Accepted Manuscript

Titanium stable isotope investigation of magmatic processes on the Earth and Moon

Marc-Alban Millet, Nicolas Dauphas, Nicolas D. Greber, Kevin W. Burton, Chris W. Dale
et al.

PII: S0012-821X(16)30264-3

DOI: http://dx.doi.org/10.1016/j.epsl.2016.05.039

Reference: EPSL 13872

To appear in: *Earth and Planetary Science Letters*

Received date: 22 July 2015

Revised date: 28 April 2016

Accepted date: 24 May 2016

<figure><img src="image_1.png" /></figure>

Please cite this article in press as: Millet, M.-A., et al. Titanium stable isotope investigation of magmatic processes on the Earth and Moon.
*Earth Planet. Sci. Lett.* (2016), http://dx.doi.org/10.1016/j.epsl.2016.05.039

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## Highlights

* First study of the Ti stable isotope composition of terrestrial and lunar rocks

* Ti isotopes display large variations during magma differentiation

* The terrestrial and lunar mantles display similar mantle Ti stable composition
1
Titanium stable isotope investigation of magmatic
processes on the Earth and Moon

3

4

5 Marc-Alban Millet¹,²*, Nicolas Dauphas², Nicolas D. Greber², Kevin W.

6 Burton¹, Chris W. Dale¹, Baptiste Debret¹, Colin Macpherson¹, Geoffrey M.

7 Nowell¹ and Helen M. Williams¹

8

9

10 ¹ Durham University, Department of Earth Sciences, South Road, Durham DH1 3LE,
11 United Kingdom

12

13 ² Origins Laboratory, Department of the Geophysical Sciences and Enrico Fermi
14 Institute, University of Chicago, 5734 S Ellis Avenue, Chicago IL, 60637, USA.

15

16 * Corresponding author: m.a.millet@durham.ac.uk

17

18

19

20

21 Keywords: Stable Isotopes, Magma differentiation, Bulk Silicate Earth, Lunar basalts
22 **Abstract**

23

24 We present titanium stable isotope measurements of terrestrial magmatic samples
25 and lunar mare basalts with the aims of constraining the composition of the lunar and
26 terrestrial mantles and evaluating the potential of Ti stable isotopes for understanding
27 magmatic processes. Relative to the OL-Ti isotope standard, the δ⁴⁹Ti values of
28 terrestrial samples vary from -0.05 to +0.55‰, whereas those of lunar mare basalts
29 vary from -0.01 to +0.03‰ (the precisions of the double spike Ti isotope
30 measurements are ca. ±0.02‰ at 95% confidence). The Ti stable isotope
31 compositions of differentiated terrestrial magmas define a well-defined positive
32 correlation with SiO₂ content, which appears to result from the fractional
33 crystallisation of Ti-bearing oxides with an inferred isotope fractionation factor of
34 Δ⁴⁹Tiₗoxide-melt = -0.23‰ × 10⁶/T². Primitive terrestrial basalts show no resolvable
35 Ti isotope variations and display similar values to mantle-derived samples (peridotite
36 and serpentinites), indicating that partial melting does not fractionate Ti stable
37 isotopes and that the Earth's mantle has a homogeneous δ⁴⁹Ti composition of
38 +0.005 ± 0.005 (95% c.i., n=30). Eclogites also display similar Ti stable isotope
39 compositions, suggesting that Ti is immobile during dehydration of subducted
40 oceanic lithosphere. Lunar basalts have variable δ⁴⁹Ti values; low-Ti mare basalts
41 have δ⁴⁹Ti values similar to that of the bulk silicate Earth (BSE) while high-Ti lunar
42 basalts display small enrichment in the heavy Ti isotopes. This is best interpreted in
43 terms of source heterogeneity resulting from Ti stable isotope fractionation
44 associated with ilmenite-melt equilibrium during the generation of the mantle source
45 of high-Ti lunar mare basalts. The similarity in δ⁴⁹Ti between terrestrial samples and
46 low-Ti lunar basalts provides strong evidence that the Earth and Moon have identical
47 stable Ti isotope compositions.
## 1. Introduction

Much of the stable isotope variation seen on Earth is mass-dependent in nature, scaling as a function of the difference in mass of the isotopes involved. Departures from such mass-dependent isotope fractionation are, however, commonplace in meteorites (Clayton, 1993; Dauphas et al., 2002) and one of the most intriguing features of the Earth-Moon system is the similarity of their mass-independent isotope signatures (e.g.: O: Clayton et al., 1973; Cr: Lugmair and Shukolyukov, 1998 Ti: Zhang et al., 2012;). This similarity is difficult to reconcile with simple giant impact models that predict that most of the Moon's mass should consist of impactor material, because in this case lunar rocks should have inherited the isotope composition of the impactor, which Pahlevan and Stevenson (2007) argued should have been different from that of the Earth. Different scenarios were proposed to explain this similarity ranging from isotope equilibration between Earth's mantle and the protolunar disk (Pahlevan and Stevenson, 2007), impact of a Mars-size body with a fast-spinning Earth (Cúk and Stewart, 2012), impact between two similar-sized protoplanets (Canup, 2012), "hit-and-run" impact (Reufer et al., 2012), or an Earth-like impactor (Dauphas et al., 2014a, Mastrobuono-Battisti et al., 2015). The wide range of physical processes invoked in each of these different models highlights the need to develop new tools to evaluate the consequences of each model on the chemical and isotopic evolution of the Moon with respect to the Earth.

Interestingly, and in contrast to many mass-independent isotope systems, the mass dependent stable isotope compositions of a number of elements in terrestrial and lunar samples display significant variations (e.g.: O and Fe: Liu et al., 2010; Poitrasson et al., 2004; Weyer et al., 2005; Mg: Sedaghatpour et al., 2013; Zn: Paniello et al., 2012; Li: Seith et al., 2006; Cl: Sharp et al., 2010; Cr: Bonnand et al., 2016). The stable isotope compositions of these elements have the potential to shed new light on the evolution of the Earth-Moon system as their variations are diagnostic features of a variety of processes such as volatile element depletion, core formation,
and magma ocean crystallization. However, many of these non-traditional stable isotope systems are affected by multiple processes, which complicate their interpretation.

Titanium has the advantage of being both lithophile and extremely refractory. It is therefore unlikely to have been affected by formation of the lunar and terrestrial cores, or volatile element depletion during the giant impact. It thus has the potential to provide unambiguous constraints on lunar magma ocean crystallization. Despite its abundance in igneous rocks and the extensive use of TiO₂ concentrations in high-temperature geochemistry, Ti stable isotopes have received very limited attention to date (Millet and Dauphas 2014; Zhang et al., 2014). Indeed, most of the studies so far focused on the detection of Ti isotopic anomalies to learn about nucleosynthetic processes, genetic relationships between planetary bodies, early solar system processes and/or cosmogenic effects (e.g.: Niederer et al., 1980; Niemeyer, 1988; Trinquier et al., 2009; Zhang et al., 2011; 2012). Titanium exists in several coordinations in magmatic systems: it is predominantly present in 5-fold coordination in all silicate melts but is also present in 4- and 6-fold coordinations in silicic and mafic melts respectively (Farges et al., 1996; Farges and Brown, 1997). In addition, 5-fold coordinated titanium transitions to 6-fold coordination during crystallisation of Ti-bearing oxides (Farges and Brown, 1997). As stable isotope theory (Schauble, 2004) predicts that stable isotope fractionation will be driven by major contrasts in elemental bonding environment, the coordination behaviour of Ti raises the possibility that Ti stable isotopes may serve as a tracer of magmatic processes. Furthermore, as Ti is refractory and comparably immobile in fluids, Ti stable isotopes are also likely to be comparatively resistant to overprinting. A potential complexity however, lies with the presence of Ti³⁺ in reduced magmas and planetary bodies such as the Moon. fO₂ estimates for the lunar mantle range from the iron-wüstite buffer to 2 log units below it (i.e. IW-2), with a most likely value of ~IW-1 (Wadhwa, 2008). Experimental work has shown that in these conditions, the lunar mantle could contain
up to 10% Ti³⁺ (Krawczynski et al., 2009). However, direct measurements of Ti³⁺/ΣTi (Simon et al., 2014) in pyroxenes and olivines from lunar basalts have not detected any Ti³⁺.

In this contribution, high precision Ti stable isotope measurements are presented for a range of terrestrial and lunar magmatic rocks with three aims. The objectives of this contribution are to (i) evaluate the effect of fractional crystallisation and magma differentiation on the Ti stable isotope composition of silicate melts; (ii) determine the Ti stable isotope composition of the terrestrial mantle and assess its homogeneity and (iii), use the Ti stable isotope composition of lunar basalts to investigate the magmatic evolution of the Moon.

## 2. Samples and methods

### 2.1 Samples

In addition to the dataset already presented in Millet and Dauphas (2014), we have analysed 34 terrestrial samples reflecting a global coverage from a diverse range of geodynamic contexts. We have also analysed 8 primitive lunar mare basalts covering the range of TiO₂ concentrations observed in lunar rocks.

Mid-Ocean Ridge Basalts (MORB): Seven MORB samples were analysed in order to constrain the composition of the upper mantle. The sample suite includes glasses from the mid-Atlantic ridge, the East-Pacific Rise (Batiza and Niu, 1992), the Pacific-Cocos-Nazca triple junction (Puchelt and Emmerman, 1983) as well as the South-West Indian Ridge (Escrig et al., 2004), including a single MORB sample displaying Sr, Pb and Os isotope compositions characteristic of the Dupal anomaly (MD57 9-1).

Island Arc Basalts. The island arc basalts studied here mainly comprise basalts sampled from oceanic arcs in order to limit the potential for crustal contamination and assimilation processes. In addition to the New Britain basalt data published by Millet
and Dauphas (2014), 3 arc lavas from the Mariana arc were studied that were previously analysed for major and trace elements by Yi et al. (2000). A single basalt from the Izu-Bonin arc, rock standard JB-2, was also analysed.

*Intraplate basalts*. The intraplate basalts mainly comprise ocean island basalts (OIBs) from four different localities. Single samples from Hawaii (BHVO-2 rock standard) and Easter Island (sample 17678, Baker et al., 1974) were analysed as well as samples from the Cape Verde (São Nicolau island; Millet et al., 2008) and Azores archipelagos (São Miguel island: Turner et al., 1997; São Jorge island: Millet et al., 2009). In addition, a single continental flood basalt (the USGS rock standard BCR-2) was also analysed.

*Eclogites*. The three eclogites measured here were collected from the Zermatt-Saas Fee ophiolite, Switzerland. They consist of 2 basaltic eclogites and 1 gabbroic eclogite, based on their mineralogy (Dale et al., 2007). Estimates of peak metamorphic conditions for these samples range between 2-3 GPa and 550 to 630 °C (Barnicoat and Fry, 1986; Reinecke, 1991). Major and trace element data (Dale et al., 2007) show that all the samples selected suffered loss of volatile and fluid-mobile elements during metamorphic dehydration.

*Mantle samples*. The mantle-derived samples are of two types. First, a single orogenic peridotite from the Beni Bousera massif was analyzed (GP13, Pearson et al., 2004). The other 3 mantle-derived samples are serpentinites from various Western Alps ophiolites. These samples originate from the lithospheric mantle section of a subducted slab and have been chosen because they record different metamorphic conditions during subduction. These have partly (BCh9, MM15) to fully (LZ14b) re-equilibrated in the antigorite (high temperature and pressure variety of serpentine) stability field during subduction (see Debret et al., 2014 for details).
160

161 *Differentiated magmas*. In order to assess the effect of magmatic differentiation, rock standards of varying SiO₂ content (AGV-1 andesite: 58.8wt%, G-2 granite: 69.1wt%, RGM-1 rhyolite: 73.4wt%) and a basaltic andesite from São Miguel Island (54.5wt% SiO₂, Azores archipelago) were analysed. Although not cogenetic, these samples span almost the full range of silica content observed in terrestrial rocks and show a continuous decrease in TiO₂ content with increasing SiO₂. To complement this sample set, 6 cogenetic samples from Agung volcano (Bali, Indonesia) related to each other by fractional crystallisation were analysed. The SiO₂ concentration of these samples ranges from 54 to ca. 63wt% and TiO₂ shows a continuous decrease from 0.92 to 0.60wt% over that range (Dempsey, 2012).

171

172 *Lunar samples*. Nine lunar mare basalts were analysed. Specifically, we selected 5 high-Ti basalts (TiO₂ ranging from 12.2 to 13.4wt%) and 3 low-Ti basalts (TiO₂ ranging from 1.9 to 3.3wt%). A single green glass sample was also analysed (sample 15426, TiO₂ = 0.5 wt%).

176

177 **2.2. Methods**

178 Samples were processed and measured using the double-spike method of 179 Millet and Dauphas (2014), which is briefly outlined here. Between 10 to 50 mg of 180 rock powder (for whole rock samples) or glass chips (for MORB glasses) were 181 digested in a 1:1 mixture of concentrated HF and HNO₃ for 48 hours. After careful 182 evaporation, samples were then taken up in nitric acid and dried down 3 times before 183 being taken up in 6 M HCl and checked for residual solids. If samples were fully 184 digested, approximately 30 mg of H₃BO₃ was added to the solution in order to ensure 185 that any potential fluorides (for which Ti has a strong affinity) having escaped visual 186 inspection are re-dissolved and thus ensure that all the Ti contained in the sample is 187 in solution. An aliquot containing 5 to 20 µg of Ti is then taken and mixed with a ⁴⁷Ti-
¹⁸⁸ ⁴⁹Ti double spike in appropriate proportions (the Ti concentrations of the various rocks analysed were known from prior work). Purification of Ti was achieved following the procedure designed by Zhang et al. (2011) and using a two-step chemistry. Firstly, the samples were passed through a TODGA cartridge to remove almost all sample matrix. Samples were then further purified using AG1-X8 resin in order to separate Mo (thus limiting potential doubly charged interferences on Ti isotopes) and remaining Ca. It is notable that lunar samples and BHVO-2 rock standard were digested at the Origins Laboratory of the University of Chicago by flux fusion method using LiBO₃ as a fluxing agent. This was to ensure that potential refractory phases were fully digested. After fusion, samples are dissolved in 3M HNO₃ before an aliquot containing 20µg of Ti is taken and spiked in ideal proportions. Chemical purification of Ti was achieved using the procedure outlined above.

201 Isotope ratio measurements were carried out on Neptune Plus MC-ICP-MS installed in the Arthur Holmes laboratory at Durham University and at the Origins Laboratory of the University of Chicago. The samples were injected into the plasma torch using an Aridus II desolvating nebuliser and the isotope measurements were performed in medium resolution mode. Despite being double-spiked, all sample measurements were bracketed by measurements of double-spiked standards (OL-Ti) measured at the same concentration level in order to account for small unresolved polyatomic interferences on ⁴⁷Ti and ⁴⁸Ti produced in the mass spectrometer. Raw data is then processed offline using in-house double-spike deconvolution codes written in Mathematica®. All data is expressed as δ⁴⁹Ti, which is the deviation of the ⁴⁹Ti/⁴⁷Ti ratio of samples relative to that of the OL-Ti standard (Millet and Dauphas, 2014). This standard is available from the corresponding author upon request.

213 Finally, it should be noted that double-spike measurements rely on the assumption that the 4 isotopes used in the double-spike deconvolution routine are related to by mass-dependent stable isotope fractionation only. Double-spike method
can provide inaccurate result if any of the isotopes involved is affected by mass-independent variation. To this day, no mass-independent variations have been found in terrestrial samples for Ti isotopes, but significant $^{50}$Ti anomalies have been detected in lunar samples related to cosmic ray exposure. These anomalies could lead to potentially inaccurate results if not accounted for. For this reason, our double-spike deconvolution procedure uses $^{46}$Ti, $^{47}$Ti, $^{48}$Ti, and $^{49}$Ti that are only related to each other by mass-dependent fractionation in both terrestrial and lunar rocks.

223

## 4. Results

Data for all standards and samples analysed during the course of this study are given in Table 1. For samples measured once only, quoted errors are internal errors, whereas for samples measured multiple times, 95% confidence intervals were calculated using the Isoplot software (Ludwig, 2003). Repeated digestion and analysis of BCR-2, BHVO-2 (HF or flux fusion digestion) and JB-2 reference materials, previously analysed by Millet and Dauphas (2014) at the Origins Laboratory (University of Chicago), were conducted in order to certify data quality and the absence of inter-laboratory bias (Fig. 1). The average Ti isotope compositions of standards analysed at Durham University is in excellent agreement with that of the Origins Laboratory data and demonstrates that the 2σ reproducibility of our $\delta^{49}$Ti measurements is ca. ±0.020‰. In addition, we have carried out replicate digestions and analysis of the RGM-1 rock standard (only in Durham) which all display values within analytical uncertainty (0.548±0.014‰, 95%c.i., n=4). Finally, for BHVO-2 rock standard, measurements of flux fusion digestions carried out in Chicago displays value within analytical uncertainty of HF-HNO$_3$ digestions performed in both Durham and Chicago (Greber et al., 2016; Fig. 1).

225

226

227

228

229

230

231

232

233

234

235

236

237

238

239

240

241

The $\delta^{49}$Ti values of all terrestrial samples ranges from -0.046 to +0.548‰.

242 Samples of differentiated magmas display a large range in Ti isotope compositions
243 and an overall enrichment in isotopically heavy Ti isotopes (+0.054 to +0.548‰)
244 relative to basalts and mantle-derived rocks (-0.046 to +0.049‰), which display more
245 homogeneous compositions. In the MORB, OIB and IAB sample subsets, there
246 appear to be no differences related to sampling localities.

247 Unlike primitive terrestrial samples, primitive lunar basalts display small but
248 resolvable Ti stable isotope heterogeneity, which relates to their TiO₂ content. High-
249 Ti basalts display overall heavier δ⁴⁹Ti values (+0.011‰ to +0.033‰) relative to
250 those of low-Ti lunar basalts and the single green glass sample analysed (-0.008 to
251 +0.011‰).

252

253 **5. Discussion**

254 Significant variations exist in the Ti stable isotope compositions of terrestrial
255 and lunar magmatic samples (Table 1). Below, we discuss the probable causes of Ti
256 stable isotope fractionation and evaluate the degree of homogeneity of the Earth's
257 mantle in terms of Ti stable isotopes. We then focus on the variability observed in the
258 Ti stable isotope composition of primitive lunar mare basalts and discuss the
259 implications of these data for the magmatic evolution of the Moon.

260

261 **5.1 Ti stable isotope fractionation during magma differentiation**

262 A striking feature of our terrestrial sample dataset is the significant enrichment
263 of differentiated magmatic rocks in isotopically heavy Ti isotopes relative to more
264 primitive magmatic rocks. This enrichment is correlated with SiO₂ content, suggesting
265 a relationship with fractional crystallisation processes (Fig. 2). Titanium is highly to
266 moderately incompatible in most of the silicate minerals typically involved in magma
267 differentiation (pyroxenes, olivine, plagioclase, micas, quartz) regardless of the
268 composition of the melt. Mass balance considerations therefore suggest that it is
269 unlikely that the crystallisation of silicate mineral phases could have modified the Ti
270 isotopic composition of the remaining melt.
271 Titanium behaviour during fractional crystallisation is mainly controlled by the
272 crystallisation of Fe-Ti oxides that can occur at all stages of magma differentiation.
273 The coordination of Ti in silicate melts and oxide minerals was investigated by
274 Farges et al. (1996) and Farges and Brown (1997). In silicate melts, Ti is
275 predominantly present in 4, 5, and 6-fold coordinations, whereas Fe-Ti oxides
276 exclusively accommodate 6-folded Ti (Farges et al., 1996). As stable isotope theory
277 (Schauble, 2004) predicts that isotopically heavy species will be preferentially
278 concentrated in low-coordination, high-force constant (stronger and stiffer) bonding
279 environments, at equilibrium Fe-Ti oxides might thus display enrichments in
280 isotopically light Ti relative to the co-existing melt. The progressive crystallisation of
281 Fe-Ti oxides during magmatic differentiation could potentially, therefore, drive the
282 residual melt to heavy isotope compositions.

283 In order to test this hypothesis, we have modelled the evolution of a primitive
284 basaltic melt to a silicic composition by fractional crystallisation using Rhyolite
285 MELTS software (Gualda et al., 2012). The composition of the starting material is set
286 to be similar to that of arc basalts (see Fig. 2 caption), crystallisation occurs at 1 kbar
287 and the relative oxygen fugacity is assumed to be constant at the NNO buffer. At
288 each calculation step, the compositions and respective mass of melt and minerals in
289 equilibrium are calculated. Steps are defined by a 5°C incremental temperature
290 decrease from the liquidus temperature (here 1118°C). The calculation stops when
291 the temperature reaches 900°C, at which point ca. 70% of the liquid has crystallised
292 and the SiO₂ content of the remaining liquid has reached 66 wt%. The Ti isotope
293 evolution of the melt is calculated, at each step, using a Rayleigh distillation law and
294 a bulk Ti stable isotope fractionation factor between the residual melt and the
295 minerals removed. As silicate minerals have a negligible effect on the budget of Ti in
296 differentiating magmas, their contribution to the Ti isotope evolution of the residual
297 melt must also be negligible. Consequently, the bulk melt-residue Ti stable isotope
298 fractionation factor is equal to that between Fe-Ti oxides and melt
(Δ⁴⁹Tiᵢₗₐ₉₆-ₐ₈₄) weighted according to the relative proportion of Ti sequestered by oxides at each temperature step. Assuming that the fractionation is equilibrium in nature, the model reproduces the observed trend in δ⁴⁹Ti vs. SiO₂ well for an empirical value of Δ⁴⁹Tiᵢₗₐ₉₆-ₐ₈₄ = -0.23‰ × 10⁶/T² (with T in K) indicating that fractional crystallisation of Fe-Ti oxides is a viable mechanism to explain the progressive enrichments of magmas in isotopically heavy Ti with increasing silica content. The predicted fractionation between oxide and melt at 900 °C is -0.167‰ for δ⁴⁹Ti. For comparison, Nuclear Resonant Inelastic X-ray Scattering (NRIXS) force constant measurements predict an equilibrium iron isotope fractionation between ilmenite (Krawczynski et al., 2014) and Fe²⁺-bearing basaltic glasses (Dauphas et al. 2014b) at the same temperature of -0.09‰ for δ⁵⁶Fe (i.e., -0.05 ‰/amu). The equilibrium Ti isotope fractionation factor between oxide and melt that is needed to account for Ti isotope variations in silicic rocks is thus very reasonable. It should be noted that our preferred model does not account for any potential dependence of the Ti stable isotope fractionation factor on melt or Fe-Ti oxide chemical composition. Experimental work as well as measurements of oxide minerals in well-characterised sample suites will be needed to ascertain the oxide-melt isotopic fractionation factor inferred from measurements of differentiated terrestrial rocks.

The origin of the heavy Fe isotopic composition of silicic rocks with >70wt% has been the subject of some debate. The explanations proposed thus far include iron mobilization by exsolved fluids (Poitrasson and Freydier, 2005; Heimann et al., 2008; Telus et al., 2012), Soret migration (Lundstrom, 2009), and fractional crystallization (Telus et al. 2012; Sossi et al. 2012; Dauphas et al. 2014b). For most silicic rocks, fluid exsolution is unlikely to be the culprit because more fluid mobile Zn does not display correlated fractionation and iron isotope enrichments also affect anhydrous A-type granites (Telus et al., 2012; Sossi et al. 2012). The finding of significant isotope variations for a fluid-immobile element like Ti in silicic magmatic
rocks supports the view that fractional crystallization can drive stable isotope
fractionations in transition metals. Iron and titanium stable isotope variations in
magmatic rocks probably bear considerable insights into the processes governing
magmatic evolution but equilibrium fractionation factors between minerals and melts
need to be better known to reap the fruits of those studies.

331

332 5.2 The Ti stable isotope composition of terrestrial basalts and the Bulk Silicate Earth

333 Basaltic magmas display less variation in δ⁴⁹Ti relative to differentiated
334 magmas (ca. 0.1‰ compared to ca. 0.6‰; Fig. 3). The variation observed is
335 statistically significant (relative to analytical precision) and may be related to
336 heterogeneity of the terrestrial mantle or other processes. However, the full range of
337 basalt δ⁴⁹Ti variability is also observed in subduction zone basalts at the high end of
338 the SiO₂ range (52-54 wt%). Two of the basaltic lavas from the Marianas and one
339 New Britain sample display the heaviest Ti stable isotope compositions of all the
340 basalts analysed, whereas the JB-2 rock standard (Izu-Bonin arc) displays the
341 lightest composition of the entire sample set. Other basalt samples from the New
342 Britain and Mariana arcs, which have lower silica contents, all display δ⁴⁹Ti values
343 within analytical uncertainties, indicating that these three samples may have already
344 been affected by the early onset of Fe-Ti oxide fractionation and are thus not
345 representative of the Ti stable isotope composition of the sub arc mantle anymore.

346 Fractional crystallisation cannot however explain the composition of the JB-2 rock
347 standard, as no mineral phase typically involved during basaltic melt differentiation
348 (e.g. olivine, pyroxenes, plagioclase, spinel, garnet) has been identified to
349 preferentially incorporate the heavy isotopes of titanium relative to melt at
350 equilibrium. As an arc basalt, JB-2 is expected to display anomalously low
351 abundances of fluid immobile elements such as Ti and Nb relative to elements of
352 similar incompatibility (Sm and Tb for Ti, Th and La for Nb; $Ti/Ti^* = Ti_N / \sqrt{Sm_N \times Tb_N}$
and $Nb/Nb^* = Nb_N/\sqrt{Th_N \times La_N}$, concentrations are normalised to the primitive mantle composition of McDonough and Sun, 1995) in normalised trace element patterns. For example, compiled data for the Izu-Bonin arc eruptives (obtained from Georoc database) show that at similar $SiO_2$ contents to JB-2 (53.2wt%), Izu-Bonin magmas typically have Ti/Ti* of 0.7 to 0.85 and Nb/Nb* of 0.16 to 0.35. While JB-2 displays the strong negative Nb anomaly ($Nb/Nb^* = 0.21$) expected for island arc lavas, it does not display a negative Ti anomaly ($Ti/Ti^* = 1.02$). This feature can be explained by oxide accumulation in this sample. In contrast to Ti, Nb does not partition into Fe-Ti oxides (Nielsen and Beard, 2000) and therefore the accumulation of Fe-Ti oxides would lead to decoupling of Ti and Nb concentrations as well as enrichment in light isotopes of titanium.

Once samples that are affected by Fe-Ti oxide crystallisation or accumulation are removed (i.e. most silicic lavas from the Mariana arc, the most silicic New Britain sample and JB-2 rock standard), the average $\delta^{49}$Ti values for island arc (+0.007±0.010‰, 95% c.i., n=8), intraplate (+0.009±0.019‰, 95% c.i., n=7) and mid-ocean ridge basalts (+0.001±0.008‰, 95% c.i., n=7) are all within statistical error of each other and within our analytical uncertainty (see Fig.1). Titanium behaves as a moderately incompatible element during melting of either spinel or garnet-bearing peridotite (Prytulak and Elliott, 2007) and is efficiently extracted from the mantle during basalt genesis. Mass-balance considerations thus dictate that basalts should record the Ti isotopic composition of their mantle source, unless melting occurs under Ti-oxide saturation. Our data on primitive basalts therefore indicates that the Earth's upper mantle has a homogeneous Ti stable isotope composition, within the precision of our measurements. In addition, the fact that the average value of our mantle samples (Beni Bousera Orogenic peridotite and Western Alps serpentinites) is indistinguishable from the average of all the primitive basalts analysed ($\delta^{49}$Ti$_{basalts}$=+0.006±0.006‰, 95% c.i., n=22 vs. $\delta^{49}$Ti$_{mantle}$ =+0.013±0.023‰, 95% c.i.,
n=4) suggests that the partial melting of typical mantle lithologies (olivine + clinopyroxene + orthopyroxene + spinel and/or garnet) does not fractionate Ti stable isotopes.

In the bulk silicate Earth, Ti is mainly hosted in the mantle with only a minor proportion stored the continental crust (McDonough, 1991). It is therefore possible to calculate the Ti stable isotope composition of the Bulk Silicate Earth as the variance-weighted mean of primitive basalts, eclogites and mantle-derived samples analysed in this study, which yields a $\delta^{49}$Ti value for the BSE (bulk silicate Earth) of +0.005±0.005‰ (95% c.i., n=29). It is worth noting that measurement of a fully processed OL-Ti standard (digestion + chemical purification) shows that the analytical method used generates accurate results at the ~±0.012‰ precision level (Millet and Dauphas, 2014), so it is unknown if the $\delta^{49}$Ti of the BSE value is accurate at the ±0.005‰ level.

It is notable that primitive island arc basalts and MORBs display the same Ti stable isotope compositions. The mobility of Ti and other high-field-strength elements in subduction zone systems is a highly debated subject that mainly revolves around the stability of Ti-bearing phases during metamorphism of the down-going slab and the nature and composition of fluids in associated dehydration (e.g. Kessel et al., 2005). Experimental constraints have shown that rutile has low solubility in pure-H₂O fluids (Audetat and Keppler, 2005) and that the rutile-aqueous fluid partition coefficients for Ti, Nb and Ta are very high (Brenan et al., 1994). However, the presence of rutile in fluid-related mineral veins of deep-subducted rocks (Gao et al., 2007) indicates potential mobility of Ti and other HFSE in subduction zone environments. This is corroborated by recent experimental work on fluorine and chlorine-rich fluids (Rapp et al., 2010) and albite-saturated fluids (Antignano and Manning, 2008) in which Ti and other HFSE appear to be orders of magnitude more soluble than in pure-H₂O fluids. It is unclear in what speciation Ti is present in such
fluids and it is thus hard to predict the direction and magnitude of stable isotope
fractionation between rutile and co-existing fluids. Nevertheless, the bonding
environment of Ti in halogen or Na-bearing fluids will be different to that of rutile and
it is expected that this process will generate stable isotope fractionation. The identical
average Ti stable isotope composition of MORBs and island-arc basalts therefore
indicates that, even if Ti is mobilised during dehydration in subduction zones, this
process does not affect the budget of Ti across the whole mantle wedge. Moreover,
the fact that the eclogites and subduction-related serpentinites measured during this
study display the same Ti isotope composition as MORBs (δ⁴⁹Tieclogites=-
0.014±0.015‰, 95% c.i., n=3 and δ⁴⁹Tiserpentinites=+0.016±0.043‰, 95% c.i., n=3)
despite having been extensively dehydrated argues against significant mobility of Ti
in subduction zones and hints that recycling of oceanic lithosphere may not generate
detectable Ti stable isotope variation within Earth's mantle.

420

421
5.3 Titanium stable isotope heterogeneity of the lunar mantle as a consequence of
422
the LMO crystallisation

Although limited in range, it appears that lunar mare basalts display slightly
variable Ti stable isotope composition (Fig. 3). This variability is due to the heavier
isotope composition of the high-Ti lunar basalts (Fig. 3). Titanium is an extremely
refractory element that only partitions into gaseous phases at extremely high
temperatures (Zhang et al., 2014). It is therefore extremely unlikely that the
difference between the two classes of lunar basalts is a consequence of evaporation
during eruptive processes, neither can it be due to silicate-metal segregation as core
formation in both the Earth and Moon took place under conditions that were too
oxidizing for titanium to partition into the core (Wade and Wood, 2001). The titanium
stable isotope variability observed in lunar basalts must therefore be related to
processes that took place during the magmatic evolution of the Moon.
The source of High-Ti basalts are thought to be genetically linked to ilmenite-bearing cumulates formed between 95% and 99.5% solidification of the LMO (e.g. Snyder et al., 1992) as opposed to the sources of Low-Ti basalts that are formed earlier (<85% solidification) in the solidification sequence of the LMO before the onset of ilmenite crystallization (e.g. Hallis et al., 2014). Other minerals present in lunar mantle sources are olivine and pyroxenes, in which Ti is highly to moderately incompatible. Low-Ti lunar basalts display $\delta^{49}$Ti values that are indistinguishable from that of the Earth's mantle. Results on terrestrial basalts undersaturated in Fe-Ti oxides ($SiO_2$ < 52wt% in fig. 2) as well as mantle samples shows that olivine and pyroxenes do not fractionate Ti stable isotopes during fractional crystallisation or partial melting (Farges et al., 1996). Thus, low-Ti lunar basalts provide the best estimate of the Ti stable isotope composition of the lunar mantle, which is $\delta^{49}$Ti=-0.003±0.014‰ (n=4) which is within statistical error to that of the Earth's mantle.
Importantly, preliminary data for ordinary, enstatite and carbonaceous chondrites (Williams et al., 2015; Greber et al., 2016) also show uniform $\delta^{49}$Ti within error of the Earth's mantle ($\delta^{49}$Ti=+0.004±0.010‰; Greber et al., 2016). Taken overall, the data for terrestrial, lunar and meteoritic samples available to date is consistent with recent models arguing for an Earth-like composition of the Moon-forming impactor (Dauphas et al., 2014a; Mastrobuono-Battisti et al., 2015).

As a Fe-Ti oxide, ilmenite hosts six-fold Ti in its structure and is thus expected to preferentially incorporate light isotopes relative to melts at equilibrium as can be inferred from measurements of terrestrial silicic rocks. This behaviour provides an opportunity to test the models proposed for the generation of high-Ti lunar basalts. If all models agree on the involvement of an ilmenite-pyroxene cumulate layer in the late stages of LMO solidification, the details of the signatures of those cumulates were incorporated in high-Ti mare basalts is still a matter of debate. One view argues that high-Ti lunar basalts are derived from low-Ti lunar basalts through assimilation of
461 ilmenite and pyroxene during magma ascent to the surface (Wagner and Grove,
462 1997). However, experimental work regarding the kinetics of assimilation of pyroxene
463 and ilmenite suggests that ilmenite does not dissolve fast enough relative to
464 pyroxene to generate compositions similar to high-Ti basalts (Van Orman and Grove,
465 2000) and would instead generate melt composition too rich in calcium. The other
466 view involves the fertilisation of deep-seated ilmenite-free cumulates (i.e. formed
467 earlier during the solidification of the LMO) by sinking ilmenite-bearing cumulates
468 (Hess and Parmentier, 1995; Hallis et al., 2014). Concerns regarding the buoyancy of
469 magmas generated by such sources as well as the elevated viscosity of the ilmenite-
470 rich layer have lead to the formulation of a third model. Instead of solid-state mixing,
471 this last model argues for fertilisation of the ilmenite-free cumulates by negatively
472 buoyant partial melts of the ilmenite-rich layer formed as a result of the late heavy
473 bombardment (Van Orman and Grove, 2000; Elkins-Tanton et al., 2002, 2004).

474 In order to test the solid-state mixing and the partial-melt fertilisation models,
475 we modelled the evolution of the LMO during the crystallisation of the ilmenite-rich
476 layer and examined its impact on the Ti stable isotope composition of high-Ti lunar
477 melts. We assume a Ti concentration of 3wt% at 95% PCS and mineral proportions
478 similar to Snyder et al. (1992). At all steps of calculations, we calculate the Ti
479 concentrations as well as the δ⁴⁹Ti of the residual melt, instantaneous cumulates as
480 well as the bulk cumulate. We use the oxide-melt Ti isotope fractionation factor
481 calculated in section 5.1 as a proxy for the ilmenite-melt fractionation factor. A simple
482 Rayleigh distillation model (T=1125°C; Van Orman and Grove, 2000) allows us to
483 estimate the Ti stable isotope composition of the ilmenite-rich layer and evolving
484 magma (Fig. 4). Calculations are stopped at 99.5% solidification, at which point
485 0.25wt% Ti remains in the liquid. Overall, modelled ilmenite-bearing cumulates vary
486 from -0.107‰ to +0.175‰. The remaining liquid at the end of ilmenite crystallisation
487 is strongly enriched in heavy isotopes (δ⁴⁹Ti=+0.288‰).
For the solid-state mixing model, we followed calculations by Hallis et al. (2014) who argued that high-Ti lunar basalt sources formed by mixing ilmenite-free cumulate made at 80% PCS with cumulates formed at 95%PCS (earliest ilmenite cumulates) in 80:20 proportions, this mixture is then topped up with 1% of trapped instantaneous residual liquid (TIRL at 95%PCS). Assuming Ti concentrations of 0.5wt% and 6wt% for the respective cumulates and 2.6wt% for the liquid (based on our LMO solidification model), it is possible to estimate a δ⁴⁹Ti of the high-Ti lunar basalts mantle sources of -0.068‰. Interestingly, solid-state mixing models for the generation of high-Ti lunar basalts imply that ilmenite is exhausted during the generation of high-Ti melts (Ringwood and Kesson, 1976; Elkins-Tanton et al., 2002).
Mass balance suggests that melts generated this way should display a δ⁴⁹Ti identical to that of their source and therefore lighter than low-Ti lunar basalts, contrary to our measurements. However, if ilmenite was not exhausted during partial melting, high-Ti lunar basalts should display heavier isotope composition than this modelled source.
Uncertainty on the kinetics of ilmenite dissolution during lunar mantle melting to make high-Ti mare basalts makes further modelling difficult. Nevertheless, if ilmenite was a residual phase during generation of high-Ti magma, one may expect a negative relationship between Ti content and δ⁴⁹Ti of high-Ti lunar basalts. Such relationship does not appear in our data but this may be due to the relatively homogeneous TiO₂ content of the samples measured here (12.2 to 13.4wt%) compared to the full range shown by high-Ti lunar basalts.

An alternative to direct mixing of isotopically light ilmenite-bearing cumulates into the sources of low-Ti lunar basalts is the fertilisation of these sources by negatively buoyant partial melts of these cumulates. This partial melting may have occurred as a result of the late heavy bombardment, directly through shock melting but also by adiabatic melting cause by the incurred mantle convection. Regardless of the process involved, our results predict that these partial melts should be enriched in
heavy isotopes of Ti relative to the residue. Quantitative constraints on the
composition of these melts is currently lacking but as such, this model may provide a
straightforward way to generate lunar mantle source with both elevated Ti content
and heavy Ti stable isotope composition

519

520 **6. Conclusion**

This study presents the first investigation of the stable Ti isotope compositions of
terrestrial and lunar igneous rocks. The main conclusions of this study are:

• The Ti stable isotope composition of terrestrial magmas shows a large variability that appears to be positively correlated with SiO₂ content. This is most likely the result of crystallisation of isotopically light Ti-oxides during magmatic differentiation with Δ⁴⁹Tiᵢₗₓide-melt = -0.23‰ × 10⁶/T²; -0.106‰ at 1,200 °C). This observed fractionation is in agreement with the relative coordination of Ti between oxide minerals and silicate melts but further experimental data is needed to ascertain this value and assess whether it is equilibrium or kinetic in nature.

• The average δ⁴⁹Ti values of primitive mid-ocean ridge, island-arc and intraplate basalts are identical within our analytical uncertainties (ca. ±0.020‰). In addition, a preliminary set of mantle-derived samples also show δ⁴⁹Ti values within error of primitive terrestrial basalts. This demonstrates that little Ti stable isotope fractionation occurs during partial melting, and suggests that Earth's mantle (and the by extension the bulk Earth because no Ti is in the core) has a homogeneous Ti stable isotope composition of δ⁴⁹Tiₐ₃ₐ₆=+0.005±0.005‰ (95% c.i., n=29).

• The lack of any significant fractionation of Ti stable isotopes between MORBs, and island-arc basalts, as well as eclogites and serpentinites form subduction zones argues against a significant mobility of Ti in fluorine or chlorine bearing fluids across the mantle wedge during dehydration of downgoing slabs.
541 • Finally, primitive lunar basalts possess Ti stable isotope compositions ranging from terrestrial values to slightly enriched in heavy isotopes (δ⁴⁹Ti up to +0.033±0.015‰). Low-Ti lunar basalts all display δ⁴⁹Ti values within error of the terrestrial mantle value whereas high-Ti lunar basalts display small but distinct enrichment in isotopically heavy Ti. The heavy δ⁴⁹Ti values recorded in high-Ti mare basalts indicate that their mantle source regions may have been fertilised either by ilmenite cumulates formed in the latest stages of the LMO or by negatively buoyant partial melts of ilmenite-bearing cumulates during the Late Heavy Bombardment.

549

550 Acknowledgements

551 MAM acknowledges financial support from Durham University and a Marie Curie 552 COFUND International Junior Fellowship held at Durham University. ND was 553 supported by grants from NSF (EAR 1144429, 1444951, 1502591) and NASA 554 (NNX12AH60G, NNX14AK09G, NNX15AJ25G). NDG was supported by the Swiss 555 National Science Foundation (grant P2BEP2-158983). HMW acknowledges support 556 from NERC Advanced Fellowship NE/F014295/2 and the ERC (Consolidator Grant 557 "HabitablePlanet" 306655).

558
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836 **Captions**

837

838 Table 1: Ti stable isotope composition of samples measured during the course of this
839 study. Data presented for samples replicated multiple times are weighted means and
840 95% c.i. calculated using Isoplot (Ludwig, 2003). For rock standards BHVO-2, BCR-2
841 and JB2, averages represent the combination of data obtained by Millet and
842 Dauphas (2014) and this study. For BHVO-2, measurement made after flux fusion
843 digestion carried out in Chicago is represented as a filled diamond.

844

845 **Figure 1:** δ<sup>49</sup>Ti values of BHVO-2, BCR-2 and JB2 basaltic rock standards measured
during the course of this study. Filled symbols represent data generated at the
846 Origins Laboratory and published in Millet and Dauphas (2014), while open symbols
847 represent data obtained at Durham University. Data for all rock standards are in
848 excellent agreement between the two laboratories.

850

851 **Figure 2:** Evolution of δ<sup>49</sup>Ti values with SiO<sub>2</sub> concentration (wt%) of magmatic rocks.
852 Arc basalts are in black squares whereas samples form the Agung volcano
853 differentiation suite are in white squares. Other differentiated magmatic rocks are in
854 grey circles. The budget of titanium during magma differentiation is controlled by the
855 ongoing fractional crystallisation of isotopically light Ti oxides. Black curve represents
856 the fractional crystallisation model generated using Rhyolite-Melts (see full text for
857 details) and designed to fit the Agung Volcano cogenetic sample set. Starting melt
858 composition (in wt%) is set at SiO<sub>2</sub>: 50; TiO<sub>2</sub>: 1; Al<sub>2</sub>O<sub>3</sub>: 18.5; FeO<sub>tot</sub>: 11; MnO: 0.2;
859 MgO: 5.5; CaO: 9; Na<sub>2</sub>O: 2.5; K<sub>2</sub>O: 0.7; P<sub>2</sub>O<sub>5</sub>: 0.2; H<sub>2</sub>O: 1.8. Best fit for the data is
860 obtained using a Ti stable isotope fractionation factor between oxides and melts of
861 Δ<sup>49</sup>Ti<sub>oxide-melt</sub> = -0.23 × 10<sup>6</sup>/T<sup>2</sup>.

862
Figure 3: Ti stable isotope compositions (δ⁴⁹Ti relative to OL-Ti) of terrestrial basalts, mantle-derived samples and eclogites as well as lunar mare basalts. Weighted means and 95% c.i. are calculated using Isoplot (Ludwig, 2003). Data for New Britain samples are from Millet and Dauphas (2014). Samples represented with open symbols show signs of fractionation or accumulation of Ti oxides based on TiO₂ concentrations and are removed from the calculation of the ⁴⁹Ti value of the BSE. Once these samples are removed, the δ⁴⁹Ti terrestrial igneous rocks show no resolvable variation between sampling location or geological context or petrographic type, indicating that the BSE has homogeneous Ti stable isotope composition of +0.005±0.005 ‰ relative to OL-Ti (95% c.i.). Low-Ti lunar basalts (red) show δ⁴⁹Ti values within error of the BSE whereas High-Ti mare basalts (green) show values ranging from BSE to enriched in heavy isotopes.

875

Figure 4: Rayleigh distillation model showing the effect of ilmenite crystallisation during the late stages of solidification of the Lunar Magma Ocean. The ilmenite-melt Ti stable isotope fractionation factor is assumed to be equal to the oxide-melt fractionation factor calculated from the Agung volcano differentiation suite. Temperature is set at 1125°C and mineral proportions are taken from Snyder et al. (1992). Kds for Ti for the mineral phases are set as follows: Ilmenite: 20; Clinopyroxene: 0.15; Pigeonite: 0.15 and Plagioclase: 0.

883
<table><caption>884 Table 1</caption><thead><tr><th>Type/Locality</th><th>Sample</th><th>Lab</th><th>SiO<sub>2</sub></th><th>TiO<sub>2</sub></th><th>MgO</th><th>δ<sup>40</sup>Ti</th><th>95% c.i.</th><th>n</th></tr></thead><tbody><tr><td colspan="9"><strong>MORB</strong></td></tr><tr><td>North Atlantic</td><td>A127D8-2</td><td>Dur</td><td>50.08</td><td>0.81</td><td>9.43</td><td>-0.003</td><td>0.020</td><td>1</td></tr><tr><td>North Atlantic</td><td>A127D11-1</td><td>Dur</td><td>51.1</td><td>1.17</td><td>8.55</td><td>0.006</td><td>0.020</td><td>1</td></tr><tr><td>EPR</td><td>R94-2</td><td>Dur</td><td>51.0</td><td>1.33</td><td>7.59</td><td>0.002</td><td>0.028</td><td>1</td></tr><tr><td>EPR</td><td>R82-1</td><td>Dur</td><td>49.3</td><td>1.06</td><td>9.17</td><td>0.002</td><td>0.014</td><td>1</td></tr><tr><td>Pacific</td><td>Sonne12<br>42a</td><td>Dur</td><td>49.93</td><td>1.5</td><td>8.23</td><td>0.005</td><td>0.021</td><td>1</td></tr><tr><td>Indian</td><td>MD57 9-1</td><td>Dur</td><td>51.87</td><td>0.98</td><td>8.88</td><td>-0.010</td><td>0.025</td><td>1</td></tr><tr><td>Indian</td><td>MD57 10-1</td><td>Dur</td><td>50.69</td><td>1.68</td><td>6.84</td><td>0.011</td><td>0.040</td><td>1</td></tr><tr><td colspan="9"><strong>Island Arc Basalts</strong></td></tr><tr><td>Japan</td><td>JB-2</td><td>Dur/OL</td><td>53.2</td><td>1.19</td><td>4.66</td><td>-0.046</td><td>0.009</td><td>3</td></tr><tr><td>Marianas</td><td>ALV1846-9</td><td>Dur</td><td>49.67</td><td>0.57</td><td>6.95</td><td>0.008</td><td>0.017</td><td>1</td></tr><tr><td>Marianas</td><td>1883-5</td><td>Dur</td><td>52.26</td><td>0.77</td><td>5.46</td><td>0.049</td><td>0.013</td><td>1</td></tr><tr><td>Marianas</td><td>1885-6</td><td>Dur</td><td>54.6</td><td>0.79</td><td>5.8</td><td>0.036</td><td>0.018</td><td>1</td></tr><tr><td colspan="9"><strong>Intraplate</strong></td></tr><tr><td>Columbia River</td><td>BCR-2</td><td>Dur/OL</td><td>54.1</td><td>2.26</td><td>3.59</td><td>-0.015</td><td>0.005</td><td>12</td></tr><tr><td>Easter</td><td>17678</td><td>Dur</td><td>47.79</td><td>2.97</td><td>7.79</td><td>-0.004</td><td>0.020</td><td>1</td></tr><tr><td>Hawaii</td><td>BHVO-2</td><td>Dur/OL</td><td>49.9</td><td>2.73</td><td>7.23</td><td>0.020</td><td>0.006</td><td>12</td></tr><tr><td>Cape Verde</td><td>SN01</td><td>Dur</td><td>41.01</td><td>2.88</td><td>15.09</td><td>0.005</td><td>0.022</td><td>1</td></tr><tr><td>Cape Verde</td><td>SN10</td><td>Dur</td><td>39.88</td><td>2.90</td><td>12.50</td><td>0.008</td><td>0.022</td><td>1</td></tr><tr><td>Azores</td><td>S1</td><td>Dur</td><td>45.31</td><td>4.09</td><td>7.76</td><td>0.026</td><td>0.027</td><td>1</td></tr><tr><td>Azores</td><td>S3</td><td>Dur</td><td>46.43</td><td>3.59</td><td>8.34</td><td>0.037</td><td>0.018</td><td>1</td></tr><tr><td>Azores</td><td>SJ52</td><td>Dur</td><td>44.35</td><td>3.78</td><td>8.20</td><td>0.017</td><td>0.018</td><td>1</td></tr><tr><td colspan="9"><strong>Eclogites</strong></td></tr><tr><td>Zermatt-Saas</td><td>SO241ii</td><td>Dur</td><td>47.14</td><td>2.75</td><td>5.20</td><td>-0.015</td><td>0.025</td><td>1</td></tr><tr><td>Zermatt-Saas</td><td>SO241v 2</td><td>Dur</td><td>49.04</td><td>2.72</td><td>5.51</td><td>0.004</td><td>0.031</td><td>1</td></tr><tr><td>Zermatt-Saas</td><td>SO284viix</td><td>Dur</td><td>49.9</td><td>0.7</td><td>8.8</td><td>-0.025</td><td>0.025</td><td>1</td></tr><tr><td colspan="9"><strong>Mantle-derived</strong></td></tr><tr><td>Alpine Serpentinite</td><td>Bch9</td><td>Dur</td><td>40.21</td><td>0.09</td><td>36.04</td><td>0.012</td><td>0.031</td><td>1</td></tr><tr><td>Alpine Serpentinite</td><td>MM15</td><td>Dur</td><td>39.53</td><td>0.06</td><td>36.34</td><td>-0.003</td><td>0.028</td><td>1</td></tr><tr><td>Alpine Serpentinite</td><td>LZ14b</td><td>Dur</td><td>39.03</td><td>0.07</td><td>38.13</td><td>0.030</td><td>0.023</td><td>1</td></tr><tr><td>Beni Bousera Peridotite</td><td>GP13</td><td>Dur</td><td>44.91</td><td>0.14</td><td>39.79</td><td>0.007</td><td>0.022</td><td>1</td></tr><tr><td colspan="9"><strong>Differentiated magmas</strong></td></tr><tr><td>Basaltic andesite</td><td>S19</td><td>Dur</td><td>54.49</td><td>2.11</td><td>6.38</td><td>0.103</td><td>0.022</td><td>2</td></tr><tr><td>Andesite</td><td>AGV1</td><td>Dur</td><td>58.84</td><td>1.05</td><td>1.53</td><td>0.084</td><td>0.029</td><td>2</td></tr><tr><td>Granite</td><td>G2</td><td>Dur</td><td>69.14</td><td>0.48</td><td>0.75</td><td>0.459</td><td>0.027</td><td>2</td></tr><tr><td>Rhyolite</td><td>RGM1</td><td>Dur</td><td>73.4</td><td>0.27</td><td>0.275</td><td>0.548</td><td>0.014</td><td>4</td></tr><tr><td>Agung</td><td>AGU03</td><td>Dur</td><td>58.01</td><td>0.76</td><td>2.92</td><td>0.120</td><td>0.027</td><td>1</td></tr><tr><td>Agung</td><td>AGU16</td><td>Dur</td><td>59.93</td><td>0.7</td><td>2.3</td><td>0.179</td><td>0.029</td><td>1</td></tr><tr><td>Agung</td><td>AGU20</td><td>Dur</td><td>62.69</td><td>0.6</td><td>1.77</td><td>0.248</td><td>0.032</td><td>1</td></tr><tr><td>Agung</td><td>AGU21</td><td>Dur</td><td>54.08</td><td>0.92</td><td>4.13</td><td>0.054</td><td>0.026</td><td>1</td></tr><tr><td>Agung</td><td>AGU23</td><td>Dur</td><td>56.42</td><td>0.76</td><td>2.95</td><td>0.108</td><td>0.029</td><td>1</td></tr><tr><td>Agung</td><td>AGU25</td><td>Dur</td><td>65.16</td><td>0.55</td><td>1.46</td><td>0.259</td><td>0.019</td><td>1</td></tr><tr><td colspan="9"><strong>Moon</strong></td></tr><tr><td>High-Ti</td><td>70017</td><td>OL</td><td>--</td><td>13.3</td><td>--</td><td>0.015</td><td>0.011</td><td>5</td></tr><tr><td>High-Ti</td><td>70215</td><td>OL</td><td>--</td><td>13.0</td><td>--</td><td>0.033</td><td>0.015</td><td>5</td></tr><tr><td>High-Ti</td><td>71055</td><td>OL</td><td>--</td><td>13.4</td><td>--</td><td>0.011</td><td>0.017</td><td>5</td></tr><tr><td>High-Ti</td><td>72155</td><td>OL</td><td>--</td><td>12.2</td><td>--</td><td>0.023</td><td>0.009</td><td>5</td></tr><tr><td>High-Ti</td><td>75075</td><td>OL</td><td>--</td><td>13.4</td><td>--</td><td>0.018</td><td>0.014</td><td>5</td></tr><tr><td>Low-Ti</td><td>15597</td><td>OL</td><td>--</td><td>1.85</td><td>--</td><td>-0.008</td><td>0.010</td><td>5</td></tr><tr><td>Low-Ti</td><td>12009</td><td>OL</td><td>--</td><td>3.30</td><td>--</td><td>-0.008</td><td>0.013</td><td>5</td></tr><tr><td>Low-Ti</td><td>15016</td><td>OL</td><td>--</td><td>2.30</td><td>--</td><td>-0.008</td><td>0.019</td><td>4</td></tr><tr><td>Green Glass</td><td>15426</td><td>OL</td><td>--</td><td>0.50</td><td>--</td><td>0.011</td><td>0.017</td><td>4</td></tr></tbody></table>
<figure><img src="image_2.png" /></figure>
<figure><img src="image_3.png" /></figure>
<figure><img src="image_4.png" /></figure>
<figure><img src="image_5.png" /></figure>