# Extreme reduction: Mantle-derived oxide xenoliths from a hydrogen-rich environment

W.L. Griffin, S.E.M. Gain, F. Camara, L. Bindi, J. Shaw, Olivier Alard, M.

Saunders, J.-X. Huang, V. Toledo, S.Y. O'Reilly

► To cite this version:

W.L. Griffin, S.E.M. Gain, F. Camara, L. Bindi, J. Shaw, et al. Extreme reduction: Mantle-derived oxide xenoliths from a hydrogen-rich environment. Lithos, 2020, 358-359, pp.105404. 10.1016/j.lithos.2020.105404. hal-02910902

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Extreme reduction: Mantle-derived oxide xenoliths from a hydrogen-rich
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Abstract

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Coarse-grained xenoliths of hibonite + grossite + Mg-Al-V spinel from Cretaceous
pyroclastic rocks on Mt Carmel, N. Israel, and from Sierra de Comechingones,
Argentina, include spherules, rods and dense branching structures of native
vanadium and V-Al alloys. Microstructures suggest that vanadium melts became
immiscible with the host Ca-Al-Mg-Si-O melt, and nucleated as droplets on the
surfaces of the oxide phases, principally hibonite. Many extended outward as rods or
branching structures as the host oxide crystal grew. The stability of V⁰ implies
oxygen fugacities ≥9 log units below the Iron-Wustite buffer, suggesting a hydrogen-
dominated atmosphere. This is supported by wt%-levels of hydrogen in gasses
released by crushing, by Raman spectroscopy, and by the presence of VH₂ among
the vanadium balls. The oxide assemblage formed at 1400-1200 °C; the solution of
hydrogen in the metal could lower the melting point of vanadium to these
temperatures. These assemblages probably resulted from reaction between
differentiated mafic melts and mantle-derived CH₄+H₂ fluids near the crust-mantle
boundary, and they record the most reducing magmatic conditions yet documented
on Earth.

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## Introduction

Vanadium is a mildly incompatible, refractory but lithophile transition element; it also
can be mildly siderophile, and may have partitioned into Earth's core at high
pressure and low oxygen fugacity ($fO_2$; Huang et al., 2015). Its valence states
normally range from +5 to +2. At 1500°C the $V^{5+}/V^{4+}$ transition occurs at $fO_2$ 5 log
units above the Iron-Wustite buffer ($\Delta$IW+5), and the $V^{3+}/V^{2+}$ transition at $\Delta$IW-6;
native vanadium ($V^0$) becomes stable at $ca~\Delta$IW-9, an $fO_2$ similar to the early solar
nebula (Grossman et al., 2008). Natural $V^0$ has been reported previously only from
sublimates in volcanic fumaroles (Ostrooumov and Taran, 2016), and in ophiolitic
chromitites (Ifandi et al., 2018).

Here we describe $V^0$ and V-Al alloys trapped as melts in hibonite ($CaAl_{12}O_{19}$)
and grossite ($CaAl_4O_7$) that crystallized in Cretaceous super-reduced melt-fluid
systems in northern Israel (Griffin et al., 2019) and Argentina. Sampling and
analytical methods are described in the Supplementary Data.

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## Background

Most of the material described here comprises xenoliths from the pyroclastic ejecta of Cretaceous (99-85 Ma) intraplate basaltic volcanoes exposed on Mt Carmel, in northern Israel (Fig.1; Griffin et al., 2016, 2018a, 2019a, references therein). At least eight of these volcanoes erupted over an area of ca 150 km², and their xenoliths provide snapshots of similar melt-fluid systems, sampled at different stages of their evolution. Aggregates of skeletal corundum crystals (Carmel
Sapphire™) occur as xenoliths in the tuffs, or re-deposited in nearby alluvial placers, mainly in the Kishon River. Melt pockets trapped within and between the corundum crystals contain mineral assemblages requiring high T (>1450-1200 °C), moderate P (ca 1 GPa) and extremely low fO₂ (see below). Detailed evidence for the natural magmatic (i.e. non-anthropogenic, non-electrostatic) origin of the samples has been presented by Griffin et al. (2019b).

Paragenetic studies (Griffin et al., 2018a, 2019; Xiong et al., 2017) suggest that both the low fO₂ and the crystallization of skeletal corundum reflect the interaction of differentiated mafic magmas with CH₄+H₂ at high fluid/melt ratios. These fluid-melt systems were sampled by eruption of the host basalts, but were not directly genetically related to them.

The earliest paragenesis in the melt pockets of these xenoliths is corundum + tistarite (Ti₂O₃) + carmeltazite (ZrAl₂Ti₄O₁₁; Griffin et al., 2018b) + Mg-Al-Ti spinel in a Ca-Mg-Al-Si-O glass. The crystallization of these Ti³⁺-bearing phases requires fO₂ at least 6 log units below the Iron-Wustite buffer (ΔIW-6; Griffin et al., 2016). The original silicate melts were depleted in Fe and Si by the exsolution of immiscible Fe-Ti-C-Si melts, which also occur abundantly as inclusions in the corundum; these melts crystallized moissanite (SiC), TiB₂, TiN and khamrabaevite (TiC) (Griffin et al., 2016). The metal-silicide melts efficiently scavenged Fe and heavier transition elements, so that no Fe-bearing oxides or silicates are found in the melt pockets. The assemblages described here thus formed from residual melts after extensive crystal fractionation and sequential immiscibility, which enriched the residua in minor elements (REE, Zr, Ti, V, Mn, Sc), Ca and Al. Hibonite first appears late in the crystallization sequence, together with corundum, carmeltazite, Mg-Al spinel, TiN, Fe-Ti silicides, TiC and glass.
Coarser-grained (to cm-sized crystals; Fig. 2) intergrowths of hibonite, grossite and spinel, as grains up to 2.5 cm across, also occur in placers of the Kishon and Yoqneam Rivers that drain Mt Carmel. They suggest that hibonite-bearing melts like those trapped in the corundum aggregates evolved even further beneath some volcanic centres. In these xenoliths, Vº and V-Al alloys occur as spherical, rod-like and branching inclusions, mainly in hibonite.

At Mt Carmel, the late crystallization of dmisteinbergite (a polymorph of anorthite) in the melt pockets reflects the peritectic reaction Liq + Crn → An, equivalent to the incongruent melting of anorthite (An → Liq + Crn). This well-studied reaction is experimentally constrained to pressures (P) >0.9 GPa and temperatures (T) of ca 1450 °C (Goldsmith, 1980). Thermodynamic modelling in the CaO-Al₂O₃-SiO₂ system (Ottonello et al., 2013) suggests that grossite may not be stable at P >1 GPa. The same modelling shows that the melt-pocket glasses, which clearly were crystallizing corundum (Griffin et al., 2016, 2018a), cannot be in equilibrium with corundum at P <1 GPa. These P-T estimates place the proposed site of fluid-melt interaction in the uppermost mantle; the crust-mantle boundary in the area has been geophysically defined at 20-25 km depth (Segev and Rybakov, 2011). Comparisons of mineral parageneses in the melt pockets with experimental studies suggest that T decreased from >1450°C to ca 1200°C as fO₂ dropped from ΔIW-6 to ΔIW≤-9 (Griffin et al., 2018a, 2019a; references therein).

We also describe very similar assemblages have been found in an as-yet undocumented locality in the Sierra de Comechingones, Argentina. These samples, obtained from mineral collectors, probably have been taken from similar basaltic cones known to exist in the area.
106 Petrography

107 The petrography and parageneses of the coarse intergrowths from Mt Carmel
108 are described in detail elsewhere (Griffin et al., 2019a). Resorbed remnants of
109 corundum in hibonite (Mt Carmel) document the peritectic reaction Crn + Liq → Hbn.
110 The latest paragenesis (Fig. 3a) is grossite + spinel + krotite (CaAl₂O₄) + fluorite ±
111 the undescribed oxyfluoride phase Ca₄Al₆F₂O₁₂ ± fluorine-rich perovskite
112 (CaTi(O,F)₃). The Argentinian samples comprise platy cm-sized crystals of V-rich
113 hibonite, in a matrix of grossite, dellagiustaite (Al₂V²⁺O₄ spinel, Cámara et al. 2018),
114 and other phases including Ca₄Al₆F₂O₁₂ (Figs 3b,c).

115 In the Mt Carmel xenoliths, Vº occurs mainly as spheroidal to amoeboid balls
116 up to mm size in hibonite crystals; some balls show exsolution of Al-rich vs Mn-rich
117 vanadium alloys. More commonly, Vº forms droplets on (0001) planes, or rods
118 normal to (0001); these rods commonly are necked down to produce linear trains of
119 droplets normal to (0001) (Fig 3a). The inclusions are zonally distributed in some
120 crystals (Fig. 4). In the Sierra de Comechingones samples, Vº forms elongate
121 inclusions in hibonite, normal to (0001), necked down at irregular intervals; some are
122 partially empty (Fig. 5; Fig. SD1). Spheroidal to amoeboid rods up to 0.1-0.2 mm
123 across are mantled by dellagiustaite or V³⁺/V²⁺-dominant spinel (Fig. 6). Some
124 examples contain tiny precipitates of Cu, and larger inclusions (Figs 6b,c) can be
125 vesicular, attesting to the presence of volatiles.

126 In the Mt Carmel material, some spherical Vº inclusions develop into “dense
127 branching structures” (DBS; Goldenfeld, 1989) that grew roughly normal to (0001) of
128 the host hibonite crystal (Fig. 7; Fig. SD2); each filament terminates in a single
129 euhedral crystal, showing that growth took place in contact with melt. Clusters of
130 fine branches terminate at planes parallel to (0001), suggesting pauses in the growth
of the hibonite crystal. Some branches continued to grow when crystal growth resumed, but new droplets of Vº also nucleated on the new crystal plane and grew independently of the older branches. In the example shown in Fig. 7, this stop-start process occurred at least four times. The 3D-µCT images (Fig. SD2) show that the outer parts of many filaments contain no solid phases, even where they terminate below the present crystal face. These "empty" portions are continuous with the vanadium-filled branches, and make up approximately 3% of the total volume of the branching structures. In the Mt Carmel material, small bulbous protuberances on the sides of some spheroidal Vº inclusions enclosed in grossite consist of VH₂, the first reported natural metal hydride (Bindi et al., 2019).

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## 142 Mineral Chemistry

The smallest Vº inclusions in hibonite from Mt Carmel contain 0.5-2.0% Si and 1.5-2.5% Mn (the oxide matrix contains no Si). Most larger inclusions have less Si, 1-2% Cr, 0.5-3.5% Mn, and up to ca 4 wt% Al; irregular balls included in grossite are the latest to form, and contain up to 15 wt% Al (Table 1). Many of the EMP analyses total <100%, which may reflect H in solid solution (see below). The discovery of vanadium dihydride (VH₂; Bindi et al. 2019) among these balls suggests that the deficit could be due to hydrogen in solid solution.

EBSD imaging confirms that Vº in the larger inclusions is cubic, and a TEM diffraction pattern matches the <110> zone axis of V (James and Straumanis, 1960). Three inclusions of Al-rich vanadium analysed by single-crystal XRD have space group *Im*-3*m*, consistent with the Al-V phase diagram (Kroupa et al., 2017). The unit cell increases from 3.032 Å to 3.068 Å as Al increases from ca 2 at.% to >20 at.% (Fig. 8).
V⁰ in the Sierra de Comechingones hibonite contains almost no Al, but has more Cr than the Mt Carmel examples, and similar Si contents (0.5-2.2 wt%); vanadium-oxide phases analysed by single-crystal XRD yielded either a tetragonal or a monoclinic cell, corresponding to two polymorphs of synthetic V₁₄O₆□₈: a β-phase (space group I4/mmm), and a γ-phase (space group C2/m) with ordered vacancies. The coexistence of these two phases constrains the temperature to <1350 °C (Davydov and Rempel, 2009). These oxides suggest an fO₂ somewhat higher than in the Mt. Carmel rocks, but the presence of V²⁺ in spinel (dellagiustaite) requires fO₂ = ΔIW between -6 and -9. Because dellagiustaite rims are found at the interface between V⁰ and hibonite and grossite, this texture might represent a later subsolidus high-T diffusion corona formed at higher fO₂.

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## 168 Analysis of included fluids

To evaluate the composition of fluids that might be enclosed in the observed voids associated with the vanadium rods and branching structures, four 2-3 mm grains of the hibonite-grossite aggregates and four corundum aggregates from Mt Carmel were analysed for H, C, N, and S contents by ignition in an O₂-He atmosphere at ca 1150 °C, followed by chromatographic analysis (see Supplementary Data). Contents of H, C, S and N are given in Table 2. The hibonite samples contain high concentrations of gas, averaging 1120 ppm C and 2800 ppm H₂ by weight. This is 4-5 times the concentrations found in the corundum aggregates, even though corundum aggregates appear to have more void space than the hibonite crystals. However, in both cases the included fluids are comprised of >95 at.% hydrogen, assuming that oxygen is at most a minor constituent, as consistent with the required low fO₂.

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Raman Spectroscopy

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Raman spectra of hibonite from both localities show sharp peaks in the range 0-1000 cm⁻¹, indicating that the hibonite is well-crystallized, in contrast to most meteoritic hibonite (e.g. Ma, 2010; Fig. 9). Peak positions agree very well with those observed by Hofmeister et al. (2004) for synthetic hibonite (Table 3). We observe a strong polarization effect on particular Raman shifts. Changes in the Raman shift and intensity of the bands can be ascribed to the small amount of V in the hibonite structure. A broad band at ca 1470 cm⁻¹ is present only in N-S orientation of the crystals, along with two shoulders at 1256 and 1588 cm⁻¹. These are difficult to attribute, as plausible groups are v₃ modes of CO₃ groups, but these are usually very weak in carbonates. Another plausible assignment is to -CH₂ and -CH₃ deformations. In the 3000-4150 cm⁻¹ spectral range, we observe (also polarized) Raman shifts corresponding to the stretching of O-H bonds. This implies the presence of H in a nominally anhydrous phase. However, the bonding is not well ordered and the bands are rather broad. Bands are also observed in hibonite from Mt. Carmel but at different Raman shifts. The only mode observed in hibonite from both localities is the Raman shift at ca 3950 cm⁻¹, which corresponds to one OH bond with the hydrogen far from any donor oxygen. In the Mt. Carmel hibonite, the band at 3819 cm⁻¹ is very strong and this points to a higher content of H diffused within the lattice and defects in the hibonite. This is consistent with the chromatography measurements reported above. The intensity of the bands changes across the crystal, suggesting that the H is related to the degassing of V⁰ inclusions, producing voids by volume reduction, as seen in the 3D µ-CT images (Figs SD-1, SD-2), and diffusing the H across the hibonite host.
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207 **Discussion and Conclusions**

208 The structures of the V⁰ inclusions (Figs 2-7) clearly suggest the coexistence
209 of an immiscible metallic melt and a low-αSiO₂ silicate melt that was crystallizing Ca-
210 Al oxides and Mg-Al-V spinel. They are best explained by the nucleation of drops of
211 molten vanadium on the faces of oxide crystals exposed to the silicate melt. Many of
212 these drops apparently attracted more V⁰ as the crystal grew, producing larger and
213 elongated inclusions, continuously in contact with the surrounding mixture of
214 immiscible melts. These inclusions are thus analogous in many ways to the CO₂-
215 filled cylindrical inclusions of sulphide melts in clinopyroxene megacrysts from some
216 alkali basalts (Andersen et al., 1987).

217 The dense branching structures (DBS) shown in Figure 7 probably evolved in
218 a different way. They are not crystallographically controlled, but radiate outwards
219 from initial melt droplets (Fig. 7a; Fig. SD2). The formation of such DBS typically
220 reflects diffusion-controlled interface motion such as the growth of a crystal into a
221 fluid under conditions of mild supercooling (Goldenfeld, 1989). Since each branch of
222 the DBS would need to be growing in contact with a fluid, these structures probably
223 grew from the initial ball on a crystal surface, outward into the melt surrounding the
224 hibonite crystal, only to be encased in hibonite as the crystal surface advanced
225 further out into the melt. The obvious planes in Figures 3 and 7b thus could
226 represent stages where the hibonite overgrew and terminated some parts of the
227 expanding DBS, but presented a stable surface for the nucleation of new Vº droplets.
228 These structures thus reflect an interplay between the growth rates of the Vº DBS
229 and the hibonite crystal, involving diffusion of several chemical species.
Very similar structures, described as “trees” or “cones” have been observed in some high-quality synthetic (HPHT) diamonds grown in metal fluxes (Palynov et al., 1997; Shiryaev et al., 2018); the material in the structures corresponds to Fe alloys or Fe oxides, related to the growth media and/or the sample container. The main differences between these inclusions and the Vº DBS described here is that many of the branches apparently consist of individual inclusions, and none of the high-quality images provided by Shiryaev et al. (2018) appear to grow from individual larger balls of the metallic material, such as those illustrated in Fig. 7. Both Palynov et al. (1997) and Shiryaev et al. (2018) interpret the structures as bunches of defects related to dislocations in the structure of the host diamond, implying that the metallic phases have nucleated on the defects in a semi-continuous manner as the host crystal grew. In contrast the spherical morphology of most Vº inclusions described here argues for separation of an immiscible melt, nucleation on the crystal surface and its trapping by the host oxide,

However, a defect-related control is plausible in the case of the Vº inclusions described here. Tiny balls of Vº appear to have nucleated in abundance on the faces of growing hibonite (and grossite) crystals, but most of those that grow outward form single rods (Fig. 3-5), with tendency to neck down as seen in fluid-inclusion trails in many minerals. The sites of such nucleation may have represented single dislocations or other imperfections. Only a few inclusions have developed into the spectacular DBS shown in Fig. 7; these may have been the sites of more complex dislocation structures, such as subgrain boundaries (Shiryaev et al., 2018).

Inclusions of Vº occur mainly in hibonite, to a lesser extent in grossite and spinel, and rarely in corundum mantled by hibonite; the metallic melt apparently began separating from the parental oxide melt near the temperature of the
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corundum-hibonite peritectic, driven by a continuing decrease in fO₂ (V²⁺→V⁰).

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Vanadium melts at >1900°C, and experimental data in binary systems suggest that the observed levels of Al, Si, Mn and Cr might lower the melting temperature only by <100°C (Table 1a, b). Since the estimated temperature of the Cor +L → Hib peritectic is close to 1300°C in the CaO-Al₂O₃-MgO-SiO₂ system (de Aza et al., 2000; Jung, 2010), any immiscible V-dominated melt would be expected to freeze upon unmixing, which apparently did not happen.

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However, the fO₂ required for the presence of a V⁰ melt (≤ΔIW - 9) implies a hydrogen-rich environment, similar to that of the early solar nebula in which hibonite-bearing Calcium-Aluminum Inclusions formed (Grossman et al., 2008). The included gasses (Table 2) are dominated by hydrogen, consistent with the inferred low fO₂, the evidence of hydrogen defects in the Raman spectra of hibonite, and the presence of VH₂ among the V⁰ balls (Bindi et al., 2019).

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In a hydrogen atmosphere, the melting point of vanadium can be lowered by >1000°C as several wt % H₂ dissolves in the melt (Fukai, 2012); solubility increases dramatically with increasing pressure. H₂ is expelled during solidification of the melt, but the solid V⁰ phase can retain significant levels of hydrogen in solution to temperatures below 1000 °C. The consistently low analytical totals of the V⁰ balls from both localities may reflect this solid solution.

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The outer portions of many DBS (Mt Carmel), and some rods (Argentina), contain no solid phases, even though the empty branches clearly are continuous with the vanadium-filled ones. Furthermore, some of the larger droplets in the DBS now comprise an irregular lump of V⁰ within a subspherical void (Figures 7, SD2). We suggest that these void spaces reflect the exsolution of H₂ from the solidifying V⁰ (and contracting) melts in the inner parts of the structures. This could lead to the

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280 remelting and expulsion of Vº from the outer parts of the branches, closer to the front
281 of hibonite nucleation, and perhaps to production of the observed VH₂ phase.
282 Similarly, the vesicles observed in some Vº grains in the hibonites from Sierra de
283 Comechingones (Fig. 5b) may represent the exsolution of H₂ from a cooling melt.
284 The rimming observed around Vº inclusions in the Argentinian samples could be
285 related to solid-state diffusion of O and Al into Vº, forming both the non-
286 stoichiometric vanadium oxides and the rims of V³⁺/V²⁺-dominant spinels, while H₂
287 from former VH₂ inclusions diffused into the hibonite host.

288 The petrographic and paragenetic similarities between the assemblages from
289 Mt Carmel and Sierra de Comechingones strongly suggest similar conditions of
290 formation. At present more constraints are available on the Mt Carmel xenoliths,
291 which apparently crystallized late in the pre-eruption evolution of fO₂ in magma-fluid
292 systems. This evolution ultimately led to the most reducing conditions yet found on
293 Earth, with free hydrogen as the fluid phase. The observed abundance of carbon in
294 the assemblages suggests an important role for CH₄ in this process. At Mt Carmel,
295 such conditions existed for at least 10 m.y. over an area of ca 150 km², in the
296 uppermost part of a thin mantle lithosphere, and high fluid/rock ratios would be
297 required to maintain such reduced volumes within the mantle. This suggests the
298 derivation of abundant CH₄±H₂ fluids from the deeper mantle, which in turn seems to
299 imply that the sublithospheric mantle beneath this area was metal-saturated (i.e. fO₂
300 = IW), such that any C-O-H fluid issuing from it was dominated by CH₄±H₂. This
301 conclusion has implications for understanding melting and metasomatic processes in
302 the mantle (Griffin et al., 2018).

303 As discussed elsewhere (Xiong et al., 2017; Griffin et al., 2018) there is a
304 significant, if scattered, body of evidence in the published literature that similar
reducing conditions may exist elsewhere on Earth, in connection with several types
of explosive volcanism. Finally, the Al-rich late-forming Vº melts reported here
represent the first *in situ* observation of Al alloys on Earth, and may add further
credence to reports of Alº in many other localities (cf. Dekov et al., 2009).

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310 Acknowledgements

311 We thank Luca Velentini (U. Padova) for the 3D-µCT images of hibonite from
312 Sierra de Comechingones. We also thank Prof. Andrey Shiryaev and an anonymous
313 referee for useful comments and suggestions, and Prof. Shiryaev for providing high-
314 resolution images of structures in diamonds. Instruments used at Macquarie
315 University are funded by DEST Systemic Infrastructure Grants, ARC LIEF,
316 NCRIS/AuScope, industry partners and Macquarie University. LB thanks MIUR-
317 PRIN2017, project "TEOREM deciphering geological processes using Terrestrial and
318 Extraterrestrial ORE Minerals", prot. 2017AK8C32. We acknowledge the scientific
319 and technical assistance of the Microscopy Australia at the Centre for Microscopy,
320 Characterization and Analysis, a facility funded by The University of Western
321 Australia, and State and Commonwealth Governments. This is contribution XXXX
322 from the ARC Centre of Excellence for Core to Crust Fluid Systems and XXXX from
323 the GEMOC Key Centre.

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385 Huang, J-H., Huang, F., Evans, L., Glasauer, S. 2015. Vanadium: Global (bio)geochemistry. Chem.Geol. 417, 68-89.

387

388 Ifandi, E., Zaccarini, F., Tsikouras, B., Grammatikopoulos, T., Garuti, G., Karipi, S., Hatzipanagiotou, K. 2018. First occurrences of Ni-V-Co phosphides in chromitite from the Agios Stefanos mine, Othrys ophiolite, Greece. Ofioliti 43, 131-145.

391

392 James, W.J., Straumanis, M.E. 1960. Lattice parameter and expansion coefficient of vanadium. Jour. Electr. Soc. 107, 69

394

395 Jung, I-H. 2010. Overview of the applications of thermodynamic databases to steelmaking processes. CALPHAD 34, 332-362.

397

398 Kroupa, A., Mazalova, M., Richter, K.W. 2017. The reassessment of the Al-V system and new assessment of the Al-Si-V system. CALPHAD 59, 47-60.

400
401 Liu, C., Eleish, A., Hystad, G., Golden, J.J., Downs, R.T., Morrison, S.M., Hummer, D.R., Ralph, J.P., Fox, P., Hazen, R.M. 2018. Analysis and visualization of vanadium mineral diversity and distribution. Am. Mineral. 103, 1080-1086.
402
403
404
405 Ma, C. 2010. Hibonite-(Fe), (Fe,Mg)Al₁₂O₁₉, a new alteration mineral from the Allende meteorite. Am. Mineral. 95, 188-191.
406
407
408 Ostrooumov, M., Taran, Yu. 2016. Vanadium, V – a new native element mineral from the Colima volcano, State of Colima (Mexico), and implications for fumarole gas composition. Min. Mag. 80, 371-382.
409
410
411
412 Ottonello, G., Attene, M., Ameglio, D., Belmonte, D., Zuccolini, M.V., Natali, M. 2013. Thermodynamic investigation of the CaO–Al₂O₃–SiO₂ system at high P and T through polymer chemistry and convex-hull techniques. Chem. Geol. 346, 81-92.
413
414
415
416 Pal'ynov, Y.N., Khokhryakov, A.F., Borzodov, Y.M., Sokol, A.G., Gusev, V.A., Rylov, G.M., Sobolev, N.V. 1997. Growth conditions and real structure of synthetic diamond crystals. Russian Geology and Geophysics 38, 920-945.
417
418
419
420 Segev, A., Rybakov, M. 2011. History of faulting and magmatism in the Galilee (Israel) and across the Levant continental margin inferred from potential field data. Jour. Geodyn. 51, 264-284.
421
422
423 Shiryaev, A.A., Zolotov, D.A., Suprun, O.M., Ivakhnenko, S.A., Averin, A.A., Buzmakov, A.V.,
424 Lysakovskyi, V.V., Dyachkova, I.G., Asadchikov, V.E. 2018. CrystEngComm 20, 7700-
425 7705.

426

427 Xiong, Q., Griffin, W.L., Huang, J-X., Gain, S.E.M., Toledo, V., Pearson, N.J., O'Reilly, S.Y.
428 2017. Super-reduced mineral assemblages in "ophiolitic" chromitites and
429 peridotites: The view from Mt Carmel. Eur. J. Mineral. 29, 557-570.

430

431

432 **Figure Captions**

433

434 Fig. 1. Location maps showing the location of the Mt Carmel volcanic field. (a) Regional
435 setting; box shows Mt Carmel area; (b) Geological map indicating location of Cretaceous
436 volcanic centers.

437

438 Fig. 2. Combined phase map and backscattered electron (BSE) image of typical hibonite-
439 grossite-spinel assemblage from Mt Carmel, with drop-like inclusions of native vanadium
440 (V⁰).

441

442 Fig. 3. (a) Transmitted-light photo (plane polars) of hibonite-grossite-spinel aggregate from
443 Mt Carmel, showing inclusions of V⁰ in platy hibonite crystals (ca 5 mm across). Note
444 alternating layers with abundant or scarce inclusions, elongation of some inclusions parallel
445 to c-axis, and necking down of inclusions. (b) Hibonite from Sierra de Comechingones, with
446 vanadium rods normal to (0001) in (c) a grossite-Ca₄Al₆F₂O₁₂ matrix with skeletal F-rich
perovskite and dellagiustaite ($V^{3+}/V^{2+}$ spinel), some rimming $V_{14}O_6$. Each figure is 1.7 mm wide.

449

Fig. 4. BSE image of polished section of a hibonite crystal from Mt Carmel (blue prism shows orientation derived from EBSD), with zonal distribution of $V^0$ inclusions.

452

Figure 5. (a) 3D-µCT image of hibonite fragment (length 600 µm) from Sierra de Comechingones, showing necking-down of cylindrical rods of $V^0$. (b) false-coloured image of (a) showing that some $V^0$ rods (red) are partially empty (green) and some are rimmed by dellagiustaite.

457

Fig. 6. (a) Reflected-light image of a polished rock chip from Sierra de Comechingones, showing a $V_{14}O_6$ inclusion rimmed by dellagiustaite. Red arrow points to a very thin grain of $V^{3+}_{1.78}Al_{0.21}V^{2+}_{1}O_4$; the composition of the dellagiustaite rim is $Al_{1.04}V^{3+}_{0.96}V^{2+}_{0.92}Mg_{0.08}O_4$. (b) BSE image of $V^0$ inclusion with abundant fluid inclusions (Table 1b, analysis b). (c) BSE image of $V^0$ inclusion (Table 1b; points 2 and 3 correspond to analyses c and d) with Cu precipitates, rimmed by euhedral Ti-rich dellagiustaite.

464

Figure 7. (a) Transmitted-light and BSE images of a spiral dense branching structure of $V^0$ in hibonite. (b) 3D-µCT image of vanadium DBS in a hibonite crystal from Mt Carmel; each cluster is ca 1 mm high and grows normal to the (0001) face of the hibonite. Red-orange colours show $V^0$, while green colours represent voids.

469

Figure 8. Al content vs unit cell dimension of inclusions of Al-bearing vanadium in grossite.
471

472 Figure 9. Raman spectra of hibonite from Sierra de Comechingones and Mt Carmel.

473

474

475

476

477

478

479

480

481

482

483

<table><caption>Table 1. Mean major-element compositions of vanadium phases, Mt Carmel.</caption><thead><tr><th rowspan="3">wt%</th><th colspan="2">vanadium</th><th colspan="2">V<sub>9</sub>(Mn,Al)</th><th colspan="2">V<sub>4</sub>Al</th></tr><tr><th colspan="2">EMP</th><th colspan="2">EMP</th><th colspan="2">EMP</th></tr><tr><th>n=5</th><th>stdev</th><th>n=5</th><th>stdev</th><th>n=7</th><th>stdev</th></tr></thead><tbody><tr><td>Si</td><td>0.42</td><td>0.22</td><td>0.52</td><td>0.11</td><td>0.38</td><td>0.04</td></tr><tr><td>Zr</td><td></td><td></td><td>0.01</td><td>0.01</td><td></td><td></td></tr><tr><td>Ti</td><td>0.01</td><td>0.02</td><td>0.15</td><td>0.11</td><td>0.20</td><td>0.03</td></tr><tr><td>Al</td><td>0.32</td><td>0.5</td><td>3.94</td><td>1.27</td><td>11.58</td><td>0.25</td></tr><tr><td>Cr</td><td>1.6*</td><td></td><td>0.71*</td><td></td><td>0.44</td><td></td></tr><tr><td>V</td><td>95.2</td><td>3.23</td><td>79.83</td><td>6.86</td><td>83.37</td><td>0.76</td></tr><tr><td>Fe</td><td>0.03</td><td>0.01</td><td>0.01</td><td>0.01</td><td></td><td></td></tr><tr><td>Mn</td><td>1.21</td><td>0.63</td><td>12.20</td><td>6.55</td><td>2.09</td><td>0.07</td></tr><tr><td>Ca</td><td>0.09</td><td>0.06</td><td>0.41</td><td>0.35</td><td>0.21</td><td>0.17</td></tr><tr><td></td><td>98.9</td><td></td><td>97.78</td><td></td><td>98.28</td><td></td></tr></tbody></table>

* EDS data

485

<table><caption>Table 2. Gas analyses of hibonite and corundum aggregates from Mt Carmel.</caption><thead><tr><th></th><th>sample</th><th>N [ppm]</th><th>C [ppm]</th><th>H [ppm]</th><th>S [ppm]</th></tr></thead><tbody><tr><td>hibonite xl</td><td>982</td><td><strong>60</strong></td><td><strong>1041</strong></td><td><strong>1572</strong></td><td><strong>73</strong></td></tr><tr><td>hibonite aggregate</td><td>982</td><td><strong>41</strong></td><td><strong>1165</strong></td><td><strong>2360</strong></td><td><strong>64.3</strong></td></tr></tbody></table>

487
<table><caption>**Table 3.** Observed Raman shifts in hibonite from Sierra di Comechingones (Argentina) and Mt. Carmel (Israel)</caption><tr><td>hibonite</td><td>1124B</td><td><strong>122</strong></td><td><strong>1165</strong></td><td><strong>3293</strong></td><td><strong>113.1</strong></td></tr><tr><td>hibonite</td><td>1124B-dup.</td><td><strong>43</strong></td><td><strong>1196</strong></td><td><strong>3925</strong></td><td><strong>228.7</strong></td></tr><tr><td>corundum semi-transparent</td><td>1124</td><td><strong>156</strong></td><td><strong>329</strong></td><td><strong>554</strong></td><td><strong>166</strong></td></tr><tr><td>corundum-brecciated</td><td>1124</td><td><strong>21</strong></td><td><strong>152</strong></td><td><strong>63</strong></td><td><strong>22.6</strong></td></tr><tr><td>corundum with glass</td><td>1124</td><td><strong>26</strong></td><td><strong>415</strong></td><td><strong>677</strong></td><td><strong>115.7</strong></td></tr><tr><td>corundum</td><td>1124</td><td><strong>226</strong></td><td><strong>347</strong></td><td><strong>667</strong></td><td><strong>222.5</strong></td></tr></table>

<table><thead><tr><th>Synthetic<br>(Hofmeister<br>et al. 2004)</th><th>Sierra de<br>Comechingones</th><th>Sierra de<br>Comechingones</th><th>Mt Carmel</th></tr><tr><th>Orientation<br>v (cm<sup>-1</sup>)</th><th>lath E-W<br>v (cm<sup>-1</sup>)</th><th>lath N-S<br>v (cm<sup>-1</sup>)</th><th>lath N-S<br>v (cm<sup>-1</sup>)</th></tr></thead><tbody><tr><td></td><td></td><td><strong>85</strong></td><td><strong>84.3</strong></td></tr><tr><td></td><td></td><td></td><td>102.1</td></tr><tr><td>193.7</td><td><strong>130.8</strong></td><td><strong>130.8</strong></td><td>127</td></tr><tr><td><strong>209.4p</strong></td><td>205.4</td><td><strong>205.4</strong></td><td><strong>208.9</strong></td></tr><tr><td><strong>250.5p</strong></td><td></td><td></td><td></td></tr><tr><td>274.0</td><td>268.6</td><td>275.8</td><td>273.9</td></tr><tr><td><strong>331.9</strong></td><td><strong>318.9</strong></td><td>323.7</td><td></td></tr><tr><td></td><td>358.3</td><td><strong>345.7</strong></td><td><strong>349.6</strong></td></tr><tr><td><strong>399.4</strong></td><td></td><td></td><td></td></tr><tr><td>449.5</td><td></td><td>438</td><td></td></tr><tr><td>457.8</td><td>457</td><td><strong>453.8</strong></td><td><strong>451.9</strong></td></tr><tr><td>489.1</td><td>485.4</td><td><strong>485.4</strong></td><td><strong>485.4</strong></td></tr><tr><td>529.5</td><td>522</td><td>531</td><td>534.7</td></tr><tr><td>564.5</td><td>576.3</td><td></td><td></td></tr><tr><td></td><td></td><td>593.6</td><td></td></tr><tr><td>624.5</td><td></td><td></td><td></td></tr><tr><td>640.4</td><td></td><td>643.7</td><td>643.3</td></tr><tr><td>683.8</td><td>685.6</td><td>681.6</td><td><strong>678.9</strong></td></tr><tr><td><strong>726.3</strong></td><td></td><td></td><td></td></tr><tr><td>741.0</td><td><strong>744.9</strong></td><td></td><td></td></tr><tr><td>773.9</td><td><strong>767</strong></td><td></td><td>762.6</td></tr><tr><td><strong>795.8</strong></td><td></td><td>804.3</td><td>808</td></tr><tr><td><strong>837.3</strong></td><td></td><td>838.</td><td>837.4</td></tr><tr><td>872.9</td><td><strong>870</strong></td><td><strong>871.7</strong></td><td><strong>875.7</strong></td></tr><tr><td><strong>910.0</strong></td><td>905.6</td><td>909.9</td><td>902.4</td></tr><tr><td></td><td></td><td>3314</td><td>3437</td></tr><tr><td></td><td></td><td></td><td>3497</td></tr><tr><td></td><td>3640</td><td>3623</td><td></td></tr><tr><td></td><td>3682</td><td>3681</td><td></td></tr></tbody></table>
3945 **3819**
3954
4063

*bold = stronger bands*

**Appendix. Sampling and analytical methods**

*Sampling*

The Shefa Yamim exploration project is aimed at the discovery of economically viable placer deposits of gemstones (mainly sapphire, ruby, hibonite, moissanite) and other commodities within the drainage basin of the Kishon River. Over >10 years, the project has sampled the Cretaceous pyroclastic centers on Mt Carmel (Fig. SD1: Rakefet, Har Alon, Bat Shelomo, Muhraka and Beit Oren complexes) and adjacent areas (Ein Ha-Shofet), and minor and major drainages in the Yaz'rael Valley. Samples range in size from several kg to >1000 tonnes. All samples were run through a static grizzly screen to remove pieces larger than 100mm in diameter. Rock samples from the vents were coarsely crushed and then treated in the same way as alluvial samples. The <100mm fraction was washed in a scrubber that breaks up any clods. The <0.5mm component is suspended in the wash water and pumped to settling ponds; fractions larger than 25 mm are used to backfill exploration pits. Samples in the +8mm-16mm and +16mm-24mm size fractions are sorted by hand on a picking belt. The +0.5 -8mm component of the sample is washed and classified into 5 fractions: 0.5-0.7mm, 0.7-1mm, 1-2mm, 2-4mm, 4-6mm, 6-8mm. These fractions are transferred to a pulsating jig plant for gravity separation. Samples in the 2mm-8mm size fractions are visually inspected after the jigging process and sorted in the recovery laboratory. The three smallest size fractions are jigged separately. The heavy concentrate in the center of the jig pan is collected and dried; material on the outer part of the jig pan is discarded. The sorters in the laboratory have demonstrated their efficiency in identifying and recovering a wide range of mineral species, including garnet (pyrope), ilmenite, spinel, chrome-diopside, diamond, moissanite, sapphire, ruby, Carmel Sapphire™, hibonite, rutile and zircon. The material described here is dominated by samples from the Rakefet Magmatic Complex, and
523

the alluvial deposits of the Kishon River mid-reach, because these localities have been most
thoroughly bulk-sampled, providing abundant material.

524

525

The unprocessed heavy mineral concentrates of several samples were hand-picked
under a binocular microscope in the CCFS laboratories. Several rock samples also have been
processed by SelFrag (electrostatic disaggregation) techniques at CCFS, sieved and hand-
picked after magnetic and heavy-liquid separation.

526

527

528

529

530

531

## Analytical Techniques

532

### SEM, EMP

533

Samples were mounted in epoxy blocks, polished and coated with carbon. A Zeiss
EVO MA15 scanning electron microscope (SEM) at the Geochemical Analysis Unit (GAU),
Macquarie University, Sydney, Australia was used to capture Backscattered Electron (BSE)
images and Energy Dispersive X-ray spectrometry (EDS) was used to map the elemental
composition of the samples, and to analyze phases of interest. An accelerating voltage of 15
keV and a beam current of 1 nA was used.

534

535

536

537

538

539

Major and minor elements were determined by electron microprobe (EMP) using a
CAMECA SX100 equipped with five wavelength-dispersive spectrometers at the Macquarie
University GeoAnalytical (MQGA; formerly GAU), Macquarie University, Sydney, Australia.
Analyses were performed using a focused beam (1-2 µm) with an accelerating voltage of 15
keV and a beam current of 20 nA. Standards were a suite of natural and synthetic minerals.
Peak counting varied between 10 – 20 s, and background was counted for 5 – 10 s on either
side of the peak. Oxygen and carbon were measured directly; matrix corrections were
carried out by the ZAF software.

540

541

542

543

544

545

546

547

Major and minor elements also were determined by electron microprobe (EMP)
using a JEOL 8530F electron microprobe equipped with five wavelength dispersive
spectrometers at the Centre for Microscopy Characterisation and Analysis (CMCA), The
University of Western Australia, Perth, Australia. Analyses were preformed using a fully
focussed beam with an accelerating voltage of 15 keV, a beam current of 15 nA. Standards
were a suite of natural and synthetic minerals. On-peak counting times were 30 seconds for
all elements. X-ray intensity data was corrected for Time Dependent Intensity (TDI) loss (or
gain) using a self-calibrated correction for Si kα, Ti kα, Mn kα, Na kα, K kα and mean atomic

548

549

550

551

552

553

554
555

number (MAN) background corrections were used throughout (Donovan & Tingle, 1996;

556

Donovan et al., 2016). Unknown and standard intensities were corrected for deadtime.

557

Oxygen was measured, matrix corrections and ZAF correction were applied throughout.

558

559

### Transmission Electron Microscopy

560

FIB foils from the region of interest were prepared for TEM using a dual-beam FIB

561

system (FEI Nova NanoLab 200). High Angle Annular Dark Field Scanning Transmission

562

Electron Microscopy (HAADF-STEM) imaging and element mapping were carried out using

563

an FEI Titan G2 80-200 TEM/STEM with ChemiSTEM Technology operating at 200 kV at the

564

Centre for Microscopy Characterisation and Analysis (CMCA), The University of Western

565

Australia, Perth, Australia. The element maps were obtained by energy dispersive X-ray

566

spectroscopy using the Super-X detector on the Titan with a probe size ~1 nm and a probe

567

current of ~0.25 nA. Total acquisition times of 20-30 minutes were used to obtain good

568

signal-to-noise ratios. Electron diffraction was carried out using a field limiting aperture that

569

selected an area approximately 400nm in diameter.

570

571

### 3D-µCT

572

X-ray micro-computed tomography (µCT) of Mt Carmel samples was done at the

573

CMCA, Perth. The grain was mounted on a finger of rigid PVC plastic with double-sided tape

574

and a separate tomogram was collected for each grain using a µCT system (Versa 520 XRM,

575

Zeiss) running Scout and Scan software (v11.1.5707.17179, Zeiss). Samples were scanned at

576

50 kV and 77 µA with the source detector positions set to -17.7 and 117 mm, respectively.

577

An LE1 source filter was applied to screen out low energy X-rays. A 0.4X objective lens was

578

used together with 2x camera binning to achieve an isotropic voxel resolution of 9 µm. Each

579

scan comprised 2501 projections through 360o with an exposure of 3s for each projection.

580

Raw data were reconstructed using XMReconstructor software (v11.1.5707.17179, Zeiss)

581

following a standard centre shift and beam hardening correction. The standard 0.7 kernel

582

size recon filter setting was also used. The visualization and analysis of data generated from

583

µCT scans were performed using Avizo (v8.1.1, FEI) software.

584

Micro-CT images of hibonite from Sierra de Comechingones were obtained with a

585

Skyscan1172, Bruker micro CT, with a pixel size of 2.90 µm and a X-ray source at 74 KV and

586

133 µA and an Al-foil of 0.125 mm as a filter. Object to source distance was 46.2 mm and
camera to source distance was 276.56 mm. Raw data were reconstructed using CTVox
volume rendering software from Bluescientific

## Released-Gas Analysis

H, C, N, and S contents were obtained using a vario EL cube elemental analyzer (Elementar, Langenselbold, Germany). Aliquots of finely powdered sample, weighing between 80 and 120 mg, were mixed with WO3 flux (sample:flux ratio between 1/1.2 and 1/1.5), packed in Sn-foils, hand pressed, and then ignited in an O₂-He gas atmosphere furnace at around 1150 °C. The produced gases were trapped and sequentially released into a set of chromatographic columns for the analysis of N (no trapping), then C, H, and S. Each sample was measured for 9 min, and released gases were sequentially analyzed with a thermal conductivity detector. Reference materials JP-1, BE-N, PM-S and BAM-U110, were repeatedly measured in the same analytical conditions and yielded average values (Table XX) well within error of accepted or published values as tabulated in the GeoReM database (http://georem.mpch-mainz.gwdg.de/).

<table><caption>**Table SD-1.** Reference materials for CHNS analyzer.</caption><thead><tr><th></th><th>H</th><th>C</th><th>N</th><th>S</th><th>S</th></tr><tr><th></th><th>TCD</th><th>TCD</th><th>TCD</th><th>TCD</th><th>IR</th></tr></thead><tbody><tr><td colspan="6"><strong>JP-1 Peridotite massif (JGS)</strong></td></tr><tr><td><em>n</em></td><td>4</td><td>12</td><td>14</td><td>14</td><td>14</td></tr><tr><td>Average (µg g<sup>-1</sup>) ± SD</td><td>3195 ± 170</td><td>763 ± 82</td><td>91 ± 23</td><td>27 ± 14</td><td>26 ± 7</td></tr><tr><td>RSD %</td><td>5</td><td>11</td><td>26</td><td>51</td><td>27</td></tr><tr><td colspan="6"><strong>BHVO-2 Basalt, Hawaiian Volcanic Observatory (USGS)</strong></td></tr><tr><td><em>n</em></td><td>5</td><td>16</td><td>12</td><td>12</td><td>8</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>181 ± 45</td><td>316 ± 60</td><td>219 ± 34</td><td>191 ± 41</td><td>179 ± 42</td></tr><tr><td>RSD %</td><td>25</td><td>19</td><td>15</td><td>21</td><td>23</td></tr><tr><td colspan="6"><strong>PM-S Microgabbros (SARM)</strong></td></tr><tr><td><em>n</em></td><td>25</td><td>26</td><td>24</td><td>26</td><td>22</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>689 ± 70</td><td>992 ± 55</td><td>154 ± 38</td><td>1128 ± 51</td><td>1165 ± 119</td></tr><tr><td>RSD %</td><td>10</td><td>6</td><td>15</td><td>5</td><td>10</td></tr><tr><td colspan="6"><strong>BE-N Altered Basalts (SARM)</strong></td></tr><tr><td><em>n</em></td><td>14</td><td>20</td><td>17</td><td>21</td><td>8</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>2771 ± 534</td><td>2301 ± 147</td><td>197 ± 42</td><td>301 ± 37</td><td>298 ± 23</td></tr><tr><td>RSD %</td><td>19</td><td>6</td><td>21</td><td>12</td><td>8</td></tr><tr><td colspan="6"><strong>BAM-U110</strong></td></tr><tr><td><em>n</em></td><td>13</td><td>18</td><td>18</td><td>17</td><td>–</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>12,258 ± 1758</td><td>72,340 ± 2640</td><td>4237 ± 165</td><td>9114 ± 1082</td><td>–</td></tr><tr><td>RSD %</td><td>14</td><td>4</td><td>4</td><td>12</td><td>–</td></tr></tbody></table>

*n* denotes the number of measurements performed; average is the arithmetic means of the *n* values measured and SD is the related standard deviation; RSD % is relative standard deviation expressed in %; TCD: Thermal Conductivity Detector; IR: Infra Red detector.
607

608 Raman Spectroscopy

609 Raman spectra were collected with and Horiba LabRAM HR Evolution microscope with a
610 Sincerity OE detector, using 532-nm laser the Department of Erath Sciences, University of
611 Milan. Reflected and transmitted light optics were used to select analysis locations. Hibonite
612 from Sierra di Comechingones was studied from a polished thin section. Hibonite from Mt.

613 Carmel was studied using a raw fragment. The analysis was performed using a 100x
614 objective and 600 (500 nm) grating. Laser power was attenuated to 50% of the power. To
615 maximize the signal-to-noise ratio, two spectra were collected using an exposure time of 30
616 s, and then merged together at the end of the acquisition. Spectra were taken at two
617 polarizations at right angles to check for changes in intensity of the raman shifts. The system
618 was calibrated using the 520 cm⁻¹ Raman line of metallic silicon. Spectra were processed
619 using LabSpec v.6 (Horiba ®). Peak positions were obtained by fitting Gaussian peaks.

620 Observed peaks are reported in Table 3 and compared with Raman peaks reported by
621 Hofmeister et al. (2004) for synthetic hibonite. Two spectral ranges are reported in Figures
622 3a (50-1000 cm⁻¹) and 3b (3000-4150 cm⁻¹). Raw Raman spectra have been deposited as
623 supplementary material.

624

625 Supplementary Figures

626 Figure SD1. Animated 3D-μCT image of 7b.

627

628 Figure SD2. 3D animation showing successive slices through the structures shown in Figure
629 7b (rotated 90°). Left, grey-scale image; V⁰ appears as bright tones, and voids in black.

630 Right, image from Fig. 7b, to show correlation of voids with green colours. Note that some
631 large vanadium “drops” in the middle parts of the structures actually are comprised of voids
632 (green) surrounding an irregular fragment of V⁰ (yellow). This is interpreted as the result of
633 exsolution of hydrogen from the crystallizing vanadium melt.

634
<figure><img src="image_2.png" /><figcaption>Figure 1</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>Figure 2</figcaption></figure>

<figure><img src="image_4.png" /><figcaption>Figure 3</figcaption></figure>

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<figure><img src="image_5.png" /><figcaption>Figure 4</figcaption></figure>

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<figure><img src="image_6.png" /><figcaption>Figure 5</figcaption></figure>

<figure><img src="image_7.png" /></figure>

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<figure><img src="image_8.png" /><figcaption>Figure 6</figcaption></figure>
<figure><img src="image_9.png" /><figcaption>Figure 7</figcaption></figure>

<figure><img src="image_10.png" /><figcaption>Figure 8</figcaption></figure>

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Arbitrary intensity

<figure><img src="image_11.png" /><figcaption>Figure 9</figcaption></figure>

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657
Appendix Sampling + Methods

**Methods**

*Sampling*

The Shefa Yamim exploration project is aimed at the discovery of economically viable placer deposits of gemstones (mainly sapphire, ruby, hibonite, moissanite) and other commodities within the drainage basin of the Kishon River. Over >10 years, the project has sampled the Cretaceous pyroclastic centers on Mt Carmel (Fig. SD1: Rakefet, Har Alon, Bat Shelomo, Muhraka and Beit Oren complexes) and adjacent areas (Ein Ha-Shofet), and minor and major drainages in the Yaz'rael Valley. Samples range in size from several kg to >1000 tonnes. All samples were run through a static grizzly screen to remove pieces larger than 100mm in diameter. Rock samples from the vents were coarsely crushed and then treated in the same way as alluvial samples. The <100mm fraction was washed in a scrubber that breaks up any clods. The <0.5mm component is suspended in the wash water and pumped to settling ponds; fractions larger than 25 mm are used to backfill exploration pits. Samples in the +8mm-16mm and +16mm-24mm size fractions are sorted by hand on a picking belt. The +0.5 -8mm component of the sample is washed and classified into 5 fractions: 0.5-0.7mm, 0.7-1mm, 1-2mm, 2-4mm, 4-6mm, 6-8mm. These fractions are transferred to a pulsating jig plant for gravity separation. Samples in the 2mm-8mm size fractions are visually inspected after the jigging process and sorted in the recovery laboratory. The three smallest size fractions are jigged separately. The heavy concentrate in the center of the jig pan is collected and dried; material on the outer part of the jig pan is discarded. The sorters in the laboratory have demonstrated their efficiency in identifying and recovering a wide range of mineral species, including garnet (pyrope), ilmenite, spinel, chrome-diopside, diamond, moissanite, sapphire, ruby, Carmel Sapphire™, hibonite, rutile and zircon. The material described here is dominated by samples from the Rakefet Magmatic Complex, and the alluvial deposits of the Kishon River mid-reach, because these localities have been most thoroughly bulk-sampled, providing abundant material.

The unprocessed heavy mineral concentrates of several samples were hand-picked under a binocular microscope in the CCFS laboratories. Several rock samples also have been processed by SelFrag (electrostatic disaggregation) techniques at CCFS, sieved and hand-picked after magnetic and heavy-liquid separation.
691

692

693 Analytical Techniques

694 SEM, EMP

695 Samples were mounted in epoxy blocks, polished and coated with carbon. A Zeiss

696 EVO MA15 scanning electron microscope (SEM) at the Geochemical Analysis Unit (GAU),

697 Macquarie University, Sydney, Australia was used to capture Backscattered Electron (BSE)

698 images and Energy Dispersive X-ray spectrometry (EDS) was used to map the elemental

699 composition of the samples, and to analyze phases of interest. An accelerating voltage of 15

700 keV and a beam current of 1 nA was used.

701 Major and minor elements were determined by electron microprobe (EMP) using a

702 CAMECA SX100 equipped with five wavelength-dispersive spectrometers at the Macquarie

703 University GeoAnalytical (MQGA; formerly GAU), Macquarie University, Sydney, Australia.

704 Analyses were performed using a focused beam (1-2 µm) with an accelerating voltage of 15

705 keV and a beam current of 20 nA. Standards were a suite of natural and synthetic minerals.

706 Peak counting varied between 10 – 20 s, and background was counted for 5 – 10 s on either

707 side of the peak. Oxygen and carbon were measured directly; matrix corrections were

708 carried out by the ZAF software.

709 Major and minor elements also were determined by electron microprobe (EMP)

710 using a JEOL 8530F electron microprobe equipped with five wavelength dispersive

711 spectrometers at the Centre for Microscopy Characterisation and Analysis (CMCA), The

712 University of Western Australia, Perth, Australia. Analyses were preformed using a fully

713 focussed beam with an accelerating voltage of 15 keV, a beam current of 15 nA. Standards

714 were a suite of natural and synthetic minerals. On-peak counting times were 30 seconds for

715 all elements. X-ray intensity data was corrected for Time Dependent Intensity (TDI) loss (or

716 gain) using a self-calibrated correction for Si kα, Ti kα, Mn kα, Na kα, K kα and mean atomic

717 number (MAN) background corrections were used throughout (Donovan & Tingle, 1996;

718 Donovan et al., 2016). Unknown and standard intensities were corrected for deadtime.

719 Oxygen was measured, matrix corrections and ZAF correction were applied throughout.

720

721 Transmission Electron Microscopy
722 FIB foils from the region of interest were prepared for TEM using a dual-beam FIB
723 system (FEI Nova NanoLab 200). High Angle Annular Dark Field Scanning Transmission
724 Electron Microscopy (HAADF-STEM) imaging and element mapping were carried out using
725 an FEI Titan G2 80-200 TEM/STEM with ChemiSTEM Technology operating at 200 kV at the
726 Centre for Microscopy Characterisation and Analysis (CMCA), The University of Western
727 Australia, Perth, Australia. The element maps were obtained by energy dispersive X-ray
728 spectroscopy using the Super-X detector on the Titan with a probe size ~1 nm and a probe
729 current of ~0.25 nA. Total acquisition times of 20-30 minutes were used to obtain good
730 signal-to-noise ratios. Electron diffraction was carried out using a field limiting aperture that
731 selected an area approximately 400nm in diameter.

732

733 3D-µCT (add method for SdC material)

734 For X-ray micro-computed tomography (µCT) at the CMCA, the grain was mounted
735 on a finger of rigid PVC plastic with double-sided tape and a separate tomogram was
736 collected for each grain using a µCT system (Versa 520 XRM, Zeiss) running Scout and Scan
737 software (v11.1.5707.17179, Zeiss). Samples were scanned at 50 kV and 77 µA with the
738 source detector positions set to -17.7 and 117 mm, respectively. An LE1 source filter was
739 applied to screen out low energy X-rays. A 0.4X objective lens was used together with 2x
740 camera binning to achieve an isotropic voxel resolution of 9 µm. Each scan comprised 2501
741 projections through 360o with an exposure of 3s for each projection. Raw data were
742 reconstructed using XMReconstructor software (v11.1.5707.17179, Zeiss) following a
743 standard centre shift and beam hardening correction. The standard 0.7 kernel size recon
744 filter setting was also used. The visualization and analysis of data generated from µCT scans
745 were performed using Avizo (v8.1.1, FEI) software.

746

747 *Released-Gas Analysis*

748 H, C, N, and S contents were obtained using a vario EL cube elemental analyzer
749 (Elementar, Langenselbold, Germany). Aliquots of finely powdered sample, weighing
750 between 80 and 120 mg, were mixed with WO3 flux (sample:flux ratio between 1/1.2 and
751 1/1.5), packed in Sn-foils, hand pressed, and then ignited in an O2-He gas atmosphere
752 furnace at around 1150 °C. The produced gases were trapped and sequentially released into
753 a set of chromatographic columns for the analysis of N (no trapping), then C, H, and S. Each
sample was measured for 9 min, and released gases were sequentially analyzed with a thermal conductivity detector. Reference materials JP-1, BE-N, PM-S and BAM-U110, were repeatedly measured in the same analytical conditions and yielded average values (Table XX) well within error of accepted or published values as tabulated in the GeoReM database (http://georem.mpch-mainz.gwdg.de/).

<table><caption>**Table SD-1.** Reference materials for CHNS analyzer.</caption><thead><tr><th></th><th>H<br>TCD</th><th>C<br>TCD</th><th>N<br>TCD</th><th>S<br>TCD</th><th>S<br>IR</th></tr></thead><tbody><tr><td colspan="6"><strong>JP-1 Peridotite massif (JGS)</strong></td></tr><tr><td><em>n</em></td><td>4</td><td>12</td><td>14</td><td>14</td><td>14</td></tr><tr><td>Average (µg g<sup>-1</sup>) ± SD</td><td>3195 ± 170</td><td>763 ± 82</td><td>91 ± 23</td><td>27 ± 14</td><td>26 ± 7</td></tr><tr><td>RSD %</td><td>5</td><td>11</td><td>26</td><td>51</td><td>27</td></tr><tr><td colspan="6"><strong>BHVO-2 Basalt, Hawaiian Volcanic Observatory (USGS)</strong></td></tr><tr><td><em>n</em></td><td>5</td><td>16</td><td>12</td><td>12</td><td>8</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>181 ± 45</td><td>316 ± 60</td><td>219 ± 34</td><td>191 ± 41</td><td>179 ± 42</td></tr><tr><td>RSD %</td><td>25</td><td>19</td><td>15</td><td>21</td><td>23</td></tr><tr><td colspan="6"><strong>PM-S Microgabbros (SARM)</strong></td></tr><tr><td><em>n</em></td><td>25</td><td>26</td><td>24</td><td>26</td><td>22</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>689 ± 70</td><td>992 ± 55</td><td>154 ± 38</td><td>1128 ± 51</td><td>1165 ± 119</td></tr><tr><td>RSD %</td><td>10</td><td>6</td><td>15</td><td>5</td><td>10</td></tr><tr><td colspan="6"><strong>BE-N Altered Basalts (SARM)</strong></td></tr><tr><td><em>n</em></td><td>14</td><td>20</td><td>17</td><td>21</td><td>8</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>2771 ± 534</td><td>2301 ± 147</td><td>197 ± 42</td><td>301 ± 37</td><td>298 ± 23</td></tr><tr><td>RSD %</td><td>19</td><td>6</td><td>21</td><td>12</td><td>8</td></tr><tr><td colspan="6"><strong>BAM-U110</strong></td></tr><tr><td><em>n</em></td><td>13</td><td>18</td><td>18</td><td>17</td><td>—</td></tr><tr><td>Average (µg g<sup>-1</sup>)</td><td>12,258 ± 1758</td><td>72,340 ± 2640</td><td>4237 ± 165</td><td>9114 ± 1082</td><td>—</td></tr><tr><td>RSD %</td><td>14</td><td>4</td><td>4</td><td>12</td><td>—</td></tr></tbody></table>

n denotes the number of measurements performed; average is the arithmetic means of the n values measured and SD is the related standard deviation; RSD % is relative standard deviation expressed in %; TCD: Thermal Conductivity Detector; IR: Infra Red detector.

761

762

763

764

765

## Raman Spectroscopy

766 Raman spectra were collected with and Horiba LabRAM HR Evolution microscope with a 767 Sincerity OE detector, using 532-nm laser the Department of Erath Sciences, University of 768 Milan. Reflected and transmitted light optics were used to select analysis locations. Hibonite 769 from Sierra di Comechingones was studied from a polished thin section. Hibonite from Mt. 770 Carmel was studied using a raw fragment. The analysis was performed using a 100x 771 objective and 600 (500 nm) grating. Laser power was attenuated to 50% of the power To 772 maximize the signal-to-noise ratio, two spectra were collected using an exposure time of 30 773 s, and then merged together at the end of the acquisition. Spectra were taken at two
774 polarizations at right angles to check for changes in intensity of the raman shifts. The system

775 was calibrated using the 520 cm⁻¹ Raman line of metallic silicon. Spectra were processed

776 using LabSpec v.6 (Horiba ®). Peak positions were obtained by fitting Gaussian peaks.

777 Observed peaks are reported in Table 3 and compared with Raman peaks reported by

778 Hofmeister et al. (2004) for synthetic hibonite. Two spectral ranges are reported in Figures

779 3a (50-1000 cm⁻¹) and 3b (3000-4150 cm⁻¹). Raw Raman spectra have been deposited as

780 supplementary material.

781

782

783

784