Oxygen Isotopic Variations in the Calcium, Aluminum-rich Inclusion–forming Region Recorded by a Single Refractory Inclusion from the CO3.1 Carbonaceous Chondrite Dar al Gani 083

Calcium, aluminum-rich inclusions (CAIs) are the oldest solids dated that formed in the solar system. Most CAIs in unmetamorphosed chondritic meteorites (chondrites; petrologic type ≤3.0) have uniform solar-like 16O-rich compositions (Δ17O ∼ −24‰) and a high initial 26Al/27Al ratio [(26Al/27Al)0] of ∼(4–5) × 10−5, consistent with their origin in a gas of approximately solar composition during a brief (<0.3 Ma) epoch at the earliest stage of our solar system. The nature of O-isotope heterogeneity in CAIs (Δ17O range from ∼−24 up to ∼+5‰) from weakly metamorphosed chondrites (petrologic type >3.0) remains an open issue. This heterogeneity could have recorded fluctuations of O-isotope composition of nebular gas in the CAI-forming region and/or postcrystallization O-isotope exchange of CAI minerals with aqueous fluids on the chondrite parent asteroids. To obtain insights into possible processes resulting in this heterogeneity, we investigated the mineralogy, rare-earth element abundances, and O- and Mg-isotope compositions of a CAI from the CO3.1 chondrite Dar al Gani 083. This concentrically zoned inclusion has a Zn-hercynite core surrounded by layers of (from core to edge) grossite, spinel, melilite, and Al-diopside. The various phases have heterogeneous Δ17O (from core to edge): −2.2 ± 0.6‰, −0.9 ± 2.1‰, −13.7 ± 2.1‰, −2.6 ± 2.3‰, and −22.6 ± 2.1‰, respectively. Magnesium-isotope compositions of grossite, spinel, melilite, and Al-diopside define an undisturbed internal Al–Mg isochron with (26Al/27Al)0 of (2.60 ± 0.29) × 10−6. We conclude that the variations in Δ17O of spinel and diopside recorded fluctuations in O-isotope composition of nebular gas in the CAI-forming region prior to injection and/or homogenization of 26Al at the canonical level. The 16O depletion of grossite and melilite resulted from O-isotope exchange with asteroidal fluid, which did not disturb Al–Mg isotope systematics of the CAI primary minerals.

Calcium, aluminum-rich inclusions are the oldest solids dated that formed in the solar system (Connelly et al. 2012), and therefore represent the cosmochemical time zero (t 0 ) of its formation.Based on the mineralogy, petrography, traceelement abundances, and isotopic compositions, it is thought that CAIs formed during a relatively short period of time (<0.3Ma) in a gas of approximately solar composition close to the young Sun and were subsequently dispersed throughout the protoplanetary disk by still unknown mechanism(s).This includes various models like the X-wind, turbulent diffusion, disk wind, or a viscous expanding disk (e.g., Shu et al. 1996;McKeegan et al. 2000McKeegan et al. , 2011;;Brownlee et al. 2006;Krot et al. 2009;Ciesla 2010;Wielandt et al. 2012;Yang & Ciesla 2012;Desch et al. 2018;Kawasaki et al. 2020;Dunham et al. 2022).However, refractory inclusions formed at this location are not all the same, as different mineralogical and isotopic types exist, e.g., "normal" CAIs and AOAs (Grossman & Steele 1980), fractionated and unidentified nuclear effects (FUN and UN) CAIs (Clayton & Mayeda 1977;Wasserburg et al. 1977;Krot et al. 2014;Park et al. 2017;Kööp et al. 2018), spinel-hibonite inclusions (SHIBs) and platy hibonite crystals (PLACs; Ireland 1988;Liu et al. 2009Liu et al. , 2012;;Kööp et al. 2016aKööp et al. , 2016b)).But all these inclusions have in common that they show mass-independent isotopic anomalies in many elements, and therefore provide an opportunity for examining the formation and evolution of the earliest isotopic reservoir(s) in the protoplanetary disk.
In unmetamorphosed chondrites (petrologic type 3.0), most CAIs have the inferred initial 26 Al/ 27 Al ratio [( 26 Al/ 27 Al) 0 ] of ∼5.0 × 10 −5 , named the canonical (MacPherson et al. 1995;Makide et al. 2009;Larsen et al. 2011;Makide et al. 2013;Ushikubo et al. 2017;Kawasaki et al. 2020;MacPherson et al. 2020).These "canonical" or "normal" CAIs have Original content from this work may be used under the terms of the Creative Commons Attribution 4.0 licence.Any further distribution of this work must maintain attribution to the author(s) and the title of the work, journal citation and DOI.
To get insight into the origin of O-isotopic heterogeneity within CAIs of weakly metamorphosed chondrites, we investigated the mineralogy, petrography, REE abundances, and O-and Al-Mg isotope compositions of a grossite-bearing and Zn-rich CAI from the CO3.1 chondrite Dar al Gani (DaG) 083.This special CAI may have recorded both the early fluctuations in O-isotope composition of nebular gas in the CAI-forming region and the subsequent O-isotope exchange with an 16 O-poor aqueous fluid on the CO chondrite parent asteroid.

Samples and Methods
The polished section of DaG 083 containing Zn-rich CAI was provided by the Institut für Planetologie of the University of Münster.The mineralogy and petrography of the CAI were studied using a JEOL 6610-LV scanning electron microscope (SEM) at the Interdisciplinary Center for Electron Microscopy and Microanalysisat the University of Münster.
Quantitative chemical compositions of the CAI minerals were obtained using a JEOL JXA 8530F electron microprobe operated at a 15 kV accelerating voltage and a beam current of 15 nA.For all elements, the measurement time was 15 s and 10 s on peak and background, respectively.Natural and synthetic reference materials were used for calibration.The matrix corrections were done according to the Uq(z) procedure (Armstrong 1991).
Oxygen isotopic analyses were performed at the Swedish Museum of Natural History (SMNH) in Stockholm and at the Hawai'i Institute of Geophysics and Planetology (HIGP), University of Hawai'i at Mānoa.At SMNH, the O-isotopic composition was obtained using a Cameca IMS-1280-HR ion probe.Analyses of 16 O − , 17 O − , and 18 O − were performed by utilizing a 10 keV Cs + primary ion beam with a beam current of 45 nA sputtering an analysis area of 6 × 6 μm 2 (20× 4 s).Presputtering of 8 × 8 μm 2 before each analysis was carried out to increase secondary ion yields.Negative oxygen ions were collected by Faraday cups (FCs) and corrected for internal mass fractionation by analyzing in-house San Carlos olivine, which was premeasured by laser fluorination mass spectrometry at the University of Göttingen.Baselines were determined separately with 400 s integration times.The mass resolving power (MRP) was set to ∼7000 for 17 O − and ∼2500 for 16 O − and 18 O − .
At HIGP, the O-isotope analyses were carried out with a Cameca IMS-1280 ion microprobe using the method described by Nagashima et al. (2015).A primary Cs + ion beam focused to 12 μm with ∼25 pA was used.The 16 O − was measured on a FC and 17 O − and 18 O − were simultaneously measured on the electron multipliers.Instrumental fractionation was corrected by using terrestrial standards: San Carlos olivine (olivine, melilite), augite (high-Ca pyroxene), and Burma spinel (spinel and grossite).The MRP was set to ∼2000 for 16 O − and 18 O − and ∼5500 for 17 O − , sufficient to separate interfering 16 OH − .
The Al-Mg isotope measurements were carried out at HIGP with the Cameca IMS-1280 ion microprobe, and the empirical fractionation factors (combining the instrumental and intrinsic fraction) were experimentally determined and used to calculate δ 26 Mg * using the setting and methods described in detail in MacPherson et al. (2020).
Rare-earth elements were measured by laser ablation inductively coupled mass spectrometry (LA-ICP-MS) at the Institute for Mineralogy, University of Münster.The LA-ICP-MS consists of an Analyte G2 Excimer laser ablation system and a Thermo Scientific ELEMENT 2TM ICP-MS.Aluminum was used as the internal standard (see Ebert &Bischoff 2016 andEbert et al. 2019 for more details).

Mineralogy and Petrography
The DaG 083 CAI studied is a mineralogically zoned object (Figures 1 and 2).Its well-rounded, very-fine-grained core is  dominated by Zn-hercynite (ZnO = 7.4 wt%; Table 1), which accounts for, at ∼110 μm in diameter, about 38 vol% of the whole object.Microprobe analyses of the core have low totals (∼92 wt%), possibly reflecting either its high porosity and/or the presence of unmeasured elements (e.g., H, S, Cl).The Zn-hercynite core is surrounded by layers of grossite and spinel 14-15 and 4-7 μm in thickness, respectively.A dark gap between the grossite layer and the Zn-rich hercynite interior is probably caused by plucking material during sample preparation.The grossite layers and inner part of the spinel layer contain abundant submicron inclusions of perovskite (Figure 2).In addition, grossite along thin cracks is corroded by a light gray Zn-bearing phase (indicated by arrows in Figure 2).This phase has lower Zn and Fe and higher Al content and total (∼96 wt%) than the CAI interior.The spinel layer is surrounded by a layer of variable thickness of Al-diopside with very minor melilite (high-Ca phase in Figure 1(c)).

Oxygen Isotopes
Oxygen isotopes were measured during two different sessions, at SWMN and HIGP.The results are listed in Table 2 and plotted in Figure 3.The Δ 17 O values for grossite measured in both sessions are similar.The CAI is isotopically heterogeneous.Zn-hercynite, grossite, and melilite have similar 16 O-poor compositions: Δ 17 O = −2.2± 0.6‰, −0.9 ± 2.1‰, and −2.6 ± 2.3‰, respectively.Aluminum-diopside has a solar-like Δ 17 O value of −22.6 ± 2.0‰, whereas spinel has an intermediate Δ 17 O value, of ∼−13.7 ± 2.1‰.The spinel measurements were only taken from the outer and perovskitefree part of the spinel layer to avoid a mixing of the two different phases (Figure 2(b)).

Magnesium Isotopes
Aluminum and magnesium isotopes were measured in grossite, melilite, spinel, and Al-diopside.The results are shown in Figure 4.The resolvable excess of 26 Mg, ∼55 ± 11‰, was found only in grossite that has a 27 Al/ 24 Mg Note.bdl = below detection limit; nd = not detected; Zn-hrc = Zn-hercynite; alt.region = altered regions, regarding to the clouds in Figure 2.  d * of 0.24 ± 0.46‰.

Rare-earth Elements
Rare-earth elements were measured in Zn-rich hercynite and grossite.The results are listed in Table 3; the CI-chondritenormalized values are plotted in Figure 5.Both minerals have similar REE patterns: Gd shows the highest enrichment (35× CI); Eu is strongly depleted (<0.1×CI); and Sm, Eu, Ho, and Er have a light depletion to their neighboring elements.

Primary and Secondary Minerals in DaG 083 CAI
The most remarkable mineralogical feature of the DaG 083 CAI is its multilayered rim sequence of grossite, nearly Fe-free spinel, and Al-diopside with minor melilite that surrounds the core, composed of nearly Mg-free Zn-hercynite.The rim sequence is generally consistent with a condensation origin (condensation of spinel prior to melilite, often observed in CAIs, could be attributed to kinetic reasons) from a cooling gas of solar composition (Yoneda & Grossman 1995;Ebel & Grossman 2000;Petaev & Wood 2005).Except for the strong Ce depletion (Figures 5(a), (b)), the REE pattern of the Zn-hercynite core and grossite shows similarities to a Group II pattern (Figure 5(b)), which can only be created during a condensation process.However, the strong depletion of Ce is a striking feature of HAL-like objects (e.g., Davis et al. 1982;Hinton et al. 1988;Ireland et al. 1988Ireland et al. , 1992;;Russell et al. 1998).Such a Ce depletion indicates incomplete condensation of the light rare-earth elements (LREEs) into the phase (Davis et al. 2018), and this can happen by condensation of perovskite in an oxidized nebula (Davis & Hinton 1986).Thereby, Ce becomes tetravalent and might be more easily implemented into perovskite, resulting in a positive Ce anomaly for perovskite and a depletion in the following phases.A condensation origin of the Zn-rich core and grossite layer is favored, but a later evaporation event under oxidizing conditions could also lead to a Ce depletion (Ireland 1988).
However, the Zn-rich hercynite core cannot be explained by either evaporation or condensation, because Zn is a moderately volatile element with a condensation temperature of ∼700 K, and Fe cannot condense into CAI-like minerals even under highly oxidizing conditions (Ebel & Grossman 2000).Therefore, Zn-hercynite is not in equilibrium with primary CAI minerals, grossite, spinel, melilite, and Al-diopside, which are predicted to condense above 1300 K (e.g., Petaev & Wood 2005); it must have a different origin.
Zn-bearing hercynite has been previously described in several weakly metamorphosed chondrites, Lance (CO3.5),Y-81020 (CO3.05),Colony (CO3.1), and Efremovka (CV3.1),where its formation was attributed to metasomatic reactions either in the solar nebula (Fahey et al. 1994;Russell et al. 1998) or on the chondrite parent asteroids (Brearley & Krot 2013;Krot et al. 2019b;Han et al. 2023).Our textural observations indicate that Zn-hercynite is a secondary mineral that most likely replaced primary Mg-free/-poor refractory mineral(s), e.g., corundum (Al 2 O 3 ), hibonite (CaAl 12 O 19 ), grossite, or krotite (CaAl 2 O 4 ).Because CO chondrites experienced metasomatic alteration in the presence of an aqueous solution Corundum and hibonite are resistant to metasomatic alteration, as indicated by (i) the presence of unaltered hibonite in association with Zn-hercynite in a CAI from the CO3.5 chondrite Lance (Fahey et al. 1994), and (ii) the common presence of micron-sized corundum grains in acid-resistant residues of the CO, CV, and ordinary chondrites metasomatically altered to various degrees (e.g., Makide et al. 2013;Needham et al. 2017).Therefore, Reactions (R1) and (R2) seem unlikely.Although replacement of grossite by Zn-hercynite has been described in grossite-bearing CAIs from Efremovka (Han et al. 2023) and partial replacement of the grossite layer by Zn,Fe,Ca,Al-oxide is observed in the DaG 083 CAI (Figure 2), Reaction (R3) is also unlikely, because the grossite layer largely avoided alteration, whereas the Zn-hercynite core contains no relict primary phase.Replacement of krotite by Zn-hercynite (Reaction (R4)) was described in a grossite-krotite-bearing CAI from the CO3.05 chondrite Y-81020 (Figure 12 in Krot et al. 2019b).We suggest that this reaction most likely explains the origin of Zn-hercynite in DaG 083 CAI.However, it is not clear if the inferred metasomatic alteration took place in situ on the DaG 083 parent body or whether it occurred in a preexisting planetesimal, and therefore we cannot completely exclude a different precursor phase for the Zn-hercynite.3).CAIs are thought to have formed in a localized disk region close to the proto-Sun in a gas of approximately solar composition, subsequently transported outward, and accreted in different planetesimals throughout the disk (Shu et al. 1996;McKeegan et al. 2000;Brownlee et al. 2006;Krot et al. 2009;Ciesla 2010;Wielandt et al. 2012;Yang & Ciesla 2012;Desch et al. 2018;Dunham et al. 2022).The 16 O-rich, solar-like composition of Al-diopside is consistent with this scenario and supports a common formation region of refractory inclusions from different chondrite groups (e.g., McKeegan et al. 1998;Rout et al. 2009;Ebert et al. 2018;Shollenberger et al. 2018;Krot 2019;Render et al. 2019;Ebert et al. 2020;Render et al. 2022).Yet, melilite, grossite, and Zn-rich hercynite are depleted in 16 O.Similar O-isotope heterogeneity was observed in a Zn-hercynite-bearing grossite-krotite-rich inclusion from Y-81020 (CO3.05;Krot et al. 2019b).Such an O-isotopic heterogeneity within refractory objects is common in refractory inclusions from CO, CV, ordinary, and Rumuruti chondrites of petrological types 3.05, where melilite, anorthite, perovskite, Zr-and Sc-rich oxides and silicates are systematically 16 O-depleted relative to corundum, hibonite, spinel, Al,Tidiopside, and forsterite, all of which have solar-like Δ 17 O (Clayton et al. 1977;Wasson et al. 2001;Rout et al. 2009;Simon et al. 2011;Krot et al. 2019aKrot et al. , 2019b;;Simon et al. 2019;Ebert et al. 2020;Krot et al. 2021Krot et al. , 2022b)).In contrast, the vast majority of refractory inclusions in unmetamorphosed carbonaceous chondrites (petrologic types 3.0) have uniform, typically solar-like O-isotopic compositions (McKeegan et al. 1998;Itoh et al. 2007;Makide et al. 2009;Bodénan et al. 2014;Krot et al. 2017aKrot et al. , 2019b;;Ushikubo et al. 2017).Based on these observations, it is suggested that O-isotope exchange during fluid-rock interaction played an important role in producing O-isotope heterogeneity within refractory objects (Krot 2019).Within the Zn-rich CAI, the phases typically affected by parent-body alteration are enriched in 18 O and 17 O.Therefore, the 16 O-depletion for melilite, grossite, and Zn-rich hercynite within the Zn-rich CAI is not a primary signature of the CAI formation region but rather was generated by fluid-assisted parent-body alteration.This alteration process pushed the Δ  However, fluid-assisted thermal metamorphism cannot be the reason for the intermediate Δ 17 O composition of spinel, because oxygen self-diffusion in spinel is very slow (Ryerson & McKeegan 1994) and this mineral retained its initial O-isotope compositions even in heavily altered meteorites, like the Allende (CV > 3.6) and CK3.7-3.8 chondrites (Krot et al. 2021;MacPherson et al. 2022;Krot et al. 2023), and in relict grains in chondrules (Krot et al. 2017b), including the Al-rich chondrule from DaG 083 itself (Ebert et al. 2022).As the spinel measurements are not (or only very minorly) affected by perovskite (Figure 2(b)), the O-isotope composition of spinel can be seen as its pristine composition.
We conclude that the spinel and Al-diopside layers of DaG 083 CAI sampled nebular reservoirs with different Δ 17 O, ∼−14‰ and ∼−22‰, providing direct evidence for variations in the O-isotope compositions of nebular gas in the CAIforming region.Subsequently, the CAI experienced metasomatic alteration in the presence of an aqueous solution that resulted in formation of Zn-hercynite and O-isotope exchange in grossite and melilite.; the inferred ( 26 Al/ 27 Al) 0 is (2.6 ± 0.3) × 10 −6 , which is much lower than the canonical value (Figure 4).The magnesium isotopes within the grossite of DaG 083 (CO3.1) are most likely undisturbed during fluid-assisted thermal metamorphism, in contrast to the oxygen isotopes.Simon et al. (2019) showed that grossite avoided redistribution of 26 Mg and the ( 26 Al/ 27 Al) 0 in CAIs within DOM 08004 (CO3.1) are indistinguishable from those in the DOM 08006 (CO3.0)CAIs.Therefore, if the Mg-isotopes are unaffected by secondary processes, two possible scenarios are possible: (i) 26 Al was initially uniformly distributed in the CAI-forming region at the canonical level, (5.25 ± 0.12) × 10 −5 , and the DaG 083 CAI formed ∼2.9 Ma after t 0 ; (ii) 26 Al was heterogeneously distributed in the CAI-forming region, and the inferred ( 26 Al/ 27 Al) 0 provides no chronological information.
It is commonly assumed that formation of 26 Al-poor CAIs predates formation of CAIs with the canonical 26 Al/ 27 Al ratio (e.g., Krot et al. 2012 and references therein).Krot et al. (2020) suggested a variable O-isotope composition of the nebular gas in the CAI-forming region before the injection of 26 Al, based on grossite-bearing CAIs that have lower 26 Al/ 27 Al values (<10 −6 ) in combination with an homogeneous O-isotopic composition of individual objects but large inter-CAI variations of Δ 17 O (−40‰ to −5‰).A similar conclusion was reached by Kööp et al. (2016a), who observed that PLAC-like inclusions with Δ 17 O ∼−25‰ have the lowest 48 Ca and 50 Ti isotopic variations, whereas the isotopic variations increase with increase of Δ 17 O (−25‰ to −17‰).The investigated PLAC-like inclusions all have low or undetectable 26 Mg * excesses.Thus, the authors assumed that the solar nebula was isotopically heterogeneous shortly after the collapse of the protosolar cloud and before 26 Al was injected.The variations in these inclusions' O-isotopic compositions originated from the presence of 16 O-poor (Δ 17 O −17‰) primordial dust relative to the primordial gaseous reservoir (CO + H 2 O; Δ 17 O < −35‰).This implies that the PLAC-like inclusions were formed before the majority of CAIs condensed out of the solar nebula, making them somewhat older (Kööp et al. 2016a).In contrast, SHIBs having solar-like O-isotopic compositions and ∼canonical ( 26 Al/ 27 Al) 0 show, in general, an homogeneous isotopic composition regarding Ti and Ca (Kööp et al. 2016b).Therefore, the Ti and Ca isotope anomalies observed in PLAClike inclusions were diluted and homogenized in SHIBs to similar values as known from CAIs, hinting at homogenization of an initially heterogeneous solar nebula.
Thus, there is much evidence that the CAI formation region was isotopically more heterogeneous at the beginning, before 26 Al was injected and, therefore, before the majority of normal CAIs were formed.During this time, temporal variations of Δ 17 O are possible (Lyons & Young 2005), but the mineralogy of the DaG 083 CAI follows with grossite → spinel → melilite → Al-diopside, nearly the predicted condensation sequence from a nebula with solar-like composition (Ebel & Grossman 2000), and a large time gap between the condensation of spinel and Ca pyroxene seems unlikely.Thus, this CAI had to be transported, or the formation region had to change, from a 16 O-depleted to a 16 O-enriched region between the crystallization of spinel and Al-diopside.
We conclude that the DaG 083 CAI formed by evaporation and condensation processes in the CAI-forming region and subsequently experienced metasomatic alteration in the presence of aqueous solutions on the CO chondrite parent asteroid.The spinel and diopside recorded variations in O-isotope composition of nebular gas in the CAI-forming region, most likely prior to injection and homogenization of 26 Al in this region.The melilite and grossite experienced O-isotope exchange with the aqueous solution and Zn-hercynite formed as a result of replacement of the preexisting refractory phase, possibly krotite.

Figure 2 .
Figure 2. Backscattered electron image of the multilayered rim around DaG 083 CAI.(a) Light gray alteration regions in grossite, indicated by arrows, follow tiny cracks.These regions are enriched in Zn, but not as much as the interior of the CAI.Tiny white inclusions in the spinel and grossite are perovskite grains.(b) An arrow marks the in situ measurement hole for O-isotopes from spinel.Only the outer and perovskite-free part of the spinel rim were measured to avoid a mixing of the two phases.(c) A close-up of the melilite and cpx on the outer part of inclusion.

Figure 5 .
Figure 5. CI-normalized REE abundances in the DaG 083 CAI (this study) and (a) HAL-type inclusions (Hinton et al. 1988) and (b) average Group II pattern from various Allende CAIs (Stracke et al. 2012 and references therein).The arrows mark values below the detection limit.CI values from Lodders et al. (2009).

Table 1
Chemical Compositions (in wt%) of Mineral Phases in the DaG 083 CAI

Table 2
Oxygen Isotopic Composition (in ‰) of Different Mineral Phases Studied a Measured at the Swedish Museum of Natural History in Stockholm.Znhrc = Zn-hercynite.ratio of ∼2900.Note that no proper Al-Mg standard exists for grossite.Therefore, proper correction for an Al/Mg relative sensitivity factor of grossite cannot be made.The internal isochron has a ( 26 Al/ 27 Al) 0 of (2.60 ± 0.29) × 10 −6 with an initial Mg 26 0