Indigenous nitrogen in the Moon: Constraints from coupled nitrogen–noble gas analyses of mare basalts

Earth and Planetary Science Letters 431 (2015) 195–205

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Earth and Planetary Science Letters www.elsevier.com/locate/epsl

Indigenous nitrogen in the Moon: Constraints from coupled nitrogen–noble gas analyses of mare basalts Evelyn Füri a,∗ , Peter H. Barry b,c , Lawrence A. Taylor b , Bernard Marty a a b c

Centre de Recherches Pétrographiques et Géochimiques, CNRS-UL, 15 rue Notre Dame des Pauvres, BP 20, 54501 Vandoeuvre-lès-Nancy, France Planetary Geosciences Institute, Department of Earth Planetary Sciences, University of Tennessee, Knoxville, TN 37996-1410, USA Department of Earth Sciences, University of Oxford, Oxford OX1 3AN, UK

a r t i c l e

i n f o

Article history: Received 18 March 2015 Received in revised form 30 July 2015 Accepted 13 September 2015 Available online 1 October 2015 Editor: C. Sotin Keywords: Moon mare basalts anorthosites nitrogen noble gases volatile origin

a b s t r a c t Nitrogen and noble gas (Ne–Ar) abundances and isotope ratios, determined by step-wise CO2 laserextraction, static-mass spectrometry analysis, are reported for bulk fragments and mineral separates of ten lunar mare basalts (10020, 10057, 12008, 14053, 15555, 70255, 71557, 71576, 74255, 74275), one highland breccia (14321), and one ferroan anorthosite (15414). The mare basalt sub-samples 10057,183 and 71576,12 contain a large amount of solar noble gases, whereas neon and argon in all other samples are purely cosmogenic, as shown by their 21 Ne/22 Ne ratios of ≈0.85 and 36 Ar/38 Ar ratios of ≈0.65. The solar-gas-free basalts contain a two-component mixture of cosmogenic 15 N and indigenous nitrogen (<0.5 ppm). Mare basalt 74255 and the olivine fraction of 15555,876 record the smallest proportion of 15 Ncosm ; therefore, their δ 15 N values of −0.2 to +26.7h (observed at the low-temperature steps) are thought to well represent the isotopic composition of indigenous lunar nitrogen. However, δ 15 N values ≤−30h are found in several basalts, overlapping with the isotopic signature of Earth’s primordial mantle or an enstatite chondrite-like impactor. While the lowest δ 15 N values allow for nitrogen trapped in the Moon’s interior to be inherited from the proto-Earth and/or the impactor, the more 15 N-enriched compositions require that carbonaceous chondrites provided nitrogen to the lunar magma ocean prior to the solidification of the crust. Since nitrogen can efficiently be incorporated into mafic minerals (olivine, pyroxene) under oxygen fugacities close to or below the iron-wustite buffer (Li et al., 2013), the mare basalt source region is likely characterized by a high nitrogen storage capacity. In contrast, anorthosite 15414 shows no traces of indigenous nitrogen, suggesting that nitrogen was not efficiently incorporated into the lunar crust during magma ocean differentiation. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The recent discoveries of H-bearing species in lunar volcanic glasses (LVGs) (Saal et al., 2008), olivine-hosted melt inclusions (Hauri et al., 2011), and apatites (McCubbin et al., 2010) indicate that at least some parts of the lunar mantle contain a significant amount of indigenous ‘water’. These findings provide tight constraints on the formation conditions and subsequent evolution of the Earth–Moon system. Models that explain the origin of the Moon by a giant impact between the proto-Earth and a planetesimal predict extensive melting and partial vaporization of the silicate material that enters the proto-lunar orbit (e.g., Canup, 2004; Nakajima and Stevenson, 2014). The depletion of the Moon in the moderately volatile element K, and all elements with

*

Corresponding author. Tel.: +33 3 83 59 42 10; fax: +33 3 83 59 42 22. E-mail address: [email protected] (E. Füri).

http://dx.doi.org/10.1016/j.epsl.2015.09.022 0012-821X/© 2015 Elsevier B.V. All rights reserved.

higher volatility than K (Albarède et al., 2014; Hauri et al., 2015; Ringwood and Kesson, 1977), is thought to be a direct consequence of these high-temperature processes. The existing impact models are, therefore, unable to account for the presence of water in the lunar interior, because hydrogen, the lightest of all elements, is expected to have been quantitatively lost during this period. Consequently, unless the impact-generated temperatures have been overestimated (“cold start” scenario; Hauri et al., 2015), transfer of water from both the Earth and the impactor to the Moon can be ruled out. As an alternative, volatile-rich material might have been delivered to the lunar magma ocean (LMO) after the Moonforming impact, before the solidification of the insolating lunar crust. The deuterium/hydrogen ratio indicates that lunar water (i.e., δ DMoon ≈ −100 to +200h; Füri et al., 2014; Saal et al., 2013; Tartèse and Anand, 2012; Tartèse et al., 2013) is isotopically comparable to water trapped in carbonaceous chondrites (δ DCC ≈ −175 to +300h; Robert, 2003). This suggests that the accretion of chondritic matter provided water to an initially dry Moon. Sim-

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Table 1 Sample description and previously published cosmic ray exposure (CRE) ages. Sample ID

Description

CRE age (Ma)

References

10020,223 10057,183

Fine-grained, low-K, high-Ti ilmenite basalt Fine-grained, high-K, high-Ti ilmenite basalt

122–132 34–58

12008,48 14053,241

Low-K, low-Ti, olivine vitrophyre basalt Al-rich, low-Ti basalt

50 16–26

14321,1318 15414,212

Clast-rich breccia Coarse-grained ferroan anorthosite (98 vol.% plagioclase) Coarse-grained, low-Ti, olivine-normative basalt

22–26 80–120

Guggisberg et al. (1979) Guggisberg et al. (1979); Hintenberger et al. (1971); Hohenberg et al. (1970); Marti et al. (1970) Stettler et al. (1973) Eugster et al. (1984); Husain et al. (1972); Stettler et al. (1973); Turner et al. (1971) Burnett et al. (1972), Turner et al. (1971) Eugster et al. (1984); Husain et al. (1972); Stettler et al. (1973) Husain et al. (1972); Marti and Lightner (1972); Podosek et al. (1972); York et al. (1972)

Fine-grained, high-Ti ilmenite-basalt Coarse-grained, high-Ti ilmenite basalt Fine-grained, high-Ti ilmenite basalt Vesicular, coarse-grained, high-Ti ilmenite basalt

310 n.d. n.d. 16–18

Schaeffer and Schaeffer (1977)

Fine-grained, high-Ti ilmenite basalt

25–43

Eberhardt et al. (1974); Eugster et al. (1977)

15555,876 15555,950 15555,955 15555,958 70255,46 71557,13 71576,12 74255,180 74255,185 74275,240

ilarly, water and other volatile elements (e.g., C, N) trapped in Earth’s mantle have been proposed to originate from the addition of a chondritic ‘late veneer’ (Albarède et al., 2013) or from the accretion of a few ‘wet’ chondritic planetesimals (Marty, 2012; Morbidelli et al., 2000). According to the post-giant-impact volatile accretion scenario, the Moon might also have acquired a significant amount of nitrogen with a chondritic isotope signature, due to the fact that carbonaceous chondrites are marked by high abundances of nitrogen (i.e., ∼1000 ppm; Kerridge, 1985). Igneous rocks recovered at the surface of the Moon during the Apollo missions contain nitrogen only at the (sub-)ppm level (Becker et al., 1976; DesMarais, 1978, 1983; Funkhouser et al., 1971; Mathew and Marti, 2001; Mortimer et al., 2015; Müller et al., 1976), and previous attempts at determining the isotopic signature of indigenous lunar nitrogen have been unsuccessful in many instances. At such low concentrations, analyses are prone to terrestrial atmospheric contamination. In addition, any contribution of nitrogen implanted by solar-wind (SW) irradiation (Füri et al., 2012; Wieler, 2002) and/or of cosmogenic 15 N produced insitu, by the interaction of high-energy galactic cosmic rays and low-energy solar-flare protons with oxygen in the mineral lattice during space exposure (e.g., Clayton et al., 1977; Mathew and Murty, 1993), will mask the indigenous nitrogen isotope composition. From the study of a lunar breccia with a short cosmic ray exposure (CRE) age, Becker et al. (1976) determined the δ 15 N value of indigenous nitrogen to be about +10h (where δ 15 N is the permil deviation from the atmospheric 15 N/14 N ratio). Murty and Goswami (1992) investigated nitrogen in the lunar meteorite MAC88105 and obtained a δ 15 N value of +17 ± 3.4h, whereas Kerridge et al. (1991) studied LVGs from the Apollo 17 double drive tube 74001/74002, and interpreted nitrogen trapped on the surfaces of the glass beads (δ 15 N ≈ +14h) to be representative of the elusive indigenous component. A similar average value of +13 ± 1.2h was proposed by Mathew and Marti (2001), who considered a variety of samples including LVGs, a mare basalt, and anorthosites. More recently, Mortimer et al. (2015) argued for a lower δ 15 N value of +0.35h, determined by high-resolution stepped-combustion of a suite of six powdered lunar basalts. In summary, previous results indicate that indigenous lunar nitrogen is isotopically heavier than nitrogen in Earth’s primordial mantle (δ 15 Nmantle ≤ −40 to −5h; e.g., Cartigny and Marty, 2013; Javoy et al., 1986; Mohapatra et al., 2009; Palot et al., 2012). These findings appear inconsistent with lunar nitrogen being directly inherited from proto-Earth mantle material, and, instead, agree with

70–90

Eugster et al. (1977); Mörgeli et al. (1977)

the notion that 15 N-rich chondritic matter provided volatiles to the Moon and Earth’s surface following the giant impact. In order to improve our understanding of the origin of nitrogen in lunar basalts and, ultimately, of the abundance and isotopic composition of nitrogen in the Moon’s interior, we investigate the coupled nitrogen and noble gas (Ne–Ar) signatures of various lunar rocks, using the ultra-low blank laser-extraction technique developed at CRPG (Hashizume and Marty, 2004; Humbert et al., 2000). Since solar and cosmogenic noble gas components in extraterrestrial matter are characterized by distinct isotope and/or elemental abundance ratios, we employ neon and argon to resolve measured volatile abundances into constituent components, and thus, to assess the extent of sample exposure. In this way, we are able to identify samples that are least affected by SW-derived and cosmogenic nitrogen contributions, which allows us to constrain the concentration and isotopic signature of any existing indigenous nitrogen. Finally, we explore whether the chondritic projectiles that apparently delivered the Moon’s water also represent a suitable source for lunar nitrogen. 2. Samples and analytical techniques Nitrogen and noble gas (Ne and Ar) abundances and isotope ratios of twelve Apollo rocks (i.e., 10 mare basalts, 1 breccia, and 1 anorthosite; Table 1) were determined by CO2 laser-extraction, static-mass spectrometry at CRPG (Hashizume and Marty, 2004; Humbert et al., 2000) during two analytical sessions in March and August 2014. Small rock fragments or mineral separates, between 0.8 and 3.6 mg in mass (Supplementary Table S1), were heated with a continuous-mode infrared CO2 laser (λ = 10.6 μm). Modulating the power of the laser and monitoring the heating procedure using a camera (Humbert et al., 2000) permitted application of several temperature steps. Fusion was achieved in the third heating step, and 14 sub-samples were reheated after being melted (heating steps 4 to 5) to ensure complete gas extraction. The extracted gas fraction was split into two calibrated volumes for specific noble gas and nitrogen purifications. Neon was inlet into the VG5400 mass spectrometer first, and the gas was exposed to a charcoal finger at 77 K and a SAESTM AP-10 getter during analysis to minimize the contribution of doubly charged 40 Ar and CO2 to the 20 Ne and 22 Ne signals, respectively. The 40 Ar++ contribution to 20 Ne was reduced to virtually zero by measuring the signal at a slightly higher mass (+0.005 amu) than the peak center. The CO++ signal 2 of all samples was comparable to the mass-spectrometer back-

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197

Fig. 1. Three-isotope plots of neon extracted by step-wise heating from a) mare basalts containing a small fraction of SW-derived and/or adsorbed atmospheric noble gases released during the first heating step; b) mare basalts containing a large amount of solar noble gases; c) mare basalts and mineral separates that are dominated by cosmogenic noble gases (i.e., 21 Ne/22 Necosm ≈ 0.85; Leya et al., 2001); and d) highland breccia 14321 and anorthosite 15414. The neon isotope compositions of modern solar wind (SW; Heber et al., 2009) and Earth’s atmosphere (Air; Ozima and Podosek, 2002) are shown for comparison. Error bars represent 1σ uncertainties.

ground, and the contribution of CO++ was corrected as part of the 2 blank. With the laser off, blanks averaged 2.0 × 10−16 mol 20 Ne and 2.5 × 10−17 mol 36 Ar in March, and 1.9 × 10−16 mol 20 Ne and 1.6 × 10−17 mol 36 Ar in August. Procedural blanks (laser on) may be higher due to the release of noble gases from the container walls during sample heating; however, several extracted gas fractions are comparable within a factor of ≤2 to the measured blanks, arguing against a significant underestimation of the true blank values. Furthermore, heating steps for which the blank contribution to the measured noble gas abundances represents ≥25% are not considered for further discussion. The nitrogen aliquot was purified in a glass line, using a CuO furnace cycled between 723 and 1073 K and a U-shaped cold trap held at 93 K. Nitrogen isotope data were collected on the VG5400 mass spectrometer using a Faraday cup for masses 28 (14 N14 N) and 29 (15 N14 N), and an electron multiplier for masses 29 (15 N14 N) and 30 (15 N15 N) (see Hashizume and Marty, 2004 for details). Measured blanks averaged 6.5 × 10−13 mol N2 and 7.6 × 10−13 mol N2 in March and August, respectively. Isotope ratios of nitrogen (15 N/14 N) are expressed in the delta (δ ) notation, where δ 15 N = [(15 N/14 N)sample /(15 N/14 N)std − 1] × 1000, in h, and the standard is atmospheric N2 with 15 N/14 N = 0.003676. All nitrogen data have been corrected for blank contributions, CO and C2 Hx interferences, and instrumental mass fractionation. Corrected 28/29 ratios are, within error, identical to the measured values, implying that any organic contribution (e.g., from the release of hydrocarbons) to the measured signals is negligible. 3. Results Neon, argon, and nitrogen isotope and abundance results for 21 sub-samples are reported in Supplementary Table S1. Samples analyzed in March and August are designated by the superscripts ‘a’ and ‘b’, respectively. The data include replicate analyses of sev-

eral samples, as well as analyses of mineral separates (15555,876 olivine and pyroxene, and 15414,212 plagioclase). 3.1. Neon and argon Bulk fragments of mare basalts 12008, 14053, 70255, 71557, 71576b , 74255, and 74275 display similar Ne–Ar release patterns. Neon and argon extracted during the first heating step show 20 Ne/22 Ne and 36 Ar/38 Ar ratios above the cosmogenic endmember, interpreted to reflect the release of a small fraction of SWderived noble gases and/or adsorbed air, whereas noble gases released at higher temperatures are cosmogenic (i.e., 21 Ne/22 Necosm ≈ 0.85; Leya et al., 2001; Fig. 1a). In contrast, mare basalt samples 10057,183 and 71576,12a show much higher 20 Ne/22 Ne ratios of 12.86 and 11.85 (Fig. 1b), as well as 36 Ar/38 Ar ratios of 5.25 and 5.27, in the first heating step. These values are comparable to the isotopic composition of SW noble gases (20 Ne/22 NeSW = 13.78 ± 0.03 and 36 Ar/38 ArSW = 5.47 ± 0.01; Heber et al., 2009), and imply that the large amount of 20 Ne and 36 Ar in these samples (Supplementary Table S1) can be explained by the presence of solar gases. The solar gas component is depleted in light isotopes compared to modern SW, as observed in previous studies of lunar samples (e.g., Füri et al., 2014), due to depth-dependent isotope fractionation upon implantation of SW and removal of nearsurface-sited solar gas by ion sputtering (Grimberg et al., 2006; Wieler et al., 2007), and/or preferential diffusive loss of light noble gas isotopes (Eberhardt et al., 1970; Signer et al., 1977). With progressive heating of samples 10057,183 and 71576,12a , the release of solar 20 Ne and 36 Ar decreases (Supplementary Table S1), and the neon and argon isotope ratios trend towards the isotope signature of the cosmogenic endmember (Fig. 1b). Mare basalt 10020 and all sub-samples of mare basalt 15555 (876/950/955/958) record cosmogenic neon and argon isotope ratios throughout all heating steps, and similar values are also obtained for olivine and pyrox-

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ene separates of 15555,876 (Fig. 1c). Significantly, replicates generally show comparable noble gas abundances; however, sample 15555,958 released a considerably lower amount of argon compared to the other three 15555 sub-samples, consistent with the chemical heterogeneity previously observed in small-sized subsamples of this coarse-grained basalt (Ryder and Schuraytz, 2001). Highland breccia 14321 and anorthosite 15414 show low concentrations of 20 Ne and 36 Ar, and both neon and argon are characterized by cosmogenic isotope compositions (Fig. 1d and Supplementary Table S1). However, the breccia records an unusually high 40 Ar/36 Ar ratio of 15 200 to 20 400. Since the breccia contains the highest abundance of potassium of all the samples studied here (i.e., [K] ≤ 0.47 wt.%; Turner et al., 1971), the elevated 40 Ar content can be explained by in-situ production from the decay of 40 K. Assuming a formation age of 3.95 Ga (Turner et al., 1971), radioactive decay generated ∼12 000 × 10−12 mol 40 Ar∗ /g, which is in good agreement with the measured abundances of 12 500 to 14 400 × 10−12 mol 40 Ar/g. 3.2. Nitrogen Nitrogen concentrations in solar-gas-free samples range from 2.59 to 17.8 × 10−9 mol N2 /g (equivalent to between <0.1 and 0.5 ppm N; Supplementary Table S1). These values are comparable to the abundances previously observed in igneous lunar rocks (Becker et al., 1976; DesMarais, 1978, 1983; Funkhouser et al., 1971; Mathew and Marti, 2001; Mortimer et al., 2015; Müller et al., 1976). Higher nitrogen abundances are found in basalt samples 10057,183 (0.8 ppm) and 71576,12 (1.4–2.1 ppm), which record a contribution of solar noble gases, as well as in breccia 14321 (0.7–1.2 ppm). The olivine fraction of mare basalt 15555,876 contains a comparable amount of nitrogen (0.7 ppm), suggesting that olivine represents an important carrier phase of nitrogen in lunar basalts. While a few samples appear to release a small amount of adsorbed atmospheric noble gases in the first heating step (Fig. 1a), the δ 15 N value of the first gas fraction is distinct, in most cases, from the isotopic composition of air (δ 15 Nair = 0h), as well as of gaseous N2 used in the Apollo sample storage cabinets (δ 15 NN2 = −1.2h; R. Zeigler, pers. comm.). The δ 15 N value of nitrogen extracted by progressive heating increases significantly for most mare basalts and the anorthosite and reaches values of up to ≥1000h upon sample melting (Fig. 2), due to the release of cosmogenic 15 N at high temperatures (Hashizume et al., 2002). Thus, the δ 15 N pattern demonstrates that nitrogen released in the low-temperature steps is least affected by a 15 Ncosm contribution. Mare basalts 71576 and 74255, as well as breccia 14321, show relatively small δ 15 N variations (Figs. 2a and 2c), indicating that the relative proportion of cosmogenic 15 N is smaller in these samples.

Fig. 2. Evolution of the δ 15 N value with progressive heating from a) mare basalts, b) various sub-samples and mineral fractions (olivine and pyroxene) of mare basalt 15555, and c) breccia 14321, anorthosite 15414, and plagioclase from 15414.

4. Discussion Indigenous nitrogen originating from the Moon’s mantle may be masked in igneous rocks collected at the lunar surface by implanted SW-derived N and ‘planetary’ (meteoritic or cometary) N, as well as by cosmogenic 15 N produced in-situ by spallation reactions during exposure to cosmic ray irradiation for millions of years. While stepped-heating extraction techniques allow for a first-order resolution of the nitrogen content into constituent components (i.e., surface-sited, indigenous, cosmogenic; DesMarais, 1978, 1983; Mathew and Marti, 2001; Mortimer et al., 2015), coupled nitrogen–noble gas analyses provide a means to quantify the contribution of trapped (solar and planetary) volatiles and cosmogenic nuclides. Thus, the extent of sample exposure can be assessed before constraining the abundance and isotopic signature of any indigenous lunar nitrogen.

4.1. Cosmogenic neon and argon and cosmic ray exposure (CRE) ages All rocks returned by the Apollo missions have been exposed to cosmic ray irradiation at the Moon’s surface for millions of years. Consequently, their initial 21 Ne/22 Ne and 38 Ar/36 Ar isotope ratios will have increased with time due to the production of cosmogenic nuclides (e.g., 21 Ne and 38 Ar) by high-energy galactic (GCR) and solar cosmic rays (SCR), which can penetrate lunar matter to depths of up to several meters and ∼2 cm, respectively (e.g., Wieler, 2002). The absolute abundance of cosmogenic noble gases depends on the duration of space exposure and the nuclide production rate, which itself is a function of the chemical composition and the shielding depth.

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Table 2 Cosmogenic

21

Ne and

38

Ar abundances, and estimated proportions of

15

199

Ncosm .

Sample

21

Necosm (10−12 mol/g)

38

Arcosm (10−12 mol/g)

21

Ne/38 Ar

10020,223a 10057,183b 12008,48a 14053,241b 14321,1318a 14321,1318b 15555,876b 15555,876b olivine 15555,876b pyroxene 15555,950b 15555,955b 15555,958b 70255,46a 71557,13a 71576,12a 71576,12b 74255,180b 74255,185b 74275,240b 15414,212b 15414,212b plagioclase

5.69 2.91 4.51 1.34 1.15 1.13 4.08 7.65 6.42 4.21 3.62 4.25 13.87 8.13 5.58 5.27 0.80 0.94 1.52 1.70 3.48

6.28 3.10 2.78 0.92 1.13 0.86 2.32 0.94 2.61 2.90 1.88 0.47 15.35 7.28 6.57 5.71 0.65 0.78 1.53 4.84 8.55

0.9 0.9 1.6 1.5 1.0 1.3 1.8 8.2 2.5 1.5 1.9 9.0 0.9 1.1 0.8 0.9 1.2 1.2 1.0 0.4 0.4

15

Ncosm c (%) 26–41 4–10 57–95 31–75 3–6 2–4 55–100 10–18 23–42 40–79 25–47 44–85 82–100 37–54 6–9 8–12 5–9 3–6 n.d. (>100) 43–86

a

Analyzed in March 2014. Analyzed in August 2014. The proportion of 15 Ncosm in each sample is estimated using a production rate ( P 15 ) of (5.4 ± 0.9) × 10−12 g 15 N (g rock)−1 Ma−1 (Mathew and Murty, 1993) together with the published CRE ages given in Table 1. For samples 71557 and 71576, the CRE ages of 164 ± 18 and 122 ± 9 Ma, respectively, derived in this study are used (see Section 4.4). Uncertainties on CRE ages (where estimated), 14 N concentrations, and 15 N/14 N ratios are considered. b

c

Given that, in most cases, the neon isotope signatures observed in this study agree with the cosmogenic endmember composition (Fig. 1), the amount of 21 Necosm represents 100% of the total 21 Ne abundance (Table 2). Mare basalts 10057,183 and 71576,12a contain 71% and 86% 21 Ne of cosmogenic origin, respectively, assuming a simple two-component mixture of cosmogenic and solar neon (see Füri et al., 2014 for details on the component deconvolution). It should be noted that 70 to 95% of the total 21 Necosm content in mafic rocks and minerals was released at the third heating step (Fig. 3a), confirming that cosmogenic noble gases are efficiently extracted only during melting. Fig. 4a shows that there is a clear linear correlation between the 21 Necosm contents of the basalt and breccia samples and previously determined CRE ages (Table 1). The data are consistent with an average 21 Ne production rate ( P 21 ) of 4.6 × 10−14 mol (g rock)−1 Ma−1 , in good agreement with the GCR production rate of 4.5 to 5.4 (×10−14 ) mol (g rock)−1 Ma−1 calculated using the 2π exposure model from Leya et al. (2001) for low shielding and an average mare basalt composition (Warren and Taylor, 2014). Anorthosite 15414 is expected to be depleted in 21 Necosm because of its lack in the major target elements Na and Mg, whereas the small (≤2 cm) mare basalt 12008 may contain an excess of 21 Necosm produced by SCRs. Based on the cosmogenic 21 Ne contents, the CRE ages of mare basalts 71557 and 71576 – which had hitherto not been determined – are estimated to be of the order of 182 Ma and 117–125 Ma, respectively (Fig. 4a). Very similar CRE ages of 174 ± 16 Ma and 127 ± 10 Ma are obtained for basalts 71557 and 71576, respectively, by directly comparing their 21 Ne contents with the production rates derived from Leya et al. (2001) for the major element compositions given in the Apollo 17 sample catalogue. However, these two small-sized rake samples might record a contribution of 21 Ne produced by SCRs; therefore, the CRE ages derived here must be regarded as an upper limit of the true exposure age. Similarly, the amount of cosmogenic 38 Ar – which is predominantly produced by spallation from the target element Ca – represents 100% of the total 38 Ar content of the majority of the samples studied here (Table 2). Mare basalt samples 10057,183 and 71576, 12a contain ∼65 and ∼47% 38 Ar of solar origin, re-

Fig. 3. Release fraction of cosmogenic a) 21 Ne and b) 38 Ar with progressive heating from various sub-samples and mineral separates (olivine and pyroxene) of mare basalt 15555. Between 83 and 95% of the total amount of 21 Necosm , and 81 to 95% of the cosmogenic 38 Ar content, was released during melting (heating step 3). 21 Ne and 38 Ar show similar release patterns, ruling out significant isotope fractionation during step-wise heating.

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Fig. 4. Abundances of a) 21 Necosm and b) 38 Arcosm in the mare basalts, breccia 14321, and anorthosite 15414 as a function of previously published cosmic ray exposure (CRE) ages (Table 1). The observed correlations are consistent with production rates of 4.6 × 10−14 mol 21 Ne (g rock)−1 Ma−1 and 5.0 × 10−14 mol 38 Ar (g rock)−1 Ma−1 , i.e., P 21 / P 38 = 0.9. CRE ages of 146–182 and 114–131 Ma are derived for samples 71557 and 71576, respectively.

spectively, and mare basalts 71576,12b and 74275,240 appear to contain 5 to 10% solar argon as well. Again, assuming published CRE ages to be correct (Table 1), an average 38 Ar production rate ( P 38 ) of 5.0 × 10−14 mol (g rock)−1 Ma−1 provides a good fit to the data (Fig. 4b). This value is comparable to the nominal value of 4.9 × 10−14 mol (g rock)−1 Ma−1 (i.e., 1.4 × 10−8 cm3 STP (g Ca)−1 Ma−1 ) from Turner et al. (1971) for a CaO content of 11 wt.%, which is typical for lunar basalts (Warren and Taylor, 2014). Basalt 15555 is characterized by a quite variable but lower CaO content of only ∼9 wt.% (Ryder and Schuraytz, 2001, and references therein), which may explain its depletion in 38 Arcosm . Based on the observed relationship, the cosmogenic 38 Ar content of mare basalts 71557 and 71576 can be explained by exposure to GCRs over a period of 146 Ma and 114–131 Ma, respectively, which is in good agreement with the 21 Ne CRE age estimates. 4.2. Solar noble gases and nitrogen in basalts 10057,183 and 71576,12 SW-derived volatiles are implanted into the top few tens of nanometers of lunar material. Hence, only fragments derived from the uppermost surface of lunar rocks are expected to contain solar noble gases. In this study, two mare basalt samples have clearly been identified to contain a large amount of solar neon and argon. Basalt 71576 is a small (size: ∼3 cm, mass: 23.5 g) rake sample. Sub-sample 71576,12 is a 1 g fraction of the specimen, and is therefore likely to contain surface-sited noble gases. In contrast, sample 10057,183 is from a depth of ∼2 cm within the original

11 cm sized rock. The presence of solar neon and argon in the studied fragment is enigmatic, but could possibly be attributed to a contribution of SW-loaded dust, introduced into the sample interior during preparation. Alternatively, basalt 10057 may contain ‘primordial’ solar-like noble gases derived from the lunar mantle, if the Moon trapped solar volatiles during its formation, as has been inferred for Earth (e.g., Marty, 2012). This interpretation, however, remains highly speculative because the Moon is depleted in volatile elements, and indigenous noble gases have not been detected in lunar rocks to this date. Since sample 71576,12 has presumably been exposed to SW irradiation, it is inferred to contain SW-derived nitrogen as well. The relative abundances of solar neon and argon detected in 71576,12a imply that the 20 Ne/36 Ar elemental ratio of the implanted solar gas is highly fractionated (20 Ne/36 Artr ≈ 8.9) compared to the signature of modern solar wind (20 Ne/36 ArSW = 42.1; Heber et al., 2009). Similar 20 Ne/36 Artr values have been found in previous studies of lunar soils and LVGs (Eberhardt et al., 1970; Eugster et al., 1980). These low values can be explained by depthdependent mass fractionation upon implantation of SW, due to the higher energies and greater penetration depths of heavier isotopes (Grimberg et al., 2006), combined with removal of near-surfacesited SW gas by ion sputtering (Wieler et al., 2007), and/or by preferential diffusive loss of light noble gas isotopes (Eberhardt et al., 1970; Signer et al., 1977). Compared to the inert noble gases, solar nitrogen might be preferentially retained because it can form chemical compounds such as ammonium and nitride by reaction with hydrogen (DesMarais et al., 1974; Müller, 1974); in addition, it may react with cations such as silicon and titanium containing free radicals (Hashizume et al., 2002). As a consequence, the 14 N/20 Ne and 14 N/36 Ar elemental ratios of the trapped solar component are expected to be higher than the SW values (14 N/20 NeSW ≈ 1 and 14 N/36 ArSW ≈ 28 to 37; Anders and Grevesse, 1989; Marty, 2012). Indeed, based on analyses of single lunar soil grains, Füri et al. (2012) and Hashizume et al. (2002) argued for a 14 N/36 Ar ratio up to 5 times greater than the SW value. Nonetheless, for fragment 71576,12a , solar 14 N is estimated to represent only <4% of the total amount of 14 N or <8% of the 14 N release at low temperatures (with 36 Artr = 36 Arm –36 Arcosm ≈ 31 × 10−12 mol/g and 14 N/36 Artr ≤ 185). Such a small proportion of implanted solar nitrogen appears consistent with the observation that the N isotope composition of this basalt falls into the range of values recorded by solar-gas-free mare basalts, even though SW nitrogen is characterized by a significantly lower isotope signature of −407 ± 7h (Marty et al., 2011), and any implantation fractionation would further lower the δ 15 N value. Solar nitrogen must therefore be masked by ‘planetary’ and/or indigenous nitrogen. 4.3. ‘Planetary’ nitrogen in basalt 71576,12 and breccia 14321 The isotopic composition of nitrogen trapped in low-antiquity lunar soils is best explained by a continuous addition of chondritic micro-meteorites (Füri et al., 2012; Hashizume et al., 2002; Wieler et al., 1999), which record a mean δ 15 N value of +22h (Marty et al., 2005). This influx results in the presence of 1.5 to 2 wt.% carbonaceous chondrite-like material in the lunar regolith (Keays et al., 1970). Due to interactions with the solar wind and the electromagnetic field, incoming ‘planetary’ nitrogen can be photo-ionized and implanted onto grain surfaces activated by sputtering at the lunar surface (Manka and Michel, 1970, 1971). An addition of planetary nitrogen must be considered for the mare basalt sub-sample 71576,12, because the Ne–Ar data demonstrate that it has been directly exposed to the lunar surface environment. Since ∼95% of the surface-sited (solar) noble gases are released at low temperatures, the δ 15 N value of the first two heating steps may provide insight into the origin of the implanted planetary component. The δ 15 N

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values range from −1.7 to +43.9h (Fig. 2a and Supplementary Table S1), comparable to the nitrogen isotope signature of chondritic (micro-)meteorites (δ 15 NCC = −45h to +170h; Kerridge, 1985; Kung and Clayton, 1978; Pearson et al., 2006). Therefore, since the contribution of solar nitrogen has been found to be negligible (Section 4.2), any excess nitrogen implanted onto the surface of mare basalt 71576,12 is likely derived from a chondritic source. Addition of terrestrial atmospheric nitrogen by the Earth wind (Ozima et al., 2005) can be ruled out, because the basalt has only been exposed at the lunar surface within the past ∼130 Ma (Section 4.1), when the global geomagnetic field prevented a significant N+ escape flux from Earth’s atmosphere. Breccia 14321 contains more nitrogen than the solar-gas-free mare basalts. Siderophile element concentrations indicate an addition of meteoritic material to 14321 (Morgan et al., 1975); therefore, the breccia might also contain meteoritic nitrogen, possibly in vesicles within shock-produced glass (Turner et al., 1971). The δ 15 N values vary between −7.6 to +34.4h in the first two heating steps, and agree with the range of values recorded by carbonaceous chondrites. Therefore, nitrogen in this sample appears to originate from the impacting chondritic body that caused the brecciation. The same process could also explain the large amount of nitrogen (1.3 to 1.9 ppm) observed in breccias 60025, 67915, and 68815 (Becker et al., 1976; Mathew and Marti, 2001). The δ 15 N values of +10 to 13h in these brecciated highland rocks might therefore reflect a planetary, rather than an indigenous, nitrogen component. In contrast to breccias and regolith samples, mare basalts are strongly depleted in siderophile elements (e.g., iridium; Warren and Taylor, 2014), ruling out a contamination by ‘planetary’ material. 4.4. Cosmogenic nitrogen (15 Ncosm ) The 15 N/14 N ratio of all samples analyzed in this study must have increased during space exposure due to the production of cosmogenic 15 N through the 16 O(p, 2p)15 N and 16 O(p, pn) 15 O spallation reactions (e.g., Clayton et al., 1977; Mathew and Murty, 1993). Since oxygen is the major target for the production of 15 N, and its abundance varies little in silicate minerals, the cosmogenic 15 N production rate ( P 15 ) is relatively insensitive to variations in rock mineralogy. Empirical estimates of P 15 , either derived from bulk-rock and soil analyses ([3.6 ± 0.8] × 10−12 g 15 N (g rock)−1 Ma−1 (Becker et al., 1976); 4.1 × 10−12 g 15 N (g rock)−1 Ma−1 (DesMarais, 1983); [5.8 ± 0.6] × 10−12 g 15 N (g rock)−1 Ma−1 (Mathew and Marti, 2001)) or from single ilmenite grain analysis ([4.7 ± 0.9] × 10−12 g 15 N (g rock)−1 Ma−1 (Hashizume et al., 2002)), are in reasonably good agreement, especially when considering the variety of samples studied and their potentially different shielding histories. These estimates are also comparable to the theoretically calculated value of (5.4 ± 0.9) × 10−12 g 15 N (g rock)−1 Ma−1 at low shielding (i.e., 0 to ∼40 g/cm2 ) from Mathew and Murty (1993). Furthermore, assuming that the 15 N content of mare basalt 70255, which records the longest CRE age of 310 Ma, is dominated by 15 Ncosm , a similar cosmogenic 15 N production rate of (4.9 ± 0.4) × 10−12 g 15 N (g rock)−1 Ma−1 can be derived here. Using the theoretically calculated P 15 value from Mathew and Murty (1993), together with the previously determined CRE ages given in Table 1 (and the CRE ages of 164 ± 18 and 122 ± 9 Ma for samples 71557 and 71576, respectively, determined in this study), the fraction of cosmogenic 15 N is estimated to range from 2 to 100% for the samples analyzed here (Table 2). This implies that the 15 N content of the majority of the mare basalts is inconsistent with a purely cosmogenic origin, and, instead, requires the presence of an indigenous nitrogen component. The lunar highland anorthosite 15414, on the other hand, contains virtually no indigenous nitrogen.

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4.5. Isotopic composition of indigenous lunar nitrogen With the exception of the mare basalt fragments 10057,183 and 71576,12, as well as breccia 14321, the samples analyzed here can be assumed to contain a simple two-component mixture of cosmogenic 15 N and indigenous nitrogen. Cosmogenic 15 N is predominantly released during sample melting (Fig. 2), whereas indigenous nitrogen is released at temperatures between ∼600 and 1100 ◦ C (Mathew and Marti, 2001; Mortimer et al., 2015). Therefore, the δ 15 N values observed at the first two heating steps are used to constrain the isotopic signature of indigenous lunar nitrogen. Nitrogen extracted at low temperatures from the mare basalts is characterized by a variable isotopic composition, ranging from −45.8 to +147.1h (or up to +444.3h in sample 71557,13; Supplementary Table S1). Elevated δ 15 N values of > +100h are only observed in samples that record a large proportion of 15 Ncosm (e.g., 12008 and 70255; Table 2) or a significant release of cosmogenic noble gas isotopes during the first two heating steps (e.g., 30% in 71557); therefore, these values are not representative of the indigenous isotopic signature. Mare basalt 74255,180 and the olivine fraction of 15555,876 show δ 15 N values of −0.2 to +26.7h (Supplementary Table S1). Since these samples record the smallest proportion of cosmogenic nitrogen out of all the solar- and planetarygas-free samples (Table 2), together with a small release (≤10%) of cosmogenic noble gas isotopes at low temperatures, their δ 15 N signature is thought to well represent the isotopic composition of the indigenous component. These results are in excellent agreement with the range of δ 15 N values (i.e., ∼0 to +20h) previously proposed for lunar nitrogen (Becker et al., 1976; Kerridge et al., 1991; Mathew and Marti, 2001; Mortimer et al., 2015; Murty and Goswami, 1992). However, several samples (14053 and 15555,876/955/958) record significantly lower δ 15 N values. Since these basalts contain no traces of solar noble gases, and any release of 15 Ncosm would act to increase the nitrogen isotope ratio, their light nitrogen isotope signatures possibly indicate that nitrogen in the mare basalt source(s) is isotopically heterogeneous. In summary, the current study provides further evidence for the presence of a small amount (<1 ppm) of indigenous nitrogen in lunar mare basalts, with δ 15 N values varying between −46 and +27h. Since nitrogen isotopes do not significantly fractionate during degassing (Cartigny et al., 2001), these values are assumed to closely approximate the nitrogen isotopic composition of the mare basalt mantle source(s). Thus, lunar nitrogen encompasses the range of isotope signatures recorded by Earth’s (relic) primordial mantle (δ 15 Nmantle ≤ −40 to −5h; e.g., Cartigny and Marty, 2013; Javoy et al., 1986; Mohapatra et al., 2009; Palot et al., 2012), enstatite chondrites (δ 15 NEC = −30 ± 10h; Kung and Clayton, 1978), and carbonaceous chondrites (δ 15 NCC,mean = +20 ± 20h; e.g., Kerridge, 1985; Marty, 2012). 4.6. Abundance and origin of nitrogen in the Earth–Moon system Accurately assessing the Moon’s nitrogen inventory is extremely difficult due to the lack of primitive mantle-derived samples. Furthermore, the degree of partial melting of the mare basalt source(s), the extent of melt degassing, as well as the behavior of nitrogen during these processes, are not well-constrained. Under oxidizing conditions, nitrogen dissolves as a N2 molecule into cavities of the silicate network, whereas under the reducing conditions of the lunar mantle ( f O2 ≈ IW − 2 to IW; Wadhwa, 2008), the nitrogen solubility in basaltic melts increases as nitrogen dissolves as N3− and becomes chemically bonded with atoms of the silicate melt network (Li et al., 2013, 2015; Libourel et al., 2003). Due to the more-reducing conditions, lunar melts are therefore predicted to preferentially retain dissolved nitrogen

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Fig. 5. Hydrogen (D/H) and nitrogen (15 N/14 N) isotope ratios of the Moon (this study; Becker et al., 1976; Füri et al., 2014; Kerridge et al., 1991; Mathew and Marti, 2001; Mortimer et al., 2015; Murty and Goswami, 1992; Saal et al., 2013; Tartèse and Anand, 2012; Tartèse et al., 2013), carbonaceous chondrites (Kerridge, 1985), Earth’s surface, and Earth’s mantle (Cartigny and Marty, 2013; Clog et al., 2013; Deloule et al., 1991; Mohapatra et al., 2009; Palot et al., 2012). Adapted from Marty (2012).

compared to terrestrial mantle melts. However, the concentration of indigenous nitrogen in lunar mare basalts (<1 ppm) is comparable to that of terrestrial mid-ocean ridge and ocean-island basalts (Marty and Dauphas, 2003; Marty and Humbert, 1997; Marty and Zimmermann, 1999). Provided that lunar and terrestrial melts experienced a comparable degree of partial melting and degassing, this either implies that the abundance of nitrogen in the mare basalt source region(s) is lower than in Earth’s mantle, or that nitrogen is efficiently retained in the residual solids of the lunar mantle. Thus, while at least some parts of the lunar mantle must contain indigenous nitrogen, the nitrogen content of the mare basalt source region(s) cannot be definitely reconstructed at this stage. Nonetheless, the 15 N/14 N ratio can be used to trace the origin of the Moon’s nitrogen. Fig. 5 illustrates that hydrogen isotope ratios of lunar samples overlap with the terrestrial signature (Füri et al., 2014; Saal et al., 2013; Tartèse and Anand, 2012), allowing for ‘water’ trapped in the Moon’s interior to be inherited from the proto-Earth in a “cold start” scenario (Hauri et al., 2015; Tartèse et al., 2014). Similarly, the low nitrogen isotope ratios observed in some mare basalts resemble that of Earth’s primordial mantle, which has been proposed to retain isotopically light nitrogen from the protosolar nebula (Marty, 2012) or from a contribution of enstatite chondrite-type material (Javoy et al., 1986). Moreover, an impactor with a composition similar to enstatite chondrites (Meier et al., 2014) would also be characterized by δ 15 N values of −30 ± 10h (Kung and Clayton, 1978). Thus, 15 N-depleted nitrogen in the Moon might originate from the proto-Earth and/or the impactor; since nitrogen is relatively refractory under reducing conditions, it could be efficiently retained within the proto-lunar magma disk. However, many lunar basalts, as well as Earth’s surface reservoirs (crust, sediments, and the atmosphere) and present-day mantle, show higher δ 15 N values than the relic primordial mantle (e.g., Cartigny and Marty, 2013). Thus, the Earth–Moon system requires a contribution of 15 N-enriched material. Indeed, the overall abundance patterns and isotopic compositions of nitrogen and other volatile elements (H, C, noble gases) in the bulk Earth can be explained by the addition of a chondritic ‘late veneer’ (Albarède et al., 2013) or by the accretion of a few ‘wet’ chondritic planetesimals (Marty, 2012). These chondritic projectiles also represent a suitable source for lunar volatiles since the Moon’s hydrogen and nitrogen

isotope signatures fall into the range of values observed in carbonaceous chondrites (Fig. 5). Thus, chondritic volatiles might have been delivered to Earth and the LMO after the Moon-forming impact, before the solidification of the insolating lunar crust ∼4.4 Ga ago (Gaffney and Borg, 2014; Norman et al., 2003). Comets record a 3-fold 15 N enrichment compared to lunar and terrestrial nitrogen (δ 15 Ncomet ≈ +1000h; Füri and Marty, 2015, and references therein), and can therefore be ruled out as major contributors to the nitrogen budget of the Earth–Moon system. The abundance of other elements that were added by the influx of chondritic matter allows to place limits on the lunar nitrogen inventory. For example, the amount of highly siderophile elements (HSEs) requires that 1.3 to 3.5 × 1016 tons of chondritic material was admixed to the lunar mantle (Bottke et al., 2010; Day et al., 2007; Schlichting et al., 2012). If these chondritic projectiles that delivered most of the Moon’s HSEs also provided it with nitrogen, the maximum nitrogen influx is of the order of 3.5 × 1013 tons, assuming a nitrogen concentration of 1000 ppm for carbonaceous chondrites (Kerridge, 1985). Thus, the average nitrogen content of a 400 km-deep LMO (Shearer et al., 2006) is estimated to be ≤1 ppm – similar to the abundance of indigenous nitrogen observed in lunar mare basalts and comparable to the nitrogen content of the bulk Earth (1.6 ppm; Marty, 2012). Alternatively, in order to explain the postulated water concentration of 100 to 300 ppm in the LMO (Elkins-Tanton and Grove, 2011; Hui et al., 2013), between 0.25 and 4.2 × 1017 tons of chondritic material would be required (Füri et al., 2014; Hauri et al., 2015; Tartèse and Anand, 2012, 2014). In this case, the LMO may have acquired up to ∼10 ppm nitrogen. The post-giant-impact volatile accretion scenario requires a mechanism for efficient trapping or ingassing of nitrogen released upon impacts of chondritic projectiles on the Moon. Prior to the formation of the anorthositic crust, incoming objects were directly assimilated by the LMO. For oxygen fugacities close to or below the iron-wustite buffer, nitrogen of chondritic origin can be dissolved into, and retained by, the magma ocean melt (Li et al., 2013, 2015; Libourel et al., 2003). Nitrogen was then transported to greater depths by convection, where, at higher temperatures and pressures, it could have been incorporated into crystallizing olivine, pyroxene, and garnet (Li et al., 2013). Here, olivine has been identified as a carrier phase of indigenous lunar nitrogen, and experimental studies have revealed that terrestrial upper mantle minerals (forsterite, diopside, enstatite, pyrope) can store large amounts of nitrogen (i.e., tens or up to hundreds of ppm; Li et al., 2013), likely because nitrogen is incorporated in the 3-oxidation state as ammonia (NH3 ) or ammonium (NH+ 4 ). In addition, terrestrial ilmenite has been shown to contain high nitrogen abundances, possibly because nitrogen replaces oxygen in its crystal structure (Madiba et al., 1998). Consequently, the mafic mare basalt source region(s) in the Moon’s interior may be characterized by a significantly large nitrogen storage capacity. In contrast, the pristine highland anorthosite 15414 studied here shows virtually no traces of 14 N, suggesting that nitrogen was not efficiently incorporated into the lunar crust during crystallization of the LMO. After the LMO solidification, upon impacting the solid crust, chondritic meteorites are (partially) vaporized. Hashimoto et al. (2007), Schaefer and Fegley (2010), and Yabuta et al. (2014) argued that impacts of carbonaceous chondrites on planetary surfaces generate vapor plumes containing nitrogen predominantly in the form of N2 , together with a high abundance of reduced H–C–S gas species. Thus, equilibration between an impact-generated melt and a reducing vapor cloud with a N2 partial pressure of ≤10 bar (Libourel et al., 2003) can potentially result in the N abundances of ∼1 ppm observed in shocked or brecciated anorthosites by Becker et al. (1976) and Mathew and Marti (2001). Consequently, this scenario could explain why anorthositic samples that contain

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shock-produced glass – such as 68815 (Becker et al., 1976), 60255 and 67951 (Mathew and Marti, 2001), as well as highland breccia 14321 (this study) – hold nitrogen with a chondritic isotope composition. 5. Conclusions Stepped-heating methods allow for a first-order resolution of the nitrogen content of lunar samples into constituent components (i.e., surface-sited, indigenous, cosmogenic). However, great care must be taken when interpreting the nitrogen systematics of extraterrestrial samples whose exposure histories are not well constrained. Only by coupling nitrogen analyses with noble gas isotope measurements can samples affected by an ‘exogenic’ (atmospheric, solar and/or planetary) nitrogen contribution be identified. In this study, two mare basalt samples have been found to contain a large amount of solar neon and argon; therefore, these samples are likely to contain surface-sited solar and/or planetary nitrogen. Since all other mare basalts are dominated by cosmogenic noble gases, nitrogen can be assumed to represent a twocomponent mixture between cosmogenic 15 N and indigenous nitrogen. The isotopic composition of indigenous nitrogen extracted at low temperatures varies between −45.8 and +26.7h, possibly suggesting that nitrogen in the Moon’s interior is isotopically heterogeneous. The lowest δ 15 N values are comparable to those of Earth’s relic primordial mantle, and they are consistent with an enstatite chondrite-like impactor. Thus, nitrogen could have been transferred from both the Earth and the impactor to the Moon. Nonetheless, a contribution of 15 N-rich material to the Earth–Moon system is required to explain the δ 15 N values ≥ 0h observed in many terrestrial and lunar samples. The data are consistent with a post-giant-impact volatile accretion scenario whereby the LMO trapped a few ppm nitrogen from the assimilation of carbonaceous chondrites, and the reducing conditions during LMO crystallization favored the incorporation of nitrogen into mafic minerals (olivine, pyroxene, garnet). In this way, the mare basalt source region(s) acquired a significant amount of nitrogen originating from a chondritic reservoir. In contrast, pristine highland rocks such as anorthosite 15414 may be 14 N-free. Only brecciated anorthosites may be able to trap chondritic nitrogen within shock-produced melt by equilibration with reducing vapor plumes generated by impacts of chondritic projectiles on the Moon. Acknowledgements This work was supported by the European Research Council under the European Community’s Seventh Framework Programme (FP7/2007–2013 grant agreement no. 267255 to BM), by the CNRSINSU Programme de Planétologie (EF), and by the NASA Cosmochemistry Grant NNX11AG58G (LAT). We thank three anonymous reviewers for thorough reviews, and C. Sotin for editorial handling. This is CRPG-CNRS contribution 2400. Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2015.09.022. References Albarède, F., Albalat, E., Lee, C.-T.A., 2014. An intrinsic volatility scale relevant to the Earth and Moon and the status of water in the Moon. Meteorit. Planet. Sci. http://dx.doi.org/10.1111/maps.12331. Albarède, F., Ballhaus, C., Blichert-Toft, J., Lee, C.-T., Marty, B., Moynier, F., Yin, Q.-Z., 2013. Asteroidal impacts and the origin of terrestrial and lunar volatiles. Icarus 222, 44–52. http://dx.doi.org/10.1016/j.icarus.2012.10.026.

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