Jérôme Aléon
email: jerome.aleon@mnhn.fr

Johanna Marin-Carbonne

Kevin D Mckeegan

Ahmed El Goresy

O, Mg, and Si isotope distributions in the complex ultrarefractory CAI Efremovka 101.1: Assimilation of ultrarefractory, FUN, and regular CAI precursors

Oxygen, magnesium, and silicon isotopic compositions in the mineralogically complex, ultrarefractory (UR) calcium-aluminum-rich inclusion (CAI) E101.1 from the reduced CV3 chondrite Efremovka confirm that E101.1 is a compound CAI composed of several lithological units that were once individual CAIs, free-floating in the solar protoplanetary disk. Each precursor unit was found to have had its own thermal history prior to being captured and incorporated into the partially molten host CAI.

) Spinel clusters are 16 O-rich (∆ 17 O ~ -22‰) and have Mg isotope systematics consistent with extensive equilibration with the host melt. This includes (i) δ 25 Mg values varying between +2.6 ‰ and +6.5 ‰ close to the typical value of host melilite at ~+5‰, and (ii) evidence for exchange of radiogenic 26 Mg with adjacent melilite as indicated by Al/Mg systematics. The spinel clusters may represent fine-grained spinel-rich proto-CAIs captured, partially melted, and recrystallized in the host melt. Al/Mg systematics indicate that both the sinuous pyroxene fragments and spinel clusters probably had canonical or near-canonical 26 Al contents before partial equilibration. (3) The main CAI host (∆ 17 O ≤ -2‰) had a complex thermal history partially obscured by subsequent capture and assimilation events. Its formation, referred to as the "cryptic" stage, could have resulted from the partial melting and crystallization of a 16 Orich precursor that underwent 16 O-depletion and a massive evaporation event characteristic of F and FUN CAIs (Fractionated with Unknown Nuclear effects). Alternatively, a 16 O-rich UR !" ! ′ ≤ 5.7 × 10 -5 ) suggest that the previously reported age difference of as much as 300,000 years between interior CAIs and their WLRs may be an artifact resulting from Mg isotopic perturbations, possibly by solid state diffusion or mixing between the interior and condensate components. The isotopic systematics of E101.1 imply that 16 O-rich and 16 O-poor reservoirs coexisted in the earliest solar protoplanetary disk and that igneous CAIs experienced a 16 Odepletion in an early high temperature stage. The coagulation of various lithological units in E101.1 and their partial assimilation supports models of CAI growth by competing fragmentation and coagulation in a partially molten state. Our results suggest that chemical and isotopic heterogeneities of unclear origin in regular CAIs may result from such a complex aggregation history masked by subsequent melting and recrystallization.

precursor may have coagulated with a 16 O-poor FUN CAI having 48 Ca anomalies, as indicated by perovskite, before subsequent extensive melting. The Al/Mg systematics (2.4 × 10 -5 ≤ ( Al !" / Al)

!" ! ′ ≤ 5.4 × 10 -5 , where ( Al !" / Al)

Introduction

As the first solids to have formed in our solar system, Calcium-Aluminum-rich Inclusions (CAIs) from chondritic meteorites provide a snapshot of the state of the solar system at its very beginning, probably as it was a dense protoplanetary disk, when the protosolar cloud envelope was still collapsing [START_REF] Yang | The effects of disk building on the distributions of refractory materials in the solar nebula[END_REF][START_REF] Mishra | Timing and extent of Mg and Al isotopic homogenization in the early solar system[END_REF][START_REF] Taillifet | Origin of refractory inclusion diversity by turbulent transport in the inner solar nebula[END_REF]. CAIs are of interest for astrophysics because their conditions of formation correspond to those prevailing in the region of telluric planet formation, within the first 1 AU from the protosun. However, in spite of almost five decades of intense study, the astrophysical context of CAI formation is still debated (e.g., [START_REF] Shu | X-rays and fluctuating X-winds from protostars[END_REF][START_REF] Itoh | Contemporaneous formation of chondrules and refractory inclusions in the early Solar System[END_REF][START_REF] Boss | Cosmochemical consequences of particle trajectories during FU Orionis outbursts by the early Sun[END_REF][START_REF] Yang | The effects of disk building on the distributions of refractory materials in the solar nebula[END_REF][START_REF] Taillifet | Origin of refractory inclusion diversity by turbulent transport in the inner solar nebula[END_REF]. Particular CAI properties that would be potentially useful for constraining astrophysical models include (1) their 16 O and 26 Al isotopic anomalies (e.g., [START_REF] Itoh | Contemporaneous formation of chondrules and refractory inclusions in the early Solar System[END_REF][START_REF] Mishra | Timing and extent of Mg and Al isotopic homogenization in the early solar system[END_REF][START_REF] Aléon | Oxygen isotopes in the early protoplanetary disk inferred from pyroxene in a classical type B CAI[END_REF], (2) their high temperature gasliquid-solid thermal histories (e.g., [START_REF] Richter | Conditions in the protoplanetary disk as seen by the type B CAIs[END_REF][START_REF] Shahar | Astrophysics of CAI formation as revealed by silicon isotope LA-MC-ICPMS of an igneous CAI[END_REF], and (3) their redox state and the oxygen fugacity in the ambient medium where they formed (e.g., [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF][START_REF] Simon | Valence of titanium and vanadium in pyroxene in refractory inclusion interiors and rims[END_REF]Grossman et al., 2008a;[START_REF] Paque | XANES and Mg isotopic analyses of spinels in Ca-Al-rich inclusions: Evidence for formation under oxidizing conditions[END_REF]. Up to now, most studies have been focused on one or two of these properties in isolation, however it has been recognized recently that many CAIs are compound objects that appear to have formed by aggregation of multiple precursors identified in the petrography (e.g., El Goresy et al., 2002;[START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Macpherson | Wellresolved variations in the formation ages for Ca-Al-rich inclusions in the early solar system[END_REF][START_REF] Ivanova | A compound Ca-Alrich inclusion from CV3 chondrite Northwest Africa 3118: Implications for understanding processes during CAI formation[END_REF] or inferred from the isotopic record (e.g., [START_REF] Simon | Calcium and titanium isotope fractionation in refractory inclusions: tracers of condensation and inheritance in the early solar protoplanetary disk[END_REF]. Such CAIs most likely grew by coagulation-fragmentation either in the solid or in a partially molten state [START_REF] Charnoz | Growth of calciumaluminum-rich inclusions by coagulation and fragmentation in a turbulent protoplanetary disk: Observations and simulations[END_REF], thus adding complexity in the interpretation of differing petrological, chemical and isotopic properties in terms of astrophysical environments and processes that could produce mixing of refractory precursors.

The goal of our study is to use multiple tracers of the conditions of formation of CAIs on a well characterized compound object in order to disentangle the various steps of its formation history and to shed light on its precursors and on the astrophysical origin of the 16 O and 26 Al isotopic anomalies. We focused on Efremovka 101.1 (hereafter E101.1), a CAI previously studied in detail for its petrography, trace element and, to a limited extent, 26 Al content by El Goresy et al. (2002). Notably, it contains several lithological units interpreted to be CAI xenoliths trapped in a host inclusion. This petrographic stratigraphy provides an opportunity to unravel the sequence of events that led to the CAI formation, growth, and thermal processing, whereas in many large CAIs extensive melting has left only cryptic traces of original formation events. We performed an in-situ systematic O, Mg, Si isotope study of all lithological units coupled with complementary petrographic observations, in order to have five isotopic tracers (O, Mg and Si mass fractionations, 16 O excesses and radiogenic 26 Mg excesses, hereafter 26 Mg*) to study the thermal history of all components, their relationships, and the interplay between thermal history and the record of the 16 O and 26 Al isotopic anomalies.

Sample and Methods

CAI Efremovka 101.1

E101.1 is a ~1.6 mm diameter Compact Type A (CTA) inclusion (Fig. 1a) from the Efremovka reduced CV3 chondrite, previously studied by El Goresy et al. (2002) for mineralogy, petrography, and Rare Earth Elements (REE) abundances. Details of its lithological units, including those defined by El Goresy et al. (2002) are given in supplementary information (Fig. S1) and listed in Table 1. Table 1 also summarizes the terminology and CAI-related acronyms used throughout the manuscript. E101.1 is dominated by melilite of typical composition between Åk 25 and Åk 30 and it contains numerous clusters of coarse-grained spinel (Fig. S1A). E101.1 exhibits two major properties that are uncommon among CAIs: [START_REF]A CAI with ultrarefractory REE patterns and enormous enrichments of Sc, Zr and Y in fassaite and perovskite[END_REF] the presence, in the same half of the inclusion, of several lithological units suggesting incorporation of xenoliths that were once individual CAIs free-floating in the nebula (El Goresy et al., 2002, Fig. S1B), some of which are associated with oxidized secondary phases, and (2) systematic ultrarefractory (UR) REE patterns associated with presence of an aluminous clinopyroxene containing up to 12.9 wt% Sc 2 O 3 and up to 5.4 wt% ZrO 2 . E101.1 is surrounded by a complete Wark-Lovering rim (WLR, [START_REF] Wark | Marker events in the early evolution of the solar system: evidence from rims on Ca-Al-rich inclusions in carbonaceous chondrites[END_REF] and exhibits portions of a forsterite-rich accretionary rim (AR, [START_REF] Krot | Forsterite-rich accretionary rims around Ca,Al-rich inclusions from the reduced CV3 chondrite Efremovka[END_REF] in its outer part (Fig. S1C). The UR-REE patterns are found throughout the inclusion in all lithological units including both rims and the xenoliths.

The most abundant xenoliths (more than 14) are made of sinuous fragments dominated by Al-diopside with chemistry similar to that often found in fine-grained spinel-rich CAIs or amoeboid olivine aggregates (Krot et al., 2004a,b) and locally including complex associations of fine-grained FeO-rich minerals such as andradite and hedenbergite with wollastonite (El Goresy et al., 2002, Fig. S1B). Fine-grained symplectite intergrowths of spinel, clinopyroxene and perovskite, reminiscent of the reaction texture found in E49, another compound CAI from Efremovka [START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF], are associated with the sinuous fragments and attributed to incorporation of small fine-grained CAI fragments (El Goresy et al., 2002).

Two areas dominated by melilite and surrounded by sinuous pyroxene fragments have been defined as subinclusions (El Goresy et al., 2002, Fig. S1B). The host inclusion contains several populations of Y-and Zr-rich perovskite grains distinguished by their ZrO 2 /Y 2 O 3 ratios. The largest perovskite grains are commonly associated with FeNi metal. The Sc-Zrrich pyroxene is systematically associated with perovskite and/or metal and defines reaction rims with the host melilite (El Goresy et al., 2002). Several regions at the interface between sinuous pyroxene and host melilite present melt pockets with compositions roughly intermediate between diopside and gehlenite (Fig. S1D). These melt pockets contain dendritic crystals of nearly pure gehlenite (Fig. S1D) and were described as an impact glass (El Goresy et al., 2002), hereafter referred to as the quenched glass. The innermost part of the host CAI is hereafter designated as the core. Three perovskites of differing ZrO 2 /Y 2 O 3 ratio were analyzed for Ca and Ti isotopes and show both excesses and depletions in 48 Ca (-7.5‰ ≤ δ 48 Ca ≤ +9.8‰) but, remarkably, all have δ 50 Ti ≈ 0‰. This unusual property is also found in FUN-type (Fractionated with isotopic anomalies of Unknown Nuclear origin) and hiboniterich CAIs from Murchison [START_REF] Kööp | A multielement isotopic study of refractory FUN and F CAIs: Mass-dependent and mass-independant isotope effects[END_REF]. The three analyzed perovskite belong to three differing ZrO 2 /Y 2 O 3 populations and their δ 48 Ca range is comparable to that of the FUN-type hibonites, apart from one having a δ 48 Ca value of +43‰ [START_REF] Kööp | A multielement isotopic study of refractory FUN and F CAIs: Mass-dependent and mass-independant isotope effects[END_REF]. El Goresy et al. (2002) reported early Mg isotopes measurements in spinel and melilite with a ~ 1‰ precision (1σ) and in anorthite with a ~4 to ~8‰ precision (1σ). Their Al/Mg systematics yielded a slightly sub-canonical isochron of slope 26 Al/ 27 Al 0 ~ 4.3 × 10 -5 in the sinuous fragments, scattered initial 26 Al/ 27 Al ratios in the host melilite and little to no 26 Mg excesses in anorthite, thus suggesting partial isotopic re-equilibration during or after 26 Al decay (El Goresy et al., 2002).

Petrographic analyses

Petrographic analyses were done using various electron probes and scanning electron microscopes (SEM) for comparison with isotopic analyses in the same phases. Maps of the distributions of Mg, Al, and Si in melilite were obtained using wavelength dispersive spectroscopy (WDS) with the Cameca SX5 and SX100 electron microprobes at the CAMPARIS facility in the Pierre et Marie Curie University in Paris. Additional chemical analyses of melilite within 10 µm of oxygen isotope analyses pits were obtained with the SX100 electron probe of the CAMPARIS facility. The small extent of some gehlenitic regions, notably around spinel, compared to the spot size of O isotopes may lead to artificial discrepancies in the relationship between melilite chemistry and O isotopes in some cases. To minimize this effect, the Åk content was determined from the Åk map in the exact location of the SIMS spots and this Åk content was preferred over the spot analyses in these cases with strong chemical gradients (Fig. S2). All electron probe analyses are available as supplementary information. Backscattered electron (BSE) images of ion microprobe crater pits, were obtained using the SEMs at the University of California Los Angeles (UCLA) and in the Geosciences Paris Sud (GEOPS) laboratory at the Orsay University, and with the field emission gun SEM (FEG-SEM) in the Geology department at the Ecole Normale Superieure (ENS) in Paris.

Ion microprobe analysis of O, Mg, and Si isotopes

All isotopic analyses were conducted using the Cameca IMS 1270 ion microprobe at UCLA in two sessions for O isotopes, one session for Mg isotopes, and two sessions for Si isotopes (Fig. 1b). Issues related with instrumental biases such as differences in instrumental mass fractionation (IMF) due to chemical variations in the target (matrix effects) and analyses overlapping on different mineral phases are discussed in supplementary information.

Uncertainties arising from these effects were taken into account and untrusted results were not considered in the discussion.

Oxygen isotopes

Oxygen isotope ratios were measured in multicollection mode using a ~ 10 µm primary Cs + beam of ~0.2 nA. Negative ions of 16 O, 17 O and 18 O were detected on an off-axis Faraday cup (FC), and on two electron multipliers (EMs, axial and an off-axis), respectively.

The mass resolving power (MRP, defined as M/∆M) was set at ~ 6500, largely sufficient to resolve the 16 OH -interference at mass 17. The normal incidence electron gun was used for charge compensation. After a 30 s presputtering with a 0.5 nA beam, 16 O, 17 O and 18 O were counted for 30 cycles of 10 s at typical count rates of 5 × 10 7 counts per second (cps), 2 × 10 4 cps and 1 × 10 5 cps for 16 O, 17 O and 18 O respectively. Liquid nitrogen was used to keep the vacuum at ~ 1 × 10 -9 Torr as a routine precaution but the plexiglas mount of E101.1 ensured limited outgassing. Data were corrected for IMF using a San Carlos olivine and a Burma spinel standard, which yielded similar IMF. All analyses were thus corrected with an average IMF. As a whole, matrix effects for O isotopes in the phases investigated here are negligible except for the silicate melt inclusion in metal and possibly for the FeO-rich silicates and the Sc-Zr-rich pyroxene (supplementary information). Results are reported using the standard δnotation, where δ 17,18 O = ( (Clayton and Mayeda, 1984) along a slope 0.52 line. They are given in ‰/atomic mass unit (‰/amu).

We obtained a total of 128 analyses, sampling most phases (melilite, spinel, forsterite, Al-diopside, FeO-rich assemblages, Sc-Zr-rich pyroxene, perovskite, impact glass, melt inclusion in metal, nepheline ± anorthite) in several lithological contexts including the host CAI, sinuous pyroxene fragments, subinclusions 1 and 2, WLR and AR. Two profiles starting from the rim were done toward the interior of the host CAI (profiles 1 and 2) (Fig. S1E). The core of the host CAI was also targeted, as well as two areas between the sinuous pyroxene fragments and the rim (areas 3 and 4) (Fig. S1E). Both spinel and melilite inside of and in proximity to spinel clusters (clusters 1 to 5) were measured (Fig. S1A); in these regions melilite is significantly more gehlenitic than the typical host melilite.

Magnesium isotopes

Magnesium isotopes were measured using a ~30 only. This instrumental mass fractionation law was preferred to that resulting from CAI evaporation in nature [START_REF] Davis | Isotopic mass fractionation laws for magnesium and their effects on 26 Al-26 Mg systematics in solar system materials[END_REF] because the instrumental effects between the standards were found to be larger than the natural effects observed in E101.1, thus yielding a potentially larger source of uncertainty if improperly corrected. The inaccuracy resulting from

this assumption is at most about 0.25‰ and is typically below 0.1‰ (Fig. S3). All isotopic results were subsequently reported in δ-notation as per mil deviations relative to the DSM3 terrestrial reference [START_REF] Galy | Magnesium isotope heterogeneity of the isotopic standard SRM980 and new reference materials for magnesium-isotope-ratio measurements[END_REF]. Measurements as an unknown of a synthetic pyroxene glass having ~ 1‰ 26 Mg excess (P1) yielded an average 26 Mg excess of 1.071 ± 0.080 ‰ (2σ). Al/Mg ratios were determined from the 27 Al + / 24 Mg + ratios corrected using the relative sensitivity factors (RSF) determined from Burma spinel (spinel), the P0 synthetic Al-Tipyroxene glass (Al-diopside, Sc-Zr-rich pyroxene, impact glass) and Madagascar hibonite (melilite). IMF for the 25 Mg/ 24 Mg ratios were calculated using Burma spinel for spinel analyses, the pyroxene P0 glass (or in one case the San Carlos pyroxene normalized to P0) for Al-diopside, Sc-Zr-rich pyroxene, impact glass and melilite and San Carlos olivine for forsterite. Burma spinel and P0 glass were previously measured relative to USNM forsterite, which composition relative to DSM3 was determined by [START_REF] Galy | Magnesium isotope heterogeneity of the isotopic standard SRM980 and new reference materials for magnesium-isotope-ratio measurements[END_REF]. We further assumed that San Carlos olivine has the same Mg isotopic composition as USNM forsterite.

Given the Earth mantle origin of both olivines, the uncertainty resulting from this assumption is much smaller than our analytical precision. In the absence of a proper melilite standard for Mg isotopes we initially assumed that hibonite could be used as a standard for the Al/Mg ratio of melilite, a reasonable approximation within uncertainties, and that the P0 glass could be used to determine the Mg isotopes IMF for melilite. Further matrix effect investigations indicate a 5 ± 0.5‰ matrix effect between P0 and Al-rich melilite for Mg isotopes, which was retroactively applied to the sample melilite data. The additional uncertainty due to this correction has been propagated and is included in the reported dataset. Al/Mg ratios for forsterite were not corrected for RSF but were close enough to 0 (average 0.015 ± 0.015, 2σ)

to have negligible effects on isochron regressions. Still, we note that the Al + /Mg + ratios on the San Carlos olivine were significantly lower than this (average 2.88 ± 0.06 × 10 -4 , 2σ), suggesting that minor amounts of Al-rich material (e.g., Al-diopside) were incorporated in the analyses in agreement with SEM imaging of the AR. The precisions obtained for Mg isotopic ratios depend on the Mg content: 2σ errors for δ 25 Mg incorporating both internal (within spot)

and external reproducibility on standards range between 1.06 ‰ and 1.75 ‰ for melilite and glass, and between 0.37 ‰ and 0.89 ‰ for other minerals. 2σ errors on δ 26 Mg* range between 0.06 and 0.95 ‰, again with larger uncertainties for the Al-rich, Mg-poor phases.

We obtained a total of 85 measurements of Mg isotope ratios from spinel, Al-diopside, Sc-Zr-rich pyroxene, impact glass and melilite. Care was taken to sample the host inclusion, the sinuous pyroxene fragments, subinclusion 1, the WLR and AR as well as five different spinel clusters. Adjacent melilite was measured in 4 of the 5 spinel clusters. Whenever possible, Mg isotopes spots were located within 20 µm of the O isotope analysis pits for a direct comparison. For example, the Mg isotope profiles 1 and 2 were started from the rim and were made in parallel with the O isotope traverse (Fig. S1E).

Silicon isotopes

Silicon isotopes were measured in multicollection mode as negative ions with a ~ 20 µm Cs + primary ion beam of 1-1.7 nA. 28 Si -and 30 Si -were detected simultaneously on two Faraday cups. 27 Al -was measured as well as a check on the composition of the analyzed mineral. The mass resolving power was set at 2400, sufficient to separate the 29 SiH - interference at mass 30. The normal incidence electron gun was used for charge compensation. Typical secondary ion intensities for 28 Si -ranged between 4 × 10 7 cps in melilite to 7 × 10 7 cps in olivine. Si isotopic compositions are reported as δ 30 Si per mil deviations relative to the NBS 28 international reference. With such conditions, a Poisson counting statistic ≤ ±0.1‰ is obtained after a few minutes integration (30 cycles of 5 s acquisition time and 30 s of presputtering). The IMF was corrected by using San Carlos olivine, P0 pyroxene glass, San Carlos pyroxene and new melilite standards specifically prepared for Si isotope analyses [START_REF] Marin-Carbonne | O, Si and Mg isotopic compositions of FUN inclusion Vigarano 1623-5[END_REF]. Matrix effects are particularly large for silicon isotopes in melilite [START_REF] Knight | Silicon isotopic fractionation of CAI-like vacuum evaporation residues[END_REF], which ionize poorly and have large IMF. This matrix effect is dependent on the Al content and was calibrated using the three melilite standards (Åk 10 , Åk 80 and a zoned crystal with Åk content ranging between 15 and 70 mol%) and the method described in [START_REF] Marin-Carbonne | O, Si and Mg isotopic compositions of FUN inclusion Vigarano 1623-5[END_REF]. Possible matrix effects in pyroxene remained within analytical uncertainty (supplementary information). Typical errors after all corrections including propagation of internal precision and external reproducibility determined on standards are 0.4-0.6 ‰ (2σ).

The mass fractionation in ‰/amu is expressed as FSi, where FSi = δ 30 Si / 0.501 where the factor 0.501 is the (M 29 Si-M 28 Si)/(M 30 Si-M 28 Si) ratio, where M x Si is the exact mass of the x Si isotope. [START_REF] Knight | Silicon isotopic fractionation of CAI-like vacuum evaporation residues[END_REF] report a mass fractionation factor of 0.518 ± 0.006 between δ 29 Si and δ 30 Si in evaporated CMAS (Ca, Mg, Al, Si) glasses analogs of CAIs, so that FSi as defined here is not strictly identical to δ 29 Si, however the difference is small with regard to our uncertainties so that it does not affect any interpretation and FSi can be compared with δ 25 Mg. Si isotopic analyses were acquired for 60 spots, preferentially in the vicinity of O and Mg isotope crater pits in order to allow as much intercomparison between the various isotopic systems as possible (within 20-30 µm). For example, Si isotope traverses were made in parallel with the O and Mg isotope profiles (Fig. S1E). Analyzed phases include olivine, pyroxene and melilite from different petrographic contexts of E101.1, including the host inclusion, sinuous pyroxene fragments, subinclusion 1, areas 3 and 4, and the WLR and AR.

Results

Petrographic observations

New electron probe analyses indicate that in spite of a typical melilite composition in the Åk 20 -Åk 30 range (El Goresy et al., 2002), melilite can reach nearly pure gehlenite composition (down to Åk 3 ) in contact with spinel, notably near the WLR as in other CAIs (e.g., [START_REF] Fahey | Microdistributions of Mg isotopes and REE abundances in a Type A calcium-aluminum-rich inclusion from Efremovka[END_REF][START_REF] Goswami | Ion microprobe studies of Efremovka CAIs: I. Magnesium isotope composition[END_REF][START_REF] Toppani | Wark-Lovering rims at the nanometer scale: a transmission electron microscopy study[END_REF][START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Kawasaki | Oxygen isotopic and chemical zoning of melilite crystals in a type A Ca-Al-rich inclusion of Efremovka CV3 chondrite[END_REF], but around spinel clusters in the interior as well (Fig. S4). These interior spinel clusters correspond to the end of spinel trails reaching the rim of the inclusion.

The electron probe mapping of Mg further shows that these spinel trails are systematically associated with aluminous melilite thus defining continuous to semi-continuous channels of aluminous melilite between the spinel clusters and the edge of the CAI (Fig. 2). This is notably the case of spinel clusters 1, 2 and 4 (Fig. 2). Being closer to the rim, cluster 3 is embedded in such a channel. Spinel clusters in the upper part of subinclusion 1 are also located at the end of such trails (Figs. 1 and2). Cluster 5 and the smaller clusters in the lower left part of Figure 1 also define a swarm of spinel associated with Al-rich melilite and connected to the edge of the CAI, although not as continuous as the upper trails (Fig. 2). Two of these small clusters are labeled as clusters 6 and 7 (Fig. 3). Detailed SEM examination of the spinels throughout the inclusion and especially of those connected to the rim via gehlenitic melilite reveals numerous signs of spinel resorption, which notably include [START_REF]A CAI with ultrarefractory REE patterns and enormous enrichments of Sc, Zr and Y in fassaite and perovskite[END_REF] rounded anhedral crystal shapes (Fig. 3a,b), (2) embayments (Fig. 3a-c), (3) a rim around spinel consisting of Al-, Ti-clinopyroxene ± spinel symplectite, locally with perovskite inclusions (Fig. 3). A mineralogical continuum is observed in these rims from pyroxene + abundant microspinel to pyroxene only to pyroxene + perovskite. The Ti content of the pyroxene can be used to map the symplectite (Fig. 3d). This symplectite is similar to that previously described in E49, another compound CTA inclusion from Efremovka [START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF]. Where the surrounding rim consists of clinopyroxene only, it is strongly reminiscent of the boundary clinopyroxene described around spinel in Type B CAIs and in the mantles of Type B1 CAIs [START_REF] Paque | The formation of boundary clinopyroxenes and associated glass veins in type B1 CAIs[END_REF]. This rim is also similar in textural relationship to the Sc-Zr-rich pyroxene rims around perovskite and metal grains (Fig. 4, El Goresy et al., 2002). As in the case of spinels, the perovskite rimmed by the Sc-Zr-rich pyroxene commonly shows embayments and anhedral shapes. FeNi metal is commonly associated with perovskite and rimmed by the Sc-Zr-rich pyroxene. It is predominantly high Ni taenite (35-51 at% Ni) but high Co kamacite (up to 8 at% Co) is also found, usually in heterogeneous grains (see supplementary informations). One grain contains 5-6 at% platinum group elements. Most grains contain inclusions of Fe-V-oxides and Ca-phosphates near their periphery. A troilite inclusion is found once. Three Si-rich glass inclusions were found in one of the metal grains (Fig. 4a).

Although E101.1 was initially described as mostly unaltered, we found nepheline associated with anorthite in replacement of melilite, notably in the vicinity of perovskite + metal nodules (Fig. 4).

Oxygen isotopes

Oxygen isotope ratios in E101.1 (Table S1) plot along the Carbonaceous Chondrites Anhydrous Minerals mixing line (Clayton and Mayeda, 1984) and span the whole range between the typical 16 O-rich composition of spinel in CAIs (∆ 17 O ~ -22 ‰) and a subterrestrial 16 O-poor composition (up to ∆ 17 O ~ -2 ‰, Fig. 5a).

Spinel and accretionary rim forsterite are uniformly 16 O-rich with ∆ 17 O ≤ -20.5 ‰.

The isotopic composition of melilite varies from significantly enriched in 16 O (∆ 17 O down to -19.8 ‰ in the WLR, Fig. 5c) to relatively 16 O-poor (∆ 17 O up to -1.9 ‰). The most 16 O-rich melilite is systematically gehlenitic melilite associated with spinel (Fig. 6), although the reverse is not systematic. This is true for melilite from the WLR and also for gehlenitic melilite associated with coarse-grained spinel clusters in the interior (Fig. 6b). Interestingly, the Åk content of melilite associated with spinel clusters define a trend similar to that in melilite as a function of distance to WLR spinel in profiles 1 and 2 (Fig. 6c). et al. (2002) to characterize several populations of perovskite (Fig. 9).

Minerals in the WLR and the forsterite-rich accretionary rim are rich in 16 O as often observed in CAIs [START_REF] Krot | Existence of an 16 Orich gaseous reservoir in the solar nebula[END_REF][START_REF] Yoshitake | Correlations between oxygen-isotopic composition and petrologic setting in a coarse-grained Ca, Al-rich inclusion[END_REF][START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Bodénan | An oxygen isotope study of Wark-Lovering rims on Ttype A CAIs in primitive carbonaceous chondrites[END_REF] ! ratio is defined as the slope of an isochron going through a given analysis and through the origin.

In the interior, spinels from subinclusion 1, cluster 1 and cluster 5 have ( Al

!" / Al) !" ! !
ratios consistent with a canonical ratio (i.e., 5.23 × 10 -5 , [START_REF] Jacobsen | 26 Al-26 Mg and 207 Pb-206 Pb systematics of Allende CAIs: Canonical solar initial 26 Al/ 27 Al ratio reinstated[END_REF] within error (weighted means of (5.0±0.4) × 10 -5 , (5.3±0.9) × 10 -5 , (5.6±0.4) × 10 -5 , 2σ, respectively, Table 2), but cluster 2 has a distinctly supracanonical weighted mean ((7.3±0.9) × 10 -5 , 2σ, Table 2). Near the CAI periphery, cluster 3 has a subcanonical ( Al !" / Al)

!" ! ! ratio of (3.6±1.2)

× 10 -5 (2σ, Table 2).

Individual ( Al !" / Al)

!" ! ! ratios in melilite range between (5.4±0.2) × 10 -5 and (2.4±0.5)

× 10 -5 (2σ) suggesting secondary perturbations owing to Mg diffusion and isotope exchange in an initially canonical melilite. No clear difference can be observed between different regions in the CAI (Fig. 10), with similar average and spread in ( Al !" / Al)

!" ! ! ratios of 3.7±0.5, 4.7±0.6, 4.1±0.9, 4.2±0.3, 4.0±0.5, 3.4±0.7, 4.0±0.7 for area 3, area 4, subinclusion 1, core, melilite in the vicinity of spinel clusters, profile 1 and profile 2, respectively (weighted mean given in units of 10 -5 ; uncertainties are 1 standard deviation here rather than 2 standard errors to take into account heterogeneities in melilite within individual areas and to emphasize the similarity of the ratios). In subinclusion 1 and spinel clusters 1 and 5, melilite associated with spinel have comparable ( Al !" / Al)

!" ! ! ratios between 4.1 × 10 -5 and 4.3 × 10 -5 , while associated spinels have higher ratios consistent with the canonical ratio (Table 2). Only melilite associated with sub-canonical spinel (cluster 3) has a ( Al !" / Al)

!" !

! ratio comparable to that of the associated spinel ((3.3±0.3) × 10 -5 and (3.6±1.2) × 10 -5 , respectively, 2σ, Table 2). This corresponds to systematic sub-canonical slopes and elevated intercepts in Al/Mg isochrons, with the highest intercept for cluster 2 (Fig. 11), except for cluster 3 where the intercept is consistent with 0. Regressions for subinclusion 1, cluster 1 and cluster 5 are identical within error (Fig. S5).

Due to large uncertainties, all estimates of the initial 26 Al/ 27 Al ratio of the sinuous pyroxene are consistent with the canonical ratio (Fig. 10) and with the average melilite value at the 2σ level.

In the WLR, ( Al !" / Al)

!" ! ! ratios range from canonical ((5.7±0.4) × 10 -5 ) to distinctly subcanonical ((3.7±0.5) × 10 -5 ) in spinel. If both spinel and pyroxene are considered together a regression yields a slope of (4.5±0.8) × 10 -5 (2σ), which becomes (4.8±0.5) × 10 -5 (2σ) if forced through the origin. This is identical to the weighted mean of ( Al !" / Al)

!" ! ! ratios of (4.8±0.2) × 10 -5 (2σ, Table 2). Given that variations in spinel should affect the determination of the slope, we also investigated the 26 Al/ 27 Al systematics of the WLR without spinel.

Pyroxene data are individually consistent with the canonical ratio and a regression yields a slope of (4.1±3.0) × 10 -5 (2σ), which becomes (6.0±0.9) × 10 -5 (2σ) if forced through the origin.

The weighted mean of 26 et al., 2012;[START_REF] Mishra | Timing and extent of Mg and Al isotopic homogenization in the early solar system[END_REF].

Mass fractionation.

δ 25 Mg values vary between 6.7±1.8 ‰ and -4.7±0.7 (2σ, Table S2). This ~ 12‰ range is unusual in normal (i.e., non-FUN) CAIs, which commonly have more homogeneous compositions (e.g., [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF][START_REF] Kita | Internal 26 Al-26 Mg isotope systematics of a Type B CAI: Remelting of refractory precursor solids[END_REF][START_REF] Bullock | Mg and Si isotopic fractionation patterns in types B1 and B2 CAIs: Implications for formation under different nebular conditions[END_REF]. The Mg isotopic compositions of the E101.1 lithological units span a large range between that of light hibonite CAIs from CM chondrites (δ 25 Mg values as low as -7‰, [START_REF] Liu | A heterogeneous solar nebula as sampled by CM hibonite grains[END_REF] and that of classical igneous Type A and B CAIs enriched in heavy isotopes (-1‰ ≤ δ 25 Mg ≤ +11‰, [START_REF] Davis | Condensation and evaporation of solar system materials[END_REF].

The most negative values are systematically found in the sinuous pyroxene fragments

Silicon isotopes

The δ 30 Si values in E101.1 span a very large range, from -9.0 ± 1.3 to +19.9 ± 1.3, which corresponds to FSi values between -4.5 ± 0.6 ‰/amu and +10.0 ± 0.6 ‰/amu where errors correspond to 2σ uncertainties (Table S3). This ~ 15‰/amu range is comparable to that of δ 25 Mg or slightly larger. Again it is highly unusual in non-FUN CAIs, which commonly have a more homogeneous composition [START_REF] Shahar | Astrophysics of CAI formation as revealed by silicon isotope LA-MC-ICPMS of an igneous CAI[END_REF][START_REF] Bullock | Mg and Si isotopic fractionation patterns in types B1 and B2 CAIs: Implications for formation under different nebular conditions[END_REF]. The lightest Si isotopic composition (most negative FSi values) corresponds to the strongest negative mass fractionation ever reported for a CAI. By contrast, the heaviest Si isotopic compositions of E101.1 reach those of the heavy (although not the heaviest) FUN CAIs (having δ 29 Si values commonly larger than +5 ‰, [START_REF] Davis | Condensation and evaporation of solar system materials[END_REF]. This corresponds to a positive mass fractionation even stronger than that for Mg.

As is the case for Mg isotopes, the most negative FSi values are also found in the sinuous pyroxene fragments (-4.5 ± 0.6 ‰/amu to -2.4 ± 0.6 ‰/amu). FSi values in the host melilite are by contrast systematically positive and range between +3.9 ‰/amu and +8.5 ‰/amu with an average and representative value at +5.9‰/amu, approaching the transition between normal and FUN CAIs [START_REF] Davis | Condensation and evaporation of solar system materials[END_REF]. Melilite in areas in close vicinity to the sinuous pyroxene fragments (sub1, area 3 and 4) have FSi values intermediate between those of the host melilite and those of the sinuous pyroxene, although they can reach the host melilite values (+1.2 ±0.4 ‰/amu ≤ FSi ≤ +8.8 ± 0.6 ‰/amu). AR olivines have slightly negative FSi values within 2σ errors of the average value (FSi between -1.9 ±0.4 ‰/amu and -0.5 ± 0.6 ‰/amu, average FSi= -1.2 ± 0.6‰/amu) and pyroxenes in the WLR have FSi values ranging between those of the AR and those of the interior melilite (-0.3±0.3 ‰/amu ≤ FSi ≤ +6.2 ± 0.6 ‰/amu) with one value at +10.0 ± 0.6 ‰/amu. The latter analysis corresponds to Ti-rich pyroxene finely intergrown with spinel, interior to the Al-diopside layer of the WLR. It must be treated with caution as the analysis overlapped significantly with spinel.

Profiles

Possible isotopic variations between the interior of the CAI and its outer part were investigated with two profiles, starting from the rim toward the interior, and with the areas 3 and 4 (see Fig. 1A). Due to large ion beam spot sizes compared to the usual WLR layers, rim analyses were acquired where the thickness allowed it. No transect was done within the WLR and the profiles started inside the spinel layer without including the WLR. The chemical composition of melilite shows large variations between Åk 20 and Åk 40 (profile 1, Fig. 13a) in the CAI interior, (also visible in the Mg map; Fig. 2). In the outermost 200 µm the composition of melilite is rather homogeneous at ~Åk 25 , except in the vicinity of the rim, where the Åk content suddenly drops towards nearly pure gehlenitic composition. This drop occurs ~30 µm inward from the WLR in profile 1 and ~70 µm inward from the rim in profile 2 (Fig. 13a). This enrichment in gehlenite toward the rim is commonly observed in Type A

CAIs (e.g., [START_REF] Fahey | Microdistributions of Mg isotopes and REE abundances in a Type A calcium-aluminum-rich inclusion from Efremovka[END_REF][START_REF] Goswami | Ion microprobe studies of Efremovka CAIs: I. Magnesium isotope composition[END_REF][START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF]Simon et al., 2011[START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF][START_REF] Katayama | Oxygen isotopic zoning of reversely zoned melilite crystals in a fluffy type A Ca-Al-rich inclusions from the Vigarano meteorite[END_REF][START_REF] Kawasaki | Oxygen isotopic and chemical zoning of melilite crystals in a type A Ca-Al-rich inclusion of Efremovka CV3 chondrite[END_REF], 2017).

The change in Mg content of melilite along profiles 1 and 2 is correlated with an isotopic evolution. Along profile 1, the δ 25 Mg value is around +4.5‰ in the interior of the CAI, where melilite is the most magnesian. In the Åk 25 zone, it is rather homogeneous around + 5.5‰ until the point where the Åk content drops (Fig. 13b). Unfortunately, it was not possible to acquire a Mg isotope analysis in the 30 µm region immediately inside the WLR without overlapping with spinel. However, because the point where the Åk content drops in profile 2 is further away from the rim at ~70 µm toward the interior, several analyses were obtained in this particular region. In the interior part of profile 2, the δ 25 Mg value barely decreases from +6.7‰ down to +5.1‰ until the 70 µm boundary (Fig. 13b). In the outer 70 µm, it decreases from + 5‰ down to + 2‰ (Fig. 13b). Such an evolution from heavy, positive values in the interior toward values closer to 0 has already been observed in some Type A CAIs [START_REF] Goswami | Ion microprobe studies of Efremovka CAIs: I. Magnesium isotope composition[END_REF][START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF], although the reverse has also been observed (e.g., CAI Leoville 144A, [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF]. Areas 3 and 4 also show an isotopic evolution with distance but are generally much closer to 0, with δ 25 Mg values evolving from about +2‰ down to ~0‰ toward the rim (Fig. 13b).

This coupled chemical -stable Mg isotopic evolution of melilite is also roughly associated with a radiogenic Mg isotopic evolution. ( Al !" / Al)

!" ! ! ratios are scattered between ~ 4 x 10 -5 and ~ 2 x 10 -5 in the interior of the CAI but increase toward the rim, notably outward of the 70 µm limit in profile 2, to reach roughly canonical values near the rim (Fig. 13c) as also found by [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF] for CAIs with supracanonical ratios in their interior.

By contrast, the Si isotopic composition does not show any appreciable evolution with distance from the interior to the rim (Fig. 13d). Profiles 1 and 2 are flat with homogeneous FSi values typical of the interior CAI near ~+6‰/amu. Areas 3 and 4 also have seemingly homogeneous compositions but at different FSi values near +3‰ (area 4) and +1‰ (area 3), although only 1 value without significant overlap is available for the latter.

Similarly to Si, the O isotope profiles do not show clear trends along the profile. ∆ 17 O values along profile 1 cluster around -6‰ with a significant spread as discussed above. Still, the outermost two analyses show ∆ 17 O values near -2‰ near the rim (Fig. 13e). Profile 2

shows comparable values (Fig. 13e) but a slight decrease in ∆ 17 O is observed toward the exterior before increasing again in the outer 70 µm. This increase and the least negative values of profile 1 may indicate a systematic increase in ∆ 17 O near the rim. Such an increase in ∆ 17 O is uncommon in Type A CAIs (e.g., [START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Kawasaki | Oxygen isotopic and chemical zoning of melilite crystals in a type A Ca-Al-rich inclusion of Efremovka CV3 chondrite[END_REF][START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF] but has been observed previously (e.g., CAI Allende A37, Simon et al., 2011[START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF]. Areas 3 and 4 show significantly lower ∆ 17 O values (near -8‰, see above), again with some scatter. Contrary to profiles 1 and 2, ∆ 17 O values tend to decrease toward the exterior in area 3 and 4. The mass fractionation relative to the CCAM line is highly variable and does not show any systematic trends (Fig. 13f).

Discussion

The coupled isotopic and petrologic observations summarized above point to a complex origin for E101.1. Here, we present model calculations to determine the nature of the E101.1 precursor materials based on observed isotopic compositions. The models reported in sections 4.1 and 4.2 are depicted schematically in Figure . 14. Theoretical endmember components are listed in Table 1; their compositions are given in supplementary information (Table S4).

Origin of the sinuous pyroxene

Thermal history of the sinuous pyroxene

The very light Si and Mg isotopic compositions of the sinuous pyroxene fragments with δ 25 Mg and FSi values both of about -3.5‰, are unusual in igneous CAIs, which are commonly enriched in heavy isotopes due to Rayleigh distillation during evaporation (e.g., [START_REF] Clayton | Isotopic variations in the rock-forming elements in meteorites[END_REF]Grossman et al., 2008b). These compositions are drastically different from those of the surrounding host CAI, which is enriched in heavy Si and Mg isotopes. This is indicative of a strong degree of disequilibrium and confirms the hypothesis inferred from textural and petrographic analysis (El Goresy et al., 2002) that the sinuous pyroxene lithologies are indeed xenolithic fragments that were captured by the main host and escaped complete assimilation.

Isotopically light values of Mg have already been reported for many platy hibonite crystals (PLACS) from the Murchison and Paris CM chondrites, with δ 25 Mg values occasionally reaching -7‰ and possibly as low as -10‰ [START_REF] Fahey | Microdistributions of Mg isotopes and REE abundances in a Type A calcium-aluminum-rich inclusion from Efremovka[END_REF][START_REF] Ireland | Correlated morphological, chemical and isotopic characteristics of hibonites from the Murchison carbonaceous chondrite[END_REF][START_REF] Ireland | Presolar isotopic and chemical signatures in hibonite-bearing refractory inclusions from the Murchison carbonaceous chondrite[END_REF][START_REF] Sahijpal | Ti, and Ca isotopic compositions and refractory trace element abundances in hibonites from CM and CV meteorites: Implications for early solar system processes[END_REF][START_REF] Liu | A heterogeneous solar nebula as sampled by CM hibonite grains[END_REF], Kööp et al. 2016a,b). The proposed interpretation for such low δ 25 Mg values is kinetic isotopic effects during condensation [START_REF] Liu | A heterogeneous solar nebula as sampled by CM hibonite grains[END_REF], at a temperature lower than the equilibrium condensation temperature. The competition between kinetic and equilibrium effects has been investigated theoretically for the evolution of Mg isotopic composition during condensation of forsterite [START_REF] Richter | Timescales determining the degree of kinetic isotope fractionation by evaporation and condensation[END_REF] and experimentally for the evaporation of a melt of CAI composition [START_REF] Richter | Elemental and isotopic fractionation of type B CAI-like liquids by evaporation[END_REF]. The amount of Mg isotope mass fractionation due to kinetic effects can reach -10 to -14‰/amu, depending on temperature. A contribution of Mg and Si condensed under thermal disequilibrium thus appears to be a possible explanation for the negative δ 25 Mg and FSi of the sinuous fragments.

A key parameter in the extent of the kinetic effects is the ratio between the condensation timescale and the timescale for the temperature change of the ambient medium [START_REF] Richter | Timescales determining the degree of kinetic isotope fractionation by evaporation and condensation[END_REF]. If the condensation is much faster than the change in ambient temperature, then condensation mostly proceeds at equilibrium and the isotopic fractionation is negligible in the considered range of temperatures. By contrast, if the change in ambient temperature is faster than condensation, then condensation proceeds at a temperature lower than expected (undercooling) and large kinetic isotopic fractionation effects are to be expected [START_REF] Richter | Timescales determining the degree of kinetic isotope fractionation by evaporation and condensation[END_REF][START_REF] Simon | Stable calcium isotopic composition of meteorites and rocky planets[END_REF][START_REF] Simon | Calcium and titanium isotope fractionation in refractory inclusions: tracers of condensation and inheritance in the early solar protoplanetary disk[END_REF]. Kinetic isotopic fractionation during condensation thus appears as a possible explanation for the negative δ 25 Mg and FSi values of the sinuous pyroxene if condensation proceeded during a rapid temperature drop [START_REF] Richter | Timescales determining the degree of kinetic isotope fractionation by evaporation and condensation[END_REF] allowing a significant degree of undercooling [START_REF] Simon | Stable calcium isotopic composition of meteorites and rocky planets[END_REF][START_REF] Simon | Calcium and titanium isotope fractionation in refractory inclusions: tracers of condensation and inheritance in the early solar protoplanetary disk[END_REF]. Either the sinuous pyroxene contains a small fraction of a strongly fractionated component or a large fraction of a component having a few per mil fractionation (e.g. a few degrees of undercooling as calculated for Ca and Ti isotopes, [START_REF] Simon | Calcium and titanium isotope fractionation in refractory inclusions: tracers of condensation and inheritance in the early solar protoplanetary disk[END_REF]. Such a temperature drop could be due to fast removal by rapid transport of condensing precursors from a region of high temperature in the protoplanetary disk, where Mg and Si start to condense in solids, to a region where the ambient temperature is suitable for pyroxene condensation. Alternatively rapid temperature fluctuations are required.

Interestingly, the sinuous pyroxene exhibits the UR REE pattern typical of E101.1, wherein the relative abundances of REEs are controlled by their respective volatilities, the most refractory REEs being the most abundant. The UR REE pattern is established by condensation as solid solutions in the most refractory minerals that can accommodate REEs, namely ZrO 2 , Y-oxides, hibonite and perovskite. As hibonite is the first major condensate, the UR pattern is established in the first few degrees below the onset of hibonite condensation [START_REF] Simon | A unique ultrarefractory inclusion from the Murchison meteorite[END_REF][START_REF] Davis | Condensation and evaporation of solar system materials[END_REF]. Such REE patterns can only be preserved in the highest temperature condensate (hibonite ?) if it is rapidly removed from the gas. One may expect a significant amount of thermal disequilibrium during this removal of the UR-carrier from the ambient gas, which possibly accounts for Ca and Ti isotopic mass fractionation [START_REF] Davis | Titanium isotopes and rare earth patterns in CAIs: Evidence for thermal processing and gas-dust decoupling in the protoplanetary disk[END_REF]Simon and De Paolo, 2010;[START_REF] Simon | Calcium and titanium isotope fractionation in refractory inclusions: tracers of condensation and inheritance in the early solar protoplanetary disk[END_REF]. Whether it also accounts for the negative Mg and Si isotopic fractionation is unclear. A separate condensation event may be required due to the large difference (~200 K) in condensation temperatures between

Ca and Ti on one hand, and Mg and Si on the other hand. Examination of UR-CAIs for which both REE patterns and Mg isotopes are available indeed shows that CAIs with UR REE patterns often have negative δ 25 Mg values between 0 and -10‰, albeit with relatively large uncertainties [START_REF] Hinton | A chemical and isotopic study of hibonite-rich refractory inclusions in primitive meteorites[END_REF][START_REF] Fahey | Hibonite-hercynite inclusion HH-1 from the Lancé (CO3) meteorite: The history of an ultrarefractory CAI[END_REF][START_REF] Simon | A unique ultrarefractory inclusion from the Murchison meteorite[END_REF][START_REF] Simon | A hibonite-corundum inclusion from Murchison: A first-generation condensate from the solar nebula[END_REF][START_REF] Liu | Isotopic records in CM hibonites: Implications for timescales of mixing of isotope reservoirs in the solar nebula[END_REF]). The only known UR CAI with δ 25 Mg > 0 ‰ (CAI 3483-3-10 in Mighei, MacPherson and [START_REF] Macpherson | Refractory inclusions in the prototypical CM chondrite, Mighei[END_REF]) is an altered fragment with a WLR. The combination of heavy Mg isotopic composition and presence of a WLR suggests that it once was an igneous CAI having undergone evaporation.

One possible scenario to explain the REE and isotopic data is that the sinuous pyroxene formed from hibonite-rich precursors that were rapidly removed from the high temperature regions where the UR REE pattern was established and then transported to a region where the temperature, about 200 K lower, was favorable for diopside condensation.

The rapid transport would have resulted in preservation of the UR REE pattern at the onset of removal and in subsequent kinetic isotopic fractionation during Si and Mg condensation. Åk 10 would still be too Si-rich to provide a match for the ∆ 17 O-FSi mixing line. Rather, this suggests that the mixing occurred between Al-diopside and a CMAS melt having isotopic compositions equal to those observed in melilite. With the final assumption that such a melt had a solar Ca/Al ratio, as expected for most CAIs (e.g., [START_REF] Simon | A preferred method for the determination of bulk compositions of coarse-grained refractory inclusions and some implications of the results[END_REF]Grossman et al., 2008b) 1). The two points in subinclusion 1 not matched by this mixing are instead well reproduced by using an endmember of almost pure gehlenitic composition. This suggests that the CMAS melt was partly crystallized as gehlenite within subinclusion 1, when the sinuous pyroxene was captured. The existence of such a partially crystallized region may explain why some sinuous pyroxene fragments distribute around subinclusion 1. The observed compositions are reproduced if 2 to 5 mol% of sinuous pyroxene was assimilated in areas 3 and 4, while this proportion reach 10 mol% in subinclusion 1.

Most likely all the sinuous pyroxene fragments derived from a single proto-CAI as indicated by ( 1) similar, unusual mineralogy and isotopic compositions and (2) the mixing lines between the sinuous fragments, areas 3 and 4, subinclusion 1 and the main host, which do not require more than one trapped xenolith. Detailed examination of the REE patterns (El Goresy et al., 2002) suggests that the sinuous pyroxene fragments are genetically related to the rest of the E101.1 CAI. Indeed, all UR CAIs have distinct REE patterns that record complexities in condensation events. These differences are seen most clearly in the abundances of light REE (LREE) relative to those of the heavy REE (HREE) and in the relative proportions of Lu, Er and Ho, the most refractory REEs. In E101.1, all units except a few Sc-Zr-rich pyroxene spots have abundant LREE relative to the HREE. In addition, Lu, Er and Ho are present in comparable abundances in E101.1. Given that these elements have different volatilities, direct condensation should result in Lu > Er > Ho [START_REF] Davis | Condensation and fractionation of rare earths in the solar nebula[END_REF][START_REF] Simon | A unique ultrarefractory inclusion from the Murchison meteorite[END_REF][START_REF] Davis | Condensation and evaporation of solar system materials[END_REF]. Their comparable or reversed abundances are usually attributed to fractional condensation and removal of an UR mineral with uniform REE activity coefficients, such as ZrO 2, that would preferentially remove Lu and then Er from the gas before establishment of the UR pattern in the main hibonite or perovskite condensate [START_REF] Simon | A unique ultrarefractory inclusion from the Murchison meteorite[END_REF]. The sinuous pyroxene fragments share both the LREE and Lu/Er/Ho features with the rest of E101.1 indicating that they may have been initially related with the host inclusion, perhaps by common UR precursors in spite of a totally different thermal history.

It has been suggested that these diopside-rich xenoliths could represent a former generation of WLR (El Goresy et al., 2002). However, our data do not allow a firm identification of the sinuous pyroxene precursor. Several features indeed compare well with WLRs such as its layered structure, the Al-diopside chemistry of pyroxene commonly found as the outer layer of WLRs and the common association of diopside with anorthite. It is not clear if the sinuous pyroxene is a former WLR fragment but several hypotheses can probably be discarded. The sinuous pyroxene is unlikely to be a fragment from the E101.1 WLR pushed inside because its Mg and especially Si isotopic composition are more negative than those found in the WLR of E101.1 and because the latter contains a Ti-rich fassaitic pyroxene layer inside diopside instead of anorthite (El Goresy et al., 2002). The sinuous pyroxene is also unlikely to be a remant of WLR associated with subinclusion 1. If our analysis is correct, subinclusion 1 is just a region of the host resulting from assimilation of a large fraction of sinuous pyroxene. Its rounded shape is probably fortuitous. Furthermore, in several places the sinuous pyroxene fragments around subinclusion 1 contain diopside in contact with subinclusion 1 and anorthite outside, which is the reverse of the commonly observed WLR sequence. It further lacks the spinel layer ubiquitous in WLR.

A similar approach can be used to investigate if the quenched glass areas with relict pyroxene and skeletal gehlenite (Fig. S1) derive from a secondary impact melting event, as initially proposed (El Goresy et al., 2002), or represent a residual melt from the trapping and partial assimilation of the sinuous fragments. Here caution must be taken as ( 1) only one Mg isotope analysis (Mg36) and no Si isotope analysis has been performed because numerous fine-grained mineral inclusions in the glass prevented more analyses, and (2) matrix effects in Mg isotopic analyses in glass relative to stoichiometric pyroxene or melilite may result in incorrect Mg isotopic ratios. If the analysis is treated as a pyroxene glass, the mixing calculation shows that the O and Mg isotopic compositions of the quench melt regions are matched by the mixing hyperbola calculated for the refractory CMAS parental-melt described above (Fig. S6), in which case the quench glass may be a residual melt. In contrast, if a matrix effect intermediate between that of pyroxene and melilite is applied to account for the glass composition intermediate between diopside and melilite, then the mixing is more consistent with mixing the sinuous pyroxene with a melilite component ~ Åk 30 in composition (Fig. S6).

In this case an impact origin is possible. Dendrites in this melt require rapid cooling, which may imply a two stage history. Note that the presence of Fe in the quenched glass (FeO up to 0.9 wt%) indicates a contribution of the FeO-rich regions associated with the sinuous pyroxene, which may have locally changed the composition of the mixture toward a less refractory and more 16 O-poor composition. The presence of FeO in the glass indicates the FeO-rich regions must have predated formation of the quenched glass (see section 4.6).

Thermal history of the main host

The main host precursor as inferred from the mixing calculation between Al-diopside and present-day host melilite is a highly refractory CMAS melt enriched in the heavy Si and Mg isotopes. The Si and Mg isotopic fractionations imply a significant amount of evaporation, but it is not clear whether this evaporation was sufficient to explain the ultrarefractory composition of the CMAS parental melt that was highly depleted in Si and

Mg. The Si isotopic composition (FSi ~+ 6.5 ‰) is comparable with that of several F-and FUN inclusions (e.g., TE, CG-14, EK-1-4-1, [START_REF] Clayton | Isotopic variations in the rock-forming elements in meteorites[END_REF][START_REF] Mendybaev | Experimental evaporation of mg-and Si-rich melts: Implications for the origin and evolution of FUN CAIs[END_REF] as well as that of evaporation residues produced in the conditions expected for the formation of FUN CAIs [START_REF] Mendybaev | Experimental evaporation of mg-and Si-rich melts: Implications for the origin and evolution of FUN CAIs[END_REF][START_REF] Mendybaev | Thermal and chemical evolution in the early Solar System as recorded by FUN CAIs: Part II -Laboratory evaporation of potential CMS-1 precursor material[END_REF]. However, the Mg isotopic composition (δ 25 Mg ~ +5‰) is closer to that of regular CAIs, whereas [START_REF] Mendybaev | Experimental evaporation of mg-and Si-rich melts: Implications for the origin and evolution of FUN CAIs[END_REF][START_REF] Mendybaev | Thermal and chemical evolution in the early Solar System as recorded by FUN CAIs: Part II -Laboratory evaporation of potential CMS-1 precursor material[END_REF] show that The results presented in Fig. 16, show first that the contribution of captured spinel does not change significantly the composition of the parental melt (step 2) relative to the products of equilibrium condensation as calculated in Grossman et al. (2008b). In each of the three cases, it is close to the first condensate expected at high temperature, near the field of melilite + krotite (CaAl 2 O 4 ) condensation. However, the heavy Si and Mg isotopes indicate that this cannot be the initial composition of the parental E101.1 host melt. Second, using any set of kinetic fractionation factors yields comparable results. All six step 3 pre-evaporative compositions calculated with any amount of spinel removed at step 2 are found to match nicely with the theoretical composition of a condensate (Fig. 16) in the hibonite + perovskite + melilite field at 10 -3 to 10 -4 bar [START_REF] Mendybaev | Thermal and chemical evolution in the early Solar System as recorded by FUN CAIs: Part II -Laboratory evaporation of potential CMS-1 precursor material[END_REF].

The precursor material of the E101.1 parental melt is thus consistent with an extremely refractory condensate, much more refractory than typical Type A CAIs, that subsequently underwent evaporation. In the model without spinel assimilation, the fractions of Si and Mg lost are 47% and 42%, respectively, using the fractionation factors of [START_REF] Knight | Silicon isotopic fractionation of CAI-like vacuum evaporation residues[END_REF] and 47% and 29%, respectively using those of [START_REF] Mendybaev | Experimental evaporation of mg-and Si-rich melts: Implications for the origin and evolution of FUN CAIs[END_REF]. Assuming that the Mg-Si isotopic decoupling is due to condensate spinel assimilation, the fraction of Mg lost reaches 66% / 50% (depending on the fractionation factor) for an initial δ 25 Mg value of +10‰ and 80% / 64% for an initial δ 25 Mg value of +15‰.

The highly refractory composition of the precursor condensate, in the stability field of hibonite, perovskite and melilite, is well in line with the presence of large perovskite grains. It also agrees with the UR REE pattern, which points to a condensation temperature range corresponding to that of hibonite and perovskite, the highest temperature major condensates that carry REE (e.g., [START_REF] Lodders | Solar system abundances and condensation temperatures of the elements[END_REF]. The excesses and depletions in 48 Ca (-7.5‰ ≤ δ 48 Ca ≤ Murchison [START_REF] Kööp | A multielement isotopic study of refractory FUN and F CAIs: Mass-dependent and mass-independant isotope effects[END_REF]. Together with the calculated hibonite-rich precursor composition and the UR pattern, often attributed to hibonite condensation (e.g. [START_REF] Davis | Titanium isotopes and rare earth patterns in CAIs: Evidence for thermal processing and gas-dust decoupling in the protoplanetary disk[END_REF], this suggests a link between the precursor of E101.1 and hibonite-rich FUN inclusions.

Whether the fractionated component (F) is initially the same as the carrier of the nuclear anomalies in 48 Ca in E101.1 is unclear because the subsequent capture of xenolithic material renders the characterization of the initial material subject to uncertainties. The precursor material of E101.1 can be viewed as an ultrarefractory FUN inclusion resulting from extensive melting and evaporation of an UR condensate precursor with isotopic anomalies of nucleosynthetic origin. Alternatively, the UR component with nuclear anomalies (the UN component) recorded in the pristine condensate perovskite may have been aggregated to a highly mass fractionated component (the F component) before they were partially melted together. In this case the F and UN components could have been initially decoupled as suggested by the existence of F CAIs without nuclear isotopic anomalies (Clayton et al., 1984;[START_REF] Komorowski | The White Angel: A unique wollastonite-bearing, mass-fractionated refractory inclusion from the Leoville CV3 carbonaceous chondrite[END_REF][START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF]. The lack of large mass fractionation in O isotopes in perovskite may support this conclusion. However, many FUN inclusions contain unfractionated components issued from the first crystallization stages before the melt was completely evaporated [START_REF] Thrane | Discovery of a new FUN CAI from a CV carbonaceous chondrite: Evidence for multistage thermal processing in the protoplanetary disk[END_REF][START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF], which may well be the case for perovskite.

Recent studies indicate that FUN CAIs are related to normal CAIs [START_REF] Park | Calcium and titanium isotopic compositions of FUN CAIs: Implications for their origin[END_REF][START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF]. On the one hand, Ca, Ti and 26 Al isotopic anomalies define a continuum between FUN inclusions and normal CAIs [START_REF] Park | Calcium and titanium isotopic compositions of FUN CAIs: Implications for their origin[END_REF][START_REF] Park | Calcium-aluminum-rich inclusions with fractionation and unidentified nuclear effects (FUN CAIs): II. Heterogeneities of magnesium isotopes and 26 Al in the early Solar System inferred from in situ high-precision magnesium-isotope measurements[END_REF]. On the other hand, the coupled petrographic and O isotope study of twelve FUN inclusions shows that they are petrographically similar to normal CAIs and exhibit O isotope mass fractionation effects of variable amplitude not correlated with the isotopic anomalies [START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF]. This was attributed to variable conditions of partial melting and evaporation and variable sampling of reservoirs with diverse presolar components [START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF][START_REF] Park | Calcium-aluminum-rich inclusions with fractionation and unidentified nuclear effects (FUN CAIs): II. Heterogeneities of magnesium isotopes and 26 Al in the early Solar System inferred from in situ high-precision magnesium-isotope measurements[END_REF]. If this is correct, and in light of the multiple aggregation history of E101.1 revealed by the present study, which probably includes assimilation of a FUN component, it is reasonable to consider that most igneous CAIs may have been formed by the aggregation and partial melting of various precursors, among which were a proportion of FUN material with inherited presolar isotopic anomalies. (ii) the diffusivity of O in solid perovskite is rapid [START_REF] Gautason | Oxygen diffusion in perovskite: Implications for electrical conductivity in the lower mantle[END_REF][START_REF] Sakaguchi | Oxygen tracer diffusion in single-crystal CaTiO 3[END_REF], a plausible interpretation is that all perovskite grains were initially 16 O-rich and experienced partial re-equilibration with the enclosing Sc-Zr-pyroxene, either sub-solidus or between solid perovskite and melt. All grains analyzed here have comparable sizes in the 20-30 µm size range and all analyses include most of the grains, so that it has not been possible to evaluate the extent of O isotopic zoning in perovskite, thus rendering the latter hypothesis impossible to confirm or negate.

Insights on

The anhedral shapes of the 16 O-rich perovskite grains are consistent with the idea that they represent partially melted relict grains. These grains belong to the population with closeto-chondritic ZrO 2 /Y 2 O 3 ratios (Per B) and to the population with ZrO 2 content near or below the detection limit of the electron probe (Fig. 9). Whereas the chondritic population can be understood as resulting from direct condensation from a solar gas, the origin of the Zr-free grains remains unclear. Because Zr-oxides have already been observed in UR CAIs [START_REF] Ivanova | Compound ultrarefractory CAI-bearing inclusions from CV3 carbonaceous chondrites[END_REF] and because Zr is an ultrarefractory element that condenses at equilibrium as ZrO 2 before corundum and hibonite condense (i.e., at higher temperature, [START_REF] Lodders | Solar system abundances and condensation temperatures of the elements[END_REF], a possibility is that the Zr-free perovskite is an UR condensate that formed after a fractionation due to previous condensation of ZrO 2 . As mentioned above such a fractional condensation is in line with the comparable abundances of Lu, Er and Ho in E101.1 instead of Lu > Er > Ho as expected from relative volatilities [START_REF] Simon | A unique ultrarefractory inclusion from the Murchison meteorite[END_REF]. Both the chondritic and the Zr-free perovskite probably represent end-member products of the condensation of UR minerals due to slight variations of the condensation conditions (temperature or pressure). We also note that some of the Zr depletion in perovskite probably arose from the late solid-state diffusion into the Sc-Zr-rich pyroxene (El Goresy et al., 2002). A reasonable interpretation is thus that the large perovskite grains were initially part of a 16 O-rich UR inclusion (or several) condensing in the hibonite + perovskite + melilite field and containing precursor UR oxides inherited from condensation steps above 1700 K (e.g., [START_REF] Simon | A unique ultrarefractory inclusion from the Murchison meteorite[END_REF], such as those described in other compound CAIs (e.g., [START_REF] Ivanova | Compound ultrarefractory CAI-bearing inclusions from CV3 carbonaceous chondrites[END_REF]. Such an UR inclusion could be the carrier of the UN isotopic anomalies suggested by 48 Ca excesses and deficits in perovskite. It may have experienced a 16 O-depletion during the melting event associated with extensive Mg, Si evaporation and isotopic fractionation that transformed the E101.1 precursor into a FUN inclusion. In this case, the 16 O-rich perovskite would represent the starting material of the FUN inclusion and melilite ± Sc-Zr-rich pyroxene would represent the final crystallization product after equilibration with the 16 O-poor gas [START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF]. The lack of strongly mass fractionated O isotope component corresponding to the heavy Si component would then be explained by the peculiar chemistry of E101.1, which lacks intermediate minerals with slow oxygen diffusivities recording the composition of the evaporating melt, such as spinel, forsterite or pyroxene. The O isotopic compositions of initial perovskite and final melilite + Sc-Zr-rich pyroxene compare well with those of initial spinel and final melilite, respectively, in some FUN inclusions having Si isotopic compositions similar to that of the E101.1 parental melt (e.g., CG14, EK1-4-1, [START_REF] Krot | Calcium-aluminum-rich inclusions with fractionation and unknown nuclear effects (FUN CAIs): I. Mineralogy, petrology and oxygen isotopic compositions[END_REF]. This scenario would also agree with the residual O isotope mass fractionation observed in Sc-Zr-rich pyroxene and some of the 16 O-poor melilite (Fig. 7). Alternatively, the 16 O-rich UR perovskite-rich precursor could have coagulated with a F(UN)-like proto-CAI previously mass fractionated and both were partially melted together.

Extraneous origin of the spinel clusters

Beside evidence of perovskite partial melting, there are numerous examples of spinel resorption (e.g., rounded shapes, embayments) in E101.1. These features, as well as the presence of the pyroxene-rich symplectite around spinels, and the chemical and isotopic zonation of melilite (enriched in gehlenite and often in 16 O at the contact with spinel; Fig. 6) are indicative of a coronitic reaction resulting in spinel dissolution. Interestingly, a spinel + liquid = anorthite + melilite + clinopyroxene reaction has been observed in the CMAS system [START_REF] O'hara | Diopside + spinel equilibria, anorthite and forsterite reaction relationships in silica-poor liquids in the system CaO-MgO-Al 2 O 3 -SiO 2 at atmospheric pressure and their bearing on the genesis of melilitites and nephelinites[END_REF][START_REF] Schairer | Critical planes and flow sheet for a portion of the system CaO-MgO-Al 2 O 3 -SiO 2 having petrological applications[END_REF][START_REF] Salmon | Phase equilibria of the join akermanite-anorthite-forsterite in the system CaO-MgO-Al 2 O 3 -SiO 2 at atmospheric pressure[END_REF][START_REF] Onuma | Study of the system CaMgSi 2 O 6 -CaFe 3+ AlSiO 6 -CaAl 2 SiO 6 -CaTiAl 2 O 6 : II. The join CaMgSi 2 O 6 -CaAl 2 SiO 6 -CaTiAl 2 O 6 and its bearing on the Ca-Alrich inclusions in carbonaceous chondrite[END_REF]. Although it has not been observed directly, such a reaction has been inferred during closed-system crystallization of Type B CAI liquids, based on the proportions of spinel as a function of temperature [START_REF] Stolper | Crystallization sequences of Ca-Al-rich inclusions from Allende: An experimental study[END_REF]. A similar reaction in the highly silica-poor system of Type A CAIs may therefore result in a comparable reaction where anorthite is absent in favor of highly gehlenitic melilite and highly aluminous pyroxene (reaction 1).

(1) spinel + liquid = gehlenite + clinopyroxene

We suggest that the coronitic textures observed around spinel are the signature of this reaction. It would explain all resorption features in spinel, the presence of the pyroxene ± spinel symplectite and the coupled enrichment in 16 O and Al 2 O 3 in melilite at the contact with spinel (Fig. 6).

Whether reaction (1) occurred in closed or open-systems cannot be deduced from the coronitic textures alone. Additional insights come from Mg isotopes (Fig. 11,12). Most spinel clusters have δ 25 Mg values near 5-6‰, close to each other and within error of the associated melilite (Fig. 12). However some heterogeneities point to incomplete homogenization.

Although the exact spinel-melilite difference can possibly be subject to small uncorrected matrix effects, the range of δ 25 Mg values in spinel from 2.6‰ (cluster 5) to 6.6‰ (cluster 4)

with an intermediate 4.3‰ (cluster 3) is beyond 2σ analytical uncertainties (< 1‰). The Al/Mg systematics of spinel clusters and associated melilite are mostly comparable from cluster to cluster (Table 2, Fig. 11) but heterogeneities exist as well. Intercepts in the Al/Mg regressions from the various clusters (Fig. 11) yield a range from 0.11 ± 0.28 ‰ to 0.87 ± 0.22 ‰, again beyond 2σ errors. Finally, spinel clusters from subinclusion 1 have Mg isotope systematics similar to the others in terms of radiogenic Mg and spinel-melilite stable isotope equilibration, albeit for a much different δ 25 Mg value near 0‰ (Fig. 12), which has been shown to result from mixing between the host inclusion and the sinuous pyroxene fragments (Fig. 15).

In a first interpretation, most melilite and spinel have identical Mg isotopic composition and Al/Mg systematics. In this hypothesis, supracanonical spinel in cluster 2 (Table 2, Fig. 11) and light melilite near cluster 5 (analysis Mg58) must be considered special as well as subinclusion 1, possibly another xenolith of its own in this case. Our favorite interpretation is however that these heterogeneities result from incomplete homogenization between spinel clusters and the host melilite/melt, including in subinclusion 1, where the analyzed spinel seem to belong both to the spinel swarms and to subinclusion 1 (Fig. 1,2). In this hypothesis, the supracanonical spinel in cluster 2 results from diffusive exchange with surrounding melilite (see section 4.6) and the light δ 25 Mg value at 3.7‰ near cluster 5 (Mg58) may indicate a contribution of light Mg from spinel, in agreement with O isotopes indicating a contribution of spinel-derived 16 O-rich oxygen (Fig. 6).

We therefore suggest that all spinel clusters initially had a stable Mg isotopic composition different from that of the host E101.1 CAI, and thus a different origin. They subsequently underwent Mg isotope exchange with the surrounding melilite or parental melt owing to rapid Mg self-diffusion in spinel [START_REF] Sheng | Self-diffusion of magnesium in spinel and in equilibrium melts: constraints on flash heating of silicates[END_REF] up to the point of complete or near-complete equilibration (clusters 1, 2 and 4). At 1300 to 1400°C, typical of melilite crystallization, Mg isotope exchange in a 10 µm spinel is complete in 2 to 10 hours. This isotopic disequilibrium is in line with chemical evidence for several spinel populations with varying V, Ti and Cr content (El Goresy et al., 2002). For comparable Cr content, host spinels (i.e., spinel clusters) and some spinels in subinclusion 1 have the highest V content. Because the diffusion of V in spinel is rapid compared to that of Cr [START_REF] Connolly | On type B CAI formation: Experimental constraints on fO 2 variations in spinel minor element partitioning and reequilibration effects[END_REF], this indicates a significant degree of equilibration with the host melt or melilite for these spinels.

Taken together, the coronitic textures, the distribution of gehlenitic melilite and the O and Mg isotope systematics suggest introduction of the parental material of the spinel clusters by capture of solid material into a partially molten CAI host droplet, confirming the suggestion of El Goresy et al. (2002) based on the minor element chemistry of spinel. Spinel clusters can thus be considered to represent 16 O-rich, spinel-rich proto-CAIs captured by the host inclusion and partially dissolved into the host melt. It is interesting to note that cluster 3, the least equilibrated of the upper 4 clusters (Fig. 1) is also the outermost. Clusters 1 to 4 may have been fragments of the same initial proto-CAI captured in a single event and increasingly equilibrated with melilite as they are deep within the channel.

Such a 16 O-rich spinel-rich proto-CAI may have been a fine-grained-spinel-rich CAI (FG-sp-rich CAI, e.g., Krot et al., 2004a) in which minerals other than spinel, i.e., having lower melting temperature (e.g., pyroxene ± melilite ± anorthite), would have been entirely dissolved in the host melt, possibly contributing to an additional local enrichment in Al and 16 O. In this case, the subhedral crystal shapes observed in many grains of the spinel clusters would have to result from recrystallization or annealing of fine-grained aggregates of spinel.

Unaltered FG-sp-rich CAIs are homogeneously rich in 16 O (e.g., [START_REF] Aléon | Calcium-aluminum-rich inclusions and amoeboid olivine aggregates from the CR carbonaceous chondrites[END_REF][START_REF] Aléon | Finegrained, spinel-rich inclusions from the reduced CV chondrite Efremovka: II. Oxygen isotopic compositions[END_REF][START_REF] Fagan | Oxygen isotopic alteration in Ca-Al-rich inclusions from Efremovka: Nebular or parent body setting? Meteorit[END_REF][START_REF] Bodénan | An oxygen isotope study of Wark-Lovering rims on Ttype A CAIs in primitive carbonaceous chondrites[END_REF][START_REF] Ushikubo | A long duration of the 16 O-rich reservoir in the solar nebula, as recorded in fine-grained refractory inclusions from the least metamorphosed carbonaceous chondrites[END_REF] and are commonly devoid of heavy Mg because they escaped extensive melting and evaporation. Their δ 25 Mg values are usually near 0‰ or slightly negative (e.g., [START_REF] Clayton | Isotopic variations in the rock-forming elements in meteorites[END_REF][START_REF] Macpherson | Early solar nebula condensates with canonical, not supracanonical, initial 26 Al/ 27 Al ratios[END_REF][START_REF] Ushikubo | A long duration of the 16 O-rich reservoir in the solar nebula, as recorded in fine-grained refractory inclusions from the least metamorphosed carbonaceous chondrites[END_REF]. As a result, capture of such spinel-rich material is consistent with the Mg-Si isotopic decoupling inferred in the parental CMAS melt of the host inclusion (section 4.2), as well as the lower δ 25 Mg value of cluster 5. We calculated that the contribution of the spinel-rich material should be of a few wt% to account for the Mg-Si isotopic decoupling, which agrees with the abundance of spinel as visible on the E101.1 polished section.

Removing the spinel clusters from the host CAI mineralogy leaves an initial precursor with melilite + perovskite mineralogy, essentially devoid of spinel, which agrees with the calculated precursor chemistry falling in the hibonite + perovskite + melilite domain if hibonite has been completely resorbed in favor of a more gehlenitic melilite. We conclude that the capture of extraneous fine-grained spinel-rich proto-CAIs well accounts for the properties of the spinel clusters as well as the isotopic systematics of the host parental melt.

The capture of 16 O-rich spinel-rich CAIs by a 16 O-poor melilite dominated host inclusion is analogous to the situation in E49, another CTA compound inclusion from Efremovka [START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF]. In other CAIs, spinel and melilite have been shown to occasionally be in Mg isotopic equilibrium (e.g., [START_REF] Fahey | Microdistributions of Mg isotopes and REE abundances in a Type A calcium-aluminum-rich inclusion from Efremovka[END_REF][START_REF] Kita | Internal 26 Al-26 Mg isotope systematics of a Type B CAI: Remelting of refractory precursor solids[END_REF] and in some cases show evidence of Mg isotope disequilibrium (e.g., [START_REF] Paque | XANES and Mg isotopic analyses of spinels in Ca-Al-rich inclusions: Evidence for formation under oxidizing conditions[END_REF][START_REF] Goswami | Ion microprobe studies of Efremovka CAIs: I. Magnesium isotope composition[END_REF], with spinel being either heavier or lighter than coexisting melilite by at most 1 or 2 ‰. With the caveat that matrix effects are important in Mg isotopic analysis by SIMS and need to be carefully taken into account to investigate the spinel-melilite isotopic difference at the sub-permil level (e.g., [START_REF] Kita | Internal 26 Al-26 Mg isotope systematics of a Type B CAI: Remelting of refractory precursor solids[END_REF][START_REF] Paque | XANES and Mg isotopic analyses of spinels in Ca-Al-rich inclusions: Evidence for formation under oxidizing conditions[END_REF], this suggests that spinel of extraneous origin that underwent subsequent partial to complete isotopic re-equilibration may be common in CAIs.

Formation of the margins and rims of the host CAI

Forsterite-rich accretionary rim

The multi-isotope study of E101.1 can also be used to shed light on the origin of the outermost envelopes of igneous CAIs, the WLR and the forsterite-rich accretionary rims (AR, e.g., [START_REF] Krot | Forsterite-rich accretionary rims around Ca,Al-rich inclusions from the reduced CV3 chondrite Efremovka[END_REF]. ARs are thought to be somewhat analogous to AOAs. They are composed primarily of highly magnesian olivine aggregates, commonly anhedral and porous.

Some are more compact with triple junctions suggesting secondary annealing [START_REF] Krot | Forsterite-rich accretionary rims around Ca,Al-rich inclusions from the reduced CV3 chondrite Efremovka[END_REF](Krot et al., , 2004b)). Numerous CAI-like nodules mostly composed of Ca-clinopyroxene ± anorthite ± spinel are often interspersed with forsterite (Krot et al., 2004b). The commonly held view is that AR are aggregates of fine-grained nebular condensates of relatively low temperature, in the range of forsterite condensation [START_REF] Krot | Forsterite-rich accretionary rims around Ca,Al-rich inclusions from the reduced CV3 chondrite Efremovka[END_REF], subsequently aggregated onto larger

CAIs and eventually partially reprocessed on the parent-body [START_REF] Krot | Forsterite-rich accretionary rims around Ca,Al-rich inclusions from the reduced CV3 chondrite Efremovka[END_REF]. The Mg and Si isotopic compositions of forsterite from the AR in E101.1 correspond to δ 25 Mg and FSi close to 0 ‰ or slightly negative, which is in good agreement with the expectation for equilibrium condensation (Fig. 17) and with previously published Mg isotopic analyses of AR and AOAs [START_REF] Larsen | Evidence for magnesium isotope heterogeneity in the solar protoplanetary disk[END_REF][START_REF] Macpherson | Wellresolved variations in the formation ages for Ca-Al-rich inclusions in the early solar system[END_REF][START_REF] Mishra | Timing and extent of Mg and Al isotopic homogenization in the early solar system[END_REF]. The 16 O-rich composition of forsterite indicates that the gas from which they condensed was 16 Orich as inferred from other AR and AOAs [START_REF] Krot | Existence of an 16 Orich gaseous reservoir in the solar nebula[END_REF].

Wark-Lovering rim

In contrast to the ARs, the origin of WLR is more debated. They typically consist of close to monomineralic layers mimicking the condensation sequence from spinel ± hibonite ± perovksite at the contact with the host inclusions to Al-Ti-pyroxene ± melilite ± anorthite, to Al-diopside (e.g., MacPherson, 2014) toward the edge of the CAI. In some cases, the outermost layers consist of pure diopside and/or forsterite. Different origins have been proposed, which can be grouped into two categories: condensation models and flash heating models. Condensation models include ( 1) successive deposition onto a CAI of various layers of condensation [START_REF] Wark | Marker events in the early evolution of the solar system: evidence from rims on Ca-Al-rich inclusions in carbonaceous chondrites[END_REF] with decreasing temperature from the innermost spinel ± hibonite ± perovskite to the outermost forsterite layers, (2) growth of isolated nucleus islands deposited onto the interior CAI [START_REF] Bolser | Microstructural analysis of Wark-Lovering in the Allende and Axtell CV3 chondrites: Implications for high-temperature nebular processes[END_REF] and (3) diffusive equilibration of the edge of the CAI with the gas [START_REF] Ruzicka | Mineral layers around coarse-grained, Ca-Al-rich inclusions in CV3 carbonaceous chondrites: Formation by high-temperature metasomatism[END_REF][START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF][START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF].

Light stable Mg isotopic composition supports these condensation models [START_REF] Taylor | 26 Al in Efremovka CAI E44L -Resolved time interval between interior and rim formation in a highly fractionated compact type A CAI[END_REF]Cosarinsky et al. 2005a,b;[START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF] but enhanced REE abundances in the rims and REE patterns identical to the host inclusion favor flash heating hypotheses (e.g. [START_REF] Wark | The formation of rims on calcium-aluminum-rich inclusions: Step I -Flash heating[END_REF]. These conflicting interpretations could potentially be reconciled if WLR formed by condensation of nebular gas and a melted outer layer is involved at some stage [START_REF] Toppani | Wark-Lovering rims at the nanometer scale: a transmission electron microscopy study[END_REF][START_REF] Keller | A FIB/TEM study of a complex Wark-Lovering rim on a Vigarano CAI[END_REF][START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF]. Mg isotopes and Ti valence have been used to argue that the WLR formed at a much higher fO 2 , closer to that of chondrites, than the inner host formed at a solar fO 2 [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF][START_REF] Dyl | Valence state of titanium in the Wark-Lovering rim of a Leoville CAI as a record of progressive oxidation in the early Solar Nebula[END_REF] although this is debated [START_REF] Simon | Valence of titanium and vanadium in pyroxene in refractory inclusion interiors and rims[END_REF]. Finally, Mg isotopes also suggest that WLR have crystallization ages commonly younger than their host inclusions by about 200,000 to 300,000 years [START_REF] Taylor | 26 Al in Efremovka CAI E44L -Resolved time interval between interior and rim formation in a highly fractionated compact type A CAI[END_REF]Cosarinsky et al., 2005a,b;[START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF]. Even larger time intervals of at least 590,000 and 690,000 years have recently been reported [START_REF] Mane | Resolved time difference between Calcium-Aluminum-rich inclusions and their Wark-Lovering rims inferred from Al-Mg chronology of two inclusions from a CV3 carbonaceous chondrite[END_REF] In Si. Such a scenario has already been proposed from the TEM study of a WLR from a Type B1 CAI from Vigarano [START_REF] Toppani | Wark-Lovering rims at the nanometer scale: a transmission electron microscopy study[END_REF]. The latter study also documented secondary alteration phases at the sub-µm scale. Although the NanoSIMS study of WLR from CV chondrites shows variations of 16 O content (Simon et al., 2011[START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF], the study of WLR in unaltered chondrites favors a parent-body origin for these variations [START_REF] Bodénan | An oxygen isotope study of Wark-Lovering rims on Ttype A CAIs in primitive carbonaceous chondrites[END_REF][START_REF] Krot | High-temperature rims around calcium-aluminum-rich inclusions from the CR, CB and CH carbonaceous chondrites[END_REF], in agreement with the TEM observations. Our data favor a systematic 16 Orich composition of primordial unaltered WLR (this work, [START_REF] Yoshitake | Correlations between oxygen-isotopic composition and petrologic setting in a coarse-grained Ca, Al-rich inclusion[END_REF][START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Bodénan | An oxygen isotope study of Wark-Lovering rims on Ttype A CAIs in primitive carbonaceous chondrites[END_REF][START_REF] Kawasaki | Chronological study of oxygen isotope composition for the solar protoplanetary disk recorded in a fluffy Type A CAI from Vigarano[END_REF][START_REF] Krot | High-temperature rims around calcium-aluminum-rich inclusions from the CR, CB and CH carbonaceous chondrites[END_REF]. Most likely, the gas with which the CAI equilibrated during WLR formation was 16 O-rich.

Zoning of the outer portion of the host CAI

This scenario of WLR formation explains the isotopic composition of the rim, and also the O, Mg and Si isotopic zoning observed in the outer part of CAIs (e.g., [START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF] for O). In E101.1, chemical and isotopic profiles indicate that the distribution of Mg in the outer part of the CAI (Fig. 13) was affected by a process resulting in [START_REF]A CAI with ultrarefractory REE patterns and enormous enrichments of Sc, Zr and Y in fassaite and perovskite[END_REF] an enrichment in Al or loss in Mg toward the exterior, (2) an enrichment in unfractionated Mg, (3) an enrichment in undisturbed near-canonical 26 Al. In the framework of the WLR formation by condensation from nebular gas into a partially melted CAI layer, these profiles are well understood by a global loss of interior Mg in the periphery together with a progressive enrichment in nebular Mg. Whether the observed profiles result from solid-state diffusion, partial melting and recrystallization or a combination of both is unclear. By contrast, the Si and O isotopic compositions do not show any relationship with proximity to the WLR (Fig. 13d,e). In the case of Si isotopes, this agrees with the WLR systematics, which shows a much greater contribution of interior Si than of interior Mg in the rim (Fig. 17). Interestingly, the width of the zone affected by this interaction varies between profile 1 (30-50 µm) and profile 2 (70-100 µm) (Fig. 13a-c) suggesting small variations in the thermal regime or in the CAI response to the WLR-forming event. Parameters such as temperature, chemical gradients due to differences in total and partial pressures in the gas and to differences in CAI chemistry probably induce even larger variations of thickness of this zone between different CAIs with different thermal histories. An extreme case of intra-CAI variations may be bowl-shaped

CAIs with different WLR mineralogy in their concave and convex faces suggesting different degrees of heating attributed to a lack of spinning in the nebula [START_REF] Ivanova | Plastically-deformed igneous Calcium-Aluminum-rich Inclusions from CV carbonaceous chondrites: Clues to a nature of CAI melting events[END_REF]. Inter-CAI variations of thermal regime during WLR formation probably also explain why Mg and O gradients are not always present simultaneously. Many Type A CAIs show for instance an enrichment in gehlenite toward the exterior, in some cases associated with a Mg isotopic composition evolving toward lower δ 25 Mg values [START_REF] Fahey | Microdistributions of Mg isotopes and REE abundances in a Type A calcium-aluminum-rich inclusion from Efremovka[END_REF][START_REF] Goswami | Ion microprobe studies of Efremovka CAIs: I. Magnesium isotope composition[END_REF][START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF][START_REF] Cosarinsky | Aluminum-26 model ages of hibonite and spinel from type A inclusions in CV chondrites[END_REF], in some cases associated with an enrichment in 16 O [START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Katayama | Oxygen isotopic zoning of reversely zoned melilite crystals in a fluffy type A Ca-Al-rich inclusions from the Vigarano meteorite[END_REF][START_REF] Kawasaki | Oxygen isotopic and chemical zoning of melilite crystals in a type A Ca-Al-rich inclusion of Efremovka CV3 chondrite[END_REF][START_REF] Kawasaki | Chronological study of oxygen isotope composition for the solar protoplanetary disk recorded in a fluffy Type A CAI from Vigarano[END_REF]. Similarly to E101.1, a decrease in δ 25 Mg values together with little variation in Si isotopic composition has been observed in the outer mantles of Type B1 CAIs [START_REF] Bullock | Mg and Si isotopic fractionation patterns in types B1 and B2 CAIs: Implications for formation under different nebular conditions[END_REF]. [START_REF] Bullock | Mg and Si isotopic fractionation patterns in types B1 and B2 CAIs: Implications for formation under different nebular conditions[END_REF] conclude that these profiles would be best explained by a secondary "ad-hoc" process.

We suggest that this process is in fact WLR formation. In the cases where the interior CAI is almost unaffected by WLR formation (i.e., where the transition region has a thickness near 0), the CAIs may have preserved pre-WLR zoning reflecting high temperature interaction of a solid CAI with nebular gas such as sublimation of Mg and Si from the solid state (e.g., CAI

Leoville 144A, [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF][START_REF] Shahar | Astrophysics of CAI formation as revealed by silicon isotope LA-MC-ICPMS of an igneous CAI[END_REF] or diffusion of 16 O-poor oxygen from the gas (e.g., CAI Allende A37, Simon et al., 2011[START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF]. Given the decoupling between O and Mg in E101.1, the two outermost 16 O-poor analyses in profile 1 may reflect the preservation of such an initial 16 O-poor composition of the melt due to equilibration with 16 O-poor gas before trapping of the xenoliths and WLR formation.

Al/Mg chronology and disturbances

The lack of a well-defined Al/Mg isochron in E101.1 host melilite, the evidence of isotopic disequilibrium between spinels and host melilite, and the insufficient precision on the slopes and intercept of isochrons deduced from analyses of pyroxene preclude a simple chronological analysis of E101.1 crystallization. Still, the distribution of model ( Al !" / Al)

!" ! ′ ratios with location in the CAI or between subcomponents can shed light on the origin of Al/Mg systematics perturbations and on the putative initial 26 Al/ 27 Al ratio prior to these perturbations (Table 2).

As mentioned in section 4.4, the lack of complete stable Mg isotope equilibrium between spinel clusters and adjacent melilite is visible to some extent with the radiogenic isotopes. The Al/Mg systematics in paired spinel and melilite are illustrated in Fig. 11. Three cases can be distinguished.

In cluster 2 (Fig. 11a), sub-canonical mellite is associated with supercanonical spinel.

All 4 analyses define a single linear relationship, which yields an initial 26 Al/ 27 Al slope of (2.9±0.3) × 10 -5 and a distinctly positive initial δ 26 Mg* intercept of 0.87 ± 0.22 ‰ (2σ errors).

This behavior is characteristic of a closed-system redistribution of Mg isotopes between spinel and associated melilite, pivoting around the bulk composition which would be coincident with the canonical isochron if the system initially had 26 Al/ 27 Al = 5.2 × 10 -5 .

In subinclusion 1 (Fig. 11b), spinel has a canonical abundance of 26 Mg*, but the two melilite analyses in the vicinity of spinel define a linear array with the spinel with a subcanonical slope of (3.5±0.4) × 10 -5 and a positive intercept of 0.45 ± 0.16 ‰. Such slope and intercepts are also indicative of closed-system exchange in the spinel-melilite system, although spinel seems to have been little affected by this perturbation. One possible explanation is an initially sub-canonical bulk composition. This would indicate a younger crystallization age for subinclusion 1, as young as 400 ± 100 thousand years after canonical.

However, when considering all melilite analyses in subinclusion 1, the canonical ( Al !" / Al)

!" !

′ ratio of analysis Mg50 is not in a favor of such a chronological interpretation (Fig. 10). Rather it may indicate the variable contribution of a 26 Al-poor component (see below).

In contrast, spinel and the associated melilite of cluster 3 (Fig. 11c) define a linear array with a slope of (3.3±0.4) × 10 -5 that passes through the origin (intercept = 0.11 ± 0.28 ‰). If this 2-point line is considered to be an isochron, it would yield a crystallization age of 480 ± 120 thousand years after canonical. Such an age may represent the age of cluster 3 trapping in the host CAI, however it would not be consistent with the closed-system interactions in other spinel clusters -melilite pairs implying that either the trapping of cluster 3 is a late isolated event or that the linear correlation does not correspond to the crystallization age. Another possible interpretation for this cluster closest to the edge of the CAI is an opensystem re-equilibration of both spinel and associated melilite with a reservoir of lower 26 Al content (e.g., [START_REF] Simon | Resetting, errorchrons and the meaning of canonical CAI initial 26 Al/ 27 Al values[END_REF]. This reservoir is unlikely to be the meteorite matrix as near canonical ( Al !" / Al)

!" ! ′ ratios are observed in melilite in close contact with the WLR in profile 1 and in areas 3 and 4 (Fig. 13c). Thus, it has to be exchange with a nebular reservoir of low 26 Al content.

The distribution of model ( Al !" / Al)

!" ! ′ ratios in melilite as function of distance from the rim (Fig. 13c) is helpful in interpreting the overall Al/Mg isotope systematics of E101.1.

As mentioned above, profile 2 and areas 3 and 4 display near canonical ( Al !" / Al)

!" ! ! ratios in the vicinity of the rim. Apart from one analysis with a low ( Al !" / Al)

!" !

! ratio in profile 1, a progressive decrease in ( Al !" / Al)

!" ! ! ratios toward the interior is visible everywhere in the outermost 80 µm of the CAI, down to about 3 × 10 -5 , inward of which scattered ratios are found. Interestingly, the isotopic composition of cluster 3 is consistent with this lower limit, cluster 3 being located ~ 100 µm inside the CAI. As the outer part of the CAI and the WLR cannot be older than the interior CAI, this suggests a secondary perturbation with the WLR and the outermost melilite having best preserved the initial near-canonical composition.

A possible interpretation based on the relationship between the parental E101.1 melt and FUN inclusions discussed above is that E101.1 crystallized from different proto-CAIs with different initial 26 Al contents. Indeed, many FUN inclusions are known to be devoid of or depleted in 26 Al (e.g., [START_REF] Esat | Search for 26 Al effects in the Allende FUN Inclusion C1[END_REF]Clayton et al., 1984;[START_REF] Davis | Melt solidification and late-stage evaporation in the evolution of a FUN inclusion from the Vigarano C3V chondrite[END_REF][START_REF] Holst | 182 Hf-182 W age dating of a 26 Al-poor inclusion and implications for the origin of short-lived radioisotopes in the early Solar System[END_REF][START_REF] Park | Calcium-aluminum-rich inclusions with fractionation and unidentified nuclear effects (FUN CAIs): II. Heterogeneities of magnesium isotopes and 26 Al in the early Solar System inferred from in situ high-precision magnesium-isotope measurements[END_REF] and mixing of a FUN precursor of low 26 Al content with xenoliths having canonical or near canonical 26 Al content will induce strong confusion of the Al/Mg isotope systematics in the final E101.1 CAI. These perturbations are likely to be more significant in the interior of the CAI than on the exterior, where re-equilibration with the regular canonical CAI reservoir, from which xenoliths are issued, will be favored. Such a 26 Al-poor FUN component may be the nebular reservoir required to explain the isotopic composition of cluster 3.

To summarize, the Al/Mg isotope systematics of E101.1 can be understood as resulting from the mixing of parental FUN material depleted in 26 Al with xenoliths having abundant (the sinuous pyroxene) up to canonical (the spinel clusters) 26 Al content. In the hot CAI interior, local closed-system interactions between spinel/pyroxene and melilite resulted in scattered ( Al !" / Al)

!" ! ! ratios. Closer to the outside of the CAI, open-system resetting of Mg (e.g. [START_REF] Simon | Resetting, errorchrons and the meaning of canonical CAI initial 26 Al/ 27 Al values[END_REF] due to mixing of canonical and 26 Al-poor material led to intermediate Al/Mg apparent isochrons resulting from the subsequent decay of 26 Al (cluster 3). Finally, the outermost portion of the CAI equilibrated with the ambient gas having canonical or near-canonical 26 Al content. This is consistent with the initial δ 26 Mg* of forsterite in the AR being comparable to that determined from AOAs and AR of igneous CAIs with canonical 26 Al/ 27 Al ratios [START_REF] Larsen | Evidence for magnesium isotope heterogeneity in the solar protoplanetary disk[END_REF][START_REF] Macpherson | Wellresolved variations in the formation ages for Ca-Al-rich inclusions in the early solar system[END_REF][START_REF] Mishra | Timing and extent of Mg and Al isotopic homogenization in the early solar system[END_REF].

A chronological constraint on the formation of E101.1 could in principle be obtained from its WLR formation, the last melting event suffered by the CAI. However, time intervals as small as 2-3 × 10 5 years between the formation of the interior CAI and that of the WLR [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF]Taylor et al., 2005a,b;[START_REF] Cosarinsky | Aluminum-26 model ages of hibonite and spinel from type A inclusions in CV chondrites[END_REF]) must be treated with caution. A problem is that the required high precision Al/Mg dating of individual minerals in WLR is extremely difficult to achieve due to the small size of the crystals. Large errors and uncertainties on the Mg isotopic closure for instance due to Mg self-diffusion in spinel make it difficult to obtain good chronological constraints. Here we show that the range of model initial 26 Al/ 27 Al in WLR spinel from E101.1 varies between (3.7±0.5) × 10 -5 and (5.7±0.4) × 10 -5 , i.e., corresponding to an apparent age interval of 460,000 years between the first and last spinel crystallized, although it is within error consistent with a 240,000 years interval between the last and the canonical spinels (Fig. 18). In a first approximation, this estimate is consistent with most previous results [START_REF] Simon | A short timescale for changing oxygen fugacity in the solar nebula revealed by highresolution 26 Al-26 Mg dating of CAI rims[END_REF]Taylor et al., 2005a,b;[START_REF] Cosarinsky | Aluminum-26 model ages of hibonite and spinel from type A inclusions in CV chondrites[END_REF], except those larger than 590,000 years measured in anorthite [START_REF] Mane | Resolved time difference between Calcium-Aluminum-rich inclusions and their Wark-Lovering rims inferred from Al-Mg chronology of two inclusions from a CV3 carbonaceous chondrite[END_REF] but anorthite is a mineral highly susceptible to Mg isotope re-equilibration and may not provide an accurate age of WLR formation. However, if the range in E101.1 spinel were to correspond to Mg exchange with a reservoir with different 26 Al content, the age conclusion would also be erroneous. For instance, a local and variable contribution of interior FUN material devoid of 26 Al to the WLR could potentially explain the 26 Al variations in WLR spinel. Such an explanation would be in line with the contribution of isotopically heavy Si (and occasionally Mg) from the interior (Fig. 13) and with low ( Al !" / Al)

!" ! ! ratios observed in the outer part of profile 1 (Fig. 13c).

We also considered what could be the extent of secondary Al-Mg disturbances in the WLR due to solid state exchange with a reservoir devoid of 26 Al. The range of model initial 26 Al/ 27 Al ratios corresponds to a ~20 % exchange for a 1µm spinel crystal during a 6 months heating at 1000 K (Fig. 18) using the self-diffusion coefficient of Mg in spinel from [START_REF] Sheng | Self-diffusion of magnesium in spinel and in equilibrium melts: constraints on flash heating of silicates[END_REF]. This exchange duration becomes 2.5 My at 750 K and 30 My at 720 K. The composition of the co-existing pyroxene can be further calculated in similar conditions using the self-diffusion coefficient of natural diopside averaged over the crystallographic axes as a first approximation [START_REF] Zhang | Ca-Mg diffusion in diopside: tracer and chemical inter-diffusion coefficients[END_REF] and corresponds to a 11 % exchange for 1µm pyroxene crystals at 1000 K (Fig. 18). At lower temperature and for longer durations, the amount of exchange in the pyroxene becomes closer to that in spinel (e.g., 18% at 720 K).

Within analytical error, the pyroxene data cannot be distinguished from a canonical isochron, nor from a younger value, nor from an exchanged value (Fig. 18). This calculation shows that late reheating in the solar nebula (e.g., for the highest temperatures) or parent-body metamorphism (e.g., for the lowest temperatures) are possible sources of chronological disturbance and younger apparent WLR ages because spinel may not have preserved the initial ratio. The range of composition in E101.1 spinel is a possible hint for this process, which may has affected all WLR spinels in other CAIs. We note that 750 K may still be too high for the peak metamorphic temperature undergone by Efremovka for which the maximal temperature is probably closer to ~ 600 K by analogy with Leoville and Vigarano of petrologic type and thermal history close to those of Efremovka [START_REF] Huss | Noble gases in presolar diamonds II: Component abundances reflect thermal processing[END_REF][START_REF] Bonal | Determination of the petrologic type of CV3 chondrites by Raman spectroscopy of included organic matter[END_REF], which would require unreasonably long duration of metamorphism for Mg isotope exchange in spinel. However, interpretation of WLR ages in more metamorphosed meteorites such as Allende must be treated with caution.

Given the uncertainties on the variations of 26 Al content in WLR spinel, the initial 26 Al/ 27 Al ratio and age of the WLR are probably best given by the converging values of profile 2 and area 4 where the initial value seems best preserved at the contact with the WLR.

This initial ratio of (5.2±0.1) × 10 -5 (2σ) agrees well with that obtained by a regression through WLR data if the three spinel analyses with evidence of disturbances are removed, that is (5.2±0.8) × 10 -5 (2σ) (Fig. 18). We conclude that the WLR of E101.1 most likely formed with a canonical initial 26 Al/ 27 Al ratio. If the canonical ratio corresponds to the time of 26 Al homogenization in the early solar system (e.g., [START_REF] Sahijpal | Ti, and Ca isotopic compositions and refractory trace element abundances in hibonites from CM and CV meteorites: Implications for early solar system processes[END_REF][START_REF] Mishra | Timing and extent of Mg and Al isotopic homogenization in the early solar system[END_REF][START_REF] Young | Supracanonical 26 Al/ 27 Al and the residence time of CAIs in the solar protoplanetary disk[END_REF] and the beginning of its chronological use, this constrains the epoch of E101.1 closure to capture of xenolithic materials to within a ~100,000 years interval. More generally, the measured ages of formation of WLR and their age difference with the interior may not be representative of their true age of formation and many WLR may be essentially contemporaneous with their host CAI formation within the time resolution of the 26 Al chronometer. In light of evidence for an isotopic continuum between regular CAIs and FUN CAIs indicated by Ca and Ti isotopes [START_REF] Park | Calcium and titanium isotopic compositions of FUN CAIs: Implications for their origin[END_REF], another important implication of the present work is that mixing with FUN-like precursor material during CAI formation followed by partial melting may be an important source of isotopic disturbance of the Al/Mg systematics that was previously unrecognized. In this case, many Mg isotopic disturbances, notably in melilite, that were attributed to late secondary processes such as alteration or metamorphism may in fact be of primary nebular origin.

O isotope reservoirs during E101.1 formation

reservoirs previously evidenced in CAIs [START_REF] Krot | Existence of an 16 Orich gaseous reservoir in the solar nebula[END_REF] co-existed in the earliest solar system. The formation of the WLR further constrains the end of E101.1 formation to within the first 100,000 years. At this late stage E101.1 had already experienced a 16 O-depletion, assimilation of 16 O-rich materials and had returned to the 16 O-rich environment, where the AR formed as in many CAIs (e.g., [START_REF] Krot | Existence of an 16 Orich gaseous reservoir in the solar nebula[END_REF][START_REF] Yoshitake | Correlations between oxygen-isotopic composition and petrologic setting in a coarse-grained Ca, Al-rich inclusion[END_REF] and where pyroxene from Type B CAIs crystallized [START_REF] Aléon | Oxygen isotopes in the early protoplanetary disk inferred from pyroxene in a classical type B CAI[END_REF]. As shown for several CAIs and expected in many CAIs due to similar general O isotope systematics, the 16 O-depletion was most likely associated with partial melting/crystallization of melilite [START_REF] Yurimoto | Oxygen isotope exchange between refractory inclusion in Allende and solar nebula gas[END_REF][START_REF] Aléon | Oxygen isotopes in the early protoplanetary disk inferred from pyroxene in a classical type B CAI[END_REF][START_REF] Kawasaki | Chronological study of oxygen isotope composition for the solar protoplanetary disk recorded in a fluffy Type A CAI from Vigarano[END_REF].

Although these observations may be used to constrain the timing and location of oxygen isotope variations in the solar protoplanetary disk, they nevertheless do not shed light on the carrier phases of the oxygen isotope change between solar [START_REF] Mckeegan | The oxygen isotopic composition of the Sun inferred from captured solar wind[END_REF] and planetary, nor on the physico-chemical processes at the origin of this isotopic difference between solar and planetary reservoirs. Two other observations regarding E101.1 are more pertinent in that respect. First, the primary component related to FUN inclusions is inferred to be dominantly 16 O-poor, which is in line with the proposition that FUN inclusions formed from a primary dust component of interstellar origin depleted in 16 O relative to the solar composition (Kööp et al., 2016a;[START_REF] Krot | Oxygen isotopic composition of the Sun and mean oxygen isotopic composition of the protosolar silicate dust: evidence from refractory inclusions[END_REF]. However, it is not clear if the 16 O-poor composition of the FUN component of E101.1 is primary or was acquired during melting by isotope exchange with a 16 O-poor gas of near terrestrial composition. We note that the UR perovskite grains present in the host inclusion were probably initially 16 O-rich, in agreement with the 16 O-rich composition of the sinuous pyroxene, suggesting that most UR materials were initially 16 O-rich. The association of the UR perovskite precursor with the FUN component and the associated 16 O-depletion are the most ancient events that occurred in the history of E101.1 and their record is partially obliterated by subsequent assimilations. We sluggish and a solid state secondary re-equilibration, for instance during assimilation of the sinuous pyroxene, would not be correlated with chemistry. This suggests that the 16 O-rich environment in the disk was not entirely homogeneous and that 16 O-rich and 16 O-poor reservoirs somehow interacted with each other. In the vertical zoning model inferred by [START_REF] Aléon | Oxygen isotopes in the early protoplanetary disk inferred from pyroxene in a classical type B CAI[END_REF], this may result from propagation of the surface 16 O-depletion toward the cool midplane. In the horizontal zoning models such as those inferred from self-shielding associated with the transport of isotopically heavy water ice produced in low temperature regions (e.g., [START_REF] Yurimoto | Molecular cloud origin for the oxygen isotope heterogeneity in the Solar System[END_REF][START_REF] Lyons | CO self-shielding as the origin of oxygen isotope anomalies in the early solar nebula[END_REF][START_REF] Young | Time-dependent oxygen isotopic effects of CO self shielding across the solar protoplanetary disk[END_REF], this may correspond to the mixing front, where 16 O-poor water vapor equilibrates with the 16 Orich ambient gas. Similar small variations in 16 O-rich hibonite CAIs have been found, which may also point to variable interactions between the 16 O-rich and 16 O-poor gas associated with small temperature differences (Kööp et al., 2016a).

Protoplanetary disk context of E101.1 formation

Our observations have shown that E101.1 is a compound CAI formed from the coagulation of at least three different precursors (Fig. 19). The first precursor is the ultrarefractory CAI, rich in perovskite, that constitutes the main mass of E101.1. The sinuous pyroxene was related to some extent to this initial UR CAI. It was 16 O-rich and had near canonical abundance of 26 Al. Perovskite contained 48 Ca anomalies. Relative abundances of REE are suggestive of a complex suite of condensation events, including fractional condensation of an UR precursor. The second precursor is a F(UN) inclusion with low 26 Al content. The exact conditions of association, melting and evaporation of these two precursors remain unclear because subsequent capture and partial assimilation events obscured these early high temperature stages. After coagulation of these two precursors and their subsequent common thermal history, the main E101.1 host still had a very refractory chemical composition and (i) had experienced a 16 O-depletion, (ii) a strong evaporation, (iii) carried 48 Ca isotope anomalies, (iv) had heterogeneous 26 Al content. After this cryptic stage (Fig. 19a), the main E101.1 captured spinel-rich proto-CAIs, possibly of the common fine-grained spinel-rich CAI type and assimilated the sinuous pyroxene-rich UR CAI. In which order these two capture and partial dissolution events occurred is again not completely clear. Isotopic mixing evidence favors capture of the spinel-CAIs first and sinuous pyroxene last (Fig. 19a).

Contrary to the cryptic stage, which can only be inferred from considerations on isotopic and chemical properties of the E101.1 main host, these last two stages are recorded in the petrography of E101.1 and can be visually identified (Fig. 19b). Finally, the WLR formed by condensation into a possibly melted and evaporated outer layer of E101.1 and induced a chemical and isotopic zoning of the outer margins immediately inside the WLR.

It has recently been shown that turbulent transport within the innermost 1 astronomical unit (AU) of a thermally zoned protoplanetary disk before the end of protostellar envelope collapse yields highly variable thermal histories for the various precursors [START_REF] Taillifet | Origin of refractory inclusion diversity by turbulent transport in the inner solar nebula[END_REF]. Extraction of CAIs from the hot zone in such an environment can be realized in at most a few 1000 years. As a result, the various precursors of E101.1 may have collided sequentially with each other during such a turbulent journey through the innermost portion of the hot zone of the earliest solar protoplanetary disk. Although not recognized in regular CAIs, the growth of mm to cm sized CAIs through such a series of collisions of numerous precursors may be the standard case of CAI formation since collisions with the nearest neighbors, i.e., between precursors of similar or only marginally different thermal histories, are likely to be statistically favored. The growing identification of compound CAIs supports this conclusion (e.g., El Goresy et al., 2002;[START_REF] Aléon | Finegrained, spinel-rich inclusions from the reduced CV chondrite Efremovka: II. Oxygen isotopic compositions[END_REF][START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Macpherson | Wellresolved variations in the formation ages for Ca-Al-rich inclusions in the early solar system[END_REF][START_REF] Ivanova | A compound Ca-Alrich inclusion from CV3 chondrite Northwest Africa 3118: Implications for understanding processes during CAI formation[END_REF][START_REF] Simon | Calcium and titanium isotope fractionation in refractory inclusions: tracers of condensation and inheritance in the early solar protoplanetary disk[END_REF]. The numerical study of the collisional evolution of CAIs by coagulation and fragmentation in a turbulent hot inner disk indicates rapid growth of mm-to cm-sized objects at collisional equilibrium in a plastic (probably partially molten) state [START_REF] Charnoz | Growth of calciumaluminum-rich inclusions by coagulation and fragmentation in a turbulent protoplanetary disk: Observations and simulations[END_REF]. Such a collisional evolution agrees with aggregation of multiple precursors in E101.1 in a partially molten state and provides a natural explanation for the observed spatial decoupling between different isotope systems and between isotope compositions and mineral chemistry.

Nevertheless, the UR REE and light isotopic composition in the sinuous pyroxene fragments point to a rapid removal of its precursors from the hottest region of the disk and their rapid transport toward the exterior. It remains to be demonstrated to what extent the lack of re-equilibration with the ambient gas fits with the velocities during turbulent transport and the residence time at a given temperature [START_REF] Taillifet | Origin of refractory inclusion diversity by turbulent transport in the inner solar nebula[END_REF].

Finally, isotopic equilibration between host and xenoliths during capture and assimilation in a partially molten state is a closed-system behavior that requires isolation of the CAI from the ambient gas, notably because the self-diffusion of O in Ca-Al-rich silicate melts is rapid [START_REF] Oishi | Oxygen diffusion in liquid silicates and relation to their viscosity[END_REF]. Mixing relationships are thus expected to be blurred if the partial melt equilibrates with the gas. While this seems not difficult to achieve for Mg and Si due to vastly differing abundances in the CAI and the nebular gas having low partial pressures of Mg and SiO gas, the case of O is more problematic. Physical isolation of melt pockets inside the CAI is unrealistic in light of the distribution of xenolithic fragments and mixed areas, which reach the CAI surface. Chemical isolation is thus required. One way to prevent E101.1 partial melts from equilibrating with the ambient CO or H 2 O gas during capture and assimilation (i.e., partial melting) is melting in a diffuse near-vacuum environment with low O abundance and thus low total pressure at solar H/O ratio. [START_REF] Nagahara | The role of exchange reactions in oxygen isotope fractionation during CAI and chondrule formation[END_REF] calculated that in a closed system with O abundances in the gas about 100 times lower than in a canonical nebula, the final composition of a melt exchanging with the gas is within a few ‰ of the starting composition. However, temperatures in diffuse regions such as the inner edge of the gas disk are likely to be very high and thus to favor a strong evaporation not indicated by the Mg and Si isotopic composition during assimilation of the sinuous pyroxene. In addition, in an open system where the ambient gas can be considered infinite, the partial melt will ultimately equilibrate with the gas, unless the cooling rate is rapid enough to preserve the initial composition. In this case, however, O isotope mass fractionation may be expected depending on the relative timescales of diffusion in the melt and evaporation [START_REF] Ozawa | Chemical and isotopic fractionations by evaporation and their cosmochemical implications[END_REF]. A closed-system behavior of E101.1 during partial melting and partial assimilation of sinuous pyroxene thus remains difficult to explain and to place in an astrophysical context. For that purpose, a better knowledge of isotope exchange rates with relevant astrophysical species such as CO and H 2 O in protoplanetary disk conditions (e.g., [START_REF] Yu | Experimental study of high temperature oxygen isotope exchange during chondrule formation[END_REF][START_REF] Boesenberg | Evaporation and the absence of oxygen isotopic exchange between silicate melt and carbon monoxide gas at nebular pressures[END_REF][START_REF] Boesenberg | Oxygen isotopic exchange between carbon monoxide gas and silicate melt: Implications for the early solar nebula[END_REF]DiRocco and Pack, 2015) is required.

Summary and conclusions

The multi-isotopic study of E101.1 sheds light on the nature of its various refractory !" ! ′ ratios down to 3.7 × 10 -5 in individual spinel grains).

The identification of mixing and interactions between the various precursor CAI components in E101.1 is facilitated owing to the large isotopic differences between the lithological units that result from different thermal histories (e.g., condensation vs evaporation) or from the sampling of different cosmochemical reservoirs (e.g 16 O-rich vs 16 Opoor or 26 Al-rich vs 26 Al-poor). Some of these interactions produced isotopic heterogeneities that are also seen in other CAIs with lower amplitude, and remain poorly understood. For instance, many heterogeneities in the Al/Mg systematics are attributed to poorly documented late stage perturbations, but the E101.1 results suggest that they may also result from primary mixing processes between 26 Al-rich and 26 Al-poor components that are obscured by subsequent partial melting. Compared to the E101.1 precursors that have large isotopic differences, the closer CAI precursors are in chemistry and isotopic composition, the more difficult it is to identify these effects.

The Abbreviations: sp -spinel, sin px -sinuous pyroxene, mel -melilite, fo -forsterite, mtxmatrix. S4.

Figure captions

with δ 25

 25 Mg values between -4.7 and -2.4 ‰. In most of the inclusion, melilite has positive values between 1.9‰ and 6.5‰ with a mean value around ~ +5‰. Intermediate δ 25 Mg values are found in melilite in areas 3 and 4 and in the subinclusion 1, all in close vicinity of the sinuous pyroxene (-1.5‰ ≤ δ 25 Mg ≤ +2.1‰). Spinels in clusters 1 to 4 have positive values comparable to that of the associated melilite and close to the typical host melilite value (Fig.12, Table2). In cluster 5, spinels have lower δ 25 Mg values (~ 2.6‰). One associated melilite has a comparable low value (3.7 ‰) whereas another melilite is more typical of the host melilite (6.1‰). Both spinels and melilite in subinclusion 1 have near 0‰ δ 25 Mg values. In the WLR, the δ 25 Mg values of pyroxene range between -1.2‰ and positive values approaching those of the interior melilite at + 3‰. WLR spinel and forsterite from the AR all have δ 25 Mg values around 0 ± 1.5‰.

4. 1

 1 .2. Partial assimilation of the sinuous pyroxene. The systematic comparison of the O, Si and Mg isotopic compositions of melilite throughout the inclusion shows that the melilites from areas 3 and 4 and from the subinclusion 1 have isotopic compositions intermediate between those of the Al-diopside sinuous pyroxene fragments (∆ 17 O ~ -20‰, δ 25 Mg and FSi ~-3.5‰) and the remaining melilite of the host inclusion, which cluster around ∆ 17 O ~ -6‰, δ 25 Mg ~ +5.5‰ FSi~ +6.5‰. These intermediate isotopic compositions can be reproduced by a single mixture between the Al-diopside sinuous pyroxene and the E101.1 main host (Fig.15) with the exception of two spots in subinclusion 1 (Fig.15c). The shape of the mixing hyperbolae depends on the relative concentrations of the elements, but by considering the O, Mg and Si isotope compositions we can calculate the mixing proportions in the main E101.1 host by assuming a composition for the second end-member (Fig.14). The sinuous fragments are dominated by Al-diopside with Al 2 O 3 typically in the 2-5 wt% range. For the calculation, we thus assumed that this endmember is pure diopside (CaMgSi 2 O 6 ) having the isotopic composition of the sinuous pyroxene. If we further assume that the second end-member consists only of the five elements Ca, Al, Mg, Si and O, we find that the three mixing hyperbolae cannot be reproduced simultaneously by any known mineral having the isotopic composition of melilite. For example, the ∆ 17 O-δ 25 Mg mixing line could be matched by mixing diopside with melilite in the Åk 10 -Åk 30 range with a best match for ~ Åk 20 , but even

δ

  25 Mg values between + 10 and + 15 ‰ are expected from evaporation experiments and up to +20‰ are found in natural FUN CAIs (e.g., CG-14, EK-1-4-1) with FSi comparable to that of E101.1, the lowest δ 25 Mg values being from evaporation of a forsterite-rich precursor[START_REF] Mendybaev | Thermal and chemical evolution in the early Solar System as recorded by FUN CAIs: Part II -Laboratory evaporation of potential CMS-1 precursor material[END_REF]. The larger degree of mass fractionation in Mg than in Si in these experiments is due to the evaporation of Mg as Mg and Si as SiO, which fractionates less due to its higher mass. Either the E101.1 compositions are indicative of unusual evaporation conditions or the Si and Mg isotopic fractionations were decoupled to some extent. In light of the capture and partial assimilation of the sinuous pyroxene, another capture event can possibly explain the low Mg isotopic composition if the captured material is Mg-rich and Sipoor, such as spinel, and has the usual δ 25 Mg value expected from equilibrium condensation near 0 ‰ if formed from a reservoir with normal planetary Mg isotopic composition. The CMAS composition of the parent melt predating this second capture event is hereafter referred to as the step 2 precapture melt.To understand the origin of the parent material of the E101.1 main host, we calculated what would be its pre-evaporation chemical composition and the extent of Si and Mg loss, based on the measured Si and Mg isotopic compositions and the CMAS composition of the step 1 parental melt calculated above. To account for possible spinel capture, we considered several cases: (1) no additional material, (2) assimilation of condensate spinel with δ 25 Mg = 0 ‰ so that the initial melt composition had a δ 25 Mg of +10 ‰ and (3) assimilation of condensate spinel so that the initial melt composition had a δ 25 Mg = +15 ‰. In cases of assimilating spinel mass balance considerations indicate that the +5‰ δ 25 Mg value is achieved if half of the Mg comes from spinel (case 2) or two-thirds of Mg comes from spinel (case 3). These contributions correspond to a mass fraction of spinel vs CMAS melt of 3.6 wt% and 4.8 wt%, respectively, which is not unreasonably large. The step 2 precapture composition of the parental melt is thus calculated by removing the Al and Mg contribution of these amounts of spinel from the step 1 parental composition. Pre-evaporation calculations were subsequently done in the three cases,[START_REF]A CAI with ultrarefractory REE patterns and enormous enrichments of Sc, Zr and Y in fassaite and perovskite[END_REF] assuming that Si and Mg evaporate as SiO(g) and Mg (g) , (2) using the Rayleigh equation for kinetic evaporation and (3) using either the fractionation factors determined in[START_REF] Knight | Silicon isotopic fractionation of CAI-like vacuum evaporation residues[END_REF] for a sub-liquidus temperature of Type B inclusions of 1400°C, or the fractionation factors determined by[START_REF] Mendybaev | Experimental evaporation of mg-and Si-rich melts: Implications for the origin and evolution of FUN CAIs[END_REF] for the production of FUN inclusions at 1900°C, thus yielding six possible pre-evaporation compositions referred to as step 3 compositions.

  +9.8‰) in three perovskite grains analyzed previously (ElGoresy et al., 2002) suggests a possible connection with FUN inclusions, agreeing with the strong mass fractionation by evaporation indicated by the FSi values. Although unusual, such 48 Ca anomalies without associated 50 Ti anomalies have also been found in hibonite-rich FUN inclusions from

  the nature of the UR precursor of E101.1 from perovskite Differences in the chemistries of perovskite grains in E101.1, notably in their ZrO 2 /Y 2 O 3 ratios, have been interpreted as indicating the presence of four distinct populations of perovskite(El Goresy et al., 2002). These populations include perovskite grains with[START_REF]A CAI with ultrarefractory REE patterns and enormous enrichments of Sc, Zr and Y in fassaite and perovskite[END_REF] chondritic ZrO 2 /Y 2 O 3 ratios ~2.7, suggesting an origin by condensation and (2) grains with ZrO 2 /Y 2 O 3 ratios ~ 0.4, typical of crystal-liquid fractionation indicative of crystallization from a melt. The two other perovskite grains are either (3) devoid of ZrO 2 or (4) have intermediate ZrO 2 /Y 2 O 3 ratios between 0 and ~2.7. Different perovskite populations can also be recognized based on their textural context, although it is not clear if chemical and textural populations can be linked: large anhedral perovksite grains are commonly associated with metal and/or V-Fe-oxide indicative of interactions with metal, and are rimmed by the Sc-Zrrich pyroxene (Fig.4). Such textures are suggestive of partial melting and reaction of preexisting perovskite with a silicate melt and agree with considerations from minor and trace elements (ElGoresy et al., 2002). Other grains, usually much smaller, are euhedral and enclosed within pyroxene, sometimes within pyroxene ± spinel rims around spinel (e.g., in cluster 5, Fig.3d), suggesting late crystallization either from a Ti-rich partial melt or at the solid state. The abundance of perovskite in E101.1 suggests a perovskite-rich UR precursor of bulk chemistry in the hibonite + perovskite + melilite condensation field The chemical and textural diversity of perovskite further indicates contribution of perovskite with (1) diverse condensation histories and (2) variable post-aggregation histories including reaction with a partial melt, new crystallization from a partial melt and solid state equilibration with the surrounding phases.Perovskite grains analyzed for O isotopes are large anhedral grains and have a range of ZrO 2 /Y 2 O 3 ratios. This includes (i) chondritic perovskite (per B), (ii) Zr-free perovskite (4 grains) and (iii) perovskite with intermediate ratios (2 grains). Large variations in ∆ 17 O are observed. On one hand, some grains enriched in16 O to a level approaching that of spinel (∆ 17 O = -19.2 ‰) are found in different populations (Fig.9). On the other hand, both 16 O-rich grains and 16 O-depleted grains with ∆ 17 O up to -8.9 ‰ are found in a single population (Zrfree perovskite). Because (i) the Sc-Zr-rich pyroxene that rims perovskite is among the most16 O-poor mineral in E101.1 and most likely records the composition of the melt, and because

  E101.1, the outermost WLR pyroxenes exhibit the UR-REE pattern typical of the host inclusion (ElGoresy et al., 2002), including the specific LREE/HREE and Lu/Er/Ho relative proportions, which suggests a contribution from the interior inclusion. In one case (Mg34) the δ 25 Mg value of the WLR pyroxene of E101.1 shows an enrichment of +3 ‰, close to that of interior melilite, but in all other cases, both pyroxene and spinel have δ 25 Mg values near 0 ‰, similar to those of the AR forsterite. By contrast, the FSi values span the whole range between the slightly negative values of the AR forsterite and those of the average main host melilite (Fig.17

  ), possibly up to +10 ‰. The Mg isotopic composition is in agreement with expectations from equilibrium condensation in most cases but in the case of spot Mg34 also shows the contribution from interior Mg. The Si isotopic composition indicates a more balanced mixture between Si of condensation origin, similar to Si in the AR forsterite, and Si from the interior with enrichment in the heavy isotopes due to evaporation.Both the REE patterns and the stable Mg, Si isotopic compositions thus indicate some contribution from the interior CAI. Such an influence strongly hints that the WLR did not form by successive deposition of condensate layers at decreasing temperatures or grew from isolated nucleus islands disconnected from the interior inclusion. Nevertheless, both Mg and Si do indicate a contribution of a condensate component. Although analysis Si15 in the innermost WLR pyroxene must be treated with caution, its FSi value of +10‰, heavier than in interior melilite, may be the additional signature of a strong evaporation from a melt. Thus our data support models based on open system behavior in which the WLR formed by a combination of (i) melting of the inner host that preserved the most refractory elements and a fraction of the less refractory Mg and Si and (ii) condensation into the melt of nebular Mg and

1 .

 1 components and on their mixing relationships. Possible precursors have been proposed after disentangling the effects of mixing and partial assimilation. It is found that the final E101.1 CAI results from the coagulation and partial melting of at least three different CAI components: An unusual UR CAI dominated by 16 O-rich Al-diopside (∆ 17 O ≤ -20‰) with light Mg and Si isotopic composition (mass fractionation ~ -3.5‰/amu in both cases) resulting from kinetic effects, most likely during condensation out of equilibrium. This Al-diopside is now present as multiple relict sinuous pyroxene fragments and locally encloses the 16 O-poor FeO-rich minerals (∆ 17 O up to ~ -4.5 ‰) described in El Goresy et al. (2002). 2. A melilite-rich UR and FUN host CAI with abundant perovskite, enriched in heavy Mg and Si isotopes by evaporation (δ 25 Mg and FSi up to ~+7‰ and ~+8‰, respectively). Decoupling of Mg and Si isotopes points toward dissolution of spinel with low δ 25 Mg after evaporation. Isotopic mixing with the sinuous pyroxene CAI in some regions indicates that the precursor was probably an ultra-refractory melt with little SiO 2 and barely any MgO. The subinclusion 1 identified by El Goresy et al. (2002) is most likely a region of the host that was partially crystallized as gehlenite at the time of sinuous pyroxene capture and remelted subsequently. The calculated pre-evaporation composition indicates that the initial precursor condensed as a hibonite-perovskite-melilite assemblage in agreement with the abundance of melilite and perovskite and evidence of a link with hibonite-rich FUN inclusions. 3. The origin of spinel clusters is not as clear but we favor an extraneous source based on several lines of evidence regarding the chemistry, O and Mg isotopes indicating exchange and partial re-equilibration with the host melilite. The whole CAI is surrounded by a WLR and an accretionary forsterite-rich rim, having isotopic systematics indicative of formation by condensation from a 16 O-rich gas (∆ 17 O ~ -22‰) with interactions with the evaporated interior. The Al/Mg systematics suggests an early formation (~canonical) of the WLR, approximately contemporaneous of that of the interior CAI with ( 26 Al/ 27 Al) 0 ~ (5.2±0.1) × 10 -5 , but also secondary perturbations, which could possibly yield erroneously young WLR formation ages in CAIs if not properly identified (( Al !" / Al)

  study of E101.1 also sheds light on the origin of O isotopic variations in CAIs. It has been long debated whether CAIs record primary isotopic heterogeneities in the solar protoplanetary disk or underwent secondary perturbations that reset the isotopic composition of some minerals. Recent studies show strong support in favor of the co-existence of 16 O-rich and 16 O-poor reservoirs early in the protoplanetary disk. The O isotope systematics of E101.1 clearly supports these conclusions and demonstrates that igneous CAIs underwent a 16 Odepletion at high temperature contemporaneously with their growth by coagulation.

Fig. 1 .

 1 Fig. 1. Backscattered electron (BSE) images of CAI Efremovka 101.1. a) Regions of interest.

Fig. 2 .

 2 Fig. 2. Mg Kα X-ray map of the upper portion of E101.1 in log scale. Spinel clusters are

Fig. 3 .

 3 Fig. 3. Example of spinels (sp) with clinopyroxene (cpx) rims and reaction textures. (A) BSE

Fig. 4 .Fig. 5 .

 45 Fig. 4. BSE images of perovskite-metal (pv-met) nodules rimmed by the Sc-Zr-rich pyroxene.

Fig. 6 .

 6 Fig. 6. Oxygen isotopic compositions of melilite in E101.1. (a) Δ 17 O as a function of melilite

Fig. 7 .Fig. 8 .

 78 Fig. 7. Apparent oxygen isotope mass fractionation relative to the CCAM line as a function of ∆ 17 O. All data include perovskite, olivine, nepheline, FeO-rich silicates and glass. 2σ errors.

Fig. 9 .Fig. 11 .

 911 Fig. 9. Oxygen isotopic composition of perovskite as a function of chemistry. Per E and Per B

Fig. 12 .

 12 Fig. 12. δ 25 Mg of spinel clusters and associated melilite. In subinclusion 1, only melilite data

Fig. 13 .

 13 Fig. 13. Chemical and isotopic profiles in melilite from the rim toward the interior. (A) melilite chemistry. (B) stable Mg isotopes. (C) inferred initial 26 Al/ 27 Al. (D) stable Si isotopes. (E) 16 O excess. (F) O isotopic mass fractionation (deviation from the CCAM line).

Fig. 14 .

 14 Fig. 14. Schematic description of the events taken into account to unravel E101.1 thermal

Fig. 15 .

 15 Fig. 15. Mixing calculation of sinuous pyroxene with host melt. (A) Mg isotopes vs O

Fig. 16 .

 16 Fig. 16. Comparison between the E101.1 parental CMAS melt composition (diamonds) and

Fig. 17 .

 17 Fig. 17. Stable Mg and Si isotopic compositions of the rims on E101.1. Orange box -range of

Fig. 18 .Fig. 19 .

 1819 Fig. 18. Comparison between Al/Mg systematics of WLR, theoretical chronology and

  abbreviations, min = minimum, max = maximum, wm = weighted mean, wm 2σ = 2 sigma uncertainty on the weighted mean, sub = subinclusion, sin = sinuous, Mel = melilite, Sp = spinel, Px = pyroxene, Ol = olivine, Fo = forsterite

  ′ is a model initial26 Al/ 27 Al ratio per analysis spot) are best understood if the FUN component was 26 Al-poor, as are many FUN CAIs. (4) A complete Wark-Lovering rim (WLR) surrounds E101.1. Its Mg and Si isotopic compositions indicate that it formed by interaction of the evaporated interior CAI with an unfractionated 16 O-rich condensate component. Heterogeneities in 26 Al content in WLR spinels (3.7 × 10 -5 ≤ ( Al !" / Al)

	!"
	!

  reproducibility on standards are 0.48‰ and 0.94‰ on δ 18 O and δ 17 O, respectively. Deviations from the slope 0.52 terrestrial mass fractionation line are given as ∆ 17 O where ∆ 17 O = δ 17 O -0.52 × δ 18 O. Mass fractionations are noted ∆ 18 O CCAM and are calculated respectively as isotopic deviations from the classical slope 0.94 Carbonaceous Chondrite Anhydrous Minerals mixing line

17,18

O/ 16 O samp -17,18 O/ 16 O SMOW ) / ( 17,18 O/ 16 O SMOW ) × 1000 with SMOW being the Standard Mean Ocean Water. Average 2σ uncertainties taking into account internal precision and

  µm O -primary ion beam of 12-15 nA: 27 Al and the three Mg isotopes ( 24 Mg, 25 Mg and 26 Mg) were detected as positive ions in multicollection mode by using four Faraday cups. The MRP was set at ~ 4200, sufficient to completely separate molecular and doubly charged ion interferences ( 48 Ca 2+ and 48 Ti 2+

  ).27 Al + ,26 Mg + , 25 Mg + and 24 Mg + were counted for 20 cycles of 10 s after a 30 to 50 s presputtering for spinel, pyroxene and olivine and for 100 cycles for melilite analyses except for analyses 60, 61 and 71 to 78 (20 cycles).Standards used for determination of the instrumental mass fractionation and elemental relative sensitivity factors were Burma spinel, a synthetic pyroxene glass with no 26 Mg excess (P0), olivine and pyroxene from San Carlos, and Madagascar hibonite. Measurements of standards were interspersed with sample analyses, thus defining 7 groups of data. For each group, standards data expressed as 1000 × ln[( 25,26 Mg/ 24 Mg) measured )/(25,26 Mg/ 24 Mg) true ] were used to calculate an instrumental mass fractionation law, which was used to determine the 26 Mg excesses (noted δ 26 Mg*) by difference with the expected 26 Mg due to mass fractionation

  , the data can be fit by mixing hyperbolae corresponding to an extremely refractory CMAS melt having approximately 49 wt% Al 2 O 3 , 39 wt% CaO, 10 wt% SiO 2 and 2 wt% MgO. This implies that the sinuous Al-diopside fragments are relicts from a xenolith mostly made of 16 O-rich diopside with light Mg and Si isotopic compositions that

was captured by an extensively molten, highly refractory, 16 O-poor melt. They were partially assimilated in this melt, which subsequently crystallized as melilite with intermediate isotopic compositions in area 3 and 4 and in subinclusion 1, while the remaining melilite preserved the initial melt composition upon crystallization. This melt is hereafter referred to as the step 1 parental melt (Table

Table 1 :

 1 Abbreviations, types of pyroxene, lithological units and model endmembers discussed throughout the manuscript clinopyroxene with variable Al and Ti content ranging from very Al and Ti rich to pure diopside Sinuous pyroxene Al-diopside Sc-Zr-rich pyroxene Ca-rich clinopyroxene with elevated Al, Ti, Sc and Zr content and low SiO 2 content Lithological units Sinuous pyroxene fragments Sinuous fragments dominated by Al-diopside and locally containing FeO-rich assemblages Clusters 1 to 7 Coarse-grained spinel clusters in the main host Core Central portion of the main host, mostly melilite Profiles 1 and 2 Chemical and O, Mg, Si isotopic traverses started from the rim and directed towards the center of the inclusion, ~perpendicular to the surface Areas 3 and 4 Regions between the sinuous pyroxene fragments and the rim, petrographically similar to those sampled by profile 1 and 2 but with different isotopic systematics Subinclusions 1 and 2 Subcircular melilite + spinel regions nearly enclosed in sinuous pyroxene, proposed to be xenoliths in El Goresy et al. (2002) Quenched glass Regions with chemistry intermediate between melilite and pyroxene with Schlieren-like heterogeneities and dendritic melilite micro-crystals, located at the sinuous pyroxene -host melilite interface

	Denomination	Meaning
	Abbreviations
	AOA	Amoeboid Olivine Aggregate
	AR	Accretionary Rim (forsterite-rich in E101.1)
	WLR	Wark Lovering Rim
	CAI	Ca-Al-rich Inclusion
	UR	Ultra-Refractory
	FUN	Fractionated with Unknown Nuclear effects
	FG-sp-rich CAI	Fine-Grained spinel-rich CAI with nodular texture
	CMAS	CaO + MgO + Al 2 O 3 + SiO 2 chemical system
	IMF	Instrumental Mass Fractionation
	RSF	Relative Sensitivity Factor
	Types of pyroxene
	WLR pyroxene Ca-rich Model endmembers
	Step 1 parental melt	Ideal CMAS composition required to match the observed composition of intermediate areas (3, 4 and
		subinclusion 1) upon mixing with sinuous pyroxene
	Step 2 precapture melt	Ideal CMAS composition calculated from step 1 parental melt by substraction of condensate spinel
	Step 3 preevaporation	Ideal CMAS composition calculated from step 2 precapture melt by correction for evaporation based on Mg and
	composition	Si isotopes

Table 2 :

 2 Summary of O, Mg and Si isotopic analyses in E101.1 lithological units and minerals * except AR Fo data given as δ 26 Mg excess in ‰ † CAI-like nodules within the AR # number of analyses

				∆ 17 O (‰)				δ 25 Mg (‰)			( 26 Al/ 27 Al) 0 ' [10 -5 ]**				FSi (‰)
		#	min	max	wm	wm 2σ #	min	max	wm	wm 2σ #	min	max	wm	wm 2σ #	min	max	wm	wm 2σ
	Profile 1 Mel	13 -7.44	-1.86	-5.01	0.27	9	3.69	6.28	5.04	0.57	9	2.44	4.52	3.43	0.13	8	4.98	6.73	6.29	0.15
	Profile 2 Mel	10 -9.32	-5.73	-6.93	0.31	10	1.88	6.91	5.06	0.55	10	2.91	4.97	3.95	0.11	9	5.16	6.56	5.66	0.13
	Core Mel	3	-6.6	-5.12	-6.07	0.57	2	3.99	5.01	4.50	1.23	2	3.98	4.36	4.15	0.48	4	5.19	7.66	6.59	0.20
	area 3 Mel	5 -10.18 -7.28	-8.64	0.43	5	0.88	2.01	1.43	0.47	5	2.83	4.09	3.70	0.11	1			1.24
	area 4 Mel	5	-9.59	-7.42	-8.03	0.4	5	0.54	2.15	1.49	0.47	5	3.91	5.39	4.71	0.11	3	2.8	3.53	3.25	0.25
	sub 2 Mel	2 -17.45 -12.82 -15.14
	sub 1 Sp						3	-0.77	0.58	-0.02	0.48	3	4.53	5.84	5.04	0.37
	sub 1 Mel	6 -12.18	-7.8	-8.78	0.36	4	-1.51	1.06	-0.12	0.53	4	2.96	5.04	4.07	0.13	3	3.99	5.69	4.99	0.26
	cluster 1 Sp	1			-22.83		3	6.42	6.70	6.53	0.51	3	4.93	5.72	5.30	0.89
	cluster 1 Mel*	2 -10.31 -3.68	-7.15	0.71	2	5.18	5.71	5.47	0.78	2	4.00	4.53	4.22	0.17
	cluster 2 Sp	1			-22.46		2	4.81	5.29	5.05	0.62	2	6.46	7.56	7.30	0.90
	cluster 2 Mel*	1			-6.14		2	5.02	5.77	5.40	0.75	2	3.62	3.90	3.77	0.18
	cluster 3 Sp						1			4.28		1			3.62
	cluster 3 Mel*						1			6.10		1			3.34
	cluster 4 Sp						1			6.63		1			5.30
	cluster 4 Mel*
	cluster 5 Sp	1			-21.77		2	2.61	2.65	2.63	0.52	2	5.47	5.74	5.62	0.35
	cluster 5 Mel*	6 -11.25 -5.51	-7.33	0.43	2	3.65	6.09	4.87	0.75	2	4.07	4.57	4.32	0.20	2	5.6	5.83	5.71	0.31
	Sin Px	7 -23.56 -18.53 -20.98	0.36	4	-4.65	-2.39	-3.40	0.31	4	4.42	13.20	5.21	1.28	3	-4.48	-2.37	-3.09	0.37
	Sin Px -FeO	1			-4.65
	WLR all	13 -22.23 -15.29 -19.56	0.23	12 -1.50	3.04	0.50	0.19	12	3.68	6.54	4.78	0.24
	WLR Mel	4 -19.84 -16.81 -18.45	0.45
	WLR Ol	3 -22.23 -17.97 -20.35	0.49	1			-0.37
	WLR Px	4 -21.66 -15.29 -19.32	0.36	4	-1.22	3.04	0.86	0.25	4	4.95	6.54	5.16	0.99	8	-0.34	9.97	2.72	0.16
	WLR Sp	2 -22.05 -20.88 -21.46	0.68	7	-1.5	1.25	0.19	0.32	7	3.70	5.70	4.76	0.25
	AR all	6 -22.41 -12.81 -19.00	0.35
	AR Fo CAIs in AR †	3 -22.41 -21.81 -22.18 3 -17.62 -12.81 -15.83	0.49 0.49	5	-1.49	1.39	-0.02	0.25	5	0.01	0.09	0.06	0.04	4	-1.88	-0.46	-1.34	0.24

* mel in close contact with sp cluster *

Acknowledgements

Jean-Noël Rouzaud provided access to the FEG-SEM operated by Damien Deldicque in the geology department at the ENS in Paris. Emeline Charon helped with the ENS FEG-SEM imaging and for a few oxygen isotopes analyses. George Jarzebinski maintained the UCLA IMS 1270 ion probe. Frederic Couffignal and Michel Fialin from the CAMPARIS facility helped with the electron probe analyses. Interaction with Larry Grossman was appreciated,

The coupling of O isotopes with Si and Mg isotope compositions indicates several episodes of 16 O-rich CAI proto-material dissolution into a 16 O-poor melt. This is clearly the case for (1) the 16 O-rich spinel clusters, which reacted with a 16 O-poor host melt, resulting in local 16 O-enrichments in the gehlenitic products of the reaction and (2) the 16 O-rich sinuous pyroxene fragments, which mixed with a 16 O-poor parental host melt to produce the intermediate compositions observed in areas 3 and 4 and in subinclusion 1. The strong 16 Odepletion observed in the Sc-Zr-rich pyroxene, which likely crystallized upon interaction between the host melt and UR perovskite, some of which being 16 O-rich as observed in PerB, further testifies to a strongly 16 O-depleted host melt from which most of the melilite in E101.1 crystallized. The sluggish self-diffusion of O in Ca-rich pyroxene [START_REF] Ryerson | Determination of oxygen self-diffusion in akermanite, anorthite, diopside, and spinel: Implications for oxygen isotopic anomalies and the thermal histories of Ca-Al-rich inclusions[END_REF], requires that this 16 O-depletion is primary. The residual ∆ 17 O heterogeneity within individual areas (Table 2) and the possible oxygen isotope mass fractionation relative to the CCAM line observed in the Sc-Zr-rich pyroxene and some of the 16 O-poor melilite (Fig. 7, Table 2) are well accounted for by assimilation of several 16 O-rich proto-CAIs into a dominantly 16 O-poor melt that had undergone partial evaporation.

All these observations point toward a 16 O-poor CAI precursor at the origin of the E101.1 parental melt (or a strong 16 O-depletion occurring during melting), before the host CAI trapped both the spinel clusters and the sinuous pyroxene xenoliths. The 16 O-depletion must therefore be contemporaneous with the main CAI growth stage by coagulation of proto-CAIs in the solar protoplanetary disk. The Al/Mg systematics of E101.1 further suggests that this growth occurred while the solar system 26 Al/ 27 Al ratio was canonical or near canonical. In agreement with increasing evidence [START_REF] Yurimoto | Oxygen isotope exchange between refractory inclusion in Allende and solar nebula gas[END_REF][START_REF] Aléon | Oxygen isotope heterogeneities in the earliest protosolar gas recorded in a meteoritic calcium-aluminum-rich inclusion[END_REF][START_REF] Aléon | Oxygen isotopes in the early protoplanetary disk inferred from pyroxene in a classical type B CAI[END_REF][START_REF] Kawasaki | Oxygen isotopic and chemical zoning of melilite crystals in a type A Ca-Al-rich inclusion of Efremovka CV3 chondrite[END_REF][START_REF] Kawasaki | Chronological study of oxygen isotope composition for the solar protoplanetary disk recorded in a fluffy Type A CAI from Vigarano[END_REF][START_REF] Katayama | Oxygen isotopic zoning of reversely zoned melilite crystals in a fluffy type A Ca-Al-rich inclusions from the Vigarano meteorite[END_REF][START_REF] Park | Oxygen isotopic composition of the solar nebula gas inferred from high-precision isotope imaging of melilite crystals in an Allende CAI[END_REF]Simon et al., 2011[START_REF] Simon | Oxygen isotopic variations in the outer margins and Wark-Lovering rims of refractory inclusions[END_REF], this conclusion indicates that both 16 O-rich (∆ 17 O ~ -24‰) and 16 O-poor (∆ 17 O ~ -2‰) refer to this stage as the "cryptic stage" reflecting that it is not possible to be more precise as to where, when and how the 16 O-depletion occurred.

The sinuous pyroxene fragments contain complex inclusions of secondary phases including wollastonite and oxidized FeO-rich silicates inferred to be andradite and hedenbergite from electron probe analyses by El Goresy et al. (2002). The present discussion is based on this characterization. We show that these secondary FeO-rich minerals are 16 Opoor. They are systematically located in the interior of the sinuous pyroxene and they contributed to the FeO content of the quenched melt region. The host inclusion, by contrast, contains numerous metal inclusions, some of which are quite large. Exsolution of silica, phosphates and V-Fe-oxides (Fig. 4d) from the metal grains is evidence of metamorphism, but most metal grains remain unoxidized. Silicates in the vicinity of metal grains are FeOfree. Although it is widely thought that oxidation and FeO-enrichment of CAI minerals occurred on a parent body in most cases (Brearely and Krot, 2013), these observations strongly suggest that the secondary alteration of the sinuous pyroxene in E101.1 occurred in the protoplanetary disk prior to trapping by the reduced host inclusion (El Goresy et al., 2002). Given that small refractory metal nuggets are present in the sinuous pyroxene fragments, this secondary alteration/oxidation was extremely localized. The conditions of formation of these secondary FeO-silicates are unclear and deserve a detailed study. They point toward existence of a high fO 2 reservoir in the region of CAI formation. This reservoir is not expected from thermodynamic equilibrium considerations [START_REF] Fedkin | The fayalite content of chondritic olivine: Obstacle to understanding the condensation of rocky material[END_REF]Grossman et al., 2008a) and has a near terrestrial oxygen isotopic composition Finally, the relationship between O isotopes and chemistry in sinuous pyroxene (Fig.