Franck Poitrasson
email: franck.poitrasson@get.omp.eu

Thomas Zambardi

Tomas Magna

Clive R Neal

A reassessment of the iron isotope composition of the Moon and its implications for accretion and differentiation of terrestrial planets

Keywords: iron isotopes, Moon, mare basalts, anorthosites, Mg-suite, highland rocks

come

A reassessment of the iron isotope composition of the

Moon and its implications for the accretion and differentiation of terrestrial planets

Introduction

Investigation of the stable isotope compositions of planetary bodies other than the Earth has been facilitated through the return of lunar samples by the Apollo missions, as well as the recognition of Martian meteorites in existing collections [START_REF] Mcsween | Allan Hills 77005: A new meteorite type found in Antarctica[END_REF], and meteorites coming from the differentiated asteroid Vesta [START_REF] Consolmagno | Composition and evolution of the eucrite parent body: Evidence from rare earth elements[END_REF].

Following pioneering studies of oxygen, silicon and potassium isotope compositions (e.g. [START_REF] Clayton | Oxygen isotopic compositions of Apollo 15, 16 and 17 samples, and their bearing on lunar origin and petrogenesis[END_REF][START_REF] Clayton | Oxygen isotopes in eucrites, shergottites, nakhlites, and chassignites[END_REF][START_REF] Molini-Velsko | isotopic composition of silicon in meteorites[END_REF][START_REF] Humayun | Potassium isotope cosmochemistry: Genetic implications of volatile element depletion[END_REF], the past 15 years witnessed rapid improvements in our knowledge of stable isotope compositions of a number of major and trace elements in lunar and other planetary bodies (e.g., Mg: [START_REF] Wiechert | Non-chondritic magnesium and the origins of the inner terrestrial planets[END_REF][START_REF] Bourdon | Chondritic Mg isotope composition of the Earth[END_REF][START_REF] Teng | Magnesium isotopic composition of the Earth and chondrites[END_REF][START_REF] Hin | Magnesium isotope evidence that accretional vapour loss shapes planetary compositions[END_REF]; Si: [START_REF] Georg | Silicon in the Earth's core[END_REF][START_REF] Fitoussi | Si isotope systematics of meteorites and terrestrial peridotites: implications for Mg/Si fractionation in the solar nebula and for Si in the Earth's core[END_REF][START_REF] Ziegler | Metal-silicate silicon isotope fractionation in enstatite meteorites and constraints on Earth's core formation[END_REF][START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF]; Ca: [START_REF] Simon | Stable calcium isotopic composition of meteorites and rocky planets[END_REF]; Fe: [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Schoenberg | Modes of planetary-scale Fe isotope fractionation[END_REF][START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF]; Zn: [START_REF] Moynier | Isotopic composition of zinc, copper, and iron in lunar samples[END_REF][START_REF] Paniello | Zinc isotopic evidence for the origin of the Moon[END_REF][START_REF] Kato | Extensive volatile loss during formation and differentiation of the Moon[END_REF]. These results were sometimes controversial, either due to analytical difficulties and/or the use of older generation of mass spectrometers, less precise and accurate (e.g., Mg: [START_REF] Wiechert | Non-chondritic magnesium and the origins of the inner terrestrial planets[END_REF][START_REF] Bourdon | Chondritic Mg isotope composition of the Earth[END_REF][START_REF] Chakrabarti | The isotopic composition of magnesium in the inner Solar System[END_REF][START_REF] Teng | Magnesium isotopic composition of the Earth and chondrites[END_REF]; Si: [START_REF] Georg | Silicon in the Earth's core[END_REF][START_REF] Fitoussi | Si isotope systematics of meteorites and terrestrial peridotites: implications for Mg/Si fractionation in the solar nebula and for Si in the Earth's core[END_REF][START_REF] Savage | Silicon isotope homogeneity in the mantle[END_REF][START_REF] Ziegler | Metal-silicate silicon isotope fractionation in enstatite meteorites and constraints on Earth's core formation[END_REF][START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF], or resulting from different interpretations of mass-dependent stable isotope variations (e.g., Fe: [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Schoenberg | Modes of planetary-scale Fe isotope fractionation[END_REF][START_REF] Beard | Comment on "Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF]Weyer et al., 2007;[START_REF] Poitrasson | Probes of the ancient and the inaccessible[END_REF][START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF][START_REF] Rustad | Iron isotope fractionation in the Earth's lower mantle[END_REF][START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF][START_REF] Elardo | Non-chondritic iron isotope ratios in planetary mantles as a result of core formation[END_REF][START_REF] Elardo | The effect of core composition on iron isotope fractionation between planetary cores and mantles[END_REF]. For iron, this resulted in debates on estimating the bulk stable isotope composition of planets (see Fig. 1, for the Earth and the Moon). However, the interplanetary, mass dependent stable isotope variations hold great promise in unraveling early Solar System processes, such as planetary accretion mechanisms or mantle-core differentiation [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Georg | Silicon in the Earth's core[END_REF][START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF].

Ideally, these stable isotope tracers should be little affected by common geological processes to allow easy interplanetary comparison [START_REF] Humayun | Potassium isotope cosmochemistry: Genetic implications of volatile element depletion[END_REF]. However, for those stable isotopes affected by processes such as fractional crystallization or element diffusion in solids, their fractionation mechanisms should be well understood and fully quantified to be useful for planet formation studies. Unfortunately, this is not yet the case for many of the elements under consideration. Beyond general rules common to all stable isotopes, such as the effect of temperature on the magnitude of isotopic fractionation or the stiffness of inter-atomic bonds (e.g. [START_REF] Schauble | Applying stable isotope fractionation theory to new systems[END_REF], initial studies have shown that elements such as Fe, Mg or O will not fractionate in the same way because they are located in different crystallographic coordination configurations or may be differently affected by redox processes. For this reason, it is important to determine and understand the systematics of the stable isotope fractionation during geological processes for all elements of interest. Among these, iron has attracted considerable attention given its importance and large abundances in terrestrial planets, and the discovery of different planetary Fe isotope signatures [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF]. However, with the expanding database, more complexities than originally anticipated have been highlighted. For example, [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF] clearly showed that there is a difference in the Fe isotope composition of lunar low-Ti and high-Ti mare basalts that may directly mirror that of their respective mantle sources. This supports the observations of [START_REF] Snyder | A Chemical-Model for Generating the Sources of Mare Basalts -Combined Equilibrium and Fractional Crystallization of the Lunar Magmasphere[END_REF] and [START_REF] Beard | The source region and melting mineralogy of high-titanium and low-titanium lunar basalts deduced from Lu-Hf isotope data[END_REF] who implied distinct mantle sources of low-Ti and high-Ti mare basalts on the basis of petrogenetic modeling and radiogenic isotope systematics, respectively. If true, this may eventually lead to a different estimate for the bulk lunar Fe isotope composition. This finding is of notable importance as it may provide a test for interplanetary Fe isotope variations imparted by either mantle-core differentiation or partial vaporization in the aftermath of a Moon-forming giant impact [START_REF] Poitrasson | Probes of the ancient and the inaccessible[END_REF].

In order to provide new constraints on Fe isotope fractionation during magmatic evolution of the lunar magma ocean, we have determined the Fe isotope composition in a suite of mare basalts and highland rocks. This new data set should enable a better evaluation of potential differences between various lunar reservoirs and their significance in creating bulk planetary signatures previously debated in the literature [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Beard | Comment on "Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF]Weyer et al., 2007;[START_REF] Weyer | Partial melting and melt percolation in the mantle: The message from Fe isotopes[END_REF][START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF][START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF][START_REF] Elardo | Non-chondritic iron isotope ratios in planetary mantles as a result of core formation[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF][START_REF] Elardo | The effect of core composition on iron isotope fractionation between planetary cores and mantles[END_REF]. Furthermore, the new Fe isotope data for several specimens were collected in order to (i) verify earlier Fe isotope determinations by means of low mass resolution MC-ICP-MS [START_REF] Wiesli | Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith[END_REF][START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF], and (ii) constrain potential sample heterogeneity problems (e.g., for olivine-normative basalt 15555).

Samples

The samples were selected to reflect a wide range in petrological and geochemical types [START_REF] Neal | Petrogenesis of Mare Basalts -a Record of Lunar Volcanism[END_REF][START_REF] Warren | A concise compilation of petrologic information on possibly nonmare Moon rocks[END_REF][START_REF] Papike | Lunar samples[END_REF][START_REF] Neal | Interior of the Moon: The presence of garnet in the primitive deep lunar mantle[END_REF][START_REF] Meyer | Lunar sample compendium[END_REF], as well as in their location on the Moon as sampled by the Apollo missions. Nine low-Ti and 12 high-Ti mare basalts, one KREEP basalt and one KREEP-rich impact melt, five ferroan anorthosites, two Mg-suite norites, one Mg-suite troctolite, one dunite, and one Mg-suite noritic anorthosite were analyzed. Only pristine highland rocks were selected for the new analyses reported in the present work according to criteria given in [START_REF] Warren | A concise compilation of petrologic information on possibly nonmare Moon rocks[END_REF] defined on the basis of siderophile element contents and textural characteristics. This is important when considering that meteorite bombardment affects Fe isotope signatures on the lunar surface [START_REF] Wiesli | Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith[END_REF][START_REF] Moynier | Isotopic composition of zinc, copper, and iron in lunar samples[END_REF][START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF]. These new data are compared to all previously published bulk lunar rock Fe isotope determinations, making up a Fe isotope database of 73 lunar rock samples. Meteorites of lunar origin (see their Fe isotope data in [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF] were not included in this compilation. This is due to an uncertain location of their origin on the Moon, strong shock modification and potential weathering on the Earth's surface. Most lunar meteorites were recovered from hot deserts and it has previously been shown that weathering under these conditions can significantly affect the Fe isotope signatures of bulk meteorites [START_REF] Saunier | Effect of hot desert weathering on the bulk-iron isotope composition of L6 and H5 ordinary chondrites[END_REF].

Methods

Between 3 and 30 mg of bulk-rock powder prepared from 0.1 to one gram of the original sample chip allocation were digested (See Table 1). Samples processed at GET-CNRS, Toulouse were dissolved with a mixture of concentrated HF-HNO 3 -HCl in closed high-pressure Teflon vessels in an oven at 135 C for 6 days. For samples processed at the University of Lausanne, the same acid mixture was used and beakers were placed on a hotplate at 120°C for 3 days. The resulting solutions were subsequently evaporated to dryness and re-dissolved in 6M HCl on a hotplate at 120°C for 48 hours. Dried residues were thereafter dissolved using 6M HCl prior to Fe purification by conventional anion-exchange chromatography following established analytical protocols [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF](Poitrasson et al., , 2005)).

The total procedural blanks were between 3 and 7 ng, which is more than three orders of magnitude lower than Fe levels in samples and is thus considered negligible. Iron isotope compositions were determined using a Neptune (ThermoFisher Scientific, Bremen, Germany) high mass resolution multiple-collector inductively-coupled-plasma mass spectrometer (MC-ICP-MS) housed at CNRS Toulouse, following the procedures detailed in [START_REF] Poitrasson | Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS[END_REF], which included mass bias correction of the purified Fe samples by Ni doping. This approach accurately corrects for mass bias shifts caused by residual matrix effects and results in a superior reliability of the measurements over the conventional standard-sample bracketing method (see [START_REF] Poitrasson | Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS[END_REF].

The new Fe isotope compositions are reported in Table 1 using the standard delta notation, in per mil (‰), relative to the European isotopic IRMM-14 reference material.

Analytical reproducibility was estimated on the basis of 80 analyses of an in-house hematite standard from Milhas, Pyrénées (sometimes referred to in the literature as "ETH hematite standard"), conducted over 3 years and in the same analytical sequences as the sample analyses reported in Table 1. Every sample was typically analyzed six times; the long term reproducibility of such pooled measurements can be estimated on the basis of the hematite reference material analyses pooled in groups of six individual measurements. This yielded 57 Fe = 0.758 ± 0.067‰ and 56 Fe = 0.511 ± 0.046‰ (2SD). Additionally, BHVO-2 and JB-2 basalt reference materials analyzed during this study yielded 57 Fe of 0.182 ± 0.062‰ and 0.095 ± 0.082‰ (2SD), respectively, in good agreement with literature values [START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Craddock | Iron isotope compositions of reference material, geostandards and chondrites[END_REF].

In the reminder of this paper, the uncertainties of the sample set mean values are calculated as two standard errors of the mean (2SE) rather than two standard deviations (2SD) in order to indicate how well the mean values are constrained rather than to give a measure of the spread of a set of individual measurements around the mean (see e.g. [START_REF] Miller | Statistics for analytical chemistry[END_REF]. The 2SE calculation involves the use of the t-correcting factor that takes into account the number of samples used in the calculation [START_REF] Platzner | Modern isotope ratio mass spectrometry[END_REF]. Furthermore, these 2SE uncertainties represent a direct proxy of the Student's t-test (e.g., [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF]Poitrasson et al., , 2005;;[START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF]. All literature data have been recalculated in this way to be mutually comparable. However, irrespective of whether calculated as either 2SE or 2SD, it is important to note that these estimated uncertainties only include analytical random errors but not systematic ones potentially associated with analytical or sampling biases.

Results and comparison with previous data

Thirty three new Fe isotope analyses of bulk lunar rock samples are listed in Table 1.

The low-Ti basalts and high-Ti basalts show well-defined but mutually distinct 57 Fe ranges from 0.043 ± 0.023‰ to 0.154 ± 0.071‰ and from 0.167 ± 0.118‰ to 0.354 ± 0.118‰, respectively. The new Fe isotope data for low-Ti mare basalts confirm previous observations of homogeneous 57 Fe systematics, similar to Earth [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF]. Two high-K mare basalts (10049, 10057) from the Apollo 11 high-Ti group show systematically lower 57 Fe values (≤0.187‰) compared with the rest of the Apollo 11 suite (≥0.290‰) and the entire Apollo 17 suite (≥0.213‰). KREEP basalt 15386 has 57 Fe = 0.18 0.07‰, intermediate between low-Ti and high-Ti mare basalts.

The new highland data reveal a bulk rock 57 Fe range from -0.73 ± 0.06‰ to 0.50 ± 0.06‰ (Table 1). This >1.2‰ variation is surprisingly large for igneous rocks, and is only approached by terrestrial mantle peridotites (see [START_REF] Williams | Systematic iron isotope fractionation variations in mantle rocks and minerals: The effect of partial melting and oxygen fugacity[END_REF][START_REF] Weyer | Partial melting and melt percolation in the mantle: The message from Fe isotopes[END_REF][START_REF] Zhao | Iron isotope variations in spinel peridotite xenoliths from North China Craton: implications for mantle metasomatism[END_REF][START_REF] Poitrasson | On the iron isotope heterogeneity of lithospheric mantle xenoliths: implications for mantle metasomatism, the origin of basalts and the iron isotope composition of the Earth[END_REF]. Ferroan anorthosites span a 57 Fe range from -0.39 ± 0.05‰ to 0.50 ± 0.06‰ although the negative 57 Fe datum has only been measured for sample 60025. Two Mg-suite norites 15455 and 77215 exhibit 57 Fe values of 0.07 0.02‰ and 0.05 0.05‰, respectively. Dunite 72415 yielded an extremely light 57 Fe value of -0.73 0.06‰ (Table 1 and Fig. 2). Mg-suite troctolite 76335 has 57 Fe = -0.01 0.06‰.

The two most negative values, from dunite 72415 and Fe-anorthosite 60025, were fully duplicated by new sample powder dissolutions, iron purifications by ion exchange chromatography and MC-ICP-MS measurements. In both cases, identical results were obtained so the figures reported in Table 1 represent the average of these duplicate determinations. The KREEP-rich basaltic impact melt 14310 yielded a 57 Fe = 0.13 0.13‰, similar to the value for pristine KREEP basalt 15386. [START_REF] Wiesli | Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith[END_REF][START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF]. In contrast, new 57 Fe values for Mg-suite norite 77215 are significantly lower than the value published by [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF], whereas the new 57 Fe value for high-Ti basalt 10003 is higher than that of [START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF]. A full replicate analysis of high-alumina basalt 14053 (Table 1), for which different chips were obtained in two different CAPTEM allocations (to CRN and FP), have indistinguishable Fe isotope compositions; yet the recent 57 Fe value by [START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF] is higher. This could be caused by modest sample heterogeneity because high-Al basalt 14053 was subject to thermal metamorphism in an impact melt sheet that caused native Fe to form in the outer portion of the sample [START_REF] Taylor | The most reduced rock from the moon, Apollo 14 basalt 14053: Its unique features and their origin[END_REF]. The new measurement of Mg-suite dunite 72415 yielded 57 Fe = -0.73 ± 0.06‰, which is statistically significantly lighter than the already very light previously reported values of -0.53 ± 0.03‰ [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF], and -0.60 ± 0.05‰ [START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF]. Therefore, the very light Fe isotope signature of 72415 is confirmed (Table 1 and Fig. 2).

Some new analyses reported in

Data reported for olivine-normative low-Ti basalt 15555 by [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF] and [START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF] were differing by more than 0.16‰, and therefore well outside the quoted analytical uncertainties. This possibly results from the medium-to coarse-grained nature of this rock with variable olivine modal contents between different rock chips [START_REF] Ryder | Catalog of Apollo 15 rocks[END_REF].

Our new Fe isotope determination for basalt 15555 yields a 57 Fe value intermediate to those published elsewhere [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF][START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF]. There are no published 57 Fe values yet of olivines from low-Ti basalt 15555. Taking the 57 Fe = -0.117‰ for olivine from low-Ti ilmenite basalt 12045 [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF] and assuming that the variable modal content of olivine in 15555 will not be accompanied by distinctive olivine Fe isotope compositions, we can compute a bulk rock 57 Fe ranging from 0.141 to 0.101‰ with olivine modal content varying respectively from 5% to 20%, depending on the 15555 sample allocation [START_REF] Meyer | Lunar sample compendium[END_REF]. This 0.040‰ range is four times smaller than the range of 15555 bulk rock values published so far, which is 0.164% from six different rock allocations (Table 1). Taking the other low-Ti basalt (sample 12021) olivine value of 0.08‰ published by [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF] would make the range among calculated bulk rock 57 Fe values for 15555 even smaller since it is closer to the reported bulk rock values. This suggests that other causes than solely olivine modal contents are responsible for the scatter among the published bulk rock 57 Fe values for basalt 15555.

As discussed below, it is likely that unrepresentative aliquot allocations of coarse-grained and mineralogically heterogeneous rocks such as 15555 is the main reason for this discrepancy.

The most extreme discrepancy with previous results is observed for ferroan anorthosite 60025. The new 57 Fe result is >0.6‰ lighter than the value reported in [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF]. A full replicate of the entire analytical protocol at CNRS Toulouse yielded an identical result confirming the reliability of the analytical and instrumental procedures.

Petrologic heterogeneity within 60025 with various modal content of felsic and mafic minerals depending on the sample fragment has been previously documented (e.g. [START_REF] James | Lunar ferroan anorthosite-60025petrology and chemistry of mafic lithologies[END_REF]Torcivia andNeal, 2017, 2018). This will be discussed in more detail below.

Discussion

The key role of the sample size for representative Fe isotope analysis

The inter-laboratory comparison of Fe isotope measurement procedures on wellhomogenized silicate rock reference materials commonly shows good agreement [START_REF] Beard | Comment on "Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS" by F. Poitrasson and R. Freydier[END_REF][START_REF] Poitrasson | On the iron isotope homogeneity level of the continental crust[END_REF][START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF][START_REF] Craddock | Iron isotope compositions of reference material, geostandards and chondrites[END_REF]. Hence, the analytically significant 57 Fe differences between new data and published values for lunar samples 77215, 10003, 14053, 15555, 60025 and 72415 appear to reflect sample heterogeneity rather than analytical issues. Literature data for high-Ti mare basalt 70035 also yield discrepant 57 Fe values (Fig. 2 and Supplementary Table ). Given that the entire replicated lunar sample data were obtained on powder aliquots produced from different rock chips, it is possible that for at least these seven lunar rocks, the sample mass used to prepare the powders was too small to be representative of the bulk rock 57 Fe given the variable mineral grain size. These sample sizes typically range from ~50 mg to ~1 g, but some studies report Fe isotope analyses based on the direct dissolution of rock chips of "a few milligrams" size [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF]. For the smaller sample sizes, Fe isotope variations probably reflect the relative differences in abundance of the major mineral phases sampled (e.g. olivine in basalt 15555, [START_REF] Ryder | Catalog of Apollo 15 rocks[END_REF], which may carry distinctly different Fe isotope signature [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Craddock | Mineralogical control on iron isotopic fractionation during lunar differentiation and magmatism[END_REF][START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF]. This would translate into rock chips used for the analyses having olivine modal contents well outside the range of 5 to 20% reported for 15555 [START_REF] Meyer | Lunar sample compendium[END_REF]. In contrast, the Fe isotope analysis of bulk terrestrial rocks would be performed with an aliquot of powder typically produced from the crushing and homogenization of at least several hundred grams of rock (see general rock sample preparation methods in [START_REF] Potts | A handbook of silicate rock analysis[END_REF]. Though often underestimated, such a sampling issue was previously pointed out for lunar Apollo rocks. For example, trace and even major element concentration heterogeneity issues were noted for basalt 15555 when comparing analyses based on powders made from rock aliquots of less than 1g [START_REF] Ryder | Chemical dispersion among Apollo 15 olivine-normative mare basalts[END_REF][START_REF] Ryder | Chemical variation of the large Apollo 15 olivinenormative mare basalt rock samples[END_REF]. This was also reported to bias the oxygen isotope compositions of basalts 75075 and 15555 [START_REF] Spicuzza | Oxygen isotope constraints on the origin and differentiation of the Moon[END_REF][START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF], and, more recently, the silicon isotope composition of ferroan anorthosites 65035 and 60015 [START_REF] Poitrasson | An Earth-Moon silicon isotope model to track silicic magma origins[END_REF], when powders were produced from rock chips weighing less than ~100 mg [START_REF] Armytage | Silicon isotopes in lunar rocks: Implications for the Moon's formation and the early history of the Earth[END_REF]. Moreover, for basalt 15555, an anomalously light Mg isotope signature reported recently [START_REF] Sedaghatpour | Magnesium stable isotopes support the lunar magma ocean cumulate remelting model for mare basalts[END_REF], combined with significant Fe-Mg inter-mineral diffusion affecting Mg isotope systematics in olivine crystals of 15555 [START_REF] Richter | Magnesium isotopic zoning of an olivine grain from lunar microgabbro 15555: Constraints on crystallization and cooling[END_REF], provide an additional clue for possible sample Fe isotope heterogeneity issue (see Table 1).

Our new results (Table 1 and Fig. 2) reveal an extended range of Fe isotope compositions for lunar bulk rock samples (-0.73 to 0.50‰) relative to what has been published so far [START_REF] Wiesli | Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith[END_REF][START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF][START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF] with the most extreme values observed among the highland rocks (Fig. 2 and Table 1). Most ferroan anorthosites show distinctly heavy Fe isotope composition ( 57 Fe >0.18‰), similar to high-Ti basalts (Figure 2 and see also [START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF], or even higher with ferroan anorthosite 15415 yielding 57 Fe = 0.50‰ (This study). This is consistent with the results of Fe isotope investigations of mineral separates performed by [START_REF] Craddock | Mineralogical control on iron isotopic fractionation during lunar differentiation and magmatism[END_REF] on a suite of low-Ti and high-Ti mare basalts, clearly identifying ilmenite and plagioclase as 57 Fe-enriched phases.

Two different CAPTEM allocations of 60025 deviate by >0.6‰ in 57 Fe (Table 1). Iron is predominantly hosted by pyroxene and olivine in lunar rocks [START_REF] Papike | Lunar samples[END_REF] and if these phases are isotopically fractionated as is the case for mare basalts [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Craddock | Mineralogical control on iron isotopic fractionation during lunar differentiation and magmatism[END_REF], an effect on the Fe isotope signature of bulk anorthosites may be expected since the occurrence of an isotopically fractionated Fe-rich mafic mineral phase may modify the 57 Fe values. These pilot studies on mineral separates from lunar basalts also indicated that plagioclase is isotopically heavier than coexisting olivine and pyroxene. This could indeed explain the heavy Fe isotope signature of anorthosites compared to some other mafic highland rocks. In the case of 60025, [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF] reported Fe content of 0.50 wt.%, whereas the analysis of the allocated sub-sample 60025,874 revealed 4.84 wt.% Fe, nearly one order of magnitude higher abundance (Supplementary Table ). Such a composition approaches those of mafic portions of 60025 [START_REF] Ryder | Lunar anorthosite 60025, the petrogenesis of lunar anorthosites, and the composition of the Moon[END_REF][START_REF] James | Lunar ferroan anorthosite-60025petrology and chemistry of mafic lithologies[END_REF][START_REF] Meyer | Lunar sample compendium[END_REF] and confirms the larger occurrence of olivine and pyroxene in our sample allocation, most probably leading to notably lower 57 Fe signature relative to the other 57 Fe measurements of 60025 and other anorthosites. In addition, [START_REF] Ryder | Lunar anorthosite 60025, the petrogenesis of lunar anorthosites, and the composition of the Moon[END_REF] concluded that systematic variations in the mineral chemistry of 60025 were probably a reflection that this sample was a mixture of closely related materials. Future work requires rigorous investigations by in situ isotopic work on anorthosite mineral phases (e.g., [START_REF] Oeser | Processes and time scales of magmatic evolution as revealed by Fe-Mg chemical and isotopic zoning in natural olivines[END_REF] to fully evaluate this. The influence of olivine is well illustrated at the bulk-rock scale with dunite 72415 having ~93% modal olivine [START_REF] Dymek | Comparative petrology of lunar cumulate rocks of possible primary origin: Dunite 72415, troctolite 76535, norite 78235 and anorthosite 62237[END_REF] and with reported 57 Fe values ranging from -0.5 to -0.73‰ depending on the rock chip analyzed (Table 1 and Fig. 2). Such isotopically light values in olivine could result from chemical diffusion-driven kinetic disequilibrium between olivine and the last melt olivine was in contact with [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF]. The implications of these findings are further discussed below.

Iron isotope composition of different lunar reservoirs

Bulk lunar samples may be grouped into four different sets given their petrologic characteristics (e.g., [START_REF] Papike | Lunar samples[END_REF]: Low-Ti basalts, high-Ti basalts, highland rocks and impact melts. Nine new low-Ti basalts yield a mean 57 Fe = 0.113 ± 0.023‰ (2SE) and 12 new high-Ti basalts analyses give 57 Fe = 0.273 ± 0.039‰ (2SE). These values are indistinguishable from mean 57 Fe values of 0.118 ± 0.028‰ (2SE, n = 10) and 0.291 ± 0.051‰ (2SE, n = 7), reported by [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF] for low-and high-Ti basalts, respectively.

It should be noted that our new data include a full replicate analysis of high-alumina basalt 14053 (Table 1) for which different chips were obtained in two different CAPTEM allocations. Given that these replicate analyses are indistinguishable within uncertainty, their mean value is used for the low-Ti basalt average 57 Fe reported above. The new data from this study together with previously published results by [START_REF] Wiesli | Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith[END_REF], [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF], [START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF], [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF], [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF] and [START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF] yield a cumulative mean 57 Fe low-Ti = 0.127 ± 0.012‰ (2SE, n = 27; Table 1). For replicate analyses, the average 57 Fe of independent determinations is used to avoid overweighting of samples analyzed multiple times in the calculated mean. Using the same approach, we obtain a mean 57 Fe high-Ti = 0.274 ± 0.020‰ (2SE, n = 25; Table 1). Our eleven new determinations of 57 Fe in highland rocks lead to a less well-defined mean because of the significant Fe isotope variations among the samples. With the literature data included, the mean 57 Fe highland = 0.078 ± 0.124‰ (2SE, n=15; Table 1) is calculated. Within this group, Fe-anorthosites yield a mean 57 Fe composition of 0.192 ± 0.207‰ (2SE; n = 6) whereas a 57 Fe value of 0.002 ± 0.179‰ (2SE; n = 9) is obtained for the Mg-suite rocks. With literature data included, impact melts yield a mean 57 Fe impact = 0.093±0.055‰ (2SE, n=5).

The new data imply that the Fe isotope systematics of the Moon is more complex than initially thought. As highlighted by [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF], low-Ti and high-Ti mare basalts possess distinctive Fe isotope signatures (Fig. 2). With the new and published mare basalt data included, a t-test allows us to reject the null hypothesis of a common origin of low-Ti and high-Ti basalts with a level of confidence well above 99% (See Supplementary Table ).

Furthermore, the t-tests show that both the highland rocks and the impact melts have Fe isotope composition indistinguishable from that of the low-Ti basalts, whereas their mean 57 Fe values are statistically different from that of high-Ti basalts with a degree of confidence higher than 99% (Supplementary Table ). Hence, from the viewpoint of the observed Fe isotope systematics, the different rock types may be arranged as two isotopically distinct groups: a 57 Fe-enriched end-member defined by high-Ti basalts, and a 57 Fe-depleted endmember best defined by low-Ti basalts, but also including impact melts and highland rocks, though the latter do not define a well-constrained mean 57 Fe value (Tables 1 andSupplementary Table).

Impact melts do not represent a deep lunar reservoir since their origin in meteoritic bombardment events occur at the lunar surface. Based on the review of [START_REF] Shearer | Thermal and magmatic evolution of the Moon, New Views of the Moon[END_REF], low-Ti and high-Ti basalts make up a significantly larger component of the Moon in that they cover 17% of lunar surface, but they represent just ~1% of the lunar crust volume [START_REF] Head | Lunar volcanism in space and time[END_REF]. The largest reservoir for which we have direct samples are highland rocks, representing the remaining 99% of the lunar crust, that is, ~10% of the volume of the Moon [START_REF] Papike | Lunar samples[END_REF]. However, the present work reveals that this reservoir is more heterogeneous than previously anticipated [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF], with a total 57 Fe range of ~1.2‰. Subdividing these highland rocks into two groups consisting of ferroan anorthosites and Mg-suite rocks makes no difference since both groups are also very heterogeneous isotopically, as computed above. As discussed above, this likely stems from sample aliquot allocations that are too small to produce representative 57 Fe analyses of bulk rocks given the frequently coarse-grained nature of highland rocks. Whether calculated as a highland rock group or separated as a ferroan anorthosite and Mg-suite groups, the uncertainties of the means are too high for the precision required for comparative planetology studies (i.e., <0.1‰ level; see [START_REF] Dauphas | Iron isotopes may reveal the redox conditions of mantle melting from Archean to Present[END_REF][START_REF] Poitrasson | On the iron isotope heterogeneity of lithospheric mantle xenoliths: implications for mantle metasomatism, the origin of basalts and the iron isotope composition of the Earth[END_REF]. The next step is to determine whether the rock groups sampled at the surface of the Moon (i.e., low-Ti, high-Ti mare basalts, highland rocks…) are actually representative of larger-scale lunar reservoirs. [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF] have shown that low-Ti and high-Ti mare basalts do not exhibit Fe isotope variations as a function of magmatic differentiation. This feature has also been observed for most terrestrial igneous rocks, except for highly evolved silicate compositions, such as granites and rhyolites [START_REF] Poitrasson | Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS[END_REF][START_REF] Schoenberg | Modes of planetary-scale Fe isotope fractionation[END_REF][START_REF] Heimann | The role of volatile exsolution and subsolidus fluid/rock interactions in producing high Fe-56/Fe-54 ratios in siliceous igneous rocks[END_REF][START_REF] Schuessler | Iron and lithium isotope systematics of the Hekla volcano, Iceland -Evidence for Fe isotope fractionation during magma differentiation[END_REF][START_REF] Telus | Iron, zinc, magnesium and uranium isotopic fractionation during continental crust differentiation: The tale from migmatites, granitoids, and pegmatites[END_REF][START_REF] Foden | Fe isotopes and the contrasting petrogenesis of A-, I-and S-type granite[END_REF]. Iron isotope variations in mafic rocks have been observed in the closed-system Kilauea Iki lava lake, Hawaii [START_REF] Teng | Iron isotope fractionation during magmatic differentiation in Kilauea Iki Lava Lake[END_REF], that were subsequently shown to be caused by kinetic effects [START_REF] Sio | Discerning crystal growth from diffusion profiles in zoned olivine by in situ Mg-Fe isotopic analyses[END_REF]. Therefore, the question arises whether or not mare basalts directly reflect the Fe isotope composition of their respective mantle sources. The importance of partial melting in the Earth's mantle to generate systematic Fe isotope differences in derivative melts has been debated for some time. Whereas some consider this process to play a key role [START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Weyer | Partial melting and melt percolation in the mantle: The message from Fe isotopes[END_REF][START_REF] Williams | Systematic iron isotope fractionation variations in mantle rocks and minerals: The effect of partial melting and oxygen fugacity[END_REF]2009;[START_REF] Dauphas | Iron isotopes may reveal the redox conditions of mantle melting from Archean to Present[END_REF], others suggested that partial melting of pristine peridotitic lithologies generates only minor 57 Fe variations in the basaltic melts relative to their peridotitic protoliths [START_REF] Beard | Comment on "Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF][START_REF] Zhao | Iron isotope variations in spinel peridotite xenoliths from North China Craton: implications for mantle metasomatism[END_REF][START_REF] Poitrasson | On the iron isotope heterogeneity of lithospheric mantle xenoliths: implications for mantle metasomatism, the origin of basalts and the iron isotope composition of the Earth[END_REF]. Accordingly, using Nuclear Resonant Inelastic X-ray Scattering (NRIXS) measurements of synthetic glass and associated partial melting calculations, Dauphas et al. (2014a) could explain only ~1/3 of the observed Fe isotopic difference between measured MORBs and their estimated bulk silicate mantle. The issue is essentially linked to their incorrect estimate of the bulk silicate mantle 57 Fe. If, instead of chondritic (~0‰; [START_REF] Dauphas | Iron isotopes may reveal the redox conditions of mantle melting from Archean to Present[END_REF]2014a) a 57 Fe value of ~0.1‰ is adopted for the Earth's mantle [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF]2013), then the discrepancy noted by Dauphas et al. (2014a) between their experimental results and natural observations disappears. Hence, this issue remains debated for the Earth and further experimental work is required. The putative mechanism to explain Fe isotope fractionation during partial melting proposed by [START_REF] Dauphas | Iron isotopes may reveal the redox conditions of mantle melting from Archean to Present[END_REF]2014a) hinges on the preferential incorporation of isotopically heavy Fe 3+ in melts relative to solid residuum. However, [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF] recognized that ferric iron does not occur in the lunar mantle [START_REF] Sato | Oxygen fugacity values of Apollo 12, 14, and 15 lunar samples and reduced state of lunar magmas[END_REF][START_REF] Sato | Oxygen fugacity and other thermochemical parameters of Apollo 17 high-Ti basalts and their implications on the reduction mechanism[END_REF] and that experimental studies have shown the lack of equilibrium Fe isotope fractionation between Fe metal (occurring in the lunar mantle) and the prevalent Fe 2+ species at mantle conditions [START_REF] Roskosz | Experimental quantification of the fractionation of Fe isotopes during metal segregation from a silicate melt[END_REF][START_REF] Poitrasson | No iron isotope fractionation between molten alloys and silicate melt to 2000°C and 7.7 GPa: Experimental evidence and implications for planetary differentiation and accretion[END_REF].

As a result, [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF] concluded that: 1) partial melting in the lunar mantle probably did not produce Fe isotope fractionation, and 2) Fe isotope compositions of low-Ti and high-Ti mare basalts directly mirror those of their corresponding lunar mantle sources. These conclusions have recently been challenged by [START_REF] Sedaghatpour | Magnesium stable isotopes support the lunar magma ocean cumulate remelting model for mare basalts[END_REF] on the basis of a fractional crystallization model. However, these models are frequently underconstrained, with many of the unknown parameters adjusted to fit the observed bulk-rock analyses (see review by [START_REF] Dauphas | Iron Isotope Systematics[END_REF]. As a result, these models do not provide a strong constraint on the actual process at play to explain the observed bulk-rock isotopic compositions measured. For instance, Sedaghatpour and Jacobsen (2019) used an estimated ∆ 57 Fe olivine-melt fractionation factor of -0.21‰ whereas it was shown experimentally to be at ~0‰ for mafic melts, including those having the chemistry of mare basalts [START_REF] Prissel | Experimentally determined effects of olivine crystallization and melt titanium content on iron isotopic fractionation in planetary basalts[END_REF]. We thus take this as a strong support for the conclusions derived by [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF].

Many authors have estimated the mineralogy of mantle sources of low-Ti and high-Ti basalts. For example, [START_REF] Snyder | A Chemical-Model for Generating the Sources of Mare Basalts -Combined Equilibrium and Fractional Crystallization of the Lunar Magmasphere[END_REF] and [START_REF] Beard | The source region and melting mineralogy of high-titanium and low-titanium lunar basalts deduced from Lu-Hf isotope data[END_REF] inferred an important role of ilmenite in the genesis of high-Ti basalts, although physical constraints of the lunar magma ocean evolution must also be considered (e.g. [START_REF] Elkins-Tanton | The lunar magma ocean: Reconciling the solidification process with lunar petrology and geochronology[END_REF]. [START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF] discussed that ilmenite, although potentially explaining elevated Ti contents of high-Ti basalts [START_REF] Snyder | A Chemical-Model for Generating the Sources of Mare Basalts -Combined Equilibrium and Fractional Crystallization of the Lunar Magmasphere[END_REF], would be incapable of shifting 57 Fe towards high values given that theoretical log factors (a reduced isotopic partition function factor) calculated for ilmenite were among the lowest for silicate and oxide minerals (compare Figures 1 and2 from [START_REF] Polyakov | The use of Mössbauer spectroscopy in stable isotope geochemistry[END_REF]. As a consequence, melting of an ilmenite-rich source should not result in basaltic melts with elevated 57 Fe values. However, more recent Fe isotope analyses of mineral phases of mare basalts have shown that ilmenite is in most cases isotopically heavier than coexisting pyroxene and plagioclase [START_REF] Craddock | Mineralogical control on iron isotopic fractionation during lunar differentiation and magmatism[END_REF]. To directly compare Fe isotope data for mineral separates with theoretical calculations, we have to assume that the mineral fractions analyzed by [START_REF] Craddock | Mineralogical control on iron isotopic fractionation during lunar differentiation and magmatism[END_REF] are pure and represent equilibrium mineral isotope compositions. This might not be the case for all samples in their study given the observed scatter in inter-mineral Fe isotopic fractionation, however. With this caveat in mind, it would indicate that, similar to pyrite [START_REF] Blanchard | Iron isotope fractionation between pyrite (FeS 2 ), hematite (Fe 2 O 3 ) and siderite (FeCO 3 ): A firstprinciples density functional theory study[END_REF][START_REF] Blanchard | Comment on "New data on equilibrium iron isotope fractionation among sulfides: Constraints on the mechanisms of sulfide formation in hydrothermal and igneous systems[END_REF][START_REF] Polyakov | Fe beta-factors for sulfides from NRIXS synchrotron experiments[END_REF], the Fe isotope fractionation factor for ilmenite estimated by [START_REF] Polyakov | The use of Mössbauer spectroscopy in stable isotope geochemistry[END_REF] appears to be grossly underestimated. If correct, the heavy Fe isotope signature of high-Ti mare basalts may be explained by partial melting of ilmenite in their source mantle as proposed by [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF]. They based their model on the inferences from [START_REF] Snyder | A Chemical-Model for Generating the Sources of Mare Basalts -Combined Equilibrium and Fractional Crystallization of the Lunar Magmasphere[END_REF] and [START_REF] Elkins-Tanton | The lunar magma ocean: Reconciling the solidification process with lunar petrology and geochronology[END_REF] suggesting that >78% of the early lunar magma ocean crystallized olivine and orthopyroxene, whereas an ilmenite-rich cumulate formed only after 95% lunar magma ocean crystallization. Combined with the remote sensing study of [START_REF] Giguere | The titanium contents of lunar mare basalts[END_REF] indicating that high-Ti basalts should represent ~10% of all lunar maria, [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF] estimated that only ~10% of the lunar upper mantle is akin to the source of the high-Ti mare basalts, whereas the remaining is similar to the source of the low-Ti mare basalts. Given that mare basalts likely mirror the Fe isotope composition of their source (see above), we can estimate the 57 Fe of the lunar upper mantle at 0.142 ± 0.026‰ (Fig. 2) from the newly derived mean 57 Fe values of low-Ti and high-Ti basalts, combined with the relative proportion of their mantle protoliths (of respectively 90 and 10%, with an estimated uncertainty of 10% by [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF], using a simple mass balance calculation. This value would apply to the upper ~250 km of the lunar mantle considering mare basalt petrogenesis [START_REF] Longhi | Experimental Petrology and Petrogenesis of Mare Volcanics[END_REF]. The estimated size of this reservoir from seismology extends down to 500 km, where the limit between the upper and lower lunar mantle was set [START_REF] Nakamura | Seismic velocity structure of the lunar mantle[END_REF].

It is interesting to note that lunar volcanic picritic glasses, thought to have been derived from the melting of protoliths located 360-520 km deep in the lunar mantle (Delano, 1979;Elkins-Tanton et al., 2003), show somewhat lighter Fe isotope signatures. However, the 57 Fe values scatter by over ~0.5‰, even for the same samples [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Weyer | Iron isotope fractionation during planetary differentiation[END_REF][START_REF] Moynier | Isotopic composition of zinc, copper, and iron in lunar samples[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF]. Whether this reflects Fe isotope stratification intrinsic to the lunar mantle, an effect of volcanism through fire fountaining [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF] or a sample heterogeneity issue remains to be tested.

The lunar highland rocks, and notably the anorthosites that make up ~80% of the crust [START_REF] Papike | Lunar samples[END_REF], may potentially give a more comprehensive view of the lunar mantle composition given that it was produced by magma ocean differentiation. Estimates for depth of the early lunar magma ocean may vary from 500 to 1200 km, and up to the entire Moon, depending again on whether seismic or petrological constraints are preferred [START_REF] Shearer | Thermal and magmatic evolution of the Moon, New Views of the Moon[END_REF][START_REF] Wieczorek | The constitution and structure of the lunar interior, New Views of the Moon[END_REF]. However, as discussed above, the majority of the highland samples for which Fe isotope data exist may show a potential effect of their coarse-grained, essentially mono-mineral nature. This makes the determination of the Fe isotope composition of the bulk lunar highlands reservoir based solely on mean 57 Fe values of these rocks rather imprecise (0.078 ± 0.124‰; Table 1) to provide useful constraints. Even if highland rocks are split into ferroan anorthosites and Mg-suite rocks, the large uncertainties remaining do not change this conclusion. As a parallel to the isotopically heterogeneous Earth's mantle possibly linked to partial melting events [START_REF] Williams | Systematic iron isotope fractionation variations in mantle rocks and minerals: The effect of partial melting and oxygen fugacity[END_REF][START_REF] Weyer | Partial melting and melt percolation in the mantle: The message from Fe isotopes[END_REF][START_REF] Williams | Iron isotope tracing of mantle heterogeneity within the source regions of oceanic basalts[END_REF], but most importantly due to metasomatic effects (Weyer et al., 2007;[START_REF] Zhao | Iron isotope variations in spinel peridotite xenoliths from North China Craton: implications for mantle metasomatism[END_REF][START_REF] Poitrasson | On the iron isotope heterogeneity of lithospheric mantle xenoliths: implications for mantle metasomatism, the origin of basalts and the iron isotope composition of the Earth[END_REF], a more precise 57 Fe estimate of this reservoir can only be provided by a careful consideration of the petrology of these rocks, even though metasomatic processes are less likely for the Moon [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF]. It has been shown above that plagioclase leads to heavy bulk-rock Fe isotope compositions whereas olivine tends towards isotopically light values in case of diffusion. Besides sample aliquot representativeness issues, such kinetic effects in olivine may make the bulk rock 57 Fe lighter [START_REF] Collinet | Crystallization history of enriched shergottites from Fe and Mg isotope fractionation in olivine megacrysts[END_REF]. In addition, these rocks may also represent mixed lithologies [START_REF] Ryder | Lunar anorthosite 60025, the petrogenesis of lunar anorthosites, and the composition of the Moon[END_REF][START_REF] James | Lunar ferroan anorthosite-60025petrology and chemistry of mafic lithologies[END_REF].

Plotting the bulk-rock 57 Fe versus the modal content of the other major phases from these samples showing less variable Fe isotope fractionation, i.e., pyroxenes, leads to a more precise Fe isotope composition for highland rocks (Fig. 3). By using samples with >10 modal % pyroxene, a better-constrained mean 57 Fe of 0.094 ± 0.035‰ is obtained for lunar highlands. This value essentially corresponds to the mean of the six Mg-suite norites for which Fe isotope compositions are available (Supplementary Table ), although three anorthosites and one Mg-suite troctolite also fall on this mean value (Fig. 3).

Implications for the Fe isotope composition of bulk Moon

Overall, it appears that the 57 Fe estimate of the lunar upper mantle sampled by mare basalts (0.142 ± 0.026‰) is indistinguishable from the value obtained from highland rocks (0.094 ± 0.035‰; Fig. 3) that likely sample a larger portion of the lunar mantle. If partial melting of the lunar mantle imparted no measurable Fe isotope fractionation [START_REF] Liu | Oxygen and iron isotope constraints on near-surface fractionation effects and the composition of lunar mare basalt source regions[END_REF][START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF], it is likely that the highland 57 Fe value of 0.094 ± 0.035‰ is a relevant estimate of the silicate portion of the Moon, considering the mass balance of these reservoirs. [START_REF] Schoenberg | Modes of planetary-scale Fe isotope fractionation[END_REF] and [START_REF] Elardo | Non-chondritic iron isotope ratios in planetary mantles as a result of core formation[END_REF] have hypothesized that Fe isotope fractionation between silicate portions of terrestrial planets and their metallic cores could occur. In contrast, the investigation of Fe isotope compositions in various meteorite classes led to the conclusion that this mechanism is unlikely for planetary bodies (Poitrasson et al., 2005), an inference also supported by subsequent experimental investigations [START_REF] Roskosz | Experimental quantification of the fractionation of Fe isotopes during metal segregation from a silicate melt[END_REF][START_REF] Poitrasson | No iron isotope fractionation between molten alloys and silicate melt to 2000°C and 7.7 GPa: Experimental evidence and implications for planetary differentiation and accretion[END_REF][START_REF] Hin | Experimental evidence for the absence of iron isotope fractionation between metal and silicate liquids at 1 GPa and 1250-1300°C and its cosmochemical consequences[END_REF]. This conclusion was challenged using the three-isotope experimental approach [START_REF] Shahar | Sulfur-controlled iron isotope fractionation experiments of core formation in planetary bodies[END_REF], although four out of the five more recent experimental results of [START_REF] Elardo | The effect of core composition on iron isotope fractionation between planetary cores and mantles[END_REF] showed no statistically significant Fe isotope fractionation between metal and silicate. In practice, this three-isotope methodology involves, at high temperature, a first thorough chemical and isotopic mixing between the liquid metallic alloy and the silicic melt followed by "un-mixing" along a secondary fractionation line according to [START_REF] Shahar | Equilibrium Fractionation of Nontraditional Stable Isotopes: an Experimental Perspective[END_REF] see their Fig. 2B). As a result, the remainder of the system evolution until isotopic equilibrium is very similar to timeseries experiments conducted without spike in which losses are monitored by mass balance.

Besides time series, the experiments of [START_REF] Poitrasson | No iron isotope fractionation between molten alloys and silicate melt to 2000°C and 7.7 GPa: Experimental evidence and implications for planetary differentiation and accretion[END_REF] were reversed, and the lack of Fe isotope fractionation between metal and silicate at equilibrium under planetary core formation conditions was also concluded in studies using a completely different experimental set ups [START_REF] Roskosz | Experimental quantification of the fractionation of Fe isotopes during metal segregation from a silicate melt[END_REF][START_REF] Hin | Experimental evidence for the absence of iron isotope fractionation between metal and silicate liquids at 1 GPa and 1250-1300°C and its cosmochemical consequences[END_REF] and see also results of [START_REF] Elardo | The effect of core composition on iron isotope fractionation between planetary cores and mantles[END_REF], as well as in studies using meteorite samples (Poitrasson et al., 2005;[START_REF] Chernonozhkin | Effect of parent body evolution on equilibrium and kinetic isotope fractionation: a combined Ni and Fe isotope study of iron and stony-iron meteorites[END_REF][START_REF] Chernonozhkin | Thermal equilibration of iron meteorite and pallasite parent bodies recorded at the mineral scale by Fe and Ni isotope systematics[END_REF][START_REF] Jordan | Iron isotope constraints on planetesimal core formation in the early solar system[END_REF]. Further, this three-isotope experimental methodology could also be affected by possible kinetic biases [START_REF] Bourdon | Isotope tracers of core formation[END_REF], so results obtained by this method should be interpreted with caution.

On a more theoretical side based on NRIXS measurements of high-pressure mineral phases, [START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF] proposed that a very high pressure-induced phase change in the lower mantle, at pressures above ~100 GPa, may generate Fe isotope differences between the silicate Earth and metallic core. This hypothesis does not apply to the lunar core-mantle differentiation, though, given the too low pressures occurring in the lunar interior. The origin of the lunar material in the aftermath of a Moon-forming giant impact is debated (see review by Dauphas et al., 2014b), with models involving various proportions of the impactor material and the proto-Earth mantle. However, the Moon forming material is unlikely to be inherited from the silicate portion of the Earth from a giant impact occurring at a point when the Earth was already big enough to reach 100 GPa since only a very small portion of the deep Earth mantle would show this heavy Fe isotope signature in the framework of [START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF] theory. Moreover, subsequent studies using similar NRIXS approaches did not support [START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF] inferences since they have concluded that metal-silicate at high pressure and high temperature conditions should result in undetectably small Fe isotope fractionation (Dauphas et al., 2014a;[START_REF] Liu | Iron isotopic fractionation between silicate mantle and metallic core at high pressure[END_REF][START_REF] Yang | Iron isotopic fractionation in mineral phases from Earth's lower mantle: Did terrestrial magma ocean crystallization fractionate iron isotopes?[END_REF]. Therefore, the estimate of Fe isotope composition of the lunar silicate mantle also involves its small metallic core and it is thus valid for the bulk Moon (Fig. 2). Collectively, at the current level of knowledge, the bulk lunar 57 Fe composition should be close to 0.094 ± 0.035‰, which is indistinguishable from the Earth's value of 0.10±0.03‰ (Fig. 2). This lunar estimate agrees well with the recent value proposed by [START_REF] Sossi | Chemical and isotopic kinship of iron in the Earth and Moon deduced from the lunar Mg-Suite[END_REF], though on the basis of a more limited set of lunar samples (five Mg-suite rocks consisting mostly of norites) and using different lines of reasoning. It is significantly heavier than the other recent estimate of [START_REF] Wang | The earliest Lunar Magma Ocean differentiation recorded in Fe isotopes[END_REF] that was based on a sole sample only (dunite 72415); that particular sample has been shown to yield rather heterogeneous 57 Fe values depending on the sample allocation (Fig. 2) and it was based on a corrected isotopic Fe composition using isotopic diffusion parameters in olivine that have been revised since [START_REF] Oeser | Processes and time scales of magmatic evolution as revealed by Fe-Mg chemical and isotopic zoning in natural olivines[END_REF].

Implications for the formation of the Moon

Given the previously described experimental and theoretical disagreements noted in the literature as to whether the mantle-core differentiation did generate isotopically distinct silicate and metallic reservoirs in the interior of the Earth, the bulk lunar Fe isotopic composition can constitute a good test of this idea. If correct, we should expect a silicate Earth with a heavier Fe isotope composition than that of the silicate Moon since the high pressures required for the phase transition to impart additional Fe isotope fractionation were not reached in the lunar interior. Realizing that no observation supports a Moon isotopically lighter than the silicate fraction of the Earth, [START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF] and [START_REF] Rustad | Iron isotope fractionation in the Earth's lower mantle[END_REF] concluded that a Moon-forming giant impact is required to enable the Moon to acquire the heavy Fe isotope composition of the Earth through a homogenization process such as that proposed by [START_REF] Pahlevan | Equilibration in the aftermath of the lunar-forming giant impact[END_REF]. Such a scenario implies that a large fraction of the Moon-forming material came from the proto-Earth mantle, which is not easy to achieve even using high resolution numerical simulations [START_REF] Canup | Lunar-forming impacts: High-resolution SPH and AMR-CTH simulations[END_REF]. Furthermore, and as discussed above, another constraint to this scenario is that the giant impact should have occurred sufficiently late in the proto-Earth accretion history to achieve a mantle-core boundary pressure well beyond ~100 GPa pressure at which the Fe isotope metal-silicate fractionation process proposed by [START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF] can operate. However, the current Earth's mantle-core boundary is only ~32 GPa higher than this limit, so the proportion of mantle having the phase change and associated isotopic effect proposed by [START_REF] Polyakov | Equilibrium iron isotope fractionation at core-mantle boundary conditions[END_REF] is limited. Therefore, the Earth should have nearly reached its current size and enough time should be allowed to lead the deep mantle to impart its Fe isotope composition to the shallower portions by convection before the impact. However, it appears that terrestrial mantle homogenization was not achieved ~1 Ga after Earth's formation [START_REF] Bennett | Coupled Nd-142-Nd-143 isotopic evidence for Hadean mantle dynamics[END_REF][START_REF] Touboul | W-182 Evidence for Long-Term Preservation of Early Mantle Differentiation Products[END_REF][START_REF] Debaille | Stagnant-lid tectonics in early Earth revealed by Nd-142 variations in late Archean rocks[END_REF], so well after the putative Moon-forming giant impact timeframe. Furthermore, the Moon forming material should essentially come from the proto-Earth to make this scenario tenable, which seems unlikely [START_REF] Canup | Lunar-forming impacts: High-resolution SPH and AMR-CTH simulations[END_REF]Dauphas et al., 2014b). Hence, our best estimate for the Moon 57 Fe value, which is indistinguishable from that of the Earth (Fig. 2), therefore also provides a constraint against terrestrial mantlecore Fe isotope fractionation.

However, whereas meteorite and experimental studies conclude the absence of Fe isotope fractionation during mantle-core differentiation of planetary bodies (Poitrasson et al., 2005;[START_REF] Roskosz | Experimental quantification of the fractionation of Fe isotopes during metal segregation from a silicate melt[END_REF][START_REF] Poitrasson | No iron isotope fractionation between molten alloys and silicate melt to 2000°C and 7.7 GPa: Experimental evidence and implications for planetary differentiation and accretion[END_REF][START_REF] Hin | Experimental evidence for the absence of iron isotope fractionation between metal and silicate liquids at 1 GPa and 1250-1300°C and its cosmochemical consequences[END_REF]Chernonozkhin et al., 2016Chernonozkhin et al., , 2017)), some recent experimental studies propose an opposite sense of fractionation during mantle-core differentiation, that is, with a metallic portion becoming isotopically heavier [START_REF] Shahar | Sulfur-controlled iron isotope fractionation experiments of core formation in planetary bodies[END_REF][START_REF] Elardo | Non-chondritic iron isotope ratios in planetary mantles as a result of core formation[END_REF][START_REF] Elardo | The effect of core composition on iron isotope fractionation between planetary cores and mantles[END_REF]. Such a sense of Fe isotope fractionation would yield planetary mantles isotopically lighter than chondrites. This finds neither evidence from the lunar samples analyzed so far apart from the extremely heterogeneous volcanic glasses (see above), nor is this supported by the latest planetary 57 Fe estimates (see [START_REF] Wang | Iron isotope fractionation in planetary crusts[END_REF][START_REF] Sossi | On the iron isotope composition of Mars and volatile depletion in the terrestrial planets[END_REF][START_REF] Dauphas | Iron Isotope Systematics[END_REF].

A simpler scenario can explain the heavy Fe isotope composition of the Earth, Moon and angrite parent body through the increased volatility of light Fe isotopes during the accretion of the protoplanetary disk [START_REF] Sossi | On the iron isotope composition of Mars and volatile depletion in the terrestrial planets[END_REF] and/or during accretionary impacts (see [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF]. [START_REF] Sossi | On the iron isotope composition of Mars and volatile depletion in the terrestrial planets[END_REF] proposed an effect linked to differential condensation of solids in the protoplanetary accretion disk on the basis of a positive correlation between 57 Fe and Fe/Mn estimates of planetary bodies (Fig. 4a), except chondrites. However, correlations between elemental ratios of contrasted volatilities (Rb/Sr or Mn/Na) and 57 Fe would rather suggest two different histories for the Earth, the Moon and the angrite parent body on one side and Mars, Vesta and the chondrite (CI) parent body on the other (Fig. 4b andc). Since there are no reasons to consider that Mars, Vesta and the chondrite parent body chemistries were also unaffected by the effect of differential solid condensation of the protoplanetary accretion disk, the dichotomy observed in Figs. 4b andc suggests that Fe isotopes record different accretion mechanisms. The lack of Fe isotope variation for Mars and Vesta relative to the chondrite (CI) parent body would imply accretionary processes through runaway growth, whereas the Earth, Moon and angrite parent body would be additionally affected by high-energy impacts towards the end of their accretion history [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Wang | Iron isotope fractionation in planetary crusts[END_REF]. This would produce planetary bodies enriched in heavy Fe isotopes through the loss of light Fe isotopes that do not necessarily require percent-level Fe loss, as computed by [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF], potentially leading to measurable Fe isotope variations. This is in contrast to the effect of differential condensation in the protoplanetary accretion disk that would primarily generate changes in elemental abundances, potentially accompanied by Fe isotope variations [START_REF] Sossi | On the iron isotope composition of Mars and volatile depletion in the terrestrial planets[END_REF]. According to the dynamic and thermodynamic calculation of [START_REF] Dauphas | Planetary and meteoritic Mg/Si and delta Si-30 variations inherited from solar nebula chemistry[END_REF], such a loss of light isotopes to space in the aftermath of interplanetary impacts should be difficult for small parent bodies of the size of the angrite parent body. However, more recent calculations involving the track motion of Jupiter generating gas drag from recently impacted bodies make this hypothesis feasible [START_REF] Hin | Magnesium isotope evidence that accretional vapour loss shapes planetary compositions[END_REF], although the calculations were conducted down to bodies only having twice the size of the angrite parent body (<500 km). More observational constraints are thus required to refine and reconcile these models.

Evaluation of the model with other stable isotope systems

There has been a growing body of mass-dependent stable isotope planetary studies in the recent years that should be compared to the Fe isotope systematics to produce, as far as possible, a scenario taking into account all available observations. [START_REF] Humayun | Potassium isotope cosmochemistry: Genetic implications of volatile element depletion[END_REF] failed to find K isotope fractionation among planetary bodies at the then-achievable levels of precision of ±~0.5‰ at the 95% confidence level. Two decades later and using advanced plasma source mass spectrometry techniques, [START_REF] Wang | Potassium isotopic evidence for a high-energy giant impact origin of the Moon[END_REF] determined that lunar rocks are higher by ~0.4‰ in 41 K relative to terrestrial samples and chondrites. They interpreted this new observation as tracking the effect of a high-energy, high angular momentum giant impact in a thick gas atmosphere surrounding the impacted proto-Earth [START_REF] Lock | The Origin of the Moon Within a Terrestrial Synestia[END_REF] rather than a low-energy disk equilibration model [START_REF] Pahlevan | Equilibration in the aftermath of the lunar-forming giant impact[END_REF].

The stable isotope compositions of volatile elements Ga (half-mass condensation temperature T c of 968K, [START_REF] Lodders | Solar system abundances and condensation temperatures of the elements[END_REF], Rb (T c = 800K, [START_REF] Lodders | Solar system abundances and condensation temperatures of the elements[END_REF] and Zn (T c = 726K; [START_REF] Lodders | Solar system abundances and condensation temperatures of the elements[END_REF] also yielded isotopically heavier values consistent with a giant impact Moon-forming scenario, associated with volatile loss [START_REF] Paniello | Zinc isotopic evidence for the origin of the Moon[END_REF][START_REF] Kato | Extensive volatile loss during formation and differentiation of the Moon[END_REF][START_REF] Kato | Gallium isotopic evidence for extensive volatile loss from the Moon during its formation[END_REF][START_REF] Pringle | Rubidium isotopic composition of the Earth, meteorites, and the Moon: Evidence for the origin of volatile loss during planetary accretion[END_REF]. However, as the database growths, the lunar rocks appear to be extremely heterogeneous, notably anorthosites, yielding for example a 66 Zn range of more than 15‰ [START_REF] Paniello | Zinc isotopic evidence for the origin of the Moon[END_REF][START_REF] Kato | Extensive volatile loss during formation and differentiation of the Moon[END_REF]. By comparing the stable isotope composition of the variably volatile elements K, Rb and Zn, [START_REF] Pringle | Rubidium isotopic composition of the Earth, meteorites, and the Moon: Evidence for the origin of volatile loss during planetary accretion[END_REF] noted that there is no clear systematics between the planetary isotope differences and the geochemical properties of these elements. This makes a detailed discussion, besides the general idea of a volatility effect explaining the high 41 K, 71 Ga, 87 Rb and 66 Zn for some lunar rocks relative to other planetary samples rather difficult.

Clearly, more work is needed to understand the igneous systematics of these stable isotope systems, as already shown with the comparatively more studied Fe isotope systematics (see above).

There is no consensus yet as to whether the majority of terrestrial bodies have a chondritic 26 Mg signature (see recent review by [START_REF] Teng | Magnesium Isotope Geochemistry[END_REF], and see also [START_REF] Sedaghatpour | Magnesium stable isotopes support the lunar magma ocean cumulate remelting model for mare basalts[END_REF], or if the Earth is isotopically heavier [START_REF] Wiechert | Non-chondritic magnesium and the origins of the inner terrestrial planets[END_REF][START_REF] Hin | Magnesium isotope evidence that accretional vapour loss shapes planetary compositions[END_REF]. This partly results from analytical difficulties to determine accurate 26 Mg values of silicate rock samples. Furthermore, lunar samples appear to yield heterogeneous 26 Mg values, with high-Ti basalts tending towards lighter isotopic compositions according to [START_REF] Sedaghatpour | Magnesium isotopic composition of the Moon[END_REF] and [START_REF] Sedaghatpour | Magnesium stable isotopes support the lunar magma ocean cumulate remelting model for mare basalts[END_REF] whereas [START_REF] Wiechert | Non-chondritic magnesium and the origins of the inner terrestrial planets[END_REF] found the opposite. Given these controversies, it is still difficult at this point to use Mg isotope compositions to discuss planet formation and differentiation processes.

Silicon was another element that initially generated debates for analytical reasons [START_REF] Georg | Silicon in the Earth's core[END_REF][START_REF] Fitoussi | Si isotope systematics of meteorites and terrestrial peridotites: implications for Mg/Si fractionation in the solar nebula and for Si in the Earth's core[END_REF]. Further, magma differentiation or the nature of protoliths may affect 30 Si values of igneous rocks [START_REF] Savage | High temperature silicon isotope geochemistry[END_REF], including on the Moon [START_REF] Poitrasson | An Earth-Moon silicon isotope model to track silicic magma origins[END_REF], although no difference between low-Ti and high-Ti mare basalts was observed [START_REF] Armytage | Silicon isotopes in lunar rocks: Implications for the Moon's formation and the early history of the Earth[END_REF]Fitoussi and Bourdon, 2012;[START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF]. However, most authors now agree on the relative 30 Si differences between meteorite types. They also concur that the Earth and the Moon are isotopically indistinguishable, with 30 Si close to -0.29‰ relative to NBS-28 reference material). These two bodies are also isotopically heavier than Mars, Vesta and chondrites, but lighter than angrites (see recent review in [START_REF] Poitrasson | Silicon isotope geochemistry[END_REF].

Lithium is an element seemingly not easy to use for interplanetary comparisons as it may potentially show strong 7 Li variations at the intra-mineral scale caused by thermal diffusion or fluid circulations (see recent reviews in [START_REF] Tomascak | Advances in Lithium isotope geochemistry[END_REF][START_REF] Penniston-Dorland | Lithium Isotope Geochemistry[END_REF]. Yet resolved 7 Li variations were observed for low-Ti versus high-Ti mare basalts [START_REF] Magna | New constraints on the lithium isotope compositions of the Moon and terrestrial planets[END_REF][START_REF] Seitz | Lithium isotope compositions of Martian and lunar reservoirs[END_REF][START_REF] Day | Evidence for high-temperature fractionation of lithium isotopes during differentiation of the Moon[END_REF]. Low-Ti basalts, the Earth's mantle [START_REF] Jeffcoate | Li isotope fractionation in peridotites and mafic melts[END_REF][START_REF] Magna | Links between mantle metasomatism and lithium isotopes: Evidence from glass-bearing and cryptically metasomatized xenoliths from Mongolia[END_REF][START_REF] Tomascak | Lithium isotopes in global mid-ocean ridge basalts[END_REF][START_REF] Marschall | The boron and lithium isotopic composition of mid-ocean ridge basalts and the mantle[END_REF], Mars [START_REF] Magna | Lithium isotope constraints on crust-mantle interactions and surface processes on Mars[END_REF] and Vesta [START_REF] Magna | Lithium systematics in howardite-eucritediogenite meteorites: Implications for crust-mantle evolution of planetary embryos[END_REF] all carry broadly similar 7 Li signatures at ~3-4‰.

Pilot studies of mass-dependent isotopic fractionations of Sr [START_REF] Moynier | Sr stable isotope composition of Earth, the Moon, Mars, Vesta and meteorites[END_REF] and 44 Ca [START_REF] Simon | Stable calcium isotopic composition of meteorites and rocky planets[END_REF] do not seem to reveal significant planetary differences either. This could be observed besides large stable isotope variations of chondritic samples [START_REF] Simon | Stable calcium isotopic composition of meteorites and rocky planets[END_REF][START_REF] Charlier | High temperature strontium stable isotope behaviour in the early solar system and planetary bodies[END_REF] and an effect of differentiation for igneous terrestrial rocks for Sr [START_REF] Charlier | High temperature strontium stable isotope behaviour in the early solar system and planetary bodies[END_REF]. This effect of igneous differentiation was also observed for mass-dependent 49 Ti and 53 Cr compositions on Earth, along with a dichotomy between high-and low-Ti lunar basalts (Bonnand et al., 2016a;[START_REF] Millet | Titanium stable isotope investigation of magmatic processes on the Earth and Moon[END_REF]).

Yet, the authors of these studies found similar Earth-Moon stable isotope compositions. More recently, Sossi et al. (2018) produced more lunar sample 53 Cr values that led to an even more scattered data set. After data filtering, they found an isotopically lighter Cr signature of the Moon relative to the Earth, potentially reflecting volatile loss of oxygenated Cr species from the Moon towards the end of the lunar accretion history, at the lunar magma ocean stage.

Chondrites and achondrites 53 Cr are heterogeneous, possibly as a result of magmatic differentiation for the latter (Bonnand et al., 2016b). Nevertheless, they seem to share the 53 Cr signature of the Earth overall [START_REF] Schoenberg | The stable Cr isotopic compositions of chondrites and silicate planetary reservoirs[END_REF].

To summarize, only iron and silicon have so far been found to display a mass-dependent enrichment towards heavy isotope composition of the Earth-Moon system relative to other planetary bodies. Importantly, and in contrast to the other elements considered, Fe and Si are likely the most abundant elements of the Earth's core, besides Ni (e.g., [START_REF] Allègre | Chemical composition of the Earth and the volatility control on planetary genetics[END_REF][START_REF] Mcdonough | Compositional model for the Earth's core[END_REF], even though estimates for silicon remain more uncertain (e.g., [START_REF] Hirose | Composition and State of the Core[END_REF][START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF][START_REF] Dauphas | Planetary and meteoritic Mg/Si and delta Si-30 variations inherited from solar nebula chemistry[END_REF]. As discussed above, metal-silicate partitioning during mantle core-differentiation is unlikely to be the major cause to explain the observed planetary Fe isotope compositions. Hence, it would be preferable, until proof of the contrary is provided, if all those isotopic observations were explained by a single process that some authors [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF] have proposed to interpret the estimated Fe and Si isotopic composition of planets: the partial planetary melting and vaporization associated with the Moon-forming giant impact. Experimental work has shown that the vaporization rate of a metal species from liquid metal is much larger compared to that of an oxide from a silicic melt (Wang et al., 1994). Therefore, at a given temperature of an interplanetary impact-induced vaporization, elemental Fe and Si evaporation from a metallic melt will be much more efficient than evaporation of respective oxides from the coexisting silicic melt, and thus more likely to leave a stable isotope imprint through the loss of lighter Fe and Si isotopes in space [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF]. Numerical simulation of a Moon-forming giant impact tend to exclude models yielding a too Fe-rich outer disk to take into account the estimated lower bulk Fe content of the Moon relative to the Earth [START_REF] Canup | Simulations of a late lunar-forming impact[END_REF][START_REF] Cuk | Making the Moon from a Fast-Spinning Earth: A Giant Impact Followed by Resonant Despinning[END_REF][START_REF] Reufer | A hit-and-run giant impact scenario[END_REF][START_REF] Canup | Lunar-forming impacts: High-resolution SPH and AMR-CTH simulations[END_REF].

However, as mentioned above, the amount of Fe in the lunar-forming disk does not need to be high to account for the observed mass-dependent Fe and Si isotope systematics since previous quantitative estimates showed that less than 1% of iron and silicon loss from the Earth-Moon system into space is required in this context to explain their observed heavy isotopic compositions relative to other planetary bodies and chondrites [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF][START_REF] Zambardi | Silicon isotope variations in the inner solar system: Implications for planetary formation, differentiation and composition[END_REF].

Conclusions

The new Fe isotope data for various lunar rocks clearly show isotopically distinct reservoirs. This is apparent from data for low-Ti and high-Ti mare basalts that have means significantly different to a degree of confidence well beyond 99%. Partial melting in the lunar mantle and subsequent magmatic evolution of these mafic melts probably did not induce Fe isotope fractionation. Rather, this sharp difference in Fe isotope composition is likely due to lunar mantle sources having different mineralogical compositions, with heavy Fe isotope composition of high-Ti basalts likely reflecting the occurrence of ilmenite. The highland rocks, that include ferroan anorthosites and Mg-suite rocks, also reveal a much larger isotopic scatter than previously found, of about 1.2‰. The finding of large 57 Fe differences among different CAPTEM aliquot allocations of the same samples suggest that the common 0.1-1 g sample size allocations is insufficient to produce a powder representative of these coarsegrained rocks for Fe isotope measurements. It seems that feldspars tend towards heavy 57 Fe whereas olivines tend towards light 57 Fe values, possibly as a result of diffusive processes for the latter. Excluding these effects with highland rocks having more than 10% pyroxene modal content leads to a more precise highland 57 Fe of 0.094 ± 0.035‰. This value, which is based on six Mg-suite rocks but that is also within uncertainty reflected by three anorthosites and one troctolite, is indistinguishable from the mean upper lunar mantle value defined using the Fe isotope systematics of mare basalts. It likely represents the bulk Moon 57 Fe, which also agrees within uncertainties with previous estimates for the bulk Earth.

Combined with literature data, this Earth-like Fe isotope composition of the Moon, heavier by ~0.1‰ relative to chondrite parent bodies, Mars and Vesta, can be best explained as fingerprinting the Moon-forming giant impact. This process alone provides the most straightforward scenario to explain the observed isotopic pattern among planets, without necessarily requiring a high-pressure, high-temperature metal-silicate stable isotope fractionation event.

Figure captions:

Fig. 1: Initial planetary iron isotope composition ( 57 Fe) estimates for the Earth and Moon relative to IRMM-14 compared to chondrite and achondrite meteorite groups [START_REF] Poitrasson | Does planetary differentiation really fractionate iron isotopes? Earth Planet[END_REF]. The arrows depict some more recent revisions towards putatively lighter compositions for the Earth and Moon. See text for references and discussion.

Fig. 2: Iron isotope composition of bulk lunar rocks expressed in 57 Fe relative to IRMM-14.

The group means and 2 standard error uncertainties are reported as continuous and dashed lines/gray zones, respectively. The Earth reference value ( 57 Fe = 0.10±0.03‰; from [START_REF] Wiesli | Space weathering processes on airless bodies: Fe isotope fractionation in the lunar regolith[END_REF]. h The results presented were obtained from two separate sample aliquot dissolution and Fe purification to duplicate completely the measurements. i When several powder aliquots have been analyzed, the mean value of the sample was used for the calculation. j See the supplementary table for the extensive compilation used to compute this mean and uncertainty. When several values are available for a given rock, the mean of all these determination was used in the calulation to avoid overweighting samples with multiple determinations. k The iron isotope composition and two standard error (2SE) uncertainties quoted are calculated from the number of analyses indicated (n) and using the Student's t-correcting factors [START_REF] Platzner | Modern isotope ratio mass spectrometry[END_REF]. Electronic Annex Click here to download Electronic Annex: Supplementary Table .xlsx

Fig. 4 :

 4 Fig. 4: Relations between iron isotope composition estimates of planetary bodies against their

  Fig.1

  Figure1

  Figure2

  Figure3

Figure 4

 4 Figure 4

Table 1 represent replicate data obtained on a new sample aliquot from the same Apollo rock compared to previously published values. Among these, the new 57 Fe values for several samples (low-Ti mare basalts 12045 and 14053, high-

 1

	Ti mare basalts 10057, 70135 and 74275, KREEP-rich basaltic impact melt 14310, KREEP
	basalt 15386, Mg-suite norite 15455, anorthosites 15455 and 62255) are indistinguishable
	within uncertainties from those reported elsewhere

  [START_REF] Poitrasson | Iron isotope differences between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms[END_REF] 2013) is shown for comparison. Note that highland rocks have been split into anorthosites and Mg-suite rocks. Samples having extreme values or variable Fe isotope compositions on different NASA CAPTEM aliquots are labeled, besides KREEP basalt 15386. Data are from Table 1 and Supplementary Table. Iron isotope composition estimates for the Lunar Upper Mantle (0.142±0.026‰) and Bulk Moon (0.094±0.035‰) from this work are reported at the bottom. Note that previous estimates of the Bulk Moon 57 Fe values ranged from -0.1 to +0.2%, depending on the authors (see text and Fig. 1). Fig. 3: Iron isotope composition of bulk highland lunar rocks as a function of their pyroxene modal content. At low pyroxene modal contents the Fe isotope values are scattered, possibly due to the predominance of olivine affected by diffusive effects leading to low Fe isotope values or increasing plagioclase abundance having a heavy isotope composition. At higher pyroxene modal content, though, a homogeneous Fe isotope composition emerges at 0.094±0.035% (2SE) for these highland rocks. Data are from Table 1, Supplementary Table and the Lunar Sample Compendium[START_REF] Meyer | Lunar sample compendium[END_REF].

  Table for data sources.

Table 1 :

 1 New iron isotope composition of bulk lunar samples from this study, along with previously published data on the same samples

		Rock type	Rock mass	57 Fe (‰)	2SE k	56 Fe (‰)	2SE k	n
			powdered (g)
	Low-Ti basalts
	12009,130	Olivine vitrophyre basalt	0.856	0.101	0.031	0.045	0.020
	12011,29	Pigeonite basalt	1.003	0.097	0.081	0.066	0.056
	12031,37	Pigeonite basalt	0.824	0.154	0.071	0.098	0.066
	12038,244	Feldspathic basalt	0.874	0.131	0.056	0.079	0.027
	12045,30 a	Ilmenite basalt	1.032	0.136	0.094	0.082	0.062
	12045,13 b		~1	0.138	0.078	0.092	0.052
	12054,6	Ilmenite basalt	0.740	0.135	0.078	0.081	0.054
	14053,237	Group C Al-rich basalt	0.634	0.070	0.024	0.038	0.023
	14053,263 a		1.003	0.043	0.023	0.039	0.020
	14053 c		~0.05	0.130	0.030	0.100	0.030
	15555,988 a Olivine-normative basalt	1.097	0.086	0.064	0.056	0.041
	15555,115 b		~1	0.208	0.058	0.139	0.038
	15555 d		~0.1	0.044	0.026	0.029	0.017
	15555,955 e		~0.05-0.1	0.104	0.050	0.073	0.029
	15555,139 f		0.0053	0.10	0.02	0.05	0.04
	15555 c		~0.05	0.130	0.050	0.090	0.030
	15556,198	Olivine-normative basalt	0.601	0.117	0.071	0.061	0.046
	Mean low-Ti basalts (this study) i		0.113	0.023
	Mean low-Ti basalts j		0.127	0.012			27
	High-Ti basalts
	10003,178	Low-K Group B2 basalt	0.856	0.320	0.071	0.213	0.047
	10003 c		~0.05	0.200	0.040	0.130	0.040
	10044,638	Low-K Group B1 basalt	0.827	0.354	0.118	0.235	0.072
	10049,101 a High-K Group A basalt	1.124	0.167	0.118	0.114	0.073
	10057,268	High-K Group A basalt	0.599	0.189	0.028	0.158	0.038
	10057 c		~0.05	0.210	0.060	0.140	0.050
	10058,250	Low-K Group B1 basalt	0.852	0.290	0.040	0.198	0.030
	10092,14	Low-K Group B3 basalt	0.699	0.327	0.039	0.212	0.031
	70135,93	Type A basalt	1.048	0.238	0.061	0.146	0.045
	70135 c		~0.05	0.220	0.060	0.150	0.050
	70315,30	Type B1 basalt	0.991	0.325	0.050	0.196	0.034
	74235,62	Type B2 basalt	0.595	0.253	0.086	0.166	0.053
	74245,34	Type C basalt	0.604	0.213	0.079	0.107	0.059
	74275,48	Type C basalt	0.820	0.273	0.081	0.177	0.040
	74275,240 g		0.251	0.270	0.025	0.220	0.025
	78598,6	Type A basalt	0.622	0.324	0.039	0.235	0.037
	Mean high-Ti basalts (this study) i		0.273	0.039			12
	Mean high-Ti basalts j		0.274	0.020			25
	KREEP
	basalt 15386	KREEP basalt	0.544	0.183	0.067	0.112	0.039
	15386 b		~1	0.230	0.033	0.154	0.022
	15386 d		~0.1	0.210	0.080	0.140	0.053
	Highland rocks
	15415,190 a Ferroan anorthosite	0.998	0.496	0.057	0.347	0.113
	15455,306 a Mg-suite Norite (CAN clast)	1.061	0.070	0.017	0.024	0.025
	15455 c		~0.05	0.050	0.010	0.030	0.010

Table 60015

 60015 New dissolution from a powder prepared in Toulouse. b Poitrasson et al. (2004), c Sossi and Moynier (2017), d Weyer et al. (2005), e Liu et al. (2010), f Wang et al. (2015) and g

		,785	Cataclastic ferroan anorthosite	0.160	0.054	0.038
	60025,874 h Ferroan anorthosite	0.167	-0.393	0.050	-0.276	0.037	12
	60025 b		~1	0.225	0.063	0.150	0.042
	62255,191	Ferroan anorthosite	0.139	0.265	0.119	0.181	0.088
	62255,134 b		~1	0.176	0.062	0.118	0.042
	62275,21	Ferroan anorthosite	0.144	0.272	0.026	0.064	0.042
	67955,94	Mg-suite noritic anorthosite	0.166	0.156	0.049	0.060	0.114
	77215,258 a Mg-suite noritic	1.043	0.045	0.051	0.019	0.029
	77215 b		~1	0.128	0.024	0.085	0.016
	72415,82 h	Mg-suite dunite	0.127	-0.730	0.062	-0.446	0.085	12
	72415 f		0.0011	-0.50	0.03	-0.35	0.02	11
	72415 c		~0.05	-0.600	0.050	-0.400	0.040
	76335,63 a	Mg-suite troctolite	1.012	-0.007	0.062	-0.050	0.122
	Mean highland rocks (this study) i		0.023	0.235			10
	Mean highland rocks j		0.078	0.124			15
	Impact melts and breccia
	14310	KREEP Basaltic impact melt	2.027	0.125	0.130	0.086	0.069
	14310 d		~0.1	0.153	0.039	0.102	0.026
	Mean impact melts and breccia j		0.093	0.055
	Mean Moon			0.166	0.032			73
	j :	a

Acknowledgements

-This study was made possible through three separate allocations of lunar samples to CRN, FP and TM by CAPTEM. Carole Boucayrand, Jonathan Prunier and Manuel Henry are thanked for maintaining the clean lab in excellent condition, as well as Rémi Freydier and Jérôme Chmeleff for their efforts to have the Neptune MC-ICP-MS up running. This manuscript was substantially revised and expanded following detailed comments from Steve Elardo, two anonymous referees and GCA AE Munir Humayun. This work was funded by a grant from the Programme National de Planétologie (PNP) of CNRS-INSU to FP. CRN was supported by NASA grant NASA-NNX15AH76G.