A new type of oxidized and pre-irradiated micrometeorite

Accepted Manuscript A new type of oxidized and pre-irradiated micrometeorite Carole Cordier, Bastian Baecker, Ulrich Ott, Luigi Folco, Mario Trieloff PII: DOI: Reference:

S0016-7037(18)30207-2 https://doi.org/10.1016/j.gca.2018.04.010 GCA 10731

To appear in:

Geochimica et Cosmochimica Acta

Received Date: Accepted Date:

21 September 2017 5 April 2018

Please cite this article as: Cordier, C., Baecker, B., Ott, U., Folco, L., Trieloff, M., A new type of oxidized and preirradiated micrometeorite, Geochimica et Cosmochimica Acta (2018), doi: https://doi.org/10.1016/j.gca. 2018.04.010

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A new type of oxidized and pre-irradiated micrometeorite Carole Cordiera*, Bastian Baecker b,c, Ulrich Ottb,c,d, Luigi Folcoe, Mario Trieloffb

a

Univ. Grenoble Alpes, Univ. Savoie Mont Blanc, CNRS, IRD, IFSTTAR, ISTerre, 38000 Grenoble,

France b

Institut für Geowissenschaften, Klaus-Tschira-Labor für Kosmochemie, Universität Heidelberg, Im

Neuenheimer Feld 234-236, 69120 Heidelberg, Germany c

Max-Planck-Institut für Chemie, Hahn-Meitner-Weg 1, 55128 Mainz, Germany

d

Magyar Tudományos Akadémia Atommagkutató Intézet, Bem tér 18/c, 4026 Debrecen, Hungary

e

Dipartimento di Scienze della Terra, Università di Pisa, Via S. Maria 53, 65126 Pisa, Italy

* Corresponding author: [email protected]

Abstract This paper investigates the mineralogy and noble gas composition of a unique micrometeorite from the Transantarctic Mountains, #45c.29. The magnetite rim and the particle interior with olivine, pyroxene and magnetite relict grains (30 to 250 µm in size) set in a vesicular mesostasis are typical features of coarse-grained, partially melted micrometeorites. Particle #45c.29 stands out from other micrometeorites of this type by the texture of the mesostasis made of abundant plagioclase and augite laths, the remarkably high Ni contents in magnetite and olivine relict grains, and by the similarly high abundance of cosmogenic noble gases (21Necos up to 1.62 x 10-7 cm3 STP/g and 38Ar up to 7.2 x 10-8 cm3 STP/g). The high Ni content of Fa26 olivine relict grains (NiO ~ 0.65 wt%), the high Ni (NiO ~ 0.8 wt%) and Ti (TiO2 ~ 0.3 wt%) contents of magnetite relicts, and the oxygen isotope composition of a sample of the particle (δ18O ~ 2.3 ‰, δ17O ~ -1.5 ‰), suggest a parentage with rare equilibrated CK chondrites. Pyroxene and plagioclase are not expected to crystallize during atmospheric entry of micrometeoroids. Their occurrence in #45c.29 may be explained by the Ca-, Al- and Na- rich composition of its precursor - in agreement with the high abundance of plagioclase reported in the matrix of CK chondrites - if combined with a relatively low cooling rate and, therefore, unusual atmospheric entry parameters (velocity/angle) of the micrometeoroid. Given these specific entry parameters, the particle has recorded unique information on mineralogical and textural transformations of micrometeoroids during atmospheric entry, with solid-state oxidation of the olivine relict grains in the igneous rim, and partial melting of relict mineral phases and relict/melt reactions in the particle interior. The cosmogenic 21Ne/22Ne ratio of 0.94 ± 0.02 is incompatible with major production by cosmogenic ray irradiation of a small particle in space. We propose that micrometeorite #45c.29 mostly records an earlier irradiation stage, in a meteoroid or more likely near the surface (< 20 cm in depth) of an asteroid. In contrast, most of the other

unmelted and scoriaceous micrometeorites analyzed for noble gases – if coming from asteroidal sources of the Main Belt – seem to have sampled deeper parts of their parent body, where they were shielded from cosmic rays and from where they were excavated during highenergy disruptive processes.

Keywords: micrometeorite; interplanetary dust; CK chondrites; oxidation; noble gases; cosmogenic Ne and Ar; pre-irradiation

1. Introduction

Micrometeorites are microscopic (typically < 2 mm in size) extraterrestrial particles that are recovered at the surface of the Earth, mainly from polar areas and deep-sea sediments (see Folco and Cordier, 2015, for a recent review). The classification of micrometeorites (Genge et al., 2008) is mainly based on the degree of preservation of their pristine texture and mineralogy, which decreases with increasing melting during atmospheric entry from unmelted micrometeorites, to partially melted (i.e., scoriaceous) micrometeorites and cosmic spherules. The affinity of unmelted and scoriaceous micrometeorites, in which relict grains are preserved, with known meteorite types is mainly determined through textural and compositional data of relict mineral phases (e.g., Kurat et al., 1994; Genge et al., 1997; Genge, 2008; Gounelle et al., 2009; Taylor et al., 2012; van Ginneken et al., 2012; Imae et al., 2013). In that respect, fine- and coarse-grained unmelted micrometeorites related to chondritic meteorites contain abundant dehydrated phyllosilicates and/or olivine and pyroxene relict grains, in some instances still arranged in chondrules. In the more heated scoriaceous micrometeorites, these relict phases are usually set in glass produced during partial melting at the eutectic (Genge, 2017) plus microphenocrysts of olivine and magnetite. Relicts of plagioclase are rare, even if plagioclase is a common constituent of most meteorites – especially the most metamorphosed ones (e.g., Rubin, 1997). This is easily understood since, in general, plagioclase disappears more rapidly than ferromagnesian silicate phases and oxides as melting proceeds (Genge, 2017). Plagioclase (and pyroxene) crystallites grown during atmospheric entry are also usually lacking in the micrometeorite record as their crystallization is impeded by the fast cooling experienced by the micrometeorites during atmospheric entry (Taylor and Brownlee, 1991). Plagioclase thus occurs rarely in micrometeorites and so far, it has only been observed as relict grains in the less heated

unmelted particles, usually derived from ordinary chondrites (Genge, 2008; van Ginneken et al., 2012) or from differentiated parent bodies (e.g., Gounelle et al., 2009; Badjukov et al., 2010). In this study, we report an in-depth petrographic, isotopic and noble gas study of a new type of coarse-grained, scoriaceous micrometeorite from the Transantarctic Mountain collection (Rochette et al., 2008) - named #45c.29. This particle consists of olivine, pyroxene and magnetite relict grains set in a fine-grained mesostasis made of oriented plagioclase microlaths with interstitial pyroxene. The assemblage and abundance of plagioclase and pyroxene are similar to differentiated unmelted micrometeorites described in the literature (Gounelle et al., 2009; Badjukov et al., 2010), whereas the mineral composition and bulk oxygen isotopes point towards a chondritic origin. We show that plagioclase and pyroxene in the mesostasis conspicuously crystallized during atmospheric entry, a process never described before in the micrometeorite literature. Besides adding to the micrometeorite inventory, our study of #45c.29 brings important new information on the partial melting of mineral phases and reactions between relict grains and melt under oxidizing conditions during atmospheric entry. This particle also stands out from other micrometeorites in terms of noble gases, with abundances of cosmogenic 21Ne and 38Ar and a cosmogenic 21Ne/22Ne ratio that strongly suggest that its precursor had a unique irradiation history amongst the known micrometeorite population.

2. Sample and analytical techniques The micrometeorite #45c.29 was found at the top glacial surface of the Miller Butte granitic nunatak (71°54.213’S, 160°29.389’E) in the Transantarctic Mountains (TAM), during the 2009 Programma Nazionale di Ricerche in Antartide (PNRA) expedition. Detailed descriptions of the Transantarctic Mountain micrometeorite traps are given in Rochette et al.

(2008). The onset of the micrometeorite accumulation in the traps is estimated to be as old as ca. 1 Ma, based on the cosmogenic nuclide exposure age of the host glacial surface (van der Wateren et al., 1999) and on the occurrence in the trap of ca. 0.8 Ma old Australasian microtektites (Folco et al., 2008). The particle was found in the 400-800 µm magnetic extract. The micrometeorite #45c.29 was manually split into several fragments used for petrographic observations and mineral composition analysis, noble gas analysis and bulk oxygen isotope analysis. The largest fragment was embedded in epoxy, sectioned and polished for petrographic observations using the SEM-EDS facilities at Siena University and ISTerre (Grenoble Alpes University) and the FEG-SEM JEOL 6500F at Istituto Nazionale di Geofisca e Vulcanologia (INGV, Rome). Modal abundances were measured from the backscattered electron image shown in Figure 1, using the Image-Pro Plus software. Mineral and glass compositions were determined by Electron Probe Micro-Analyses (EPMA) using three different microprobes, a JEOL JXA 8200 Superprobe at INGV (Rome), a Cameca SX 100 at the National History Museum (NHM) of London, and a JEOL JXA-8230 at ISTerre. Analytical conditions were 15 kV accelerating voltage and 7 nA beam current (INGV) and 15 kV accelerating voltage and 20 nA beam current (NHM and ISterre). In addition, high precision analyses of minor elements in representative olivine grains were obtained on the JEOL JXA-8230 microprobe at ISTerre by combining high beam current (300 nA), long counting time (9 minutes) and correction of the instrumental drift by standard bracketing (San Carlos olivine standard USNM 111312/444 and XEN internal laboratory standard; Batanova et al., 2015). ZAF correction was applied at NHM and ISTerre and PAP correction was applied at INGV. We only considered analyses with oxide sum ranging from 97 to 102 and with cation sum in the structural formula differing by less than 0.05 a.p.f.u. from the ideal cation sum. Element maps were also obtained using wavelength-dispersive

spectrometers (WDS) at NHM, London, with current conditions similar to those used for spot analyses, and at ISTerre, Grenoble, with 20 kV, accelerating voltage, 500 nA beam current and 400 ms dwell time. The triple oxygen isotope composition of an 11-µg sample of #45c.29 was measured by laser fluorination gas mass spectrometry at the Centre for Stable Isotope Research and Analyses (Göttingen, Germany). The analytical procedure is detailed in Pack and Herwartz (2014). The fragment was fluorinated with purified F2 gas and the released O2 was purified from excess F2 by reaction with NaCl and cryogenic separation of the resulting Cl2. Sample O2 was analyzed on a ThermoElectron MAT 253 gas mass spectrometer. Oxygen isotopic analyses are reported in standard δ notation, in ‰ vs. V-SMOW: δ18O (‰) = [(18O/16Osample/18O/16OV-SMOW) - 1] x 1000 and similarly for δ17O using the 17O/16O ratio. 17O, which represents the deviation from the terrestrial fractionation line, was calculated as 17O = δ17O – 0.52 δ18O (Clayton, 1993). Precision is typically ± 0.22 ‰ for δ18O, ± 0.15 ‰ for δ17O and ± 0.07 ‰ for 17O. Terrestrial samples of similar weight were measured during the same analytical session as #45c.29 and their composition plots on the terrestrial fractionation line, suggesting that the method is applicable for small samples. We analyzed the noble gases, from He to Xe, in two different pieces of #45c.29 of roughly similar size (~ 300 µm x 200 µm). The mass of each aliquot was measured using a microbalance Mettler-Toledo: 22.1 µg for sample #1 and 24.2 µg for sample #2. The noble gas analyses were performed at MPI for Chemistry (Mainz) using CO2 laser heating (MIR 1030, New Wave Research, wavelength of 10.6 μm, i.e. in the mid infrared range) coupled with isotope measurements on a high sensitivity Noblesse noble gas mass spectrometer (Nu Instruments) with multiple ion counting (8 channeltrons, with an additional Faraday cup used for 4He and 40Ar). We used different heating procedures for the two samples, each based on a number X of 1-minute heating steps at different Y laser current and referred hereafter as X x Y

Watt. For sample #1, the first heating step (2 x 5 Watt) was followed by a second step of 1 x 30 Watt. A third extraction (3 x 5 to 30 Watt) released an additional small amount of 3He. For sample #2, detectable amounts of gas were released during the first step (3 x 0.91 Watt) but not during the second step (3 x 5 Watt). Multi-ion counting was applied in the analysis of Kr and Xe. A difference correction scheme was applied to minimize the cross-talk correction between the different detectors, the corrections being usually much smaller than analytical uncertainties. For xenon, we did not use the “internal” method for multi-ion counting, in contrast to what has been done by Baecker (2014) for other Transantarctic Mountain micrometeorites, because this approach yielded some unrealistic calibration factors for sample #2. The results reported for #45c.29 are those obtained by the “classical” offline calibration method and are a weighted average calculated from the results obtained using the two different collector configurations with which data were taken. Signals measured for the rare isotopes 124Xe and 126Xe were small (~1 cps) and the 124Xe/132Xe and 126Xe/132Xe ratios have, therefore, large uncertainties. More experimental details are given in Baecker (2014).

3. Results

3.1 Petrography

The particle #45c.29 is a 710 µm x 590 µm sub-rounded complete micrometeorite (Fig. 1A, inset) showing an almost continuous shell of microscopic magnetite crystals, known in the literature as the magnetite rim (e.g., Genge et al., 2008). A backscattered electron (BSE) image of half of the particle is shown in Figure 1A and a simplified sketch of the main constituents is shown in Figure 1B. In the petrographic description below, the terms “dark

gray”, “light gray” and “bright” refer to the BSE contrast of olivine and pyroxene phases, which is a good proxy for the Mg/Fe ratio (decreasing from dark to bright). The degree of weathering is low - corresponding to level 1a in the classification scheme of van Ginneken et al. (2016), with no visible encrustation and minor loss of olivine grains in the outermost parts of the particle only. Three domains characteristic of scoriaceous micrometeorites are identified. From core to rim, these are: the vesicular particle interior, the igneous rim and the magnetite rim. The magnetite rim is visible on the edges of the section, except on the lower edge which corresponds to the splitting plane. It consists of a 5 µm-thick alignment of micron-sized magnetite crystals (Figs. 2A and B). The igneous rim (10 vol.%) extends over ~ 15 µm between the particle interior and the magnetite rim, except on the right side of the section where the particle interior is directly in contact with the magnetite rim. This igneous rim is made of abundant glass with minor euhedral magnetite and skeletal crystallites which were too small to be analyzed, but which could be olivine or pyroxene. The spatial distribution of the crystallites is heterogeneous, with abundant crystallites on the upper side (Fig. 2B) and almost none on the left side of the section (Fig. 2C). In addition, the igneous rim contains very dark rounded or euhedral grains speckled with numerous oriented tiny Fe-oxide inclusions (Figs. 2C and D). When euhedral (Fig. 2D), the shape of these grains suggests that they were formerly olivine but their dark BSE contrast suggests a Mg-rich composition compared to the olivine grains of the particle interior. The particle interior (90 vol.%) is made of coarse olivine (40 vol.%) and pyroxene (8 vol.%) grains together with rare oxide grains (2 vol.%), set in a fine-grained mesostasis (40 vol.%). The mesostasis is made of oriented plagioclase laths with interstitial pyroxene, irregularly shaped Fe-oxides and variable amounts of glass (Figs. 3A and B). Elongation of the plagioclase laths defines domains with distinct orientations (Figs. 1 and 3B). Glass-rich

areas are observed near large olivine and pyroxene grains (Fig. 1). These glassy pockets contain euhedral to skeletal olivine grains, dendritic pyroxene crystallites, minor Fe-oxide grains and, contrary to the igneous rim, they also contain randomly orientated plagioclase laths (Figs. 3A and C). Some of the glassy pockets are, however, almost devoid of crystallites (Fig. 3D). On the lower part of the section (Fig. 1), two large (up to 50 µm) aggregates made of magnetite and Al-spinel intergrowth show vermicular rims and are associated with small blebs of Fe-Ni metal, less than 1 µm in diameter (Fig. 3E). Small spherical vesicles are dispersed throughout the particle interior but are concentrated in glass-rich pockets near olivine and pyroxene grains (Fig. 1). The lower edge of the section, which corresponds to the splitting plane, shows a lobate morphology, resulting from the fragmentation of the particle through large vesicles. The left-hand side of the particle is made of a large (250 x 150 µm in size) olivine grain (Ol3, Fig. 1). The grain is fractured and the open fractures are filled with glass and magnetite (Fig. 3F). The olivine grain is also decorated by fine trails of oxide inclusions. When in contact with the igneous rim, the olivine has a dark gray shade on BSE images of Figures 2A and 2C (i.e. relatively Mg-rich composition), with numerous tiny Fe-oxide inclusions, preferentially aligned along the crystallographic direction of the olivine crystals. The oxide inclusions are systematically associated with small domains that show black in BSE images (Fig. 2C). With increasing distance from the igneous rim, the olivine composition goes back to a less Mg-rich composition (i.e. light gray shade on BSE images) and Fe-oxide grains cluster in melt-filled fractures, together with rare euhedral olivine grains (Figs. 2A and E). In addition to mesostasis and glassy pockets, pyroxene occurs as large (30 µm) crystals surrounded by glass and as a large aggregate in the center of the particle (Fig. 1). This aggregate is made of two dark nuclei surrounded by vesicular melt, speckled with small

oxides and rich in spherical vesicles and cup-shaped pores less than 1 µm in size (Fig. 3H). The dark nuclei and the vesicular melt are rimmed by large pyroxene margins (Fig. 3G), themselves not homogeneous (Fig. 3G). A thin bright overgrowth with dendritic shape is also commonly observed around the large pyroxene grains (Figs. 3D, G and H).

3.2 Mineral composition

Representative mineral compositions are given in Table 1 and the whole dataset is available in Supplementary Table S1. X-ray element maps are shown in Figures 4 and 5. In the particle interior, olivine has high Fa content (26.5 ± 1.6 mol.%) and high Ni content (0.67 ± 0.20 wt%). In the igneous rim, the oxide-bearing olivine grains are more magnesian, with up to 90 mol.% Fo (semi-quantitative EDS analysis, not shown here). We also analyzed the bulk composition of the assemblage of Mg-rich olivine and small Fe-oxide grains with a 10µm electron beam. The composition is not stoichiometric but it nevertheless gives a Fa content (Fa22) in the range of the olivine in the inner part of the particle. This similarity is also obvious in the Fe and Mg distribution maps shown in Figure 4 where oxide-bearing Mg-rich olivine grains of the igneous rim cannot be discriminated from olivine of the particle interior. The two nuclei in the pyroxene aggregate are made of Mg-rich pigeonite (Fs2-5 Wo7-10) while the aggregate rims and other grains have augite composition (Fs2-12Wo30-45). Pigeonite shows relatively large ranges of Fe/Mg and Fe/Mn atomic ratios, from 0.02 to 0.12 and from 5 to 50, respectively. Figure 4 shows the X-ray element maps of most of the studied fragment and Figure 5 a close-up view of the large central pyroxene aggregate. The pigeonite nuclei observed in BSE images (Figs. 1A and 3G) are well discriminated using the Mg (and Fe) Kα X-ray maps, but they are not recognizable using the distribution of other elements. The chemical maps rather show the occurrence of a large (>200 µm-long) pigeonite grain

surrounded by wide augite rims in the center of the particle. This indicates that the vesicular melt surrounding the nuclei has a pigeonite composition, but it is slightly enriched in Fe compared to the nuclei. The contact between the vesicular melt and the augite rims is not clear in BSE images (Fig. 3H) and is mainly identified based on the Ca and Ti distribution maps as well as the distribution of small vesicles and pores (Figs. 3G and 5). The augite rims are not homogeneous as shown in the BSE images of Figure 3. From dark to light gray augite, the Fe/Mg atomic ratio increases from 0.03 to 0.20, and the TiO2 and Al2O3 contents decrease (Table 1 and Fig. 5). The two vermicular aggregates of oxide grains consist of intergrowths of magnetite and Mg-bearing hercynite. Magnetite in the intergrowths is rich in TiO2 (0.3 to 0.4 wt%), MgO (2.1 to 2.5 wt%), Cr2O3 (3.8 to 6 wt%) and NiO (0.7 to 0.9 wt%) compared to magnetite grains grown during atmospheric entry (e.g., Toppani and Libourel, 2003; van Ginneken et al., 2012). Small FeNiS spherical blebs (ca 25 wt% Ni, SEM semi-quantitative analysis) are observed in inclusions and around the vermicular oxide grains. In the mesostasis, plagioclase ranges in composition from bytownite to anorthite (An 87 to An95). It has high FeO (1.2 to 1.8 wt%) and MgO (0.7 to 1.2 wt%) contents, even if we cannot exclude that these high contents are mixed analyses with adjacent pyroxene due to the relative small size of the analyzed grains compared to the beam spot. Pyroxene in the mesostasis was too small and dendritic to be analyzed, but we estimated its composition from the bulk composition of the mesostasis determined by defocused-beam (~10 µm in diameter) microprobe analysis. Using elements that are not partitioned into plagioclase (Ti, Cr), we found that pyroxene represents 40 vol.% of the analyzed area, which is consistent with the petrographic observations. From this, we back-calculated the composition of the pyroxene using the composition of the bulk mesostasis, that of plagioclase and the proportion of

pyroxene. The resulting stoichiometric pyroxene is Fs20Wo30-35 and has approximately the same element contents as the augite rim (Table 1). Glass in the igneous rim and in the inner part of the particle has a similar composition (Table 1 and Fig. 6). In particular, both have high Al2O3 contents (12-17 wt%, Fig. 6A) compared to mineral phases. Hence, the glass distribution is clearly visible in the Al Kα X-ray map of Figure 4 (green color). Consistently with petrographic observations, the glass is localized in the igneous rim and along olivine and pyroxene grain boundaries and in their fractures. The composition of the glass in the particle interior is independent from the nature of the adjacent relict grains (Table 1). The TiO2 content (Fig. 6B) is slightly lower in the igneous rim (0.48 wt%) than in the particle interior (> 0.60 wt%). The bulk mesostasis has higher Al2O3 and lower MgO contents than the glass of the igneous rim and glassy pockets (Fig. 6).

3.3. Oxygen isotopes

The three-oxygen isotope analysis yields δ18O = 2.28 ± 0.22 ‰, δ17O = -1.54 ± 0.15 ‰ and 17O = -2.73 ‰. This composition plots below the terrestrial fractionation line (TFL), in between the Carbonaceous Chondrite Anhydrous Mineral Line defined from anhydrous minerals in CAIs (CCAM, Clayton et al., 1977) and the Primitive Chondrule Mineral Line defined from minerals – relicts apart - in chondrules (PCM; Ushikubo et al., 2012, Fig. 7). A thin igneous rim is preserved on one side of the sample analyzed for oxygen isotopes (Fig. 7A, inset), but it represents a negligible fraction of the analyzed mass. We thus assume that the measured composition is representative of the particle interior and of the micrometeoroid before Earth’s atmospheric entry. This is consistent with the relatively low δ18O value that suggests that the analyzed sample did not undergo significant oxygen exchange with Earth’s

atmosphere (δ18O =23.5 ‰, Thiemens et al., 1995) nor significant mass-dependent fractionation due to evaporation during atmospheric entry heating. These conclusions are confirmed by mass balance calculation (Fig. 7B). If the precursor of #45c.29 plots on the PCM (as most of the anhydrous minerals in chondrules of carbonaceous chondrites), our calculation suggests that atmospheric O2 may contribute up to 13% to the oxygen isotopic composition of #45c.29. This is a maximum estimate as bulk chondrites - as well as their components - are shifted to the right of PCM line, and because it does not include the effect of mass fractionation due to atmospheric entry heating (Fig. 7B). We note, however, that the modeled contribution of atmospheric oxygen is similar to the modal abundance of magnetite and igneous rims (10 %) in the sample used for petrographic observations.

3.4. Noble gases

Noble gas amounts (cm3 STP) and concentrations (cm3 STP/g) for the two samples of #45c.29 are summarized in Table 2. Below, the results for #45c.29 are compared with noble gas data obtained on 29 other TAM micrometeorites and reported by Baecker (2014), as well as those reported in the literature for micrometeorites from other Antarctic collections (Osawa et al., 2000; Osawa and Nagao, 2002a and 2002b; Osawa et al., 2003a and 2003b; Osawa et al., 2010) and for interplanetary dust particles (IDPs, Nier and Schlutter 1993; Pepin et al., 2000).

3.4.1. Helium

The 3He concentration of (5.9 ± 0.1) x 10-8 cm3 STP/g in sample #1 is the highest of all TAM micrometeorites. The 3He/4He ratio (~ 149 x 10-4) is also higher than in all other

analyzed TAM and Antarctic micrometeorites. It is also higher than in most IDPs, where it is only exceeded by a special IDP with a 3He/4He ratio of 200 x 10-4 (Nier and Schlutter, 1993). In contrast, sample #2 has only half as much 3He, together with an order of magnitude higher 4

He, resulting in a significantly lower 3He/4He ratio (~ 6 x 10-4). Comparison with

cosmogenic 21Ne and 38Ar demonstrates that the bulk of cosmogenic 3He (more than 80 %) must have been lost, even in sample #1 in which the high 3He is almost completely of cosmogenic origin. Hence, helium will not be discussed further.

3.4.2. Neon

In the classic three-isotope plot of 20Ne/22Ne versus 21Ne/22Ne of Figure 8, the three analyses obtained on #45c.29 (two from sample #1 and one from sample #2) show more cosmogenic compositions, with low 20Ne/22Ne (4.5 to 8.3) and high 21Ne/22Ne ratios (0.34 to 0.64), compared to most micrometeorites from TAM and other Antarctic collections. We calculated the abundances of cosmogenic 21Ne and trapped 20Ne, as well as trapped 20Ne/22Ne ratios (Table 3), assuming that neon in #45c.29 samples is a mixture of cosmogenic and trapped Ne. The trapped Ne, in turn, was considered a mixture of solar wind proper (SW; Heber et al., 2012) and fractionated solar wind (FSW; Benkert et al., 1993; Wieler et al., 2007). The cosmogenic 21Ne/22Ne ratio was determined graphically in Figure 8 as 0.94 ± 0.02, namely from the dashed regression line through the #45c.29 samples and considering a cosmogenic 20Ne/22Ne of 0.80 ± 0.03; the latter ratio is typical for cosmogenic production by galactic cosmic rays (GCR) in chondritic meteorites (“GCR chondrites” in Fig. 7; Leya and Masarik, 2009; Trappitsch and Leya, 2013). Even if production of cosmogenic Ne in micrometeoroids is expected to result from solar cosmic radiation (SCR) rather than GCR, the obtained concentrations are robust with respect to the choice of component compositions.

Inferred trapped 20Ne/22Ne ratios (~11.8, Table 3) and 20Ne concentrations (~ 1.4 x 106

cm³ STP/g) are in the upper range of those calculated for TAM micrometeorites (Baecker,

2014). Sample #1 shows the strongest cosmogenic contribution to Ne with 21Necos = (1.62 ± 0.04) x 10-7 cm3 STP/g (Table 3). This concentration is significantly higher than those calculated for almost all Antarctic micrometeorites, the exception being the scoriaceous micrometeorite M240410 (Osawa et al., 2010; Table 4). Sample #2 is less shifted towards the cosmogenic end-member in Figure 8, and the concentration of cosmogenic 21Ne is only about one third that of sample #1 (Table 3).

3.4.3. Argon The concentrations of cosmogenic 38Ar and trapped 36Ar in #45c.29 samples were calculated from their 38Ar/36Ar ratios, assuming a mixture of cosmogenic and trapped Ar, as for Ne. We considered a 38Ar/36Ar ratio of 0.65 in the cosmogenic end-member, as typical for chondritic meteorites (Eugster, 1988), and a trapped ratio of 0.188 (Supplementary Material Figure S2). Contrary to Ne, the calculated value of cosmogenic 38Ar in the samples depends strongly on the choice of trapped and cosmogenic 38Ar/36Ar ratios. Production of 38Ar in chondritic micrometeorites by GCR leads to a cosmogenic end-member with 38Ar/36Ar ratio of ~ 2.35 as predicted by Trappitsch and Leya (2013), while production by SCR for an energy spectrum with a rigidity of 100 MV gives a ratio of ~ 0.91. These two scenarios lead to 38Ar cosmogenic concentrations that are about 5% lower and ~11% higher, respectively, than those given in Table 3 and discussed below. The assumed composition of the end-member components is not critical for sample #2 (with its high 38Ar/36Ar ratio, Supplementary Material Figures S2), but it is for sample #1 as well as for low detections of cosmogenic 38Ar reported in other micrometeorites.

Sample #2 has the highest 38Ar/36Ar ratio (0.35, Table 2 and Supplementary Material Figure S2), which corresponds to a cosmogenic 38Ar concentration of ~7.2 x 10-8 cm3 STP/g (Table 3). This is clearly higher than in other TAM and Antarctic micrometeorites, including the gas-rich scoriaceous micrometeorite M240410 (~4.8 x 10-8 cm3 STP/g; Osawa et al., 2010). The abundance of cosmogenic Ar in sample #2 even exceeds that of cosmogenic 21Ne with a (21Ne/38Ar)cos ratio of ~ 0.74 (Table 3). This low ratio suggests low Mg/Ca ratio in this fragment, as Mg and Ca are the major elements involved in the production of cosmogenic Ne and Ar, respectively. The concentration of cosmogenic 38Ar calculated for sample #1 is only half that in sample #2 (Table 3). The abundances of trapped 36Ar in #45c.29 (5.2 and 3.5 x 107

cm³ STP/g, Table 3) are high compared to other micrometeorites, while concentrations of

40

Ar (2.8 and 4.6 x 10-5 cm³ STP/g) are in the range typical for micrometeorites.

3.4.4. Krypton The concentrations of 84Kr measured in samples #1 and #2 (7.2 and 9.4 x 10-9 cm³ STP/g, Table 2) are amongst the highest measured for TAM micrometeorites (Baecker, 2014). Values of 80Kr/84Kr (0.043 in sample #1 and 0.044 in sample #2) and 86Kr/84Kr ratios (0.332 and 0.318) show possible enrichments compared to common trapped components (e.g., 80

Kr/84Kr = 0.039 and 86Kr/84Kr = 0.310 in the Q (P1) component; Busemann et al., 2000).

The measurement uncertainties are uncomfortably large for these isotopic ratios and do not allow to assess the origin of these possible enrichments.

3.4.5. Xenon Concentrations of 132Xe in #45c.29 (1.69 and 1.29 x 10-9 cm3 STP/g, Table 2) are in the lower range of concentrations reported for TAM micrometeorites (Baecker, 2014). Because of small count rates at the rare 124Xe and 126Xe isotopes, the 124Xe/132Xe and

126

Xe/132Xe ratios have large uncertainties, but clearly are low and do not indicate significant

cosmogenic contributions. Both samples show, however, high abundances of the heavy isotopes 134Xe and 136Xe compared to air and other known trapped components, with 134

Xe/132Xe > 0.40 and 136Xe/132Xe > 0.36 (Table 2). In addition, sample #2 also shows excess

129

Xe from 129I radiogenic decay (129Xe/132Xe = 1.053 ± 0.016, Table 2).

4. Discussion

4.1 Atmospheric processing of #45c.29

4.1.1 Melting and oxidation

The coarse-grained, scoriaceous micrometeorite #45c.29 underwent melting and oxidation during pulse heating resulting in the formation of the magnetite and igneous rims as commonly observed in micrometeorites (Fig. 2, Genge et al., 2005). However, vesicles and glassy pockets also occur in the particle interior, typically around pyroxene and olivine grains (Fig. 1). These pockets differ from the igneous rim because they usually contain plagioclase laths in addition to the common olivine/pyroxene crystallites and small Fe-oxide grains (Fig. 3C). The compositions of the glass in the igneous rim and in the glassy pockets are, however, similar except for the depletion in TiO2 observed in the igneous rim (Fig. 6B). The composition of some glassy pockets requires, in addition, a contribution from olivine or lowCa pyroxene, consistent with the location of these pockets close to olivine and pyroxene grains (Fig. 6D). We thus believe that both igneous rim and glassy pockets are quench products of the melt produced during atmospheric entry heating and thus that melting extended into the inner part of the particle. We do not observe depletion in volatile elements,

such as Na, from the inner glassy pockets to the outer igneous rim (Fig. 6C), which suggests that Na was not severely lost by evaporation during atmospheric entry. As Na volatility decreases with increasing oxygen fugacity (Wulf et al., 1995; Hewins et al., 2005), the glass composition indicates that #45c.29 was partially melted under oxidizing conditions. Melting and crystallization under oxidizing conditions are usually associated with ablation spheres, namely glassy spheres that separate from the fusion crusts of meteorites during ablative flight and that decelerate at relatively low altitude compared to micrometeoroids (Genge and Grady, 1999). Such an origin can, however, be ruled out for #45c.29 since it bears the magnetite rim typical of micrometeorites. Oxidation of #45c.29 is also documented by the occurrence of small and rounded forsterite grains with abundant Fe-oxide intergrowths in the igneous rim (Figs. 2C and D). These grains are similar to natural and experimentally oxidized olivine grains at temperatures above 600 °C (Haggerty and Baker, 1967; Champness, 1970; Goode, 1974) and to some olivine relict grains observed in meteorite fusion crusts (Genge and Grady, 1999), IDPs (Rietmeijer, 1996) and at the periphery of single mineral micrometeorites (Taylor et al., 2012). The small black dots visible near Fe-oxides in BSE images (Fig. 2C) were too small to be analyzed accurately but could correspond to the silica phase formed during olivine oxidation. The oxide grains in the olivines are oriented, suggesting that oxidation occurred at solid-state (topotactic reaction) with Fe-oxides that formed along the planar defects of olivine (Champness, 1970). Hence, rounded olivine grains of the igneous rim are relict grains that did not totally melt, but underwent severe oxidation. The large olivine grain observed on the left-hand side of the section shown in Figure 1 is also oxidized for over 20 µm from the igneous rim, with abundant Fe-oxide intergrowths all elongated in the same direction and enrichment in MgO of the olivine (Figs. 2A and C). More inward in this grain, the abundance of Fe-oxide intergrowths progressively decreases

(although they are still present) and coarse and euhedral Fe-oxide grains rather cluster in wide melt-filled fractures. Adjacent to these fractures, olivine shows cellular shapes (i.e., cuspate contacts with the melt) and it is enriched in MgO over a layer of few microns (Fig. 9). The cellular texture shows that the thin Mg-rich margin grew by reaction between olivine relict grains and the melt that infiltrated into fractures (Walton and Shaw, 2009): disequilibrium between Fe-rich olivine and oxidized melt leads to incongruent dissolution of olivine, oxidation of Fe and crystallization of magnetite, and heterogeneous nucleation and growth of new Mg-rich olivine in equilibrium with the melt. The small euhedral olivine grains observed in the melt-filled fractures have similar Mg-rich composition (Fig. 2F) and some of them are zoned with a less magnesian relict core, which has similar BSE contrast as the large olivine grain away from cuspate margin. This indicates that they progressively formed by the necking off of the reactive cellular margins into the fractures. The reaction between Fe-rich olivine and oxidized melt could be, therefore, an alternative process for the formation of Mg-rich olivine “microphenocrysts” in some micrometeorites, and it would begin as soon as micrometeoroids undergo a small degree of melting and oxidation. This process differs from gradual melting of olivine and subsequent microphenocryst nucleation and growth, which produce the rounded olivine relict grains commonly reported in micrometeorites (Figs. 9A and B). Further olivine/melt reaction leads to the coalescence of cellular margins and to the dark and mottled aspect of olivine like that observed in the olivine relict grain in Figure 2G and close to the large vesicles in Figure 9C. There, the olivine is associated with two large aggregates of magnetite and hercynite with a vermicular rim rich in melt inclusions (Fig. 9D). The association of magnetite and spinel, the Cr-rich composition (Cr2O3 ~ 4 wt%) of magnetite and the disequilibrium rims of the magnetite grains suggest they are relict grains that underwent partial melting during atmospheric entry and not the products of metal or

sulfide oxidation. Partial melting of magnetite implies that the temperature rose up to 1400°C (Presnall, 1995), and combined with the significant oxidation of surrounding olivine, this shows that vesicles act as efficient carriers of heat and oxygen through micrometeoroids during atmospheric entry. The temperature reached in the inner part of the micrometeoroid was also sufficient (> 1300 °C; Presnall, 1995) to trigger the disequilibrium melting of relict pyroxene, as evidenced by the corroded pigeonite nuclei set in vesicular melt with pyroxene composition (Figs. 3G and H). In the pigeonite cores, element ratios that involve iron, which has two different valence states, vary significantly (e.g, Fe/Mg and Fe/Mn) whereas other cation ratios are constant (e.g., Cr/Mg and Ca/Mg ratios, Figs. 10A and B). These chemical trends suggest that partial melting of pigeonite occurred under oxidizing conditions and that the initial pigeonite was probably richer in FeO. This conclusion is supported by the occurrence of tiny Fe-oxide grains in the vesicular melt. The composition of the relict grains can thus be modified during their partial melting and caution should be taken when using mineral chemistry of relict grains in precursor identification studies.

4.1.2 Crystallization of plagioclase- and pyroxene-bearing mesostasis

One striking characteristic of particle #45c.29 is the occurrence of a mesostasis made of plagioclase laths and augite crystallites, with minor interstitial tiny oxide grains and glass (Fig. 3B), and locally, olivine microphenocrysts in the particle interior. Augite also forms large grains together with margins around the vesicular melt formed by pigeonite melting (Fig. 3G and Table 1). Augite partially to totally includes small equant Mg-rich olivine microphenocrysts (Figs. 3D and H), which suggests that it crystallized from the silicate melt during atmospheric entry. The origin of the chemical variations observed in the augite

margins of the large central aggregate (Figs. 3D and G) is not clear, but it could result from the coalescence of zoned microphenocrysts of augite, as suggested by the organization of the variations along streaks perpendicular to the crystal border. Locally, the large augite grains and margins also show a very thin overgrowth that is very bright on BSE images and that highlights a late change in crystallization conditions, such as melt composition or oxygen fugacity. This bright rim shows the same dendritic shape as the pyroxene crystallites of the mesostasis (Fig. 3D). Such mesostasis, rich in plagioclase and pyroxene crystallites and with nearly holocrystalline texture, has never been described before in micrometeorites, even in those that derive from plagioclase-pyroxene rich precursors such as the differentiated cosmic spherules (Cordier et al., 2012). Indeed, the nucleation and growth of pyroxene – and even more of plagioclase - are unexpected during atmospheric entry as they require cooling rates much lower than predicted for micrometeoroids (Taylor and Brownlee, 1991). Several observations attest, however, that the mesostasis of #45c.29 crystallized from the silicate melt during atmospheric entry: (1) as described above, the continuity between the edges of the large augite grains and the small dendritic laths of the mesostasis, which suggests they both formed contemporaneously (Fig. 3D); (2) the gradient observed in the shape of pyroxene laths that range from subhedral in the mesostasis to dendritic close to glassy pockets, igneous rim and magnetite rim (Figs. 3C and 2D); (3) the absence of small (<10µm) vesicles within the mesostasis and their concentration in glassy pockets (Fig. 1) that suggest that when solidification proceeded, bubbles were pushed away from the crystallizing assemblage; (4) the spatial association of mesostasis and large (>50µm) vesicles on the lower edge of the particles that shows they formed at the same time. How augite and plagioclase crystallized in #45c.29 is not completely clear, but we postulate that their crystallization resulted from a combination of factors. As shown in Figure

6D, the glass in #45c.29 has high Ca/Mg, Al/Mg and Na/Mg ratios (Ca/Mg>0.4, Al/Mg>0.9 and Na/Mg>0.05) compared to igneous rims of other micrometeorites (Ca/Mg< 0.2, Al/Mg<0.3 and Na/Mg<0.03; Genge, 2006). The coupled enrichment in refractory (Ca and Al) and volatile (Na) elements shows that the precursor material of the mesostasis had an unusual composition, enriched in plagioclase and Ca-pyroxene, compared to the precursors of other micrometeorites (Fig. 6D). Such an unusual composition, combined with a relatively low cooling rate, may explain the plagioclase and pyroxene crystallization, suggesting unusual entry parameters (i.e., slow velocity/shallow entry angle) of the micrometeoroid. The abundance of plagioclase-pyroxene bearing mesostasis in #45c.29 suggests that the particle interior underwent substantial melting. The external glass-rich rim that we named igneous rim is, therefore, not strictly a rim produced by surface melting and surrounding an unmelted core, as described by Genge (2006). In #45c.29, the difference in texture between particle interior and rim would rather indicate difference in cooling rate and degree of oxidation.

4.2 Which precursor material for micrometeorite #45c.29?

The coarse-grained, scoriaceous micrometeorite #45c.29 shows a combination of petrologic features that are rare, if not unique, amongst micrometeorites and that suggest a parent body that is not frequently sampled by the micrometeorite population so far known. Such features include: the high NiO content of olivine relict grains (NiO = 0.6 to 0.7 wt%; Fig. 10D); the high Cr, Ni and Ti contents of the magnetite relict grains (Figs. 10E and F) and their association with Al-spinel; and the occurrence of a plagioclase-rich mesostasis. Association of Ni-rich olivine with Cr- and Ni-bearing magnetite grains has only been reported in one unmelted micrometeorite before (particle M7, Kurat et al., 1994) and it is

characteristic of CK chondrites. In these chondrites, equilibrated olivine with Fa contents (Fa31±3) close to that of olivine in #45c.29 (Fa25-30) occurs within the matrix, at the border of unequilibrated POP chondrules and in coarse-grained aggregates (Kallemeyn et al., 1991; Tomeoka et al., 2001; Chaumard and Devouard, 2016). Olivine grains in equilibrated CK chondrites have low MnO contents compared to olivine from equilibrated ordinary chondrites and they have remarkably high NiO contents, ranging from a typical value of ~ 0.5 wt% up to 0.8 wt% in Dhofar 015 (Ivanova et al., 2000). Some of them are darkened by curvilinear trails or blebs of Fe-oxide (Rubin, 1992; Tomeoka et al., 2001), such as observed in olivine of #45c.29, away from melt-filled fractures (e.g., Ol1; Fig. 3F). Magnetite grains are abundant in CK chondrites (~ 4 vol.%) and they show exsolution lamellae of ilmenite and aluminous spinel and are rich in Ti, Cr and Ni, such as observed in #45c.29 (Figs. 10E and F; Geiger and Bischoff, 1995; Greenwood et al., 2010; Dunn et al., 2016). Pentlandite, which usually forms coarse rims around magnetite aggregates in CK chondrites, is not observed in #45c.29, but small Ni-rich sulfide blebs (Ni ~ 25wt%; qualitative EDS analysis) occur in inclusions and close to the vermicular magnetite grains (Fig. 3E). This setting could be explained by degassing and oxidation of pentlandite (T ~ 700°C; Greshake et al., 1998) in the precursor at the high temperature attested by the incipient melting of magnetite (> 1400°C; Presnall, 1995). On the other hand, Mg-rich pigeonite such as observed in #45c.29 (Fe/Mg < 0.12) are not reported in CK chondrites, which rather include low-Ca and high-Ca pyroxenes, both with Fe/Mg > 0.2 (Figs. 10A and B). We showed that pigeonite relicts in #45c.29 suffered from severe melting and oxidation during atmospheric entry heating, and that these processes changed the pyroxene composition; in particular, they decreased the Fe/Mg ratio (Figs. 10A and B). We cannot thus exclude that pyroxene composition was different from pigeonite before atmospheric entry. Given that only a dozen analyses of pyroxene – often average

compositions – in CK chondrites have been published (Noguchi, 1993; Kurat et al., 2002), it is also possible that the chemical range of CK pyroxenes shown in Figures 10A and B is not representative. For example, Noguchi (1993) reported, yet in graphical form only, some pyroxene grains in Karoonda (CK4) with compositions (Fs<10Wo8, CaO up to 10 wt% and Al2O3 = 2 wt%; see his Figures 2b and 11) that resemble pigeonite in #45c.29 (Fs2-10 Wo7-11, CaO = 3.5 to 6 wt% and Al2O3 = 1.5 to 2 wt%; inset of Fig. 10A). The coarse-grained, scoriaceous micrometeorite #45c.29 could thus derive from a lithology not frequently sampled by CK chondrites. Abundant coarse-grained plagioclase with variable composition (An17 to An100) and minor high-Ca pyroxene (Fs4-11Wo42-53) usually occur in equilibrated CK chondrites (Noguchi, 1993). Melting of these mineral phases during atmospheric entry could explain the high Ca, Al and Na content of the glass and mesostasis of #45c.29 (Fig. 6D). In turn, this suggests that #45c.29 derives from an equilibrated CK chondrite (CK4 to CK6) source, as unequilibrated CK chondrites are rich in phyllosilicates (~ 65%) and olivine (5%) but poor in plagioclase and high-Ca pyroxene (Martin et al., 2013). An equilibrated precursor is also consistent with the homogeneous, equilibrated composition of olivine and the low TiO2/Cr2O3 ratio of the magnetite grains in #45c.29 (Fig. 10E). Parentage with thermally-metamorphosed carbon-poor chondritic material is also consistent with the coarse grain size of the olivine relict grains and with the mineralogical evidence for oxidation of #45c.29, as carbon pyrolysis would have maintained reducing conditions in the melt formed during atmospheric entry (Brownlee et al., 1983). In conclusion, particle #45c.29 most likely derives from equilibrated CK chondrites, in which equilibrated Ni-rich olivine is associated with pyroxene, magnetite, and plagioclase. This is consistent with the oxygen isotope composition obtained on one sample of #45c.29 that falls within the field of CK/CV chondrites in Figure 7.

4.3 The message from noble gases

4.3.1 Trapped noble gases

The trapped noble gas inventories of the two individual samples of #45c.29 are rather similar: concentrations of trapped 20Ne, 36Ar, 84Kr and 132Xe differ by less than 50%. This similarity extends also to the isotopic composition of the heavy gases Kr and Xe, which differs from the pattern generally observed in other unmelted and scoriaceous micrometeorites from the TAM collection (Baecker, 2014) and from other Antarctic collections (Osawa et al., 2000; Osawa and Nagao, 2002a and 2002b; Osawa et al., 2003a and 2003b; Osawa et al., 2010). The most striking observation is the enrichment in #45c.29 of the heavy 134Xe and 136Xe isotopes compared to air and/or other well-characterized trapped components, such as solar wind and meteoritic Q-gas (Fig. 11). In contrast, most of the other micrometeorites (for which data of sufficient precision exist) generally plot close to the Q component composition, more or less like the classical “AVCC” (“average carbonaceous chondrites Xe”; Ott, 2014). The position of #45c.29 samples in the 134Xe/132Xe and, in particular, 130Xe/132Xe versus 136Xe/132Xe diagrams of Figure 11 suggests that the enrichment in heavy Xe results from mass fractionation of trapped fractionated air. Presence of fractionated air Xe is further supported by the 84Kr/132Xe ratios of samples #1 and #2 (~4.3 and ~7.3, Table 2), which are much higher than in typical meteoritic trapped components (e.g., 84Kr/132Xe < 1 for meteoritic Q-gas; Busemann et al., 2000). Terrestrial atmosphere has a much higher ratio (84Kr/132Xe = 27.8; Ozima and Podosek, 2002) and high Kr/Xe is thus generally regarded a sign of contamination by air (Schelhaas et al., 1990; Scherer et al., 1994; Schwenzer et al., 2012). Atmospheric contamination is also supported by the petrographic

evidence of oxidation (Section 4.1.1) and by the oxygen isotopic composition of #45c.29 (Supplementary Material Figure S2). Sample #2 may still contain trapped meteoritic Xe, as suggested by its slightly low 136Xe/132Xe and high 130Xe/132Xe ratios compared to sample #1 (Fig. 11), and by the apparent excess of 129Xe from extinct 129I. The trapped 20Ne/22Ne ratio inferred from the regression line in Figure 8 (20Ne/22Ne ~ 11.8; Table 3) suggests a predominantly fractionated solar wind origin for the trapped Ne component, with a concentration at the high end among the TAM micrometeorite collection (Baecker, 2014). Trapped 36Ar is also high (> 3.5 x 10-7 cm³ STP/g) compared to other TAM micrometeorites, and it is associated with lower-than-air 40Ar/36Ar (Supplementary Material Figure S2), which shows that most of 36Ar is extraterrestrial. The ratio of trapped 20Ne to 36Ar of ~ 3 to 4 is ten times lower than in solar wind (20Ne/36Ar = 42.1; Heber et al., 2009). This could indicate that only 10% of trapped 36Ar is of solar wind origin and the remaining 90% (~ 4 x 10-7 cm³ STP/g) is “planetary” Ar. This does not match, however, with the measured concentration of 132Xe which is ~ 3 times lower than predicted (5 x 10-9 cm³ STP/g) by assuming the 36Ar/132Xe ratio of the dominant trapped planetary component in primitive meteorites (Q (P1) with 36Ar/132Xe ~ 80; Busemann et al., 2000). This leads us to consider that trapped Ne has been lost preferentially compared to trapped Ar during (or prior to) atmospheric entry, a possibility that we also suggest for cosmogenic Ne compared to cosmogenic Ar (Section 4.3.2). Our conclusions that Xe (and Kr) is mostly trapped fractionated air and could be fractionated from 36Ar during atmospheric entry make impossible further estimation of the original inventory of extraterrestrial Ar, Kr, and Xe and limit comparison with known meteorite types. We note, nevertheless, that low concentrations of primordial (i.e. non-solar wind) trapped gases are consistent with an origin from an equilibrated type of meteorite parent body, as indicated by petrographic data (Section 4.2).

4.3.2 Cosmogenic abundances and ratios in coarse-grained micrometeorites

The two fragments of #45c.29 analyzed for noble gases show distinct cosmogenic compositions for Ne and Ar. Sample #1 has high 21Necos and its cosmogenic 21Ne/38Ar ratio (~ 4.6, Table 3) is in the typical range of chondritic meteorites (~5 to 10; Supplementary Material Table S3). Sample #2 has extremely high 38Arcos and a low (21Ne/38Ar)cos ratio of 0.74 compared to values measured in sample #1 and in chondritic matter (Supplementary Material Table S3). The two samples must, therefore, have had distinct chemical compositions, as the major target elements are different for spallogenic Ne (Mg) and spallogenic Ar (Ca). This chemical heterogeneity between the two samples is consistent with the occurrence of domains with distinct mineralogy in the fragment shown in Figure 1, which is twice as large as the samples analyzed for noble gases. Since sample #1 has a typical chondritic (21Ne/38Ar)cos ratio, its mineralogy is probably dominated by Mg-rich minerals like in most micrometeorites. It would be equivalent to the large olivine relict grain observed on the left-hand side of the particle in Figure 1. On the other hand, the very low (21Ne/38Ar)cos ratio of sample #2 suggests it contains a Ca-rich component. By analogy with the fragment shown in Figure 1, sample #2 is probably mainly made of pyroxene-plagioclase mesostasis (Table 2) that has inherited its cosmogenic signature from the Ca-rich phases of the CK chondrite precursor. The chemical heterogeneity of the two samples is challenging as the chemical composition controls the cosmogenic production rates (Supplementary Material Table S3). These rates are needed in combination with the concentrations of cosmogenic products to make use of the Poynting-Robertson effect and infer the place of origin in the Solar System (distance from the Sun) and the travel time of the micrometeoroid (e.g., Osawa and Nagao,

2002a; Trappitsch and Leya, 2013). We tried to solve the problem by modeling mixtures of Fa26 olivine for the Mg-rich mineral phase and mesostasis for the Ca-rich component. For this purpose, we used production rates (Supplementary Material Table S3) calculated from the spreadsheets of Trappitsch and Leya (2013). In applying the Poynting-Robertson effect, we have to estimate the size and density of #45c.29 before its atmospheric entry. We consider two extreme cases in Supplementary Material Figure S4. In panels A and B, we assume that #45c.29 did not suffer from ablation and that its precursor had the same diameter (~ 650 µm, i.e., the average diameter of #45c.29) and density (2.37 g/cm³, calculated from the dimensions and weights of the noble gas samples) as the micrometeorite. In Figures S4C and D, we consider an alternative case, with severe ablation and loss of 75 % of the initial mass (Rudraswami et al., 2016). This changes only slightly the travel time and inferred place of origin but, otherwise, it does not affect our conclusions. With no mass loss assumed, abundances of cosmogenic 21Ne and 38Ar in sample #1 can be both reproduced if i) its precursor was made of 80 wt. % of Fa26 olivine and 20 wt. % of Ca-rich mesostasis and ii) was exposed to cosmic ray as micrometeoroid during ~145 Ma. Such an exposure corresponds to a place of origin at ~16.6 AU. We then tried to match the 21

Ne and 38Ar of sample #2, using the same cosmic ray exposure (CRE) age and place of

origin, but with a different proportion of olivine and mesostasis. For cosmogenic 38Ar, this can be done with an olivine/mesostasis ratio of 35/65, but there is no solution for 21Ne. A possible explanation is that the mineralogy of the two samples is more complex than assumed in our simple picture. We also tested the use of pure forsterite for the Mg-rich end-member and Ca-Al-rich inclusion (CAI) or eucrite compositions for the Ca-rich component, but results are qualitatively similar. Another possibility might be the loss of cosmogenic Ne (but not Ar), as already invoked by Osawa et al. (2010) to explain the longer CRE ages they calculated

from Ar compared to Ne. Experiments such as those performed by Füri et al. (2013) on millimeter-sized CI-like particles also suggest severe loss of cosmogenic Ne during atmospheric entry, even when the heating pulse is short. Unfortunately, Füri et al. (2013) did not give information on the behavior of cosmogenic 38Ar. Mesostasis can confidently be assumed to be less retentive for cosmogenic Ne than olivine relict grains, and about 60% loss from mesostasis – which primarily affects sample #2 - gives a common (but not perfect) solution to 21Ne as well as 38Ar in both samples. The inferred CRE age is ~ 140 Ma and place of origin ~ 16 AU (Table 4 and Supplementary Material Figs. S4A and B). Comparable CRE age and place of origin are obtained when ablation of 75% of the initial mass is considered (Supplementary Material Figs. S4C and D). The common solution for samples #1 and #2 indicates that #45c.29 micrometeoroid would have been released from its parent body well beyond the orbit of Saturn. This may be taken to indicate a cometary origin, but significant dust production by comets starts only when they enter the inner Solar System. We will show in Section 4.3.3 that the (21Ne/22Ne)cos ratio of #45c.29 is also inconsistent with a long irradiation as a micrometeoroid and rather points to a more complex irradiation history. In any case, the contrasting abundances of cosmogenic 21Ne and 38Ar in samples #1 and #2 show that cosmogenic noble gases measured on fragments of coarse-grained, partially melted micrometeorites can lead to wrongly assigned cosmogenic production rates, and thus to incorrect exposure ages and places of origin. First, mineral phases can be heterogeneously distributed within coarse-grained micrometeorites and because each of them has distinct abundances in cosmogenic noble gases, the measured composition can be completely different from that describing the respective bulk micrometeorite. Then, the thermal effect of atmospheric entry is also not homogeneously distributed throughout the coarse-grained,

partially melted particles, with variations of the relative abundance of relict grains versus melted mesostasis and thus of the loss of cosmogenic noble gases.

4.3.3 Pre-irradiation scenario

The cosmogenic dataset obtained on #45c.29 is unique as we obtained three measures of the Ne isotope ratios on two distinct samples. The array of the three data points in Figure 8 corresponds to a mixing line between the trapped and cosmogenic components and allows to estimate the (21Ne/22Ne)cos ratio in #45c.29 fairly accurately to 0.94 ± 0.02 (see Section 3.4.2). This value can be compared with the (21Ne/22Ne)cos ratios predicted for irradiation of micrometeoroids during space exposure (Trappitsch and Leya, 2013). Results of the comparison are shown in Supplementary Material Figure S5 for different release distances of the micrometeoroids and for various mixtures of olivine and mesostasis. In all cases, the predicted (21Ne/22Ne)cos ratio is far below (< 0.83) the observed value of ~ 0.94 and this difference in Ne isotope ratio cannot be accounted for by the loss of cosmogenic Ne as invoked for the abundances of 21Ne and 38Ar. It rather suggests that the irradiation story was more complex. Given the igneous and magnetite rims, there is no doubt that #45c.29 was a micrometeoroid at the time it encountered the Earth’s atmosphere. The isotopic composition of cosmogenic Ne suggests, however, that most of the cosmogenic production happened while #45c.29 was still part of a larger body. Indeed, in micrometeoroids, there is no shielding and cosmogenic Ne is produced by primary galactic and solar cosmic rays. At depth in meteoroids or asteroids, on the other hand, cosmogenic production by secondary particles produced by primary galactic cosmic rays is efficient, while contributions from solar cosmic rays become negligible (Leya and Masarik, 2009). The cosmogenic Ne isotope ratio may thus tell us about the actual setting of the spallation, namely

if it occurred during the travel of the micrometeoroid or before, during pre-irradiation in a larger body. We explore two different settings for pre-irradiation of #45c.29: in a meteoroid (Supplementary Material Fig. S6) and close to the surface of an asteroid (Fig. 12 and Supplementary Material Fig. S7). In doing so, we neglected the contribution from the irradiation during the micrometeoroid travel, which we calculated as realistically less than 10% (for an origin at 2.5 AU and at least 20% of mesostasis in sample #1). For pre-irradiation in carbonaceous chondrite meteoroids, we used the model of Leya and Masarik (2009) and meteoroid radii of 20 and 65 cm. For pre-irradiation in an asteroid, we considered production at different depths (0 to 120 cm) using the 2π model of Leya et al. (2001) for cosmogenic Ne isotope ratios, and, since these authors did not give information on 38Ar production from Ca, the model of Hohenberg et al. (1978) for the cosmogenic 21Ne/38Ar ratio. For sample #1, which was probably dominated by olivine and for which data do not point toward important loss of cosmogenic Ne during atmospheric entry, both pre-irradiation in meteoroid and at shallow depth (~10 cm) in an asteroid reproduce the cosmogenic 21

Ne/38Ar and 21Ne/22Ne ratio for a mesostasis fraction of ~30% (Fig. 12 and Supplementary

Material Figs. S6 and S7). For sample #2, pre-irradiation in the asteroid also reproduces – within uncertainties – the 21Ne/38Ar and 21Ne/22Ne ratios considering a mesostasis fraction of ~ 0.8 and taking into account 60% loss of cosmogenic Ne from the mesostasis (Fig. 12). Preirradiation in a meteoroid, on the other hand, reproduces the cosmogenic Ne/Ar ratio if similar loss is considered, but it does not lead to a match for the 21Ne/22Ne ratio (Fig. S6). These are not unique solutions, however, and other combinations of loss, depth and mesostasis fractions may be possible. We prefer the scenario of pre-irradiation near the surface of an asteroid as it better fits the data and it is the simplest scenario: cosmogenic Ne and Ar are produced close to the

surface of the parent body, whereas the trapped solar wind signature is acquired during the travel time of the micrometeoroid from its parent asteroid to the Earth. If irradiation occurred in a meteoroid, on the other hand, this one must have disintegrated at some time before entering Earth's atmosphere in order to allow the micrometeoroid to collect solar wind.

4.3.4 Origin of micrometeoroids with long CRE age

In addition to #45c.29, only five other particles amongst the 246 Antarctic micrometeorites analyzed for noble gases so far have significant concentrations of 21Necos, and therefore long inferred CRE ages and large source distances (Table 4). Other micrometeorites have undetectable or low contents of 21Necos and accordingly low nominal CRE ages (< 1 Ma; Osawa et al., 2000; Osawa and Nagao, 2002a, 2002b; Osawa et al., 2003a and b; Osawa et al., 2010; Baecker, 2014). It has been proposed that micrometeorites with significant concentration of 21Necos may sample sources located at heliocentric distances larger than 30 AU, i.e. Kuiper Belt objects (Osawa et al., 2000; Osawa and Nagao, 2002a; Osawa et al., 2010; Table 4). However, numerical simulations have shown that only 20% of the Kuiper Belt dust feed the inner solar system – the other 80 % being trapped or ejected from the solar system by giant planets - and that most of the dust is then destroyed by mutual collisions in the inner solar system (Liou et al., 1996). The combined study of petrologic and noble gas features of the coarse-grained, scoriaceous micrometeorite #45c.29 shows that the high 21Necos and long apparent CRE age (~140 Ma, Table 4) can rather result from preexposure to cosmic rays, possibly close to the surface (depth on the order of 10 cm) of a parent asteroid. Such a complex irradiation scenario can possibly also explain the long CRE age estimated for some other micrometeorites; however, this cannot be verified for the other

five micrometeorites with long exposure age, since we do not have an estimate of the cosmogenic end-member 21Ne/22Ne ratio as good as we have for #45c.29. If the assumption that high 21Necos is indicative of pre-irradiation close to the surface of asteroids or in meteoroids, only a small fraction of micrometeorites (6 out of 246, i.e. 2.5 %) appear to have recorded such process. One possible explanation is that micrometeorites with low 21Necos suffered from significant degassing and loss of cosmogenic Ne during atmospheric entry (Füri et al., 2013). In our scenario, 60% of cosmogenic Ne was lost from mesostasis during atmospheric entry, but #45c.29 still contains large amounts of 21Necos (Table 3). This suggests that micrometeorites with low 21Necos may be those that virtually lost all their Ne during atmospheric entry. This is likely for cosmic spherules but even so, the proportion of particles with high 21Necos does not exceed 5 % of the total number (n=111) of less heated, i.e. unmelted and partially melted, micrometeorites analyzed for noble gases. This proportion is still far less than the proportion of micrometeorites of the same size range (50 – 2300 µm) that have oxygen isotope compositions that point toward ordinary chondrite parentage and origin in the inner asteroid belt at ~ 2.5 AU (~20 %; Cordier and Folco, 2014). An alternative possibility is that most micrometeorites with low 21Necos were buried deeply (> 1 to 2 m) within their asteroid parent before being ejected in space as dust particles during hypervelocity impacts. This conclusion is supported by Astronomical observations (Infrared Astronomical Satellite, IRAS) that identified bands of dust formed by mutual collisions and catastrophic disruption of Main Belt asteroids and that provided evidence that these bands replenish the zodiacal cloud around 2 A.U (e.g., Sykes, 1990; Dermott et al., 1984, 2002). On the other hand, micrometeorites (such as apparently #45c.29) that are ejected by meteoroid disruption or by surface erosion of asteroids, for example an increase in the asteroid rotation rate and in the dust fall off, appear to be uncommon.

5. Conclusions

We report on the discovery of a new type of coarse-grained, scoriaceous micrometeorite made of Fa26 olivine, Fs5 Wo10 pigeonite and magnetite relict grains set in a fine-grained mesostasis of calcic plagioclase, augite, Fe-oxide and glass. This particle is related to equilibrated CK chondrites based on its oxygen isotopic composition and on the high Ni contents of the olivine and magnetite relict grains. The occurrence of plagioclase and pyroxene microphenocrysts in #45c.29 is probably due to a combination of factors, such as the composition of the precursor material (Ca-, Al- and Na-rich) and an unusual, relatively low, cooling rate of the particle during atmospheric entry. The latter may be related to an uncommon trajectory of the #45c.29 micrometeoroid during atmospheric entry, as also indicated by the severe oxidation of the whole particle. For example, olivine relicts were transformed, at the solid state, into forsterite and Fe-oxide intergrowths in and close to the igneous rim. More inward in the particle, the oxidation progressed through the melt, which reacted with olivine relict grains to form Mg-rich margins, equant olivine microphenocrysts and small Fe-oxide crystals. These features show that significant oxidation is not restricted to ablation spherules. In terms of noble gases, the coarse-grained, scoriaceous micrometeorite #45c.29 is unique and stands out from all other micrometeorites. Solar-wind implanted Ne is present, with a concentration on the high end relative to other micrometeorites from the Transantarctic Mountain sample suite. Trapped Ar is also present and it is of solar-wind and/or primordial origin. Xenon may be largely trapped mass-fractionated air collected during the passage through Earth’s atmosphere, as in some other scoriaceous micrometeorites. These features do not allow identification of a clear connection to known meteorite types, but are consistent with the suggested origin from an equilibrated chondritic precursor.

Concentrations of cosmogenic noble gases are remarkably high and the abundance ratio (21Ne/38Ar)cos indicates extremely different Mg/Ca ratio in the two analyzed samples of #45c.29. Cosmogenic noble gas abundances on fragments of coarse-grained, partially melted micrometeorites can thus lead to wrongly assigned cosmogenic production rates, and thus to incorrect exposure ages and inferred places of origin. For example, if the cosmogenic 21Ne and 38Ar concentrations of #45c.29 are interpreted in terms of irradiation as a micrometeoroid particle, an origin from beyond Saturn is required (~16 AU). On the other hand, analysis of noble gases on different fragments allows the cosmogenic isotope ratios to be estimated more accurately. In the case of #45c.29, the high cosmogenic 21Ne/22Ne ratio primarily records an early irradiation stage, in a meteoroid or close to the surface (~10 to 20 cm in depth) of an asteroid, before the release of the micrometeoroid. In this respect, #45c.29 stands out from most other unmelted and scoriaceous micrometeorites of the same size range since it is expected that a significant part of the large (> 50 µm) dust captured by the Earth’s atmosphere is debris of catastrophic collisions and disruptions of Main Belt asteroids.

Acknowledgments: This work was supported by the Italian Ministero della Istruzione, Università e Ricerca (MIUER) through the Programma Nazionale delle Ricerche in Antartide (PNRA) "Meteoriti Antartiche" project (grant ID: PNRA1600029) and the Progetti di Rilevenza Nazionale "Cosmic Dust" project (grant ID: PRIN201520158W4JZ7). Carole Cordier is part of Labex OSUG@2020 (ANR10 LABX56). Bastian Baecker was supported by DFG grants OT 171/5-1 and 5-2. Authors are grateful to Andreas Pack (Georg-August-Universität Göttingen, Germany) for measuring the oxygen isotope composition of the particle and to Valentina Batanova (ISTerre, France), John Spratt (Natural History Museum, London, UK) and Andrea Cavallo (Certema, Italy) for their help during EPMA analytical sessions. Support

by the Fundation Klaus Tschira Stiftung is acknowledged. The constructive comments of Matthew J. Genge and two anonymous reviewers have helped to improve the manuscript. We also thank Alexander N. Krot for his editorial handling.

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Figure captions: Figure 1. Backscattered electron image of the coarse-grained, scoriaceous micrometeorite #45c.29 (A) and interpretative sketch (B). Inset in (A) is the backscattered electron image of the whole micrometeorite.

Figure 2. Backscattered electron images of the igneous rim in the coarse-grained, scoriaceous micrometeorite #45c.29. (A) Large olivine relict grain (Ol3, Fig. 1B) in contact with the igneous rim. (B) Close-up of the igneous rim. (C) Oxidation of olivine relict grain (Ol3) in contact with igneous rim. Inset: Close-up view of oxide inclusions in olivine, systematically associated with small black domains. Note the difference of crystallite abundance in the igneous rim between (B) and (C). (D) Oxidized olivine relict in the igneous rim. (E) Cuspate margins in olivine relict grain (Ol3) near fractures filled with glass and Feoxides. Olivine is decorated by fine trails of Fe-oxide grains and is rimmed by Mg-rich oxidefree cuspate margins. (F) Close-up view of the cuspate margins and melt-filled fractures. Mgrich margins are progressively necked off into the melt (black circles), giving rise to olivine “microphenocrysts” (arrowed). (G) Mottled aspect of olivine relict grain formed by the coalescence of cuspate margins. Abbreviations: Ol: olivine; Px: pyroxene; Ox: Fe-oxide; Fo: forsterite; MR: magnetite rim; IR: igneous rim. Labels of the crystal are given in Fig. 1. Scale bars are in µm.

Figure 3. Backscattered electron images of the interior of the coarse-grained, scoriaceous micrometeorite #45c.29 made of olivine and pyroxene grains in a fine-grained mesostasis of plagioclase and pyroxene laths that define domains with distinct orientations (A). (B) Fine-grained mesostasis with euhedral plagioclase laths and interstitial pyroxene, Fe-

oxide and glass. (C) Glassy pockets near olivine relict grain with plagioclase laths, pyroxene crystallites, small euhedral olivine and rare Fe-oxide grains. Plagioclase and pyroxene laths lose their preferential orientation in the glass pockets. (D) Augite grains with complex compositional patterns and olivine inclusion (outlined in yellow) in the Fe-rich pyroxene rim. This rim is in textural continuity with pyroxene crystallites of the mesostasis (arrowed). Note the small, equant Mg-rich olivine grains lining the Fe-rich olivine relict. (E) Vermicular oxide aggregate (magnetite + hercynite) with small blebs of FeNi sulfide. (F) Olivine relict grains decorated with trails of tiny Fe-oxides that are crosscut by melt-filled fractures. (G) Two nuclei of pigeonite relicts separated by vesicular glass with pyroxene composition, and surrounded by augite rims. The shape of the original pigeonite grain deduced from chemical maps is outlined (dotted line). (H) Close-up of the pigeonite nuclei and vesicular glass. Small Mg-rich olivine grains (outlined in yellow) occur in the augite rim. Abbreviations: GP: glass pockets; Ol: olivine; Px: pyroxene; Aug: augite; Pgt: pigeonite; Pl: plagioclase; Gl: glass; Labels of the crystal are given in Fig. 1. Scales are given in µm.

Figure 4. X-Ray WDS map of a portion of the coarse-grained, scoriaceous micrometeorite. The vesicles have been colored in black to avoid misleading interpretation. Pyroxene grains are identified from the Al distribution map and their contours are reported in all maps.

Figure 5. X-Ray WDS map of the pyroxene aggregate with pigeonite nuclei surrounded by augite rims. The vesicles and fractures have been colored in black to avoid misleading interpretation.

Figure 6. Compositions of igneous-rim glass, glassy pockets and mesostasis in the coarse-grained, scoriaceous micrometeorite #45c.29. Bulk EPM analyses are given in Table 1 and Supplementary Material Table S1. In (D), these compositions are compared with igneous rims of other micrometeorites (IR; Genge, 2006), CK chondrites (CK; Greenwood et al., 2010) and with mixing products of mineral phases in CK chondrites (An75 plagioclase, Fa31 olivine, Fs26 Wo2 low-Ca pyroxene, Fs10Wo44 high-Ca pyroxene and magnetite; Noguchi, 1993, Chaumard et al., 2014; Chaumard and Devouard, 2016). The melting of An75 plagioclase (Pl), high-Ca pyroxene (High-Ca px) and magnetite (Mag) in 57:38:5 proportions can reproduce the composition of the mesostasis.

Figure 7. Oxygen isotope composition of the analyzed sample of #45c.29. (A) Comparison to chondrites in the δ17OSMOW vs δ18OSMOW (‰) diagram (fields drawn after Franchi, 2008). The analytical uncertainty is contained in the symbol size. The black solid line labelled TFL is the Terrestrial Fractionation Line (Clayton, 1993), the dashed line labelled CCAM is the Carbonaceous Chondrite Anhydrous Mineral Line (Clayton et al., 1977), and the dotted line labelled PCM is the Primitive Chondrule Mineral Line (Ushikubo et al., 2012). Inset: photo of the analyzed fragment of #45c.29 (weight: 11 µg). IR: igneous rim. (B) Estimation of the contribution of atmospheric oxygen to the oxygen isotope composition of the sample of #45c.29 analyzed for oxygen isotopes, in the ∆17O vs δ18OSMOW (‰) diagram. The composition of Earth stratosphere is from Thiemens et al. (1995), composition of CK chondrites and their constituents are from Clayton and Mayeda (1999), Greenwood et al. (2010), Ivanova et al. (2000), Nakamura et al. (2000), Russell et al. (2004) and Zipfel et al. (2000). Mass-dependent fractionation during atmospheric entry (evaporation, loss of Fe-Ni metal beads) would shift the isotope composition to the right side of the diagram (gray arrow).

Figure 8. Three-isotope plot of 20Ne/22Ne vs. 21Ne / 22Ne in the coarse-grained, scoriaceous micrometeorite #45c.29, compared with other unmelted (Un.) and scoriaceous (Sc.) micrometeorites from the Transantarctic Mountain collection (TAM MMs: Baecker, 2014) and from other Antarctic collections (Osawa et al., 2002a; Osawa et al., 2010). Individual releases are shown for sample #1 and the long-dashed red line is a fit (Isoplot model 1; Ludwig, 2008) through the three data points. Other micrometeorites with high 21

Necos are labeled. Error bars are not shown when included in the symbol size. Trapped

components plotted for reference are SW (Heber et al., 2012), Ne-B (Black, 1972), FSW (Benkert et al., 1993; Wieler et al., 2007), Ne-HL (Huss and Lewis, 1994), and Earth atmosphere EA (Eberhardt et al., 1965). Predicted compositions are shown for Ne cosmogenic production in micrometeoroids made of Fa26 olivine or mesostasis by solar (“SCR MM”; with rigidities of 75, 100 and 125 MV) and galactic cosmic rays (“GCR MM”; Trappitsch and Leya, 2013 with updates, primarily for 20Ne, provided by R. Trappitsch; see Supplementary Material Table S3). The range of GCR production in chondritic meteorites (“GCR chondrites”) is based on Leya and Masarik (2009).

Figure 9. Backscattered electron images of (A) margins in an oxide-bearing Mg-rich olivine grain in the coarse-grained, scoriaceous micrometeorite #45c.29 (Ol3, Fig. 1B): the contact between olivine core and its margin is not rounded but cuspate; (B) contact between rounded relict grain and skeletal margin in a porphyritic cosmic spherule from the Transantarctic Mountain micrometeorite collection; (C) mottled olivine and vermicular magnetite+spinel association near the lower edge of the fragment (vesicle plane); (D) closeup view of the vermicular magnetite (light gray) + hercynite (dark gray) association.

Figure 10. Comparison of mineral chemistry in the coarse-grained, scoriaceous micrometeorite #45c.29 and CK chondrites. (A-B) Pyroxene. Inset in (A) shows the pyroxene quadrilateral diagram, with pyroxenes in CK chondrites (gray diamonds) reported in Figure 2 of Noguchi (1993), together with relict pigeonite in #45c.29. (C-D) Olivine. (E-F) Magnetite. Mineral compositions in CK chondrites are from Noguchi (1993), Geiger and Bischoff (1995), Tomeoka et al. (2001), Chaumard et al. (2014), Chaumard and Devouard (2016), and Dunn et al. (2014). Ranges of NiO and FeO in olivine and magnetite grains in the CK3 chondrites Dhofar 015 are based on information given in abstract form by Ivanova et al. (2000). Composition of relict grains in micrometeorites are from Steele (1992), Kurat et al. (1994), Beckerling and Bischoff (1995), Genge et al. (2005), Gounelle et al. (2005), Cordier et al. (2011), van Ginneken et al. (2012), Imae et al. (2013), Rudraswami et al. (2015, 2016).

Figure 11. Three-isotope plots 134Xe/132Xe (A) and 130Xe/132Xe (B) versus 136Xe/132Xe in the two analyzed samples of the coarse-grained, scoriaceous micrometeorite #45c.29. For sample #1 the two individual gas release steps are shown in addition to the total. Data for other unmelted (Un.) and scoriaceous (Sc.) micrometeorites from the Transantarctic Mountains (Baecker, 2014) are included for comparison. “Trapped” compositions Q (Busemann et al., 2000) and Earth’s atmosphere (EA; Ozima and Podosek, 2002) are shown for reference. The mixing dashed lines show the effects of addition of nucleosynthetic Xe-HL (Huss and Lewis, 1994) and addition of xenon produced by fission of 244Pu (Ozima and Podosek, 2002). The solid lines show the effect of mass-dependent fractionation (favoring the heavy isotopes) acting on Q and EA trapped components.

Figure 12. Predicted cosmogenic (21Ne/22Ne)cos and (21Ne/38Ar)cos ratios in 2π geometry by GCR are compared with the range of these ratios in samples #1 and #2 of the

coarse-grained, scoriaceous micrometeorite #45c.29 (arrows). Predicted ratios are shown for depths of 10 cm and 50 cm. The dashed curves with large symbols pertain to different fractions of mesostasis (fmeso) in a Fa26-olivine/mesostasis mixture. The dotted lines with the small open symbols show what is expected if predictions for a depth of 10 cm are coupled with 60% loss of cosmogenic Ne from mesostasis (but no loss of cosmogenic Ar).

Al2O3 (wt%)

Na2O + K2O (wt%)

21

17

13 9 47.5

1.5 C

Glass - igneous rim Glassy pockets Mesostasis

A

48.5

49.5 SiO2 (wt%)

50.5

1.3 1.1 0.9 0.7 47.5

51.5

3.0

0.9 B

2.5 0.7

48.5

49.5 SiO2 (wt%)

50.5

51.5

toward An75

D Mag

+10% Mag

60% Pl

(Al/Mg)at

TiO2 (wt%)

2.0

0.5

1.5 1.0

Ol and low-Ca px IR

0.5 0.3 47.5

CK

48.5

49.5 SiO2 (wt%)

50.5

51.5

0.0 0

0.5

High-Ca px

1.0 1.5 (Ca/Mg)at

2.0

2.5

10

δ17OSMOW (‰) M

M A C P CC

A

CI

R

5

CR clan LLL H EH,EL

0

CM TFL

IR CO

-5

#45c.29

CV+CK

100 µm

PC

#45c.29

-2 -3

1 3% air O

Mix

8 7 % or O s cur e r p li ne g n i

20

Earth stratosphere

mass-dependent fractionation

-4

CK chondrites Chondrules in CK Olivine in CK Magnetite in CK Plagioclase in CK

-5 -6 -5

15

AM

5 10 18 δ OSMOW (‰) M

0

CC

-10 -5 Δ17O 0 B -1

0

5 10 18 δ OSMOW (‰)

15

20

25

14 12

SW B

20

Ne/22Ne

FSW

10 8 6 4 2 0 0.0

EA HL

#2 #1(1)

Sc. MM M240410 #45b.10 #45b.17 #45b.08

0.2

0.4 0.6 21 Ne/22Ne

#45c 29 TAM Un. MMs TAM Sc. MMs Antarctic MMs Trapped components GCR chondrites GCR MM (Fa26) GCR MM (Meso) SCR MM (Fa26) SCR MM (Meso)

#1(2)

0.8

1.0

Pyroxene 100

0.03

A

#45c.29 Pigeonite Augite CK chondrites Low-Ca px High-Ca px Px relicts in MM

B

80

Cr/Mg

Fe/Mn

0.02 60

40

0.01 20

Oxidation & Fe-depletion

Wo

0.00

0

0.0

0.1

0.2 Fe/Mg

0.3

0.0

En

0.1

Fs

0.2

0.3

0.4

Fe/Mg

Olivine 1.0

C 0.8

0.6

D

0.4

Dhofar 015

0.8

NiO (wt%)

MnO (wt%)

1.0

CC CK UOC EOC #45c.29 Ol relicts in MM

0.6

0.4

0.2

0.2

0.0

0.0 0

1.0

10

30

20 FeO (wt.%)

40

0

Magnetite 1.0

E

20 FeO (wt.%)

30

40

Field for CC CK #45c.29

F

CK3 TiO2 (wt%)

CK3

TiO2 (wt%)

10

0.5

0.5

CK4-6

CK4-6 0

0

2

4 Cr2O3 (wt%)

6

0

0

0.2

0.4 0.6 NiO (wt%) Dhofar 015

0.8

1.0

0.44

(A)

Xe/132Xe

0.42 0.40

134

0.38

#2

EA Q

#45c 29 steps #45c 29 totals TAM Un. MMs TAM Sc. MMs Trapped components Fractionation To Xe-HL To Pu fission

0.36

Xe/132Xe

0.17

130

0.16 0.15

#1

(B) Q

EA

0.14 0.13 0.32

0.34 136

0.36 Xe/132Xe

0.38

100

(21Ne/22Ne)cos

(21Ne/38Ar)cos

1.0

10 #1

#45c.29: 0.94 ± 0.02

1

0.9

#2

0.0 21

0.95

0.2

0.4 0.6 f-mesostasis

Ne/38Ar, 10 cm 21 Ne/38Ar, 10 cm + Ne loss 21 Ne/38Ar, 50 cm

21

0.8

1.0

Ne/22Ne, 10 cm 21 Ne/22Ne, 10 cm + Ne loss 21 Ne/22Ne, 50 cm

Al2O3 (wt%)

Na2O + K2O (wt%)

21

17

13 9 47.5

1.5 C

Glass - igneous rim Glassy pockets Mesostasis

A

48.5

49.5 SiO2 (wt%)

50.5

1.3 1.1 0.9 0.7 47.5

51.5

3.0

0.9 B

2.5 0.7

48.5

49.5 SiO2 (wt%)

50.5

51.5

toward An75

D Mag

+10% Mag

60% Pl

(Al/Mg)at

TiO2 (wt%)

2.0

0.5

1.5 1.0

Ol and low-Ca px IR

0.5 0.3 47.5

CK

48.5

49.5 SiO2 (wt%)

50.5

51.5

0.0 0

0.5

High-Ca px

1.0 1.5 (Ca/Mg)at

2.0

2.5

10

δ17OSMOW (‰) M

M A C P CC

A

CI

R

5

CR clan LLL H EH,EL

0

CM TFL

IR CO

-5

#45c.29

CV+CK

100 µm

PC

#45c.29

-2 -3

1 3% air O

Mix

8 7 % or O s cur e r p li ne g n i

20

Earth stratosphere

mass-dependent fractionation

-4

CK chondrites Chondrules in CK Olivine in CK Magnetite in CK Plagioclase in CK

-5 -6 -5

15

AM

5 10 18 δ OSMOW (‰) M

0

CC

-10 -5 Δ17O 0 B -1

0

5 10 18 δ OSMOW (‰)

15

20

25

14 12

SW B

20

Ne/22Ne

FSW

10 8 6 4 2 0 0.0

EA HL

#2 #1(1)

Sc. MM M240410 #45b.10 #45b.17 #45b.08

0.2

0.4 0.6 21 Ne/22Ne

#45c 29 TAM Un. MMs TAM Sc. MMs Antarctic MMs Trapped components GCR chondrites GCR MM (Fa26) GCR MM (Meso) SCR MM (Fa26) SCR MM (Meso)

#1(2)

0.8

1.0

Pyroxene 100

0.03

A

#45c.29 Pigeonite Augite CK chondrites Low-Ca px High-Ca px Px relicts in MM

B

80

Cr/Mg

Fe/Mn

0.02 60

40

0.01 20

Oxidation & Fe-depletion

Wo

0.00

0

0.0

0.1

0.2 Fe/Mg

0.3

0.0

En

0.1

Fs

0.2

0.3

0.4

Fe/Mg

Olivine 1.0

C 0.8

0.6

D

0.4

Dhofar 015

0.8

NiO (wt%)

MnO (wt%)

1.0

CC CK UOC EOC #45c.29 Ol relicts in MM

0.6

0.4

0.2

0.2

0.0

0.0 0

1.0

10

30

20 FeO (wt.%)

40

0

Magnetite 1.0

E

20 FeO (wt.%)

30

40

Field for CC CK #45c.29

F

CK3 TiO2 (wt%)

CK3

TiO2 (wt%)

10

0.5

0.5

CK4-6

CK4-6 0

0

2

4 Cr2O3 (wt%)

6

0

0

0.2

0.4 0.6 NiO (wt%) Dhofar 015

0.8

1.0

0.44

(A)

Xe/132Xe

0.42 0.40

134

0.38

#2

EA Q

#45c 29 steps #45c 29 totals TAM Un. MMs TAM Sc. MMs Trapped components Fractionation To Xe-HL To Pu fission

0.36

Xe/132Xe

0.17

130

0.16 0.15

#1

(B) Q

EA

0.14 0.13 0.32

0.34 136

0.36 Xe/132Xe

0.38

100

(21Ne/22Ne)cos

(21Ne/38Ar)cos

1.0

10 #45c.29: 0.94 ± 0.02

#1

1

0.9

#2

0.0 21

0.95

0.2

0.4 0.6 f-mesostasis

Ne/38Ar, 10 cm 21 Ne/38Ar, 10 cm + Ne loss 21 Ne/38Ar, 50 cm

21

0.8

1.0

Ne/22Ne, 10 cm 21 Ne/22Ne, 10 cm + Ne loss 21 Ne/22Ne, 50 cm

Table 1. Representative EPMA composition of mineral phases, glass and mesostasis in the coarse-grained, scoriaceous micrometeorite #45c.29. The whole dataset is given in Supplementary Table S1. Mineral Textural positiona SiO2 K2O V2O3 NiO Phase

TiO2 Total

Al2O3 Fa wt%

Cr2O3 Fs

Fe2O3b Wo

FeO An

MnO Ab

MgO

CaO

Na2O

mol.% Relict OL-1 38.2 0.12 0.05 0.04 n.c. 24.3 0.21 37.0 0.10 n.d. n.d. n.d. 0.66 100.7 27 Olivine Relict OL-5 37.5 0.09 0.05 0.04 n.c. 26.3 0.24 34.6 0.15 n.d. n.d. n.d. 0.64 99.5 30 Olivine Relict OL-7 38.0 0.07 0.21 0.19 n.c. 25.4 0.22 36.6 0.07 n.d. n.d. n.d. 0.64 101.3 28 Olivine Relict OL-3 38.4 0.08 0.09 0.04 n.c. 22.7 0.18 38.6 0.03 n.d. n.d. n.d. 0.69 100.8 25 Olivine Relict OL-3 38.2 0.04 0.06 0.03 n.c. 24.3 0.20 37.2 0.03 n.d. n.d. n.d. 0.68 100.7 27 Augite Rim light gray 53.1 0.36 2.85 0.74 n.c. 7.46 0.13 20.5 14.6 0.26 0.02 n.d. 0.09 100.1 12 30 Augite Rim dark gray 51.0 1.02 7.53 1.05 n.c. 1.41 0.21 19.0 19.2 0.09 0.01 n.d. 0.02 100.4 3 41 Augite Mesostasis c 52.4 0.94 5.65 0.60 n.c. 11.2 0.23 14.6 12.7 n.c. n.c. n.c. n.c. 97.7 21 30 Pigeonite Relict nuclei 57.3 0.35 1.61 0.74 n.c. 1.22 0.15 34.8 4.03 < 0.01 < 0.01 n.d. 0.01 100.2 2 8 Pigeonite Relict nuclei 56.3 0.39 1.89 0.79 n.c. 2.94 0.16 31.6 5.19 0.02 0.01 n.d. 0.03 99.3 5 10 Plagioclase Mesostasis 45.5 0.11 30.7 0.16 n.c. 1.77 0.06 1.33 17.0 1.34 < 0.04 n.d. < 0.08 98.0 87 12 Plagioclase Mesostasis 43.2 0.04 32.4 < 0.06 n.c. 1.4 < 0.06 0.72 18.5 0.47 < 0.04 n.d. < 0.08 96.9 96 4 Magnetite Aggregate 0.15 0.26 2.11 3.76 59.5 27.1 0.05 1.74 <0.01 n.d. n.d. 0.36 0.87 95.9 Magnetite Aggregate 0.31 0.33 2.52 3.98 59.7 26.8 0.05 2.48 <0.01 n.d. n.d. 0.36 0.81 97.3 Glass Igneous rim 51.0 0.47 11.9 0.07 n.c. 13.2 0.20 7.70 12.6 1.11 0.06 n.d. < 0.08 98.4 Glass Pocket, close to px 50.7 0.75 14.8 0.19 n.c. 6.14 0.11 10.5 16.9 0.95 0.08 n.d. < 0.08 101.2 Glass Pocket, close to ol 49.3 0.77 14.2 0.63 n.c. 8.65 0.11 9.85 15.4 0.97 0.10 n.d. < 0.08 100.0 Mesostasis d 47.8 0.37 22.0 0.24 n.c. 6.26 < 0.07 6.59 16.0 1.07 0.08 n.d. < 0.08 100.6 n.c.: not calculated; n.d.: not determined a) Labels of the olivine grains are given in Fig. 1B. For augite, the terms dark and light gray refer to contrast on BSE images. b) FeO and Fe2O3 were calculated using the stoichiometric method developed by Droop (1987). c) Composition of pyroxene in the mesostasis has been determined by mass-balance calculation (see text for further information). d) Mesostasis was analyzed with a defocused beam of 10 µm. Olivine

Table 2. He, Ne, Ar, Kr and Xe results for samples #1 and #2 of the coarse-grained, scoriaceous micrometeorite #45c.29. Specimen

Unit #1

#1

#1

#1

#2

Extraction stage [1] [2] total 4 He 10-12 cm3 STP 75.2 (5.5) 12.1 (4.3) 1064 (45) 3 He 10-12 cm3 STP 0.97 (2) 0.30 (1) 0.04 (1) 0.639 (28) 4 He 10-8 cm3 STP/g 340 (25) 55 (20) 395 (32) 4396 (190) 3 He 10-8 cm3 STP/g 4.37 (10) 1.38 (4) 0.16 (4) 2.64 (12) 3 He/4He 10-4 128.3 (9.7) 251.9 (89.9) 6.01 (36) 22 Ne 10-12 cm3 STP 1.93 (5) 4.49 (10) 4.04 (6) 22 Ne 10-8 cm3 STP/g 8.74 (25) 20.32 (48) 16.68 (28) 20 Ne/22Ne 7.25 (10) 4.51 (5) 5.30 (13) 8.26 (8) 21 Ne/22Ne 0.405 (9) 0.642 (101) 0.571 (7) 0.340 (5) 36 Ar 10-12 cm3 STP 5.27 (17) 6.70 (22) 9.66 (38) 36 Ar 10-8 cm3 STP/g 23.8 (8) 30.3 (1.0) (1.3) 39.9 (1.6) 40 Ar 10-8 cm3 STP/g 1829 (149) 1019 (148) 2848 (210) 4575 (155) 38 Ar/36Ar 0.2380 (48) 0.2512 (112) 0.2454 (66) 0.3474 (30) 40 Ar/36Ar 74.9 (4.5) 31.4 (4.6) (3.3) 108 (3) 84 Kr 10-12 cm3 STP 0.137 (5) 0.022 (4) 0.228 (16) 84 Kr 10-8 cm3 STP/g 0.621 (25) 0.099 (16) 0.940 (65) 80 Kr/84Kr 0.037 (2) 0.080 (8) 0.043 (2) 0.044 (2) 82 Kr/84Kr 0.196 (10) 0.199 (18) 0.197 (9) 0.182 (4) 83 Kr/84Kr 0.203 (8) 0.191 (14) 0.201 (7) 0.195 (4) 86 Kr/84Kr 0.330 (15) 0.345 (26) 0.332 (14) 0.318 (6) 132 Xe 10-12 cm3 STP 0.033 (5) 0.005 (1) 0.031 (1) 132 Xe 10-8 cm3 STP/g 0.147 (21) 0.022 (5) 0.129 (6) 124 Xe/132Xe 0.0024 (16) 0.0028 (70) 0.0053 (8) 126 Xe/132Xe 0.0035 (5) 0.0060 (25) 0.0030 (5) 128 Xe/132Xe 0.062 (3) 0.080 (10) 0.069 (3)

[3] 87.3 (7.0) 1.30 (2)

5.90 (11) 149.4 (12.3) 6.42 (10) 29.06 (54)

11.97 (27) 54.2

50.5 0.159 (6) 0.720 (30)

0.037 (5) 0.169 (21) 0.0024 (17) 0.0040 (6) 0.064 (3)

129

Xe/132Xe

0.896 (18)

1.374 (57)

0.961 (17)

0.135 (4)

0.140 (14)

0.136 (4)

0.785 (7)

0.765 (36)

0.782 (8)

0.401 (9)

0.412 (28)

0.402 (8)

0.375 (25)

0.377 (11)

1.053 (16) 130

Xe/132Xe 0.147 (4)

131

Xe/132Xe 0.798 (5)

134

Xe/132Xe 0.404 (7)

136

Xe/132Xe

0.377 (12) 0.361 (6) Uncertainties in the last digits are given in parentheses.

Table 3. Trapped and cosmogenic Ne and Ar in samples #1 and #2 of the coarse-grained, scoriaceous micrometeorite #45c.29. 20 21 38 Netrapped Necos 36Artrapped Arcos (20Ne/22Ne)trapped (21Ne/38Ar)cos #1 142 (8)16.22 (39) 51.9 (1.3) 3.51 (35) 11.64 (43) 4.62 (47) #2 133 (8) 5.32 (14) 35.2 (5.0) 7.22 (33) 11.97(21) 0.74 (4) Concentrations are expressed as 10-8 cm3 STP/g. Uncertainties in the last digits are given in parentheses. See text for details of the component resolution.

Table 4. Cosmogenic Ne in micrometeorites with long cosmic ray exposure age. Sample Typea 21 Necosd

Ref.b CRE aged

Compos. Source regiond

Diameterc CRE agee

Weightc Density Source regione µm µg AU Ma

g/cm³ 10-8 cm3 STP/g Ma AU F97AC013 Un.MM 1 CI 70 0.5 2.78 8.6 102 ± 21 34 - 42 157 48.0 F97AC003 Un.MM 1 CI 100 0.5 0.95 2.2 24 ± 11 16 - 26 38.5 34.0 F97AC001 Un.MM 1 CI 110 2.0 2.87 13 159 ± 16 30 - 33 237 46.3 M240410 Sc.MM 2 CI 114 1.9 2.46 32.6 393 ± 27 52 ± 3 615 79.2 #45b.17#1 Sc.MM 3 L 555 224 2.50 4.41 9±1 4.3 ± 0.2 23.9 6.9 #45c.29 #1 Sc.MM 4 Fa26/Mes 647 336 2.37 16 140 16 (a) Un.MM = unmelted micrometeorite; Sc.MM = scoriaceous micrometeorite. (b) References: 1: Osawa and Nagao (2002a); 2: Osawa et al. (2010); 3: Baecker (2014); 4: this study. (c) Average diameter and weight of whole micrometeorites (for #45b.17#1 and #45c.29#1, extrapolated from the density of the fragment analyzed for noble gases). (d) CRE ages and orbital distance of the source as published. (e) CRE ages and orbital distance of the source recalculated using the same model for cosmogenic production rates as for #45c.29 (Trappitsch and Leya, 2013, see Supplementary Material Table S3).