# Preaccretionary origin of clay minerals in CI chondritic objects? Insights from the Orgueil clay mineralogy and iron oxidation state

J.-C. Viennet, K. Amano, L. Delbes, B. Lavina, F. Hubert, S. Bernard, S.

Yaroslavtsev, D. Bessas, A. Chumakov, J.-P. Celse, et al.

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► To cite this version:

J.-C. Viennet, K. Amano, L. Delbes, B. Lavina, F. Hubert, et al. Preaccretionary origin of clay minerals in CI chondritic objects? Insights from the Orgueil clay mineralogy and iron oxidation state. Earth and Planetary Science Letters, 2025, 671, pp.119678. 10.1016/j.epsl.2025.119678. hal-05337094

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# Preaccretionary origin of clay minerals in CI chondritic objects? Insights from the Orgueil clay mineralogy and iron oxidation state

J.-C. Viennet<sup>a,*</sup>, K. Amano<sup>a,b</sup>, L. Delbes<sup>a</sup>, B. Lavina<sup>b,c</sup>, F. Hubert<sup>d</sup>, S. Bernard<sup>a,b</sup>, S. Yaroslavtsev<sup>e</sup>, D. Bessas<sup>e</sup>, A. Chumakov<sup>e</sup>, J.-P. Celse<sup>e</sup>, M. Gounelle<sup>a</sup>, L. Remusat<sup>a,b</sup>, M. Roskosz<sup>a</sup>

## ARTICLE INFO

Edited by: Dr O Mousis

**Keywords:**

CI Materials
Orgueil meteorite
Phyllosilicates
X-Ray diffraction
Mössbauer
Redox ratio

## ABSTRACT

Phyllosilicates are the main phases in carbonaceous Ivuna-type (CI) materials and major carriers of water (structural OH) as well as of some organic matter adsorbed on their surfaces. Understanding their formation is therefore crucial for reconstructing the history of water and carbon in the early Solar System. While typically interpreted as products of extensive aqueous alteration under asteroidal conditions, their presence in CI chondrites, whose bulk composition matches that of the solar photosphere, remains a paradox.

To investigate their formation processes, we characterized the crystal structure and proportion of phyllosilicates in the Orgueil CI meteorite using X-ray diffraction (XRD) profile modeling on oriented preparations as well as Fe valences by Mössbauer spectroscopy at room and helium temperatures. The XRD modeling confirms the presence of discrete smectite and serpentine, along with two randomly stacked (R0) mixed-layer minerals (MLMs): serpentine-mica (90 %-10 %) and serpentine-smectite-mica (50 %-30 %-20 %), the latter being dominant across all size fractions. Mössbauer spectroscopy reveals that discrete serpentine and the R0 serpentine-mica MLM are Fe²⁺-rich, while the discrete smectite and the R0 serpentine-smectite-mica MLM contain 80 % Fe³⁺.

The coexistence of these distinct Fe valences, the unusual serpentine-smectite-mica clay structure, the presence of mica layers within MLMs, and their structural homogeneity across size fractions appear to be incompatible with low-temperature aqueous alteration at equilibrium. Instead, they support a pre-accretionary origin for phyllosilicates, likely involving gas-grain interactions or condensation process.

## 1. Introduction

Among the rocky objects in the Solar System, Ivuna-type carbonaceous materials (CI) stand out as particularly significant. The CI chondrites are considered the most chemically primitive rocks in the Solar System as indicated by their bulk compositions, which closely resemble that of the solar photosphere (Lodders, 2003; Asplund et al., 2009). Recent analyses of terrestrially unaltered fragments from the asteroid Ryugu, returned by the Hayabusa2 mission, reveal chemical, isotopic, and mineralogical properties akin to those of CI meteorites, providing the best match to the Solar System bulk composition (Yokoyama et al.,

2022; Ito et al., 2022; Nakamura et al., 2022). Similarly, preliminary results from the OSIRIS-REx mission suggest that Bennu samples also closely align with this composition (Lauretta et al., 2024).

The CI materials are also mineralogically and petrographically distinct from other chondrites. They lack chondrules and refractory inclusions (Krot et al., 2014) and contain rare anhydrous silicates, such as olivine and pyroxene (Aléon-Toppani et al., 2024; Alfing et al., 2019; Nakamura et al., 2022). Instead, CI materials are mainly composed of phyllosilicates (a mixture of serpentine and smectite, >80 Vol. %), magnetite (< 10 Vol. %), sulfides (~ 5 % Vol. %) and carbonates (dolomite, calcite, breunnerite; < 5 Vol. %) (King et al., 2015; Nakamura

<footnote><sup>a</sup> Muséum National d'Histoire Naturelle, Sorbonne Université, UMR CNRS 7590, Institut de minéralogie, de physique des matériaux et de cosmochimie, Paris, France</footnote>

<footnote><sup>b</sup> Center for Advanced Radiation Sources, The University of Chicago, Chicago, IL 60637, USA</footnote>

<footnote><sup>c</sup> Advanced Photon Source, Argonne National Laboratory, Argonne, IL 60439, USA</footnote>

<footnote><sup>d</sup> Institut de Chimie des Milieux et Matériaux de Poitiers (IC2MP), UMR 7285 CNRS, Université de Poitiers, F-86073 Poitiers Cedex 9, France.</footnote>

<footnote><sup>e</sup> European Synchrotron Radiation Facility, F-38043 Grenoble, France</footnote>

<footnote>* Corresponding author at: Muséum National d'Histoire Naturelle, Sorbonne Université, CNRS, UMR 7590, Institut de Minéralogie, Physique des Matériaux et Cosmochimie, Paris, France.</footnote>

<footnote>E-mail address: jean-christophe.viennet@mnhn.fr (J.-C. Viennet).</footnote>
et al., 2022). Sulfates and ferrihydrite (< 10 Vol. %) in CI meteorites are also present and are interpreted as products of terrestrial alteration, resulting from Fe-sulfide oxidation and carbonate dissolution (Gounelle and Zolensky, 2001; King et al., 2020; Roskosz et al., 2024; Leroux et al., 2024).

Among all carbonaceous objects, CI materials are the most water-rich, with water primarily hosted as hydroxyl groups (OH) in the structure of phyllosilicates (Yokoyama et al., 2022). They also contain the highest concentrations of organic matter, some of which is adsorbed onto the interlayer surfaces of phyllosilicates (Viennet et al., 2023; Bahae-eddine et al., 2024). Consequently, understanding the formation history of phyllosilicates in CI materials is essential for reconstructing the history of water and carbon in the Solar System.

The phyllosilicates in CI materials are commonly interpreted as the products of extensive aqueous alteration of anhydrous silicates, occurring in a closed system at equilibrium within the parent body of CI chondrites and carbonaceous (C type) asteroids (Zolensky et al., 1989; Brearley, 2006; Zolotov, 2012; Bland and Travis, 2017; Nakamura et al., 2022). This alteration is thought to have occurred in closed system at equilibrium because of the chemically primitive nature of CI chondrites (i.e., 1:1 match with solar photospheric abundances - Lodders, 2003) and at low temperatures (< 100°C), as evidenced by the absence of Mg-chlorite (Nakamura et al., 2022). Other studies, such as Clayton and Mayeda, (1984), based on a model using CM2 oxygen isotopic compositions, proposed a slightly higher temperature range (100-150°C) for the alteration processes.

However, petrological evidence challenges the hypothesis of low temperature aqueous alteration in a closed system at equilibrium. Tomeoka and Buseck (1988) suggested that fine phyllosilicates formed through the aqueous alteration of coarse-grained ones. Studies of Ryugu grains (Aléon-Toppani et al., 2024; Leroux et al., 2024), propose a two-stage alteration process in which fine phyllosilicates may originate from the alteration of GEMS (glass with embedded metal and sulfides, fine-grained aggregates) or from nanometer-sized assemblages of amorphous silicates and Fe-based nanophases. The nature of the primary silicates from which coarse phyllosilicates formed was not discussed by Tomeoka and Buseck (1988). LeRoux et al. (2024) hypothesized that coarse phyllosilicates formed via the replacement of anhydrous crystalline silicates, such as olivine and pyroxene, whereas Aléon-Toppani et al. (2024) suggested that they precipitated in fractures, veins, or pores under more oxidative conditions involving Mg-rich anhydrous minerals. Additionally, anhydrous silicates, even at the micrometer scale, do not show features of aqueous alteration (Reid et al., 1970; Alfing et al., 2019; Nakamura et al., 2022; Kawasaki et al., 2022; Aléon-Toppani et al., 2024). This has been attributed either to varying degrees of aqueous alteration in highly localized environments (i.e., anhydrous particles that escaped alteration) or to the incorporation of chemically distinct particles into the parent body surface materials through impact processes or low-temperature mechanical mixing.

Altogether, these observations raise critical questions about the processes of phyllosilicate formation in CI materials. While a global equilibrium alteration model has been widely accepted, growing evidence suggests that alteration occurred in multiple stages within localized microenvironments (Reid et al., 1970; Tomeoka and Buseck, 1988; Alfing et al., 2019; Kawasaki et al., 2022; Nakamura et al., 2022; Aléon-Toppani et al., 2024; Leroux et al., 2024). However, the nature of the primary minerals, the alteration conditions (e.g., temperature, redox state, and water-rock ratios), and the underlying mechanisms remain poorly constrained. Consequently, reconciling the highly altered mineralogy of CI chondrites with their status as the most chemically primitive chondritic materials remains a paradox in cosmochemistry (Lodders, 2003; Brearley, 2006; Weisberg et al., 2006; Krot et al., 2014).

As phyllosilicates constitute the dominant mineral phases in CI materials (> 80 Vol. % - King et al., (2015)), refining their crystal chemistry is crucial. This includes determining the nature of individual layers, their stacking mode, and their relative proportions within a crystal to

better understand their origin. Of particular importance is the occurrence of mixed-layer minerals or nanoscale intergrowths of serpentine and smectite, which have been reported at the nanoscale (Bass, 1971; Tomeoka and Buseck, 1988; Leroux et al., 2024).

Classical X-ray diffraction (XRD) methods applied to CI meteorites by King et al., (2015) cannot resolve the proportions of serpentine and smectite layers or the stacking sequence within phyllosilicates. Here, we overcome this limitation by modeling the XRD profile obtained from oriented preparations of the Orgueil CI meteorite, which were previously disaggregated and size-sorted in decreasing granulometry. This approach has proven effective for quantitatively determining the crystal chemistry of clay minerals, as demonstrated in soil science (Hubert et al., 2012; Viennet et al., 2015; Bakker et al., 2019) and clay-rich sedimentary/diagenetic rocks (Sakharov and Lanson, 2013 and references therein). In addition, room- and helium-temperature Mössbauer measurements were used to determine the Fe valence in these clay minerals, providing new insights into their redox conditions of formation (see sections 3.4, 4.2 and 4.3). This dual approach enables us to unravel the complex structure of Orgueil clay minerals, which cannot be explained by low-temperature aqueous alteration of anhydrous minerals or amorphous phases on the parent body. Instead, our results consistently point to a pre-accretionary origin of the phyllosilicates in Orgueil and possibly other CI materials, either through condensation or nebular gas-grain interactions (see section 4.3).

## 2. Materials and Methods

### 2.1. Meteorite of Orgueil

A 500 mg sample of the Orgueil CI meteorite was extracted from specimen #234 housed at the Muséum National d'Histoire Naturelle, France. The sample consists of several millimeter-sized grains.

### 2.2. Sample preparation

The sample was first disaggregated and dispersed through five Na-saturation cycles using a 1 mol·L⁻¹ NaCl solution combined with sonication. It was then washed by dialysis in pure water (Spectra/Por® Membrane, MWCO: 1 kDa). After washing, the >50 µm fraction was isolated by sieving the Na-saturated sample, dried at 50°C, and stored in a refrigerator. The <50 µm fraction was subsequently separated into five sub-fractions (<0.02, 0.02-0.05, 0.05-0.2, 0.2-2, and 2-50 µm) following a centrifugation protocol developed in soil science (Hubert et al., 2012; Viennet et al., 2015; Bakker et al., 2019). Each sub-fraction was collected through repeated centrifugation cycles until the supernatant became clear (9 steps). Finally, all sub-fractions were Ca-saturated through five saturation cycles using a 1 mol·L⁻¹ CaCl₂ solution and washed in pure water by dialysis (Spectra/Por® Membrane, MWCO: 1 kDa). The sub-fractions were then dried and weighted.

### 2.3. X-ray diffraction analysis and profile modeling of 00l reflections

#### 2.3.1. XRD acquisition

Because *hk* reflections (i.e., those corresponding to Miller indices) are similar among different clay minerals, the identification of clay mineral phases relies mainly on their 00l reflections. These 00l reflections correspond to the interlayer distances along the *c* axis, enabling the identification of the nature of the layers. For this, oriented preparations are essential because they promote the preferential deposition of clay minerals on their basal surfaces, thereby enhancing the 00l X-ray diffraction intensities in the geometry used. As a result, a series of 00l reflections, free of *hk* reflections, are obtained with order of magnitude stronger intensities as compared with randomly oriented powder preparations.

To further distinguish clay minerals natures, additional procedures, including Ca²⁺ saturation of the interlayer cations under ambient
conditions and ethylene glycol solvation, were conducted. Ca-saturation ensures a uniform interlayer cationic composition, resulting in consistent interlayer distances for smectite layers under specific measurement conditions. Ethylene glycol solvation induces characteristic shifts in 00l reflection of smectite layers, aiding in the precise identification of clay mineral phases.

The Ca-saturated sub-fractions were first dispersed in distilled water, and oriented preparations were obtained by pipetting suspensions onto silicon wafers, followed by drying at room temperature. The air-dried preparations (hereafter referred to as Ca-AD specimens) were initially analyzed by X-ray diffraction (XRD). Ethylene glycol solvation was subsequently performed by exposing the same slides to ethylene glycol vapor for 48 hours (preparations referred to as Ca-EG), followed by XRD analysis. XRD patterns were acquired using a Panalytical X'pert Pro MPD two-circle diffractometer operated at IMPMC (Paris, France). The instrumental settings included a divergence slit of 0.25°, anti-scatter slit of 2.3°, Soller slits of 2.3° and 0.5 rad, and the 2θ angular range of 2-65° was scanned using CoKα₁,₂ radiation (step size: 0.033°2θ; counting time per step: 300 s). To ensure data reliability and improve the signal-to-noise ratio, measurements were repeated three times, and the final XRD patterns were calculated as the average of the three scans.

### 2.3.2. XRD profile modeling of 00l reflections

The Sybilla software (Chevron™; Aplin et al., 2006) was used to fit experimental XRD profiles of 00l reflections (3.5-40°2θ CuKα1,2) through trial-and-error. For software compatibility, experimental CoKα₁,₂ patterns were converted to CuKα₁,₂ radiation. This fitting method (Sakharov and Lanson, 2013), has been widely applied to characterize clay minerals in diverse geological contexts, including weathering, diagenesis, and hydrothermal processes (Bakker et al., 2019; Hubert et al., 2012; Viennet et al., 2015; Sakharov and Lanson, 2013 and references therein).

The abundance and composition of clay minerals, whether as discrete phases or mixed-layer minerals (MLMs), were refined using Markovian statistics to model the stacking sequences of layers, characterized by the Reichweite parameter (R; Sakharov and Lanson, 2013 and references therein). The Reichweite parameter quantifies the extent of ordering, reflecting how the arrangement of adjacent layers influences the nature of a given layer. When R = 0 (i.e., R0), this indicates a random stacking of layers, whereas R = 1 (i.e., R1) implies that the occurrence probability of a given layer type depends on the nature of its nearest neighbor. In a two-component R1 structure, the junction probability Pii between two successive identical layers i (which are present in minor proportion in the multilayer model, i.e., Wi < 0.5) determines the stacking sequence. For Pii = 0, the stacking exhibits the maximum possible degree of ordering (R1-MPDO). MLMs were defined by the number, type, relative proportion, and stacking sequence of their layers, all considered adjustable parameters. Coherent scattering domain sizes (CSDSs) were modeled using a lognormal distribution and treated as adjustable parameters (Sakharov and Lanson, 2013 and references therein). Instrumental parameters (e.g., beam divergences, goniometer radius) and structural models were integrated into the refinement process.

The layer thicknesses for serpentine and mica layers were fixed at 7.25 and 9.98 Å, respectively, for all measurement conditions. For Ca-saturated expandable layers in air-dried (AD) preparations, interlayer thicknesses depend on hydration states: ~9.8 Å for dehydrated (S0w), ~12.5 Å for mono-hydrated (S1w), and ~15.0 Å for bi-hydrated (S2w) layers (Ferrage, 2005). After ethylene glycol (EG) solvation, smectite layers expand, incorporating one (S1g, 12.9 Å) or two (S2g, 16.8 Å) EG sheets into their interlayers (Ferrage, 2005).

The multi-specimen method (Sakharov and Lanson, 2013, and references therein) was used to ensure consistency across diffraction patterns obtained from different treatments (Ca-AD, Ca-EG). This approach leverages contrasts in layer thickness and structure factors, influenced by interlayer cation type, amount, and position, to refine structural

models. The fitting strategy aimed at developing a single structural model per sample that could describe all diffraction patterns (Ca-AD, Ca-EG). Model validity was assessed based on (i) consistent phase proportions across treatments and (ii) structural model coherence across particle size fractions. Fit quality was quantified using the unweighted R-factor (Rp; Sakharov and Lanson, 2013, and references therein).

### 2.4. Middle infrared analysis

Fourier-transform infrared (FTIR) spectra have been recorded in transmission mode on the 400-4000 cm⁻¹ range with a 4 cm⁻¹ resolution using a Nicolet 6700 FTIR spectrometer fitted with a KBr beamsplitter and a DTGS-KBr detector. The spectra were obtained by averaging 200 scans from KBr pellets, prepared by mixing 1 mg of sample with 150 mg of KBr, followed by drying at 120°C for two days. The measurements were performed on the spectroscopic platform at IMPMC, Paris, France.

### 2.5. Mössbauer spectroscopy measurements and analysis

The Mössbauer spectroscopy measurements were performed at the ESRF nuclear resonance beamline ID14 (previously ID18) (Kupenko et al., 2024). This technique is based on the synchrotron Mössbauer source (SMS) enabling conventional (energy-domain) spectroscopy. Experiment was carried out in 7/8 + 1 bunch mode of the accelerator. The stability of nuclear monochromator (conditions of ⁵⁷FeBO₃ crystal) was periodically checked during the experiment by measuring the standard absorber (Yaroslavtsev and Chumakov, 2022). The beam size was focused down to ~10×10 µm using Kirkpatrick-Baez mirrors. The sample thickness was ~300 µm. The time dependence of the Doppler velocity has a sinusoidal shape to avoid the slippage of ⁵⁷FeBO₃ crystal. The unfolded spectra were registered in 1024 channels. α-Fe was used as a reference sample for calibration. Measurements were carried out at room temperature (~298 K) and at various temperatures below with the lowest one at 4 K. Low temperature measurements were performed using He-gas exchange-cryostat available at the beamline (Van Der Linden et al., 2016).

The Mössbauer spectra were processed with the SYNCmoss software (Yaroslavtsev, 2023), specially designed for the fitting of SMS spectra. SYNCmoss allows to take into account the actual instrumental function, which is extracted from an experimental spectrum of a standard sample, in the fitting procedure using full transmission integral.

A “bulk” sample of Orgueil CI meteorite was measured for determining the redox ratio of phyllosilicates in Orgueil CI meteorite.

## 3. Results

### 3.1. Particle size distribution

Sequential particle size distribution of the Orgueil meteorite shows that the majority of its mass is concentrated in the 2-50 µm size fraction, which represents 31 Wt. %. The 0.2-2 µm and 0.05-0.2 µm fractions account for 29 Wt. % and 23 Wt. %, respectively. The finer fractions, 0.02-0.05 µm and < 0.02 µm, contribute 10 Wt. % and 5 wt. %, respectively. The coarsest fraction, > 50 µm, represents only 2 Wt. % of the total mass.

### 3.2. Qualitative description of the XRD patterns

The XRD patterns of the sub-fractions obtained from oriented preparations are presented in Fig. 1. The 2-50 µm fraction is dominated by magnetite, as indicated by its characteristic XRD peaks at 4.83, 2.96, 2.53, 2.42, 2.10, and 1.71 Å (PDF Card # 04-005-4319), marked with grey stars. Additionally, small XRD peaks at 2.64 and 2.06 Å indicate the presence of minor amounts of pyrrhotite 4C (COD Card # 2104739), marked with grey circles. Clay minerals are present in small quantities,
<figure><img src="image_4.png" /><figcaption>**Fig. 1.** Experimental XRD patterns the sub-fractions obtained under the Ca-Ad state (black line) and the Ca-Eg state (red line). The peak position are indicated in angströms. The relative mass of each sub-fraction shown was normalized to 100 %.</figcaption></figure>

as evidenced by the weak intensity of their 0211 reflection at 4.57 Å and the displacement of XRD peaks between the Ca-AD and Ca-EG treatments, showing the presence of smectite layers. Magnetite is also present in smaller amounts in the other size fractions, as suggested by the weak intensity of its XRD peaks, while pyrrhotite reflections are absent in the other fractions.

In the 0.2-2 µm fraction, the (0211) reflection of clay minerals is visible, suggesting poor particle orientation due to their large size. This reflection is absent in the < 0.2 µm fractions, indicating better particle orientation, which enhances the intensities of their 00l reflections compared to the fractions coarser than 0.2 µm. Regardless of particle size, the presence of swelling layers, corresponding to smectite, is confirmed by the shift of XRD patterns between the Ca-AD and Ca-EG states. In the Ca-AD state, the rational series of peaks at 15.00, 4.99, and 3.01 Å could correspond to the 001, 003, and 005 reflections of smectite, respectively. After Ca-EG treatment, these peaks shift to 17.16, 8.20, 5.61, 3.41, and 2.81 Å, forming an irrational series, which suggests the smectite is not discrete but part of a mixed-layer mineral (MLM). A discrete Ca-smectite swelling at 17.16 Å would instead produce 00l reflections at 8.58, 5.72, 4.29, 3.43, and 2.86 Å. The observed shifts and positions indicate a smectite-rich MLM likely containing serpentine layers. Specifically, the 002 reflection at 8.20 Å shifts toward the 001 reflection of serpentine, and the 004 reflection at 3.41 Å shifts toward the 002 reflection of serpentine. Note that in the < 0.02 µm fraction, the XRD peak at ~26.16 Å corresponds to a superstructure only visible in the Ca-AD state. Such XRD peak indicates the presence of an R1-MPDO MLM, most likely with a 10 Å layer (i.e. mica) and smectite (S2w) layers. This superstructure is not visible in the Ca-EG state, which could be due to higher heterogeneity in the solvation state of smectite layers under EG, leading to limit the stacking order with 3 types of layers (along the c-axis).

Discrete serpentine is identified in the 0.05-0.2 and 0.2-2 µm fractions by nearly rational 00l reflections, which remain unchanged between Ca-AD and Ca-EG treatments at 7.28 and 3.62 Å. In contrast, these

reflections are absent in the < 0.02 µm fraction, indicating a lack of discrete serpentine in the finest fraction.

With decreasing particle size, the clay mineralogy and structure remain similar. The 00l reflections of discrete serpentine weaken, broaden, and shift slightly to lower angles as particle size decreases. For the smectite-rich MLM, 00l reflections also broaden with decreasing particle size, indicating a reduction in CSDS. In the < 0.02 µm fraction, the 001 and 002 reflections shift slightly to lower angles, reflecting the apparent irrationality associated with reduced CSDS. This suggests that the composition and proportion of layers constituting the MLM do not vary with particle size.

Qualitative analysis of the XRD patterns indicates that magnetite is primarily present in the > 2 µm size fraction, along with minor pyrrhotite and clay minerals. The < 2 µm size fractions are dominated by clay minerals with minor magnetite. While the clay mineral structure remains consistent across size fractions, the proportion of discrete serpentine decreases, becoming nearly absent in the < 0.02 µm fraction. Alongside discrete serpentine, a smectite-rich MLM with potential serpentine layers is present in all size fractions.

Table 1

### 3.3. XRD profiles modeling

The final modeled XRD patterns for the Ca-AD and Ca-EG states are presented in Fig. 2, with the detailed contributions of each clay mineral phase shown in Fig. 3. The fitting procedure provided parameters such as coherent scattering domain sizes (CSDS), the nature of the mixed-

<table><caption>**Table 1**
Proportion of the relative mass of each sub-fraction normalized to 100 % given a total sample recovery of 92 %.</caption><thead><tr><th>Size fraction</th><th>&lt; 0.02 µm</th><th>0.02-0.05 µm</th><th>0.05-0.2 µm</th><th>0.2-2 µm</th><th>2-50 µm</th><th>&gt; 50 µm</th></tr></thead><tbody><tr><th>Wt. %</th><td>5</td><td>10</td><td>23</td><td>29</td><td>31</td><td>2</td></tr></tbody></table>
<figure><img src="image_5.png" /><figcaption>Fig. 2. Comparison of calculated and experimental XRD patterns obtained after Ca-Ad and Ca-Eg states for the different sub-fractions. The experimental and calculated XRD patterns are shown as crosses and solid lines, respectively. Difference plots are shown at the bottom of the figure.</figcaption></figure>

layer minerals (MLMs), and the proportions of each phase, which are listed in Tables 2 and 3. Up to four clay mineral phases, including both discrete phases and R0 MLMs, were required to reproduce the full set of experimental XRD patterns. Note that an R1 MLM was used to fit the superstructure of the < 0.02 µm in the Ca-AD state. The multispecimen method was used to ensure consistency and reliability of the model. Details of the fitting strategy are presented in the supplementary file.

Table 4

Based on the same crystal structure model for Ca-AD and Ca-EG XRD patterns (Table 3), the proportions of the clay phases in between the two treatments are very closed with a maximum relative variation of 2 % per clay phase, even in the < 0.02 µm fraction for which the contribution of the R0 serpentine-smectite-mica MLM in Ca-EG (84 % Wt.) represents the sum of R0 serpentine-smectite-mica MLM (65 % Wt.) and R1-MPDO serpentine-smectite-mica (18 % Wt.) in Ca-AD. This serpentine-smectite-mica MLM with a 50-30-20 composition (i.e. mixed layer mineral randomly interstratified consisting of 50 % of serpentine layers, 30 % of smectite layers and 20 % of mice layers) is the dominant clay phase (84 %Wt.) of the < 0.02 µm fraction. It remains the dominant clay phase in the 0.02-0.05 and 0.05-0.2 µm fractions, accounting for 83 % and 65 %, respectively. The finest fraction also contained a discrete smectite phase accounting for 14 % and a R0 serpentine-mica MLM with a 90-10 composition that constitutes the remaining 2 %. For coarser size fractions > 0.02 µm, the proportion of discrete smectite decreases slightly, while the proportion of R0 serpentine-mica MLM increases, reaching 12 % in the 0.05-0.2 µm fraction. Discrete serpentine was also introduced to reproduce the experimental patterns in the 0.02-0.05 and 0.05-0.2 µm fractions, accounting for 2 % and 9 %, respectively.

### 3.4. Mössbauer measurements and modelling results

The presence of ferrihydrite necessitated performing Mössbauer measurements at 4 K to determine the Fe valence in clay minerals (see supplementary files for details regarding the fitting strategy). Due to beam time constraints, a full study of Mössbauer spectra acquisition across various temperatures for all size fractions was not feasible. Instead, we analyzed the evolution of Mössbauer spectra behavior at 4 K as a function of particle size for the 0.05-0.2 µm fraction, which contains less magnetite than the bulk fraction but significant contributions from

discrete serpentine, discrete smectite, and the R0 serpentine-smectite-mica MLM. Additionally, we analyzed the < 0.02 µm fraction, which lacks discrete serpentine, and one bulk sample of Orgueil.

Figure 4 presents the Mössbauer data and their fitting obtained at 4 K. The ferrihydrite contribution was modeled using a many-state relaxation model at 4 K to fit superparamagnetic behavior of ferrihydrite nanoparticles (see supplementary file), enabling the determination of the redox ratio of clay minerals (see Table 5). For the < 0.02 µm size fraction, the redox ratio of discrete smectite and the R0 serpentine-smectite-mica MLM was calculated using two doublets, one representing the Fe²⁺ contribution and the other representing the Fe³⁺ contribution, resulting in a redox ratio of 80 %, calculated as N<sub>Fe²⁺</sub>/(N<sub>Fe³⁺</sub> + N<sub>Fe²⁺</sub>), where N is the number of Fe ions. For the 0.05-0.2 µm size fraction, the spectra were fitted with three doublets for clay minerals: two corresponding to the same contributions as in the < 0.02 µm fraction and an additional doublet representing Fe²⁺ in the discrete serpentine. Notably, no Fe³⁺ contribution from the discrete serpentine was detected within the sensitivity limits of the Mössbauer measurements.

Overall, the redox ratio of clay minerals was 80 % for the < 0.02 µm size fraction (discrete smectite and R0 serpentine-smectite-mica MLM) and 60 % for the 0.05-0.2 µm fraction, with the discrete serpentine being entirely Fe²⁺. For the bulk fraction, the Mössbauer spectrum was fitted using the same parameters as the 0.05-0.2 µm fraction while allowing the Fe²⁺ contribution from discrete serpentine to vary. The redox ratio of phyllosilicates in Orgueil obtained in the bulk sample was determined to be 53 %.

### 3.5. MIR results

Fig. 5 presents the mid-infrared (MIR) transmission spectra of the sub-fractions. The band at 3690 cm⁻¹ (corresponding to 3687 cm⁻¹) is attributed to the in-phase stretching of the interlayer OH groups in Mgserpentine, while the band at 612 cm⁻¹ corresponds to the longitudinal (LO) and transverse (TO) modes of in-plane OH bending in lizardite, with some octahedral Mg²⁺ substituted by Fe²⁺ (Baron and Petit, 2016). The bands at 1098 and 930 cm⁻¹ are assigned to Si-O stretching vibrations in serpentine.

Characteristic bands of saponite (i.e. Mg-rich smectite) are also observed, including an OH stretching band near 3679 cm⁻¹ and an OH
<figure><img src="image_6.png" /><figcaption>**Fig. 3.** Contributions of the different clay phases to the calculated and the experimental XRD patterns obtained after Ca-Ad and Ca-Eg states for the different sub-fractions.</figcaption></figure>

bending band near 667 cm⁻¹. Additionally, in the finest size fraction, a broad band centered at 3623 cm⁻¹ appears, which could be attributed to Fe³⁺-OH stretching vibrations in smectite with low Fe³⁺ content in octahedral positions (Fox et al., 2021). The band at 1012 cm⁻¹ corresponds to the Si-O stretching vibration of saponite. The band at 798 cm⁻¹ is linked to Si-O-Al deformation modes, indicating a low amount of six-coordinated Al in the octahedral sheets. The band at 589 cm⁻¹ likely represents [6]Fe³⁺-O<sub>apical</sub>-Si⁴⁺ coupled lattice bending vibrations with a possible contribution from [4]Fe(III) (Baron et al., 2016).

As the particle size decreases, there is a notable reduction in the intensities of bands related to serpentine layers and an increase in the bands associated with saponite layers, consistent with the XRD results.

Furthermore, the intensities of bands corresponding to Fe³⁺ increase with particle size, which is consistent with the Mössbauer findings.

## 4. Discussion

### 4.1. Advanced XRD profile modeling for precise phyllosilicate characterization

A novel approach was applied in the present study to characterize the clay mineralogy of Orgueil, according to recent advancements in soil science (Bakker et al., 2019; Hubert et al., 2012; Viennet et al., 2015). Bass, (1971) was the first to unambiguously identify the crystalline
<caption>**Table 2**</caption>

Composition and structural parameters for the mineral phases obtained from the XRD profile modeling as a function of the particle size and treatments. Note that the CSDS parameter is given in layer per crystal. Serp/Sm/Mica represents the proportion of serpentine, smectite and mica layers within a crystal structure. Paa: probability that a serpentine layer follows a serpentine layer, Pba: probability that a smectite layer follows a serpentine layer and Pca: probability that a mica layer follows a serpentine layer.

<table><caption>**Table 3**</caption><thead><tr><th rowspan="2">Mineral Phase</th><th rowspan="2">Fractions</th><th colspan="2">&lt; 0.02 µm</th><th colspan="2">0.02 - 0.05 µm</th><th colspan="2">0.05 - 0.2 µm</th></tr><tr><th>State</th><th>Ca-AD</th><th>Ca-EG</th><th>Ca-AD</th><th>Ca-EG</th><th>Ca-AD</th><th>Ca-EG</th></tr></thead><tbody><tr><td rowspan="3">Smectite<br>R0 Serp/Sm/<br>Mica</td><td>CSDS</td><td>2</td><td>2</td><td>2</td><td>3</td><td>3</td><td>3</td></tr><tr><td>CSDS</td><td>5</td><td>5</td><td>6</td><td>7</td><td>9</td><td>10</td></tr><tr><td>Serp/Sm/<br>Mica</td><td>50/<br>30/<br>20</td><td>50/<br>30/<br>20</td><td>50/<br>30/<br>20</td><td>50/<br>30/<br>20</td><td>50/<br>30/<br>20</td><td>50/<br>30/<br>20</td></tr><tr><td rowspan="4">R1 - MPDO<br>Serp/Sm/<br>Mica</td><td>CSDS</td><td>10</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Serp/Sm/<br>Mica</td><td>50/<br>30/<br>20</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td></tr><tr><td>Paa/Pba/<br>Pca</td><td>0.0/<br>1.0/<br>1.0</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td></tr><tr><td>CSDS</td><td>5</td><td>5</td><td>6</td><td>6</td><td>7</td><td>7</td></tr><tr><td rowspan="2">R0 Serp/Mica</td><td>Serp/Sm/<br>Mica</td><td>90/<br>10</td><td>90/<br>10</td><td>90/<br>10</td><td>90/<br>10</td><td>90/<br>10</td><td>90/<br>10</td></tr><tr><td>CSDS</td><td>3</td><td>2</td><td>6</td><td>5</td><td>13</td><td>12</td></tr><tr><td>Serpentine</td><td>CSDS</td><td>0</td><td>0</td><td>2</td><td>2</td><td>11</td><td>9</td></tr></tbody></table>

Proportion in Wt. % of the mineral phases as a function of the particle size and treatments. R0 corresponds to a random stacking of layers while the R1 corresponds maximum possible degree of ordering.

<table><thead><tr><th rowspan="2">Fractions</th><th colspan="2">&lt; 0.02 µm</th><th colspan="2">0.02 - 0.05 µm</th><th colspan="2">0.05 - 0.2 µm</th></tr><tr><th>State</th><th>Ca-AD</th><th>Ca-EG</th><th>Ca-AD</th><th>Ca-EG</th><th>Ca-AD</th><th>Ca-EG</th></tr></thead><tbody><tr><td>Smectite</td><td>14</td><td>14</td><td>10</td><td>10</td><td>11</td><td>13</td></tr><tr><td>R0 Serp/Sm/Mica</td><td>65</td><td>84</td><td>82</td><td>83</td><td>65</td><td>66</td></tr><tr><td>R1-MPDO Serp/Sm/<br>Mica</td><td>18</td><td>-</td><td>-</td><td>-</td><td>-</td><td>-</td></tr><tr><td>R0 Serp/Mica</td><td>3</td><td>2</td><td>6</td><td>5</td><td>13</td><td>12</td></tr><tr><td>Serpentine</td><td>0</td><td>0</td><td>2</td><td>2</td><td>11</td><td>9</td></tr></tbody></table>

material in Orgueil as consisting primarily of smectite and serpentine, and highlighted the extremely thin nature of smectite and serpentine layers, with a secondary constituent likely being randomly interlayered in smectite in minor amounts. These findings are consistent with the results obtained here for the < 0.05 µm size fractions. Plus, although not discussed, a XRD peak at low angle corresponding to a superstructure was also present in the AD state (Bass 1971). Such superstructure might represent the coherent stacking of 7 Å and 10 Å layers already observed in high-resolution transmission electron microscopy (HRTEM) study (Fig.9.b in Tomeoka and Buseck, 1988).

The XRD profile modelling results provide precise determination of the structure of phyllosilicates and a detailed quantification of their mineralogical composition. These results refine and complement previous HRTEM studies (Golden et al., 1994; Ming et al., 1992; Tomeoka and Buseck, 1988) by offering a broader and more quantitative perspective on the mineralogy of the samples. The HRTEM studies have described both coarse and fine clay minerals, comprising intergrowths of layers with d-spacings of ~7 Å, ~10 Å, and wavy layers with d-spacings up to 15 Å, interpreted as serpentine (~7 Å) and smectite (~10-15 Å) layers (Tomeoka and Buseck, 1988). The present XRD results, reporting 7.25 Å, 9.98 Å, and 15 Å layers, are fully consistent with these HRTEM studies while providing a more precise characterization of the mineralogical composition.

Unlike TEM, XRD allows for the application of treatments such as Ca-AD and Ca-EG to distinguish the nature of the layers based on their XRD behavior enabling the unambiguous distinction of smectite and mica layers. Such distinction by TEM is highly challenging because under vacuum collapsed smectite layers without interlayer H₂O should exhibit d-spacings of ~10 Å, making it difficult to differentiate between a 10 Å mica layer and a collapsed smectite layer (Amouric and Olives, 1991; Golden et al., 1994; Huff et al., 1988; Ming et al., 1992). However, dehydration of smectite may be incomplete under TEM vacuum conditions (Kagawa et al., 2020) and smectites may store organic molecules within their interlayer spaces (Bahae-eddine et al., 2024; Viennet et al., 2023). In addition, at the HRTEM scale, mica layers appear as straight layers with 10 Å d-spacings, whereas smectite layers are imperfect and curved, with d-spacings ranging from 10 to 13 Å (Ahn and Peacor, 1986). Accordingly, the 10 to 15 Å layers observed in the calcium-aluminum-rich inclusions of the Allende CV meteorite were interpreted as intergrowths of mica and smectite layers (Tomeoka and Buseck, 1982). The well-defined 10 Å layers observed by Tomeoka and Buseck, (1988) in CI Orgueil meteorite could, therefore, correspond to mica layers, while wavy layers >10 Å represent smectite layers. In addition, the TEM analysis of phyllosilicates of CI meteorites after organic molecule intercalation in the interlayer space of smectite has led to >12 Å layers being interpreted as smectite and the few remaining 10 Å layers as mica (Golden et al., 1994; Ming et al., 1992). Thus, the combination of straight layers at 10 Å and wavy > 10 Å layers could indicate the presence of both smectite and mica-type layers, consistent with the results from XRD profile modeling.

## 4.2. Oxidation state of Orgueil CI meteorite

The general consensus is that clay minerals in Orgueil are more oxidized than those in Ryugu (Leroux et al., 2024; Nakamura et al., 2022; Roskosz et al., 2024). It remains unclear whether the clay minerals in CI chondrites are more oxidized than those from Ryugu due to terrestrial oxidation, evidenced by the presence of ferrihydrite, maghemite components in magnetites and sulfates (see supplementary files; Gounelle and Zolensky, (2001); King et al., (2020); Roskosz et al., (2024); Leroux et al., (2024)) or because they originally formed under more oxidizing conditions (Leroux et al., 2024; Roskosz et al., 2024). One of the main issues lies in the difficulty of determining the redox ratios of clay minerals in CI chondrites. Determining the redox ratio of clay minerals of Orgueil using XANES is particularly difficult due to their fine-scale associations with Fe-sulfur, magnetite and ferrihydrites (Leroux et al., 2024; Tomeoka and Buseck, 1988), the presence of the latter preventing a proper interpretation of XANES data (Garenne et al., 2019; Leroux et al., 2024). But this is not the case for cryogenic Mössbauer data collected at 4 K since this technique allows to isolate the signal from the contribution of ferrihydrites.

Thanks to cryogenic Mössbauer data collected at 4 K on the size fractions of Orgueil, the redox ratio of clay minerals in bulk Orgueil was determined to be 53 % (NFe³⁺/(NFe³⁺ + NFe²⁺)). This value is similar to the bulk values for pristine (terrestrially unaltered) grains returned from the Ryugu asteroid, i.e. 50-65 % (Roskosz et al., 2024). Therefore, clay minerals in Orgueil are not overoxidized as compared to Ryugu, contrary to the common assumption. Additionally, the Fe³⁺ in the tetrahedral sheet for the discrete smectite and/or R0 serpentine-smectite-mica MLM cannot be attributed to terrestrial oxidation because clay minerals cannot accommodate Fe²⁺ in tetrahedral sites due to its large ionic size.

Plus, some grains of Ryugu exhibit redox differences (Nakamura et al., 2022; Roskosz et al., 2024), with serpentine-rich grains showing less than 35 % of Fe³⁺ and smectite-rich grains showing more than 65 % of Fe³⁺ (Roskosz et al., 2024). This is consistent with the present study reporting the pure Fe²⁺ nature of serpentine and the more oxidized nature of discrete smectite and R0 serpentine-smectite-mica MLM containing up to 80 % Fe³⁺. We suggest that these differences in redox ratios
<caption>**Table 4**</caption>

Models Mössbauer parameters at 4 K obtained on Orgueil bulk, 0.05-0.2 µm and < 0.02 µm fractions. The parameter corresponds to the ratio between lines due to texture/polarization (in isotropic case =0.5). The R parameter is relaxation rate (frequency). The Alpha parameter is the magnetic anisotropy energy divided by temperature energy. Note that the term “Fix” indicate that some parameter was fixed during Mössbauer fitting as explained in the text.

<table><thead><tr><th>Size<br>Fraction</th><th>Parameters</th><th>Fe<sup>2+</sup> Discrete serpentine + R0<br>MLM serpentine-mica</th><th>Fe<sup>2+</sup> Clays<br>Discrete smectite + R0 MLM<br>serpentine-smectite-mica</th><th>Fe<sup>3+</sup> Clays<br>Discrete smectite + R0 MLM<br>serpentine-smectite-mica</th><th>Ferrihydrite<br>(Bulk: +Site B<br>magnetite)</th><th>Magnetite<br>(Site A)</th></tr></thead><tbody><tr><td rowspan="10">Bulk</td><td>δ IS<br>(mm/s)</td><td>1.249<br>±0.022</td><td>1.309<br>Fix</td><td>0.363<br>Fix</td><td>0.464<br>±0.005</td><td>0.510<br>±0.019</td></tr><tr><td>Δ QS<br>(mm/s)</td><td>2.816<br>±0.042</td><td>2.792<br>Fix</td><td>0.992<br>Fix</td><td>-0.056<br>±0.005</td><td>-0.378<br>±0.04</td></tr><tr><td>Width<br>(mm/s)</td><td>0.175<br>±0.103</td><td>0.230<br>Fix</td><td>0.787<br>Fix</td><td>-</td><td>-</td></tr><tr><td>H (T)</td><td>-</td><td>-</td><td>-</td><td>58.740<br>±0.201</td><td>48.003<br>±0.314</td></tr><tr><td>A</td><td>0.5<br>Fix</td><td>0.5<br>Fix</td><td>0.5<br>Fix</td><td>A=0.495<br>±0.008</td><td>A=0.562<br>±0.032</td></tr><tr><td>R</td><td>-</td><td>-</td><td>-</td><td>6.823<br>±0.893</td><td>0.0<br>±0.001</td></tr><tr><td>Alpha</td><td>-</td><td>-</td><td>-</td><td>2.113<br>±0.059</td><td>17.125<br>±5.332</td></tr><tr><td>Area<br>(%)</td><td>3.5<br>±0.5</td><td>1.4<br>±0.1</td><td>5.5<br>±0.5</td><td>78.7<br>±1.3</td><td>10.9<br>±1.3</td></tr><tr><td rowspan="7">0.05-0.2<br>μm</td><td>δ IS<br>(mm/s)</td><td>1.273<br>±0.022</td><td>1.309<br>Fix</td><td>0.363<br>Fix</td><td>0.469<br>±0.008</td><td>-</td></tr><tr><td>Δ QS<br>(mm/s)</td><td>2.958<br>±0.044</td><td>2.792<br>Fix</td><td>0.992<br>Fix</td><td>-0.073<br>±0.016</td><td>-</td></tr><tr><td>Width<br>(mm/s)</td><td>0.193<br>±0.098</td><td>0.230<br>Fix</td><td>0.787<br>Fix</td><td>-</td><td>-</td></tr><tr><td>H (T)</td><td>-</td><td>-</td><td>-</td><td>57.466<br>±0.141</td><td>-</td></tr><tr><td>A</td><td>0.5<br>Fix</td><td>0.5<br>Fix</td><td>0.5<br>Fix</td><td>0.481<br>±0.008</td><td>-</td></tr><tr><td>R<br>(frequency)<br>Alpha (KV/<br>kT)</td><td>-</td><td>-</td><td>-</td><td>R=4.023<br>±0.429<br>Alpha=1.814<br>±0.033</td><td>-</td></tr><tr><td>Area<br>(%)</td><td>4.8<br>±0.6</td><td>2.8<br>±0.2</td><td>11.2<br>±0.6</td><td>80.5<br>±0.9</td><td>-</td></tr><tr><td rowspan="7">&lt; 0.02 μm</td><td>δ IS<br>(mm/s)</td><td>-</td><td>1.309<br>±0.049</td><td>0.363<br>±0.046</td><td>0.496<br>±0.022</td><td>-</td></tr><tr><td>Δ QS<br>(mm/s)</td><td>-</td><td>2.792<br>±0.098</td><td>0.992<br>±0.078</td><td>-0.085<br>±0.042</td><td>-</td></tr><tr><td>Width<br>(mm/s)</td><td>-</td><td>0.230<br>±0.168</td><td>0.797<br>±0.112</td><td>-</td><td>-</td></tr><tr><td>H (T)</td><td>-</td><td>-</td><td>-</td><td>59.159<br>±0.360</td><td>-</td></tr><tr><td>A</td><td>-</td><td>0.5<br>Fix</td><td>0.5<br>Fix</td><td>0.56<br>±0.017</td><td>-</td></tr><tr><td>R<br>(frequency)<br>Alpha (KV/<br>kT)</td><td>-</td><td>-</td><td>-</td><td>3.709<br>±0.897<br>1.461<br>±0.069</td><td>-</td></tr><tr><td>Area<br>(%)</td><td>-</td><td>4.3<br>±0.9</td><td>16.7<br>±1.2</td><td>78.9±1.4</td><td>-</td></tr></tbody></table>

can be attributed to varying proportions of clay mineral types rather than intrinsic differences in the redox states of clay minerals themselves, as also noted by Roskosz et al., (2024). This might also explain the Fe redox variability at micrometer scale (Leroux et al., 2024).

Finally, CI chondrites are generally thought to have formed under more oxidizing conditions than other chondrite classes, as indicated by their position in the Urey-Craig diagram and also their high iron oxidation index (Brearley et al., 1998; Garenne et al., 2019), our results indicate that the redox conditions governing the formation of clay minerals in CI chondrites were actually not significantly more oxidizing. Moreover, CI materials may, in fact, have been even more reduce than those of other carbonaceous chondrites as detailed and demonstrated by Amano et al. (2025).

### 4.3. Formation of clay minerals in Orgueil CI meteorite

The absence of structural evolution in clay minerals as a function of particle size contradicts the hypothesis that fine-grained phyllosilicates

formed through the alteration of coarse-grained phyllosilicates (Tomeoka and Buseck, 1988), as previously proposed (Aléon-Toppani et al., 2024; Leroux et al., 2024). Indeed, such formation processes, common in soil science, are typically characterized by structural evolution in clay minerals as particle size decreases (Hubert et al., 2012; Viennet et al., 2015). These results align with the observed iron valence in clay mineral phases sorted by particle size and type, revealing distinct redox ratios: coarse-grained phyllosilicates correspond to Fe<sup>2+</sup>-rich discrete serpentine, while finer grains correspond to Fe<sup>3+</sup>-rich R0 MLMs and discrete smectite. These redox differences appear to be incompatible with an aqueous alteration at equilibrium, suggesting instead distinct formation (micro)environments (Aléon-Toppani et al., 2024; Leroux et al., 2024; Nakamura et al., 2022; Reid et al., 1970).

The results presented here challenge also the traditional views on the origins of clay minerals in CI materials. The unique structure of clay minerals, composed of R0 serpentine-smectite-mica MLMs, is not observed in low-temperature aqueous alteration environments, neither on Earth nor in extraterrestrial contexts. Furthermore, the presence of
<figure><img src="image_7.png" /><figcaption>Fig. 4. Experimental and modeled Mössbauer spectra and phase contributions at 4 K.</figcaption></figure>

<table><caption>**Table 5**
Redox ratio of clay minerals determined by Mössbauer spectroscopy at 4 K.</caption><thead><tr><th>Size fraction</th><th>%Fe<sup>3+</sup>/(Fe<sup>2+</sup>+Fe<sup>3+</sup>)<br>Discrete serpentine + R0 MLM serpentine-mica</th><th>%Fe<sup>3+</sup>/(Fe<sup>2+</sup>+Fe<sup>3+</sup>)<br>Discrete smectite + R0 MLM serpentine-smectite-mica</th><th>Overall<br>%Fe<sup>3+</sup>/(Fe<sup>2+</sup>+Fe<sup>3+</sup>) in clay minerals</th></tr></thead><tbody><tr><td>Bulk</td><td>0 %</td><td>80 %</td><td>53 %</td></tr><tr><td>0.05-0.2 µm</td><td>0 %</td><td>80 %</td><td>60 %</td></tr><tr><td>&lt; 0.02 µm</td><td>-</td><td>80 %</td><td>80 %</td></tr></tbody></table>

mica-type layers cannot be explained by low-temperature conditions. On Earth, such layers typically form under high-temperature conditions in magmatic and metamorphic rocks (Bailey, 1984; Mottana et al., 2002). Since magmatic or metamorphic processes are not expected in CI objects, their formation raises questions. Mica-type layers (e.g., Na-phlogopite or margarite) have been previously observed in the CV Allende meteorite, specifically in calcium-aluminum-rich inclusions

(CAIs) or amoeboid olivine aggregates (Hashimoto and Grossman, 1987; Keller and Buseck, 1991). Petrographic and thermodynamic evidence suggests these layers formed in the solar nebula via gas-solid interactions (Hashimoto and Grossman, 1987; Keller and Buseck, 1991) while, Tomeoka and Buseck, (1982) reported intergrowths of mica and montmorillonite in Allende CAIs and attributed their formation to low-temperature metamorphism involving aqueous alteration. Note
<figure><img src="image_8.png" /><figcaption>Fig. 5. MIR spectra of the sub-fractions obtained in transmission mode in the range 4000-500 cm-1 (a) and 1300-400 cm-1 (b).</figcaption></figure>

that, even low-temperature metamorphism is not expected for CI chondritic objects. The presence of mica layers within MLM along with the unique R0 serpentine-smectite-mica MLM structure, absent in low temperature aqueous alteration process, likely points to a pre-accretionary origin for these phyllosilicates, possibly through gas-grain interactions or direct condensation within the solar nebula.

A nebular origin for phyllosilicates has been proposed in various chondritic objects. Characteristics such as sharp contacts between phyllosilicates and anhydrous silicates, the absence of alteration fronts, the presence of mica layers, deuterium to hydrogen ratios in the structural OH groups, and specific petrological features have led some authors to suggest that phyllosilicates could form via nebular processes (Bischoff, 1998; Brearley, 2014; Ciesla et al., 2003; Fanale and Cannon, 1974; Hashimoto and Grossman, 1987; Keller and Buseck, 1991; Lau-retta et al., 2000).

The origin of phyllosilicates, whether in the solar nebula or within parent bodies, has historically been debated from a macroscopic thermodynamic point of view. Anders, (1971) proposed that CI chondrites closely approximate the condensable fraction of primordial solar-system matter. Lewis, (1972), through thermodynamic calculations, showed that phyllosilicates (e.g., talc) condensate in the solar nebula. Later, Barshay and Lewis, (1976), along with references therein, refined equilibrium condensation diagrams and demonstrated that phyllosilicates (e.g., serpentine) could condense at low temperatures under equilibrium conditions. Note that, the classical condensation sequences (Grossman, 1972), which assume equilibrium, did not account for their formation although Grossman and Larimer, (1974 and references therein) showed that phyllosilicates might form via reactions between magnesium silicates and gaseous H₂O at low temperatures. Such models, however, neglect the kinetic inhibition of phyllosilicate formation. Indeed, part of these minerals may form through gas-grain interactions between H₂O and silicates (Prinn and Fegley, 1989) and are therefore controlled by kinetic factors such as solid-state diffusion rates, particle size, and crystallinity (see also Fegley, (2000, and references therein)). Using the Simple Collision Theory (SCT) model, Prinn and Fegley, (1989) and later Fegley, (2000, and references therein) showed that phyllosilicate formation (i.e. serpentine) via water vapor interaction with silicates (i.e. olivine) would be kinetically inefficient within the lifetime of the nebula. Nonetheless, more recent thermo-kinetic studies suggest that phyllosilicates could indeed form by gas-grain interactions within a million years via water vapor alteration of silicates, a timescale faster than the nebula lifetime (D'Angelo et al., 2019; Thi et al., 2020). Moreover, local concentration of dust and gas or shock waves compressing and heating locally the nebular gas and dusts in icy regions

would drastically speed up the formation of nebular phyllosilicates (Ciesla et al., 2003). Also, Yamamoto and Tachibana, (2016) successfully demonstrated the experimental formation of serpentine from water vapor (pH₂O = 1 bar) and amorphous olivine-like material at 100°C within a few hours. Using the experimentally derived thermodynamic data, they applied the SCT model to evaluate kinetics of serpentine formation from amorphous olivine under more expected nebula conditions (i.e. total pressure of solar nebula of 10⁻³ bar) and concluded that this process could occur within the solar nebulae. Finally, Woitke et al., (2018) also explored the formation of minerals at temperatures as low as 100 K in a solar composition gas, with and without equilibrium condensation by thermodynamic simulation. Their findings indicate that phyllosilicates are the most abundant condensates at low temperatures: micas (e.g., phlogopite) dominate at ~500 K, while serpentine (e.g., lizardite) becomes more abundant between 200-400 K. Although experimental formation of clay minerals through condensation might have been successful (John and Tibbetts, 2020), the experimental conditions applied do not appear to be directly relevant to nebular environments.

## 5. Conclusion

The refinement of Orgueil clay mineralogy through XRD profile modeling and Mössbauer spectroscopy at room and helium temperatures, combined with particle size separation, provides a comprehensive characterization of the crystal structure, proportions, and Fe redox state of phyllosilicates. XRD profile modeling show the presence of four clay phases: discrete smectite and serpentine, along with two randomly stacked (R0) mixed-layer minerals (MLMs) serpentine-mica (90 %-10 %) and serpentine-smectite-mica (50 %-30 %-20 %). Regardless of particle size, the R0 serpentine-smectite-mica MLM is the dominant phase, and the crystal structure of clay minerals remains unchanged. Cryogenic Mössbauer data collected at 4 K reveal that Orgueil and Ryugu share a similar redox state, with Fe²⁺-rich discrete serpentine and R0 serpentine-mica MLM, while discrete smectite and R0 serpentine-smectite-mica MLM contain 80 % Fe³⁺. The observation that Orgueil and Ryugu silicates share similar redox properties demonstrate that clay minerals of Orgueil have not undergone significant terrestrial oxidation, in contrast to sulfides.

The CI parent bodies have undergone a complex history, involving accretion, disruption, and re-accumulation, as evidenced by their rubble-pile structures and the presence of numerous clasts in both returned samples and CI meteorites (e.g. Nakamura et al., 2022). Aqueous activity on these parent bodies may explain for features such as
the evaporite sequences recently observed in Bennu grains (McCoy et al., 2025) and the fluid inclusions reported in a Ryugu pyrrhotite crystal (Nakamura et al., 2022). These secondary processes may have also contributed to the partial oxidation of Fe-bearing phases, shifting their redox state toward more Fe³⁺-rich compositions.

However, a growing body of evidence, including thermodynamic modeling, suggests that the origin of phyllosilicates in CI materials may not be exclusively due to parent-body aqueous alteration, but could instead be pre-accretionary. For example, the homogeneous crystal structure across particle sizes, the differences in Fe redox states among phyllosilicate types, and the presence of mica layers in MLMs are inconsistent with a low temperature aqueous alteration at equilibrium under asteroidal conditions. Instead, as proposed by Reid et al., (1970), Orgueil might be considered as a complex low-temperature mechanical mixture, the chemistry of which reflects the contributions of materials from different environments. We propose that most of the phyllosilicates formed in the protoplanetary disk either by gas-grain alteration or condensation processes as shown in recent thermodynamic models. A pre-accretionary origin for these clay minerals could also explain the coexistence of unaltered high-temperature minerals and phyllosilicates within CI objects. For instance, altered amorphous phases coexist with unaltered, micrometer-sized olivine or pyroxene crystals and micro-chondrules (Reid et al., 1970; Nakamura et al., 2022; Aléon-Toppani et al., 2024; Leroux et al., 2024). Along the same line, the finding of pyrrhotite containing pentlandite exsolution textures directly points to high-temperature nebular processes (Visser et al., 2019). Finally, the observed redox variation among phyllosilicates may reflect formation in distinct environments and could serve as constraints on local formation conditions such as temperature, C/O ratios, and emplacement within the pre-accretionary disk and rings, which ultimately control the final Fe³⁺/Fe²⁺ ratio.

## Declaration of generative AI and AI-assisted technologies in the writing process

During the preparation of this work the authors used ChatGPT in order to assist with grammar correction and English language polishing. After using this tool/service, the author(s) reviewed and edited the content as needed and take(s) full responsibility for the content of the publication.

## CRediT authorship contribution statement

**Viennet:** Writing – review & editing, Writing – original draft, Visualization, Validation, Software, Resources, Project administration, Methodology, Investigation, Funding acquisition, Formal analysis, Data curation, Conceptualization.
**K. Amano:** Writing – review & editing, Formal analysis, Data curation.
**L. Delbes:** Writing – review & editing, Data curation.
**B. Lavina:** Writing – review & editing, Data curation.
**F. Hubert:** Writing – review & editing, Methodology, Formal analysis, Data curation.
**S. Bernard:** Writing – review & editing.
**S. Yaroslavtsev:** Writing – review & editing, Software, Methodology, Data curation.
**D. Bessas:** Writing – review & editing, Methodology, Data curation.
**A. Chumakov:** Writing – review & editing, Methodology, Data curation.
**J.-P. Celse:** Methodology, Data curation.
**M. Gounelle:** Writing – review & editing.
**L. Remusat:** Writing – review & editing, Funding acquisition.
**M. Roskosz:** Writing – review & editing, Writing – original draft, Validation, Project administration, Methodology, Investigation, Funding acquisition, Formal analysis, Data curation.

## Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

## Acknowledgements

We acknowledge financial support from the DIM-Map ACAV⁺ (project PARYUGU - PI: M.R) and by the European Research Council for funding via the ERC project HYDROMA (grant agreement No. 819587 - PI: L.R). J.-C. Viennet gratefully acknowledges financial support from the ATM 2022 project "Orgueil" (OE-7590). We thank Elisabeth Malassis for administrative simplification. The authors wish to acknowledge the workforce of the spectroscopic and X-ray diffraction facilities at IMPMC. The authors are grateful to the Chevron Energy Technology Company, a division of Chevron U.S.A. Inc. who freely allowed the use of the Sybilla software for academic purposes. A big thanks goes to Marek Szczerba for its help for computing modelling. The TEM facility at the CCM (Lille University) is supported by the Chevreul Institute, the European FEDER and Région Nord-Pas-de-Calais. Authors acknowledge the European Synchrotron Radiation Facility (ESRF) for provision of synchrotron radiation facilities under proposal number ES-1595 at nuclear resonance beamline ID14. The authors would like to thank the Muséum National d'Histoire Naturelle for sharing a part of the Meteorite of Orgueil (specimen number #234).

## Supplementary materials

Supplementary material associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2025.119678.

## Data availability

The data are available in the data.InDoRES repository: https://doi.org/10.48579/PRO/V8ZI1H.

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