIs.

El Djouf 001 MK #5 (Fig. 1a–b and EA1) is a hibonite–grossite CAI that is also 16O-rich (Fig. 11b). However,unlike the other 16O-rich CAIs, its magnesium isotopesare unfractionated.

CAI Gao-Guenie (b) #3 shows extreme fractionation ofmagnesium isotopes with D25Mg of 22–24& (Fig. 12). Thisinclusion also has mass fractionated oxygen, with meliliteplotting �7&/amu and spinel plotting 10–12&/amu tothe right of the CCAM line (Fig. 11b).

Two of the CAIs associated with chondrule materials,Acfer 209 PL91165 #1 (Fig. 8) and Gao-Guenie (b) #4

(Fig. 10), show relatively large variations in D25Mg of anor-thite (Fig. 12). This spread probably reflects incompletehomogenization between isotopically heavy relict CAIsand isotopically normal chondrule materials. Relict spinelgrains in Gao-Guenie (b) #4 are fractionated by 3–6&/amu; they also show highly-fractionated oxygen-isotopecompositions (Fig. 11). In Acfer 209 PL91165 #3 (Fig. 9),both anorthite and spinel are essentially unfractionated.

3.3.2. 26Al–26Mg systematics

Aluminum–magnesium evolution diagrams for individ-ual inclusions having uniformly 16O-rich compositions(Fig. 11a) are shown in Figs. 13 and 14. All of these CAIs,except El Djouf 001 MK #5, show excesses of radiogenic26Mg (26Mg*) correlated with 27Al/24Mg ratio. For sevenof these inclusions, the spread in 27Al/24Mg ratios is suffi-cient to produce linear arrays that can be interpreted asinternal isochrons (Fig. 13). These inclusions give inferredinitial 26Al/27Al ratios ((26Al/27Al)0) ranging from(4.43 ± 0.24) � 10�5 to (5.43 ± 0.62) � 10�5. For theremaining 16O-rich inclusions, the regression lines areforced through the origin and represent model isochrons.The (26Al/27Al)0 ratios inferred from the internal and modelisochrons are similar and consistent with the canonical va-lue of (4.5–5) � 10�5 (Fig. 14). In fact, the mean value forthe (26Al/27Al)0 ratios from the seven internal isochrons,

Fig. 6. BSE images of a melilite–spinel CAI Gao-Guenie (b) #3 (a), a melilite–spinel CAI in an amoeboid olivine aggregate Gao-Guenie (b) #6

(b and c), and the melilite-rich CAIs GRA 95229-18 #2 (d) and Temple Bar MK #4 (e). Regions outlined in ‘‘b” is shown in detail in ‘‘c”. Al,Ti-Di, Al,Ti-diopside; Al-Di, Al-diopside; An, anorthite; Au, remnants of gold coating; Di, diopside; Mel, melilite; Ol, forsteritic olivine; Pv,perovskite; Sod, sodalite; Sp, spinel.


(4.7 ± 0.12) � 10�5, is virtually identical with the mean ofthe (26Al/27Al)0 ratios from the seven model isochrons,(4.68 ± 0.13) � 10�5. In CAIs having a large spread in27Al/24Mg ratios, the inferred (26Al/27Al)0 are slightly be-low than ratio reported by Jacobsen et al. (2008)[(5.23 ± 0.13)) � 10�5], and are clearly resolved from thesupra-canonical values of (5.85–7) � 10�5 reported forCAIs from metamorphosed CV chondrites (Bizzarroet al., 2004; Young et al., 2005; Taylor et al., 2005; Thraneet al., 2006; Cosarinsky et al., 2007).

One CAI plotted in Fig. 14, GRA 95299-18 #2, de-serves special comment. For this inclusion, spinel, melilite,

and secondary sodalite were measured using monocollec-tion mode. Spinel and melilite fall along a linear arraywith a slope of (5.0 ± 0.5) � 10�5 (Fig. 14f). Sodalite,which is secondary, falls well off of the linear array(Fig. 14f). Because the spinel grain is small and was mea-sured using monocollection, its d26Mg value has largeuncertainties. If we calculate a model isochron with meli-lite that is forced through the origin, the same slope re-sults, but the uncertainty is much lower. Fig. 14g showsthe error envelopes resulting from the two different calcu-lations for the slope of the isochron. This illustrates theimportance of high-precision measurements of phases with

Fig. 7. BSE image of an anorthite-rich CAI fragment GRA 95229-18 #2. Al,Ti-Di, Al,Ti-diopside; An, anorthite; Sp, spinel.

Fig. 8. Combined elemental map in Mg (red), Ca (green), and Al Ka (blue) X-rays (a), elemental map in Al Ka X-ray (b), and BSE images (cand d) of an igneous, anorthite-rich CAI Acfer 209 PL91165 #1. Regions outlined in ‘‘a” are shown in detail in ‘‘c” and ‘‘d”. White numberswith black circles and black numbers with white circles in ‘‘a” show the analyzed spots for magnesium and oxygen isotopic measurements,respectively. The CAI is surrounded by a mantle of high-Ca pyroxene and low-Ca pyroxene with poikilitically enclosed olivine grains,anorthite, and Fe,Ni-metal. Anorthite in the CAI core is granular; anorthite in the chondrule-like mantle is lath-shaped. An, anorthite; FeS,iron sulfide; Hpx, high-Ca pyroxene; Lpx, low-Ca pyroxene; Met, Fe,Ni-metal; Ol, olivine; Sp, spinel; W and white veins are FeO-richproducts of terrestrial weathering.


Fig. 9. Combined elemental map in Mg (red), Ca (green), and AlKa (blue) X-rays (a), and BSE image (b) of the anorthite-spinel-Al,Ti-diopside CAI Acfer 209 PL91165 #3 associated with chon-drule material. White numbers with black circles in ‘‘a” show theanalyzed spots for magnesium isotopic measurements. Blacknumbers with white circles in ‘‘b” show the analyzed spots foroxygen isotopic measurements. Squares in ‘‘b” show sputteredregions for oxygen isotopic measurements. The CAI portionconsists of spinel, Al,Ti-diopside, and anorthite; spinel grainscontain rare inclusions of perovskite. The chondrule portion iscomposed of ferromagnesian low-Ca pyroxene, high-Ca pyroxene,and olivine, and Fe,Ni-metal. Al,Ti-Di, Al,Ti-diopside; An, anor-thite; Di, diopside; FeS, iron sulfide; Hpx, high-Ca pyroxene; Lpx,low-Ca pyroxene; Met, Fe,Ni-metal; Ol, olivine; Pv, perovskite; W,FeO-rich terrestrial weathering products.


low 27Al/24Mg ratios when high-precision isochrons areneeded. For subsequent discussions, we will use the modelisochron for this CAI.

The hibonite–grossite-rich CAI El Djouf 001 MK #5

shows apparent deficits in d26Mg in both grossite and hib-onite and no evidence of 26Mg* (Table 3; Fig. 14i). Theapparent deficit in 26Mg could, in principle, actually be anexcess of 25Mg, but this inclusion shows no evidence formass fractionation in magnesium (Fig. 12). It thus has low-er 25Mg/24Mg ratio than most CR2 CAIs, so the isotope ef-fect is most likely a true deficit of 26Mg.

Fig. 15 shows Al–Mg evolution diagrams for grossite inthe grossite-rich CAIs GRA 95229-17 #7, GRA 95229-17

#8, and GRA 95229-31 #3. Model isochrons for grossite in

these CAIs give systematically lower (26Al/27Al)0 values[(4.2 ± 0.1) � 10�5, (4.3 ± 0.2) � 10�5, and (4.3 ± 0.1) �10�5, respectively than those for other minerals (Table 3;Figs. 13 and 14)]. Also note that for GRA 95229-17 #7 andGRA 95229-17 #8, two different (26Al/27Al)0 ratios can be in-ferred considering data for grossite and melilite separately,with melilite giving higher values (Table 4; Fig. 16). Thesesystematic differences almost certainly reflect an incorrect va-lue of the sensitivity factor for grossite for which we had noappropriate standards. The sensitivity factor for grossitewas assumed to be the same as for hibonite.

Gao-Guenie (b) #3, the inclusion with highly-fraction-ated oxygen and magnesium isotopes (�11 and 23&/amu,respectively; Figs. 11 and 12), has an apparent deficit in26Mg (or excess of 25Mg). The large degree of mass fraction-ation in magnesium could produce an apparent isotopicanomaly in 26Mg, if the wrong fractionation correction isused. However, the observed 26Mg deficit is too large to beeliminated by any reasonable mass fractionation law. Weconclude that the isotope effect is real, but because of the highdegree of mass fractionation, we cannot uniquely assign theanomaly to either 25Mg or 26Mg. Our standard data-reduc-tion procedure assigns it to 26Mg. Fig. 17 shows an Al–Mgevolution diagram for Gao-Guenie (b) #3. A regressionthrough the spinel and melilite data give an inferred(26Al/27Al)0 ratio of (2.0 ± 1.7) � 10�5 with an intercept ofd26Mg = �2.6 ± 0.4&. Although these characteristics areunusual, objects with similar systematics have been reportedbefore. For example, a hibonite-bearing microspherule fromColony showed an inferred (26Al/27Al)0 ratio of(4.4 ± 2.4) � 10�6 with an intercept of d26Mg = �2.6 ±1.3& (Russell et al., 1998).

The compound CAI–chondrule objects, Acfer 209

PL91165 #1, Acfer 209 PL91165 #3, and Gao-Guenie (b)

#4, have low inferred (26Al/27Al)0, indistinguishable from0: (0.10 ± 0.35) � 10�5, (0.37 ± 0.33) � 10�5, and (0.12 ±0.15) � 10�5, respectively (Fig. 18a, c, and e). Note thatthe large uncertainties on the inferred (26Al/27Al)0 resultmainly from propagation of errors from measurements ofthe anorthite standard, which has much higher 27Al/24Mgratio than plagioclase in these compound objects (�330and <160, respectively). If errors from the anorthite stan-dard are not considered, the excess of 26Mg can be resolvedin Gao-Guenie (b) #4 as well (Fig. 18f). Data for relict spi-nel grains in Acfer 209 PL91165 #3 also show resolvable ex-cess of 26Mg (Table 3; Fig. 19). Forced through origin,these data define a model isochron corresponding to the(26Al/27Al)0 of (4.4 ± 0.7) � 10�5, which is much steeperthan the model isochron based on plagioclase data for thisobject (0.37 ± 0.33) � 10�5.

4. DISCUSSION

4.1. Pristine nature of CAIs in CR2 chondrites

Many CAIs in CR2 chondrites are mineralogically pris-tine objects that lack preterrestrial secondary alterationminerals. Several CAIs, including two inclusions withsecondary sodalite ± carbonates replacing melilite, containsecondary phyllosilicates (Figs. 6d, EA2). Because

Fig. 10. Combined elemental map in Mg (red), Ca (green), and Al Ka (blue) X-rays (a), and BSE image (b) of an anorthite-spinel CAI Gao-

Guenie (b) #4 associated with chondrule material. White numbers with black circles in ‘‘a” and black crosses in ‘‘b” show the location of ionprobe spots for magnesium isotopic measurements. Black numbers with white circles in black squares in ‘‘b” show the location of ion probespots for oxygen isotopic measurements. Al,Ti-Di, Al,Ti-diopside; An, anorthite; Di, diopside; FeS, troilite; Hpx, high-Ca pyroxene; Lpx,low-Ca pyroxene; Met, Fe,Ni-metal; Ol, olivine; Pv, perovskite; Sp, spinel; W, FeO-rich terrestrial weathering products.


phyllosilicates and carbonates are common minerals in aqu-eously altered meteorites (e.g., CI and CM), and becausemost CR2 chondrites experienced mild aqueous alteration(Weisberg et al., 1993, 2006; Krot et al., 2006a), we inferthat these secondary minerals most likely formed duringaqueous alteration, possibly on the CR parent asteroid.In contrast to CAIs from metamorphosed CV chondrites(e.g., Allende, Efremovka, Vigarano), CR2 CAIs do notcontain secondary nepheline, andradite, hedenbergite, mag-netite, sulfides, and ferrous olivine. This indicates that CR2CAIs avoided iron–alkali metasomatic alteration that af-fected most CAIs in CV chondrites (Krot et al., 1998 andreferences therein). As a result, oxygen and magnesium-iso-tope compositions of CR2 CAIs must have preserved pri-mary, nebular records of their formation processes suchas evaporation, condensation, and melting.

Although preterrestrial secondary minerals are rare inCAIs from CR2 chondrites, products of terrestrial weather-ing are common in the host meteorites, and are occasionally

observed in the CAIs. For the most part, these consist ofiron oxides, which form veins that cross-cut the CAIs(e.g., Figs. 1 and 6). In addition, incipient replacement ofmelilite grains by calcium-poor aluminosilicates is occasion-ally observed (e.g., Fig. 2b).

4.2. Oxygen-isotope compositions of CR2 CAIs: Implications

for understanding the oxygen-isotope composition of the Sun

and the nature of oxygen-isotope exchange in CAIs

4.2.1. Uniformly 16O-rich CAIs

Most CR2 CAIs are 16O-rich and have very similar, 16O-rich (D17O � �23&) compositions (Fig. 11a and d). Thedata in Fig. 11a exhibit a small spread along the CCAMline that is slightly greater than expected from counting sta-tistics. Although our data are not extensive enough toclearly determine the nature of the spread, examination ofFig. 11d suggests that most inclusions are isotopicallyhomogeneous and that much of the spread is due to slight

Table 2Oxygen-isotope compositions (in &) of individual minerals in CR2CAIs.

Phase/No. d18O ± 2r d17O ± 2r D17O ± 2r16O-rich CAIs

Acfer 097 PL92521 #1

Mel #1 �44.3 ± 0.6 �47.0 ± 0.9 �24.0 ± 0.9Mel #2 �44.1 ± 0.6 �46.9 ± 0.9 �24.0 ± 1.0Sp #3 �43.7 ± 1.1 �47.5 ± 1.0 �24.8 ± 1.2

Acfer 097 PL92521 #2

Mel #1 �42.7 ± 0.6 �48.2 ± 0.9 �26.0 ± 1.0Mel #2 �42.0 ± 0.6 �46.2 ± 0.9 �24.3 ± 0.9Mel #3 �43.2 ± 0.6 �46.2 ± 0.9 �23.7 ± 1.0Mel #4 �42.6 ± 0.6 �45.9 ± 0.9 �23.7 ± 1.0

Acfer 209 PL91165 #2

Mel #1 �44.3 ± 0.6 �46.9 ± 0.9 �23.8 ± 1.0Mel #2 �43.7 ± 0.7 �46.3 ± 0.9 �23.6 ± 1.0Mel #3 �43.8 ± 0.6 �47.6 ± 0.9 �24.9 ± 0.9Sp #4 �40.7 ± 1.1 �45.1 ± 1.0 �23.9 ± 1.2

Asuka 881828-61-4 #1

Mel #1 �44.4 ± 1.3 �47.2 ± 1.1 �24.1 ± 1.3Mel #2 �41.3 ± 1.3 �44.2 ± 1.1 �22.7 ± 1.3Mel #3 �44.7 ± 1.3 �47.9 ± 1.1 �24.7 ± 1.3

EET 92042-22 #4

Mel #1 �44.3 ± 1.3 �46.8 ± 1.1 �23.8 ± 1.3Mel #2 �44.0 ± 1.3 �46.8 ± 1.1 �23.9 ± 1.3Mel #3 �44.6 ± 1.3 �46.7 ± 1.1 �23.5 ± 1.3

EET 96286-6 #7Mel #1 �41.9 ± 1.3 �45.3 ± 1.1 �23.5 ± 1.3Mel #2 �42.2 ± 1.3 �45.3 ± 1.0 �23.4 ± 1.2Mel #3 �41.6 ± 1.3 �44.6 ± 1.1 �23.0 ± 1.3

EL Djouf 001 PL91172 #1

Mel #1 �43.3 ± 1.3 �46.1 ± 1.1 �23.6 ± 1.3Mel #2 �42.7 ± 1.3 �44.8 ± 1.1 �22.6 ± 1.3Mel #3 �42.2 ± 1.3 �44.4 ± 1.1 �22.4 ± 1.3

Gao-Guenie (b) #1

Mel #1 �43.0 ± 0.5 �46.3 ± 1.0 �23.9 ± 1.0Sp #2 �37.4 ± 0.9 �42.9 ± 1.1 �23.4 ± 1.2

Gao-Guenie (b) #2

Mel #1 �42.2 ± 0.5 �45.3 ± 1.1 �23.4 ± 1.1Sp #2 �41.8 ± 0.9 �45.5 ± 1.1 �23.8 ± 1.2

Gao-Guenie (b) #6

Mel #1 �42.7 ± 0.4 �45.6 ± 1.0 �23.4 ± 1.0Mel #2 �40.5 ± 0.4 �43.3 ± 1.0 �22.3 ± 1.0Sp #3 �42.7 ± 0.9 �46.5 ± 1.1 �24.3 ± 1.2

Gao-Guenie (b) #10

Grs #1 �36.1 ± 0.5 �39.3 ± 1.0 �20.6 ± 1.0Grs #2 �39.1 ± 0.6 �42.5 ± 1.1 �22.1 ± 1.1

GRA 95229-17 #7

Mel #1* �45.7 ± 1.3 �48.0 ± 1.5 �24.2 ± 1.6Mel #2* �44.7 ± 1.2 �46.5 ± 1.4 �23.3 ± 1.5Mel #3* �44.1 ± 1.2 �47.5 ± 1.5 �24.5 ± 1.6Grs #4* �47.6 ± 1.2 �47.7 ± 1.6 �23.0 ± 1.8Grs #5* �45.0 ± 1.1 �46.0 ± 1.4 �22.6 ± 1.5

GRA 95229-17 #8

Mel #1* �41.3 ± 1.1 �43.5 ± 1.8 �22.0 ± 1.9Mel #2* �45.3 ± 1.2 �47.5 ± 1.5 �23.9 ± 1.6Mel #3* �43.5 ± 1.1 �45.6 ± 1.5 �23.0 ± 1.6

Table 2 (continued)

Phase/No. d18O ± 2r d17O ± 2r D17O ± 2r

Hib #4* �45.8 ± 1.1 �45.8 ± 1.5 �22.0 ± 1.6Grs #5* �43.4 ± 1.7 �44.4 ± 1.9 �21.9 ± 2.0

GRA 95229-18 #2

Mel #1 �43.6 ± 0.6 �45.4 ± 0.7 �22.7 ± 0.8

GRA 95229-18 #22

An #1 �43.4 ± 0.8 �46.2 ± 0.8 �23.6 ± 0.9An #2 �42.2 ± 0.9 �45.4 ± 0.7 �23.4 ± 0.9Sp #3 �39.7 ± 0.6 �44.6 ± 0.6 �23.9 ± 0.7

GRA 95229-31 #3

Sp #1* �38.2 ± 2.2 �41.6 ± 1.8 �21.7 ± 2.1Mel #2* �42.4 ± 1.3 �44.2 ± 1.5 �22.2 ± 1.6Grs #3* �43.1 ± 1.2 �44.5 ± 1.5 �22.1 ± 1.6Grs #4* �45.2 ± 1.3 �45.6 ± 1.5 �22.1 ± 1.7

Temple Bar MK #4

Mel #1 �39.1 ± 1.3 �42.3 ± 1.1 �22.0 ± 1.3Mel #2 �40.1 ± 1.3 �43.0 ± 1.1 �22.1 ± 1.3Mel #3 �43.5 ± 1.3 �46.6 ± 1.1 �24.0 ± 1.3Mel #4 �42.3 ± 1.3 �45.3 ± 1.1 �23.3 ± 1.3

El Djouf 001 MK #5

Hib #1* �37.6 ± 1.2 �41.3 ± 1.5 �21.8 ± 1.6Hib #2* �38.9 ± 1.1 �42.8 ± 1.4 �22.5 ± 1.5Hib #3* �39.4 ± 1.5 �43.6 ± 1.5 �23.1 ± 1.7Hib #4* �39.5 ± 1.1 �44.0 ± 1.6 �23.5 ± 1.7Grs #5* �39.1 ± 1.1 �44.3 ± 1.4 �24.0 ± 1.5

CAI with highly-fractionated oxygen and magnesium

Gao-Guenie (b) #3

Mel #1 �24.2 ± 0.5 �33.1 ± 1.0 �20.5 ± 1.0Mel #2 �23.8 ± 0.5 �32.8 ± 1.0 �20.4 ± 1.0Mel #3 �24.3 ± 0.9 �32.5 ± 0.8 �19.8 ± 0.9Sp #4 �18.4 ± 0.9 �30.7 ± 1.1 �21.2 ± 1.2Sp #5 �16.5 ± 0.9 �29.5 ± 1.0 �20.9 ± 1.1Sp #6 �16.8 ± 0.9 �28.4 ± 0.8 �19.6 ± 0.9Sp #7 �17.1 ± 0.8 �28.6 ± 0.8 �19.7 ± 0.9

16O-depleted CAI–chondrule compound objects

Acfer 209 PL91165 #1

An #1 �31.8 ± 1.1 �35.0 ± 1.0 �18.5 ± 1.1An #2 �28.5 ± 1.0 �32.1 ± 1.0 �17.3 ± 1.2An #3 �36.4 ± 1.0 �39.9 ± 1.0 �20.9 ± 1.1An #4 �32.4 ± 1.0 �35.0 ± 1.0 �18.2 ± 1.1Lpx #5 �16.5 ± 0.6 �19.8 ± 0.9 �11.2 ± 1.0Hpx #6 �17.1 ± 0.6 �20.6 ± 0.9 �11.7 ± 1.0Lpx #7* �19.2 ± 1.7 �21.5 ± 1.6 �11.6 ± 1.9Lpx #8* �17.9 ± 1.7 �20.4 ± 1.6 �11.1 ± 1.8Lpx #9* �18.3 ± 1.7 �19.4 ± 1.6 �9.9 ± 1.8Hpx #10* �18.0 ± 1.4 �22.1 ± 1.6 �12.7 ± 1.8Ol #11* �16.2 ± 1.0 �19.9 ± 1.4 �11.5 ± 1.5Ol #12* �17.4 ± 1.0 �19.5 ± 1.2 �10.5 ± 1.3

Acfer 209 PL91165 #3

An #1 �13.9 ± 1.0 �17.3 ± 1.0 �10.1 ± 1.1An #2 �14.2 ± 1.0 �17.8 ± 1.0 �10.4 ± 1.1An #3 �18.4 ± 1.0 �23.1 ± 1.0 �13.5 ± 1.1Sp #4 �38.1 ± 1.1 �41.3 ± 1.0 �21.5 ± 1.2Sp #5* �42.2 ± 1.4 �43.7 ± 1.4 �21.7 ± 1.6Sp #6* �34.7 ± 1.4 �39.4 ± 1.5 �21.4 ± 1.7Sp #7* �37.8 ± 1.3 �41.1 ± 1.6 �21.5 ± 1.7Sp #8* �40.9 ± 1.2 �44.8 ± 1.4 �23.5 ± 1.6

(continued on next page)


Table 2 (continued)

Phase/No. d18O ± 2r d17O ± 2r D17O ± 2r

Lpx #9* �6.8 ± 1.7 �10.3 ± 1.4 �6.7 ± 1.7Hpx #10* �8.5 ± 1.6 �11.9 ± 1.4 �7.5 ± 1.6Hpx #11* �4.6 ± 1.8 �7.1 ± 1.5 �4.8 ± 1.8Ol #12* �22.3 ± 0.8 �23.7 ± 1.8 �12.1 ± 1.8

Gao-Guenie (b) #4

An #1 4.1 ± 0.7 �0.2 ± 1.0 �2.4 ± 1.1An #2 5.3 ± 0.8 �0.6 ± 1.1 �3.4 ± 1.1An #3 4.7 ± 0.7 �0.2 ± 1.0 �2.6 ± 1.1An #4 2.4 ± 1.2 �0.5 ± 1.1 �1.7 ± 1.2Sp #5 16.2 ± 0.9 �0.3 ± 1.1 �8.7 ± 1.2Sp #6 16.2 ± 0.5 �0.9 ± 0.8 �9.3 ± 0.9Lpx #7 1.4 ± 0.9 �1.2 ± 0.9 �1.9 ± 1.0

An, anorthite; Grs, grossite; Hib, hibonite; Hpx, high-Ca pyroxene;Lpx, low-Ca pyroxene; Mel, melilite; Ol, Olivine; Sp, spinel.D17O = d17O – 0.52 � d18O.

* Data acquired using primary ion beam rastered over 5 � 5 lm2

area (for details see section Oxygen-isotope measurements). Errorsare 2r.


differences in D17O from inclusion to inclusion. These differ-ences could reflect variations in composition of the CAIprecursors or nearly complete isotopic exchange betweenthe CAI melts and the surrounding gas of slightly differentcomposition.

The isotopic homogeneity of the CR2 CAIs is incontrast with CAIs from the oxidized CV chondrites ofthe Allende-like subgroup, that experienced more extensivefluid-assisted thermal metamorphism than the oxidizedBali-like and the reduced CVs (Krot et al., 1998). MostCAIs in the Allende-like meteorites show evidence for oxy-gen-isotope exchange with an 16O-poor reservoir, shiftingthe compositions of melilite and anorthite towards the TFline, close to the compositions of secondary mineralsproduced during fluid-assisted thermal metamorphism(magnetite, fayalite, hedenbergite, andradite, nepheline,sodalite, wollastonite) (e.g., Clayton et al., 1977; Yurimotoet al., 2008 and references therein; Krot et al., 2006a,2008a). In contrast, several melilite-rich CAIs in Kaba (onlyfive inclusions have been measured so far), the least meta-morphosed CV chondrite of the Bali-like subgroup (Bonalet al., 2006), largely retained their original, 16O-rich compo-sitions (Nagashima et al., 2007a).

The mean D17O value for 16O-rich CR2 CAIs [excludinghighly-fractionated CAI Gao-Guenie (b) #3 and three CAIsassociated with chondrule materials (Table 2)] is�23.3 ± 1.9&. This value is indistinguishable from themean D17O value (�23.6 ± 1.1&) obtained for spinel,Al,Ti-diopside, and forsterite in coarse-grained, igneousCAIs from CV chondrites measured by the UH Camecaims-1280 (MacPherson et al., 2008). Based on these obser-vations, we infer that the CR2 and CV CAIs originatedin an isotopically similar reservoir. This reservoir is similarin composition to that of the Sun inferred from measure-ments of the solar wind implanted into lunar-soil metal(Hashizume and Chaussidon, 2005) and into the collectorreturned by the GENESIS mission (McKeegan et al.,2008). It has been recently proposed that oxygen-isotope

composition of the protoplanetary disk evolved with timefrom 16O-rich (D17O � �25&) to 16O-poor (D17O � 0&)as a result of CO self-shielding and radial transport ofgas and solids in the disk (e.g., Clayton, 2002; Yurimotoand Kuramoto, 2004; Lyons and Young, 2005; Ciesla andCuzzi, 2006; Yurimoto et al., 2007). We suggest that oxy-gen-isotope compositions of the CR2 CAIs, as well as thepre-exchange compositions of the CV CAIs, could have re-corded the primordial composition of the inner protoplan-etary disk – region where most CAIs probably formed(Krot et al., 2009).

4.2.2. Isotopically heterogeneous CAI–chondrule compound

objects

About 5% (8 out of 166) of the CR2 CAIs identified inthis study were remelted with addition of chondrule mate-rial and can be classified as CAI–chondrule compound ob-jects. All three CAI–chondrule compound objects measuredby SIMS [Acfer 209 PL91165 #1, Acfer 209 PL91165 #3,Gao-Guenie (b) #4] have heterogeneous oxygen-isotopecompositions. Spinel grains in Acfer 209 PL91165 #3

largely retained their original 16O-rich signature(D17O � �22&) and plot along CCAM line (Fig. 11c andd). Spinel grains in Gao-Guenie (b) #4 are both highly en-riched in 18O, implying mass-dependent fractionation, and16O-depleted (D17O � �9&) relative to most CR2 CAIs.Anorthite, high-Ca pyroxene and low-Ca pyroxene in allthree objects are 16O-depleted to varying degrees (Fig. 11cand d). The mineralogy, petrography and oxygen-isotopecompositions of the CR2 compound objects indicate thatspinel grains in these objects are relict, whereas anorthiteand high-Ca pyroxene crystallized from the chondrulemelts. Prior to incorporation into the host chondrule melt,spinel in Gao-Guenie (b) #4 had a highly-fractionated oxy-gen composition. Oxygen-isotope heterogeneity among sili-cate minerals in Acfer 209 PL91165 #1 and #3 may indicatethat relict CAIs in these objects experienced partial meltingand dilution by the 16O-depleted host chondrule melt. Inaddition, because the gas phase in the chondrule-formingregion is likely to have been 16O-poor (e.g., Krot et al.,2006b,c), and because isotopic exchange between gas andmelt in this setting is likely to be rapid (Yu et al., 1995),the chondrule melts could have exchanged oxygen to a sig-nificant degree with the surrounding 16O-poor nebular gas.Both dilution by 16O-poor melt and isotopic exchange with16O-poor gas previously have been invoked to explain oxy-gen-isotope heterogeneity of the CAI–chondrule compoundobjects in other chondrite groups (Russell et al., 2005; Krotet al., 2005c, 2008a).

Although compound, CAI–chondrule objects, like thosedescribed here, and chondrules with only a chemical signa-ture of CAI-like precursors (e.g., volatility fractionatedtrace element patterns or Ca, Al-rich bulk compositions)have been previously described in many chondrite groups,such objects are quite rare (Russell et al., 2005 andreferences therein; Krot et al., 2005c, 2007, 2008a). Theseobservations indicate that only a small fraction of CAIswere recycled during chondrule formation.

Because CAIs in most chondrite groups originated in asimilarly 16O-rich reservoir (Yurimoto et al., 2008 and

Table 3Magnesium-isotope data for CR2 CAIs.

Phase/no. 27Al/24Mg ±2r d26Mg* (&) ±2r D26Mg (&) ±2r D25Mg (&) ±2r16O-rich CAIs

Acfer 097 PL92521 #1

Mel #1 54.1 1.5 19.0 2.0 28.3 1.1 4.7 0.8Mel #2 76.6 2.2 26.9 2.4 35.9 1.3 4.5 1.0Mel #3 47.3 0.8 17.6 1.9 24.3 0.9 3.4 0.9Mel #4 105.0 1.7 37.1 2.5 46.1 1.2 4.4 1.1Mel #5 41.0 0.6 14.1 2.0 21.6 1.0 3.8 0.9Sp #6� 2.63 0.02 1.2 0.3 3.2 0.3 1.0 0.1Sp #7� 2.65 0.02 1.1 0.3 3.1 0.3 1.0 0.1

Acfer 097 PL92521 #2

Mel #1 13.1 0.2 4.9 1.5 13.7 0.9 4.5 0.6Mel #2 16.9 0.2 5.9 1.4 9.9 0.8 2.0 0.6Mel #3 15.8 0.2 7.0 1.5 7.1 0.8 0.0 0.7Mel #4 16.1 0.2 6.0 1.9 12.2 1.1 3.2 0.8Mel #5 17.3 0.2 7.0 1.3 6.8 0.8 �0.1 0.6

Asuka 881828-61 #4

An #1 312.8 33.2 103.4 5.3 99.9 3.0 �1.7 2.1An #2 24.3 2.9 7.5 2.5 8.7 1.7 0.6 0.9An #3 133.1 14.4 45.7 3.4 38.4 2.1 �3.6 1.3Mel #4 7.4 0.1 1.9 1.1 10.8 0.5 4.6 0.5Mel #5 15.5 0.2 4.6 1.3 11.1 0.6 3.3 0.6

EET 92042-22 #4

Mel #1 15.1 0.2 5.1 1.1 13.1 0.5 4.1 0.5Mel #2 18.8 0.3 6.0 1.1 13.9 0.5 4.0 0.5Mel #3 15.2 0.2 5.0 1.0 12.3 0.5 3.7 0.5Mel #4 14.2 0.2 4.2 1.1 13.5 0.5 4.8 0.5Mel #5 16.7 0.2 5.5 1.1 13.2 0.5 3.9 0.5Mel #6 19.4 0.3 5.5 1.3 12.7 0.6 3.7 0.6

EET 96286-6 #7

Mel #1 10.2 0.2 3.6 0.8 8.2 0.5 2.4 0.4Mel #2 10.2 0.2 3.3 0.9 7.5 0.5 2.1 0.4Mel #3 9.4 0.1 3.3 1.0 7.9 0.5 2.4 0.5Mel #4 10.4 0.2 3.1 1.0 8.3 0.5 2.7 0.4Mel #5 9.2 0.1 2.8 1.1 8.3 0.5 2.8 0.5Mel #6 9.4 0.1 3.1 0.9 7.3 0.5 2.2 0.4

El Djouf 001 PL91172 #1

Mel #1 18.0 0.3 5.7 1.2 13.6 0.6 4.0 0.5Mel #2 11.7 0.2 4.7 1.0 10.9 0.5 3.1 0.4Mel #3 16.7 0.2 4.9 1.3 13.0 0.7 4.1 0.6Mel #4 18.8 0.3 6.2 1.2 12.6 0.6 3.2 0.6Mel #5 14.8 0.2 5.6 1.3 11.4 0.6 2.9 0.6Mel #6 16.9 0.2 5.5 1.1 13.5 0.6 4.1 0.5Mel #7 18.7 0.3 6.3 1.2 12.9 0.6 3.4 0.5Mel #8 19.7 0.3 6.9 1.4 13.8 0.7 3.5 0.6Mel #9 19.1 0.3 6.6 1.3 13.1 0.6 3.3 0.6

Gao-Guenie (b) #1

Mel #1 23.4 0.3 8.9 1.9 11.5 1.0 1.3 0.8Mel #2 24.7 0.4 9.8 1.6 10.6 1.0 0.4 0.7Mel #3 20.3 0.3 7.5 1.9 11.3 1.0 2.0 0.8Mel #4 20.2 0.4 7.5 1.6 10.3 1.0 1.5 0.6Mel #5 21.5 0.4 7.2 2.0 11.2 1.1 2.1 0.8Sp #6� 2.57 0.01 0.7 0.4 4.0 0.2 1.7 0.1

Gao-Guenie (b) #2

Mel #1 14.0 0.2 5.4 1.3 9.2 0.5 1.9 0.6Mel #2 14.1 0.3 4.4 1.3 9.6 0.5 2.6 0.6

(continued on next page)


Table 3 (continued)

Phase/no. 27Al/24Mg ±2r d26Mg* (&) ±2r D26Mg (&) ±2r D25Mg (&) ±2r

Mel #3 8.7 0.2 2.2 1.4 8.3 0.6 3.1 0.6Mel #4 23.6 0.4 7.2 1.6 13.6 0.6 3.2 0.7Mel #5 7.8 0.2 2.1 1.6 5.2 0.7 1.6 0.7Mel #6 17.4 0.3 5.3 1.6 11.1 0.8 2.9 0.7Sp #7� 2.55 0.01 0.7 0.4 6.0 0.2 2.7 0.2

Gao-Guenie (b) #6

Mel #1 14.4 0.2 5.1 1.3 8.7 0.5 1.8 0.6Mel #2 18.0 0.3 6.4 1.3 10.0 0.5 1.8 0.6Mel #3 13.9 0.2 5.2 1.4 9.8 0.5 2.3 0.6Mel #4 13.8 0.2 5.4 1.3 9.2 0.5 2.0 0.6Mel #5 10.2 0.1 3.9 1.2 7.4 0.4 1.8 0.6Mel #6 15.3 0.2 5.0 1.3 10.7 0.5 2.9 0.6Sp #7� 2.53 0.01 0.8 0.4 5.9 0.2 2.6 0.1

GRA 95229-17 #7

Mel #1 41.4 0.6 15.3 1.7 31.0 0.7 7.9 0.8Mel #2 30.5 0.4 10.4 1.5 17.0 0.8 3.3 0.7Mel #3 22.9 0.3 8.7 1.5 25.9 0.8 8.7 0.7Mel #4 24.8 0.4 9.2 1.6 15.3 0.8 3.1 0.7Mel #5 97.1 1.6 34.9 2.3 35.0 1.1 0.1 1.0Grs #6 3983.6* (188.1) 1193.6* (16.8) (1,159.0) (6.0) (�8.1) (3.6)Grs #7 1895.6* (90.4) 559.4* (10.0) (545.4) (5.4) (�4.6) (2.8)Grs #8 2362.4* (119.3) 705.8* (14.3) (675.4) (10.7) (�9.2) (2.7)Grs #9 2813.0* (134.9) 848.0* (14.2) (809.0) (6.3) (�10.9) (3.5)

GRA 95229-17 #8

Mel #1 33.8 0.6 12.8 1.4 42.2 0.6 14.8 0.6Mel #2 10.0 0.2 2.3 1.5 33.3 0.7 15.8 0.7Mel #3 34.2 0.7 13.5 1.6 41.0 0.7 13.8 0.7Mel #4 8.7 0.1 1.9 1.8 32.6 0.8 15.7 0.9Mel #5 18.2 0.4 6.0 1.4 41.4 0.6 17.9 0.7Mel #6 14.1 0.2 4.8 1.3 35.4 0.5 15.5 0.6Grs #7 1163.0* (82.9) 350.8* (18.2) (362.4) (17.4) (4.4) (2.0)Grs #8 529.3* (26.8) 164.9* (5.4) (167.7) (3.8) (1.2) (1.7)Grs #9 451.2* (24.0) 141.4* (6.1) (140.6) (4.0) (�0.4) (2.1)

GRA 95229-18 #2

Mel #1 15.8 0.2 6.4 1.4 8.7 0.6 1.2 0.6Mel #2 15.9 0.2 5.4 1.3 8.1 0.6 1.4 0.6Mel #3 27.7 0.4 10.0 1.6 12.5 0.7 1.3 0.8Mel #4 23.8 0.5 8.4 2.3 11.4 0.9 1.5 1.1Sp #5 2.57 0.01 1.2 1.7 �0.8 0.7 �1.0 0.8Sod #6 96.8* (10.7) 15.9* (3.3) (10.2) (1.8) (�2.9) (1.4)

GRA 95229-18 #22

An #1 952.5 103.1 299.1 7.2 309.6 5.9 4.1 1.6An #2 932.4 98.8 297.2 5.6 304.4 3.4 2.9 1.8An #3 576.0 61.0 181.5 3.5 190.1 1.9 3.7 1.3An #4 588.3 62.7 190.3 4.7 191.1 2.9 0.4 1.6Sp #5� 2.51 0.02 0.8 0.3 9.6 0.3 4.5 0.1

GRA 95229-31 #3

Grs #1 608.5* (29.0) 189.2* (7.5) (183.3) (3.6) (�2.5) (2.8)Grs #2 581.8* (27.7) 178.9* (5.1) (173.1) (2.7) (�2.5) (1.9)Grs #3 362.5* (17.2) 109.5* (3.1) (102.0) (1.6) (�3.5) (1.2)Grs #4 432.0* (23.6) 130.0* (5.4) (127.7) (3.7) (�1.0) (1.7)Grs #5 1069.7* (53.3) 338.7* (7.4) (324.3) (5.4) (�5.6) (1.9)

Temple Bar MK #4

Mel #1 13.2 0.2 4.2 1.2 14.0 0.5 5.0 0.5Mel #2 13.4 0.2 5.0 1.1 13.9 0.5 4.5 0.5Mel #3 14.3 0.2 5.2 1.1 14.6 0.6 4.8 0.5Mel #4 15.3 0.2 5.5 1.2 14.1 0.6 4.4 0.5


Table 3 (continued)

Phase/no. 27Al/24Mg ±2r d26Mg* (&) ±2r D26Mg (&) ±2r D25Mg (&) ±2r

Mel #5 15.4 0.2 5.3 1.2 13.9 0.6 4.4 0.5Mel #6 13.3 0.2 4.8 1.1 13.5 0.6 4.4 0.5

El Djouf 001 MK #5

Hib #1 139.3 6.6 �2.5 1.5 �2.5 0.7 0.015 0.7Hib #2 150.3 7.1 �2.2 1.5 �3.1 0.6 �0.5 0.7Grs #3 65.3* (3.1) �4.4* (1.6) (�8.1) (0.7) (�1.9) (0.8)Grs #4 766.8* (37.7) �5.4* (4.2) (�6.4) (1.9) (�0.5) (1.9)Grs #5 558.9* (27.5) �1.5* (3.3) (�5.4) (1.5) (�2.0) (1.5)Grs #6 91.9* (4.4) �1.4* (2.3) (�3.6) (1.2) (�1.1) (1.0)

CAI with highly-fractionated oxygen and magnesium

Gao-Guenie (b) #3

Mel #1 12.0 0.2 �1.1 1.8 42.3 1.1 22.1 0.8Mel #2 5.4 0.1 �1.8 1.3 44.5 0.9 23.5 0.5Mel #3 11.6 0.2 �0.9 1.4 43.2 0.9 22.5 0.6Sp #4� 2.45 0.01 �2.4 0.6 42.2 0.5 22.7 0.2Sp #5� 2.51 0.01 �2.2 0.6 44.7 0.5 23.9 0.2

16O-depleted CAI–chondrule compound objects

Acfer 209 PL91165 #1

An #1 91.4 9.9 �0.7 4.4 9.0 2.0 5.0 2.0An #2 25.6 2.7 0.9 4.1 �1.4 1.9 �1.2 1.9An #3 171.8 18.4 2.1 5.1 8.9 2.4 3.5 2.3

Acfer 209 PL91165 #3

An #1 95.5 10.1 3.0 4.2 4.8 1.9 0.9 1.9An #2 103.6 11.0 2.4 4.2 1.0 1.9 �0.7 1.9An #3 86.3 9.1 2.6 4.1 3.2 1.8 0.3 1.9An #4 64.3 6.8 1.3 4.3 4.5 2.0 1.6 1.9Sp #5� 2.50 0.01 0.9 0.3 �0.5 0.2 �0.7 0.1Sp #6� 2.40 0.01 0.8 0.3 0.3 0.2 �0.2 0.1Sp #7� 2.45 0.01 0.7 0.2 1.2 0.2 0.2 0.1Sp #8� 2.56 0.01 0.8 0.3 1.6 0.2 0.4 0.1SP #9� 2.53 0.01 0.8 0.3 1.1 0.2 0.2 0.1

Gao-Guenie (b) #4

An #1 125.2 13.3 1.4 2.2 4.5 1.6 1.5 0.8An #2 83.2 8.8 0.3 2.5 9.5 1.8 4.7 0.9An #3 76.0 8.1 0.5 2.2 4.7 1.6 2.2 0.8An #4 72.9 7.7 0.9 2.2 1.3 1.6 0.2 0.7An #5 115.7 12.3 1.3 2.2 4.5 1.6 1.6 0.8Lpx #6� 0.19 0.01 0.2 0.2 2.9 0.1 1.4 0.1Sp #7� 2.47 0.01 0.0 0.6 7.0 0.5 3.6 0.2Sp #8� 2.49 0.01 0.1 0.6 8.7 0.5 4.4 0.2SP #9� 2.59 0.01 0.1 0.6 11.9 0.5 6.0 0.2

�Data acquired using multicollection mode. The others were acquired using monocollection mode.*Data for grossite and sodalite listed in parentheses were corrected for instrumental mass fractionation (IMF) using hibonite and anorthite

standards, respectively. An, anorthite; Grs, grossite; Hib, hibonite; Lpx, low-Ca pyroxene; Mel, melilite; Sod, sodalite; Sp, spinel.


references therein) and only a small number of CAIs appearto have been affected by chondrule melting events in an16O-poor gaseous reservoir, the commonly observed oxy-gen-isotope heterogeneity in CAIs from metamorphosedCV chondrites, with melilite and anorthite being systemat-ically 16O-depleted compared to spinel, Al-diopside,hibonite, and forsterite, is puzzling. Although gas–solidexchange between 16O-rich melilite and anorthite and16O-poor gas in the solar nebula (Dyl et al., 2008) cannotbe excluded, the lack of oxygen-isotope heterogeneity inCAIs from CR2 (this study) and CO3.0 chondrites (Itohet al., 2004) makes this hypothesis unlikely. We suggest that

16O-depletion of melilite and anorthite in CAIs from meta-morphosed CV chondrites is due to fluid-solid isotopeexchange on the CV asteroidal body (Krot et al., 2008a)rather than gas–melt exchange hypothesized by Yurimotoet al. (1998) and Itoh and Yurimoto (2003). This conclusiondoes not preclude that some CV CAIs experienced oxygen-isotope exchange during remelting in an 16O-poor gas (e.g.,Yurimoto et al., 1998; Kim et al., 2002; Harazono andYurimoto, 2003; Krot et al., 2005c, 2008a; Thrane et al.,2008). Instead it implies that such remelting is unlikely tobe the dominant process responsible for oxygen-isotopeheterogeneity in CV CAIs.

Fig. 11. Oxygen-isotope compositions of individual CAI minerals from CR2 chondrites. In ‘‘a–c”, the data are plotted as d17O vs. d18O, whered17,18O, [(17,18O/16O)sample/(

17,18O/16O)SMOW � 1] � 103; SMOW, Standard Mean Ocean Water. Error bars (2r, �1& and 0.5& for d17O andd18O, respectively) are not shown for clarity. Carbonaceous Chondrite Anhydrous Mineral (CCAM) line and the Terrestrial FractionationLine (TFL) are shown for reference. In ‘‘d”, the same data are plotted as D17O, d17O � 0.52 � d18O; errors are 2r. Each column representsdata for an individual inclusion. Most CAIs are uniformly 16O-rich (D17O 6 �20&), indicating formation in an 16O-rich gaseous reservoirwithout subsequent isotopic exchange. Data for an igneous melilite–spinel-rich CAI Gao-Guenie (b) #3 define a mass-dependent fractionationline parallel to the TFL. CAIs associated with chondrule materials are 16O-depleted to varying degrees, suggesting melting and oxygen-isotopeexchange during chondrule formation. Spinel in one of these CAIs, Gao-Guenie (b) #4, is highly fractionated. An, Anorthite; Grs, grossite;Hib, hibonite; Hpx, high-Ca pyroxene; Lpx, low-Ca pyroxene; Mel, melilite; Ol, olivine; Sp, spinel.


4.3. Magnesium-isotope compositions of CR2 CAIs and their

implications

4.3.1. The initial 26Al/27Al ratio in the Solar System

The (26Al/27Al)0 values inferred from the internal andmodel isochrons of the majority of 16O-rich CR2 CAIsare in the range of (4.4–5.4) � 10�5. Considering the errors,all of these values are consistent with the canonical value of

(4.5–5) � 10�5 (MacPherson et al., 1995). So-called ‘‘supra-conical” (26Al/27Al)0 ratios ranging from 5.85 � 10�5 up to�7 � 10�5 have been reported for whole-rock CAIs frommetamorphosed CV chondrites (e.g., Young et al., 2005;Thrane et al., 2006). The most precise of these determina-tions, which is based on CAIs from four CV chondrites, im-plies a ratio of (5.85 ± 0.05) � 10�5 (Thrane et al., 2006).On the other hand, Jacobsen et al. (2008) determined a

Table 4D17O values and the inferred (26Al/27Al)0 in CR2 CAIs.

Sample D17O ± 2r (&) (26Al/27Al)0 � 10�5 Dt (My)

16O-rich CAIs

Internal isochronsAcfer 097 PL92521 #1 �24.3 ± 1.0 4.89 ± 0.23 �0.09 ± 0.05Asuka 881828-61-4 #1 �23.8 ± 2.0 4.73 ± 0.44 �0.05 ± 0.10Gao-Guenie (b) #1 �23.6 ± 0.7 5.43 ± 0.62 �0.20 ± 0.13Gao-Guenie (b) #2 �23.6 ± 0.5 4.53 ± 0.75 �0.01 ± 0.19Gao-Guenie (b) #6 �23.3 ± 2.0 5.14 ± 0.75 �0.14 ± 0.17GRA 95229-17 #7 melilite �24.0 ± 1.3 4.97 ± 0.49 �0.10 ± 0.11GRA 95229-18 #22 �23.7 ± 0.5 4.43 ± 0.24 0.02 ± 0.06

Model isochronsAcfer 097 PL92521 #2 �24.4 ± 2.2 5.44 ± 0.60 �0.20 ± 0.12EET 92042-22 #4 �23.7 ± 0.4 4.44 ± 0.39 0.01 ± 0.10EET 96286-6 #7 �23.3 ± 0.5 4.60 ± 0.56 �0.02 ± 0.14El Djouf 001 PL91172 #1 �22.9 ± 1.2 4.75 ± 0.34 �0.06 ± 0.08GRA 95229-17 #7 grossite �22.8 ± 0.5 4.16 ± 0.11 0.08 ± 0.03GRA 95229-17 #8 melilite �23.0 ± 1.9 5.12 ± 0.37 �0.14 ± 0.08GRA 95229-17 #8 grossite �21.9 ± 2.0 4.32 ± 0.17 0.04 ± 0.04GRA 95229-18 #2 �22.7 ± 0.8 5.06 ± 0.54 �0.12 ± 0.12GRA 95229-31 #3 �22.0 ± 0.5 4.29 ± 0.11 0.05 ± 0.03Temple Bar MK #4 �22.8 ± 1.9 4.98 ± 0.46 �0.11 ± 0.10El Djouf 001 MK #5 �23.0 ± 1.7 no excess of 26Mg* –

CAI with highly-fractionated oxygen and magnesiumGao-Guenie (b) #3 �20.3 ± 1.2 1.96 ± 1.71 –

16O-depleted CAI–chondrule compound objects model isochronsAcfer 209 PL91165 #1 �20.9 to �9.9 0.10 ± 0.35 P2.42Acfer 209 PL91165 #3 chdondrule �13.5 to �4.8 0.37 ± 0.33 2.63 ± 2.33Acfer 209 PL91165 #3 relict spinel �23.5 to �21.4 4.37 ± 0.66 0.03 ± 0.17

Internal isochronGao-Guenie (b) #4 �9.3 to �1.7 0.12 ± 0.15 P2.98

Relative 26Al–26Mg age Dt (My) = 1/k � ln[(26Al/27Al)Canonical /(26Al/27Al)Sample], where k is the decay constant of 26Al = ln(2)/0.73 and(26Al/27Al)Canonical = 4.5 � 10�5. Negative and positive values for Dt indicate older and younger than the CAI corresponding to the canonicalvalues, respectively. All errors are 2r.


well-defined whole-rock isochron for Allende CAIs thatgave a (26Al/27Al)0 ratio of (5.23 ± 0.13) � 10�5. Thewhole-rock isochron actually defines the time when the bulkCAI becomes a closed system for the aluminum–magne-sium-isotope systematics. The whole-rock data may reflectthe time of aluminum–magnesium fractionation in theCAI precursors relative to the bulk solar nebula. In con-trast, the internal and model isochrons that we have mea-sured most likely represent CAI crystallization ages.While these two distinct events (both of which could havehappened multiple times) may have been very close in time,it is possible that the aluminum–magnesium fractionationof the precursors was in fact measurably earlier than thecrystallization of CAIs minerals. The internal isochronsthat give (26Al/27Al)0 ratios in excess of the canonical valuehave been reported in several CV CAIs by Young et al.(2005) and Cosarinsky et al. (2007). In these cases, the inter-nal isochrons date the crystallization of the CAIs, as ourdata do, so direct comparison is possible. Among the fifteen16O-rich inclusions measured in this study, the two highestinferred (26Al/27Al)0 ratios are found in Gao-Guenie (b) #1

and Acfer 097 PL92521 #2 [both (5.4 ± 0.6) � 10�5] (Figs.

13 and 14). This value is within 2-sigma of the highest su-pra-canonical (26Al/27Al)0 ratios from internal isochrons re-ported by Cosarinsky et al. (2007), and also within errors ofthe canonical range. All other (26Al/27Al)0 ratios inferredfor the 16O-rich CR2 CAIs are less than or equal to5.1 � 10�5. Our data therefore provide no compelling evi-dence in support of supra-canonical (26Al/27Al)0 ratios inCR2 CAIs.

4.3.2. The duration of CAI formation

There is considerable uncertainty about the duration ofCAI formation. On the basis of a tight fit to a whole-rock iso-chron, Jacobsen et al. (2008) suggested that CAI formation inCV chondrites lasted only�20,000 years. On the other hand,Hsu et al. (2000) reported three internal isochrons [(4.6–5) � 10�5, (4.3 ± 0.1) � 10�5, and (3.3 ± 0.8) � 10�5] in anAllende Type B CAI composed of three texturally distinctigneous regions – spinel-free islands, spinel-rich core, andmelilite mantle – and concluded that this CAI recorded mul-tiple heating events during 300,000 years of its evolution.Magnesium-isotope measurements of mixed phases and bulksample of eight coarse-grained CAIs from CV chondrites by

Fig. 12. Mass fractionation of individual minerals in CAIs from CR2 chondrites corrected for instrumental mass fractionation (see text fordetails). Because no appropriate standards were used to correct grossite, sodalite and high-Ca pyroxene for IMF, data for these minerals arenot included in this figure. Errors for standards represent 2r error over one year of measurements. 2r errors for anorthite, hibonite andmelilite standards using monocollection mode are 0.89, 0.66, and 1.34&, respectively. 2r error for spinel and low-Ca pyroxene standards usingmulticollection mode are 0.64 and 0.1&, respectively. Errors for CAI minerals represent 2r error over a measurement session. An, anorthite;Hib, hibonite; Lpx, low-Ca pyroxene; Mel, melilite; Sp, spinel.


Young et al. (2005) show a range of (26Al/27Al)0 from <4 to7 � 10�5. Young et al. concluded that CV CAIs experiencedprolonged thermal history (�0.3 My) in the solar nebula.Cosarinsky et al. (2007) reported significant variations inthe inferred (26Al/27Al)0 ratios [from 3.8 to 6.1 � 10�5] in aset of CV CAIs and their Wark–Lovering rims, and con-cluded that formation and thermal processing of CV CAIslasted for at least 0.5 My.

Before we discuss the implication of our data for con-straining the duration of CAI formation, it is importantto define the term ‘‘CAI formation”.

Most CAIs are surrounded by a single- (typically, Al-diopside) or multilayered rim sequence composed of spi-nel ± hibonite ± perovskite, melilite ± anorthite, Al-diop-side, and forsterite (Wark and Lovering, 1977). Althoughthe formation of these so-called Wark–Lovering rims couldbe a logical choice for the end of CAI formation, someCAIs surrounded by Wark–Lovering rims have an addi-tional, outermost layer composed of 16O-rich forste-rite ± anorthite ± Al-diopside ± Fe,Ni-metal (MacPhersonet al., 1985; Krot et al., 2001a). These so-called forsterite-rich accretionary rims (MacPherson et al., 1985) are miner-alogically and isotopically similar to amoeboid olivineaggregates (AOAs), another type of refractory inclusionscomposed of forsterite ± Fe,Ni-metal and Al-diop-side + spinel + anorthite ± melilite inclusions (Krot et al.,2004). The bulk compositions of both CAIs and AOAsare depleted in volatile and moderately volatile elementsrelative to solar system composition, suggesting formationat high ambient temperature, possibly near the protoSun(Krot et al., 2009). These observations imply a close spatial

and temporal relationship between CAIs, AOAs, and for-sterite-rich accretionary rims on CAIs. Therefore, for thisdiscussion, we will define the CAI formation interval asthe period between when the first CAI precursors were pro-cessed in the solar system and the time when the ambienttemperature in the CAI-forming region dropped belowthe condensation temperature of forsterite. This would tothe first order mark the time when the refractory forsteritelayers were added to CAIs and when AOAs formed. Thetime at which the temperature dropped below the forsteritecondensation temperature cannot be determined by the cur-rent tools of radiochronology. However, the condensationof forsterite apparently took place close in time to evapora-tion, condensation and melting events experienced by CAIsthat can be dated, and provides a petrographic marker thatcan be widely applied. The timescales inferred using thisdefinition represent an upper limit of the duration of CAIformation. Petrographic, compositional, and isotopic dataimply that this limit is not too different from the values re-turned by the 26Al–26Mg clock for CAIs from primitive(unmetamorphosed) chondrites.

Another issue must be addressed before one can estimatethe duration of CAI formation. In order to use any isotopicchronometer, the system being used must have been isoto-pically homogenized by the event that is being dated. Thus,in the case of the 26Al–26Mg system, the 26Al/27Al and the26Mg/24Mg ratios must have been homogenized to a suffi-cient degree that any variations reflecting initial heterogene-ity must be much smaller than variations reflectingradioactive decay over the timescale being dated. Perhapsthe best direct test of the degree of isotopic heterogeneity

Fig. 13. Al–Mg evolution diagrams for 16O-rich CAIs from CR2 chondrites. The regression lines corresponding to (26Al/27Al)0 representinternal isochrons. Gray-dotted lines indicate 2r errors on the slopes. Dashed-dotted line with the supra-canonical value of(26Al/27Al)0 = 5.85 � 10�5 and black-dotted line with the canonical value of (26Al/27Al)0 = 4.5 � 10�5 are shown for the reference. Isochronfor the grossite-bearing melilite-rich CAI GRA 95229-17 #7 is based only on measurements of melilite grains. An, anorthite; Mel, melilite; Sp,spinel.


Fig. 14. Al–Mg evolution diagrams for 16O-rich CAIs from CR2 chondrites. The regression lines corresponding to (26Al/27Al)0 are forcedthrough origin and represent model isochrons. Data point for secondary sodalite in GRA 95229-18 #2 (see ‘‘f”) was not considered forcalculating the model isochron of this CAI. Gray-dotted lines indicate 2r errors on the slopes. Dashed-dotted line with the supra-canonicalvalue of (26Al/27Al)0 = 5.85 � 10�5 and black-dotted line with the canonical value of (26Al/27Al)0 = 4.5 � 10�5 are shown for the reference.Isochron for the grossite-bearing melilite-rich CAI GRA 95229-17 #8 is based only on measurements of melilite grains. An, anorthite; Grs,grossite; Hib, hibonite; Mel, melilite; Sod, sodalite; Sp, spinel.


is the tightness of fit of the whole-rock isochrons for Al-lende inclusions produced by Thrane et al. (2006) and Jac-obsen et al. (2008). If the entire uncertainty in the isochronswere in fact due to isotopic heterogeneity in the system, thiswould imply a <3% heterogeneity in the 26Al/27Al ratio.Thus, inclusions formed at the same time from the isotopicreservoir described by the whole-rock isochrons of Thraneet al. (2006) and Jacobsen et al. (2008) would have hadthe same 26Al/27Al ratio within 3%.

We now turn to the formation interval for CAIs. Basedon the above definition, we can significantly restrict therange of proposed formation intervals found in the

literature. For example, the uncertainty in the Jacobsenet al. (2008) whole-rock CAI isochron implies either theCAI precursors formed in less than 20,000 years or theyformed in a single nearly instantaneous event from a reser-voir with less than 3% heterogeneity in the 26Al/27Al ratio(note, however, that there is a discrepancy between thewhole-rock isochrons for the CV CAIs reported by Thraneet al., 2006 and Jacobsen et al., 2008). Isotopic heterogene-ity fades into insignificance as the timescale increases.

The internal and model isochrons reported here mostlikely reflect the time of crystallization of the CR2 CAIs.The spread in (26Al/27Al)0 ratios between the highest

Fig. 15. Al–Mg evolution diagrams for grossite-rich CAIs fromCR2 chondrites. Data only for grossite are plotted. The sensitivityfactor for grossite is assumed to be the same as for hibonite (Weberet al., 1995). The regression lines corresponding to (26Al/27Al)0 areforced through origin and represent model isochrons. Gray-dottedlines indicate 2r errors on the slopes. Dashed-dotted line with thesupra-canonical value of (26Al/27Al)0 = 5.85 � 10�5 and black-dotted line with the canonical value of (26Al/27Al)0 = 4.5 � 10�5

are shown for reference.


(5.4 � 10�5) and lowest (4.4 � 10�5), if taken at face value,implies a time interval of 220,000 years. The range in thedata is considerably greater than one would expect based

on the individual uncertainties, implying that much of thevariation is real. If the uncertainties are included, the min-imum formation interval permitted by our data is �95,000and the maximum is �400,000 years.

4.3.3. The time interval between CAI and chondrule

formation

Three CR2 CAIs measured here experienced meltingand oxygen-isotope exchange in an 16O-poor gaseous reser-voir during chondrule formation. The inferred (26Al/27Al)0

of these CAIs (<0.7 � 10�5) indicate melting >2 My aftercrystallization of CAIs with the canonical (26Al/27Al)0.The inferred low values of (26Al/27Al)0 of the CR2 CAIsmelted with chondrule materials are consistent with the26Al–26Mg measurements of CR2 chondrules, which indi-cate that most CR2 chondrules formed �2.5 to >4 My afterthe CAIs (Nagashima et al., 2007b, 2008). The >2 My agedifference is also generally consistent with the 207Pb–206Pbage difference of �2.4 ± 0.7 My between CV CAIs withthe canonical (26Al/27Al)0 ratio and CR2 chondrules (Ame-lin et al., 2002, 2006), further supporting the chronologicalsignificance of 26Al–26Mg systematics.

4.4. Gao-Guenie (b) #3: First FUN CAI in CR2 chondrites?

The 16O-rich spinel–melilite CAI, Gao-Guenie (b) #3,with highly-fractionated oxygen- and magnesium-isotopecompositions (�11 and 23&/amu, respectively), has mag-nesium with an apparent deficit in 26Mg (or excess of25Mg), and shows evidence for live 26Al. The large degreeof mass fractionation and the nonlinear isotopic anomalyin magnesium suggest that Gao-Guenie (b) #3 is a FUNCAI (Fractionation and Unidentified Nuclear effects; Was-serburg et al., 1977). Additional isotope measurementsshould be made to search for nonlinear isotopic anomaliesand to evaluate the degree of mass fractionation in otherelements. The relatively high inferred (26Al/27Al)0 ratio(2.0 ± 1.7 � 10�5) in Gao-Guenie (b) #3 is unusual for aFUN inclusion, but is not unprecedented. Vigarano 1623-5, a forsterite-rich CAI with highly mass-fractionated oxy-gen, magnesium, and silicon and nuclear anomalies in manyelements, also exhibits evidence for 26Al at a level similar tothat inferred from Gao-Guenie (b) #3 (McKeegan et al.,2005). If Gao-Guenie (b) #3 is confirmed to be a FUNCAI, it will be the first such inclusion from CR2 chondrites.

Gao-Guenie (b) #3 differs from most other FUN inclu-sions in two ways. First, the mass-fractionation line alongwhich the minerals plot lies at a D17O value of�20.3 ± 1.2& (Fig. 11b) rather than at D17O � �23.5&,where spinel, Al,Ti-diopside, and forsterite in most FUNinclusions from CV chondrites plot (Davis et al., 2000; Krotet al., 2008b; Thrane et al., 2008; Hiyagon and Hashimoto,2008). Second, the degree of oxygen-isotope fractionationof these minerals within an individual FUN inclusion isgenerally consistent with the inferred crystallization se-quence of coarse-grained igneous CAIs (compact Type A,Type B and forsterite-bearing Type B): spinel ? forste-rite ? Al,Ti-diopside (Stolper, 1983). In Gao-Guenie (b)

#3, however, spinel, which appears to have crystallized first(euhedral spinel grains are enclosed by melilite), is more

Fig. 16. Al–Mg evolution diagrams of the CAI GRA 95229-17 #7. ‘‘a” and ‘‘c” and ‘‘b” and ‘‘d” are uncorrected raw data and corrected datafor melilite and grossite, respectively. The sensitivity factor for grossite is assumed to be the same as for hibonite (Weber et al., 1995). Grs,grossite; Mel, melilite.


fractionated than melilite (Fig. 11b), the opposite of what isexpected from a melt that is crystallizing as it evaporates.These two differences imply that Gao-Guenie (b) #3 hada different and probably more complex history than FUNinclusions from CV chondrites. The fact that Gao-Guenie

(b) #3 is less 16O-rich than both the majority of CAIs fromCR2 chondrites and most FUN CAIs from CV chondritessuggests that the precursor material of Gao-Guenie (b) #3

was already 16O-depleted prior to melting event that re-sulted in evaporation and isotopic fractionation. The factthat spinel is more fractionated than melilite could indicatethat the CAI melt began to experience back reaction withan unfractionated gas phase toward the end of its crystalli-zation, lowering the degree of mass fractionation in theresidual melt.

4.5. Implications of negative d26Mg and absence of 26Al in

some CAIs

Two 16O-rich CAIs, El Djouf 001 MK #5 and Gao-Gue-

nie (b) #3 exhibit distinctly negative d26Mg values (Fig. 14i,Fig. 17). One of these, Gao-Guenie (b) #3, also has evidenceof live 26Al. Other rare CAIs (platy hibonites and pyroxene-hibonite microspherules) with similar apparent 26Mg defi-

cits, some with evidence for 26Al and some without, havebeen found in CM and CO chondrites (e.g., Ireland, 1988;Russell et al., 1998; Liu et al., 2006). Many FUN inclusionsalso show 26Mg deficits (Lee and Papanastassiou, 1974;Esat et al., 1978; Clayton et al., 1984; Papanastassiou andBrigham, 1989; Loss et al., 1994; Thrane et al., 2008). Theseinclusions pose a problem for a simple picture of the chro-nology of CAI formation.

First, we discuss the apparent deficits of 26Mg. The mag-nesium-26 in the solar system consists of a majority ofatoms that were synthesized as 26Mg in various stellarsources and atoms that decayed from 26Al either in theinterstellar medium or in the solar system. If the anomaliesare truly associated with deficit of 26Mg, they could reflect alow contribution from 26Al. However, in a solar composi-tion, with Al/Mg � 0.1, 26Al/27Al ratios of several � 10�3

would be required to eliminate the inferred deficits of26Mg* in these inclusions. It is not plausible that the solarsystem had such a high 26Al/27Al ratio. Thus, we conclude,as have previous authors (e.g., Wasserburg et al., 1977;Russell et al., 1998) that the 26Mg deficits are unrelated tothe presence or absence of 26Al and that the magnesium-isotope anomalies reflect formation from material withslightly nonsolar isotope ratios. The origin of the nonsolar

Fig. 17. Three-isotope magnesium diagram (a) and Al–Mg evolu-tion diagram (b) of the melilite–spinel CAI Gao-Guenie (b) #3

characterized by highly-fractionated oxygen isotopes (see Fig. 11b).Errors are 2r. Mass-dependent fractionation curve (dashed curvein ‘‘a”) is shown for reference.


magnesium is not easy to determine, but the association ofmagnesium-isotope anomalies with anomalies in titaniumand calcium (e.g., Ireland, 1988; Lee et al., 1978) suggest thatall of these anomalies reflect incomplete mixing of nucleosyn-thetic components that contributed to the solar system.

Among inclusions with deficits of 26Mg, some show noevidence of live 26Al (e.g., Fig. 14h), while others showapparent isochrons, but lower inferred (26Al/27Al)0 ratios,ranging from �4.4 � 10�6 to �2 � 10�5, than the majorityof CAIs (Fig. 17; Ireland, 1988; Russell et al., 1998; McKe-egan et al., 2005). These observations confirm the decou-pling of 26Al and magnesium-isotope anomalies inferredabove. Inclusions without evidence for 26Al could haveformed late, after the 26Al decayed. In this study, threeCAIs with associated with chondrule material have low in-ferred (26Al/27Al)0 ratios, and these are interpreted as hav-ing been remelted during chondrule formation �2 My afterthe CAIs originally formed. These inclusions have also ex-changed oxygen with an 16O-poor reservoir. In contrast, El

Djouf 001 MK #5 and the FUN inclusions formed in an16O-rich environment (D17O � �23.5&), like the majority

of CAIs (Fig. 11b; Krot et al., 2008b). If one accepts theargument that the oxygen-isotope composition of the solarnebula evolved from the 16O-rich composition observed inCAIs to compositions near the TF line (e.g., Krot et al.,2005a; Aleon et al., 2007), then the oxygen data suggest thata small subset of CAIs formed early, either without or withvery low abundances of 26Al. This implies some sort of het-erogeneity in the 26Al distribution in region where CAIsformed.

There are two fundamental classes of isotopic heteroge-neity to consider. First, 26Al could have been heteroge-neously distributed on a spatial scale that was large withrespect to individual inclusions, so inclusions formed inone area got 26Al and those formed in another area didnot. Another version of this class is that the abundance of26Al in an area changed with time. The second class of het-erogeneity is at the scale of individual grains in the solarnebula. If 26Al were preferentially located in grains of a par-ticular type, and processes in the solar nebula could en-hance or deplete the abundance of those grains relative tothe rest of the material in the nebula, the result would bedifferent initial abundances of 26Al, even though the origi-nal material from the pre-solar cloud was well-mixed. It isbeyond the scope of this paper to discuss these different op-tions in detail; they will be discussed in a subsequent paper.But it is important to note that a small subset of CAIsapparently formed early but without the canonical abun-dance of 26Al. For this subset of inclusions, 26Al cannotbe used as a chronometer. It appears that oxygen isotopesprovide a way to distinguish between initial heterogeneityin the abundance of 26Al and difference in (26Al/27Al)0

caused by the passage of time.

4.6. Correlation of (26Al/27Al)0 and D17O in CR2 CAIs

As discussed above, the 16O-rich CR2 CAIs show asmall range of D17O values outside the range of experimen-tal uncertainties (Fig. 11a and d). The same inclusions alsoshow a small spread in inferred (26Al/27Al)0 ratios, slightlylarger than expected from the uncertainties if all inclusionsformed with the same (26Al/27Al)0 ratio. These observa-tions allow us to search for a possible evolution of oxy-gen-isotope composition of the inner solar system (regionwhere CAIs probably formed) with time. In Fig. 20, weplotted the average value of D17O from minerals analyzedwithin an individual CAI vs. the inferred (26Al/27Al)0 inthese objects. CAIs with isochrons inferred from grossiteare not plotted. For compound CAI–chondrule objectsthe range in D17O is shown rather than the average value.The observed differences between the oxygen- and magne-sium-isotope compositions of CR2 CAIs and CAI–chon-drule compound objects (this study), and between CR2CAIs and CR2 chondrules (Nagashima et al., 2007a,b)are generally consistent with evolution of the oxygen-iso-tope composition of the inner solar nebula with time as in-ferred in CO self-shielding models of Yurimoto andKuramoto (2004) and Lyons and Young (2005). However,the data for the fifteen CAIs for which we have both oxy-gen and magnesium data [excluding grossite, and twoisotopically anomalous CAIs, El Djouf 001 MK #5 and

Fig. 18. Al–Mg evolution diagrams of CAIs associated with chondrule materials. Line with (26Al/27Al)0 of 1 � 10�5 is shown for reference. In‘‘a”, ‘‘b”, ‘‘c”, and ‘‘d”, the regression lines are forced through the origin and represent model isochrons In ‘‘e” and ‘‘f”, the regression linesrepresent internal isochrons. Errors are 2r. In ‘‘a”, ‘‘c”, and ‘‘e”, errors from the terrestrial anorthite standard are propogated for anorthite inthe CAI–chondrule objects analyzed; in ‘‘b”, ‘‘d”, and ‘‘f”, errors from the anorthite standard are not propagated for the analyzed samples.


Gao-Guenie (b) #3] show no correlation between D17O and(26Al/27Al)0 (Fig. 20). This suggests that there was no sig-nificant evolution of D17O in the CAI-forming region overthe duration of CR CAI formation, which is estimated tobe between 100,000 to 400,000 years. If the CAIs truly re-cord no variation in oxygen-isotope composition, it mightsuggest a brief duration of CAI formation consistent withThrane et al. (2006) and Jacobsen et al. (2008). Gao-Guenie

(b) #3 may show the first indication of evolution of

oxygen-isotope composition. Because this inclusion haslower (26Al/27Al)0 [(2.0 ± 1.7) � 10�5] than the canonicalvalue in most CR2 CAIs and is slightly 16O-depletedcompared to them, it may have experienced multistageformation and recorded the evolution of oxygen-isotopecompositions in the early solar nebula over 0:9�0:7

þ2:2 My.Whether this formation occurred in the CAI- orchondrule-forming region is unclear, because no peripheralportion of the inclusion is preserved.

Fig. 19. Al–Mg evolution diagram of relict spinel (a) and three-isotope oxygen diagram (b) of individual minerals from Acfer 209

PL 91165 #3, CAI associated with chondrule materials. In ‘‘a”, theregression line is forced through the origin and represents a modelisochron. Errors are 2r. Dashed-dotted line with the supra-canonical value of (26Al/27Al)0 = 5.85 � 10�5 and black-dottedline with the canonical value of (26Al/27Al)0 = 4.5 � 10�5 areshown for the reference. An, anorthite; Hpx, high-Ca pyroxene;Lpx, low-Ca pyroxene; Ol, olivine; Sp, spinel.


4.7. Comparison of CAIs from CR-clan meteorites

Based on chemical, mineralogical and bulk oxygen-iso-tope similarities, Weisberg et al. (1995, 2001, 2006) andKrot et al. (2002) grouped the CR, CH and CB chondritestogether into a CR-clan of meteorites, suggesting a petroge-netic kinship. Subsequently, Krot et al. (2005b, 2008c,d),Krot and Nagashima (2008), and Krot (2008) confirmedthat CH and CB chondrites have a close genetic relation-ship. Both groups contain abundant nonporphyritic(cryptocrystalline and skeletal) magnesian chondrules,chemically zoned metal grains, and compositionally uni-form metal–sulfide nodules that appear to have formed inan impact generated gas–melt plume (Krot et al., 2005b).However, the CR chondrites are almost certainly distinct;they do not contain such chondrules and metal grains.The CR chondrules are primarily type I porphyritic

chondrules (Weisberg et al., 1995) and many, if not most,of the metal grains were melted during CR chondrule-form-ing event(s) (Connolly et al., 2001). CR chondrites havearound 30% matrix (e.g., McSween, 1977), whereas the neb-ular fines in CB and CH chondrites reside primarily inchondrite lithic clasts (e.g., Bischoff et al., 1993b; Krotet al., 2002; Greshake et al., 2002).

The mineralogy and oxygen- and magnesium-isotopecompositions of CR2 CAIs described here provide no sup-port for a genetic link between CR and the metal-rich CHand CB chondrites either. (i) CAIs in CR2 chondrites aredominated by melilite-rich (melilite ± pyroxene + spinel)inclusions (AOAs are the most common refractory inclu-sions in CR2 chondrites, but they were largely neglectedin this study); other types of CAIs – hibonite-rich, gros-site-rich, and anorthite-rich – are rare. In contrast, CAIsin CH and CB chondrites are dominated by grossite-rich,hibonite-rich and pyroxene � spinel � forsterite ± melilitetypes of inclusions; AOAs are quite rare (Bischoff et al.,1993b; Weber and Bischoff, 1994; Krot et al., 2001b,2008d). (ii) Although most CAIs in CH chondrites are uni-formly 16O-rich, more than 50% of them lack evidence for26Mg* (Kimura et al., 1993; Weber et al., 1995; Krotet al., 2009). CAIs in CB chondrites are 16O-depleted andhave no evidence for 26Mg* (Krot et al., 2001b; Gounelleet al., 2007). In contrast, nearly all CR2 CAIs are 16O-richand have the canonical (26Al/27Al)0.

To summarize, the high-temperature components(CAIs, chondrules, and metal grains) of CR2 chondritesconsist primarily of early-formed pristine nebular materials,while the CH and CB chondrites contain a major compo-nent of material apparently produced during a late-stageplanetary-scale collision. These observations provide noindication of a petrogenetic relationship between CR chon-drites and metal-rich CH and CB chondrites. We concludethat grouping the CH and CB chondrites together with CRchondrites into the CR clan is, therefore, both unjustifiedand misleading.

5. CONCLUSIONS

1. Most CAIs in CR2 chondrites are mineralogically pris-tine; only few of them contain secondary sodalite, car-bonates, and phyllosilicates – products of aqueousalteration. It seems that CR2 CAIs preserved essentiallyunchanged records of their nebular formation region(s).

2. Spinel, hibonite, grossite, anorthite, and melilite in mostCR2 CAIs have 16O-rich (D17O = �23.3 ± 1.9&) com-positions and show no evidence for postcrystallizationisotopic exchange, commonly observed in CAIs fromCV chondrites.

3. The inferred (26Al/27Al)0 in the majority of 16O-richCAIs are consistent with the canonical value of (4.5–5) � 10�5 and a short duration (<0.5 My) of CAI forma-tion. Our data do not support the ‘‘supra-canonical”(26Al/27Al)0 value of (5.85–7) � 10�5 inferred from‘‘whole-rock” and mineral isochrons of CAIs from meta-morphosed CV chondrites.

Fig. 20. (26Al/27Al)0 vs. D17O values in individual CAIs from CR2 chondrites. Filled and open circles correspond to the internal and modelisochrons, respectively. White rectangles show the range of oxygen-isotope compositions from CAI associated with chondrule materials. Thedot-filled region shows a range of oxygen-isotope compositions of CR chondrules (data from Krot et al., 2006b and Connolly et al., 2008).The light-grey area shows a range of (26Al/27Al)0 in CR2 chondrules (data from Nagashima et al., 2008). The grey regions correspond to thecanonical and supra-canonical (26Al/27Al)0 of (4.5–5) � 10�5 and (5.85–7) � 10�5, respectively. Errors are 2r.


4. Two CR2 CAIs, El Djouf 001 MK #5 and Gao-Guenie (b)

#3, show deficits of 26Mg relative to normal magnesium,suggesting formation from material with slightly nonsolarmagnesium composition. One of these inclusions, El

Djouf 001 MK #5, shows no evidence for 26Al, whileGao-Guenie (b) #3 apparently formed with 26Al at a leveljust under half of that in most CR2 CAIs. These inclusionsapparently formed early but did not acquire the canonicallevel of 26Al. Similar inclusions are found occasionally inseveral chondrite classes and, for these relatively rareinclusions, 26Al cannot be used as a chronometer.

5. The 16O-rich spinel–melilite CAI Gao-Guenie (b) #3 hashighly-fractionated oxygen- and magnesium-isotopecompositions (�11 and 23&/amu, respectively) andshows deficit in 26Mg. This CAI formed with a lower-than-canonical abundance of 26Al [(2.0 ± 1.7) � 10�5],like FUN CAIs in CV chondrites. Additional isotopemeasurements (e.g., Ca, Ti, Si) are required to confirmthat this CAI is a truly FUN inclusion.

6. Eight of 166 CAIs identified (�5%) in 47 sections of 15CR2 chondrites experienced incomplete melting duringchondrule formation. Three igneous CAI–chondrulecompound objects measured show large variations in

oxygen-isotope compositions (D17O ranges from�23.5& to �1.7&). This oxygen-isotope heterogeneityreflects the presence of relict CAI grains and incompletemixing of the 16O-enriched CAI melt and 16O-depletedferromagnesian silicate melt, and, possibly, exchangewith an 16O-poor (D17O > �5&) nebular gas duringmelting. The inferred low (26Al/27Al)0 of the CAI–chon-drule compound objects (<0.7 � 10�5) indicate melting>2 My after crystallization of CAIs with the canonical(26Al/27Al)0 and evolution of the oxygen-isotope compo-sition of the inner solar nebula on a similar timescale.

7. Because CAIs in CR and CV chondrites appear to haveoriginated in a similarly 16O-rich reservoir and only asmall number of CAIs were affected by chondrule melt-ing events in an 16O-poor gaseous reservoir, the com-monly observed oxygen-isotope heterogeneity in CAIsfrom metamorphosed CV chondrites is most likely dueto fluid–solid isotope exchange on the CV asteroidalbody rather than gas–melt exchange in the solar nebula.This conclusion does not preclude that some CV CAIsexperienced oxygen-isotope exchange during remelting,instead it implies that such remelting is unlikely to bethe dominant process responsible for oxygen-isotope


heterogeneity in CV CAIs. Experiments are needed totest this hypothesis.

8. The mineralogy and oxygen- and magnesium-isotopecompositions of CAIs in CR chondrites are differentfrom those in the metal-rich CH and CB carbonaceouschondrites. These observations and a growing body ofother data provide no justification for grouping CR,CH and CB chondrites into the CR clan.

ACKNOWLEDGMENTS

We thank Harold Connolly, Andy Davis, Yunbin Guan, andSara Russell for critical comments and suggestions which helpedto improve the manuscript. We thank M. Killgore for providingsamples of El Djouf 001 and Temple Bar for this study. This workwas supported by NASA grants NNX07AI81G andNNX08AH91G (A.N. Krot, P.I.), NAG5-4212 (K. Keil, P.I.),NNX08AG58G (G.R. Huss, P.I.), and NNH04AB47I (I.D.Hutcheon, P.I.) and by the LLNL Institute of Geophysics andPlanetary Physics. This work was performed under the auspicesof the U.S. Department of Energy by the University of California,Lawrence Livermore National Laboratory under Contract DE-AC52-07NA27344. This is Hawai‘i Institute of Geophysics andPlanetology publication No. 1797 and School of Ocean and EarthScience and Technology publication No. 7727.

APPENDIX A. SUPPLEMENTARY DATA

Supplementary data associated with this article can befound, in the online version, at doi:10.1016/j.gca.2009.01.042.

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Associate editor: Sara S. Russell

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