# Metasomatic Alteration of Type 3 Ordinary and Carbonaceous Chondrites

A N Krot, M I Petaev, L. Piani, Y. Marrocchi, W. Fujiya, O V Pravdivtseva, E. Dobrică, L G Vacher, A J King, M. Lee, et al.

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A N Krot, M I Petaev, L. Piani, Y. Marrocchi, W. Fujiya, et al., Metasomatic Alteration of Type 3 Ordinary and Carbonaceous Chondrites. Space Science Reviews, 2025, 221 (1), pp.7. 10.1007/s11214-025-01135-z. hal-04960286

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# Metasomatic Alteration of Type 3 Ordinary and Carbonaceous Chondrites

A.N. Krot¹ ID · M.I. Petaev² · L. Piani³ · Y. Marrocchi³ · W. Fujiya⁴ · O.V. Pravdivtseva⁵ · E. Dobrică¹ · L.G. Vacher⁶ · A.J. King⁷ · M. Lee⁸ · E. Van Kooten⁹ · B. Jacobsen¹⁰ · C.M.O'D. Alexander¹¹ · A. Bischoff¹² · A.J. Brearley¹³ · C. Le Guillou¹⁴ · L. Remusat¹⁵ · J. Leitner¹⁶,¹⁷ · G.R. Huss¹

Received: 2 April 2024 / Accepted: 28 December 2024
© The Author(s) 2025

## Abstract

Metasomatism refers to the process during which a pre-existing rock undergoes compositional and mineralogical transformations associated with chemical reactions triggered by the reaction of fluids which invade the protolith. It changes chemical compositions of minerals, promotes their dissolution and precipitation of new minerals. In this paper, we review metasomatic alteration of type 3 ordinary (H, L, LL) and carbonaceous (CV, CO, CK) chondrites, including (i) secondary mineralization, (ii) physicochemical conditions, (iii) chronology ($^{53}$Mn-$^{53}$Cr, $^{26}$Al-$^{26}$Mg, $^{129}$I-$^{129}$Xe) of metasomatic alteration, (iv) records of metasomatic alteration in H, O, N, C, S, and Cl isotopic systematics, (v) effects of metasomatic alteration on O- and Al-Mg-isotope systematics of primary minerals in chondrules and refractory inclusions, and (vi) sources of water ices in metasomatically altered CV, CO, and ordinary chondrites, and outline future studies.

**Keywords** Carbonaceous chondrites · Ordinary chondrites · Metasomatic alteration · Water · Oxygen · Hydrogen · Nitrogen and sulfur isotopes · $^{53}$Mn-$^{53}$Cr chronology · $^{129}$I-$^{129}$Xe chronology

## 1 Introduction

There are several classes of chondritic meteorites comprising a number of groups: carbonaceous (CI, CM, CR, CO, CV, CK, CR, CH, CB, CY, CL), ordinary (H, L, LL), enstatite (EH, EL), Kakangari (K), and Rumuruti (R) (Krot et al. 2014a). Some chondrites cannot be classified into the existing groups and are named ungrouped or unique, e.g., Acfer 094 (Newton et al. 1995). CV chondrites are subdivided into three subgroups, oxidized Allende-like (CV<sub>oxA</sub>), oxidized Bali-like (CV<sub>oxB</sub>), and reduced (CV<sub>red</sub>), which may or may not represent fragments of the same parent asteroid (Weisberg et al. 2006; Gattacceca et al. 2020). Virtually all chondrites experienced aqueous alteration and/or thermal metamorphism in their parent asteroids. Based on the degree of aqueous alteration and thermal metamorphism, chondrites are classified into petrologic types from 1 to <3 and from 3 to 6, respectively. Based on the distribution of chromium in chondrule ferroan olivines (Grossman and Brearley 2005), Raman parameters of matrix organic matter (Bonal et al. 2006, 2007), sulfur
abundance in matrix (Grossman and Brearley 2005), composition of FeNi-metal (Kimura et al. 2007), matrix modal mineralogy (Brearley 1993; Greshake 1997), and abundance of presolar grains and noble gas components (Huss and Lewis 1995; Nguyen et al. 2010; Leitner et al. 2016; Nittler et al. 2018), type 3 chondrites are further subdivided into subtypes from 3.0 to 3.9; even more detailed classification is used for type 3.0–3.1 chondrites. We note that this classification scheme is rather simplistic, because aqueous alteration and thermal metamorphism are closely related processes and virtually all thermally metamorphosed type 3 chondrites experienced aqueous alteration (Bonal et al. 2020). The least altered and the least thermally metamorphosed chondrites are type 3.00. There are only a few type 3.00 chondrites known, including Semarkona (LL), Northwest Africa (NWA) 8276 (L), Dominion Range (DOM) 08006 (CO), and Acfer 094 (C-ungrouped) (Fig. SM1; Newton et al. 1995; Grossman and Brearley 2005; Alexander et al. 2018a; Davidson et al. 2019; Zanetta et al. 2022; Krämer Ruggiu et al. 2022; Barosch et al. 2022).

Here we review metasomatic alteration of type 3 ordinary (H, L, LL) and carbonaceous (CV, CO, CK) chondrites, including (i) secondary mineralization, (ii) physicochemical conditions and (iii) chronology ($^{53}$Mn-$^{53}$Cr, $^{26}$Al-$^{26}$Mg, $^{129}$I-$^{129}$Xe) of metasomatic alteration; (iv) records of metasomatic alteration in H, O, N, C, S, and Cl isotope systematics; (v) effects of metasomatic alteration on O- and Al-Mg-isotope systematics of primary minerals in chondrules and refractory inclusions; (vi) sources of water ices in metasomatically altered CV, CO, and ordinary chondrites, and outline future studies.

# 2 Primary and Secondary Mineralization in Type 3 Chondrites

## 2.1 Primary Mineralogy of Chondritic Components in Primitive Chondrites

Chondrites are composed of several major components, including refractory inclusions [Ca,Al-rich inclusions (CAIs) and amoeboid olivine aggregates (AOAs) – aggregates of CAIs and fortsterite+FeNi-metal condensates], chondrules, matrix, FeNi-metal, Fe$\pm$Ni-sulfides, and organic matter. Refractory inclusions consist of minerals that formed by gas-solid condensation above $\sim$ 1300 K or by melting of gas-solid condensates in a gas of approximately solar chemical composition, i.e., highly reduced: H$_2$O/H$_2$ ratio of $\sim$6.6$\times$10$^{-4}$ corresponding to $\sim$ 7 log units below the iron-wüstite (IW) buffer (e.g., Krot 2019). As a result, primary oxide and silicate minerals in CAIs and AOAs from type 3.00 chondrites are Fe- and Na-free (Table SM1). Pristine refractory inclusions contain no oxidized and low-temperature secondary minerals, such as magnetite, nepheline, sodalite, etc.

Pristine chondrules consist of minerals (Table SM2) formed by melting of solid precursors during transient heating events at $\sim$ 1600–2000 K in a gas much more oxidized than solar, IW–3 to IW–1, and contain no secondary minerals such as magnetite, phyllosilicates, nepheline, etc. Chondrules experienced gas-melt interaction before and during crystallization resulting in re-condensation and zoned distribution of moderately volatile elements, such as Si, Mn, Cr, Na, and K (Libourel et al. 2006). Because these elements are also mobile during aqueous/metasomatic alteration (Grossman and Brearley 2005), this may complicate distinguishing nebular and asteroidal records in chondrules.

Pristine matrices are composed predominantly of amorphous ferromagnesian silicates, sub-micron forsterite and enstatite grains, FeNi-metal, Fe-sulfides, organics, presolar grains, and have abundant porosity (Table SM2; Greshake 1997; Grossman and Brearley 2005; Davidson et al. 2019). Because crystalline silicates are common in pristine matrices but virtually absent in interstellar medium (ISM), some fraction of matrix materials must have
experienced thermal processing in the protoplanetary disk (PPD), most likely during formation of refractory inclusions and chondrules. We note that thermal processing of amorphous silicates in the PPD (i) below 1100 K that resulted in formation of crystalline silicates and/or (ii) below 900 K that may have modified their primordial oxygen-isotope composition without formation of crystalline silicates are also possible (Ishizaki et al. 2023; Sakurai et al. 2023; Yamamoto et al. 2024). The abundance of primordial, molecular cloud material in pristine chondrite matrices that avoided thermal processing in the PPD remains controversial (Alexander et al. 2017). Due to its extremely fine-grained nature and high porosity, matrix can be easily altered; e.g., amorphous material even in the most primitive carbonaceous chondrite Acf er 094 experienced hydration (Matsumoto et al. 2019).

Refractory inclusions, chondrules, and matrices having different mineralogies, textures, chemistry, and grain-sizes responded differently to metasomatic alteration.

## 2.2 Initial Stages of Alteration of Chondritic Components in Type 3.0–3.1 Chondrites

The initial stages of metasomatic alteration are largely recorded by matrices and chondrules in ordinary, CV, and CO chondrites of petrologic types 3.0–3.1. The petrographic and chemical effects of metasomatic alteration include (i) dissolution of chondrule mesostasis and plagioclase, (ii) zoning of Na and Ca in chondrule mesostasis and plagioclase, (iii) replacement of chondrule mesostasis by phyllosilicates, and (iv) redistribution of matrix nanosulfides (Grossman et al. 2000; Grossman and Brearley 2005; Lewis and Jones 2019). Typical secondary minerals formed during the initial stages of metasomatic alteration of type 3 chondrites are listed in Table SM3; their textural occurrences are shown in Figures SM2–SM7.

FeNi-metal experienced carburization to form FeNi-carbides (Figs. SM2a,b; Hutchison et al. 1987; Alexander et al. 1989; Taylor et al. 1981; Shibita 1996; Krot et al. 1997; Keller 1998; Krot and Todd 1998; Davidson et al. 2019):

$$ (R1) \quad CO(g) + H_2O(g) \rightarrow CO_2(g) + H_2(g) $$

$$ (R2) \quad 3Fe(s) + CO_2(g) + 2H_2(g) \rightarrow Fe_3C(s) + 2H_2O(g) $$

FeNi-metal and carbides were subsequently oxidized to form magnetite (Fig. SM2):

$$ (R3) \quad Fe_3C(s) + 6H_2O(g) \rightarrow Fe_3O_4(s) + 6H_2(g) + CO_2(g) $$

$$ (R4) \quad 3Fe(s) + 4H_2O(g) \rightarrow Fe_3O_4(s) + 4H_2(g) $$

Note that reactions R1–4 do not require aqueous fluid and may have occurred either in the solar nebula or the chondrite parent bodies. For example, FeNi-carbides described in anhydrous interplanetary dust particles (IDPs) could be by-products of Fischer-Tropsch synthesis in the solar nebula (Christoffersen and Buseck 1983; Bradley et al. 1984; Bradley 1994). Oxidation of metal to magnetite is thermodynamically possible in the canonical solar nebula below < 400 K (Fegley and Prinn 1989) and was experimentally studied by Hong and Fegley (1998) (see Sect. 5.1). There are, however, mineralogical evidence suggesting formation of FeNi-carbides and magnetite in the chondrite parent asteroids (Krot et al. 1997). (i) FeNi-carbides replace metal nodules in chondrules and form veins crosscutting their fine-grained rims (Figs. SM2c–f). (ii) Two types of magnetite can be distinguished: (a) “dirty”, Cr-bearing magnetite that formed by *in situ* oxidation of Cr-bearing FeNi-metal nodules (Figs. SM2c, 4b):

$$ (R5) \quad 3Fe(s) + 4H_2O(aq, g) \rightarrow Fe_3O_4(s) + 4H_2(g) $$
and (b) “clean”, Cr-free magnetite overgrowing oxidized nodules (Figs. SM2c, 4b) and crosscutting fine-grained rims around chondrules (Figs. SM2c–f, 5b,e). This magnetite formed by precipitation of Fe dissolved in aqueous solution:

$$ (R6) \quad 3\text{Fe}^{2+}_{(\text{aq})} + 4\text{H}_2\text{O}_{(\text{aq})} \rightarrow \text{Fe}_3\text{O}_{4(\text{s})} + 6\text{H}^{+}_{(\text{aq})} + \text{H}_2(\text{g}), $$

Marrocchi et al. (2016) interpreted clean Cr-free magnetite overgrowing opaque nodules in chondrules from the oxidized Bali-like CV3.1 chondrite Kaba as igneous, i.e., as a result of crystallization FeSO melts. Formation of FeSO melt requires drastic oxidizing conditions, too oxidizing even for chondrule formation, IW+1 to IW+2 (see Sect. 2.1). The lack of magnetite-bearing chondrules in the reduced CV3 chondrites and the differences in O-isotope compositions between Cr-free magnetite and chondrule phenocrysts in Kaba chondrules (Marrocchi et al. 2016; Krot et al. 2022a) contradict this interpretation (see Sect. 7.1.2).

In addition, framboidal magnetite occasionally observed in matrix of the oxidized CV3.1 chondrites Kaba and Mokoia also formed by direct precipitation from aqueous fluid (Keller and Buseck 1990; Tomeoka and Buseck 1990). The existence of aqueous solution during alteration of type 3.0–3.1 chondrites is supported by hydration and replacement of amorphous silicates by phyllosilicates and ferroan olivine (Fig. SM3).

Subsequently, the Cr-free magnetite is preferentially replaced by Mn-bearing fayalite (Fig. SM4):

$$ (R7) \quad 2\text{Fe}_3\text{O}_{4(\text{s})} + 3\text{SiO}_{2(\text{aq})} + 3\text{xMn}^{2+}_{(\text{aq})} + 2\text{H}_2(\text{g}) \rightarrow 3(\text{Fe}_{2-x}\text{Mn}_x)_2\text{SiO}_{4(\text{s})} + 3\text{xFe}^{+}_{(\text{aq})} + 2\text{H}_2\text{O}_{(\text{aq})} $$

Isolated fayalite grains in matrices and pore space of AOAs and fine-grained CAIs, as well as fayalite grains overgrowing chondrule olivines and fine-grained matrix-like rims (Figs. SM5–6) apparently formed by direct precipitation from aqueous solutions:

$$ (R8) \quad x\text{Fe}^{2+}_{(\text{aq})} + (2-x)\text{Mn}^{2+}_{(\text{aq})} + \text{SiO}_{2(\text{aq})} + 2\text{H}_2\text{O}_{(\text{aq})} \rightarrow (\text{Fe,Mn})_2\text{SiO}_{4(\text{s})} + 4\text{H}^{+}_{(\text{aq})} $$

In Kaba and MacAlpine Hills (MAC) 88107 (CO3.1-like), fayalite often associates with hedenbergite (Fig. SM5) that may have formed by direct precipitation from aqueous fluid and by replacement of fayalite:

$$ (R9) \quad \text{Fe}^{2+}_{(\text{aq})} + \text{Ca}^{2+}_{(\text{aq})} + 2\text{SiO}_{2(\text{aq})} + 2\text{H}_2\text{O}_{(\text{aq})} \rightarrow \text{CaFeSi}_2\text{O}_{6(\text{s})} + 4\text{H}^{+}_{(\text{aq})} $$

$$ (R10) \quad \text{Fe}_2\text{SiO}_{4(\text{s})} + \text{SiO}_{2(\text{aq})} + \text{Ca}^{2+}_{(\text{aq})} \rightarrow \text{CaFeSi}_2\text{O}_{6(\text{s})} + \text{Fe}^{2+}_{(\text{aq})} $$

Manganese-bearing fayalite is critical for ⁵³Mn-⁵³Cr dating of metasomatic alteration of UOCs, COs, and CVs (see Sect. 6.1). It is also important for constraining O-isotope composition of aqueous fluid, and, potentially, temperature of metasomatic alteration (see Sect. 7.1). We note, however, that reactive relationships between fayalite and magnetite (Fig. SM4) suggest these minerals may not have been in isotopic equilibrium with the same fluid. Therefore, temperature estimates based on O-isotope compositions of fayalite and magnetite must be taken with caution.

In the reduced CV3s (Efremovka, Leoville, Vigarano) and MAC 88107, magnetite, fayalite, and hedenbergite associate with kirschsteinite (Fig. SM7) (Krot et al. 2000; Ganino and Libourel 2017; MacPherson et al. 2017) that may have formed by replacement of magnetite and by direct precipitation from aqueous solution:

$$ (R11) \quad \text{Fe}_3\text{O}_{4(\text{s})} + \text{Ca}_{(\text{aq})}^{2+} + \text{SiO}_{2(\text{aq})} + \text{H}_2(\text{g}) + 2\text{H}^{+}_{(\text{aq})} \rightarrow \text{CaFeSiO}_{4(\text{s})} + 2\text{Fe}^{2+}_{(\text{aq})} + 2\text{H}_2\text{O}_{(\text{aq})} $$

$$ (R12) \quad \text{Ca}^{2+}_{(\text{aq})} + \text{Fe}^{2+}_{(\text{aq})} + \text{SiO}_{2(\text{aq})} + 2\text{H}_2\text{O}_{(\text{aq})} \rightarrow \text{CaFeSiO}_{4(\text{s})} + 4\text{H}^{+}_{(\text{aq})} $$

Kirschsteinite (kirschsteinite-monticellite solid solution) grains in CVs are often chemically-zoned (Fig. SM7b) suggesting evolution of chemical composition of aqueous fluid:
$$ (R13) \quad \mathrm{CaFeSiO}_{4(\mathrm{s})} + x\mathrm{Mg}^{2+}_{(\mathrm{aq})} \rightarrow \mathrm{Ca}(\mathrm{Fe}_{1-x}\mathrm{Mg}_x)\mathrm{SiO}_{4(\mathrm{s})} + x\mathrm{Fe}^{2+}_{(\mathrm{aq})} $$

Kirschsteinite is a critical mineral for ⁵³Mn-⁵³Cr chronology of metasomatic alteration of the reduced CVs (see Sect. 6.2), and potentially important for constraining O-isotope composition of aqueous fluid in their parent body. However, no SIMS (secondary ion mass-spectrometry) standards for kirschsteinite are currently available.

In contrast to matrices and chondrules, most primary minerals in refractory inclusions from 3.0–3.1 chondrites show little evidence for metasomatic alteration. The only exception are Ca-aluminates, grossite and krotite, which can be replaced by hercynite that is commonly Zn-bearing (Fig. SM8; Krot et al. 2019b; Ebert et al. 2024; Han et al. 2024):

$$ (R14) \quad \mathrm{CaAl}_{2}\mathrm{O}_{4(\mathrm{s})} + x\mathrm{Fe}^{2+}_{(\mathrm{aq})} + (1-x)\mathrm{Zn}^{2+}_{(\mathrm{aq})} \rightarrow (\mathrm{Fe,Zn})\mathrm{Al}_{2}\mathrm{O}_{4(\mathrm{s})} + \mathrm{Ca}^{2+}_{(\mathrm{aq})} $$

$$ (R15) \quad \mathrm{CaAl}_{4}\mathrm{O}_{7(\mathrm{s})} + 2x\mathrm{Fe}^{2+}_{(\mathrm{aq})} + (2-2x)\mathrm{Zn}^{2+}_{(\mathrm{aq})} + \mathrm{H}_{2}\mathrm{O}_{(\mathrm{aq})} \rightarrow 2(\mathrm{Fe}_{x}\mathrm{Zn}_{1-x})\mathrm{Al}_{2}\mathrm{O}_{4(\mathrm{s})} + \mathrm{Ca}^{2+}_{(\mathrm{aq})} + 2\mathrm{H}^{+}_{(\mathrm{aq})} $$

## 2.3 Advanced Stages of Alteration in Type 3 Chondrites of Higher Petrologic Subtypes

Studies of secondary mineralization in type 3 carbonaceous chondrites as a function of their petrologic subtypes are largely focused on the formation of nepheline in refractory inclusions and chondrules in CO3.0–3.8 s, where it formed by replacement of melilite, plagioclase, and chondrule mesostasis (Kojima et al. 1995; Tomeoka and Itoh 2004; Itoh et al. 2004; Brearley and Krot 2012; Wick and Jones 2012). There is a correlation between the amounts of nepheline in the CO chondritic components and the degree of thermal metamorphism of their host meteorites (Kojima et al. 1995; Itoh and Tomeoka 1998; Russell et al. 1998; Tomeoka and Itoh 2004), suggesting the nephelinization resulted from fluid-assisted thermal metamorphism. Mineralogical changes during metasomatic alteration of CV3 and CK3 chondrites are poorly known, because (i) well established metamorphic sequence among CV3s is currently absent (Bonal et al. 2006) and (ii) CKs of petrologic type <3.7 are not known.

Bridges et al. (1997) and Lewis et al. (2022) showed that evidence of metasomatism is present throughout the metamorphic sequence in OCs, and thus was a ubiquitous process in the OC parent bodies. In types 3.0–3.8, chondrule glassy mesostasis experienced dissolution and replacement by phyllosilicates, which survived only in type 3.0–3.1. Oxidation of phosphorous exsolved from FeNi-metal resulted in formation of merrillite and apatite; apatite typically replaces merrillite. Primary calcic plagioclase in chondrules was altered to various degrees by sodalite, scapolite, and nepheline in petrologic types 3.2–3.9, and to albite in types 3.6–5 (see also Bischoff and Keil 1983, 1984, 2012). Plagioclase also developed alteration features such as Ca-Na zoning (Grossman et al. 2000), micropores, and alteration lamellae in types 3–4. Sodic plagioclase is present in minor amounts as a primary phase, but also formed from the crystallization of chondrule mesostasis glass (types 3.2–3.9), and predominantly through albitization reactions in calcic plagioclase (types 3.6–5). K-feldspar occurs in albite in types 3.6–6 as fine-scale exsolution lamellae and as larger patches. Not all K-feldspar in OCs is secondary: e.g., coarse-grained K-feldspar (Or95) occurs in achondritic clasts of unknown origin in Adzhi-Bogdo (LL3–6) breccia (Bischoff et al. 1993). In type 4–6 Ls and LLs, mesostasis glass crystallized into secondary plagioclase of diverse composition (An<85). Equilibration of plagioclase is observed through type 5 and by type 6, plagioclase is chemically equilibrated at ~An11. Apatite evolves to Cl-apatite with a minor F and no OH. There are, however, exceptions: e.g., apatite in Renchen (L5–6) contains 3.4 wt% F and only 0.47 wt% Cl (Bischoff et al. 2019). Based on these observations, Lewis
et al. (2022) concluded that aqueous alteration during prograde metamorphism of OCs produced most of the alteration and equilibration features in their plagioclase. During retrograde metamorphism, high temperature, short duration infiltration of anhydrous, alkali- and halogen-bearing fluids caused incorporation of potassium into plagioclase that subsequently exsolved.

## 2.4 Metasomatic Alteration of Igneous CAIs in Allende (CV<sub>oxA</sub> > 3.6)

Coarse-grained, cm-sized igneous CAIs are present exclusively in CV and CK chondrites. Several types of igneous CAIs are recognized: Compact Type A (CTA), Type B (B), Forsterite-bearing Type B (FoB), and Type C (C) (Fig. SM9). These types consist of similar primary minerals, but have different modal mineralogy and bulk chemical compositions (MacPherson 2014). Because of their large sizes, the common presence in a widely available CV3 chondrite fall Allende (CV<sub>oxA</sub>>3.6), and the importance for studying the early solar system chronology, the secondary mineralogy mainly of Allende igneous CAIs is characterized in detail (e.g., Krot et al. 2021).

Most of the secondary minerals in Allende igneous CAIs are Mg-rich and anhydrous (Table SM4). As a result, both nebular and asteroidal models have been proposed to explain their formation. According to the nebular models (Hashimoto and Grossman 1987; McGuire and Hashimoto 1989; Keller and Buseck 1991; Ushikubo et al. 2007a,b), secondary minerals formed during gas-solid reactions in the solar nebula. According to the asteroidal models (Brearley and Krot 2012; Krot et al. 2021), these minerals formed during aqueous fluid – rock interaction in the CV parent asteroid. Both nebular and asteroidal processes have been invoked to explain alteration of CV CAIs by Fagan et al. (2007).

Secondary minerals providing clear evidence for involvement of the CO₂- and H₂O-bearing fluid in their formation include clintonite, montmorillonite, margarite, biopyriboles, calcite, and tilleyite (Table SM4; Fig. SM10). Involvement of aqueous solutions is also indicated by the presence of veins of secondary minerals crosscutting primary anorthite and melilite, and by dissolution textures of anorthite and melilite with open space filled by secondary minerals (Fig. SM11).

Allende igneous CAIs are surrounded by Wark-Lovering (WL) rims that originated in the CAI-forming region (Wark and Lovering 1976), which are in turn surrounded by fine-grained matrix-like rims composed of secondary ferroan olivine and nepheline (Fig. SM9e), and by a layer of secondary CaFe-rich silicates (salite-hedenbergite pyroxenes, andradite, and wollastonite (Fig. SM9f). The altered regions of primary melilite are depleted in Ca and enriched in Si. There is a zoned distribution of secondary Na- and Cl-bearing minerals in the Allende CAIs: wadalite/adrianite is found almost exclusively in âkermanite-rich cores of Type B CAIs, whereas sodalite and nepheline are concentrated in their gehlenitic peripheries (Figs. SM9d,e). These observations provide clear evidence for an open-system behavior of several elements (Si, Na, Cl, Ca, and Al) during *in situ* metasomatic alteration of these CAIs: Si, Na, and Cl were added, whereas Ca and some Al were removed from the inclusions. The mobility of Al during the alteration is consistent with the presence of grossular, kushiroite, and Na-rich melilite in veins crosscutting primary anorthite grains inside the inclusions (Fig. SM11c). The outward transport of dissolved Al from CAIs and the influx of Na, K, and Si with an aqueous fluid equilibrated with the Allende matrix could explain the abundant nepheline grains in fine-grained rims around altered Allende CAIs (Fig. SM9e):

$$ (R16) \quad \mathrm{Na}^{+}_{(\mathrm{aq})} + \mathrm{Al}^{3+}_{(\mathrm{aq})} + \mathrm{SiO}_{2}(\mathrm{aq}) + 2\mathrm{H}_{2}\mathrm{O}_{(\mathrm{aq})} \rightarrow \mathrm{NaAlSiO}_{4}(\mathrm{s}) + 4\mathrm{H}^{+}_{(\mathrm{aq})} $$

The zoned distribution of nepheline and sodalite in Allende CAIs and the lack of sodalite in fine-grained rims around these CAIs may indicate an earlier formation of nepheline than
sodalite, consistent with observations of Che and Brearley (2021). Calcium released during alteration of melilite interacted with Fe-, Mg- and Si-enriched fluid that most likely was in equilibrium with the Allende fine-grained matrix (Krot et al. 2001; Ganino and Libourel 2017). This process resulted in precipitation of CaFe-rich silicates (hedenbergite, andradite, wollastonite, and salite-hedenbergite pyroxene) at the boundary between fine-grained rims and the Allende matrix:

$$ (R17) \quad Ca^{2+}_{(aq)} + Fe^{2+}_{(aq)} + 2SiO_{2}(aq) + 2H_{2}O_{(aq)} \rightarrow CaFeSi_{2}O_{6(s)} + 4H^{+}_{(aq)} $$

$$ (R18) \quad 3Ca^{2+}_{(aq)} + 2Fe^{3+}_{(aq)} + 3SiO_{2}(aq) + 6H_{2}O_{(aq)} \rightarrow Ca_{3}Fe_{2}Si_{3}O_{12(s)} + 12H^{+}_{(aq)} $$

$$ (R19) \quad Ca^{2+}_{(aq)} + SiO_{2}(aq) + H_{2}O_{(aq)} \rightarrow CaSiO_{3(s)} + 2H^{+}_{(aq)} $$

$$ (R20) \quad Ca^{2+}_{(aq)} + 0.5Fe^{2+}_{(aq)} + 0.5Mg^{2+}_{(aq)} + 2SiO_{2}(aq) + 2H_{2}O_{(aq)} \rightarrow Ca(Fe,Mg)Si_{2}O_{6(s)} + 4H^{+}_{(aq)} $$

Secondary mineral assemblages in Allende igneous CAIs are largely controlled by the elements mobilized by the fluid and chemical composition of melilite replaced, that ranges from nearly pure gehlenite (Åk~2) to Åk~80. For example, gehlenitic melilite in CTA CAIs and in mantles of Type B1 CAIs is often crosscut by veins of grossular and anorthite with minor spinel and corundum (Figs. SM11a, 12). Mass-balance equations describing these reactions are:

$$ (R21) \quad 2Ca_{2}Al_{2}SiO_{7(s)} + 3SiO_{2} \rightarrow Ca_{3}Al_{2}Si_{3}O_{12(s)} + CaAl_{2}Si_{2}O_{8(s)} $$

$$ (R22) \quad 3Ca_{2}Al_{2}SiO_{7(s)} + Ca_{2}MgSi_{2}O_{7(s)} + SiO_{2} \rightarrow Ca_{3}Al_{2}Si_{3}O_{12(s)} + MgAl_{2}O_{4(s)} + 2CaO $$

$$ (R23) \quad 3Ca_{2}Al_{2}SiO_{7(s)} + 3SiO_{2} \rightarrow 2Ca_{3}Al_{2}Si_{3}O_{12(s)} + Al_{2}O_{3} $$

$$ (R24) \quad 3Ca_{2}Al_{2}SiO_{7(s)} + 2Ca_{2}MgSi_{2}O_{7(s)} + H_{2}O \rightarrow Ca_{3}Al_{2}Si_{3}O_{12(s)} + CaMg_{2}Al_{4}SiO_{10}(OH)_{2} + 6CaO + 3SiO_{2} $$

Åkermanite-rich melilite in cores of Type B1 CAIs and in Type C CAIs is replaced by coarse-grained grossular, monticellite, wollastonite, Al-diopside, and minor forsterite (Figs. SM13–14):

$$ (R25) \quad Ca_{2}MgSi_{2}O_{7(s)} + Ca_{2}Al_{2}SiO_{7(s)} + SiO_{2} \rightarrow Ca_{3}Al_{2}Si_{3}O_{12(s)} + CaMgSiO_{4(s)} $$

$$ (R26) \quad Ca_{2}MgSi_{2}O_{7(s)} + 2Ca_{2}Al_{2}SiO_{7(s)} + SiO_{2} \rightarrow Ca_{3}Al_{2}Si_{3}O_{12(s)} + CaMgSiO_{4(s)} + CaAl_{2}SiO_{6(s)} + CaO $$

$$ (R27) \quad 3Ca_{2}MgSi_{2}O_{7(s)} + Ca_{2}Al_{2}SiO_{7(s)} + 2Ca_{2}Al_{2}Si_{2}O_{8(s)} \rightarrow 3Ca_{3}Al_{2}Si_{3}O_{12(s)} + CaMgSiO_{4(s)} + Mg_{2}SiO_{4(s)} $$

$$ (R28) \quad Ca_{2}Al_{2}SiO_{7(s)} + 3Ca_{2}MgSi_{2}O_{7(s)} + SiO_{2} \rightarrow Ca_{3}Al_{2}Si_{3}O_{12(s)} + 3CaMgSiO_{4(s)} + 2CaSiO_{3(s)} $$

Because of large compositional variations of primary melilite in the Allende igneous CAIs, and, possibly, of local variations of fluid composition, different assemblages of secondary minerals can be observed even within a single inclusion (Fig. SM13).

Secondary minerals in Allende igneous CAIs recorded multistage metasomatic alteration. For example, (i) melilite in a FoB CAI was replaced by grossular, monticellite, and wollastonite (Fig. SM14a). In the peripheral part of the same CAI, grossular and monticellite were subsequently replaced by Al-diopside (Fig. SM14b):
$$ (R29) \quad \underset{\text{grl}}{\text{Ca}_3\text{Al}_2\text{Si}_3\text{O}_{12(\text{s})}} + \underset{\text{mnl}}{\text{CaMgSiO}_{4(\text{s})}} \rightarrow \underset{\text{di-CTs}_{\text{s}}}{\text{CaMgSiO}_{4(\text{s})}} + \text{CaAl}_2\text{SiO}_{6(\text{s})} + 2\text{CaO} + 2\text{SiO}_2 $$

(ii) Secondary mineral assemblage of monticellite, Al-diopside, and wollastonite replacing ˚akermanitic melilite in a Type B2 CAI is crosscut by a monomineralic vein of wollastonite (Fig. SM14c):

$$ (R30) \quad \text{CaO} + \text{SiO}_2 \rightarrow \underset{\text{wol}}{\text{CaSiO}_{3(\text{s})}} $$

(iii) Secondary anorthite in the peripheral portions of CAIs is often replaced by sodalite, nepheline, and ferroan olivine (Figs. 6c,d in Krot et al. 2021):

$$ (R31) \quad \underset{\text{an}}{\text{CaAl}_2\text{Si}_2\text{O}_{8(\text{s})}} + 3\text{Na}_2\text{O} + 2\text{NaCl} \rightarrow \underset{\text{sod}}{\text{Na}_8(\text{Al}_6\text{Si}_6\text{O}_{24})\text{Cl}_2(\text{s})} + \text{CaO} $$

$$ (R32) \quad \underset{\text{an}}{\text{CaAl}_2\text{Si}_2\text{O}_{8(\text{s})}} + \underset{\text{nph}}{\text{Na}_2\text{O}} \rightarrow \underset{\text{sod}}{2\text{NaAlSiO}_{4(\text{s})}} + \text{CaO} $$

$$ (R33) \quad \text{MgO} + \text{FeO} + \text{SiO}_2 \rightarrow \underset{\text{fa-fo}_{\text{s}}}{(\text{Fe, Mg})_2\text{SiO}_{4(\text{s})}} $$

Secondary minerals also provide evidence for local variations in chemical composition of metasomatic fluid. These include (i) ferroan grossular overgrowing Fe-poor/free grossular (Fig. SM15a); (ii) ferroan monticellite overgrowing Fe-poor/free monticellite (Fig. SM15b), and (iii) hedenbergite and andradite overgrowing Fe-poor/free Al-diopside (Fig. SM15c). These late-formed Fe-bearing/rich secondary minerals are typically Ti-free. Only rare occurrences of Ti-rich garnet, hutcheonite, overgrowing grossular have been described in the Allende Type B CAIs (Fig. SM15d). These observations suggest that mobility of titanium in the Allende aqueous fluid was very limited.

## 2.5 Metasomatic Alteration of Igneous CAIs in CK3.7–3.8 Chondrites

Coarse-grained igneous CAIs in CK3.7–3.8s have similar textures, primary mineralogies, and isotopic compositions to those from Allende (Chaumard et al. 2014; Shollenberger et al. 2018; Krot et al. 2024, 2023), but experienced higher degree of metasomatic alteration under much more oxidizing conditions, FMQ+2 to FMQ+4.5 vs. IW to ~FMQ, respectively (Righter and Neff 2007).

Only three igneous CAIs (CTA, B, FoB) have been studied in detail in CK3s so far (Fig. SM16; Krot et al. 2024). Secondary minerals identified in these CAIs are listed in Table SM5. In contrast to the commonly observed secondary monticellite, Na-melilite, nepheline, sodalite, and wollastonite in Allende igneous CAIs, none of these minerals have been identified in the CK3 CAIs studied. In addition, secondary grossular, Al-diopside, spinel, and olivine in CK3 CAIs are enriched in FeO, and secondary plagioclase contains significant Na₂O content (up to 85 mol% Ab).

Melilite is almost completely replaced by Fe-bearing grossular, Fe-bearing Al-diopside, clintonite, CaNa-plagioclase, FeMg-spinel, FeMg-olivine, and wadalite; anorthite is less altered (Figs. SM17–19). Primary Ti-rich pyroxenes (grossmanite and fassaite) are corroded by FeAl-diopside, ilmenite, titanite, FeMg-spinel, Ti-bearing grossular, and clintonite (Figs. SM17f, 18c–e). Low-Ti Al-diopside in a FoB CAI shows little evidence for alteration (Fig. SM19). Primary spinel and forsterite experienced Fe-Mg interexchange to various degrees. Coarse secondary grossular and Al-diopside in the CAIs cores are often chemically-zoned: the grain edges are enriched in TiO₂ (up to 3 and 7 wt%, respectively) suggesting local mobilization of Ti during metasomatic alteration (Figs. SM17c,d). Some grossular grains are overgrown by ferroan grossular with a sharp compositional boundary between them suggesting multistage formation of this mineral (Figs. SM17e, 18f).
The CK igneous CAIs are crosscut by veins of FeMg-olivine, CaNa-plagioclase, FeAl-diopside, Ti-bearing grossular, and occasionally contain geodes composed of these minerals (Figs. SM18b,f). These CAIs are also surrounded by a Ca-rich layer composed of chemically-zoned FeAl-diopside, Fe-grossular, and andradite; edges of the grossular and Al-diopside grains contain significant TiO₂ contents (up to 0.7 wt%) suggesting removal of Al and Ti from the host inclusions.

## 2.6 Metasomatic Alteration of Dark Inclusions in Allende

Chondritic lithic fragments, commonly observed in CV3s and named dark inclusions (DIs), appear to have experienced much higher degree of the metasomatic alteration than the host meteorites suggesting their alteration predated final accretion of the CV parent asteroid. Although originally the origin of DIs was attributed to nebular processes – condensation, aggregation, and alteration – under highly-oxidized, non-solar conditions (Kurat et al. 1989; Johnson et al. 1990), it was subsequently shown that alteration of DIs resulted in extensive mobilization of Ca, Na, K, and S and could have occurred in the CV parent asteroid in zones of high fluid permeability (Krot et al. 2001; Brearley and Krot 2012) prior to local impact brecciation and mixing. In this section, we will touch on only one aspect of Allende DIs – redistribution of Ca and formation of Ca-rich rims around them.

Figure SM20 shows three completely altered Allende DIs, 4302 – 1, IV – 1, and 3b – 1, which consist entirely of secondary minerals – ferroan olivine, nepheline, and CaFe-pyroxenes, and contain chondrule pseudomorphs (Fig. SM20d). In addition, the inclusion 4301 – 2 is crosscut by numerous CaFe-pyroxene veins. The DIs are surrounded by a thick layer of hedenbergite+wollastonite+andradite. High abundance of CaFe-silicate objects is also observed in the Allende host outside the CaFe-rich rim of 4302 – 1. The thickness of the CaFe-silicate layer around DIs is inversely proportional to the abundance of CaFe-pyroxenes inside their host inclusions suggesting loss of Ca from the previously altered DIs and its precipitation along the boundary with the Allende meteorite. Therefore, the Allende rimmed DIs, like CaFe-silicate layers around Allende CAIs (Fig. SM9), provide an important evidence for *in situ* alteration of Allende.

## 3 Modification of Organic Matter During Metasomatic Alteration

Organic material occurs in ordinary and carbonaceous chondrites (Yang and Epstein 1983; Robert et al. 1987; Alexander et al. 1998, 2007, 2012; Aponte et al. 2017, 2019), and mainly consists of insoluble organic matter (IOM); the soluble fraction (SOM) is very low in abundance. The IOM is located within the matrix, and is finely distributed through the mineral phases. The elemental and isotopic composition of IOM evolves with increasing petrologic type (see also Sects. 7.2 & 7.3), as well as its structural organization.

### 3.1 Abundance, Elemental and Isotopic Compositions of Insoluble Organic Matter

In comparison to lower grade (type 2 and 1) chondrites (CIs, CMs, and CRs), the IOM contents of CO3s, CV3s and UOCs are lower, which is mainly due to lower proportions of matrix in them. When normalized to the matrix contents, the IOM carbon content of the least metamorphosed type 3 chondrites is roughly comparable to those of CMs and CIs at ~2 wt% (Grady et al. 1989; Alexander et al. 2007). However, the H/C, N/C and O/C ratios
of IOM decrease with increasing petrologic type, indicating progressive loss of these hetero-atoms as the structure becomes more aromatic in response to thermal stress (Remusat et al. 2019). For terrestrial samples, such as kerogens in sedimentary rocks, this process is called carbonization (Vandenbroucke and Largeau 2007).

The carbon-isotope compositions of IOM become progressively heavier with heating intensity in the parent body (Alexander et al. 2007). This is consistent with the usual process of “cracking” that organic macromolecules experience under thermal stress (Sackett 1978) in which the residual IOM gets heavier with preferential loss of ¹²C-bearing functional groups (the ¹³C bonds are stronger than ¹²C bonds).

The nitrogen-isotope compositions of IOM are more variable, but scatter around δ¹⁵N of ~ 0‰ and do not clearly correlate with the petrologic type. Heating experiments performed on the IOM from CIs and CMs also do not show evolution of N-isotope composition (Remusat et al. 2019).

The hydrogen-isotope compositions of IOM in UOCs exhibit a dramatic increase in δD (up to ~ 12,000‰) with increasing of petrologic type (Alexander et al. 2007). Flash pyrolysis to 600 °C of the IOM from the L3.3 chondrite Grosvenor Mountains (GRO) 95502 resulted in a sharp increase in the δD of the residue (Remusat et al. 2016). This could be due to (i) the presence of minute amounts of extremely D-rich and thermally resistant organic moieties in the IOM of UOCs (Aléon 2010; Remusat et al. 2016) or (ii) the loss of thermally labile functional groups that are hydroscopic and/or easily exchange hydrogen isotopes (e.g., OH or acid groups) during IOM isolation (Alexander et al. 2017). Hydrogen-isotope fractionation associated with the progressive oxidation of Fe by H₂O and loss of H₂ are also proposed to explain the progressive D-enrichment of OC IOM (Alexander et al. 2010). Relative to IOM in UOCs, the δD values are much lower in the IOM of COs and CVs. In the reduced CVs, δD only slightly increases with petrologic type (Alexander et al. 2007). This would imply that the putative D-rich, thermally resistant component in the UOC IOM is not present in the IOM of the COs and CVs. It is also problematic for the Fe oxidation explanation for the D-enrichments in the UOC IOM since there is clear petrologic evidence for oxidation of Fe in the COs and CVs (see Sect. 2.2).

## 3.2 Structural Organization of IOM: A Sensitive Tool of the Peak Metamorphic Temperature

Raman spectroscopy has turned out to be the instrument of choice to evaluate the extent of thermal metamorphism in type 3 chondrites (Michel-Levy and Lautie 1981; Quirico et al. 2003, 2009, 2014; Bonal et al. 2006, 2007, 2016, 2020, 2007). As for terrestrial sample analogues, the chondritic IOM is described as a disordered macromolecule and the so-called Raman “defect” bands are sensitive to the structural organization degree and to the size of the polyaromatic units. As largely documented for terrestrial samples (see Buseck and Beyssac 2014 for a review), with increasing peak metamorphic temperature, the defects progressively diminish and the size of the aromatic domains increases. Raman spectroscopy thus allows for classification of the petrologic grade of chondrites independently of the mineralogy. However, a precise cross comparison among chondrite groups is possible only if the accreted precursors were comparable, and if the conditions during metamorphism (duration, T, P, fO₂, etc.) were similar (Beyssac et al. 2003; Quirico et al. 2009). Unfortunately, Raman spectroscopy cannot be used as a thermometer, since we have a limited knowledge on the nature of the accreted organic precursors, and that simulating long term evolution over geological timescales in the laboratory is highly challenging. An attempt has been made to
use Raman spectra to estimate peak metamorphic temperatures in chondrites by correlat-
ing them with independent temperature estimates for specific meteorites (Busemann et al.
2007), but Bonal et al. (2016) questioned this approach.

The structural evolution of IOM under temperature increase can also be documented by high resolution TEM (transmission electron microscopy), which reveals the size of the aromatic units and the spacing between aromatic layers in them as the IOM becomes increasingly carbonized (Rouzaud and Oberlin 1990). It has been shown that the typical size of the aromatic units increases in COs, CVs, and UOCs between type 3 and 3.6, from dimensions of ~1 nm to ~5–10 nm, and the mean interlayer spacing decreases to values that are close to those of graphite (Smith and Buseck 1981; Rietmeijer and MacKinnon 1985; Remusat et al. 2008; Le Guillou et al. 2012). Nuclear magnetic resonance (NMR) spectroscopy shows that it is the aliphatic functional groups that are lost first during metamorphism, while in carbon XANES (X-ray absorption near edge structure) spectra the increasing absorption intensity of an exciton is directly linked to the growth of aromatic units (Cody et al. 2008).

## 3.3 Soluble Organic Matter

In type 3 chondrites, the abundance of soluble organic compounds is much lower than in meteorites of lower petrologic types and much more restricted in their diversity (Schmitt-Kopplin et al. 2012). As a result, the SOM in type 3 chondrites has received much less attention compared to low petrologic type counterparts. A few amino acids were reported in COs and CVs, with concentrations around ppm level (~10× lower than in CMs, ~300× lower than in CRs). The predominance of amine terminal amino acids in COs and CVs is interpreted as the result of metasomatism in their parent bodies (Burton et al. 2012). A few monocarboxylic acids ranging from C1 to C6 have been reported in Allende, whereas acetic acid only has been detected in investigated COs and CVs (Aponte et al. 2019). However, the C-isotope compositions of the monocarboxylic acids in Allende indicate that they are dominated by terrestrial contamination. Aliphatic amines are also much less abundant in CVs and COs compared to CIs, CMs and CRs, with concentrations being one to three orders of magnitude lower (Aponte et al. 2017). In both cases, a lower abundance is likely the result of metasomatism in the parent body. Free polycyclic aromatic hydrocarbons (PAHs) show a similar pattern of decreasing abundance with increasing petrologic type (Elsila et al. 2005). In both COs and CVs, the lower abundances of SOM are almost certainly the result of metasomatic alteration in their parent bodies.

An extensive series of Mg-metalorganic compounds has been identified in thermally altered type 3 chondrites, but are virtually absent in type 1 and 2 chondrites (Ruf et al. 2017). This class of compounds is believed to be thermally stable, and hence can be used as a fingerprint for metasomatism in the parent bodies.

## 3.4 Open Questions

There is a clear effect of thermal metamorphism on organics in type 3 chondrites, which evolve chemically, isotopically and structurally. This transformation is interpreted as the result of long-term peak metamorphic temperature, likely related to ²⁶Al radioactive decay. It serves as a signature of the petrologic type and is a precious tool to classify chondrites.

Some questions remain open:

(i) The evolution mechanism (loss of hetero-atoms and structural reorganization) is assessed from measurements obtained on the bulk IOM but has not so far been studied in situ. This could reveal the organic-inorganic association and the role of the former on the maturation using scanning transmission X-ray microscopy (STXM) and TEM. This is, however,
made difficult by the increasing scarcity of organic particles within the matrix of higher petrologic types.

(ii) The origin of the precursors of OM in different chondrite groups and whether CVs, COs, and UOCs accreted the same mix of organic precursors remain unclear.

(iii) Whereas the role of fluids during metamorphism has been documented in UOCs, CVs and COs for inorganic phases, it is difficult to tell whether or not such fluids played a role on the organic evolution, or if temperature alone is responsible for the observed transformation. Experimentally identifying specific reaction pathways related to the presence of oxidizing fluids at temperature above 300 °C could potentially offer a way to resolve this question.

## 4 Modification of Presolar Grains Abundances and Abundances of Isotopically Anomalous Noble Gas Components Associated with Them During Aqueous/Metasomatic Alteration

Presolar grains are stellar condensates that form in the outflows of evolved stars (red giants, asymptotic red giants) and in the ejecta of stellar explosions (novae, supernovae), were incorporated in the protosolar molecular cloud (PMC), and can be found in chondrites and IDPs (e.g., Lewis et al. 1987; Tang and Anders 1988; Choi et al. 1998a; Zinner et al. 2005; Zinner 2014; Floss and Haenecour 2016). Presolar grains are distinguished from solar system material by their highly anomalous isotopic compositions and have been identified *in situ* with ion imaging techniques using either a NanoSIMS or SIMS+SCAPS (stacked CMOS-type active pixel sensor) detector. They are typically ~100–600 nm in size and include SiC, graphite, spinel, corundum, hibonite, and silicates (e.g., Nagashima et al. 2004; Zinner 2014; Nittler and Ciesla 2016).

In chondrites, presolar grains occur exclusively in interchondrule matrix and fine-grained chondrule rims, with apparently no systematic differences in the abundances among these locations (Leitner et al. 2016; Floss and Haenecour 2016). There are, however, indications of different alteration paths for interchondrule matrix and fine-grained rims in CRs and COs, resulting in differing presolar silicate abundances in matrix and rims (Leitner et al. 2016; Floss and Haenecour 2016).

Investigation of the abundance and distribution of presolar grains in different chondrite groups and petrologic types can shed light on possible heterogeneities in the PMC and/or PPD, and effects of metasomatic alteration and thermal metamorphism in the chondrite parent asteroids. To compare the distribution of presolar grains identified by ion imaging in matrices of different chondrites, they are commonly reported as matrix-normalized abundances. Presolar nanodiamond, SiC, and graphite contain relatively high abundance of isotopically anomalous noble gas components. As a result, the abundance of these presolar grains can be also inferred from whole-rock noble gas measurements (Huss 1990; Huss and Lewis 1995; Huss et al. 2003). Alexander (2005) argued that the relative abundances of presolar grains and their matrix-normalized absolute abundances are roughly CI-like in all chondrites. However, presolar SiC grain abundance estimates based on noble gas contents are inconsistent with this: they are typically lower (Huss and Lewis 1995; Huss et al. 2003). Alexander et al. (1990) and Russell et al. (1997) hypothesized that the presence of oxidants and a fluid to enhance diffusion and reaction may make it possible to degas/destroy the SiC grains at relatively low temperatures. Later Riebe et al. (2016) found that SiC grains in samples of Tagish Lake (C2ung) that experienced aqueous alteration to different degrees have different releases of noble gases when subjected to HNO₃ during Closed System Step
Etching (CSSE), a commonly used technique for extracting noble gases. This finding is in agreement with studies by Wieler et al. (1992), Floss and Stadermann (2009a), and Davidson et al. (2014) who showed that in CR2s the concentration of SiC grains determined by NanoSIMS raster ion imaging is higher than SiC concentrations derived from noble gas concentrations. Davidson et al. (2014) suggested that aqueous alteration in the CR parent body might have affected and degassed the outermost gas-rich layers of SiC grains, resulting in the lower estimate of SiC abundance based on noble gas concentrations.

Presolar silicate grains (typically amorphous), the most abundant type of presolar grains, make up to percent levels (on average ~ 400 ppm) of anhydrous IDPs (Floss et al. 2006; Busemann et al. 2009; Davidson et al. 2012; Alexander et al. 2017) and have a factor of a few lower abundance in type 3 carbonaceous chondrites (Nittler et al. 2018). The highest abundance of presolar silicates reported are in Acfer 094 and Chwichiya 002 (C3.00; 150–200 ppm), CO3.00 chondrites Allan Hills (ALHA)77307 (190 ppm) and DOM 08006 (240 ppm), minimally altered CRs (160–220 ppm), and the CM chondrite A-12169 (~275 ppm) (Nguyen et al. 2007, 2010, 2009b, 2009, 2015, 2018; Nittler et al. 2018, 2021; Davidson et al. 2019; Smith et al. 2023). Recent studies of UOCs showed similarly high presolar silicate abundances, although their distribution appears to be more heterogeneous. Presolar silicate abundances in relatively pristine matrices Semarkona (LL3.00), MET 00526 (L/LL3.05) and Queen Alexandra Range (QUE) 97008 (L3.05) are 150, 55–275, and 145 ppm, (see Fig. SM21 and references therein).

A systematic study of the inferred abundances of presolar SiC, diamond, and graphite in meteorites of various class and petrologic type showed correlations between the abundance patterns within a meteorite group and the metamorphic history of the host meteorite (Huss and Lewis 1995). Alexander et al. (1990) confirmed for a set of OCs the presence of isotopically anomalous phases only for specimens of petrologic type 3.6 or below, with destruction during metamorphism being the most likely and straightforward explanation. Besides SiC, presolar refractory oxide grains have been identified in grain separates from Tieschitz (H/L3.6) (Nittler et al. 1997, 2008). When compared the fraction of presolar Al-oxides among the total number of analyzed Al-oxide grains for the Tieschitz separates, 0.27±0.04% (Nittler et al. 1997) and 1.59<sup>+0.44</sup><sub>-0.35</sub>% (Nittler et al. 2008), respectively, with those for ordinary chondrites of lower petrologic types, e.g., Krymka (LL3.2) with 0.85±0.11%, no significant difference is evident: i.e., the occurrence of refractory presolar grains appears to be, on a first order estimate, unchanged up to a petrologic type of ~3.6.

Presolar silicates are sensitive indicators of aqueous alteration and thermal metamorphism (e.g., Zinner 2014) leading to rapid decreases in their abundance, decreases in the ratio of presolar silicates to oxides, and increases in the FeO contents in presolar silicates. However, abundances of presolar grains, especially of silicates, in chondrites of petrologic types >3.0 are not as thoroughly studied as in carbonaceous chondrites of lower petrologic types. Most of the available data are from OCs with petrologic types ranging from 3.00 to 3.2 (Fig. SM21).

Generally, the study of presolar grain abundances has the potential to support the identification of more complex, and potentially small, alteration processes affecting the fine-grained fractions of chondritic meteorites. Statistically significant variations of presolar silicate abundances could, in combination with other micro- and nanoanalytical techniques, provide additional information on the alteration state of a given area of a meteorite. The presolar silicate-to-oxide ratio can also be of limited value. Typically, presolar oxide abundances are very low (a few ppm), resulting in low grain numbers, which in turn are responsible for large errors for the respective ratios. Thus, statistically significant silicate-to-oxide ratios are only available for very thoroughly studied meteorites, or if the specimen contains
unusually high abundances of oxides (e.g., Leitner et al. 2012). A better approach might therefore be to monitor the presolar SiC abundances simultaneously, which are typically in the range of a few tens of ppm (e.g., Davidson et al. 2014); however, this requires even more effort, since both O-anomalous and C-anomalous grain abundances have to be measured. Such investigations are rather time-intensive, even more so, if presolar grain abundances are low. In addition, investigation of presolar silicates and oxides by AES (Auger Electron Spectroscopy, e.g., Floss and Stadermann 2009b; Bose et al. 2010) and TEM-EELS (Transmission Electron Microscopy-based Electron Energy Loss Spectroscopy) can yield information on their Fe contents and the oxidation state of iron in these grains, and thus, allow inferences about the alteration state of their host materials.

# 5 Experimental Studies and Physicochemical Analysis of Metasomatic Alteration Reactions

There are experimental and theoretical approaches in understanding physicochemical conditions of metasomatic alteration that affected type 3 ordinary, CO, and CV chondrites. Both are discussed below.

## 5.1 Experiments

Experimental studies are largely focused on the origin of FeNi-carbides (Llorca and Casanova 1998), magnetite (Hong and Fegley 1998), fayalite (Dobrică et al. 2022), and phyllosilicates (Tomeoka and Kojima 1995; Kojima and Tomeoka 1999, 2000) commonly observed in chondrules and matrices of type 3.0–3.2 chondrites, and alteration of melilite, anorthite, and enstatite described in refractory inclusions and chondrules in type 3 chondrites of higher petrologic subtypes (Nomura and Miyamoto 1998; Ohnishi and Tomeoka 2007; Ichimura et al. 2017). There are only a few experimental studies focused on the effects of alteration on noble gases concentrations and isotopic compositions (Nakasyo et al. 2000). Various types of starting materials were used in the experiments: synthetic and natural (meteorite fragments) samples, powder and block reactants, single crystals or polymineralic samples.

Llorca and Casanova (1998) studied reaction between kamacite grains and H₂+CO gas mixture in a nebular-type environment (H₂/CO = 250:1; $5 \times 10^{-4}$ atm total pressure (P), and temperature (T) of 473K). Carbon deposition, hydrocarbon production in the C₁–C₄ range, and the formation of an ε-carbide phase occurs when FeNi-metal particles are exposed to the gas mixture for 10³ hours, suggesting that gas-solid reactions in the solar nebula during CO hydrogenation represent a plausible mechanism for the formation of FeNi-carbides and carbonaceous material in IDPs.

The calculated magnetite formation temperatures in the canonical solar nebula are 357–384K, depending on the distribution of C between CO and CH₄, but formation of magnetite in the solar nebula may be kinetically inhibited (Fegley and Prinn 1989). Hong and Fegley (1998) studied oxidation of FeNi-metal from the Gibeon meteorite to magnetite via the net reaction 3Fe(s) + 4H₂O(g) → Fe₃O₄(s) + 4H₂(g) at ambient atmospheric P, at T = 364–715K, and a gas mixture with H₂/H₂O molar ratios of ~4–41. The obtained data were used to model the reaction time of FeNi-metal grains in the solar nebula as a function of grain size and T. Although the calculated reaction times for 0.1–1 µm radius metal grains are within theoretical lifetimes of the solar nebula (0.1–10 Ma), these estimates may provide only the lower limits (Hong and Fegley 1998) suggesting that formation of magnetite under nebular conditions (lower T, P, and higher H₂/H₂O ratio) seems unlikely.
Dobrică et al. (2022) performed a hydrothermal experiment using a mixture of FeO-rich amorphous silicates, Fe-metal powder, and water as starting components that was heated at T = 493K for 6 days. Multiple mineral phases were observed in the experimental products, including elongated porous crystals of pure fayalite, greenalite [(Fe²⁺,Fe₃⁺)₂-₃Si₂O₅(OH)₄], Fe-rich amorphous silicate, and iron oxide. The experiment supports a hydrothermal growth of fayalite in metasomatically altered type 3 chondrites (Zolotov et al. 2006). The lack of Mg in starting amorphous silicates is unfortunate, because matrices of primitive chondrites are dominated by FeMg-amorphous silicates which are often replaced by fayalite (Figs. SM3c,d).

Ohnishi and Tomeoka (2007) performed hydrothermal alteration reactions of enstatite at variable pH (0–14), T (373–573K), and reaction times (24–336 h). Under acidic conditions, no significant alteration occurred, whereas under neutral and alkaline conditions serpentine and saponite formed in different proportions. The amounts of phyllosilicates tend to increase with increasing pH, T, and run duration. There is also a tendency for saponite to form at higher pH and T and under longer run durations compared to serpentine. Ohnishi and Tomeoka (2007) concluded that alteration of CVs occurred at higher T and more alkaline conditions than that of CIs.

Tomeoka and Kojima (1995) and Kojima and Tomeoka (1999, 2000) performed long-time (7 weeks) hydrothermal alteration experiments of Allende at 723K, 800 bars pressure, under acidic and neutral conditions. It was found that pH and precursor mineralogy are important factors controlling phyllosilicate mineralogy. (i) *Under acidic conditions*, phyllosilicates replacing low-Ca pyroxene are dominated by a chlorite-group phyllosilicate containing 2–10 wt% Al₂O₃; serpentine and talc are minor. Phyllosilicates replacing matrix olivine consist of chlorite and serpentine and contain 8–10 wt% Al₂O₃. Phyllosilicates replacing chondrule mesostasis are dominated by a high-Al (12–20 wt% Al₂O₃) phyllosilicate. CaFe-rich pyroxenes in the matrix are extensively replaced by saponite. Matrix altered under acidic conditions is depleted in Ca, Na, K, and S compared to the Allende matrix. (ii) *Under neutral conditions*, phyllosilicates produced are mainly low-Al chlorite containing up to 2.5 wt% Na₂O; high-Al phyllosilicates were not observed. Diopside in CAIs and CaFe-rich pyroxenes in the matrix remain largely unaltered. Melilite and anorthite in CAIs are replaced by Al-saponite. Most common phyllosilicates in naturally altered CV3 are Fe-rich saponite and intergrowths of Na-mica with serpentine or saponite; chlorite is minor. Matrix altered under neutral conditions shows no significant differences in bulk major element compositions compared to the Allende matrix. By comparing the mineralogy and chemical compositions of naturally and artificially altered CVs, Kojima and Tomeoka (1999, 2000) concluded that alteration of CVs took place under variable pH.

In the experiments by Nomura and Miyamoto (1998) and Ichimura et al. (2017), pure gehlenite, diopside, spinel, and CaNa-plagioclase mixed with SiO₂, ±Al₂O₃ were heated in alkaline and acid solutions at T of ~473–673K. Spinel and diopside showed no evidence for alteration. Gehlenite was replaced by hydrogrossular [Ca₃Al₂(SiO₄)₃-ₓ(OH)₄x], hydrosodalite [Na₈Al₆Si₆O₂₄(OH)₂×2H₂O], analcime (NaAlSi₂O₆×H₂O), tobermorite (Ca₅Si₆O₁₈H₂×4H₂O) and nepheline hydrate (NaAlSiO₄×0.5H₂O); CaNa-plagioclase was replaced by Na-zeolites and analcime. Upon heating above 673 K, these minerals were transformed into grossular, dmisteinbergite, nepheline, sodalite, and wollastonite. Although none of the intermediate hydrated secondary phases have been observed in Allende CAIs, their formation cannot be excluded (e.g., Enokido et al. 2023). Studies of CAIs from CVs less metamorphosed than Allende are required to test this hypothesis.

Hydrothermal experiments performed on Allende at T of 473 K and water/rock of 1/50 showed that the concentration of noble gases decreased (Nakasyo et al. 2000). The loss of
noble gases also affected the isotopic composition of xenon, apparently due to the preferen-
tial alteration of nanodiamond containing mineral phase/phases relative to the trapped xenon
component carrier ones.

**5.2 Physicochemical Analysis of Metasomatic Alteration Reactions in CV Chondrites**

5.2.1 Transport of Elements in Aqueous Solution and Steam During Metasomatic Alteration

It is inferred that CVs experienced fluid-rock interaction during progressive metamorphism (Krot et al. 1995, 1998). The common presence of phyllosilicates in Kaba (CV<sub>OxB</sub>3.1) and nearly complete absence of hydrous minerals in Allende (CV<sub>OxA</sub>>3.6) suggest the aqueous alteration in the CV parent asteroid occurred at progressively increasing temperatures and was followed by thermal metamorphism resulting in loss of aqueous solutions and dehydration of phyllosilicates in the CV<sub>OxA</sub> chondrites.

Thermodynamic modeling of metasomatic alteration in a multisystem compositionally
similar to the CVs (Petaev and Mironenko 1997) shows that there are two distinct stages
of the alteration process. At low water/rock ratio, <0.1–0.3, no aqueous solution is present
in the system, because all water is consumed by the oxidation and hydration reactions. Hy-
drogen released by these reactions forms large volumes of gaseous phase that contains trace
amounts of volatile anionic elements (H, O, S, Cl, and C). At this stage, an aqueous alter-
ation could proceed as an isochemical process without appreciable redistribution of cationic
elements. At higher water/rock ratio, an aqueous solution coexists with the altered rock and
is capable of redistributing cationic elements such as Na, Ca, Mg, Fe, Si, and Al among CV
chondrite components.

The source of water in chondrite parent bodies was probably water ice. Upon heating, ice sublimates, melts and evaporates to form liquid water and steam, if the temperature exceeds boiling point (Fig. SM22). Liquid water and steam react with minerals and amorphous phases and leach some elements. Initially pure water is transformed into aqueous solution with SiO₂, Fe²⁺, Mn²⁺, Mg²⁺, Ca²⁺, Cl⁻, Na⁺, HCO₃⁻ etc. dissolved in it. Aqueous solution can transport dissolved elements at all temperatures. The ability of steam to transport elements depends on *P* and *T*. Supercritical fluid can efficiently transport elements, but measured Allende tensile strength values implies that at *P*<sub>gas</sub> > 100 bars the rock would break and gas would be released (Grimm and McSween 1989). Therefore, the presence of supercritical fluid in the CV asteroid seems unlikely. Low density subcritical “wet” steam (P<sub>H2O</sub> = P<sub>H2O,equil</sub>, *T* < 647 K) or superheated “dry” steam (P<sub>H2O</sub> < P<sub>H2O,equil</sub>, *T* > 647 K) can transport some elements dissolved in the solution.

The concentration of a solute in steam depends on its *T*, *P*, and density. The distribution coefficients of solutes between the “parental” aqueous solution and the steam are low (Fig. SM23). Steam condensation would produce a “secondary” solution nearly free of solutes which cannot be efficiently transported by the steam. Solubility of iron in water under oxidizing conditions is low; therefore, equilibrium concentration of iron in the steam is close to 0. Calcium concentration in the steam is expected to be also low because of low distribution coefficients of CaCl₂ and CaSO₄ between steam and aqueous solution. Therefore, neither Fe nor Ca can be transported by steam. Note that rapid mobilization of Ca in aqueous solution in Allende at temperature < 473 K was experimentally demonstrated (Jones and Brearley 2006).

Metasomatic alteration was a non-equilibrium, multistage process with many reactions observed not going to a completion (see Sects. 2.2–2.6). Therefore, we used equilibrium approach in evaluating intensive parameters of the observed reactions assuming they reached
steady state. The calculated values of fugacities and activities of gaseous and aqueous species for each reaction are compared with each other and with independent estimates of these parameters when available. Gaseous and solid solutions were treated as ideal. For aqueous species, at any T, total pressure (P<sub>tot</sub>) ≈ pressure of saturated water vapor (P<sub>sat H2O vapor</sub>); i.e., gaseous phase involved in alteration reactions is nearly pure H<sub>2</sub>O steam equilibrated with an aqueous solution.

### 5.2.2 Fe-C-O-H System: Oxidation and Carburization of FeNi-Metal and Formation of FeNi-Carbides and Magnetite

Oxidation and carburization of FeNi-metal could have taken place by a gaseous phase only, without aqueous solutions (see Sect. 2.2). The gas-phase in the Fe-C-O-H system may contain up to three major C-bearing species (CO, CO<sub>2</sub>, and CH<sub>4</sub>), depending on its chemical composition and P-T conditions. Carbon monoxide, CO, is dominant at high T, while methane, CH<sub>4</sub>, is the most thermodynamically stable C-bearing species at low T and moderate P characteristic of asteroidal alteration of type 3 chondrites as well as at low T and low P characteristic of the solar nebula. Therefore, if chemical equilibrium among gaseous species is reached in H-dominated nebular or asteroidal gases, then essentially all C will form CH<sub>4</sub>. However, in the cooling solar nebula the conversion of CO to CH<sub>4</sub> is kinetically inhibited (Fegley and Prinn 1989), thus CO remains the most abundant C-bearing species even at low T. Studies of hydrothermal gases in the H<sub>2</sub>O-H<sub>2</sub>-CO<sub>2</sub>-CO-CH<sub>4</sub> system suggest that in hydrothermal natural systems formation of CH<sub>4</sub> is also kinetically inhibited (Chiodini and Marini 1998). Therefore, in the analysis of solid-gas equilibria, CH<sub>4</sub> can be neglected.

The partitioning of C between gaseous CO and CO<sub>2</sub> is governed by the redox reaction:

$$ (R34) \quad CO(g) + H_2O(g) = CO_2(g) + H_2(g). $$

Concentration of either gaseous species can be used as an intensive parameter along with the T and P<sub>tot</sub> of the system. Under oxidizing conditions necessary to stabilize magnetite, the CO<sub>2</sub>/CO ratio in the gaseous phase is ~ 1 at high T; below ~900K, it is rapidly increasing with falling T. Here we choose the H<sub>2</sub>O/H<sub>2</sub> ratio and the concentration of CO<sub>2</sub> (X<sub>CO2</sub>) as intensive parameters. Then the CO<sub>2</sub>/CO ratio in the gaseous phase can be calculated using the equation:

$$ (R35) \quad \log (CO_2/CO) = 2041.5/T - 1.90 + \log (H_2O/H_2). $$

There are four Fe-carbides – orthorhombic Fe<sub>2</sub>C (η-carbide), hexagonal Fe<sub>7</sub>C<sub>3</sub> (ε-carbide), monoclinic Fe<sub>5</sub>C<sub>2</sub> (χ-carbide or Hägg-carbide), and orthorhombic Fe<sub>3</sub>C (θ-carbide or cementite, Ni-free analogue of meteoritic cohenite); their thermodynamic properties are poorly known (see Gibbs free energies of Fe carbides in Supplementary Material to Sect. 5.2.2).

The observed replacement of metal by magnetite and carbides in the Fe-C-O-H system can be described by the following chemical reactions:

$$ (R36) \quad Fe(s) + 4 H_2O(g) = Fe_3O_4(s) + 4H_2(g) $$

$$ (R37) \quad 3Fe(s) + CO_2(g) + 2H_2(g) = Fe_3C(s) + 2H_2O(g) $$

$$ (R38) \quad 5Fe(s) + 2CO_2(g) + 4H_2(g) = Fe_5C_2(s) + 4H_2O(g) $$

Similar reactions can be written for mutual replacements of metal, magnetite, cohenite, Hägg-carbide, wüstite (Fe<sub>0.947</sub>O), and siderite (FeCO<sub>3</sub>) (see Fe-C-O-H system: Oxidation and carburization of FeNi-metal and formation of FeNi-carbides and magnetite in Supplementary Material to Sect. 5.2.2).
Phase relations in the Fe-C-O-H system at different partial pressure of CO₂ (P<sub>CO2</sub>) as a function of T and H₂O/H₂ ratio are shown in Figure SM24. Low- and mid-T portions of the diagrams have the same topology at all P<sub>CO2</sub>: large stability fields of magnetite and Hägg-carbide are terminated by siderite at low T. As P<sub>CO2</sub> increases, the stability fields of siderite and Hägg-carbide increase at the expense of metal and magnetite; the Hägg-carbide—siderite—magnetite invariant point moves toward higher T and more oxidizing conditions (higher H₂O/H₂ ratios).

The topology of the high-T portions of the diagrams is P<sub>CO2</sub>-dependent. At p<sub>CO2</sub> = 0.05, the Hägg-carbide, unaltered (low-Ni) metal, magnetite, and wüstite are stable at T = 812K, H₂O/H₂ = 0.30, and CO₂/CO = 1.21. At lower P<sub>CO2</sub> values (Figs. SM24a,b), the invariant assemblage consists of metal, Hägg-carbide, and magnetite; the T and H₂O/H₂ ratio of the invariant point are decreasing as the P<sub>CO2</sub> decreases. If P<sub>CO2</sub> exceeds 0.05 bar, the invariant assemblage of metal, Hägg-carbide, and magnetite becomes unstable; metal and Hägg-carbide coexist with wüstite instead (Fig. SM24c). A decrease in X<sub>Fe</sub> has the opposite effect on the phase boundaries due to enlargement of the metal stability field; e.g., altered metal is in equilibrium with magnetite and metastable Hägg-carbide even at P<sub>CO2</sub> = 0.05 (Fig. SM24d). A decrease in X<sub>mgt</sub> would increase the magnetite stability field, but the amount of impurities in magnetite is so small, that the X<sub>mgt</sub> never drops below ~0.98; this would result in a negligible shift (0.008 log unit) of the magnetite phase boundaries.

High-T portions of all diagrams include metastable fields of cohenite and/or Hägg-carbide (Fig. SM24). Cohenite is metastable at all T. Hägg carbide becomes metastable relative to unaltered metal (X<sub>Fe</sub> = 0.95) and carbon above ~700K according to reaction:

$$ (R39) \quad \mathrm{Fe_5C_{2(s)}} = 5\mathrm{Fe(s)} + 2\mathrm{C(s)} $$

The assemblage of altered metal (X<sub>Fe</sub> = 0.5) and Hägg-carbide is metastable above ~650K.

The lack of wüstite in FeNi-metal-carbide-magnetite assemblages places upper limits on the T (812K), P<sub>CO2</sub> (0.05 bar), and H₂O/H₂ ratio (0.30) of the initial stage of alteration of type 3 chondrites. However, at such conditions the Hägg-carbide would be metastable. To form a stable mineral assemblage of unaltered metal, magnetite, and Hägg-carbide, the T should be below 700K and P<sub>CO2</sub> in the gaseous phase less than ~0.002 bar (Fig. SM24a). No cohenite will be formed at this stage of alteration. Because both the T and H₂O/H₂ ratio of the invariant assemblage of metal, magnetite, and Hägg-carbide are decreasing with decreasing P<sub>CO2</sub> values, this assemblage can be formed either in a nebular (low P<sub>CO2</sub> values) or asteroidal (higher P<sub>CO2</sub> values) environment.

### 5.2.3 Ca-Mg-Si-O-H System: Formation of Pure Magnetite, Fayalite, Hedenbergite, Andradite, and Kirschsteinite

The formation conditions of these minerals have been first evaluated by Krot et al. (1998). Due to space limitation, an updated discussion of the reactions involved and their physicochemical parameters is presented in the Supplementary Material to Sect. 5.2.3. Overall, Fayalite is stable in equilibrium with relatively reducing gaseous phase and Fe²⁺-rich aqueous solution. An increase of Ca²⁺ concentration in aqueous solution will result in formation of either hedenbergite or kirschsteinite at high or low SiO₂ concentrations in aqueous solution, respectively. Magnetite requires oxidizing gaseous phase and Fe²⁺-rich aqueous solution. Andradite would be the only mineral equilibrated with oxidizing gaseous phase and Ca²⁺-rich aqueous solution.

Thermochemical equilibria in closed gas-solution-rock system (Zolotov et al. 2006; Jogo et al. 2009) suggest that Fayalite content in olivine is a function of T, P, water/rock ratio,
and ratios of activities of Mg and Fe in aqueous solution (Fig. SM26). Pure fayalite is stable at lower *T* and Fe²⁺/Mg²⁺ ratio in aqueous solution, and higher water/rock ratio and *P* than FeMg-olivine. In equilibrium, fayalite, magnetite, CaFe-pyroxenes, FeNi-carbides, and phyllosilicates coexist.

### 5.2.4 Ca-Fe-Mg-Si-O-H System: Formation of Ca-Rich Rims Around Allende Dark Inclusions

Petrographic observations (see Sect. 2.6) indicate that Ca-rich rims around Allende DIs grew in situ, at the interface between the DIs and the host Allende after the breccia formation of Allende. Because the growth of the Ca-rich rims requires significant transport (re-mobilization) of Ca, Fe, Si, and Mg, aqueous solutions must have played an important role in this process.

Physicochemical conditions of Ca-rich rim formation have been first evaluated by Krot et al. (2001) and are reviewed in some detail in the *Supplementary Material to Sect. 5.2.4*. The analysis of observed mineral-solution reactions revealed that the aqueous solutions locally equilibrated either with DIs or Allende matrix should have had Fe²⁺/Ca²⁺ ratios different by a factor of ~1.2. This difference will result in diffusional fluxes of Ca²⁺ ions from the DIs towards the host Allende and Fe²⁺ ions in opposite direction, which will drive the reactive system toward equilibrium. Such compositional gradients and diffusional exchange of Ca and Fe between the solutions could exist only if the solutions are stagnant in pores of the host Allende and Allende DIs. Mixing of such solutions at the Allende-DI interface will increase concentrations of both Ca²⁺ and Fe²⁺ ions above the solubility limit of a CaFe-rich phase, which will precipitate in the cavities along the interface as a rim.

The rim formation will continue until matrix olivines in the DI and Allende become equilibrated, or the Ca-rich objects in the DI are exhausted, or when an aqueous solution dries out. Petrographic observations indicate that neither complete exhaustion of the DIs in Ca, nor complete equilibrium between the DIs and Allende have been achieved. As a result, the growth of the Ca-rich rims was probably terminated by disappearance of the aqueous solution.

### 5.2.5 Ca-Fe-Mg-Zn-H-O System: Formation of Zn-Bearing Hercynite and Ilmenite

Substitution of Ca by Fe±Zn in krotite, grossite, and perovskite in CAIs indicates arrival Fe- and Zn-bearing aqueous solution from the matrix, perhaps, before the formation of CaFe-rich silicate rims around CAIs. As the Ca²⁺/Fe²⁺ ratio in the solution diminishes due to gradual removal of Ca²⁺, krotite and grossite are altered first, followed by perovskite, well before the formation of Fe-bearing grossular (Figs. SM29a,b). An accompanying change of Mg²⁺/Fe²⁺ ratio should increase Fe content in spinel and olivine first, followed by formation monticellite-kirschsteinite solid solution (Fig. SM29c).

### 5.2.6 Alteration of Igneous CAIs in Allende

The altered patches and veins in mantles and cores of the Allende igneous CAIs are dominated by Ca, Al, Mg-rich phases which are enriched in SiO₂ and depleted in CaO compared to the primary phases (Tables SM1, 4). Such mineralogy requires an efficient addition of Si to and removal of Ca from CAIs. Most secondary FeMg-silicates in CAI interiors — forsterite, monticellite, grossular, and Al-diopside — are essentially Fe-free compared to the ferroan olivine, kirschsteinite, andradite, and salite—hedenbergite pyroxene in fine-grained rims and
neighboring matrix. It appears that for some reason the Fe, abundant in an aqueous solution circulating through the Allende matrix (Krot et al. 1998; Ganino and Libourel 2017), has not been efficiently delivered to the interiors of the Allende igneous CAIs, which prior to the alteration were very compact. The same may be true for Na and Cl which are predicted to have high concentrations in an aqueous solution equilibrated with matrix (Petaev and Mironenko 1997), but in wadalite replacing âkermanitic melilite in Type B CAIs, these elements are decoupled from each other.

High concentrations of SiO₂, Na, and Cl are expected in both steam and aqueous solution. Therefore, these elements can be supplied to CAIs by steam equilibrated with aqueous solution without addition of Fe and Ca insoluble in steam. Reactions involving SiO₂, CO₂, Na, and Cl do not require aqueous solution. Reactions involving Ca, Fe, and Mg need aqueous solution for addition or removal of these elements. This may explain Fe-rich secondary minerals in matrix and FeO-poor secondary minerals in CAIs. Sodium and chlorine are not expected to be decoupled, unless steam formed at early stage of alteration contains only Cl-bearing species such as HCl and NH₄Cl which may be present in accreted ices (Zolotov and Mironenko 2007). This may explain formation of Na-free wadalite in CAI cores.

The inferences about the involvement of an aqueous solution or H₂O steam or both in the alteration process can be made based on a physicochemical analysis of the observed and/or possible alteration reactions and the textural relationships and mineral crystallization sequences in altered refractory inclusions. The textures and mineral chemistry of altered areas leave no doubt that an aqueous alteration was a non-equilibrium, multistage process, with many observed reactions not going to completion (see Sect. 2.4). Given the multitude of such non-equilibrium effects, considering them would make the problem intractable. Instead, an equilibrium approach in evaluating intensive parameters of the observed alteration reactions is employed here, assuming that they have reached a steady state. Then, in order to evaluate the plausibility of these reactions, the calculated values of fugacities and activities of different gaseous and aqueous species for each reaction are compared (for details see Supplementary Material to Sect. 5.2.6.). Since large number of cations and anions were involved in alteration reactions, the observed reactions can be split into several subsystems based on chemical elements added to and/or removed from primary mineral assemblages: (1) SiO₂, (2) CO₂, (3) SiO₂, CO₂, and Ca, and (4) Na±Cl and Ca±SiO₂. The results of modeling are shown in Figures SM30–33.

(1) The reactions involving addition of either gaseous H₄SiO₄ or aqueous SiO₂ (Fig. SM30), can proceed forward at any *T* and *P* because the required H₄SiO₄ fugacities in steam and activities of SiO₂ in aqueous solution are either much lower than or comparable with the solubility of SiO₂ in the steam or in the aqueous solution formed by its condensation.

Of special value here are the reactions describing possible alteration of nepheline (NaAlSiO₄) to analcime (NaAlSi₂O₆(s) × H₂O):

$$ (R55_{aq}) \text{ NaAlSiO}_{4(\text{s})} + \text{SiO}_{2(\text{aq})} + \text{H}_{2}\text{O}_{(\text{liq})} = \text{NaAlSi}_{2}\text{O}_{6(\text{s})} \times \text{H}_{2}\text{O}_{(\text{s})} $$

$$ (R55_{g}) \text{ NaAlSiO}_{4(\text{s})} + \text{H}_{4}\text{SiO}_{4(\text{g})} = \text{NaAlSi}_{2}\text{O}_{6(\text{s})} \times \text{H}_{2}\text{O}_{(\text{s})} + \text{H}_{2}\text{O}_{(\text{g})} $$

Experimental studies at *T* – *P*<sub>H₂O</sub> conditions identical to those considered here (Wilkin and Barnes 1998) showed that analcime is a stable phase, which can be dissolved or replaced by other phases if activities of SiO₂ and cations in the aqueous solution change. Therefore, the lack of analcime in the CV chondrites effectively places an upper limit on the H₄SiO₄ fugacity in the steam (Fig. SM30a) and *a*SiO₂ in aqueous solution (Fig. SM30b) implying that during CAIs’ alteration, neither steam nor aqueous solution was SiO₂-saturated. Because the upper limit of aqueous *a*SiO₂ is 3–6 orders of magnitude higher than that in
the aqueous solution made by *in situ* condensation of steam (Fig. SM30b), a solution with a higher concentration of SiO₂, tentatively labeled as “external” aqueous solution, can be involved in CAI’s alteration. Such a solution must come to a CAI from outside (matrix?) as a liquid (Krot et al. 2001; Ganino and Libourel 2017). The lower limits on the H₄SiO₄ fugacity and aSiO₂ is set by reactions of melilite replacement by anorthite+grossular and by grossular+monticellite+wollastonite, respectively, which are widespread in the Allende igneous CAIs (Figs. SM12a, 14a).

(2) The phase equilibria involving loss or gain of CO₂:

$$ (R56) \quad 3\text{CaCO}_3\text{(s)} + 2\text{CaSiO}_3\text{(s)} = \underset{\text{calcite}}{\text{Ca}_5\text{Si}_2\text{O}_7(\text{CO}_3)_{2\text{(s)}}} + \underset{\text{wollastonite}}{\text{CO}_2\text{(g)}} $$

$$ (R57) \quad \underset{\text{tilleyite}}{\text{Ca}_5\text{Si}_2\text{O}_7(\text{CO}_3)_{2\text{(s)}}} = \underset{\text{spurrite}}{\text{Ca}_5\text{Si}_2\text{O}_8(\text{CO}_3)_{\text{(s)}}} + \text{CO}_2\text{(g)} $$

$$ (R58_{aq}) \quad \text{CaCO}_3\text{(s)} + \text{SiO}_2\text{(aq)} = \text{CaSiO}_3\text{(s)} + \text{CO}_2\text{(g)} $$

$$ (R58_g) \quad \text{CaCO}_3\text{(s)} + \text{H}_4\text{SiO}_4\text{(g)} = \text{CaSiO}_3\text{(s)} + \text{CO}_2\text{(g)} + 2\text{H}_2\text{O}\text{(g)} $$

and the occurrence of tilleyite, wollastonite, and calcite within the same CAI (Fig. SM10a,c) allow to estimate of XCO₂ (PCO₂) in steam as well as the activities of Ca²⁺ and HCO₃⁻ ions in aqueous solution (Fig. SM31). In Figure SM31, reactions involving melilite are shown for comparison; dashed lines denote metastable reactions. Relative stabilities of calcite and wollastonite are greatly dependent on the presence or lack of aqueous solution. If steam does not condense, stabilization of calcite requires high amounts of CO₂ (Fig. SM31a). If steam condenses, even very low (ppm) levels of CO₂ would result in precipitation of calcite. The general lack of calcite and common presence of wollastonite among melilite alteration products suggests that at the initial stages of melilite alteration silica was delivered as H₄SiO₄. Precipitation of calcite requires an aqueous solution with concentrations of [HCO₃⁻] ≈ 2×[Ca²⁺] of ~10⁻⁵–10⁻⁴ mol/kgH₂O (Fig. SM31b) and rather low mole fraction of CO₂ in the steam (Fig. SM31), which seems to be reasonable for a C-bearing system such as carbonaceous chondrites. At such conditions, the observed replacement of melilite by grossular+monticellite+forsterite can readily proceed, providing the released CaO reacts with CO₂ to form calcite, regardless of whether an aqueous solution saturated in calcite is present or not.

(3) Many alteration reactions require addition of CO₂±SiO₂ accompanied by removal of Ca±SiO₂ (Fig. SM32). While both CO₂ and SiO₂ can be added or removed in either steam (Figs. SM32a,b) or aqueous solution (Figs. SM32c,d), the very low solubility of Ca in steam requires the presence of an aqueous solution, which would absorb the released Ca as either dissolved Ca bicarbonate, Ca²⁺ + 2HCO₃⁻ or calcite. If SiO₂ is delivered in the steam, the latter must condense to form an aqueous solution, initially Ca-free, before alteration reactions can proceed. In such a case, the amount of aqueous solution is expected to be rather small, with its Ca absorption capacity being limited. The calcium concentration in the aqueous solution will affect the sequence of alteration reactions, with the alteration of gehlenitic melilite to anorthite being the most sensitive to the Ca build up. If such aqueous solution was not effectively drained before becoming calcite saturated, one would expect to see widespread calcite associated with altered silicates. While calcite is present at least in some altered Allende CAIs, it tends to occur in nodules hinting for an effective removal of Ca²⁺ and HCO₃⁻ with drained solution, which can accumulate in cavities and fill them with calcite and other CO₂-bearing minerals upon evaporation of water to form nodules.

(4) The formation of nepheline and sodalite (Fig. SM33) requires an addition of Na and Cl and removal of Ca (reactions R31 and R32) and addition or removal SiO₂ (reactions R59aq and R60aq):

$$ (R59_{aq}) \quad 2\text{Ca}_2\text{Al}_2\text{SiO}_7\text{(s)} + 3\text{SiO}_2\text{(aq)} + 2\text{Na}^{+}\text{(aq)} = \underset{\text{Ca}}{\text{Ca}_3\text{Al}_2\text{Si}_3\text{O}_{12}\text{(s)}} + \underset{\text{Na}}{2\text{AlSiO}_4\text{(s)}} + \underset{\text{Na}^{+}\text{(aq)}}{\text{Ca}^{2+}\text{(aq)}} $$
$$ (R60_{aq}) \ 2\text{CaAl}_2\text{Si}_2\text{O}_{8(\text{s})} + 2\text{Na}^{+}_{(\text{aq})} + 2\text{CO}_{2(\text{g})} + \text{H}_2\text{O}_{(\text{aq})} = 2\text{NaAlSiO}_{4(\text{s})} + \text{Al}_2\text{O}_{3(\text{s})} + 2\text{Ca}^{2+}_{(\text{aq})} + 2\text{HCO}_3^-_{(\text{aq})} + 2\text{SiO}_{2(\text{aq})} $$

Sodium, chlorine, and SiO₂ can be transported by either steam or aqueous solution, but the necessity to remove Ca requires the presence of aqueous solution. As discussed above, the latter can be formed either by steam condensation or arrive from the matrix. The substitution of anorthite by sodalite at higher T requires a solution with rather low Na (and Cl) concentrations that can be formed by the steam condensation. At lower T, an 'external' aqueous solution with higher Na and Cl concentration is required. Other reactions can proceed only in the presence of 'external' aqueous solution with high concentrations of Na and Cl. It is interesting to note that the presence of widespread Cl-bearing and Na-free wadalite in many altered Type B CAIs (Fig. SM9d) points to the decoupling of Na and Cl during at least early stages of aqueous alteration. In this case, the Cl could be supplied by the steam containing HCl and NH₄Cl but neither NaCl nor NaOH.

# 6 Chronology of Metasomatic Alteration

The chronology of metasomatic alteration provides important constraints on the setting of alteration (nebular vs. asteroidal), the accretion time, and the thermal histories of chondrite parent asteroids.

## 6.1 $^{53}$Mn-$^{53}$Cr Chronology of Fayalite

Doyle et al. (2015) reported high precision SIMS analyses of nearly pure fayalite in EET 90161 (L3.05), A-881317 (CV<sub>oxB3</sub>) and MAC 88107 (CO3.1-like). To ensure proper correction for relative sensitivities (RSF) between $^{55}$Mn⁺ and $^{52}$Cr⁺ ions, synthetic fayalite (Fa₉₉) doped with Mn and Cr was used as a standard. The excess of $^{53}$Cr in fayalite from these meteorites correlates with the $^{55}$Mn/$^{52}$Cr ratio, indicating in situ decay of $^{53}$Mn. The inferred initial $^{53}$Mn/$^{55}$Mn ratios [$^{53}$Mn/$^{55}$Mn]<sub>0</sub> in fayalite from EET 90161, A-881317 and MAC 88107 are (4.35±1.23)×10⁻⁶, (3.07±0.44)×10⁻⁶, and (2.58±0.21)×10⁻⁶, respectively (Fig. SM34). Note that since RSF strongly depends on fayalite content in olivine (Fig. SM35), the reported ($^{53}$Mn/$^{55}$Mn)<sub>0</sub> in fayalite using San Carlos olivine (Fa₁₀) as a SIMS standard are incorrect (Table SM6).

$^{53}$Mn is a short-lived radionuclide, therefore a time anchor with a known $^{53}$Mn/$^{55}$Mn ratio and U-corrected Pb-Pb absolute age is required to translate the inferred ($^{53}$Mn/$^{55}$Mn)<sub>0</sub> to absolute ages. The D'Orbigny angrite with the measured U-corrected Pb-Pb absolute age of 4563.37±0.25 Ma (Brennecka and Wadhwa 2012) is often used as the time anchor. We note, however, that two different ($^{53}$Mn/$^{55}$Mn)<sub>0</sub> ratios have been reported for this meteorite, (3.24±0.04)×10⁻⁶ (Glavin et al. 2004) and (3.54±0.18)×10⁻⁶ (McKibbin et al. 2015). The former is based on whole-rock and mineral separates measurements using Thermal Ionization Mass Spectrometry (TIMS); the latter is based on in situ measurements of olivine and kirschsteinite using SIMS. When anchored to ($^{53}$Mn/$^{55}$Mn)<sub>0</sub> of D'Orbigny reported by Glavin et al. (2004), these ratios correspond to fayalite formation at 2.4<sup>+1.8</sup><sub>-1.3</sub>, 4.2<sup>+0.8</sup><sub>-0.7</sub>, and 5.1<sup>+0.5</sup><sub>-0.4</sub> Ma after CV CAIs, respectively. If we use ($^{53}$Mn/$^{55}$Mn)<sub>0</sub> reported by McKibbin et al. (2015), these ages will be younger by ~0.5 Ma.
## 6.2 $^{53}$Mn-$^{53}$Cr Chronology of Kirschsteinite

MacPherson et al. (2017) reported high precision SIMS analyses of kirschsteinite (CaFeSiO₄) in the reduced CV3s Vigarano and Efremovka. To ensure proper correction for relative sensitivities between $^{55}$Mn⁺ and $^{52}$Cr⁺ ions, synthetic kirschsteinite doped with Mn and Cr was used as a standard. Well-resolved $^{53}$Cr excesses correlate with $^{55}$Mn/$^{52}$Cr indicative of *in situ* decay of $^{53}$Mn; the inferred initial ratio ($^{53}$Mn/$^{55}$Mn)₀ is (3.71±0.50)×10⁻⁶ (Fig. SM36). When anchored to $^{53}$Mn-$^{53}$Cr relative and U-corrected $^{207}$Pb-$^{206}$Pb absolute ages of the D'Orbigny angrite (Glavin et al. 2004; Brennecka and Wadhwa 2012), this ratio corresponds to kirschsteinite formation 3.2<sup>+0.8</sup><sub>-0.7</sub> Ma after CV CAIs. If we use ($^{53}$Mn/$^{55}$Mn)₀ reported by McKibbin et al. (2015) as an anchor, this age will be younger by ~0.5 Ma. The kirschsteinite data are consistent within error with the data for aqueously-formed fayalite from the CV<sub>oxB</sub>3 chondrite A-881317 as reported by Doyle et al. (2015), supporting the idea that CaFe-silicates in CV3s are cogenetic with fayalite and formed during metasomatic alteration in the CV parent asteroid(s).

## 6.3 $^{26}$Al-$^{26}$Mg Chronology of Secondary Minerals in CV3 Igneous CAIs

There are a number of secondary minerals in the Allende igneous CAIs with high $^{27}$Al/$^{24}$Mg ratio, >20, including anorthite, corundum, grossular, nepheline, and sodalite, and therefore suitable for search for a resolvable excess of radiogenic $^{26}$Mg ($^{26}$Mg*) using *in situ* SIMS measurements. Most analyses of these minerals showed no detectable $^{26}$Mg* [( $^{26}$Al/$^{27}$Al)₀ <5×10⁻⁶] suggesting formation > 3 Ma after primary CAI minerals (e.g., Hsu et al. 2006). This formation age is consistent with $^{53}$Mn-$^{53}$Cr ages of secondary fayalite and kirschsteinite in CV3s, and, thus suggesting metasomatic alteration of Allende CAIs occurred in an asteroidal setting. However, Fagan et al. (2007) reported resolvable $^{26}$Mg* in several grossular and plagioclase analyses of Type B and CTA CAIs in Allende which correspond to the inferred ($^{26}$Al/$^{27}$Al)₀ of ~4.5×10⁻⁵ (see their Fig. 8). Fagan et al. (2007) concluded that Allende CAIs experienced a protrackted, episodic alteration: some alteration occurred during the formation of CAIs; some continued in the CV parent asteroid. This interpretation was subsequently challenged by Krot et al. (2010) who demonstrated that $^{26}$Mg* in secondary grossular could have been inherited from primary gehlenitic melilite replaced by this grossular, and, thus has no chronological significance (Fig. SM37).

Ushikubo et al. (2007a,b) reported $^{26}$Mg* in anorthite of an Allende Type B CAI corresponding to $^{26}$Al/$^{27}$Al)₀ of (1.2±0.5)×10⁻⁵ and interpreted this result as an evidence for a late-stage nebular episode of alteration. Krot et al. (2021) showed subsequently that anorthite measured by Ushikubo et al. (2007a,b) is in fact primary, igneous and its Al-Mg isotope systematics was probably disturbed during thermal metamorphism (see Sect. 6.4) and, therefore has no chronological significance.

To summarize, $^{26}$Al-$^{26}$Mg systematics of secondary minerals in Allende CAIs suggest that metasomatic alteration was a late-stage process, > 3 Ma after crystallization of primary minerals.

## 6.4 Modification of $^{26}$Al-$^{26}$Mg Systematics in Primary Minerals of CAIs and Chondrules from Metasomatically Altered Type 3 Chondrites

Metasomatic alteration could have disturbed Al-Mg isotope systematics in some primary phases in chondrules and refractory inclusions of type 3 chondrites. In type 3.0–3.1 UOCs, COs and CVs, most CAIs and AOAs contain $^{26}$Mg* due to *in situ* decay of $^{26}$Al and
define internal isochrons with the canonical ($^{26}$Al/$^{27}$Al)$_0$ of ~5×10⁻⁵ (e.g., MacPherson et al. 1995; Ito and Messenger 2010; Ushikubo et al. 2017). These also include CAIs which experienced postcrystallization O-isotope exchange in melilite, anorthite, grossite, and krotite (e.g., Simon et al. 2019), suggesting that Mg self-diffusion in these minerals under $P-T-f_{O2}$ conditions of metasomatic alteration of type 3.0–3.1 chondrites is much slower than O self-diffusion.

In CAIs from more metamorphosed chondrites such as Allende (CV>3.6), the ($^{26}$Al/$^{27}$Al)$_0$ of anorthite show a wide spread, ranging from the canonical to ~ 0, whereas other minerals in the same CAIs plot along the canonical isochron (e.g., Jacobsen et al. 2008). The observed spread in ($^{26}$Al/$^{27}$Al)$_0$ in anorthite is most likely due to redistribution of Mg isotopes through Mg-self diffusion (Fig. SM38) during thermal metamorphism that resulted in homogenization of Mg isotopes and disturbance of its $^{26}$Al-$^{26}$Mg systematics (LaTourette and Wasserburg 1998; Ito and Messenger 2010). Although both melilite and anorthite in Allende igneous CAIs experienced postcrystallization O-isotope exchange (see Sect. 7.1.4.3), the $^{26}$Al-$^{26}$Mg systematics in melilite are undisturbed, suggesting its O self-diffusion in melilite is much faster than Mg self-diffusion under $P-T-f_{O2}$ conditions of metasomatic alteration experienced by Allende.

In coarse-grained igneous CAIs from NWA 4964 (CK3.8) and NWA (CK3.7), fluid-assisted thermal metamorphism resulted in the disturbance of $^{26}$Al-$^{26}$Mg systematics to various degree in melilite, grossmanite, AlTi-diopside, and spinel (Fig. SM39; Krot et al. 2023).

In chondrules from type 3.0–3.1, $^{26}$Al-$^{26}$Mg systematics of chemically pristine plagioclase (having excess of Si and high Mg content; Tenner et al. 2019) appears to be undisturbed, even when plagioclase experienced extensive O-isotope exchange (Nagashima et al. 2017, 2018, 2018, 2021, 2022). In contrast, $^{26}$Al-$^{26}$Mg systematics of chondrule glassy mesostasis and plagioclase replaced by nepheline in weakly metamorphosed chondrites (3.05–3.1) can be reset resulting in younger crystallization ages (Van Orman et al. 2014; Nagashima et al. 2017; Piralla et al. 2023).

## 6.5 $^{36}$Cl-$^{36}$S Systematics of Sodalite and Wadalite in Allende and Ningqiang (C3-Ungrouped)

$^{36}$Cl is a short-lived radionuclide that decays to $^{36}$Ar (98.1%, $\beta^-$) and $^{36}$S (1.9%, $\epsilon$ and $\beta^+$) with a half-life of 0.3 Ma. Early $^{36}$Cl studies of early solar system materials primarily focused on the search for $^{36}$Ar excesses (Villa et al. 1981; Göbel et al. 1982; Murty et al. 1997; Rai et al. 2003) and provided no conclusive evidence for the presence of $^{36}$Cl in the early solar system. Turner et al. (2013) first reported $^{36}$Ar excesses produced by the decay of $^{36}$Cl in sodalite from the Allende CAI *Pink Angel*. The initial $^{36}$Cl/$^{35}$Cl ratio (($^{36}$Cl/$^{35}$Cl)$_0$) in sodalite that formed ~ 7 Ma after CAI formation (based on $^{129}$I-$^{129}$Xe chronology) is (1.9±0.5)×10⁻⁸. No evidence for $^{26}$Al, ($^{26}$Al/$^{27}$Al)$_0$ < 1.7×10⁻⁶, was found in the *Pink Angel* sodalite (Hsu et al. 2006). Conversely, studies of the $^{36}$Cl-$^{36}$S systematics in sodalite and wadalite in CAIs and chondrules from Allende and Ningqiang showed clear excesses of $^{36}$S correlated with $^{35}$Cl/$^{34}$S ratios, indicative for *in situ* decay of $^{36}$Cl (Lin et al. 2005; Hsu et al. 2006; Ushikubo et al. 2007a,b; Jacobsen et al. 2011). The inferred ($^{36}$Cl/$^{35}$Cl)$_0$ in these minerals ranges from (1.81±0.13)×10⁻⁵ to (3.7±0.8)×10⁻⁶. The $^{26}$Al-$^{26}$Mg systematics in the same sodalite and grossular-wadalite mineral assemblage showed no resolvable excesses of $^{26}$Al, with upper limits on ($^{26}$Al/$^{27}$Al)$_0$ ranging between 0.4×10⁻⁶ and 7.0×10⁻⁶. This corresponds to formation ages between 3 to 4 Ma after CAI formation, suggesting that sodalite and wadalite in the CV3 parent body formed after or close to the time $^{26}$Al had gone extinct.
Large excesses of ³⁶S were reported in sodalite of a heavily metasomatically altered Allende CAI *Curious Marie*. These excesses do not correlate with the ³⁵Cl/³⁴S ratios, suggesting late closed-system homogenization of ³⁶S excesses after ³⁶Cl became extinct (Tang et al. 2017). The ²⁶Al-²⁶Mg systematics of *Curious Marie* shows nearly flat, elevated ²⁶Mg* (~40‰) and extremely high bulk ²⁷Al/²⁴Mg ratio (~95). Assuming δ²⁶Mg₀ = 0, the inferred model (²⁶Al/²⁷Al)₀ is (6.2±0.9)×10⁻⁵. These observations suggest that either the secondary phases in *Curious Marie* formed early (<0.15 Ma after CAI formation) when ²⁶Al was still at the canonical level, as favored by Tang et al. (2017), or the alteration took place after ²⁶Al had completely decayed away and ²⁶Mg* had been homogenized.

No ³⁶S excesses were found in several CAIs and chondrules from CVs (Plagge et al. 2006; Nakashima et al. 2008) and in a halite from the Zag (H3–6) meteorite (Nakashima et al. 2010).

There is no agreement on the origin of the ³⁶Cl excesses. The proposed models include stellar nucleosynthesis (Lin et al. 2005; Tang et al. 2017) and irradiation by solar energetic particles (SEP) (Hsu et al. 2006; Jacobsen et al. 2011; Wasserburg et al. 2011; Leya et al. 2018). The initial abundance of ³⁶Cl and decoupling from ²⁶Al in same or co-existing minerals, as well as the “young” I-Xe ages in *Pink Angle* sodalite, indicate (i) the production of ³⁶Cl by late-stage SEP irradiation in the PPD, and (ii) the production of ³⁶Cl is unrelated to the origin of ²⁶Al (Jacobsen et al. 2011). Jacobsen et al. (2011) concluded that ³⁶Cl was largely produced by irradiation of a volatile-rich reservoir [³⁵Cl(n,γ)³⁶Cl)], possibly HCl-bearing water ices (Zolotov and Mironenko 2007), in an optically thin region of the PPD adjacent to the region in which the CV chondrite parent asteroid accreted while the Sun was a weak T Tauri star. Subsequently, ³⁶Cl accreted into the Allende together with condensed water ices. Note that Leya et al. (2018) concluded that the correlation lines in the diagram ³⁶S/³⁴S and ³⁵Cl/³⁴S are not isochrons but mixing between water that contains chlorine (likely as ice) that has been irradiated either by galactic cosmic rays or SEP and sulfur (without any chlorine) with solar isotopic composition.

More work is needed to understand the origin of ³⁶Cl in the secondary metasomatic minerals in chondrites. For example, (i) wadalite is present in a CTA CAI from NWA 4964 (CK3.8) (Fig. SM17a) but has not been analyzed for its ³⁶Cl-³⁶S systematics yet. (ii) Large excesses of ³⁶S correlating with ³⁵Cl/³⁴S ratios, with (³⁶Cl/³⁵Cl)₀ as high as (9.4±0.8)×10⁻⁵, have been observed in lawrencite (FeCl₂) in an EL3 clast from the polymict ureilite Almahata Sitta (Feng et al. 2012). These preliminary results have not been published and need to be reinvestigated; they could have implication for the origin and distribution of ³⁶Cl in the early solar system as well as for gas dissipation in the PPD (Jacobsen and Desch 2023).

## 6.6 ¹²⁹I-¹²⁹Xe Chronology of the CV, CO, and CK Chondritic Components

The ¹²⁹I-¹²⁹Xe chronometer is based on the decay of now-extinct ¹²⁹I to ¹²⁹Xe with half-life of 16.14 Ma (García-Toraño et al. 2018), where the ratio of the accumulated daughter ¹²⁹Xe to stable ¹²⁷I in the host mineral reflects the iodine isotopic ratio ¹²⁹I/¹²⁷I at closure to Xe loss. This ratio is measured as the ratio of radiogenic ¹²⁹Xe (¹²⁹Xe*) present in the sample to ¹²⁸Xe* produced by artificial neutron irradiation of ¹²⁷I. This relative chronometer can be tied to the absolute age through the Shallowater aubrite. The absolute age of Shallowater has not been measured directly due to the low U content; instead it is derived from the observed correlation between ¹²⁹I-¹²⁹Xe and ²⁰⁷Pb-²⁰⁶Pb ages in a number of samples. Currently, the re-evaluated absolute I-Xe age of Shallowater is 4562.4±0.2 Ma (Pravdivtseva et al. 2017a).

I-Xe dating cannot be done *in situ*: separation of individual minerals or chondritic components is required. The separated polymineralic chondritic components could potentially
contain multiple I-carrier phases with different closure times. The I-Xe ages have been reported for single plagioclase and pyroxene grains (Pravdivtseva et al. 2009, 2013), phyllosilicate separates (Krot et al. 2006), and magnetite separates (Pravdivtseva et al. 2017b). All these minerals occur as secondary phases in metasomatically altered type 3 chondrites (Tables SM3–5). Such phases, when they are small, cannot be cleanly separated, which could make interpretation of I–Xe ages difficult. Nevertheless, the age information in such complex samples can be successfully resolved if, during step-wise heating, radiogenic Xe from different host phases is released at significantly different temperatures, resulting in multiple I-Xe isochrons.

I-Xe ages of magnetites separated from several CV3s, CO3s, and CK3–5 s and of Allende CAIs, chondrules, matrix, and DIs from several CVs relative to Shallowater are shown in Figure SM40 and listed in Tables SM7–8. It appears that formation of magnetite in CVs, COs, and CKs started nearly contemporaneously with the formation of magnetite in Orgueil (CI1; Pravdivtseva et al. 2017b) and Murchison (CM2; Pravdivtseva et al. 2013), ~0–2 Ma prior to Shallowater.

I-Xe ages of the Allende matrix and most Allende DIs are within the range of the CV magnetite. These samples yielded high-precision I-Xe ages suggesting similar I-carrier phase. Two DIs, one from Allende and another from Efremovka have older I-Xe ages (5.7±0.4 Ma and 4.92±1.8 Ma prior to Shallowater, respectively), suggesting a different, more refractory, I-carrier phase that was apparently not affected by later metasomatic alteration.

The heavily-altered fine-grained CAIs from Allende, with the exception of the CAI *Curious Marie*, have nearly identical release profiles for radiogenic ¹²⁹Xe* and ¹²⁸Xe* suggesting the same I-carrier, most likely sodalite (Kirschbaum 1988; Pravdivtseva et al. 2003). The major release peak of ¹²⁹Xe* in *Curious Marie* is also consistent with sodalite being the I-carrier phase. However, high-resolution step-wise pyrolysis of this CAI demonstrated that the I-Xe system survived intact in two different mineral phases, suggesting that metasomatic alteration affected them to different degrees. The fission composition in low-T extractions is consistent with Xe from ²⁴⁴Pu while Xe from neutron-induced fission of ²³⁵U resided in the high-T release peak, supporting two different I-carrier mineral phases.

I-Xe ages of the Allende fine-grained CAIs are ~ 3 Ma after Shallowater suggesting the I-carrier phase (sodalite) recorded later stage of metasomatic alteration compared to the Allende matrix and DIs. I-Xe ages of an altered chondrule, an altered coarse-grained CAI, and one of the I-carrier secondary phases in *Curious Marie* have intermediate ages, ~0–1 Ma after Shallowater (Fig. SM40). Two DIs from the reduced CV3s Vigarano and Leoville have the youngest I-Xe ages, ~ 9 Ma after Shallowater. Overall, CV3 components recorded a prolonged duration of metasomatic alteration of the CV parent asteroid, ~ 11 Ma.

## 6.7 ¹²⁹I-¹²⁹Xe Chronology of Chondrules from Type 3 Ordinary Chondrites

Among the 43 LL3 chondrules studied (Gilmour et al. 2006; Fig. SM41; Table SM9 and references therein), 10 have yielded two distinct isochrons, with higher temperature releases corresponding to the older apparent I-Xe ages, probably related to chondrule formation in the lower metamorphic grade chondrules. The I-Xe ages of the oldest OC chondrules cluster tightly at 2.6±1.8 Ma after CV CAIs. Lower temperature I-Xe ages generally reflect post-formational alteration. Not only do LL chondrules tend to be younger with increasing metamorphic grade, but the spread between high- and low-temperature I-Xe ages increases also, suggesting longer or multiple alteration events for higher metamorphic grade LL meteorites. Within each OC group, the I-Xe ages of chondrules from the higher metamorphic
grades tend to be younger, indicating longer post-formational processing or multiple sec-
ondary alteration events (Caffee et al. 1982; Pravdivtseva et al. 2002, 2004).

**6.8 Accretion Time and Thermal History of CV, CO, CK, and OC Parent Asteroids**

The $^{129}$I-$^{129}$Xe ages of magnetites from CV3s, CO3s, and CK3–5 s indicate the accretion of their parent bodies no later than ~3 Ma after CV CAIs. Furthermore, the $^{129}$I-$^{129}$Xe ages of altered CAIs and DIs in CV3s suggest a prolonged duration of metasomatic alteration, longer than 11 Ma, although the carrier phases and corresponding closure temperatures of I-Xe isotopic systematics are not well-constrained. The latter requires a relatively large parent asteroid for CV chondrites.

Thermal evolution models of asteroids heated by $^{26}$Al decay by Hevey and Sanders (2006) showed that asteroids as large as 50 km in radius can sustain high T close to peak T for longer than 10 Ma. Doyle et al. (2015) modelled the thermal evolution of the CV, CO, and L chondrite parent bodies heated by $^{26}$Al decay. Peak T in asteroids with radius larger than 30 km are predominantly determined by their accretion time (the amount of $^{26}$Al) but not by the parent body size. Thus, Doyle et al. (2015) determined the accretion time of these bodies, which are assumed to be >30 km in radius, to be 2.4–2.7 Ma, 2.1–2.4 Ma, and 1.6–1.8 Ma after CV CAIs for the CV, CO, and L chondrite parent bodies, respectively, so that the central T of these bodies reach the estimated peak metamorphic T (~600 °C, ~600 °C, and ~950 °C, respectively). They showed that fayalite can form ~0.6 Ma later than the accretion of the L chondrite parent body, or ~1.5–2.5 Ma later than the accretion of the CV and CO chondrite parent bodies in their outer regions where the peak T are no higher than 300 °C and suitable for fayalite formation without later Fe-Mg interdiffusion. These ages are consistent with the $^{53}$Mn-$^{53}$Cr ages of fayalite and kirschsteinite.

**6.9 Future Studies**

I-Xe ages of the Allende CV<sub>oxA</sub>>3.6 components dominate the data set for CV3s (Fig. SM40), demonstrating a need for a broader I-Xe study of chondritic components from the reduced and oxidized CVs of different petrologic types. Although small single chondrules could yield high-precision I-Xe ages (Caffee et al. 1982), the I-Xe ages of CV, CO, and CK chondrules are virtually absent. Resonance ionization techniques, such as the RELAX instrument developed by the University of Manchester group (Gilmour et al. 2006; Crowther et al. 2008), yield greater sensitivity than conventional mass spectrometry and provide unique opportunities for small samples and single-mineral studies, which would be critical for identifying I-carrier phases in metasomatically altered chondritic components.

# 7 Records of Metasomatic Alteration in Oxygen, Hydrogen, Nitrogen, Carbon, Sulfur, and Chlorine Isotopic Systematics of Type 3 Chondrites and Their Components

**7.1 Oxygen Isotopes**

**7.1.1 Introduction**

Oxygen has three stable isotopes: $^{16}$O, $^{17}$O, and $^{18}$O. Oxygen-isotope compositions of the solar system materials are typically expressed as permil deviations from the Vienna Standard
Mean Ocean Water (VSMOW) value, δ¹⁷O and δ¹⁸O: δ⁴O = ((⁴O/¹⁶O)ₘₚₐₗ/((⁴O/¹⁶O)ₗₚₐₗ)ₗₚₘₐₗ – 1)×1000), where i = 17, 18. On a three-isotope oxygen diagram, δ¹⁷O vs. δ¹⁸O, all terrestrial samples plot along a line with slope of ~ 0.52, called the terrestrial fractionation (TF) line, whereas extraterrestrial materials deviate from this line; the deviation is expressed as Δ¹⁷O = δ¹⁷O – 0.52×δ¹⁸O. Bulk oxygen-isotope composition of chondrites is one of the most important parameters used for their classification (Clayton 2008; Yurimoto et al. 2008; Ireland et al. 2020).

Ultraviolet CO self-shielding either in the PMC and/or in the outer regions of the PPD (outside ~ 30 AU) resulting in formation of ¹⁶O-rich CO and ¹⁷,¹⁸O-enriched water followed by silicate dust – H₂Oice – COgas fractionation and evaporation is the currently favored mechanism invoked to explain the observed range of Δ¹⁷O in the solar system (Clayton 2002; Yurimoto and Kuramoto 2004; Lyons and Young 2005; Krot et al. 2020).

Chondrules (phenocrysts and mesostasis) and fine-grained refractory inclusions in unmetamorphosed chondrites (petrologic type ≤3.0) have internally uniform Δ¹⁷O and on a three-isotope oxygen diagram plot along ~ slope-1 line, named the primitive chondrule mineral (PCM) line (Fig. SM42; Ushikubo et al. 2012, 2017, 2018). Secondary minerals formed during metasomatic alteration of type 3 chondrites deviate from the slope-1 and are out of isotope equilibrium with the primary minerals (see Sects. 7.1.2 & 7.1.3). In addition, primary minerals chondrules and refractory inclusions in chondrites of petrologic type > 3.0 are often isotopically heterogeneous (see Sect. 7.1.4). The nature of this heterogeneity remains controversial. The proposed mechanisms include condensation from or gas-melt exchange with nebular gas of variable O-isotope composition, and isotope exchange with metasomatic fluids in the chondrites parent asteroids (e.g., Yurimoto et al. 2008; Kawasaki et al. 2018; Krot et al. 2022b).

In this section, we summarize O-isotope compositions of the aqueously/metasomatically-formed minerals in type 3 chondrites and discuss possible effects of metasomatic alteration on O-isotope compositions of the primary minerals in their chondrules and refractory inclusions.

## 7.1.2 Oxygen-Isotope Compositions of Magnetite and Fayalite in Type 3.0–3.1 Chondrites

Although Acfer 094 avoided metasomatic alteration in its parent asteroid, it contains intergrowths, named cosmic symplectites (COS) composed of magnetite, pentlandite, troilite, Na₂SO₄, and elemental sulfur (Sakamoto et al. 2007; Matsumoto et al. 2022). Magnetite in COS is characterized by the anomalous O-isotope composition (δ¹⁷,¹⁸O ~ +180‰) and is interpreted to have recorded O-isotope composition of the primordial water in the PPD, and, hence, providing a strong argument supporting the CO self-shielding model (Yurimoto and Kuramoto 2004; Lyons and Young 2005; Sakamoto et al. 2007). The origin of COS, however, remains unclear: they may have formed by (i) oxidation and sulfidization of FeNi-metal in the PPD or in a pre-existing planetesimal (Sakamoto et al. 2007), (ii) high temperature interaction between FeNi-metal and O-Na-S-rich material (hot corrosion) (Matsumoto et al. 2022), and (iii) decomposition of amorphous ferric-hydroxysulfate-hydrate formed by interaction of sulfuric acid and ferric iron in a cryofluid environment on the Acfer 094 parent asteroid (Hashimoto and Nakano 2023).

Oxygen-isotope compositions of magnetite and fayalite formed during metasomatic alteration of type 3 UOCs, COs, CVs, and CKs are shown in Figures SM43–45. On the diagrams δ¹⁷O vs. δ¹⁸O, magnetite in fayalite in type 3.0–3.2 UOCs plot along mass-dependent fractionation lines with Δ¹⁷O of ~ +4.3‰ which correspond to Δ¹⁷O of aqueous fluids in their parent asteroids. There are no resolvable differences in the inferred Δ¹⁷O
of aqueous fluids in H, L, and LL chondrites or among chondrites of different petrologic subtypes (Fig. SM43). Magnetite and fayalite in CO3.05–3.1 s and MAC 88107 (CO3.1-like) chondrites plot along mass-dependent fractionation lines with $\Delta^{17}$O of ~ -0.2‰ and ~ -1.8‰, respectively (Figs. SM44). In contrast, magnetite and fayalite in the CVoxB chondrites, Kaba (CV3.1), Mokoia (CV~3.6), and Asuka (A)-881317 (CV3), show significant ranges in $\Delta^{17}$O, from ~0 to ~ -4‰ from ~ +3.5 to ~ -2‰, respectively (Figs. SM45a,b), suggesting either O-isotope evolution of aqueous fluid or heterogeneous O-isotope compositions of water ices that accreted into CV parent asteroid(s).

Oxygen-isotope compositions of magnetite in metamorphosed (petrologic types 3.6–3.8) CVs and CKs are shown in Figures SM45c,d. High precision data for magnetite in CKs have very limited ranges of $\delta^{18}$O (from -2.5 to +2.5‰) and $\Delta^{17}$O (~ -3‰), possibly reflecting O-isotope equilibration with matrix silicates during thermal metamorphism. Data for magnetite from Allende and Ningqiang (C3-ungrouped) are less precise; within analytical uncertainty, their $\Delta^{17}$O values are similar to those of CK magnetites; $\delta^{18}$O values are more variable.

To summarize, there are differences in $\Delta^{17}$O of fayalite and magnetite in UOCs, COs, CVs, and CKs suggesting different $\Delta^{17}$O of aqueous fluids, and, possibly water ices, in their parent asteroids. In the CVoxB chondrites, fayalite and magnetite show a range of $\Delta^{17}$O, possibly reflecting evolution of O-isotope composition of aqueous fluids. The inferred $\Delta^{17}$O of aqueous fluids are very different from $\Delta^{17}$O of COS that is hypothesized to represent O-isotope compositions of primordial water ices in the PPD.

### 7.1.3 Oxygen-Isotope Compositions of Secondary Minerals in Igneous CAIs from Allende and CK3.7–3.8 Chondrites

Oxygen-isotope compositions of the *in situ* SIMS measured secondary minerals in coarse-grained igneous CAIs from Allende (CVoxA > 3.6) and CK3.7–3.8 s are shown in Figures SM46a and SM46b, respectively. The data plot along mass-dependent fractionation lines with indistinguishable $\Delta^{17}$O of ~-3±2‰ (average ± 2SD) and ~-3.7±1.9‰, respectively. Although all minerals were measured with matrix-matched standards, there are some differences in chemical compositions of secondary minerals in CV and CK CAIs (see above) and the standards, which may explain some differences in their $\delta^{18}$O values.

### 7.1.4 Oxygen-Isotope Heterogeneity of Refractory Inclusions and Chondrules in Metasomatically Altered Type 3 Chondrites

**Fine-Grained CAIs and AOAs** Amoeboid olivine aggregates and fine-grained CAIs in unmetamorphosed ordinary and carbonaceous chondrites (petrologic type ≤3.0) have internally uniform $\Delta^{17}$O, typically solar-like, ~ -24±2‰ (Figs. SM42, 47–49) consistent with their formation in a gas of ~ solar oxygen-isotope composition. In contrast, AOAs and fine-grained CAIs in metamorphosed chondrites are isotopically heterogeneous: corundum, hibonite, spinel, forsterite, AlTi-diopside, and enstatite have $^{16}$O-rich solar-like compositions, whereas Zr- and Sc-rich oxides and silicates, grossite, krotite, perovskite, melilite, anorthite, and perovskite are $^{16}$O-depleted to various degrees (Figs. SM47–49). The $\Delta^{17}$O values of the most $^{16}$O-depleted primary minerals overlap with $\Delta^{17}$O values of metasomatic fluids inferred from O-isotope compositions of secondary minerals (see Sects. 7.1.2 & 7.1.3). These observations and the fact that the O-isotope heterogeneity is inconsistent with a condensation sequence of minerals in refractory inclusions suggest that O-isotope heterogeneity in fine-grained CAIs and AOAs resulted from isotope exchange during aqueous fluid–rock interaction in the chondrite parent asteroids (Simon et al. 2019; Krot et al. 2019a,b,c; Ebert et al. 2020; Han et al. 2024).
**Igneous Inclusions** Coarse-grained igneous CAIs are found exclusively in CVs and CKs. Both groups experienced metasomatic alteration and thermal metamorphism and lack type 3.0 members. Like the fine-grained CAIs in metamorphosed ordinary and carbonaceous chondrites, the coarse-grained igneous CAIs in CV3.1–4 s and CK3.7–3.8 s have internally heterogeneous O-isotope compositions: hibonite, spinel, forsterite, and low-Ti Al-diopside have solar-like $^{16}$O-rich ($\Delta^{17}$O ~ -24±2‰) compositions, whereas grossite, krotite, perovskite, melilite, anorthite, and Ti-rich pyroxenes (fassaite, grossmanite) are $^{16}$O-depleted to various degrees (Fig. SM50). There is no agreement on the nature of this heterogeneity. It could have resulted from (1) O-isotope exchange during multistage, incomplete melting of CAIs in the nebular gas of variable $\Delta^{17}$O (Yurimoto et al. 1998; Ito et al. 2004; Kawasaki et al. 2015, 2018; Aléon 2016, 2018) and/or (2) during aqueous fluid-rock interaction in the CV and CK parent asteroids (Yurimoto et al. 2008; Krot et al. 2022b, 2023).

(1) Igneous CAIs experienced melting in the solar nebula and thus could have experienced gas-melt O-isotope exchange. The early generation of gaseous reservoirs with variable $\Delta^{17}$O has been advocated in several papers (Yurimoto et al. 1998; Itoh and Yurimoto 2003; Aléon et al. 2007; J. Simon et al. 2005, 2011, 2016; Kawasaki et al. 2012, 2018; Aléon 2016, 2018, 2018, 2020, 2020). If igneous CAIs experienced O-isotope exchange with the nebular gas of variable $\Delta^{17}$O, the distribution of oxygen isotopes in these CAIs must be consistent with igneous crystallization sequence studied experimentally (Stolper 1982). However, this is generally not the case. For example, in Type B CAIs eutectic melt regions composed of co-crystallized highly-åkermanitic melilite (Åk$_{65-71}$), anorthite, low-Ti (~2–5 wt% TiO$_2$) fassaite, and spinel, the spinel and fassaite are similarly $^{16}$O-rich ($\Delta^{17}$O ~ -24‰), whereas melilite and anorthite are $^{16}$O-depleted to various degrees ($\Delta^{17}$O up to ~ -1‰; see (Fig. 11a,b in Krot et al. 2022b), suggesting post-crystallization O-isotope exchange.

The only mineral phase in igneous CAIs where igneous chemical zoning correlates with O-isotope zoning is AlTi-pyroxenes: Ti-rich pyroxene cores are $^{16}$O-depleted to various degrees relative to Ti-poor pyroxene edges having solar-like $\Delta^{17}$O (Kawasaki et al. 2018; MacPherson et al. 2022; Krot et al. 2022b). However, it remains unclear whether this O-isotope zoning was established during igneous crystallization or resulted from postcrystallization O-isotope exchange. Several observations may support the latter mechanism. (i) High-Ti pyroxene inclusions inside spinel grains were isolated from interaction with external O-isotope reservoir and preserved $^{16}$O-rich compositions (Figs. SM50a,d, 51). (ii) O-isotope zoning is observed in igneously zoned Ti-rich pyroxenes in FUN (fractionation and unidentified nuclear effects) CAIs which experienced melting, evaporation and crystallization in nearly vacuum (Krot et al. 2014b; Mendybaev et al. 2013, 2017) and therefore must have avoided gas-melt isotope exchange. (iii) There is an apparent correlation of $\Delta^{17}$O in pyroxenes and thermal history of the host meteorites: pyroxenes with similar chemical compositions are more $^{16}$O-depleted in CAIs from more metamorphosed chondrites, CK3.7–3.8 and NWA 8418 (CV4) (Fig. SM51).

(2) Oxygen-isotope composition of aqueous fluid in Allende (CV$_{OXA}$>3.6) and CK3.7–3.8 s is constrained by O-isotope compositions of their secondary minerals (see Sect. 7.1.3 and Fig. SM46). Postcrystallization O-isotope exchange of igneous CAIs with the fluid appears to be consistent with several observations: (i) $\Delta^{17}$O of $^{16}$O-depleted primary minerals (melilite, anorthite, high-Ti pyroxenes, grossite, krotite) approach $\Delta^{17}$O of the fluid (Figs. SM47–49). (ii) On $\delta^{17}$O vs. $\delta^{18}$O plots (Fig. SM50), the $^{16}$O-depleted primary minerals deviate to the right from a nebular slope-1 PCM line, i.e., towards the inferred compositions of aqueous fluids (Krot et al. 2022b). (iii) Oxygen self-diffusion in anorthite under wet conditions is several magnitudes higher than under dry conditions (Fig. SM52) and may explain its $^{16}$O-poor composition.
To summarize, the coarse-grained igneous CAIs in CVs and CKs experienced post-
crystallization O-isotope exchange to various degrees with an external ¹⁶O-poor reservoir
(Δ¹⁷O ~ -3±2‰), most likely aqueous fluid in the CV and CK chondrite parent asteroids.
The isotope exchange appears to be mineralogically not petrologically controlled: i.e. vari-
ations in O-isotope compositions of minerals within individual CAIs are inconsistent with
their inferred crystallization sequence. Oxygen-isotope compositions of melilite, anorthite,
Ti-rich fassaite, perovskite, grossite, krotite, and davisite were affected by this process to
various degrees. Hibonite, spinel, low-Ti fassaite, Al-diopside, and forsterite retained their
original O-isotope compositions established during igneous crystallization of coarse-grained
CAIs. This does not preclude O-isotope exchange in some igneous CAIs during melting in
the solar nebula (e.g., some FoB and FUN CAIs having relatively ¹⁶O-depleted composi-
tions, Δ¹⁷O > -20‰, and CAIs melted during chondrule formation), but such inclusions
appear to be rare (Kawasaki et al. 2012; Krot et al. 2005, 2014b).

**Chondrules** Chondrule phenocrysts and mesostasis in the least altered and the least meta-
morphosed carbonaceous chondrite Acfer 094 (C3.00) have internally uniform O-isotope
compositions (Fig. SM42). Glassy mesostasis and plagioclase in chondrules from even
weakly altered and metamorphosed ordinary and carbonaceous chondrites (Semarkona
(LL3.00), Yamato (Y)-81020 (CO3.05), and Kaba (CV<sub>oxB</sub>3.1) are ¹⁶O-depleted relative to
chondrule phenocrysts (Fig. SM53; see also Ebert et al. 2022). The Δ¹⁷O values of mesosta-
sis and plagioclase approach ¹⁷O of aqueous fluid inferred from secondary magnetite and
fayalite (see Sect. 7.1.2) suggesting they experienced O-isotope exchange with the fluid
(Kita et al. 2010; Ebert et al. 2022; Fukuda et al. 2022; Krot et al. 2022a,b).

**7.1.5 Future Studies**

Oxygen-isotope compositions of aqueous fluid in the reduced CV3s is not known yet, but
can be constrained by measuring secondary kirschsteinite. Development of the matrix-
matched mineral standards, including kirschsteinite, for SIMS measurements is required
to constrain the δ¹⁸O values of aqueous fluids, and, potentially, crystallization temperature
of secondary mineral parageneses.

To test the hypotheses of O-isotope exchange in igneous CAIs from CVs and CKs —
gas-melt and fluid-rock — experimental data on oxygen self-diffusion in CAI minerals un-
der proper physicochemical conditions (P — T — fO₂—water/rock ratio) of metasomatic
alteration of CVs and CKs are required.

**7.2 Hydrogen Isotopes**

**7.2.1 Bulk Chondrite Hydrogen-Isotope Compositions**

The first bulk H-isotope compositions measured in UOCs were obtained about 40 years
ago (Robert et al. 1979; McNaughton et al. 1982; Yang and Epstein 1983) and shown to
present high variabilities with δD relative to SMOW, from ~ -140 up to +3000‰. UOCs
of petrologic type ≤ 3.2 generally contain the highest amounts of hydrogen (up to 0.1 wt%
H) and are enriched in deuterium compared to UOCs of higher petrologic types (Alexander
et al. 2012; Vacher et al. 2020; Grant et al. 2023; Table SM10, Fig. SM54a). The least meta-
morphosed LL3.00 Semarkona has the highest amount of hydrogen among non-weathered
UOCs (average H content of 0.07±0.04 wt%) and the highest and more variable δD value
of 2000±745‰ (Table SM10).
Metasomatized CVs, COs, and CKs are depleted in hydrogen compared to CIs, CMs, and CRs that experienced relatively low temperature aqueous alteration and present moderate δD variations with values ranging from -115 to +240‰, with the exception of one aliquot of Ornans (CO3.4) having a δD of +2150‰ (Kerridge 1985) (Table SM10 and references therein, Fig. SM54b).

### 7.2.2 Hydrogen-Isotope Compositions of Organics

As described in Sect. 3, the IOM is the major carrier of C, N and noble gases, and one of the two major carriers of hydrogen in type 3 chondrites. In the most primitive chondrites, the IOM is strongly enriched in δD compared to bulk rocks (Alexander et al. 2007, 2012, 2021) and shows heterogeneities with µm-sized areas both enriched and depleted in D compared to the average IOM value (Busemann et al. 2006; Remusat et al. 2006). With increasing temperatures during parent body alteration, the H/C ratio of the IOM decreases due to the re-organization of the organic matter structure, the loss of heteroatoms, and the development of large aromatic domains (see Sect. 3.1; Alexander et al. 2007, 2014, 2017, 2009). In carbonaceous chondrites, the hydrogen loss is associated with the destruction of µm-scale heterogeneities (Remusat et al. 2008) and likely to H-isotope fractionation (Alexander et al. 2010, 2014).

In CVs and COs, the IOM shows low-to-moderate D-enrichments relative to the bulk rocks compared to the more water-rich CIs, CMs, and CRs with δD ranging from +60 to +1400‰ and from +130 to +700‰, respectively (Alexander et al. 2007). Although there is only moderate D-increase for the CV IOM and no clear trend exists for the CO IOM δD compositions with petrological type or H contents (Alexander et al. 2007), the relatively low δD and H contents of these meteorites compared to those of CMs, CIs, and CRs could reflect the effects of parent body processes under increasing temperatures.

In UOCs, the IOM is enriched in δD (from +3000 to 12,000‰) (Alexander et al. 2010) compared to those of other chondrites and to the UOC bulk compositions (Yang and Epstein 1983; Alexander et al. 2007, 2010, 2021). However, the IOM appears to be more homogeneous than in CCs showing only rare µm-sized D-rich anomalies (Remusat et al. 2016). The presence of a positive correlation between the IOM δD and its C/H ratio (Fig. SM55a) was initially proposed to result from isotopic exchanges between IOM and aqueous fluids following D enrichments of water through Rayleigh-type fractionations after metal oxidation during alteration processes in the parent asteroid (Alexander et al. 2010). However, this hypothesis is at odds with the absence of general D-enrichments in bulk UOCs with increasing the metamorphism degree (Fig. SM55a; Vacher et al. 2020; Marrocchi et al. 2020; Piani et al. 2021; Grant et al. 2023). Moreover, the IOM of ordinary chondrites seems to contain a D-rich thermally resistant component, whose proportions would be enhanced during heating due to the loss of more D-poor organic material (Figs. SM55b,c; Remusat et al. 2016). The presence of this D-rich component in UOCs that is unsampled by carbonaceous chondrites argues for a difference in the organic reservoirs of OC and CC IOMs.

### 7.2.3 Hydrogen-Isotope Compositions of Hydrated Silicates

Like for CMs, CIs and CRs (for details see Lee et al. this journal), the δD composition of water-bearing minerals in the oxidized Bali-like CVs (Kaba, Bali, and Grosnaja) was estimated by SIMS with an analytical protocol using Cs⁺ primary ions and found to share a common δD composition of ~ -80±100‰ (Piani and Marrocchi 2018). The D-enrichment of the CV water-bearing minerals compared to those of CMs can be interpreted as the evidence that CVs accreted ices with a distinct hydrogen-isotope composition than those in
CMs or CRs. However, the additional presence of generally δD-depleted organic matter in
CVs compared to that in CMs could also be seen as the result from isotopic exchanges be-
tween water and organics during parent body metasomatism (Piani and Marrocchi 2018).
The exchange reactions could be favored in CVs compared to CMs by the higher T and/or
different alteration mechanisms, possibly implying steam instead of, or in addition to, liquid
fluids (see Sect. 5).

For COs, the δD of water-bearing minerals was measured only in DOM 08006 (CO3.00) and found to be +306±135‰ (Piani et al. 2021), a value slightly higher than for CVs. More measurements in this group of chondrites would be needed to determine if this unique value is representative for the water in all COs; only meteorites with low terrestrial weathering degrees should be selected for analyses to avoid terrestrial contamination (Alexander et al. 2018a).

Highly heterogeneous δD with extreme values up to 10,000‰ have been measured in the H-bearing matrix of Semarkona (LL3.00) (Fig. SM56; Deloule and Robert 1995; Piani et al. 2015). The highest δD values in the matrix exceed by far the isotopic composition observed in primitive OC IOMs (Alexander et al. 2007; Remusat et al. 2016) and are not associated with C-rich areas indicating that their carrier is inorganic, likely hydrous silicates (Fig. SM56; Piani et al. 2015). The identification of the extreme δD variations associated with inorganic phases in Semarkona indicate that the parent body of LL-type ordinary chondrites have likely accreted some D-rich material derived from the molecular cloud and having escaped complete homogenization with hydrogen of the PPD (Aléon 2010; Piani et al. 2021). The nature of the D-rich inorganic carrier, possibly phyllosilicates or hydrated amorphous silicates (Dobrică and Brearley 2021), has yet to be characterized.

7.3 Carbon- and Nitrogen-Isotope Compositions of the Main C- and N-Bearing Phases

The carbon-isotope compositions of carbonates in UOCs, CVs, and COs (δ¹³C range from -23 to +5‰) were determined from CO₂ released after phosphoric acid treatments of the whole rocks and found to be ¹³C-depleted compared to carbonates in CIs, CMs, and CRs (Grady et al. 1988; Alexander et al. 2015). The limited quantities of carbon measured (e.g., 0.02 wt% C in Semarkona; Alexander et al. 2015) are consistent with the absence or scarcity of carbonates observed in these chondrites (e.g., Hutchison et al. 1987). The low δ¹³C values of carbonates in CVs, COs and UOCs could be explained either by their formation at higher temperature and/or by more oxidized fluids (higher CO₂ fractions) than for the CIs, CMs, and CRs (Alexander et al. 2015).

In UOCs, the C-isotope compositions of organic matter were estimated using stepped heating combustion (Grady et al. 1989). Removing the terrestrial weathering contaminants that can constitute an important part of the low temperature released gases of chondrite finds, the δ¹³C of the indigenous organic compounds were estimated to be of ~-28‰ with no clear variation trend observed according to the chondrite petrographic type (Grady et al. 1989).

The carbon- and nitrogen-isotope compositions of CV and CO IOM range from -18 to -5‰ for δ¹³C and from -65 to +17‰ for δ¹⁵N (Alexander et al. 2007). Due to the high abundance of IOM among the known C-bearing components in CVs and COs, the bulk δ¹³C and δ¹⁵N values of CVs and COs appear to be dominated by the IOM (Greenwood and Franchi 2004; Alexander et al. 2015).

Although FeNi-carbides are commonly observed in the least metamorphosed COs, CVs, and UOCs (see Sect. 2.2), their C-isotope compositions are not known yet.
### 7.3.1 Future Studies

Trace amounts of carbonates appear to be present in CVs, COs and UOCs when measured in whole rocks (Grady et al. 1988), but no *in situ* measurement were ever reported for their carbon-isotope compositions. Such measurements should be possible with SIMS at least in Semarkona (LL3.00), in which 15 µm-sized calcites were observed (Hutchison et al. 1987) and would allow a better understanding of the conditions of secondary alteration. In addition, SIMS measurements of C-isotope compositions of FeNi-carbides in CVs, COs, and UOCs would be also important.

Although the organic matter (including both SOM and IOM) and carbonates are major carbonaceous constituents in chondrites, their reported abundances do not account for more than ~60% of the total carbon content of the rock (Smith and Kaplan 1970; Alexander et al. 2015). Investigations should aim to identify the exact nature and origin of the missing carbonaceous component, which possibly consists of some kind of organic compounds lost during the IOM isolation process (Alexander et al. 2015).

Finally, the nature of the inorganic H-bearing components at the origin of the D-rich signatures observed in UOCs and the amplitude of their modifications during parent body processes require thorough investigations that could be achieved by coupling high resolution mineralogical observations using TEM and isotope imaging using NanoSIMS.

## 7.4 Sulfur Isotopes

With four stable isotopes ($^{32}$S, $^{33}$S, $^{34}$S and $^{36}$S), sulfur can be used as a probe of the astrophysical environment of mass-independent photochemistry (Thiemens 2006). The sulfur isotope ratios are usually expressed as $\delta^{33}$S, $\delta^{34}$S, and $\delta^{36}$S ($\delta^{n}$S = [($^{n}$S/$^{32}$S)/($^{n}$S/$^{32}$S$_{V-CDT}$)]$^{n}$) × 1000; with n being the mass of the considered isotope and V-CDT: Vienna Canyon Dablo Troilite used as reference. Briefly, the extents of mass-independent anomalies for the different isotopes are thought to depend on the energy of the photons responsible for the photolysis of the major S-bearing phase in the gas of the solar nebula, H$_{2}$S (Chakraborty et al. 2013).

The highest mass-independent variations of sulfur reported so far in chondrites have been found to be associated with the $^{16}$O-poor COS in Acfer 094 (weighted means of $\Delta^{33}$S = +3.84±0.72‰ and $\Delta^{36}$S = -6.05±2.25‰ with $\Delta^{33}$S = $\delta^{33}$S+1000×[(($\delta^{33}$S/1000+1)$^{0.515}$-$^{1}$] and $\Delta^{36}$S = $\delta^{36}$S+1000×[($\delta^{36}$S/1000+1)$^{1.9}$-$^{1}$]; Vacher et al. 2021). These values are inconsistent with the energy of ultraviolet-photons coming from the protosun and could imply the presence of massive O and B stars close to the solar system's parent molecular cloud (Vacher et al. 2021). Except in COS of Acfer 094, no significant mass-dependent or mass-independent isotopic variations were reported for type 3 chondrites so far (Tachibana and Huss 2005; Marrocchi et al. 2016; Alexander et al. 2022).

Large mass-dependent isotope fractionation effects ($\delta^{34}$S from -8 to +6‰) have been described in aqueously formed sulfides from CIs, CMs, and volatile-rich clasts in polymict ureilites and HED meteorites (Visser et al. 2019). Aqueously formed magnetite and fayalite in chondrules and matrices of UOCs, CVs, and COs commonly associate with elongated sulfide (troilite or pyrrhotite?) grains, which may have formed by direct precipitation from aqueous fluid (Fig. SM57). Sulfur-isotope compositions of sulfides associated with Cr-free magnetite in Kaba chondrules and interpreted as igneous in origin (Marrocchi et al. 2016; but see Sect. 2.2 for an alternative explanation) show a small range of $\delta^{34}$S, from -1±0.4‰ to +1.4±0.4‰. Elongated sulfides associated with fayalite have yet to be analyzed.
## 7.5 Sources of Water in UOCs, COs, and CVs

The main process controlling the O-isotope composition of the CV and CO parent asteroids is related to isotopic exchange between ¹⁶O-enriched anhydrous silicates and ¹⁷,¹⁸O-enriched aqueous fluid (Fig. SM58; Marrocchi et al. 2018). The initial O-isotope composition of aqueous fluid and its abundance (water/rock ratio) in the CV and CO asteroids, however, are poorly constrained. It is commonly assumed that COS in Acfer 094 having anomalously ¹⁷,¹⁸O-enriched compositions (δ¹⁷,¹⁸O ~ +180‰; Δ¹⁷O ~ +85‰; Sakamoto et al. 2007) and sulfur mass-independent isotope anomalies (Δ³³S ~ +4‰ and Δ³⁶S ~ -6‰; Vacher et al. 2021) recorded the primordial O-isotope composition of water ice in the PPD providing evidence for the CO self-shielding in the PMC and/or outer disk (Yurimoto and Kuramoto 2004; Lyons and Young 2005; Krot et al. 2020). The ¹⁷,¹⁸O-depleted compositions of aqueously formed minerals in UOCs and CCs (see Sects. 7.1.2 & 7.1.3, and Lee et al. 2025, this collection) provide no evidence that such primordial water ice directly accreted into chondrite parent bodies, and suggest instead that prior to accretion into chondrite parent asteroids O-isotope composition of water ice evolved towards the Earth-like value (Δ¹⁷O ~ 0‰) through either partial evaporation and isotope exchange with ¹⁶O-enriched silicates and CO gas or/and isotopic evolution of water with time during CO self-shielding in the outer disk (Yurimoto and Kuramoto 2004; Lyons and Young 2005; Lyons et al. 2009; Krot et al. 2015; Alexander et al. 2017; Marrocchi et al. 2018).

Once accreted and melted, the chondritic water evolved towards ¹⁶O-rich composition by exchange with anhydrous minerals and amorphous phases (Marrocchi et al. 2018; see also Lee et al. 2025, this collection). The initial water ice that accreted into chondrite asteroids is thus expected to have had a Δ¹⁷O value equivalent to or higher than any aqueously formed minerals present in chondrites (Krot et al. 2015; Fujiya 2018; Marrocchi et al. 2018; Alexander 2019a,b). In the case of weakly metamorphosed CO3.0–3.1 s and CV<sub>oxB</sub>3.1 s, the highest Δ¹⁷O values of aqueously formed minerals ranging from ~ -2 to ~ +3.5‰ (Figs. SM44–45) put the lower limits on the initial Δ¹⁷O of water ice in COs and CVs. Alexander (2019b) suggested that (i) different carbonaceous chondrites, including CVs and COs, could have originated from a mixture of the same four components: (1) chondrule/chondrule precursor component that partially lost FeNi-metal and volatiles, (2) refractory inclusion component that has a refractory inclusion-like bulk composition, (3) anhydrous and reduced but otherwise CI-like matrix, and (4) water ice with relatively high Δ¹⁷O (+3.5‰) and δ¹⁸O (18–21‰). (ii) Variations in the amount of water ice accreted by different chondrite groups could be responsible for differences in their bulk O-isotope compositions and O-isotope compositions of aqueously formed minerals. For example, the ¹⁶O-enriched bulk compositions of COs compared to CVs would result from lower amount of the water ice component accreted by COs (Alexander 2019b). Nonetheless, the observed differences in Δ¹⁷O values of fayalite in Kaba (CV<sub>oxB</sub>3.1) and Elephant Moraine (EET) 90043 (CO3.1), ~ +3.5‰ and -0.3‰, respectively (Figs. SM45a, 44c), may suggest that CVs accreted a more ¹⁷,¹⁸O-rich water ice than COs (Marrocchi et al. 2018). Using the estimated O-isotope composition of the CV<sub>oxB</sub> aqueous fluids during precipitation of pure fayalite and the inferred water/rock ratio modeled for its formation (0.1–0.2; Zolotov et al. 2006), Marrocchi et al. (2018) calculated the initial O-isotope composition of the CV<sub>oxB</sub> aqueous fluid (δ¹⁸O ≥ +60‰, Δ¹⁷O > 15‰) suggesting a potential contribution of water ice from the outer solar system.

Hydrogen-isotope compositions estimated for hydrated minerals in CVs could also be the results of the parent body evolution from an initially D-poor CM-like water toward D-richer values through isotope exchanges with D-rich organics that would be favored by the high alteration temperatures of CV parent body (Piani and Marrocchi 2018). Alternatively,
the moderate δD values measured for both water and IOM in CVs and COs could have resulted from their early accretion in the disk at temperatures allowing a more efficient re-equilibration with the D-poor solar gas than for CIs, CMs, and CRs that accreted later or at higher heliocentric distances (Piani et al. 2021). Further isotopic measurements on natural samples and experimental constraints on isotope exchange reactions are needed to better differentiate primary isotopic signatures from secondary isotope modifications taking place in the parent bodies during metasomatic alteration.

As for weakly metamorphosed CVs and COs, the Δ¹⁷O of ices accreted in the UOC parent bodies should be equal or higher than the highest value inferred from the alteration fluids at the origin of the secondary minerals, fayalite and magnetite (i.e., Δ¹⁷O ~ +5‰; Fig. SM43). However, the ranges of δ¹⁸O and Δ¹⁷O measured in bulk UOCs cannot be reconciled with a similar water ice O-isotope composition as those suggested for CCs, and Δ¹⁷O values as high as ~16‰ were proposed for the OC initial water ice (Alexander 2019a). Additionally, the presence of D-rich IOM, including a D-rich thermally resistant component (Sect. 7.2.2), and the D-rich and isotopically heterogeneous hydrated silicates measured in Semarkona (Sect. 7.2.3) support the interpretation that OCs inherited a higher fraction of the primordial (i.e., molecular cloud) water ice and/or organics compared to CCs (Deloule and Robert 1995; Aléon 2010; Piani et al. 2015; Remusat et al. 2016). The inferred Δ¹⁷O and δD values of water ice and δD composition of IOM seem to indicate that OCs have recorded lower degrees of re-equilibration between the primordial molecular cloud molecules and protosolar H₂ than most CCs, which might be at odds with their supposed formation locations in the inner and outer solar system, respectively (Warren 2011; Kruijer et al. 2017). However, the temperatures regulating the efficiency of re-equilibration between the primordial molecules and protosolar H₂ are not only dependent on the formation distance of the chondritic parent body but also on the thermal evolution of the disk with time and the nature of interactions with the parent molecular cloud. Strong differences exist in models trying to infer the H-isotope evolution of the disk water (see the discussion in Krot et al. 2015 and references therein); with no systematic dependences on the heliocentric distance (Yang et al. 2013; Jacquet and Robert 2013). Nonetheless, although the re-equilibration timescale could be different for oxygen and hydrogen isotopes (Lyons and Young 2005; Yang et al. 2013), no model allows a comprehensive description of the O- and H-isotope distribution for water throughout the disk evolution, nor does it explain the isotopic differences between OCs and CCs. Further attempts should be made in the future (i) to bring reliable constraints on isotope distribution in planetary materials, notably by deconvoluting the isotopic contribution of parent body processes from the hydrogen- and oxygen-isotope signatures of the primary ice in chondrites, and (ii) to estimate the spatial/temporal evolution of water hydrogen and oxygen-isotope compositions in the context of the recent disk dynamical evolution models.

## 8 Conclusions

1. Type 3 ordinary and CV, CO, and CK carbonaceous chondrites experienced aqueous fluid – rock interaction at ~100–300 °C during progressive thermal metamorphism followed by loss of the fluid and nearly dry metamorphism. Both aqueous solution and steam played an important role in mobilization of rock-forming elements and resulted in metasomatic alteration of these meteorites. The alteration was multistage and produced different minerals and mineral assemblages in chondrite matrices, chondrules, and refractory inclusions reflecting differences in primary mineralogy, textures, and chemical compositions of these chondritic
components as well as temporal and local variations in fluid chemical composition. Detailed mineralogical studies of the secondary minerals and mineral assemblages in type 3 chondrites of different petrologic subtypes using FE SEM and focused ion beam (FIB) + TEM combined with physicochemical analysis and thermodynamic calculations could potentially constrain conditions during metasomatic alteration (P − T − fO₂-water/rock ratio-aSiO₂ − aCa − aFe...).

2. Experimental syntheses of secondary minerals (grossular, wollastonite, nepheline, sodalite, dmisteinbergite) observed in metasomatically altered CAIs in Allende (CV<sub>oxA</sub>>3.6) resulted in formation of intermediate hydrated phases (e.g., hydrogrossular, hydrosodalite, analcime, nepheline hydrate) not observed in Allende. Studies of CAIs from less metamorphosed CVs are required to understand how relevant these experiments are. Although a hydrothermal experiment with Fe-metal, amorphous ferroan silicates and water did produce pure fayalite commonly observed in UOCs, CVs and COs of petrologic type 3.0–3.2, an experimental synthesis of FeMg-olivine from amorphous FeMg-silicates commonly observed in matrix of primitive chondrites is needed to understand a broad range of matrix olivine compositions (Fa<sub>~40–100</sub>) in these meteorites.

3. The origin of the precursors of organic matter in different chondrite groups and whether CVs, COs, and UOCs accreted the same mix of organic precursors remain unclear. Although IOM, SOM, and carbonates are the major carbon-bearing constituents in chondrites, their reported abundances do not account for more than ~60% of the chondrites total carbon content. Identification of the exact nature and the origin of the missing carbonaceous component is urgently needed. The role of metasomatic fluids in the evolution of organic matter in UOCs, COs, and CVs is not known. *In situ* studies of IOM using STXM and TEM and experimentally identifying specific reaction pathways related to the presence of oxidizing fluids at temperature above 300 °C could potentially offer a way to shed some light on this problem.

4. The <sup>53</sup>Mn-<sup>53</sup>Cr chronology of fayalite and kirschsteinite in UOCs, CVs, and COs measured with SIMS using matrix-matched standards suggests metasomatic alteration in their parent asteroids occurred ~2–5 Ma after formation of CAIs. However, the obtained data are very limited and do not allow the duration of metasomatic alteration in the UOC, CO, and CV parent asteroids to be constrained. The <sup>129</sup>I-<sup>129</sup>Xe ages of magnetites separated from CV3s, CO3s, and CK3-5 s are similar to those from CMs and CIs, ~3-5 Ma after CAIs, suggesting nearly contemporaneous aqueous activity in the carbonaceous chondrites parent asteroids. The <sup>129</sup>I-<sup>129</sup>Xe ages of polymineralic samples (matrix, chondrules, CAIs, dark inclusions) from CV chondrites (mostly from Allende) recorded prolonged duration of their metasomatic alteration, > 11 Ma. However, the <sup>129</sup>I-carrier phases and corresponding closure temperatures of I-Xe isotope systematics of these samples are not well-constrained yet. The <sup>53</sup>Mn-<sup>53</sup>Cr and <sup>129</sup>I-<sup>129</sup>Xe chronologies of metasomatic alteration and thermal evolution models of asteroids heated by <sup>26</sup>Al decay suggest that the CV, CO, and ordinary chondrite parent bodies accreted 2.4-2.7 Ma, 2.1-2.4 Ma, and 1.6-1.8 Ma after CV CAIs, respectively.

5. The <sup>26</sup>Al-<sup>26</sup>Mg systematics of secondary minerals having high Al/Mg ratio (anorthite, grossular, nepheline, sodalite) places only lower limits on their formation ages (> 3 Ma after formation of CAIs). The reported large excesses of radiogenic <sup>26</sup>Mg in some secondary minerals were probably inherited from the primary minerals after nearly complete decay of <sup>26</sup>Al and, thus, provide no chronological information. High precision Al-Mg isotope measurements of cogenetic secondary minerals (mineral parageneses) showing a large range of Al/Mg ratio (e.g., grossular, anorthite, monticellite, forsterite, corundum) can be used to construct internal Al-Mg isochrons and potentially date their formation.
6. The $^{26}$Al-$^{26}$Mg systematics of chondrule glasses can be disturbed in metasomatically altered chondrites even of low petrologic type (3.0–3.2) diminishing their chronological significance. Although chemically primitive chondrule plagioclase in some type 3.0–3.1 chondrites experienced aqueous/metasomatic alteration resulting in formation of phyllosilicates, there is no clear evidence that this process affected its $^{26}$Al-$^{26}$Mg systematics. Therefore, the observed ranges of ($^{26}$Al/$^{27}$Al)$_{0}$ in plagioclase-bearing chondrules from type 3.0–3.1 chondrites are probably chronologically meaningful (but see Piralla et al. 2023). In chondrites of higher petrologic type, the $^{26}$Al-$^{26}$Mg systematics of plagioclase in chondrules and refractory inclusions, and even of melilite in CAIs can be disturbed and should be considered with caution in chronological interpretations.

7. The $\Delta^{17}$O values of metasomatically formed minerals in UOCs, COs, CVs and CKs range from ~ +5 to ~ -4‰ suggesting differences in $\Delta^{17}$O of aqueous fluids and, possibly, water ices that accreted into their parent asteroids. Development of the matrix-matched mineral standards for SIMS measurements is required to constrain the $\delta^{18}$O values of aqueous fluids, and, potentially, crystallization temperature of secondary mineral parageneses. The fact that the inferred $\Delta^{17}$O values of aqueous fluids in the UOCs, COs, CVs, and CKs parent asteroids are distinctly different from the hypothesized $\Delta^{17}$O of primordial water ices recorded by cosmic symplectites in Acfer 094 (~ +80‰) suggests either extensive thermal processing of water ices in the PPD accompanying by isotopic exchange with $^{16}$O-enriched CO and anhydrous silicates or/and isotopic evolution of water with time during CO self-shielding in the outer disk.

8. Chondrule phenocrysts and plagioclase or glassy mesostasis as well as refractory inclu-
sions in the least metamorphosed (type 3.00) chondrites have internally uniform $\Delta^{17}$O val-
ues. Chondrules and refractory inclusions in UOCs, COs, CVs, and CKs of higher petrologic
type (> 3.0) have internally heterogeneous $\Delta^{17}$O: glassy mesostasis, plagioclase, melilite,
grossite, krotite, Ti-rich pyroxenes, perovskite and Sc- and Zr-rich oxides and silicates are
$^{16}$O-depleted relative to olivine, low-Ca pyroxene, hibonite, spinel, and low-Ti pyroxenes
which retained their initial O-isotope compositions. The $\Delta^{17}$O values of the $^{16}$O-depleted
phases approach the inferred $\Delta^{17}$O values of aqueous fluid suggesting that the observed O-
isotope heterogeneity most likely have resulted from isotopic exchange with aqueous fluids
during metasomatic alteration. Most of the $^{16}$O-depleted minerals show no evidence for dis-
turbance of their Al-Mg isotope systematics, suggesting that Mg self-diffusion coefficients
in these minerals under P - T - $f_{O_2}$ conditions of metasomatic alteration is much slower
than oxygen self-diffusion. Therefore, CAIs with the canonical ($^{26}$Al/$^{27}$Al)$_{0}$ and internally
heterogeneous O-isotope compositions do not necessarily provide evidence for significant
fluctuations in O-isotope composition of nebular gas in the CAI-forming region. At the same
time, isotopic exchange with nebular gas of variable O-isotope composition cannot be ex-
cluded. Experimentally defined oxygen self-diffusion in CAI- and chondrule-like minerals
under the inferred P - T - $f_{O_2}$-water/rock ratio conditions during metasomatic alteration
are required to test both hypotheses.

9. Carbon-isotope compositions of carbonates and FeNi-carbides in UOCs, CVs, and COs can potentially provide constraints on the origin of water ices in their parent asteroids, but are not known yet. These phases can be analyzed *in situ* by SIMS.

10. The main process controlling the O-isotope composition of the CV and CO par-
ent asteroids is related to isotopic exchange between $^{16}$O-enriched anhydrous silicates and
$^{17,18}$O-enriched aqueous fluid. The initial O-isotope of water ice and its abundance in the
CV and CO chondrite parent asteroids are poorly constrained: its $\Delta^{17}$O value must have
been equal or higher than the highest $\Delta^{17}$O of aqueously formed minerals in CV and CO
chondrites, i.e., $> +3.2\%$ and $> -0.3\%$, respectively. The calculated water/rock ratios in
CV and CO asteroids are 0.1–0.2 and 0.01–0.1 (Marrocchi et al. 2018); these calculations, however, are highly model-dependent.

Parent asteroids of ordinary chondrites appear to have accreted low abundance of water ice, possibly, because of Jupiter blocking the inward flow of ice (Morbidelli et al. 2016), but with higher $\Delta^{17}$O value than that in CVs and COs, $\ge$ +5‰. These observations and the presence of D-rich IOM, including a D-rich thermally resistant component, and the D-rich and isotopically heterogeneous hydrated silicates measured in Semarkona (LL3.00) may indicate that (i) UOCs inherited a higher fraction of the primordial (molecular cloud) water ice and/or organics compared to CCs, and (ii) UOCs recorded lower degrees of re-equilibration between the primordial molecular cloud molecules and protosolar H$_2$ than most CCs. This interpretation appears to contradict the assumed formation locations of OCs and CCs in the inner and outer solar system (Warren 2011; Kruijer et al. 2017). Further attempts should be made to deconvolute the isotopic contribution of parent body processes from the hydrogen and oxygen isotopic signatures of the primary ice in chondrites, and to estimate the spatial/temporal evolution of water hydrogen- and oxygen-isotope compositions in the context of the recent disk dynamical evolution models (Desch et al. 2018; Burkhardt et al. 2019; Jacquet et al. 2019; Kleine et al. 2020).

**Supplementary Information** The online version contains supplementary material available at https://doi.org/10.1007/s11214-025-01135-z.

**Funding** Open Access funding enabled and organized by Projekt DEAL.

## Declarations

**Competing Interests** The authors declare that they have no conflict of interest.

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