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effects on the composition of and noble gases in lunar samples: Insights from Apollo 12018 Evelyn Füri, Laurent Zimmermann, Etienne Deloule, Reto Trappitsch

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Evelyn Füri, Laurent Zimmermann, Etienne Deloule, Reto Trappitsch. Cosmic ray effects on the isotope composition of hydrogen and noble gases in lunar samples: Insights from Apollo 12018. and Planetary Science Letters, Elsevier, 2020, 550, pp.116550. ￿10.1016/j.epsl.2020.116550￿. ￿hal- 02926751￿

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Distributed under a Creative Commons Attribution - NonCommercial - NoDerivatives| 4.0 International License Earth and Planetary Science Letters 550 (2020) 116550

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Earth and Planetary Science Letters

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Cosmic ray effects on the isotope composition of hydrogen and noble gases in lunar samples: Insights from Apollo 12018 ∗ Evelyn Füri a, , Laurent Zimmermann a, Etienne Deloule a, Reto Trappitsch b a Centre de Recherches Pétrographiques et Géochimiques, Université de Lorraine, CNRS, F-54000 Nancy, France b Lawrence Livermore National Laboratory, Nuclear and Chemical Sciences Division, 7000 East Ave, L-231, Livermore, CA 94550, USA a r t i c l e i n f o a b s t r a c t

Article history: Exposure of rocks and regolith to solar (SCR) and galactic cosmic rays (GCR) at the ’s surface Received 19 May 2020 results in the production of ‘cosmogenic’ and at a rate that depends on Received in revised form 19 August 2020 a complex set of parameters, such as the energy spectrum and intensity of the cosmic ray flux, the Accepted 21 August 2020 chemical composition, size, and shape of the target as well as the shielding depth. As the effects of Available online xxxx cosmic rays on the D production in lunar samples remain poorly understood, we determine here the D Editor: F. Moynier content and noble gas (He-Ne-Ar) characteristics of nominally anhydrous mineral (olivine and pyroxene) Keywords: grains and rock fragments, respectively, from different documented depths (0 to ≥4.8 cm) within Apollo hydrogen olivine basalt 12018. Deuterium concentrations, determined by secondary mass spectrometry, and 3 21 38 noble gases cosmogenic He, Ne, and Ar abundances, measured by CO2 laser extraction static mass spectrometry, mare basalt are constant over the depth range investigated. isotope ratios (20Ne/22Ne ≈0.86 and 21Ne/22Ne cosmic rays ≈0.85) of the cosmogenic endmember are comparable to the theoretical signature of GCR-produced neon. cosmogenic nuclides These observations indicate that the presence of significant amounts of SCR nuclides in the studied sub- exposure age samples can be ruled out. Hence, D within the olivines and pyroxenes must have been predominantly produced in situ by GCR-induced reactions during exposure at the lunar surface. Comparison of the amount of D with the 21Ne (184 ± 26 Ma) or 38Ar (193 ± 25 Ma) exposure ages yields a D − − − production rate that is in good agreement with the value of (2.17 ± 0.11) × 10 12 mol(g rock) 1Ma 1 from Füri et al. (2017). These results confirm that cosmic ray effects can substantially alter the hydrogen isotope (D/H) ratio of indigenous ‘’ in returned extraterrestrial samples and meteorites with long exposure ages. © 2020 The Author(s). Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).

1. Introduction to the lunar interior (e.g., Anand et al., 2014; Barnes et al., 2016; Desch and Robinson, 2019; Füri et al., 2014; Greenwood et al., The hydrogen isotope ratio is the key indicator for plan- 2011; Hui et al., 2017; Robinson et al., 2016; Saal et al., 2013; etary water origin(s) because different reservoirs Sharp, 2017; Singer et al., 2017; Tartèse and Anand, 2013). How- (solar, chondritic, cometary) have characteristic D/H signatures ever, the D/H ratio of mantle-derived samples does not always (e.g., Alexander, 2017; McCubbin and Barnes, 2019). Volcanic reflect the hydrogen isotope composition of the lunar mantle glass beads and olivine-hosted melt inclusions therein, the phos- source. In addition to magmatic processes (e.g., degassing; Saal et phate mineral apatite, and nominally anhydrous minerals in var- al., 2013; Tartèse and Anand, 2013), (SW) implantation ious rock types returned from the Moon by the Apollo missions and cosmic ray induced spallation reactions – triggered by solar record a wide range of δD values (where δD[] =[(D/H)sample/ (SCR) and galactic cosmic rays (GCR) that can penetrate lunar mat- −6 ter to depths of a few centimeters or several meters, respectively [(D/H)SMOW − 1] × 1000, with (D/H)SMOW = 155.76 × 10 ; Hage- mann et al., 1970), between ≤−500 and ≥+1000 (see Mc- (Reedy and Arnold, 1972) – can modify the D/H signature of in- Cubbin et al., 2015 for a review), which have been interpreted digenous ‘water’ (i.e., H, H2, and/or H2O) in lunar rocks, minerals, to reflect hydrogen or water contributions from multiple sources, and volcanic glasses. such as the solar nebula, carbonaceous chondrites, and/or comets, Since the Moon is an airless body and has no global magnetic field, SW , including and noble gas , are im- planted into the top few tens of nanometers of all rocks or regolith * Corresponding author. grains exposed to the lunar surface environment (e.g., Hashizume E-mail address: [email protected] (E. Füri). et al., 2000). Nonetheless, a contribution of SW-implanted hydro- https://doi.org/10.1016/j.epsl.2020.116550 0012-821X/© 2020 The Author(s). Published by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/). 2 E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550 gen to the measured water abundances and D/H ratios can gen- ent depths (0 to ≥4.8 cm) within Apollo olivine basalt 12018 by erally be ruled out because all recent hydrogen isotope measure- SIMS. In parallel, we analyzed the noble gas (He-Ne-Ar) contents ments of lunar samples have been carried out in situ by secondary and isotope ratios of bulk rock fragments by CO2 laser extraction ion mass spectrometry (SIMS), in the interior of lunar volcanic static mass spectrometry to quantify the abundances of cosmo- glass beads (Füri et al., 2014) or within mineral grains and melt genic noble gas nuclides (3He, 21Ne, 38Ar) at each depth and to inclusions that were never directly exposed to SW irradiation (e.g., constrain the irradiation conditions and duration. This combined Barnes et al., 2014, 2013; Boyce et al., 2010; Hui et al., 2017; data set permits to determine if depth-dependent shielding vari- Robinson et al., 2016; Saal et al., 2013; Tartèse et al., 2013; Tartèse ations result in significant inter-sample differences in the rate of and Anand, 2013). In contrast, cosmic ray produced (‘cosmogenic’ cosmogenic D production within Apollo 12018. or ‘spallogenic’) deuterium is expected to contribute significantly to the D/H ratio of water-poor lunar samples that experienced long 2. Samples and analytical techniques exposure to cosmic rays at the surface of the Moon (Füri et al., 2017). Apollo sample 12018 is a medium-grained, low-Ti olivine basalt Knowledge of the cosmogenic D production rate (P D ) and the (supplementary Table S1; Papike et al., 1976) – also described as cosmic ray exposure (CRE) age of the sample of interest is critical olivine dolerite (Cuttitta et al., 1971; Kushiro et al., 1971) or gab- for correcting measured D/H ratios for the cosmogenic contribu- bro (Megrue, 1971) –, composed of approximately 70% large olivine tion, and, ultimately, for determining the source(s) of lunar water. and pyroxene crystals set in a variolitic matrix (Walter et al., 1971). −12 −1 −1 To this date, a P D value of 0.92 to 1 × 10 mol(g rock) Ma , Ten chips were allocated for this study by NASA’s Curation and derived by Merlivat et al. (1976) and Reedy (1981), has been used Analysis Planning Team for Extraterrestrial Materials (CAPTEM); in most studies of lunar samples (e.g., volcanic glasses, melt inclu- these chips were extracted from different documented depths (0 sions, apatites, plagioclase), irrespective of their chemical composi- to ≥4.8 cm; supplementary Table S2) along a slab cut through the tion. However, the P D value depends on the abundance of various middle of the rock that was originally 8 × 6 × 6cm in size (Fig. 1). target elements (O, Mg, Si, Fe, Al), the size and shape of the object, The chips were gently crushed in an agate mortar to obtain indi- and its exposure history (Reedy, 1981). Furthermore, Greenwood vidual olivine and pyroxene grains for hydrogen isotope analyses et al. (2018)argued that mare basalt 70215 analyzed by Merlivat by SIMS as well as small rock fragments for noble gas analyses by et al. (1976) contains indigenous OH (in addition to water from static noble gas mass spectrometry. terrestrial contamination and cosmogenic D), rendering their P D Olivine and pyroxene grains, separated from eight out of the estimate unreliable. Füri et al. (2017) recently obtained a signifi- ten depth samples, were mounted in crystalbond and polished in- −12 −1 −1 cantly higher P D value of (2.17 ±0.11) ×10 mol(g rock) Ma dividually with alumina powder and to minimize contami- for the Moon’s surface from SIMS analyses of nominally anhydrous nation by terrestrial water. Subsequently, the grains were mounted olivine grains with a wide range of CRE ages, either indicating in together with two ‘dry’ standards: synthetic forsterite that previous studies had severely underestimated the production (4.5 ± 1 ppm H2O; Wetzel et al., 2015) and synthetic Suprasil of cosmogenic deuterium by GCRs, or, alternatively, hinting as the 3002 quartz glass (≤1.66 ppm H2O; Hauri et al., 2017). The ma- presence of additional deuterium produced by SCRs in the studied jor element composition of the olivine and pyroxene grains was olivines. determined by microprobe analysis at the Université de The importance of SCR effects in lunar samples can be assessed Lorraine (Service Commun de Microscopie Electronique et de Mi- through studies of tracks (e.g., Crozaz, 1980) or concen- croanalyses X, Nancy, France) with a Cameca SX100 operating at trations of (14C, 10Be, 26Al, 53Mn, 81Kr) and stable 12 nA and 15 kV. Counting times on peaks and background were noble gas nuclides (e.g., Hohenberg et al., 1978; Leya et al., 2001; 10 s for all elements. Results are reported in supplementary Ta- Nishiizumi et al., 2009; Reedy, 1980; Reedy and Arnold, 1972; ble S3 and show that the olivine grains display a range in forsterite Trappitsch and Leya, 2014). SCR production occurs almost contents (Fo = 100 × Mg / [Mg + Fe]) between Fo66 and Fo77. exclusively within the topmost 1–2 cm of exposed rock or The selected pyroxene grains can be classified as (high-Ca) pi- surfaces on the Moon because of the lower energy of SCRs com- geonites (Kushiro et al., 1971). pared to GCRs (∼10–100 MeV vs. ∼1–10 GeV) (Reedy and Arnold, In situ measurements of H and D concentrations in the Apollo 1972). Since the rates of neon isotope production by SCRs are sub- 12018 olivines and pyroxenes – as well as in Apollo 12004, 12018, stantially higher that those by GCRs within the uppermost few and 15016 olivines previously studied by Füri et al. (2017)– were g/cm2 of shielding (where ‘shielding’ depth [g/cm2] = depth [cm] performed at the CRPG (Nancy, France) using the CAMECA 1280 × density [g/cm3]), cosmogenic neon concentrations are expected HR2 ion microprobe. Samples were coated with and left in − to decrease significantly with (shielding) depth in lunar samples the instrument airlock at ∼ 6 × 10 9 mbar for several days in or- that were exposed to both solar and galactic cosmic rays (Rao et der to ensure thorough removal of any terrestrial adsorbed water al., 1994, 1993; Trappitsch and Leya, 2014). Importantly, the main before introduction into the sample chamber. In addition, a liquid target elements and , from which 21Ne and 22Ne cold trap was used to reduce the hydrogen background − are predominantly produced in lunar rocks, show a lower cross and maintain a pressure of ≤ 7 × 10 10 mbar during analyses. 21 22 section ratio for Ne/ Ne at lower energies, i.e., the region in StHs6/80-G andesitic glass (250 ± 7 ppm H2O and δD =−95 ± 2; which production from SCRs is important, than at higher energies, Jochum et al., 2006) and MON9 pyrope (56 ± 6 ppm H2O; Bell et where production from GCRs dominates. Given that the dominant al., 1995), on the sample mount of Füri et al. (2017), were used as − − fluence of the SCR particle spectrum is at much lower energies additional standards. Spot analyses of H and D secondary ions than for the GCR particle spectrum, the 21Ne/22Ne ratio of SCR were carried out in mono-collection mode on an electron multi- neon is significantly lower than that of GCR neon (Füri et al., 2017; plier at a nominal mass resolution m/m = 1600 using a 10 kV + Rao et al., 1994, 1993). Therefore, the three isotopes of neon repre- Cs primary ion beam and a normal-incidence electron gun (emis- − sent a powerful tool to recognize, and possibly quantify, different sion = 0.21 mA) for charge compensation. The 18O count rate noble gas components (SW, SCR, GCR) in samples returned from was measured for 4 s on a Faraday cup during each cycle to moni- the Moon. tor the stability of secondary ion intensities. Standards and miner- To assess the importance of SCRs and GCRs for the produc- als were pre-sputtered for 180 s over an area of 30 × 30 μm prior tion of cosmogenic deuterium in lunar samples, we determined to signal acquisition. During the pre-sputtering process, the inten- the deuterium content of olivine and pyroxene grains from differ- sity of the primary beam was measured with the primary Fara- E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550 3

Fig. 1. Apollo 12018 fragments used for this study are derived from a) slab 14; b) column 17; c) slices 49, 51, 55, and 52 (see supplementary Table S2 for details). Sample ◦ 12018,14,17,52 originates from the ‘top’ exterior surface of the rock, where the angle of exposure was ∼45 from the zenith, as indicated by cutting diagrams (Meyer, 2011). (NASA images # S70-19566, S70-19581, S70-19598). day cup; it was found to decrease only slightly (i.e., from 11.2 to gon was separated from and neon by adsorption onto a 10.0 nA and from 13.8 to 12.4 nA) over each 24-hour measurement charcoal finger at 77 K, and helium and neon were subsequently , which included both standards and samples. For analysis, trapped onto a He-cooled cryogenic trap at ∼15 K. Helium was the ∼15 μm beam was rastered over an area of 20 × 20 μm, and a first released from this trap at 34 K and analyzed in peak-jumping dynamical transfer operating system was used to compensate the mode (4He on the H2 Faraday detector, 3He on the central (Ax) primary rastering and to refocus the beam in the secondary part compact discrete dynode (CDD) detector). Neon was released from of the ion probe. To eliminate any hydrogen contamination from the cryogenic trap by increasing the temperature to 110 K, and the crater edges, a 1800 μm field aperture limiting the analyses to the amount of gas introduced into the mass spectrometer was ions from the central ∼10 μm of the beam and an electronic gate adjusted through volume dilution in the volume-calibrated purifi- − − of 80% were used. The H and D ion intensities were determined cation line to match the 20Ne signal of air standard measurements for 4 and 20 s, respectively, for 30 cycles. Under these analytical with 20Ne/22Ne = 9.80 (Fig. 2). The three isotopes of neon were conditions, typical count rates on StHs6/80-G were ∼99,000 cps analyzed in multi-collection mode (22Ne on H1 CDD, 21Ne on Ax − − for H and 16 cps for D , whereas count rates on the ‘dry’ stan- CDD, 20Ne on L2 CDD). Neon isotope analyses consisted of 5 blocks − − dards averaged 1770 ± 667 cps for H and 0.30 ± 0.14 cps for D of 30 cycles each, and peak centering was performed at the start (supplementary Table S4), in agreement with our previous study of each measurement block. A charcoal finger at 77 K and a Zr-Al − − (Füri et al., 2017). Measured H and D count rates, as well as getter at room temperature were used to minimize the contri- 40 20 22 estimated deuterium concentrations in Apollo 12018 olivines and bution of doubly charged Ar and CO2 to the Ne and Ne pyroxenes, are reported in supplementary Table S4. signals, respectively. Given the high mass resolution of the de- 40 ++ Noble gas (He-Ne-Ar) abundances and isotope ratios were de- tector in the L2 position (∼1800), Ar was partially resolved termined by CO2 laser extraction static mass spectrometry at the from the peak of interest (Honda et al., 2015; Wielandt and Storey, CRPG noble gas analytical facility (Füri et al., 2018; Humbert et al., 2019; Zhang et al., 2016); therefore, no correction was applied 20 + + 2000). For each depth sample, two separate fragments, between to the Ne signal. The CO2 signal was measured at the begin- + 1.54 and 5.88 mg in mass (supplementary Table S5), were ana- ning of each analysis, and the 22Ne signal was corrected using ++ + lyzed. The fragments were placed into different pits of the laser a CO2 /CO2 ratio of 0.4%; notably, the contribution of ++ 22 + ∼ chamber connected to the purification line of the HELIX-MC Plus CO2 to the Ne signal amounted to only 1 cps, and is there- + (Thermofisher Scientific) multi-collector noble gas mass spectrom- fore negligible. The 21Ne signal was measured at the peak center; + eter. After sample introduction, the sample chamber was baked at no hydride (20NeH ) correction was performed because the getter ◦ 110 C overnight under ultra-high vacuum to remove any terres- allowed maintaining a low hydrogen background. After releasing trial adsorbed gases. Each rock fragment was heated individually from the charcoal finger, the argon isotopes were analyzed in 40 38,36 with a continuous-mode CO2 laser (λ = 10.6μm). Two to peak-jumping mode ( Ar on the central Faraday detector, Ar three heating steps were applied by modulating the power of the on Ax CDD). laser and monitoring the heating procedure on a TV screen us- Standard HESJ (He Standard of Japan) of Matsuda et al. (2002) ∼ ◦ 3 4 ing a CCD camera. A low-temperature step ( 600 C) was aimed with a He/ He ratio of 20.63 ± 0.10 RA (where RA is the atmo- at removing surface-sited (atmospheric, solar) gases, an ‘interme- spheric 3He/4He ratio) was used as a helium standard, whereas air ◦ diate’ heating step (∼800 C) was applied for samples from the aliquots were used to determine the analytical sensitivity and re- near-surface (≤0.6 cm) to remove any remaining solar gases, and producibility for neon and argon (Zimmermann and Füri, 2015). fusion was achieved in the last heating step to release ‘volume- The reproducibility (1σ s.d.) of standard measurements was 1.2% correlated’ (cosmogenic, radiogenic) noble gas components (Curran for 4He, 0.9–1.1% for 20Ne, and 1.3% for 36Ar abundances, and 1.5% et al., 2020). Re-heating of three samples from ∼1 to 1.5 cm depth for 3He/4He, 0.2% for 20Ne/22Ne, 0.6% for 21Ne/22Ne, and 0.4% for (supplementary Table S5) demonstrated that ≥97.7% of the total 38Ar/36Ar. Reported uncertainties (2σ s.d.) for sample isotopic ra- noble gas content was extracted upon melting. The extracted no- tios take into account the standard reproducibility (in addition ◦ ble gases were purified using five hot (500 C) and cold (room to the analytical precision), even though samples and standards temperature) Ti sponge getters (Zimmermann et al., 2015). Ar- have significantly different isotope ratios for helium, neon, and ar- 4 E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550

Fig. 2. a) 20Ne/22Ne and b) 21 Ne/22Ne ratios (corrected for instrumental mass fractionation) as a function of the 20Ne signal (in counts per second, cps) for air standards (n = 37) and Apollo 12018 fragments analyzed by step-wise heating. The variable 20 Ne count rates for the air standard measurements were obtained by varying the amount of gas introduced into the mass spectrometer through volume dilution in the volume-calibrated purification line. The horizontal dashed lines indicate the 20Ne/22Ne and 21Ne/22Ne ratios of the terrestrial (20Ne/22Ne = 9.80 and 21Ne/22Ne = 0.0288–0.0290; Eberhardt et al., 1965; Györe et al., 2019; Saxton, 2020; Wielandt and Storey, 2019). Uncertainties (2σ s.d.) are smaller than symbol sizes.

− gon (Fig. 2). Procedural blanks, with the laser off, averaged 2.1 × The D signals of 13 olivine and two pyroxene grains from − − 10 16 mol 20Ne and 2.1 × 10 17 mol 36Ar. Helium blanks were the Apollo 12018 depth samples vary between 2.09 ± 0.26 and below the detection limit. Blank-corrected He-Ne-Ar abundances 2.93 ± 0.31 cps, and the Apollo 12018 olivines prepared by and isotope ratios for Apollo 12018 rock fragments are reported in Füri et al. (2017)yield similar signals of 2.22 ± 0.27 to 2.33 ± supplementary Table S5. 0.28 cps upon reanalysis for this study (supplementary Table S4). − Apollo 12004 and 15016 olivines yield different D ion intensi- 3. Results ties of 1.09–1.10 cps and 4.09–4.26 cps, respectively, because of distinct CRE ages (Section 4.3; Füri et al., 2017). All of these val- 3.1. Hydrogen and deuterium content of Apollo 12018 olivines and − ues are significantly higher than the D signals of the quartz glass pyroxenes (0.25 ± 0.09 cps) and forsterite (0.33 ± 0.11 cps) standards, in- − − The measured H count rates, normalized to either the primary dicating that the D measurements are not substantially affected − 18 − − 18 − ion beam intensity (H /Ip) or the O signal (H / O ), of the by any terrestrial contamination. Based on the known D content −9 − ‘dry’ standards (i.e., synthetic forsterite and Suprasil 3002 quartz (3.9 × 10 mol/g; Jochum et al., 2006) and the measured D glass) correspond to H O concentrations of 4.1 to 4.6 ppm when signal (16.0 ± 1.2cps), normalized to the primary ion beam in- 2 − compared to the glass and garnet standards. Given that the H2O tensity, of standard StHs6/80-G, the D signals of Apollo 12018 content of the quartz glass (≤1.66 ppm H2O; Hauri et al., 2017) olivines and pyroxenes correspond to D concentrations of 0.50 to −9 −9 cannot be clearly resolved from that of the forsterite standard (4.5 0.61 × 10 mol/g. Similar values of 0.48 to 0.62 × 10 mol/g 18 − ± 1 ppm H2O; Wetzel et al., 2015), the H2O detection limit is in- are obtained when the D abundances are calculated using the O ferred to be on the order of 4.5 ppm in this study. The majority count rate for normalization (supplementary Table S4). Within er- (n = 33 out of 43) of the lunar olivine and pyroxene grains ana- ror, these values are consistent with the D content of Apollo 12018 − −9 lyzed here yield H count rates that are slightly higher than those olivines determined previously (0.49 to 0.54 × 10 mol/g; Füri et of both ‘dry’ standards (supplementary Table S4). Their H2O con- al., 2017). Importantly, the new results show that the D concentra- centrations are estimated to vary between ∼4.5 and 9.2 ppm, i.e., tion in olivine and pyroxene is indistinguishable, and there is no measured water contents are up to ∼5 ppm above the detection D concentration gradient between the surface and ∼4.5 cm depth − limit. The elevated H signals may indicate that Apollo olivines (Fig. 3). and pyroxenes can contain several ppm of indigenous lunar wa- ter, although two previous studies suggested that indigenous water 3.2. He-Ne-Ar characteristics of Apollo 12018 rock fragments is present in very low abundance within lunar olivines (≤2 ppm) (Hui et al., 2013; Mosenfelder and Hirschmann, 2016). Alterna- At the first heating step, roughly half of the total 3He and 4He tively, the samples might be variably contaminated by terrestrial content is extracted, together with only a small fraction (∼1.5 to adsorbed water. Distinguishing between these two scenarios is not 7%, in most cases) of the total neon and argon abundance (supple- possible. This demonstrates the difficulty of accurately and pre- mentary Table S5). This indicates that noble gases in Apollo 12018 cisely quantifying trace amounts of indigenous water (hydrogen) fragments are predominantly ‘volume-correlated’, i.e., produced in in lunar samples using SIMS because terrestrial contamination (in- situ by (4He and 40Ar) or spallation reactions, troduced during sample preparation and/or analysis) might result and are only efficiently released upon melting at high temper- in ppm-level variations in measured water contents. SW-derived atures. Indigenous () noble gases have never been found in or cosmogenic hydrogen is not expected to contribute to the mea- lunar samples (Füri et al., 2018; Wieler and Heber, 2003), and are, − sured H signals because analyses were carried out in situ within therefore, ruled out as a possible component. Total concentrations mineral grains that were never directly exposed to the SW, and of radiogenic 4He and 40Ar vary by ∼25 to 30% in the twenty indistinguishable H contents were found for samples with a wide rock fragments analyzed here, whereas 3He, 21Ne, and 38Ar abun- range of CRE ages (Section 4.3; Füri et al., 2017). dances vary by ∼10 to 15% (1σ s.d.). This is consistent with the E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550 5

(see Wieler et al. (2007)for details on this component) and cos- mogenic neon produced by nuclear interactions (Fig. 4a). For all depth samples from this study, the isotopic composition of neon extracted at high temperatures (heating steps 2 and 3) falls onto the same mixing line, irrespective of the analyzed neon amount (Fig. 2). Furthermore, the measured neon isotope ratios of this cos- mogenic endmember (20Ne/22Ne ≈ 0.86 and 21Ne/22Ne ≈ 0.85) are comparable to the theoretical signature of GCR-produced neon (20Ne/22Ne ≈ 0.75 and 21Ne/22Ne ≈ 0.89), which can be calcu- lated using the 2π exposure model from Leya et al. (2001)for low shielding (0–15 g/cm2), together with the major element composi- tion of olivine basalt 12018 given in supplementary Table S1. How- ever, for most samples, neon extracted at the first heating step has a lower 21Ne/22Ne ratio of ∼0.72 (at a 20Ne/22Ne ratio of ∼0.86; Figs. 2 and 4a), comparable to the signature of neon produced by SCRs, calculated for a rigidity R0 = 100 MV and an incident parti- 2 cle flux J0 = 100 protons/s/cm (Füri et al., 2017; Trappitsch and Leya, 2014).

4. Discussion Fig. 3. Deuterium concentration in Apollo 12018 olivines and pyroxenes, as well as 3He, 21Ne, and 38Ar concentrations in rock fragments (duplicates at each depth), as a function of depth below the surface. Average values (solid lines) and their 1σ 4.1. SCR nuclides in Apollo 12018? standard deviations (shaded areas) are indicated for each nuclide. Uncertainties (2σ s.d.) of noble gas abundances are smaller than symbol sizes. The low 21Ne/22Ne ratio of neon extracted at the first heat- ing step appears to hint at the presence of SCR noble gases in notion that small-scale chemical or mineralogical heterogeneities Apollo olivine basalt 12018. Indeed, the fragments analyzed here (e.g., olivine phenocrysts, up to 1–2 mm in size) can affect the cover the entire depth range in which SCR nuclides can be pro- concentration of parent (or target) nuclides, and, correspondingly, duced. Models indicate that the abundance of SCR 21Ne should the abundances of radiogenic and cosmogenic isotopes in mg-sized drop off very rapidly within the topmost cm, whereas the concen- sub-samples (Füri et al., 2017). Nonetheless, Fig. 3 clearly shows tration of GCR 21Ne is expected to increase slightly with increasing that duplicates from a given depth contain similar gas amounts, shielding (i.e., with increasing depth below the surface) because, and there is no noble gas concentration gradient between the rock for each incident primary particle, a cascade of secondary parti- surface and 4.8 cm depth. cles is produced which can then undergo nuclear interactions (e.g., Since only a small amount of neon (and argon) is released from Reedy and Arnold, 1972; Fig. 5). Given that the neon isotope pro- Apollo 12018 rock fragments at low temperature, neon (and ar- duction rates by SCRs are significantly higher that those by GCRs gon) isotope count rates are significantly lower at the first heating at the uppermost rock surface (i.e., by about a factor of four for step compared to the second step (Fig. 2). However, Fig. 2 shows 21Ne; Leya et al., 2001; Trappitsch and Leya, 2014; Fig. 5), ‘total’ that air standards yield constant neon isotope ratios, i.e., 20Ne/22Ne (SCR + GCR) cosmogenic neon concentrations are expected to de- and 21Ne/22Ne, over the entire range of count rates measured for crease significantly with (shielding) depth in lunar samples that the samples. This demonstrates that there is no analytical bias for were exposed to both solar and galactic cosmic rays; the same is the different heating steps, and differences in neon (and 36Ar/38Ar) true for cosmogenic 3He and 38Ar. Indeed, Rao et al. (1993, 1994) isotope ratios can be attributed to variable contributions from dif- observed a systematic decrease of 3He, 21,22Ne, and 38Ar concen- ferent components (SW, SCR, GCR) in the analyzed noble gas frac- trations with increasing depth in sub-samples of Apollo 61016 and tions. 68815, particularly within the first cm below the surface, indicat- Previous analyses by single-step heating revealed that frag- ing a progression from a mixed noble gas component produced ments from “the upper pitted surface” of Apollo 12018 contained from SCRs and GCRs to a pure GCR component. It is noteworthy significant quantities of SW-derived noble gases, as indicated by that Apollo samples 61016 and 68815 both have exposure ages of 20Ne/22Ne and 36Ar/38Ar ratios of up to 11.4 and 5.4, respectively ∼2 Ma only (Rao et al., 1994, 1993), implying that any shielding (Megrue, 1971), whereas three other chips from the near-surface changes during their residence on the lunar surface are unlikely to were dominated by cosmogenic neon and argon, with much lower be important. 20Ne/22Ne and 36Ar/38Ar ratios of 1.20–1.84 and 0.772–0.98, re- Fig. 3 shows that D, 3He, 21Ne, and 38Ar concentrations in spectively (Bogard et al., 1971; Hintenberger et al., 1971) (Fig. 4). Apollo 12018 are constant over the studied depth interval, im- The results from this study show that several Apollo 12018 frag- plying that the theoretical depth-dependent ments from various depths contain a small fraction of SW-derived production rates are not appropriate for the sub-samples stud- noble gases (with 20Ne/22Ne = 2.135 to 9.708 and 36Ar/38Ar = ied here. SCR production rates depend on the rigidity (momentum 1.026 to 3.016), predominantly released during the first heat- per unit charge; R0) and flux ( J0) of SCR particles, as well as on ing step (Fig. 4; supplementary Table S5). Even fragments ex- the sample size, shape, and orientation, and its erosion rate (e.g., tracted from 1.5 to ∼3cm depth yield elevated 20Ne/22Ne and Rao et al., 1994). While there is some debate about the correct 36 38 Ar/ Ar ratios compared to the cosmogenic endmember; this choice of R0 and J0 for modeling cosmic ray interactions with lu- clearly demonstrates that some sub-samples below the rock sur- nar samples (e.g., Rao et al., 1994; Reedy, 1980), the irradiation face were contaminated by SW-loaded dust during cutting, as pre- history of Apollo 12018 at the surface of the Moon is the key viously suggested for other Apollo samples by Füri et al. (2015, unknown. Erosion by cosmic ray sputtering and by impacts of cos- 2017). mic dust and micrometeorites could have modified the outermost In a three-isotope plot of neon, the data from Bogard et al. rock surface, whereas tumbling and intermittent burial could have (1971), Hintenberger et al. (1971), and Megrue (1971)define a resulted in changes in sample orientation and variable shielding mixing line between implantation-fractionated SW-derived neon conditions (Reedy, 1980; Reedy and Arnold, 1972). Furthermore, 6 E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550

Fig. 4. a) Three-isotope plot of neon and b) 20Ne/22Ne versus 36Ar/38Ar for Apollo 12018 fragments. Step-heating data from this study are shown together with previous results obtained by single-step heating (Bogard et al., 1971; Hintenberger et al., 1971; Megrue, 1971). The neon and argon isotope compositions of modern solar wind (SW; Heber et al., 2009) and the terrestrial atmosphere (Air; Eberhardt et al., 1965; Györe et al., 2019; Saxton, 2020; Wielandt and Storey, 2019), as well as the calculated isotope ratios of cosmogenic neon produced by solar cosmic rays (SCR; Trappitsch and Leya, 2014) and galactic cosmic rays (GCR; Leya et al., 2001) for shielding between 0 and 15 g/cm2, are shown for comparison. The solid line in a) represents a mixing line between the GCR endmember and implantation-fractionated SW-derived neon. Uncertainties (2σ s.d.) of isotope ratios from this study are smaller than symbol sizes (uncertainties for previous results were not reported).

tamination, isotopically distinct from implanted SW and adsorbed terrestrial atmospheric gases (Fig. 4).

4.2. GCR noble gases and cosmic ray exposure (CRE) ages

Assuming that olivine basalt 12018 contains a binary mixture of SW-derived and cosmogenic (GCR-derived) 20,21,22Ne and 36,38Ar, the amount of cosmogenic 21Ne and 38Ar can be derived numeri- cally for each rock fragment, based on the isotope composition of the two endmembers (see Füri et al., 2014, 2018, for details on the component deconvolution). We use here 21Ne/22Ne ratios of 0.035 (measured by Megrue, 1971) and 0.89 (calculated from the model of Leya et al., 2001) for SW-derived and cosmogenic neon, respectively, and 36Ar/38Ar ratios of 5.4 (measured by Megrue, 1971) and 0.65 (i.e., the lowest ratio measured in this study) for SW-derived and cosmogenic argon (Fig. 4). Based on this two- component model, we estimate that ≥99.4% of the measured 21Ne and 38Ar in most Apollo 12018 depth samples was produced in situ by GCR-induced spallation reactions during exposure at the Fig. 5. Theoretical depth-dependent production rates of cosmogenic 21Ne by solar lunar surface. Only the three fragments with 36Ar/38Ar ratios ≥1.5 (SCR; Trappitsch and Leya, 2014) and galactic cosmic rays (GCR; Leya et al., 2001) at the first heating step contain ∼2–4% SW-derived argon. Impor- as a function of depth below the surface of Apollo olivine basalt 12018. tantly, these results are insensitive to the precise ratios assumed for the two endmembers. The measured 3He can be inferred to be surface documentation and cutting diagrams indicate that rock col- entirely of cosmogenic origin; however, since cosmogenic 3He is umn 12018,14,17, from which the fragments for this study were readily lost by diffusion (e.g., Rao et al., 1994), 3He exposure ages ◦ extracted, was oriented at an angle of ∼45 relative to the lunar are not discussed here. zenith (Fig. 1; Meyer, 2011). Therefore, even though our samples A CRE age can be derived by comparing the accumulated abun- 21 38 are thought to originate from the ‘top’ exterior surface of the rock, dance of cosmogenic noble gas nuclides ( Necosm, Arcosm) with it is possible that they have not been measurably affected by SCRs, empirical or theoretical noble gas production rates (e.g., Hohenberg as a result of the complex and prolonged (see Section 4.2) expo- et al., 1978; Leya et al., 2001; Reedy, 1981). However, it should sure history of Apollo 12018. again be emphasized that nuclide production rates are function of Overall, based on the flat D, 3He, 21Ne, and 38Ar concentration the concentration of various target elements, as well as of the sam- profiles (Fig. 3) and the neon isotope composition of the ‘volume- ple orientation and shielding during the entire exposure history at correlated’ cosmogenic component (Fig. 4a), the presence of signif- the lunar (near-)surface. Intra- and inter-sample variations in ef- icant amounts of SCR nuclides in the studied rock fragments can fective nuclide production can, therefore, be substantial (Drozd et be ruled out. Consequently, cosmogenic deuterium and noble gases al., 1974; Füri et al., 2017). Nonetheless, since the 2π exposure in Apollo 12018 must have predominantly been produced by GCRs. model from Leya et al. (2001)closely reproduces the neon iso- The origin and of the low-21Ne/22Ne neon endmember re- tope ratios of Apollo 12018 fragments, and because these authors mains enigmatic; given that this component is detected at low argued that their modeled noble gas data are consistent with ra- temperatures and represents only a small fraction of the total gas dionuclide results, we use their physical model to calculate 21Ne amount, it might represent some unidentified form of surface con- exposure ages. By assuming that the production rate of GCR 21Ne E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550 7

Fig. 7. Deuterium content of olivines and pyroxenes as a function of the cosmic 21 21 38 ray exposure age derived from the Necosm concentration in mare basalt fragments Fig. 6. Cosmic ray exposure ages derived from Necosm (T21) and Arcosm (T38) concentrations, assuming an uncertainty of 10% for the respective production rates. (T21). Results from this study for Apollo 12004, 12018, and Apollo 15016 are shown Results for the twenty Apollo 12018 rock fragments (black circles) and their average together with the data from Füri et al. (2017). The new results agree with the D ± × −12 −1 −1 value (red circle) are compared to the 81Kr-Kr exposure age from Marti and Lugmair production rate of (2.17 0.11) 10 mol(g rock) Ma (solid line) but are ∼ × −12 −1 −1 (1971)(grey diamond). (For interpretation of the colors in the figure(s), the reader inconsistent with the P D value of 1 10 mol(g rock) Ma from Merlivat is referred to the web version of this article.) et al. (1976)and Reedy (1981)(dashed line).

− − − varies between 4.8 and 6.1 × 10 14 mol(g rock) 1Ma 1 for shield- ages (T ). The results from this study are consistent with those ing between 0 and 15 g/cm2 in olivine basalt 12018 (Fig. 5), we 21 21 from Füri et al. (2017), and can be explained by in situ production obtain an average Ne CRE age (T21) of 184 ± 26 Ma for the of cosmogenic D by GCR-induced spallation reactions during expo- twenty rock fragments (Fig. 6). As discussed by Füri et al. (2017, 38 sure at the lunar surface at a time-averaged rate of (2.17 ± 0.11) × 2018), the production rate of cosmogenic Ar is debated, in part − − − 10 12 mol(g rock) 1Ma 1. It is worth noting that a slightly longer because the ‘total’ abundance of the major target element Ca is exposure duration of ∼215 Ma for Apollo 12018 would result in highly sensitive to mineralogical heterogeneities (e.g., the distri- an even better agreement with the other coupled D-21Ne data; bution of high-Ca pigeonites and plagioclase). We use here the − − − empirically-derived value of 4.6 × 10 14 mol(g rock) 1Ma 1 from this value is entirely compatible with the exposure ages found here Bogard et al. (1971) and Hintenberger et al. (1971); thus, the re- when the scatter in the data and the uncertainties are considered 38 (Fig. 6). The intercept of the best-fit line through the data indicates sulting Ar CRE ages (T38) agree, in most cases, within uncertain- ties with those derived from 21Ne (T ), and yield an average value that olivine and pyroxene grains in Apollo mare basalts with CRE 21 = ∼ × −9 −1 of 193 ± 25 Ma (Fig. 6). Hintenberger et al. (1971)reported com- ages 0 Ma contain 0.1 10 mol D (g rock) . This is con- parable exposure ages based on 3He (180 Ma), 21Ne (210 Ma), and sistent with the observation that the mineral grains studied here 38Ar (200 Ma) concentrations, and Stettler et al. (1973) obtained contain a few ppm indigenous and/or terrestrial adsorbed water a similar 38Ar CRE age of 170–180 Ma for Apollo 12018. Further- with a terrestrial-like D/H ratio (see Section 3.1). more, Marti and Lugmair (1971)determined a CRE age of 195 ± Although Reedy (1981)argued for a significantly lower P D ∼ × −12 −1 −1 16 Ma by using the 81Kr-Kr technique; this method is assumed to value of 1 10 mol(g rock) Ma for lunar mare basalts be largely independent of chemistry and shielding. We note that exposed to GCRs, his theoretical calculations clearly demonstrate Füri et al. (2017)argued for a significantly longer exposure dura- the importance of chemistry and shielding for the production of tion of 242 ± 42 Ma because their Apollo 12018 fragment con- cosmogenic nuclides. Assuming that D is produced from the target 21 21 ∼ tained a larger amount of cosmogenic Ne ( Necosm = 13.86 × elements O, Mg, Al, Si, and Fe at a ratio of 2:1:1:1:0.5, olivine −12 21 10 mol/g) compared to the Necosm concentrations measured and pyroxene (pigeonite) in Apollo 12018 (supplementary Table 21 −12 here ( Necosm = 10.0 ± 1.3 × 10 mol/g; Fig. 3), possibly as a S3) are expected to contain a similar amount of cosmogenic D after result of a higher proportion of forsteritic olivine (rich in the target ∼200 Ma of exposure, but the same would not be true for chemi- element Mg); however, the amount of cosmogenic 38Ar was com- cally distinct minerals. Furthermore, SCR effects could significantly 38 −12 parable between the two studies ( Arcosm = 10.29 × 10 mol/g increase the effective cosmogenic D production rate at the upper- − vs. 9.2 ± 1.5 × 10 12 mol/g; Fig. 3). These observations demon- most rock surface, unless a high surface erosion rate modifies the strate that, even though noble gas contents and isotope ratios can nuclide production profile, as expected for samples that experience be determined at very high precision for mg-sized samples, small- long (≥10 Ma) exposure at the surface of the Moon (Reedy, 1981). scale chemical or mineralogical heterogeneities can result in a Consequently, the chemical composition of the sample of interest range of CRE ages obtained from the nuclide accumulation method, and the irradiation conditions must be known and taken into ac- particularly for medium- or coarse-grained lunar samples such as count when correcting measured D/H ratios of water in lunar or olivine basalt 12018 (Fig. 6). other extraterrestrial materials for the cosmogenic D contribution. The production rate of cosmogenic D in other minerals has yet to 4.3. Cosmogenic deuterium in lunar olivines and pyroxenes be determined. Thus, care should be taken when applying pub- lished P D values for deriving the isotope signature of indigenous Fig. 7 shows the deuterium content of Apollo 12018 olivines water in apatite or plagioclase, particularly for water-poor samples and pyroxenes – together with the deuterium concentration in that experienced prolonged exposure to cosmic rays at the surface Apollo 12004 and 15016 olivines – as a function of their 21 Ne CRE of the Moon. 8 E. Füri et al. / Earth and Planetary Science Letters 550 (2020) 116550

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