Oxygen isotope identity of the Earth and Moon with implications for the formation of the Moon and source of volatiles
Edited by Mark Thiemens, University of California San Diego, La Jolla, CA; received January 20, 2024; accepted May 8, 2024
Significance
The isotopic similarity of Earth and Moon rocks is a significant puzzle in geo- and cosmochemistry, contradicting prevailing Moon formation models. Our data show an isotopic match at a sub-ppm level with no heterogeneity within Earth's mantle and lunar samples. We explore these findings in relation to Moon formation models and the hypothesis of a collision with a planet that lost its silicate mantle due to large impacts. Additionally, our results suggest that water on both Earth and Moon originated from a well-mixed reservoir rather than arriving late. Our study highlights the importance of samples obtained through space missions.
Abstract
The Moon formed 4.5 Ga ago through a collision between proto-Earth and a planetesimal known as Theia. The compositional similarity of Earth and Moon puts tight limits on the isotopic contrast between Theia and proto-Earth, or it requires intense homogenization of Theia and proto-Earth material during and in the aftermath of the Moon-forming impact, or a combination of both. We conducted precise measurements of oxygen isotope ratios of lunar and terrestrial rocks. The absence of an isotopic difference between the Moon and Earth on the sub-ppm level, as well as the absence of isotope heterogeneity in Earth’s upper mantle and the Moon, is discussed in relation to published Moon formation scenarios and the collisional erosion of Theia’s silicate mantles prior to colliding with proto-Earth. The data provide valuable insights into the origin of volatiles in the Earth and Moon as they suggest that the water on the Earth may not have been delivered by the late veneer. The study also highlights the scientific value of samples returned by space missions, when compared to analyses of meteorite material, which may have interacted with terrestrial water.
A common model is that our Moon formed 4.5 Ga ago by collision of proto-Earth with a Mars-sized body called Theia (1–3).
The Moon is highly depleted by several orders of magnitude in volatile elements and water when compared to chondrites and the Earth mantle (4). The apparent absence of water in lunar rocks led to the formulation of a “bone-dry Moon” (5). The upper Earth mantle contains roughly 200 µg g–1 water with likely higher, but less well-constrained concentrations in the lower mantle (6). Based on chlorine isotope fractionation in lunar basalts, Sharp et al. (7) concluded that the water concentration in the lunar mantle is approximately 0.09 µg g–1, which agrees with the estimate of Albarède of ≤1 µg g–1. The depletion of the Moon in volatile elements has two sources: i) the volatile element depletion of the building blocks of the Moon (proto-Earth, Theia) by nebular fractional condensation and ii) volatile loss during the formation of the Moon.
The paradigm of a dry Moon had been challenged by measurements of the hydrogen concentration (OH, H2O) in lunar volcanic glasses and olivine melt inclusions (8, 9), anorthosite (10), and lunar apatite (Ca5[PO4]3[OH, F, Cl]) (11–13). These studies come to the conclusion that the Moon may have a water content as high as ~50 to 250 µg g–1, which would be comparable to the Earth mantle and question the paradigm of a “bone dry” Moon.
The question about the water content of the Moon relates to the nature of the material that the Moon formed from. In the canonical model, Theia is depicted as a Mars-sized (MT ~ 0.1 M⨁) body that collided with the Earth slightly off-axis (14). An outcome of collision modeling is that Moon silicates contain a much higher mass fraction XT of Theia than Earth rocks (14). According to these models, the Moon contains up to 70 wt.% Theia, whereas the Earth contains only ~10 wt.% Theia. The suggested predominance of Theia as the source of lunar rocks bears one of the major conundrums in cosmochemistry [Isotope Crisis, (15)]. If the Moon contains 70 wt.% Theia and Earth only 10 wt.%, and if Theia was compositionally different from proto-Earth and had a different differentiation history (e.g., core formation age), one would expect a distinct compositional (chemical, isotopic) contrast between Earth and Moon rocks. Many elements (O, Si, S, Ti, Cr, and W) show distinct, partly mass-independent isotope heterogeneity among solar system bodies (planets and asteroids) but show no (O, Si, S, Ti, and Cr) or only very small (O and W) differences between the Earth and Moon (16, 17).
For the fingerprinting of different solar system reservoirs, the measurement of oxygen isotope ratios [17O/16O, 18O/16O; expressed in the form of δ17O, δ18O (18), and Δ’17O, Eq. 1] is a powerful tool. Oxygen is a major component in all rocky material of the solar system and shows large mass-dependent and mass-independent isotope variations (19). The large natural isotopic variations can be measured with very high analytical resolution. For Δ’17O, isotope variations among asteroids, planets, and the Moon span as much as ~7,500 ppm (19). These large variations can now be analyzed with a <20 ppm (2σ SD) resolution, making oxygen isotopes a sensitive tracer for reservoir mixing among solar system bodies.
If the Moon and Earth have different mixing ratios of Theia and proto-Earth, in general, an oxygen isotope difference is expected. The current situation regarding the Δ’17O of the Earth and Moon is not entirely clear. Some studies have suggested a small, mass-independent isotope difference between the Earth and Moon (20–22), whereas other studies did not resolve any difference (23–28). Also, oxygen isotope heterogeneity within lunar lithologies had been suggested (22, 29).
Two routes have been paved out of the Isotope Crisis, i.e., the isotopic identity of the Earth and Moon: i) proto-Earth and Theia were isotopically similar, and ii) proto-Earth and Theia were isotopically dissimilar, but the fraction of Theia is identical in Moon and Earth silicates.
The isotopic similarity of Theia and proto-Earth has been discussed in ref. 16 and had earlier been considered to explain the oxygen isotope similarity of the Earth and Moon (23, 30). For tungsten isotopes, not only sampling of the same reservoir, but also a similar accretion and differentiation history is required to explain the isotope similarity of Earth and Moon (31, 32) and had been suggested to be the result of happenstance (16). More recently, Olsen and Sharp (33) developed an accretion and differentiation model, which explains the tungsten isotope similarity of the Earth and Moon in light of a pebble accretion. Such a mode of accretion of Theia and proto-Earth in the same feeding zone may also explain the similarity of the Earth and Moon in oxygen, titanium, and chromium.
Several processes have been proposed to resolve the Isotope Crisis by a similar fraction XT of Theia in Earth and Moon silicates. Either a similar distribution of Theia in the Moon and proto-Earth or vigorous mixing in the aftermath of the impact would explain the isotope similarity even if Theia and proto-Earth accreted isotopically different material. An endmember would be the formation of the Moon by fission from the Earth mantle (34). In such a scenario, the fraction of Theia XT in the Earth and Moon would be identical, i.e., zero, and the Moon would entirely be made from Earth mantle material. Such a scenario would well explain the isotopic and chemical identity of the Earth and Moon. However, the fission theory has been dismissed on the basis of dynamic considerations (35), but a modified impact-triggered fission model was suggested by Čuk and Stewart (36). In their model, a fast-spinning proto-Earth was hit by small, highly energetic Theia(s). These impacts would have triggered the fission and would lead to a Moon that is predominantly made of Earth mantle material. An alternative is that Theia and proto-Earth were isotopically distinct, but that intense mixing during and in the aftermath of the giant Moon-forming impact led to the homogenization of the Earth–Moon system. Pahlevan and Stevenson (37) found solutions in their impact models that led to intense mixing of Theia and proto-Earth material and thus could explain the isotope identity even if Theia and proto-Earth were isotopically dissimilar. They found solutions for oxygen but did not consider refractory tungsten or titanium. An extreme case for intense mixing of Theia and proto-Earth before reaccretion of the Earth and accretion of the Moon has been suggested by Lock et al. (38), where the Earth and Moon formed from a homogenized donut-shaped vapor cloud (synestia) and hence have identical isotope compositions. Canup (39) modeled the collision of proto-Earth with Theia of similar mass. Their calculations suggest that for such a case, Earth and Theia material would mix well and similar compositions of lunar and Earth mantle rocks would be the outcome.
Here, we present high-precision oxygen isotope data for a diverse suite of lunar rocks and compare their compositions with isotope data for the terrestrial mantle. The data from this study are compared to reevaluated published data, and results are discussed with respect to implications on Moon forming models and the volatile inventory of the Earth and Moon.
1. Results
We have obtained data from 14 different lunar samples along with 169 analyses of terrestrial San Carlos olivine and 22 analyses of terrestrial UWG2 garnet that were measured in the same sessions as the lunar samples. Uncertainties are given as a 2σ SEM. A mean 2σ SD of 9 ppm was obtained from the replicates of San Carlos olivine (2σ SD = 9.2 ppm, N = 169) and UWG2 garnet (2σ SD = 11.8 ppm, N = 22). All data and metadata are listed in SI Appendix, Table S1 (online at the Göttingen Research Data repository at https://doi.org/10.25625/MZGCSQ).
Where data were available, we used the concentration of iridium or other highly siderophile elements to distinguish pristine from impactor contaminated lunar rocks (see SI Appendix for details). In this study, only pristine, i.e., not impactor-contaminated rocks were considered. Pristine lunar rocks are isotopically homogenous with a mean Δ’17O☾ = −51.4 ± 1.4 ppm (2σ SEM). This datum is identical to the composition of the present day Earth mantle with Δ’17O⨁ = −51.6 ± 1.0 ppm (40) and San Carlos olivine with Δ’17OSCO = −51.8 ± 0.8 ppm, which is a good approximation for the composition of the Earth upper mantle. Our high-precision data show that the Earth and Moon are within ΔΔ’17O⨁−☾ = 0.2 ± 1.6 ppm isotopically identical. No oxygen isotope heterogeneity between lunar mare basalts and volcanic glasses or anorthosites have been resolved (Fig. 1).
Fig. 1.
2. Discussion
The data from this study are compared to previous studies with high-precision triple oxygen isotope data (20–26). In order to ensure interlaboratory comparability, it is important to consider which standards and which reference line had been used in respective studies. Miller et al. (41) introduced a set of silicate samples that include the commonly used standards San Carlos olivine and UWG2 garnet (42). It has been shown by Greenwood et al. (21) and Peters et al. (40) that the Δ’17O of San Carlos olivine is representative for the Δ’17O of the Earth upper mantle (represented by residual mantle olivine and partial mantle melts). Miller et al. (41) showed that ΔΔ’17OSCO-UWG2 is to 8 ± 4 ppm. In this study, a difference of ΔΔ’17OSCO–UWG2 = 8.4 ± 1.4 ppm has been measured, confirming previous data (41). We anchor the Δ’17O of UWG2 garnet to San Carlos olivine (43) with Δ’17OUWG2 = –51.8 ppm – 8.4 ppm = –60.2 ppm. Previous high-precision triple oxygen isotope studies of lunar rocks used either UWG2 garnet (23–25), San Carlos olivine (21), or both (20, 26) as reference point(s) for the composition of the Earth. Along with information about the slope and intercept of the reference line for Δ’17O, this allows reevaluation of published data. The reevaluated literature data are illustrated along with data from this study in Fig. 2.
Fig. 2.
In line with the average of previous studies (20–26) but at the sub-ppm level, data from this study suggest that the Moon and Earth are isotopically identical with ΔΔ’17O☾–⨁ = 0.2 ± 1.6 ppm. We do not resolve any isotope heterogeneity between lunar basalt and anorthosites (29) or between lunar basalts and volcanic glasses (22) (Fig. 1). The absence of any intra-Moon Δ’17O heterogeneity is in line with high-T mantle fractionation with little fractionation in δ18O and associated high-T triple isotope exponents close to 0.528 to 0.529 (45).
The data from this study deviate from the data published by Herwartz et al. (20) (after reevaluation; Fig. 2). For this study, a new laser fluorination line was built, and the improvements are described in the Materials and Methods section and SI Appendix. The exact cause for the discrepancy, however, could not be resolved.
The rocks that were brought back by the Apollo astronauts were all collected from the lunar surface. Contamination of lunar rocks by solar wind oxygen hence may have to be considered when doing high-precision Δ’17O studies. A result of the Genesis mission is that implanted solar wind has a Δ’17O ≈ –28,000 ppm (46). The O/H ratio of the Sun is 4.3 × 10–4 (47). Volcanic glass beads on the lunar surface have reported solar wind implanted hydrogen concentrations up to 220 µg g–1 (2.2 × 10–4 mol H g–1; 2,000 µg g–1 H2O) (48). Assuming the absence of elemental fractionation, these glasses would be contaminated by 9.6 × 10–8 mol g–1 solar wind oxygen. Considering the host rock contains 45 wt.% oxygen, the contamination by solar wind would shift the Δ’17O by only 0.1 ppm. Although solar wind implantation may contribute even to a planet’s water budget (49), effects on Δ’17O of bulk lunar rocks are small and currently below the analytical resolution. The discrepancies between this and previous studies (20, 22, 29) hence are likely not due to solar wind contamination of samples.
We assume that the Δ’17O of the studied low-iridium lunar basalts, volcanic glasses, and anorthosites represents the composition of the lunar mantle. This assumption is based on the observation of Greenwood et al. (21), who measured Δ’17O for a large variety of basalts from the Earth (N = 19; 5.8 ≤ δ18O ≤ 6.3‰) and reported, within error, identical Δ’17O with ΔΔ’17Obasalt–SCO = –2 ± 4 ppm. No difference in Δ’17O is observed between mare basalt, volcanic glasses, and anorthosites (Fig. 1). Another argument for lunar volcanic rocks to be representative for the Δ’17O of the lunar mantle is that they formed at high-T in the absence of liquid water. The lunar samples are all ultramafic (volcanic glasses) and mafic (basalt, anorthosite) rocks formed by different degrees of partial melting of the ultramafic lunar mantle in combination with fractional crystallization and/or assimilation. All these igneous processes occurred at high temperatures where fractionation in δ18O is ≤1.5‰ and where the associated θ is 0.528 to 0.529 (45). Such processes hence do not shift the Δ’17O if fractionation in δ18O remains small.
Our dataset includes samples from various landing sites (Apollo 11, 12, 15, 16, 17), covering a significant portion of the lunar surface with no discernible difference in Δ’17O. Therefore, we conclude that the samples studied are the most reliable proxy for the Δ’17O of the bulk silicate Moon. However, not only for oxygen isotope studies, but also for sampling lunar mantle rocks should be a primary focus of future lunar sample return missions (50). Previous examinations of lunar meteorites, which could expand the range of locations, have revealed that terrestrial alteration (i.e., exchange with water) can pose a significant challenge for oxygen isotopes (20). As a result, such samples were not included in this study, underscoring the importance of sample-return missions.
The isotope identity of the Earth and Moon had earlier been suggested (23–26) but is now observed on the previously unresolved <1 ppm scale (Fig. 1). This considerably restricts the degree of the primordial isotope contrast between Theia and proto-Earth, as well as the heterogeneous distribution of Theia between the Earth and Moon (Fig. 3).
Fig. 3.
The isotope identity of Moon and Earth, as illustrated in diagram Fig. 3 requires either Theia and proto-Earth having the same composition (16), vigorous mixing of Theia and Earth after the Moon-forming impact (37–39), or a combination of both.
Although the striking chemical (51) and isotopic (16) similarities of the Earth and Moon can be explained by similar compositions and/or complete mixing of Theia and proto-Earth, it would also be explained by the formation of the Moon from Earth material alone. Darwin (34) suggested that the Moon formed by fission. Such a formation scenario would explain the isotopic and chemical similarity of the Earth and Moon but had been dismissed on the basis of dynamic considerations (35).
Here, we discuss a variant of a model, where the Moon is formed entirely from Earth mantle material. In this model, the Moon formed by a collision of proto-Earth with Theia had lost its silicate mantle due to collisional erosion. Such a scenario would resolve the isotope crisis (15) for all lithophile elements, including oxygen, and would also explain the pre-late-veneer tungsten isotope similarity of the Earth and Moon (52). The selective stripping of the outer silicate mantle from differentiated planetesimals, asteroids, and planets is a common process during the formation of terrestrial planets (53). It was reasoned by O’Neill and Palme (54) that the existence of ~70 different iron meteorite parent bodies attests that the stripping of planetesimal mantles was, indeed, common in the early solar system (53, 55, 56). Such pure metallic planetary embryos can form, but these are typically very small (<0.001 M⨁) (56). Such small metallic bodies would be suitable parent bodies for iron meteorites, but too small for Theia.
The high density of Mercury is explained to be the result of collisional erosion and loss of silicate (57). O’Neill and Palme (54) suggested that even the Earth lost a fraction of its mantle through collisional erosion. The removal of mantle silicate by collisional erosion becomes more difficult as the size of the body increases (53). It was shown in numerical simulations that high-energy collisions considerably increase the formation rate of daughter bodies with high metal/silicate ratio such as Mercury or the iron meteorite parent bodies (53). If impact velocities vi exceed ~2 to 4 times the escape velocities ve of the target, erosion is a likely outcome of many impacts (58). High relative velocities are more likely if giant planet migration occurred (Grand tack model) than in calm accretion scenarios (59). We suggest that such high-energy collisions during the gravitational disturbances caused by the giant planet migration may have been responsible for the removal of the silicate mantle of Theia. Because complete mantle stripping is easier for smaller planetesimals (53), a metal-rich Theia may also be explained by accretion from smaller embryos that already had lost their mantles.
The idea of a largely metallic Theia impactor, however, is particularly difficult to prove. A hint toward a largely metallic Theia could be the oversized Earth core, which is estimated to be ~6 to 10% larger than expected (54, 60). We speculate that the oversized core reflects the addition of a largely metallic Theia instead of being the result of collisional erosion of the Earth mantle (54) or being the consequence of collisional erosion during accretion (55). Venus, which is similar in size than the Earth and may have shared a similar accretionary history, does not show indication for an oversized core (Fig. 4) and thus may support the idea that the Earth oversized core is due to addition of a largely metallic Theia. However, the uncertainty intrinsic to the estimate of Venus’ core is still too large to draw a final conclusion here.
Fig. 4.
The same high-energy collisions that have stripped the mantle off Theia and/or its embryos, however, would also facilitate the postcollisional chemical and isotopic equilibration between the Earth mantle and the material that the Moon accreted from refs. 37 and 38. High-energy collisions furthermore allow smaller sizes for Theia, which would have facilitated not only the removal of the mantle and postdepositional homogenization but would also make a very similar initial isotope composition of Theia and proto-Earth more probable (26). Which one of the three processes or which combination of them led to the isotope identity of the Earth and Moon, however, remains difficult to prove.
Studies have identified high concentrations of hydrogen in lunar ultramafic volcanic glasses (8, 9). Saal et al. (8) suggested 745 ppm H2O in source melt. Assuming 20% partial melting, one gets about 150 ppm in the lunar mantle; this is in the same order of magnitude suggested from water content in melt inclusions (9) and would imply similar water content in the Moon than in the Earth upper mantle (6). Other studies, however, argue that the water content of the lunar mantle is ≤1 ppm (4, 7), which implies that the lunar mantle is heterogenous with respect to volatile elements. Our data do not resolve any difference in Δ’17O between green and orange lunar volcanic glasses and other lunar lithologies (Fig. 1). If the source region of the melts was enriched in volatiles by addition of water-rich CM chondrite material [~10 wt.% H2O, (61); Δ’17O = –2,500 ppm, (62)], the Δ’17O would be ~4 ppm lower than the bulk Moon, which would be difficult to resolve. Contamination by CI type chondrites [~28 wt.% H2O, (61); Δ’17O = 200 ppm, (62)] would result in a 0.1 ppm shift in Δ’17O and would not be resolved in our dataset. In order to identify the source of water in the volcanic glasses, a higher analytical resolution of <1 ppm in Δ’17O is required.
It was reasoned by Greenwood et al. (21) and Peters et al. (40) that oxygen isotopes constrain the origin of volatiles added by the last accreting material (late veneer) (63). The Earth is suggested to have accreted ~0.5 wt.% late veneer, while the Moon collected only ~0.02 wt.% (64) and is thus regarded as representing the pre-late-veneer Earth composition (21, 40). This discrepancy in late veneer component can be explained by a few large impacts instead of a rain of small meteorites (65). The estimates for the amount of late veneer of ~0.5 wt.% are based on platinum group element concentrations. A fraction of the platinum group elements may have merged with Earth’s core, and Earth’s mantle may actually contain as much as ~4 wt.% late veneer component (66).
The absence of an oxygen isotope difference between the Moon and Earth (this study; previous studies, Fig. 2) and between pre- and post-late-veneer Earth mantle (40) requires that the late veneer component was isotopically similar to the Earth–Moon system. Among the known meteorites, only the addition of a ~0.5 wt.% enstatite chondrite component is compatible with the observed ΔΔ’17O☾-⨁ = 0.2 ± 1.6 ppm. Enstatite chondrites have ΔΔ’17OEC-⨁ ≈ 50 ± 20 ppm (20), and late accretion of such material would only lead to a 0.25 ppm Δ’17O increase of Earth Δ’17O (0.5 wt.% late veneer) and a 0.01 ppm increase of the Moon (0.02 wt.% late veneer).
Piani et al. (67) conducted a study on the water content of enstatite chondrites, which have isotopic similarities to the Earth and the Moon. The researchers measured hydrogen concentrations and found that enstatite chondrites alone could potentially account for the water on the Earth, without the need for an additional late accreting source of water (66). This suggestion is in line with the conclusion based on oxygen isotopes that the majority of Earth’s water had accreted before the Moon formation and the late veneer (21). Our data support the idea that Earth’s water accreted early (21) and the possibility of enstatite chondrites being a source of water for the Earth (67).
The identical Δ’17O isotopic composition of the pre- and post-late-veneer Earth (40) and the Moon (this study) puts tight constraints on the type of meteorites dominating the late veneer, as its Δ’17O must have been very similar to proto-Earth. We conclude that not only the proto-Earth and Theia but also the undifferentiated meteorite class dominating the late veneer likely originated from the same planetary feeding zone with similar Δ’17O.
3. Materials and Methods
3.1. Sampling.
A total of 14 pristine lunar rocks from the NASA Apollo program were made available and analyzed for this study. The lunar samples (low-, high-Ti mare basalt, volcanic glasses, highland rocks) were bracketed with terrestrial samples (San Carlos olivine, UWG2 garnet). The classification of lunar samples as “pristine” (i.e., not impact contaminated) was based on the Ir concentration of the respective samples (for details, see also SI Appendix).
3.2. Oxygen Isotope Analyses.
All samples were analyzed by means of laser fluorination (68). For this study, a new laser fluorination line was built. Key improvements of the extraction line compared to the extraction line used in the study by Herwartz et al. (20) include a fully automated fluorination reaction by laser heating, automated and improved cleaning of the sample O2 gas (e.g., removal of N2 from sample O2), an automated introduction of the sample into the mass spectrometer, automated sample and reference gas pressure adjustment during dual-inlet measurement (SI Appendix), careful monitoring of contaminant species in the sample gas (SI Appendix), lower overall blanks (e.g., from leaks), higher number of standards to sample ratios, and matching amounts of oxygen extracted from lunar and terrestrial samples irrespective of the lithology. In the new line, we used BrF5 as fluorination agent, whereas Herwartz et al. (20) used F2.
After loading the samples (~2 mg each; amount adjusted according to the oxygen concentration in the samples) in a 14-pit Ni sample holder, samples were heated overnight in vacuum. Subsequently, samples were gently premelted in vacuum to reduce the sample surface. The lunar samples were bracketed with San Carlos olivine standards in a ratio >1. The base vacuum in the line was in the range of 4 × 10–6 mbar. Residual moisture in the sample chamber was removed by multiple prefluorinations by admitting ~50 to 100 mbar BrF5 several times for 10 to 30 min before starting fluorination of samples. After admitting ~100 to 130 mbar BrF5 to the sample chamber, the ~2.5 mm (inner diameter) sample pit was multiply scanned with the infrared laser with stepwise increase in laser power up to 45 W. Excess BrF5, SiF4, and other condensable gases were trapped in a first U-trap (trap #6) at –196 °C. The gas was then conducted through a NaCl trap (160 °C) followed by a second cold trap (trap #5) for removal of Cl2 that was also held at –196 °C. The O2 was then collected in a 4 mm (outer diameter) U-trap (trap #2) that is filled with 5 Å molecular sieve pellets and held at –196 °C. For separation of impurities, O2 gas was then released from trap #2 by heating and conducted with a 10 mL min–1 He carrier gas stream through a 3 m 1/8” packed molecular sieve column (Restek) held at 50 °C. Only the eluting O2 (separated from N2 and other impurities) was then trapped for ~1,100 s on a second 13× molecular sieve trap (trap #8) at –196 °C in front of the mass spectrometer. The separation was monitored using a Pfeiffer QMS220 quadrupole mass spectrometer. The He was then evacuated from trap #8 via the conventional Thermo MAT253Plus gas source mass spectrometer and the O2 was released into the bellows of the mass spectrometer by heating trap #8 to 50 °C using a temperature-controlled water bath. The measurements were conducted in dual-inlet mode at an intensity of 5,000 mV on the Faraday cup used for measuring the 16O2+ ion beam. The Thermo Isotope Scripting Language (ISL) acquisition script was modified so that the pressure in the bellows was adjusted during the measurement to keep the intensity ±1% constant. Usually ~40 to 60 sample-reference cycles were acquired with an integration time of 26 s. After each measurement, the sample gas was scanned from mass 12 to 98 (see online data, https://doi.org/10.25625/MZGCSQ). The main contaminants were on mass 28 (residual N2+, possibly CO+ as CO2 fragment), mas 40 (40Ar+, not separated from O2 in the gas chromatograph), and mass 44 (CO2+, likely from the reaction of organics with oxygen in the source). All mass intensities are available online. The complete sample analysis took about 2 h. Including prefluorinations, the run of an entire sample holder took about 30 h, during which no human interaction was required. The entire procedure was automatized using National Instruments LabVIEW in combination with modified Thermo ISL scripts.
The oxygen isotope ratios are expressed relative to the Vienna Standard Mean Ocean Water (VSMOW) standard in the conventional δ notation (18). Small deviations in the triple oxygen isotope ratios are expressed as Δ’17O (in ppm) relative to a slope 0.528 reference line with zero intercept (Eq. 1).
[1]
Measured data were normalized to San Carlos of the respective sessions with Δ’17O = –51.8 ppm (43). The datum is an average of studies on the Δ’17O of San Carlos olivine relative to VSMOW water (69–71). The uncertainty in Δ’17O was 8 to 10 ppm (2σ, SD). For δ18O, we measured UWG2 garnet with a composition of 5.75‰ (42), giving a δ18O = 5.27‰ for San Carlos olivine. The difference in Δ’17O between UWG2 garnet and San Carlos olivine is –8 ppm (41). All individual data are listed in SI Appendix, Table S1, available online at Göttingen Research Online/Data https://doi.org/10.25625/MZGCSQ.
Data, Materials, and Software Availability
The measurement data and metadata are available at Göttingen Research Online/Data https://doi.org/10.25625/MZGCSQ (72). All additional data needed to evaluate the conclusions in the paper are present in the paper and/or SI Appendix.
Acknowledgments
Ryan Zeigler and the entire NASA Captem team are thanked for providing samples from the Apollo program. Dennis Kohl is thanked for his help in the course of the laboratory work. The constructive comments of three anonymous reviewers and the associated editor Mark Thiemens are highly acknowledged and helped to improve this manuscript. M.F. was financially supported by the International Max-Planck-Research School for Solar System Science that is based at the Max-Planck-Institute of Solar System Research (Göttingen).
Author contributions
M.F. and A.P. designed research; M.F., S.T.M.P., D.H., P.H., and T.D.R. performed research; M.F., S.T.M.P., D.H., and T.D.R. analyzed data; and M.F., S.T.M.P., D.H., and A.P. wrote the paper.
Competing interests
The authors declare no competing interest.
Supporting Information
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Copyright © 2024 the Author(s). Published by PNAS. This open access article is distributed under Creative Commons Attribution-NonCommercial-NoDerivatives License 4.0 (CC BY-NC-ND).
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The measurement data and metadata are available at Göttingen Research Online/Data https://doi.org/10.25625/MZGCSQ (72). All additional data needed to evaluate the conclusions in the paper are present in the paper and/or SI Appendix.
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Received: January 20, 2024
Accepted: May 8, 2024
Published online: December 16, 2024
Published in issue: December 24, 2024
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Acknowledgments
Ryan Zeigler and the entire NASA Captem team are thanked for providing samples from the Apollo program. Dennis Kohl is thanked for his help in the course of the laboratory work. The constructive comments of three anonymous reviewers and the associated editor Mark Thiemens are highly acknowledged and helped to improve this manuscript. M.F. was financially supported by the International Max-Planck-Research School for Solar System Science that is based at the Max-Planck-Institute of Solar System Research (Göttingen).
Author contributions
M.F. and A.P. designed research; M.F., S.T.M.P., D.H., P.H., and T.D.R. performed research; M.F., S.T.M.P., D.H., and T.D.R. analyzed data; and M.F., S.T.M.P., D.H., and A.P. wrote the paper.
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The authors declare no competing interest.
Notes
A.P. is an organizer of this Special Feature.
This article is a PNAS Direct Submission.
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Oxygen isotope identity of the Earth and Moon with implications for the formation of the Moon and source of volatiles, Proc. Natl. Acad. Sci. U.S.A.
121 (52) e2321070121,
https://doi.org/10.1073/pnas.2321070121
(2024).
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