Serpentiniztaion of Dumont Dunite

Generation of high-Ni sulfide and alloy phases during serpentinization of the Dumont Dunite, Quebec

MICHELLE SCIORTINO

A Thesis submitted in conformity with the requirements for the degree of MASc Graduate Department of Geology, University of Toronto

Generation of high-Ni sulfide and alloy phases during serpentinization of the Dumont Dunite, Quebec

MASc Graduate Department of Geology, 2014

Michelle Sciortino

Department of Geology

University of Toronto

Abstract The Dumont Sill in the Abitibi, Quebec, is broadly zoned from gabbroic chilled margins through peridotite to a principally dunitic core. The dunite contains several layers enriched in primary sulfide minerals. Serpentinization occurred isochemically with respect to major components apart from the addition of H2O. An extensive database which includes 1420 mineralogical samples and 4400 microprobe analyses has been used to understand alteration and nickel deportment in primary and secondary minerals. In layers containing primary accumulations of magmatic sulfide serpentinization has resulted in the remobilization of nickel from olivine to enrich cumulus sulfides. In layers lacking primary sulfide accumulations, serpentinization was accompanied by the formation of awaruite. Assemblages which are completely serpentinized contain lesser nickel in silicates, higher nickel tenor sulfides, and more modal awaruite compared to those which are weakly and partially serpentinized, which are characterized by higher nickel in silicates and lower Ni tenor sulfides.

Key Words: Ultramafic Nickel, Serpentinization, Nickel Remobilization, Komatiites

Acknowledgements The author would like the thank Royal Nickel Corporation for their support in this project, specifically, Johnna Muinonen, John Korczak and Alger St-Jean for their contributions to the various aspect of this research. Also special thank-you to Stephanie Downing and Kathryn Sheridan (SGS Lakefield) for their dedication and mineralogical expertise in executing Royal Nickel Corporation’s ExplominTM program.

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Contents Abstract ...... ii Acknowledgements ...... iii Table of Abbreviations ...... vi List of Figures ...... vii List of Tables ...... viii Introduction ...... 1 Geological Setting ...... 3 Methods ...... 4 Assays ...... 4 X-Ray Diffraction ...... 5 Electron Microprobe ...... 5 Modal analysis ...... 5 Block Model Calculations ...... 7 Observations: The Dunite Subzone at Dumont ...... 7 Local geology ...... 7 Mineral Assemblages ...... 11 Facies divisions ...... 11 Ni-bearing phases in sulfide-bearing zones ...... 12 Ni-bearing phases in low-grade low-sulfide zones ...... 12 Partial serpentinization ...... 14 Completely serpentinized dunites: Fe-serpentine facies...... 15 Completely serpentinized dunite: Mg-serpentine facies ...... 15 Paragenesis ...... 17 Mineral Compositions ...... 20 Ni-rich phases ...... 20 Spinel ...... 23 Silicate and hydroxide phases ...... 24 Lithogeochemistry ...... 28 CIPW norms...... 28 Distribution of Ni ...... 28 Discussion...... 29

Intensive parameters during sill emplacement ...... 29 Primary redox state ...... 29 Ni Content of Primary Sulfide Blebs ...... 31 Metamorphism ...... 33 Modal variation ...... 33 Metasomatism and volume change ...... 36 Serpentine composition and the redox state of Fe ...... 39 Nickel remobilization ...... 41 Fluid Chemistry and Thermodynamics ...... 44 Phase relations ...... 44 Stability of Ni-Fe Alloys ...... 49 Serpentinization of the Dumont Sill ...... 50 Mass Transfer to Intercumulus Sulfide Blebs and Nickel Remobilization ...... 56 Nickel Distribution and Controls ...... 57 Implications...... 59 Conclusions and Exploration Implications ...... 60 References ...... 62 Appendix A: Methods Additional Information ...... 70 A-1: Electron Micro Probe Measurement Conditions by Lab ...... 70 Appendix A: Figures ...... 73 Appendix B: Reported Explomin Mineralogy and Electron Microprobe Data for Figure 6 ...... 79

Table of Abbreviations

aH2O activity of H2O

aSiO2 activity of SiO2 AW Awaruite BRU Brucite sulfide CNi concentration of Ni in sulfide as a weight fraction FESP Iron Serpentine FS Field Stitch HZ Heazlewoodite LDZ Lower Dunite Zone LOI Loss on ignition MGSP Magnesium Serpentine MT Magnetite N Number of Data or Points NISI Nickel in Silicates OL Olivine PN Pentlandite PS Particle Scan SG Specific Gravity SP Serpentine STDDEV Standard Deviation UDZ Upper Dunite Zone

X(NiS, FeS, FeO, NiO) mole fraction for the respective subscript

List of Figures Figure 1: Location of the Dumont Sill within the Abitibi greenstone belt 4 Figure 2: A)Geological map of the Dumont Sill 8 B) 1st vertical derivative of total magnetic field Figure 3: Level plan view at 250 m elevation 10 Figure 4: Typical Cross-Sectional View of the Dumont Deposit from Line 8200E 11 Figure 5. Textures of Ni-bearing phases 13 Figure 6: Mineral assemblages as reported by ExplominTM field stitch mineralogy 16 Figure 7. Opaque mineral paragenesis during progressive alteration 19 Figure 8: Distribution of modal mineralogy at the 250 m level from the block model 21 A) Modal abundance of Fe-serpentine B) Modal abundance of magnetite C) ExplominTM samples colored by heazlewoodite to pentlandite ratio Figure 9 A) Average pentlandite, heazlewoodite and alloy compositions for various assemblages 22 Figure 10: Comparison of Mg# vs Ni concentration in olivine and serpentine 26 Figure 11: Ni versus Cr (wt%, XRF whole-rock data). 29 Figure 12. Calculated ΔFMQ for measured pairs of olivine and chromite compositions 30 Figure 13: Calculated bulk Ni concentration in primary sulfide blebs 32 Figure 14: Changes in mineralogy with Hydration for upper dunite, low sulfide 35 Figure 15: Whole rock major elements versus serpentine for upper dunite, low sulfide samples 38 Figure 16 A) Total Fe a.p.f.u. per 5 cation formula, versus tetrahedral Si + Al for serpentine 42 B) Fe3+ versus LOI in whole rocks C) Calculated Fe3+ in serpentine vs whole rock

Figure 17 A) Log aSiO2 vs log aH2O after Frost and Beard (2007). 48

B)Temperature vs log aSiO2. Drawn after Frost and Beard (2007). o Figure 18: ΣS2 vs fO2 for Ni assemblages of interest at 300 C and 2 kbar (altered after Frost, 1985). 50 Figure 19. Distribution of sulfide minerals and Fe serpentine facies alteration in the block model (modified after Sciortino et al., 2013) 58

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List of Tables Table 1: Phase assemblages and textures in serpentinized dunite at Dumont 18 Table 2: Summary of pentlandite compositions (wt %). 19 Table 3: Average Ni content (wt%) for pentlandite in sulfide-bearing assemblages 22 Table 4: Summary of awaruite compositions (wt%) 23 Table 5: Summary of heazlewoodite compositions (wt%).. 23 Table 6: Microprobe results for chromite (wt%). 24 Table 7: Microprobe results for magnetite (wt%). 24 Table 8: Comparison of Ni content for phases of interest in sulfide-bearing and sulfide-poor samples. 25 Table 9: Microprobe results for olivine (wt%). 27 Table 10: Microprobe results for serpentine (wt%). 27 Table 11: Comparison of whole rock assays for variably serpentinized populations 38 Table 12: Average modal abundance for minerals of interest as reported by TM Explomin particle scan 54

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Introduction Serpentinites are rocks which are dominantly composed of the serpentine group minerals. In a retrograde metamorphic setting below the temperature and pressure stability of olivine, olivine will serpentinize (hydrate) in the presence of a phase containing H2O to form serpentine

[(Mg, Fe, Ni)3Si2O5(OH)4] and brucite [(Mg, Fe, Ni)(OH)2]. In these low temperature conditions Mg-Fe diffusion in olivine is slow (Evans, 2010) and cannot adjust to its equilibrium composition. Serpentine contains a greater proportion of Mg/Fe than olivine when local equilibrium is achieved between metastable olivine and newly formed serpentine. Lacking a suitable silicate host phase, Fe2+ disproportionates to form Fe0 and Fe3+. The result is the 3+ 2+ production of a ferric iron bearing phase, most commonly magnetite [(Fe2 Fe )O4] or ferric- +2 3+ 3+ 2+ iron bearing serpentine (cronstedtite [(Fe2 Fe )(Fe Si)O5(OH)4]). The oxidation of Fe in 3+ olivine to Fe in secondary minerals produces significant amounts of H2, resulting in some of the most reduced fluids observed on earth (Abrajano and Pasteris, 1989; Charlou et al., 1998; Kelly et al., 2000; Sleep et al., 2004). These conditions generate unusual secondary assemblages of rare metallic nickel-iron alloys and/or desulfurized primary nickel-sulfides (Chamberlain,1965; Frost, 1985; Abrajano and Pasteris, 1989; Kelley et al., 2001; Klien and

Bach; 2009). Continued H2 gas generation is a matter of concern for mine safety, so recognition of gas-producing lithologies is of considerable importance.

In tectonic settings where serpentinization occurs under low temperature and low water to rock ratios, in a retrograde environment, the pore fluid is governed by water-rock interaction therefore the assemblages and reactions observed in the rocks can be used to assess fluid compositions in terms of thermodynamic potentials of fS2 and fO2. Previous thermodynamic studies Frost, 1985; Klein and Bach, 2009) assessed the effects of fO2 and fS2 on nickel-sulfide mineralization, whereas previous experimental studies (Moody 1976a &a1976b, Fillidipis 1985, Janecky and Seyfried 1986, Marcaillou et al 2011) have examined how mineral assemblages vary depending on conditions during serpentinization. Although the progressive nature of serpentinization and the resulting mineral assemblages are well understood, the deportment of Ni due to its re-mobilization from primary olivine into newly formed metamorphic phases has not been well documented in literature. Furthermore, it has remained unclear how much of the disproportionation of Fe2+ to Fe0 and Fe3+ occurs in the early stage of serpentinization by via the

formation of a cronstedtite component in serpentine, and how much occurs later through the production of magnetite by the transformation of Fe2+ in brucite or a brucite-like component of serpentine into magnetite. The fine grained nature of partially serpentinized phases and the associated high variability of such, along with the consistency of thermodynamics calculations with petrological observations suggest that serpentinization occurs on a sub-grain by grain scale as a series of tiny oxidation-reduction fronts (Evans, 2013). The presence of Cl-bearing hydroxides in partially serpentinized rocks has been noted at “reaction fronts” and may be evidence for remobilization of elements originally contained in olivine to secondary phases, namely nickel (Rucklidge 1972a, Rucklidge and Patterson, 1977). A more recent experimental study by Liu et al., 2011 determined Ni solubility and speciation in hydrothermal systems and confirmed Ni mobility as chloride species and that solubility and speciation of the chloride phase was dependent on fO2, providing a mechanism for nickel transport from primary phases to secondary phases. Various serpentinized ultramafic bodies provide evidence that nickel in olivine may be re-distributed into secondary nickel sulfide minerals and Ni-Fe alloys (Keays and Kirkland, 1972; Eckstrand,1975; Donaldson, 1981; Duke, 1985; Keays and Jowitt 2013). The potential for the change in nickel deportment from primary silicates to newly formed Ni-Fe alloy and higher tenor nickel sulfides has profound implications for the exploration and the development of ultramafic hosted, low grade, disseminated, Ni-bearing bodies.

Nickel in ultramafic bodies can occur in several forms; metallic nickel and nickel sulfides which are recoverable by conventional metallurgical methods (flotation and magnetic separation) and “matrix nickel”, nickel considered to be economically “unrecoverable”, hosted within serpentine, brucite or olivine. Assays cannot differentiate recoverable nickel deportment versus unrecoverable nickel in silicates. In traditional high grade nickel sulfide deposits, the nickel potentially hosted in silicates or hydroxides is negligible compared to the measured assay grade. With the increased rarity of traditional high grade nickel-sulfide deposits, the nickel industry has turned to laterites which are capital, energy, and metallurgically intensive. As a result, the exploration and development of lower grade, disseminated nickel sulfides/alloy deposits has become increasing attractive in recent years due to their low cost structure, large deposit sizes and traditional recovery methods.

The Dumont Sill is an ultramafic intrusion which contains a large Ni-bearing dunitic body, called the dunite subzone, within which an 1179 Mt reserve has been delimited (Royal Nickel Corporation, 2013). Several investigations have addressed the genesis and economic potential of the dunite subzone (Eckstrand, 1975; Duke, 1986); however, due to the low grade nature of the dunite subzone along with economic factors and a lack of understanding of the continuity, controls and spatial distribution of nickeliferous metallic minerals, the Dumont project was idle until its recent acquisition in 2007 by Royal Nickel Corporation.

The challenge in developing Dumont type disseminated nickel-sulfide and alloy deposits is in understanding Dumont project value due to the fact that nickel assay values do not indicate recoverable nickel. The serpentinization process in the context of nickel remobilization to recoverable phases and the resulting changing nickel in silicates and needs to be understood in order to reconcile assays with mineralogy and nickel deportment. However, a gap exists in the literature between the understanding of the progressive nature of serpentinization and the resulting re-mobilization of nickel into newly formed metamorphic phases. This work uses an extensive mineralogical and lithogeochemical database to describe the effects of progressive serpentinization of the host rock and the accompanying changes in local thermodynamic quantities including aSiO2, fO2 and fS2. The results are used to define a series of key lithofacies that can be recognized in the field as being associated with H2 gas generation and as being likely or unlikely to host recoverable Ni at economic grades.

Geological Setting The Dumont Sill is a differentiated ultramafic to mafic intrusion located in the Abitibi greenstone belt of Quebec, 25 km west of the town of Amos. It is one of many komatiitic bodies found in the Abitibi greenstone belt (Fig 1.; Duke, 1986; Sproule et al., 2005). It is hosted by the lavas of the Barraute Volcanic Complex (Goodwin and Ridler 1970) which hosts at least 5 ultramafic complexes in the Amos area at approximately the same stratigraphic level (Duke 1986). All komatiites in the Abitibi greenstone belt have experienced varying degrees of chloritization, serpentinization, and carbonatization during sea-floor weathering, hydrothermal alteration, and regional metamorphism (Arndt, 1986; Arndt et al., 1989; Lahaye et al., 1995; Sproule et al., 2005) and have typically been metamorphosed to lower to upper greenschist facies, with localized amphibolite facies metamorphism. The age of the Dumont intrusion is not

explicitly known. The conformable nature of the body, together with the character of its differentiation, suggests that it was emplaced as a virtually horizontal shallow sill slightly later than its enclosing lavas at approximately 2.7 Ga (Duke, 1986; Chown et al., 1992). The host volcanic rocks were then folded and faulted and metamorphosed to greenschist facies assemblages during the amalgamation of the Superior Province ca. 2.65 to 2.7 Ga (Sproule et al., 2005; Ayer et al., 2002; Duke, 1986)

Figure 1: Location of the Dumont Sill within the Abitibi greenstone belt (after Duke, 1986). Northwest and southeastern boundaries are marked by an Archean metasedimentary belt and the Grenville Front respectively.

The sill now dips steeply to the northeast and is overlain by post-glacial clays and till varying in thickness from a maximum of 50 m over the central portion of the intrusion to less than 10 m toward the edges. The contact between the Dumont intrusion and its host rock is not observable in outcrop except at the extreme southeastern margin, where it appears conformable.

Methods

Assays To support the feasibility study that was completed in 2013 (Royal Nickel Corporation 2013), 161,703 m of drilling was completed at spacings of 100 m by 100 m or less (Fig. A-2 Appendix A). Mineralization within the dunite subzone was sampled continuously down drill holes in 1.5 m lengths and was assayed using a nitric acid and aqua regia partial digestion and ICP-AES finish for 35 different elements by ALS Global (formerly Chemex) as described by Royal Nickel Corporation, (2013). These assays are the basis for the Royal Nickel’s resource and

block models and will hereafter be referred to as resource assays. Specific gravity was measured on a 3 g split from each resource assay sample by pycnometer using either methanol or acetone as the displacing fluid. A further set of assays, referred to here as Explomin™ assays, are whole rock complete digestion analyses, completed at SGS Lakefield. Samples are digested by borate fusion and analyzed by wavelength dispersion X-ray fluorescence (Method XRF76C, SGS 2013). Assays are performed on splits from the crushed, homogenized samples used for the particle scan. TM To determine the proportions of whole-rock FeO vs Fe2O3, 20 Explomin pulps were analyzed for FeO content at Actlabs by complete dissolution in hydrofluoric acid and metavanadate titration in an open system.

X-Ray Diffraction Qualitative X-Ray diffraction (XRD) was performed on 36 metallurgical variability samples at SGS Lakefield with a BRUKER AXS D8 Advance diffractometer with 40 KeV accelerating voltage and 35 mA beam current. Scanning was performed in 0.2 s steps of 0.02° from 3 to 70° 2θ on crushed and homogenized material. Metallurgical variability samples are selected from contiguous domains of drill core showing homogeneous nickel deportment and alteration. Metallurgical variability samples comprise 10 to 50 m of either half or full NQ drill core (Royal Nickel Corporation, 2013). Mineral identification and interpretation was done by matching the measured diffraction pattern to that of single phase reference materials. The reference patterns used are those complied by the Joint Committee on Powder Diffraction Standards - International Center for Diffraction Data (JCPDS-ICDD) database.

Electron Microprobe Electron Microprobe (EMPA) measurements were completed at McGill University, Xstrata Processing Services (XPS) and at the University of Toronto as per the Appendix A-1., Appendix A, Figure A-3 displays the distribution of samples selected for electron microprobe work. Data were collected for magnetite, serpentine, pentlandite, heazlewoodite, awaruite, chromite and olivine.

Modal analysis Mineralogical sampling of drill core distributed spatially across the Dumont deposit (Fig. A-1) began in 2009. Since then 1,420 samples of ¼ core, each 1.5 m in length, have been

crushed, homogenized and analyzed for mineral modal abundances using automated quantitative mineralogy scanning (QEMSCAN), particle scanning or “PS”, on grain mounts at SGS Minerals in Lakefield, Ontario. Each such measurement includes examination of a separate intact 2 cm by 3 cm piece of core, selected to be representative of the previously described 1.5m ¼ core, prepared as a mounted puck for field image surface imaging mode (field stitch; FS) and analyzed for mineralogy. Mineralogical samples were selected to provide spatially and compositionally representative data down drill holes along the length of the deposit. Mineralogy was determined via SGS’s Explomin™ program which is a trademarked suite of packages that allows routine, automated and statistically robust documentation of modal mineralogy, alteration and textures. ExplominTM data is generated by QEMSCAN. The QEMSCAN process uses a scanning electron microscope (SEM) EVO 430, equipped with a backscattered electron (BSE) detector and an energy dispersive spectrometer (EDS) with the DiscoverTM software package to generate rastered quantitative chemical analyses combined with measured BSE intensities. The QEMSCAN system was operated with an accelerating voltage of 25 keV, probe current of 11 nA with a dwell time chosen to end after the acquisition of 1,000 X-ray counts at each spot. The point spacing in the rastered image is 2 μm. Back-scattered electron intensity calibration is performed on gold, copper and quartz standards. Three points are measured on each standard to correct for drift before and after each run to ensure that the BSE response remained stable throughout the run. A daily quality assurance and quality control (QAQC) in house monitoring standard is used (RNC-176A) in the middle of the run to ensure the BSE did not drift and return during the process. This ensures that mineral classification by BSE intensity is reproducible. Each mineral of interest is assigned a species identification profile (SIP) using mineral specimens previously identified using X-ray diffraction (XRD) and electron microprobe analyses (EMPA). The SIP was developed using both elemental x-ray counts and BSE intensity of the mineral phases in the known examples. Most mineral definitions are straightforward but in some cases a distinction must be made at an arbitrary cut off value of some parameter. For Dumont reporting purposes the mineral “Fe-serpentine” is distinguished as serpentine containing more than 2 elemental weight percent (wt %) iron (Fe). Where a phase resembling brucite contains more than 3% wt Fe it is classified by QEMSCAN as coalingite. It should be noted however that

this distinction is purely based on composition; the phase recognized as coalingite +3 (Mg10Fe 2(CO3)(OH)24·2H2O) may in fact be brucite containing a significant amount of the

Fe(OH)2 component in solid solution (i.e., Fe-brucite; amakinite). The resulting images have a resolution of 2 µm and 15 µm for the particle scan and field stitch respectively.

Block Model Calculations Images displaying block model data were created in CAE Mining’s Studio 3 software. Blocks are populated with ordinary kriging of ExplominTM mineralogy (PS) and resource assays as shown in Royal Nickel Corporation, 2013. Iso-surfaces were created by an automated command in Studio 3 which selects blocks passing a specified value which results in the formation of a wireframe containing blocks of values equal to, or larger than a the specified value. Section or plan views of iso-surfaces displayed the intersection of the iso-surface (wireframe) with the plane of interest.

Observations: The Dunite Subzone at Dumont

Local geology The Dumont Sill is approximately 7 km long and less than 1 km thick. It is composed of gabbro, peridotite, dunite and a small discontinuous layer of clinopyroxenite ( Fig. 2a). Duke (1986) provided a detailed account of lithological characteristics of the Dumont Sill. The uppermost gabbroic unit is a quartz-rich diabasic-textured gabbronorite containing approximately 40% plagioclase (Duke, 1986). The upper gabbro grades into a lower cumulate gabbronorite which lacks modal quartz. The base of the lower gabbro is marked by an increase in cumulus clinopyroxene to over 90 % in a discontinuous clinopyroxenite layer. Below the clinopyroxenite the appearance of increasing modal proportions of pseudomorphed olivine causes a transition to wehrlite, followed by extreme adcumulate dunite which makes up the greater part of the center of the sill. Despite the paucity of outcrop, the extent of the ultramafic zone has been well defined through a 161,703 m resource drilling program (Royal Nickel, 2013), and through airborne magnetometer surveys (Fig. 2b). The ultramafic zone is subdivided into a lower peridotite subzone, dunite subzone and upper peridotite subzone. The lower and upper peridotite zones are olivine-chromite cumulates with variable amounts of intercumulus clinopyroxene up to

mesocumulate textures containing a maximum of 10% pyroxene. The dunite zone comprises olivine accumulates containing trace amounts of cumulus chromite and intercumulus clinopyroxene. Offsets in the magnetic contours and internal stratigraphy of the ultramafic zone along with oriented drill hole data have provided evidence for a number of faults at a high angle to the

Figure 2: A)Geological map of the Dumont Sill The sill dips 60 to 70  to the northwest. A discontinuous clinopyroxenite unit too small to be shown on this scale is found at the contact between the gabbro and the top of the upper peridotite subunit. D1 to D7 refer to structural domains 1 to 7, separated by faults indicated by orange lines. Figure 2B: 1st vertical derivative of total magnetic field (VTEM survey). The upper and lower peridotites appear as magnetic highs compared to the magnetic low over the central dunite. (Images provided by Royal Nickel Corporation).

long axis of the sill, consistent with the northeast and east-trending regional fault pattern, which divide the sill into structural blocks or domains (Fig. 2). Structural core logging has also identified several faults parallel to the strike of the intrusion based on other offsets in mineralization and alteration, and there are probably other faults which have not yet been recognized with the available observations. The true thickness of the upper mafic and lower ultramafic zone varies between fault blocks throughout the sill. The northwestern end of the body has not been outlined precisely; however, the ultramafic zone is 6,800 m in length with an average true thickness of 450 m. The dip of the ultramafic zone also varies with location in the sill from 60° to 70°. The extent of the mafic zone is much less well defined due to the low density of drill hole data that intersects this zone and its contact with the host rock. An estimate of 200 m is based on the few complete drill hole intersections available and the locations of several outcrop. Lithological variations are evident in the variations of composition in the resource assays as illustrated in Figure 3. The transition from the dunite to the lower peridotite is gradual and are marked by decreases in nickel and Mg/Mg+Fe and increases in chromium and aluminum. The increase in chromium corresponds to an increase in cumulus chromite in the sill toward the base and upper margins (Fig. 3b). The increases in aluminum towards the upper and lower margins of the dunite against the enclosing peridotite result from the increase in intercumulus trapped liquid, which now exists as altered secondary phases such as chlorite. Lithogeochemistry shows a small increase in Mg/Mg+Fe upward from the stratigraphically lower dunite zone toward the upper dunite zone and a decrease into both the upper and lower peridotites (Figure 2a). Also shown in Figure 3 are the outlines of domains containing more than 0.1% S. These domains are in all but one case confined to the dunite subzone. The sulfide-bearing layers form the basis of the Dumont Sill’s 1.179 Gt reserve (Royal Nickel Corporation, 2013). Cumulus sulfide layers within the dunite subzone are parallel with the dip of the intrusion (Fig. 4). They are strongly affiliated with and restricted to the iron poor upper dunite except in the central south east, where two additional layers appear within the lower peridotite (Fig. 3a). A maximum of four olivine-sulfide cumulate layers occur locally within the dunite subzone but they do not extend over the entire strike length of the dunite. Figure 4 is a cross section along the local grid line 8200E illustrating the internal stratigraphy of the sill through a typical example of a well-mineralized portion of the dunite,

which is flanked top and bottom by the upper and lower peridotite zones. Contact-parallel layers or lenses of relatively sulfide-rich dunite are outlined by the presence of Ni grades in excess of 0.3% Ni.

Figure 3: Level plan view at 250 m elevation (surface is at ~315 m) showing the distribution of composition through the Royal Nickel Feasibility block model as constrained by ICP-AES analyses of whole rocks (Royal Nickel Corporation, 2013). Black outlines are the intersection isosurface of sulfur with 0.1% S cut off (ICP-AES). The thicker black outline is the pit outline proposed by the feasibility study. A). Fe content (wt %). B). Cr content (ppm).

Figure 4: Typical Cross-Sectional View of the Dumont Deposit from Line 8200E, looking northwest showing the outline of the proposed open pit from the feasibility study (Royal Nickel, 2013). Drill hole assays were interpolated into the block model as described by Royal Nickel Corporation, (2013). Locations of samples selected for microprobe analysis of olivine (Fig. 10) are projected schematically onto the section as blue (sulfide) and black (low-sulfide).

Mineral Assemblages

Facies divisions Mineralization at Dumont is generally confined to the dunite zone. Serpentinized dunite at Dumont displays a continuous spectrum of textures and mineral assemblages but it is convenient to subdivide the rocks into several distinct categories. First, we distinguish between incipiently, partially, and completely serpentinized dunite. Second, we distinguish between alteration tending to produce serpentinite containing abundant iron-rich serpentine, which we refer to as the Fe-serpentine facies (FESP), and alteration producing only lizardite-(nepouite) serpentine, which we refer to as the Mg-serpentine facies(MGSP). Lastly, we note important differences in the mineralogy of Ni-rich phases between those rocks which initially (prior to alteration) contained significant accumulations of disseminated magmatic sulfide, referred to

here as sulfide-bearing (Fig. 4), and those rocks that contained trace to no sulfides, here called low-sulfide, defined by sulfur assay less than 0.05%.

Ni-bearing phases in sulfide-bearing zones Disseminated nickel sulfide mineralization occurs in higher-grade layers (~ 0.35% to 1.5% Ni) that are parallel to the dip of the dunite subzone (Fig. 4). Sulfide mineralization in these layers is dominated by pentlandite (PN) or heazlewoodite (HZ). Lesser amounts of awaruite (AW) occur as replacements or overgrowths on pentlandite (Fig5b, 6b). Pentlandite and heazlewoodite occur as medium to coarse-grained blebs occupying intercumulus spaces within a primary magmatic texture, sometimes exhibiting secondary overgrowths within magmatic blebs (Fig. 5c,d, Fig. 6b,c). These blebs are commonly intimately associated with magnetite(MT), brucite ± chromite ± awaruite, in intercumulus spaces that are rarely up to 10 mm but generally less than 2 mm in size. In and around some of the larger regional fault zones, millerite can be found alongside magnetite and heazlewoodite (Fig 6H), however this has only been noted in two samples within the Dumont dunite and more frequently occurs in association with talc carbonate alteration (steatization) of the lower peridotite-footwall contact. Abundant very small grains of both Ni sulfide and alloy phases can be found throughout pseudomorphed olivine(OL) grains and along former grain boundaries (Fig.5d) in incipiently to partially serpentinized and Fe serpentine facies zones.

Ni-bearing phases in low-grade low-sulfide zones The silicate or matrix phase assemblages observed for low-grade layers lacking primary magmatic sulfide accumulations are consistent with those for sulfide bearing assemblages, except that they contain little to no sulfides. In incipiently serpentinized layers with low sulfide, Ni sulfide and metallic minerals are rarely observed (Fig. 6i). In partially serpentinized low- sulfide layers awaruite is observed along with trace pentlandite, which occurs at the edges of the pseudomorphed olivine boundaries (Fig. 6 I j-l.) in the Fe-serpentine facies. Ni-Fe alloys (probably awaruite) and Co-Fe alloys are also found as minute grains within the Fe-serpentine matrix in partially to nearly completely serpentinized samples, however these are present with grain sizes on the order of 1 µm and therefore smaller than the resolution of ExplominTM (Fig. 5 d,f,g). In completely serpentinized rocks dominated by Mg-serpentine the sub-µm sized alloy

Figure 5. Textures of Ni-bearing phases. A to C: Reflected light photomicrograph of carbon-coated mounted serpentinite. A) left panel: magnetite filling dilated cleavage planes in a pentlandite grain from sample EXP_394; right panel: equivalent Explomin field stitch image showing Fe-serpentine surrounding olivine grains B) Awaruite replacing pentlandite (EXP_037); assemblage described by reaction (15) C) Heazlewoodite replacing pentlandite in EXP_648; assemblage described by reaction (16). D) BSE image of submicron-scale Ni-Fe alloy inclusions (bright spots) within Fe-Serpentine matrix (light grey) in sample 11-RN-353-EXP003. Dark grey is Mg-serpentine. E) BSE image of Mg-serpentine (dark grey) in 11-RN-353-EXP002. A few small nickel-bearing grains are visible along a pseudomorphed serpentine grain boundary. F) X-Ray map for Mg, with pentlandite grain in the center. Green spots correspond to area where serpentine is Fe-rich. G) Ni X-ray map, analog to image 5f. Bright spots of high Ni counts correspond to areas of Fe-serpentine in 5f. The grain of Mg-serpentine in the lower left of 5f contains little to no Ni inclusions in 5g. PN=Pentlandite, MT=Magnetite, AW=Awaruite, HZ=Heazlewoodite, MGSP=Magnesium Serpentine, FESP=Iron Serpentine, OL=Olivine, SP=Serpentine

inclusions are not often observed and either a low-sulfidation assemblage of trace awaruite, magnetite (± heazlewoodite) or a higher sulfidation assemblage of heazlewoodite and magnetite (± awaruite) is observed (Fig. 6 m,n).

Partial serpentinization No un-serpentinized samples have been observed; the freshest samples contain approximately 50% olivine, but only 4% of the dunite subzone contains more than 20% olivine. Secondary assemblages contain variable amounts of Fe and Mg-serpentine, Fe and Mg-brucite, magnetite, ± pentlandite, ± heazlewoodite, ± awaruite. Petrographic observations and XRD results suggest lizardite is the dominant serpentine mineral along with minor chrysotile and antigorite. Pseudomorphic replacements of primary textures are well-preserved throughout most of the intrusion with minimal evidence for anisotropic strain. In some cases near faults, evidence of anisotropic strain and elongation can be seen within pseudomorphed olivine grains. The least serpentinized dunites are olivine-chromite adcumulates as the weakly serpentinized dunite generally occurs at the stratigraphically lower dunite (lower two sulfide- bearing layers in Figure 3 and 4. The least serpentinized samples contain up to 50% relict olivine centers ± minor Fe-serpentine. Trace magnetite can be found concentrated along the edges of former olivine grains, or as isolated grains throughout the serpentine matrix within former olivine grains. The intercumulus sulfide spaces are mainly pentlandite with some minor magnetite ± chromite. Incipiently and partially serpentinized samples exhibit well preserved primary magmatic textures (Fig 5a,6a, Table 1). Primary intercumulus blebs contain pentlandite with minor primary magnetite ± secondary awaruite. The largest grains of Ni-Fe alloy observed in the Dumont Sill occur in this assemblage as secondary overgrowths up to 5 mm in size on primary pentlandite (Fig 6.B&C) . A feature commonly observed in incipiently to partially serpentinized sulfide bearing samples, is dilatational veins of magnetite occurring along cleavage planes in earlier pentlandite (Fig. 5a) Partially to nearly completely serpentinized dunite contains some relict olivine surrounded by Fe-serpentine. Mg-serpentine is found further away from olivine grains, typically occupying a well-connected network lining fractures and following former olivine grain boundaries surrounding the Fe-serpentine and relict olivine. Fe and Mg brucite (Fe- brucite is identified as coalingite in the QEMSCAN mineral mode) can be found in spotty

patches within pseudomorphed grains. Minor magnetite can be found as tiny grains within or on the boundaries of the pseudomorphed olivine grains. Partially to nearly completely serpentinized dunite contains some relict olivine surrounded by Fe-serpentine. Mg-serpentine is found further away from olivine grains, typically occupying a well-connected network lining fractures and following former olivine grain boundaries surrounding the Fe-serpentine and relict olivine. Fe and Mg brucite (Fe-brucite is identified as coalingite in the QEMSCAN mineral mode) can be found in spotty patches within pseudomorphed grains. Minor magnetite can be found as tiny grains within or on the boundaries of the pseudomorphed olivine grains.

Completely serpentinized dunites: Fe-serpentine facies Completely serpentinized samples, defined by the lack of relict olivine, can be grouped into Fe-serpentine and Mg-serpentine dominated groups. Completely serpentinized samples which are Fe-serpentine dominant contain little magnetite, which occurs either within primary intercumulus sulfide blebs or in trace amounts on the edges of pseudomorphed olivine grains. The major host of iron is Fe-serpentine or Fe-brucite which occurs as replacements of relict olivine cores, or along pseudomorphed grain boundaries as brucite rings. Intercumulus primary sulfide blebs show some secondary overgrowth of magnetite and awaruite replacing pentlandite. (Fig. 6 c, d, Table 1). Extremely small nickel-rich inclusions can be found throughout pseudomorphed olivine grains and along grain boundaries. These inclusions appear to include both sulfide and alloy phases (Fig 5 f, g), tentatively identified as pentlandite and awaruite, respectively.

Completely serpentinized dunite: Mg-serpentine facies Mg-serpentine dominant, completely serpentinized samples have no olivine, and little to no Fe-serpentine. The major host for Fe is in magnetite thus abundant magnetite is found in Mg- dominant completely serpentinized samples within primary intercumulus blebs as secondary overgrowths on pentlandite or forming rings with Mg-brucite around pseudomorphed olivine Intercumulus sulfide blebs are mainly pentlandite + magnetite, or pentlandite + heazlewoodite +magnetite, or heazlewoodite + magnetite. (Fig 6. e to g, Table 1). Heazlewoodite occurs as overgrowths on pentlandite and magnetite.

Figure 6: Mineral assemblages as reported by ExplominTM field stitch mineralogy. Arrows denote progressively greater alteration. Images A to H illustrate sulfide assemblages. Images I to N represent low sulfide assemblages. A) Incipient serpentinization. A significant amount of OL remains, with trace FESP; PN+MT reside in sulfide blebs. B) Partial serpentinization: OL coexists with some FESP, MGSP and trace brucite. There is minor magnetite, and coarse AW partially replaces PN. C to D) Completely serpentinized Fe-serpentine facies. Significant Fe and Mg serpentine and Fe brucite. AW replaces PN and MT. E to G) Completely serpentinized Mg-serpentine facies: the remaining serpentine and brucite is Mg-rich. Sulfide blebs contain abundant magnetite and heazlewoodite replacements after pentlandite. H) furthur alteration, well beyond the consumption of olivine. Heazlewoodite is replaced by relatively sulfur-rich nickel sulfides such as millerite. Images I to N are the sulfide-poor, low grade, analogues of A to H serpentinization. I is the least serpentinized (incipiently serpentinized) with little to no awaruite, whereas L is the most serpentinized (completely serpentinized Mg-serpentine facies) with abundant awaruite. In low sulfide assemblages awaruite appears to be stable well beyond complete serpentinization in the Mg-serpentine facies. M to N: awaruite is converted to heazlewoodite or taenite + heazlewoodite. Bulk Ni content in sulfides wt% for A to H and reported mineralogy for Figure 6 can be found in appendix B.

Paragenesis The mineral paragenesis charts in Fig. 6 and 7 summarize the changes in opaque mineralogy accompanying progressive serpentinization for both sulfide-rich and sulfide-poor layers in the dunite subzone. Early serpentinization proceeds by the generation of Fe-serpentine (Fig. 6) and awaruite but little magnetite. Subsequent replacement of some relict olivine cores by Fe-brucite (identified by ExplominTM as coalingite) also does not generate significant amounts of secondary magnetite. At this stage there is large-scale replacement of pentlandite in primary sulfide blebs by awaruite. At a later stage of serpentinization the early-formed Fe- serpentine is replaced by Mg-serpentine + magnetite. At this stage awaruite appears no longer to be stable in sulfide-rich assemblages; it disappears from the micro-inclusion population within pseudomorphed olivine grains and is replaced by secondary pentlandite in the primary sulfide blebs. In the latest stages of alteration, pentlandite is replaced by heazlewoodite while all remaining Fe-bearing silicates are entirely replaced by Mg-serpentine, Mg-brucite, and magnetite. Awaruite is generally absent in sulfide-rich samples in the Mg-serpentine facies but in the sulfide-poor domains awaruite persists well after the replacement of pentlandite by heazlewoodite. The micro-inclusions of awaruite in sulfide-poor samples disappear in favor of the generation of smaller numbers of large awaruite grains. With continued alteration awaruite eventually disappears even in the sulfide-poor layers, being completely replaced by heazlewoodite. Although some entire samples contain fairly homogeneously distributed mineral assemblages representative of the stages of alteration described above, it is common to see large ranges of development of alteration even at the scale of individual primary olivine grains. Relict olivine cores are commonly surrounded by Fe-serpentine rimmed by Mg-serpentine. Ni mineralogy varies on sub-mm scales with the local changes in serpentine mineralogy so that, for example, finely disseminated awaruite within a Fe-serpentine domain in the heart of a partially serpentinized olivine grain may coexist with a pentlandite-heazlewoodite-magnetite assemblage in a nearby patch of cumulus sulfide, the two domains being separated by a zone of Mg- serpentine.

Table 1: Phase assemblages and textures in serpentinized dunite at Dumont.

Silicate Matrix Other Properties Serpentinizatio Proportion Density Magnetic n State Olivine Serpentine Brucite Magnetite Other Ni Phases of Dunite* (kg*m-3) Susceptibility Present trace, spots trace, spots slightly Dominantly within grains or within grains Incipient 4 >2800 <40 fractured Fe SP along or along (OL>20%) boundaries boundaries small Some to Dominantly some both Fe minor spots grains<5µm of none FESP, and Mg as within grains PN or/and AW Partial 26 2600-2800 <40 (20%

Figure 7. Opaque mineral paragenesis during progressive alteration from fresh olivine through Fe-serpentine to Mg- serpentine facies. Blue and red lines represent sulfide and low sulfide assemblages respectively. Dotted lines represent trace amounts of a phase which is only rarely observed. The orange line represents increasing Ni contents and the change from awaruite to taenite. (MT: Magnetite, PN: Pentlandite, AW: Awaruite, HZ: Heazlewoodite, ML: Millerite).

Table 2: Summary of pentlandite compositions (wt %). Values represent point data for pentlandite across all assemblages

As Ni S Fe Zn Co Cd Cu Mn AVERAGE 0.0 30.4 33.0 32.3 0.0 3.8 0.0 0.0 0.0 MAX 0.2 38.2 34.3 41.4 0.1 42.3 0.0 5.4 0.0 MIN 0.0 18.2 23.9 7.6 0.0 0.3 0.0 0.0 0.0 STDDEV 0.0 3.3 0.7 5.5 0.0 5.0 0.0 0.2 0.0 N-POINTS 862.0 1124.0 1124.0 1124.0 862.0 1124.0 33.0 1124.0 33.0

Because the incipiently, partially and completely serpentinized, Fe-serpentine bearing samples are generally accompanied by less magnetite than is the Mg serpentine facies (Fig 6 a,b), they have lower magnetic susceptibility and are represented by the lower magnetite abundance (<2.39%)in the central south east in Fig. 8b. As a result the zones rich in Fe-bearing serpentines in Fig. 8a (>10%) correspond to the magnetic low in Figure 2b, and the low magnetite abundance in Figure 8b and the pentlandite dominant sulfide mineralogy in Figure 8c. The remainder of the dunite subzone outside of the magnetite low, contains higher magnetite abundances (>2.39%) and little to no Fe-serpentine. Therefore the partially serpentinized and Fe- serpentine facies occur adjacent to each other in the small area in the central southeast with little

to no magnetite and form a small portion of the dunite subzone which is dominantly Mg- serpentine facies.

Mineral Compositions

Ni-rich phases The compositions of pentlandite, heazlewoodite and awaruite are plotted together in Figure 9. 1,124 EMP points for pentlandite were measured within 112 samples. Pentlandite is the most variable in composition of the nickeliferous opaque minerals, displaying a bimodal distribution of Ni tenor from 18 to 38 wt% Ni. The low nickel pentlandite group is associated with pentlandite + magnetite ± awaruite associations in the central southwest (Figure A-4a), whereas the high nickel pentlandite group is associated with heazlewoodite (Table 3, Fig. 9a). Compositions are consistent with equilibration at temperatures below 300 °C (Misra and Fleet, 1973). Pentlandite from sulfide-poor zones contains up to 50 wt% cobalt, reflecting decreases in Ni and Fe but the average is 3.8 wt% Co and furthermore 90% of pentlandite analyses contain < 9.5% Co. High cobalt contents in pentlandite occur at the expense of both Ni and Fe, being especially notable in samples containing heazlewoodite and having low modal abundances of sulfides. 699 alloy compositions were measured by EMP in 112 samples (Fig. 9). Over 80% of alloy grains contain between 71 and 75% Ni (Table 4), corresponding closely to the ideal awaruite stoichiometry of Ni3Fe. Awaruite at Dumont is slightly Ni poor on average and contains some cobalt and copper; it is the most copper-rich of the nickel bearing minerals (chalcopyrite is not observed). Awaruite is sometimes observed co-existing with very fine native copper; some of the higher reported copper values may be a result of fine grained inter-growths. A sub-population of Ni rich awaruite (Ni wt % 75-89) is generally found in the northern portion of the dunite (Fig. 9b) where it is associated with trace heazlewoodite in low sulfide zones. The most nickel rich Ni-Fe alloys (~90 wt% Ni) plot as taenite according to the phase diagram of Cacciamani, 2010. Heazlewoodite compositions, represented by 641 EMP analyses from 91 samples (Table 5; Fig. 9), are the least variable in composition of the nickel-bearing minerals, regardless of assemblage. The nickel content ranges from 61.14 to 74.03 wt%, with an average of 72.2% Ni.

Figure 8: Distribution of modal mineralogy at the 250 m level from the block model (surface elevation is ~315 m). Black dots show the locations of ExplominTM samples which were interpolated by ordinary kriging into the block model (Royal Nickel Corporation, 2013). A) Modal abundance of Fe-serpentine (“FESP”). Any rock containing more than 10 % Fe-serpentine is considered to be part of the Fe-serpentine facies. B) Modal abundance of magnetite. C) ExplominTM samples colored by heazlewoodite to pentlandite ratio (Hz/Pn); values from 0 to 1 are pentlandite dominant, 1 to 5 are mixed; 5 and over are heazlewoodite dominant. (Images provided by Royal Nickel Corporation)

Figure 9 A) Average pentlandite, heazlewoodite and alloy compositions for various assemblages found in layers rich in cumulus sulfide (whole-rock S > 0.1 wt%). Three compositions containing more than 5% cobalt were not plotted. B) All measured average pentlandite compositions for various assemblages found in sulfide-poor zones (whole rock S < 0.1 wt%). Compositions with high Co concentrations plot towards the S apex. Red and blue bars indicate estimated primary magmatic sulfide compositions in the upper and lower dunite zones, respectively (PN: Pentlandite, HZ:Hazelwoodite, AW:Awaruite, TN:Taenite). All assemblages occur with magnetite.

Table 3: Average Ni content (wt%) for pentlandite in sulfide-bearing assemblages. Values reported represent average of 5 to10 points across a given sample.

Assemblage MIN MAX AVG N-Samples PN+AW 25.5 33.5 27.7 29 PN+HZ 30.6 35.7 33.6 20 PN+MT 24.2 34.0 30.1 11 PN+MT+HZ+AW 29.7 32.3 31.0 6 PN=Pentlandite, Aw=Awaruite, Hz=Heazlewoodite, Mt= Magnetite

Table 4: Summary of awaruite compositions (wt%). Values represent point data for awaruite across all assemblages.

As Ni S Fe Zn Co Cd Cu Mn AVERAGE 0.01 72.85 0.05 24.93 0.02 1.00 0.00 0.48 0.00 MAX 0.12 89.86 10.40 38.25 0.20 5.05 0.06 2.64 0.02 MIN 0.00 59.03 0.00 7.12 0.00 0.02 0.00 0.00 0.00 STDDEV 0.02 3.10 0.45 3.11 0.03 0.91 0.01 0.30 0.01 N-POINTS 534 699 699 699 534 699 27 699 27

Table 5: Summary of heazlewoodite compositions (wt%). Values represent point data for heazlewoodite across all assemblages.

As Ni S Fe Zn Co Cd Cu Mn AVERAGE 0.01 72.08 26.76 0.59 0.01 0.04 0.01 0.10 0.00 MAX 0.11 74.31 28.92 6.57 0.15 1.48 0.04 11.23 0.02 MIN 0.00 61.14 21.13 0.03 0.00 0.00 0.00 0.00 0.00 STDDEV 0.02 1.01 0.45 0.72 0.02 0.11 0.01 0.48 0.01 N-POINTS 419 641 641 641 419 641 28 641 28

Figure A-4 shows the spatial distribution of Ni contents in pentlandite, awaruite and heazlewoodite in the Dumont sill. A population of Ni-poor pentlandite (20-28% wt Ni occurs dominantly in the central south east (Fig. A4-A) in the area which corresponds to the magnetic low in Fig. 2B. Populations of high Ni-awaruite (75-89 wt% Ni) and magnetite (0.3 - 0.8 wt% Ni) overlap in the northern portion of the dunite (Fig. A-4 B and C).

Spinel Chromite analyses were collected at the centers of grains (Table 6) and show compositions typical of komatiites (Barnes and Brand, 1999). Magnetite on average contains 0.07% Ni by weight. 77% of 893 points analyzed have values less than 0.06% (Table 7). A subset containing greater than 0.2% Ni in magnetite are from samples from the northeast (Fig. A- 4b) containing high Ni tenor awaruite and taenite and lacking in primary sulfide minerals. Overall, magnetite associated with low sulfide samples exhibits a higher nickel content that those with sulfides (Table 8).

Table 6: Microprobe results for chromite (wt%). Values represent point data for chromite across all assemblages.

SiO2 MgO Al2O3 NiO V2O3 ZnO TiO2 Cr2O3 FeO AVERAGE 0.15 10.98 15.74 0.08 0.11 0.09 0.36 51.01 19.97 MAX 1.13 14.16 23.19 0.16 0.19 0.2 0.82 59.18 27.24 MIN 0.06 9.64 11.38 0.04 0 0 0.1 39.89 13.19 STDDEV 0.17 1.02 3.21 0.03 0.08 0.06 0.16 5.66 2.89 N 47 47 47 47 47 47 47 47 47

Table 7: Microprobe results for magnetite (wt%). Values represent point data for magnetite across all assemblages.

SiO2 MgO FeO Cr2O3 Ni Al2O3 MnO TiO2 CoO ZnO CaO AVERAGE 0.19 0.46 91.39 0.20 0.07 0.00 0.12 0.02 0.05 0.01 0.00

MAX 9.90 61.35 94.02 15.01 1.60 0.32 1.83 0.48 0.20 0.46 0.42

MIN 0.00 0.00 10.74 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

STDDEV 0.53 2.14 3.33 0.86 0.16 0.02 0.13 0.05 0.04 0.02 0.01

N 920 920 920 920 920 893 920 920 891 880 853

Silicate and hydroxide phases Microprobe analysis of olivine grains was limited to the central south east portion of the sill (locations projected onto the section in Fig. 4), with the exception of sample EXP_648 (Fig. 5c) which is found at depth in the central region. 141 points in 14 samples were probed through the stratigraphic thickness of the sill. Points were measured in the center of and across relict olivine grains. The Ni content in the olivine measured across the dunite zone ranged from 0.29 to 0.4 wt% (Figure 10, Table 9) Within individual grains or samples there is little variation in the Ni content of olivine. Nickel concentrations are higher in olivine in the upper dunite compared with the Fe-rich, chromite-rich LDZ (Fig. 10).

Table 8: Comparison of Ni content for phases of interest in sulfide-bearing and sulfide-poor samples.

AW PN HZ SP MT Incipient and Low- partially 73.1(7) N/A N/A 0.18(6) 0.11(7) sulfide serpentinized Completely Low- serpentinized Fe- 74.4(16) N/A N/A 0.2(12) 0.11(16) sulfide serpentine facies Completely Low- serpentinized Mg- 74.2(24) N/A N/A 0.1(6) 0.14(27) sulfide- serpentine facies Incipient and sufide- partially 73.2(4) 29.3(8) 72.3(1) 0.23(6) bearing serpentinized Completely sulfide- serpentinized Fe- 72.8(13) 29.0(18) 71.9(7) 0.14(13) 0.02(18) bearing serpentine facies Completely sulfide- serpentinized Mg- 72.8(8) 33.2(14) 72.5(17) 0.1(1) 0.06(17) bearing serpentine facies HZ/PN>1 Completely sulifide- serpentinized Mg- 72.5(17) 32.7(15) 71.3(13) 0.04(2) 0.02(15) bearing serpentine facies HZ/PN<1 Low sulfide is defined as S<0.05% (ALS aqua regia, partial digestion) Incipiently and partially serpentinized, OL>5%, completely serpentinized Fe-serpentine facies, OL<=5% and FESP>=5%, completely serpentinized Mg- serpentine facies , OL<=5% and FESP<5%. For sulfide-rich groups in incipiently, partially, completely (Fe- serpentine facies) and completely (Mg-serpentine, HZ/PN<1) serpentinized rocks, a sulfur cut-off of S>=0.15% was used. For HZ/PN>1, representing heazlewoodite dominant samples, S> =0.1% was used to define the sulfide- bearing samples. Minerals are those reported by ExplominTM particle scan: HZ=Heazlewoodite, PN=Pentlandite, AW=awaruite, SP=serpentine. “N” denotes the number of samples contributing to the average. Note that for the purposes of comparison, samples with 0.05%>S>0.1% were not included here. Numbers in parenthesis denotes the number of samples

A total of 982 EMP measurements of serpentine were done in 59 ExplominTM samples. Analyses are of mesh-textured lizardite forming pseudomorphs of olivine, but not from dilatational structures like veins. Serpentine shows a range of compositional variability (Table 10), most notably in iron, magnesium and silicon. The molar ratio Mg/Mg+Fe ranges from 73 to 99, showing a wider range than its precursor olivine. Ni contents of serpentine from the same samples are variable, so data shown in Figure 10 are averages of five points for each sample; there is a weak positive correlation between Ni and Fe, where higher Mg# serpentine contained less nickel (Fig. 10).

In partially serpentinized samples, where Ni-Fe alloys occur within the serpentine matrix with grain sizes at or below the resolution of SEM imagery, some of the high Ni values reported for serpentine analyses may result from analyses whose analytical volumes (e.g., several µm in diameter) straddled very small Ni alloy inclusions. Although the chlorine content of serpentine was not measured by microprobe, spikes in the EDS spectra of fine grained Fe-enriched serpentine were frequently observed, which is consistent with the

100

Mg# 92

86 0 0.1 0.2 0.3 0.4 0.5 0.6 Ni wt %

Figure 10: Comparison of Mg# vs Ni concentration in olivine and serpentine. Mg#= molar Mg/ Fe+Mg. Olivine with Mg # less than 90 are from the stratigraphically lower dunite which occurs with abundant cumulus chromite, whereas those with Mg# greater than 90 are from the upper dunite. Serpentines show a weak anticorrelation between Mg# and Ni concentration. Closed blue diamonds and squares represent olivine compositions from sulfide bearing and low-sulfide assemblages respectively. Blue and red “X’s” represent serpentine compositions from sulfide bearing and low-sulfide assemblages respectively. Dashes represent olivine compositions along two liquid lines of descent as described in the text.

Table 9: Microprobe results for olivine (wt%). Values represent point data for olivine across all assemblages.

SiO2 MgO Al2O3 Ni TiO2 Cr2O3 FeO AVERAGE 40.75 48.87 0.04 0.29 0.01 0.06 9.53 MAX 41.77 51.07 0.26 0.40 0.04 0.66 11.97 MIN 37.99 46.36 0.01 0.21 0.00 0.01 7.00 STDDEV 0.50 1.15 0.02 0.06 0.01 0.05 1.56 N 141 141 141 141 141 141 141

Table 10: Microprobe results for serpentine (wt%). Values represent point data for serpentine across all assemblages.

SiO2 Al2O3 FeO Ni MgO MnO Cr2O3 CoO CaO TiO2 AVERAGE 38.53 0.23 2.97 0.14 38.68 0.08 0.12 0.01 0.01 0.02 MAX 43.88 12.17 10.72 1.31 45.55 1.67 2.19 1.61 0.07 0.08 MIN 3.64 0.70 2.19 0.14 2.14 0.06 0.15 0.05 0.03 0.03 STDDEV 4.58 0.00 0.57 0.03 2.75 0.00 0.00 0.00 0.00 0.00 N 982 982 982 982 982 919 982 919 982 64

presence of a chlorine-bearing hydroxide phase perhaps occurring in fine-scale intergrowths with serpentine (Rucklidge and Patterson, 1977). Brucite compositions were not measured in this study by electron microprobe; however unpublished electron microprobe analyses of brucite from Dumont contained iron contents as high as 12.5 wt% (pers. commun. Whiteman and Oliviera, 2009). Iron bearing brucite tends to occur in completely serpentinized, Fe-serpentine facies samples Figure 6 C&D). ExplominTM mineral classification groups brucite and “coalingite” separately based on iron counts, as the mineral coalingite is the oxidation product of iron bearing brucite (Mumpton and Thompson, 1966), however it is likely that much of what ExplominTM has classified as coalingite is still Fe- rich brucite, because samples containing this phase are commonly observed to undergo spontaneous oxidation upon exposure to air, which implies that they were not already oxidized to coalingite prior to sample preparation.

Lithogeochemistry

CIPW norms The serpentinization process has altered primary mineralogy. To understand the primary mineralogy prior to serpentinization, a CIPW normative calculation (Cross et al., 1903) was applied to the ExplominTM database of 1420 whole rock assays. The dunite subzone was an extreme adcumulate dunite containing an average of 97% normative olivine with minimal primary magnetite (av. 0.13%) and chromite (av. 0.8%) ± primary sulfide. The whole rock assay in low grade non-sulfide samples is approximately equivalent in nickel content and Mg# to the olivine composition, as previously suggested by Duke 1986; therefore the whole rock normative forsterite content can be used to infer the olivine forsterite content. The olivine composition varies through the stratigraphic height, increasing in Mg# and nickel content upwards through the dunite subzone (Fig. 10), and decreasing downward into the chromite bearing dunite. This is consistent with the average nickel content in low sulfide zones being higher in the upper Fe-poor upper dunite compared to the lower dunite (Table 8)

Distribution of Ni Nickel deportment in komatiites can be assessed using a plot of Ni versus Cr (Fig. 11). The ExplominTM data set from Dumont plots displays several trends. A cluster of compositions containing about 6,000 ppm Cr, 2,500 ppm Ni results from accumulation of chromite and Ni-rich olivine. Although most samples in this group are classified as un-mineralized, there is some scatter to high Ni in the mineralized subset of chromite-bearing rocks. Samples with low Cr are dunites without cumulus chromite. The large concentration of data around 1,000 ppm Cr and 3,000 ppm Ni represents the un-mineralized dunites whereas the trend to high Ni reflects accumulation of primary magmatic sulfide within the mineralized layers in the dunite zone. Although Barnes and Brand (1999) labeled the Cr-rich chromite cumulate trend to be "barren", we note that mineralized samples from the lowest sulfide-bearing layer in the central southeast of the sill occur within this group, showing in Figure 11as a scatter of points in the upper middle part of the diagram. It is also noteworthy that the pentlandite-dominant mineralized samples attain higher Ni concentrations than the heazlewoodite-dominant group, although the majority of both groups are concentrated in the range below 5,000 ppm Ni.

Non-Mineralized

Low Sulfide

5000 Sulphide-Pn Dominant

Sulphide-Hz Dominant LOG[Cr (ppm) XRF] (ppm) LOG[Cr

500 1000 10000 LOG[Ni (ppm) XRF]

Figure 11: Ni versus Cr (wt%, XRF whole-rock data). Samples are categorized as follows: Non-mineralized; Pn+Hz+Aw < 0.1 %: Low-sulfide; S < 0.05 % (ICP-41) and Pn+Hz+Aw>0.1 %: sulfide-bearing, Pn Dominant; S > 0.05% (ICP-41), Pn + Hz + Aw > 0.1 %, Pn/Hz <= 1. sulfide-Hz Dominant; S > 0.05% (ICP-41), Pn + Hz + Aw > 0.1 %, Pn/Hz > 1. The two populations of high and low chromium represent the lower dunite, which lacks chromite, and upper dunite which contain cumulus chromite. Note that the upper dunite is slightly more nickel rich.

Discussion

Intensive parameters during sill emplacement

Primary redox state

The oxygen fugacity (fO2) at the time of sulfide saturation and sill emplacement can be evaluated by considering the composition of equilibrated chromite-olivine pairs. The compositions of 29 chromite-olivine pairs were measured in six samples, whose locations are projected onto the section in Figure 4. Samples EXP_128 and EXP_041 were selected from the the part of the lower dunite zone where the least amount of serpentinization occurred and chromite occurs both as a cumulus phase and as inclusions within olivine. Here chromite occurs in primary grain-grain contact with fresh olivine. The chromite-olivine pairs in the four other samples were selected from the center of the dunite subzone where chromite is rare. Because chromite is rare in these samples and they have been relatively more

serpentinized, most chromite grains observed were surrounded by serpentine. The compositions of olivine are plotted as Mg# vs Ni (wt%) in Figure 10. The spinel-olivine-orthopyroxene oxygen geobarometer (Wood,1990; Ballhaus et al.,

1991) was applied to the olivine and chromite pairs to establish the fO2 of the komatiitic magma at the time of accumulation of olivine and spinel. The results are dispersed along T-fO2 paths interpreted to reflect local scale closed system re-equilibration along oxygen buffers (Fig. A-5). By comparison of these paths with that of the ideal quartz-fayalite-magnetite (FMQ) oxygen buffer, we are able to determine the approximate redox state of the system when olivine and chromite crystallized together from the komatiitic magma. The data are shown in Figure 12 as deviations in log units from the fO2 of FMQ (i.e., FMQ) versus Mg# of olivine. The points within samples EXP_041 and EXP_128, which contain abundant cumulus chromite, show consistent values of fO2 and temperature of emplacement within 1 log unit of FMQ whereas the samples taken from the chromite-poor part of the dunite zone show fO2 ranging from FMQ to

ΔFMQ -3, suggesting equilibration at very low fO2 higher in the stratigraphic sequence. Sulfide- poor un-mineralized samples also show lower fO2 1.00 EXP_128 0.00 EXP_275

-1.00 EXP_041 EXP_438

Δ FMQ Δ -2.00 271-EXP008

-3.00 EXP_397

-4.00 88 89 90 91 92 93 Mg # Olivine

Figure 12. Calculated ΔFMQ for measured pairs of olivine and chromite compositions. Error bars for calculated ΔFMQ represent the standard error of several pairs measured within each sample. Blue diamonds and red squares are sulfide-bearing and low-sulfide samples respectively.

Ni Content of Primary Sulfide Blebs The equilibration between sulfide liquid and olivine can be expressed in terms of an exchange coefficient KD as defined in equation 1.

1) Published experimental data (Fleet and MacRae, 1988; Gaetani and Grove, 1997; Brenan and Caciagli, 2000; Brenan, 2003) were used to establish a function for KD which is dependent sul on fS2, fO2, and the Ni content of the sulfide melt CNi :

sulfide  fO2    fO2  K D  CNi 34.7log   312 11.0log   70.8 2)   fS 2    fS 2  Sulfur fugacity was estimated by assuming that the sulfide melt was at equilibrium with a komatiite liquid, which permits the use of a simple parameterization of fS2 depending on ΔFMQ, total Fe content of the silicate melt expressed as Fe2O3*, temperature T (K) and pressure P (kbar) (equation 3; Mungall and Brenan, 2014; Mungall and Brenan in prep; Moretti and Baker 2008). sul sul 3 In the present formulation KD has a linear dependence on CNi rather than the (CNi ) dependence proposed by Barnes et al. (2013). The difference results from the existence in the existing experimental database of strongly conflicting measurements of KD at high values of Nisul. At the Ni concentrations observed in olivine at Dumont, use of the Barnes et al. (2013) sul sul cubic dependence of KD on CNi leads to unphysical results like CNi > 100%. We have used the estimated parental magma composition of Duke (1986) to estimate MgO, FeO (and Fe2O3*) in equations 1 and 2. The KD of 0.30 for FeO-MgO exchange between olivine and sulfide melt (Roeder and Emslie, 1970) was used to estimate the liquidus T along an olivine-controlled liquid line of descent for two magmas (Fig. 10), one identical to the parental magma of Duke (1986) to represent the upper dunite subzone, the other very similar but with 13.5 % FeO* and 25 wt% MgO to represent the lower dunite subzone. The results of olivine-spinel oxygen geobarometry (previous section) were used to estimate the proportions of ferric and ferrous iron (Kress and Carmichael, 1991). Comparison of the measured Ni content of olivine with equation 2 thus permitted estimation of the Ni content of the primary blebs of sulfide liquid that accumulated with olivine. The samples originating from the lower part of the dunite zone, where the original

olivine was higher in Fe, show lower Ni tenors for the original sulfide melt of 20 to 30 wt % suggesting the original sulfide bleb was equivalent to pentlandite plus minor chromite and or mss prior to serpentinization. This is consistent with the observed mineralogy in the freshest partially serpentinized lower dunite samples as reported by ExplominTM which contain pentlandite with minor magnetite ± chromite (Fig. 6 A).

The samples from the upper dunite zone have Fo90+ olivine with higher Ni and the estimated primary sulfide liquid composition is correspondingly enriched in Ni to values upwards of 35% (Fig. 10, 11). The sulfide phase assemblage after solidification but prior to serpentinization may therefore have consisted of pentlandite ± godlevskite ± heazlewoodite . Samples from the upper dunite rarely contain abundant olivine as they are more serpentinized than those originating in the lower dunite, so as a result this primary assemblage is represented in our data by only one sample out of 1420 ExplominTM analyses, EXP_648, which does indeed contain Ni rich pentlandite and heazlewoodite (Fig. 5c).

) D

Ni (wt%, Ni K 30

20 20 30 40 50 Ni (wt%, ExplominTM)

Figure 13: Calculated bulk Ni concentration in primary sulfide blebs. Vertical axis: values calculated using the measured olivine composition and the KD for Ni partition between olivine and sulfide melt as described in the text. Horizontal axis: Bulk composition calculated using modal mineralogy from field stitch ExplominTM results. The stratigraphically higher dunites (Mg# > 90; blue dots) have higher Ni contents in olivine and correspondingly higher Ni contents in sulfide as estimated by either method. Error bars for calculated bulk sulfide composition represent the standard error of several data within each sample. Bulk sulfide Ni content based on ExplominTM comprises the average of compositions of pentlandite, awaruite and heazlewoodite weighted to their modal abundances.

The estimated primary Ni content of the sulfide blebs is compared with a value determined by image analysis of the actual samples in Figure 13. The image analysis was done by assuming that primary magnetite was a minor constituent of the sulfide blebs and that any magnetite present has been added during metamorphism (see caption of Fig. 5). An average of the compositions of pentlandite, heazlewoodite, and awaruite weighted for their areal extent in each sample as determined by ExplominTM was therefore determined as an estimate of the Ni content of the sulfide blebs. The average Ni of the intercumulus sulfide blebs calculated both from ExplominTM and from mineral compositions via equations 1to 3 for EXP_648, EXP_037, EXP_275 and 282- EXP_005 supports the existence of primary sulfide blebs containing upwards of 35% Ni (Fig. 13). Such high Ni tenor has been reported for primary sulfides at Bethano, in Australia (Barnes et al, 2011b). Despite the overall agreement between modal abundance of Ni-bearing phases and the estimation of primary sulfide compositions, there are significant differences. Most of the Ni- rich sulfides in the upper dunite zone contain less Ni than anticipated based on the composition of olivine, whereas the lower tenor sulfides in the lower dunite zone appear to plot higher compared to ExplominTM mineralogy. .

Metamorphism

Modal variation During metamorphism olivine is hydrated to form serpentine and as result the hydrous component reflected in whole rock geochemical analysis by LOI increases with decreasing olivine content and increasing total serpentine content (Fig. 14A). In partially serpentinized samples there is early formation of Fe serpentine and Mg serpentine together along with increase in LOI to a maximum LOI of approximately 23% in the Fe-serpentine faces (Fig. 14 b,c). The change from Fe-serpentine facies to Mg-serpentine is accompanied by a drop in LOI and a slight drop in the average serpentine content from an average of 92% to 90% serpentine, represented by the black dots in Fig. 14A.

Partially serpentinized and Fe-serpentine facies samples contain the least amount of magnetite, however magnetite formation is not directly related to hydration (Fig. 14d) which is consistent with the observations in various other serpentinite bodies (Toft et al., 1990; Oufi et al., 2002; Bach at el., 2006, Frost and Beard, 2007; Evans et al., 2009). However there does appear

to be a slight increase in magnetite in partially serpentinized samples before the lowest magnetite values observed in the Fe-serpentine facies group. The wholesale addition of magnetite to the rock occurs after complete serpentinization or hydration at the expense of Fe-serpentine and Fe-

100 100

90 A 90 B

80 70

Fe-Serpentine MGSP 70 SP SP (TOTAL) 60 Facies Mg-Serpentine 50 Facies 60 Partially Serpentinized 40 50 0 20 40 60 10 15 20 25 OL LOI %

50 10

40 C 8 D

30 6

MT FE FE SP 20 4

10 2

0 0 10 15 LOI 20 25 10 15 LOI 20 25

50 2 E 40 1.5 F

1 CG FE FE SP 20 0.5 10

0 0 0 2 4 6 8 10 0 2 4 6 8 10 MT MT

100 14 12 90 G H

80 8

70 Fe2O3 6 SP SP (TOTAL) 4 60 2 50 0 0 2 4 6 8 10 0 2 4 6 8 10 MT MT

8 8 7 I 7 J 6 6

5 5

4 4

BRU BRU 3 3 2 2 1 1 0 0 10 15 20 25 80 85 90 95 100 LOI SP TOTAL

Figure 14: Changes in mineralogy with Hydration for upper dunite, low sulfide. Mineralogy defined by ExplominTM. Group definitions: Blue diamonds = incipiently and partially serpentinized; OL>5%: Green squares = completely serpentinized Fe-serpentine facies; OL<=5% and FESP>=5%: Red squares = completely serpentinized Mg- serpentine facies; OL<=5% and FESP<5%. LOI - Loss on ignition reported in ExplominTM whole rock geochemistry; MGSP - magnesium-serpentine; FESP - iron-serpentine; OL - olivine; SP(TOTAL) - the sum of MGSP and FESP; MT - Magnetite; CG - Coalingite (represents the probable presence of Fe-brucite); BRU - magnesium brucite. Fe2O3 in Figure 14 H is reported on an anhydrous basis. Black circle, diamond and square represent the average for the three groups: partially serpentinized, Fe-serpentine facies and Mg-serpentine facies respectively brucite referred to as coalingite (Fig. 14 e,f). Even after the modal abundance of Fe-serpentine is diminished, there appears to be continued magnetite production within the Mg-serpentine to near half of the rock mass containing serpentine content (Fig. 14G up to 4-5% magnetite, after which increases magnetite abundances can be attributed to increased whole rock iron content (Fig. 14 H).

Brucite does not occur until LOI has attained values of at least 10%, corresponding to near half of the rock mass containing serpentine. Brucite increases with increasing hydration in partially serpentinized dunite, with the trend continuing into the Fe-serpentine facies. In the completely serpentinized Mg serpentine facies brucite modes vary between 0.5 and 7 %, centered around the LOI average of 17.8% (Fig. 14I). If we consider the Mg-serpentine facies group separately (Fig. 14J), brucite abundance shows a decreasing trend with increasing amounts of serpentine.

Metasomatism and volume change

Serpentinization of dunite occurs by the addition of H2O to a rock essentially composed of (Mg, Fe) olivine to produce secondary serpentine minerals, brucite and magnetite. This process is known to produce some of the most reducing and silica under-saturated fluids on earth (Frost and Beard, 2007). In some instances serpentinization is isochemical, apart from the addition of H2O, accompanied by a volume expansion and corresponding decrease in density, whereas in others serpentinization is accompanied by for major loss of major elements and results in no volume expansion (Evans, 2008). At Dumont, there are systematic stratiform variations in primary composition and mineral assemblage, such as the presence of layers in the dunite with or without cumulus chromite or primary sulfide, which have survived serpentinization without apparent change, suggesting that serpentinization proceeded by simple addition of H2O to the rock. Furthermore, a comparison of major element geochemistry to total serpentine abundance as a proxy for serpentinization in incipiently, partially and completely serpentinized samples (excluding sulfide-bearing zones to simplify the presentation) shows no trend with increasing serpentinization ( Fig. 15 a to d). The average whole rock concentrations of NiO, MgO, SiO2 and Fe2O3 remain nearly constant in each group (Table 11). The increased scatter in composition must therefore reflect mobility of these elements on the length-scale of the 1.5 m samples but does not indicate that there was any net loss or gain of material from the system apart from the increase in H2O. The sole exception to this pattern is a notable depletion in CaO, from as high as 0.4 % in the freshest samples, to averages of about 0.1% in the serpentinites (Fig. 15 E).. The loss of CaO during serpentinization

is widely recognized, and in systems containing more CaO than the Dumont dunite it can lead to the formation of rodingites in mafic rocks outside the serpentinite (e.g., O’Hanley 1996 and Shervais et al. 2005). Rodingite is not observed at Dumont, perhaps because of the paucity of Ca-bearing minerals in the protolith as well as the absence of mafic dikes predating serpentinization..

0.6 12 0.5

0.4 A 10 B

0.3

NI% 8 0.2 Fe2O3% 6 0.1

0 4 50 60 70 80 90 100 50 60 70 80 90 100 SP (total) Sp (total)

52 44

43 50 C

42 D

41 SiO2% MgO% 46 40

44 39

42 38 50 60 70 80 90 100 50 60 70 80 90 100 SP(total) SP (total)

1 3

0.8 E 2.8 F

0.6

2.6 SG CaO% 0.4

2.4 0.2

0 2.2 50 60 70 80 90 100 50 60 70 80 90 100 SP (total) SP (total) Figure 15: Whole rock major elements versus serpentine for upper dunite, low sulfide samples. Elements measured by whole rock XRF borate fusion (SGS) normalized to anhydrous basis. Serpentinization was isochemical with respect to major elements plotted with the exception of CaO. Black circle, diamond and square represent the average for the three groups: partially serpentinized, Fe-serpentine facies and Mg-serpentine facies respectively.

Table 11: Comparison of whole rock assays for variably serpentinized populations.

Upper Dunite Zone(UDZ) FeO % MgO% SiO2% (ppm) S % N

Incipiently and Partially Serpentinized 5.72 42.59 35.05 2869 0.019 43

Completely Serpentinized (Fe- Serpentine Facies) 5.66 41.10 34.16 2690 0.018 130

Completely Serpentinized (Mg- serpentine Facies) 5.44 41.64 34.89 2701 0.022 279

Lower Dunite Zone(LDZ) FeO % MgO% SiO2% (ppm) S % N

Incipiently and partially serpentinized 7.79 40.8 34.8 2334 0.020 51

Completely serpentinized (Fe- Serpentine facies) 7.68 39.5 33.5 2433 0.021 74

Completely serpentinized (Mg- serpentine facies) 7.76 39.9 34.1 2390 0.021 143

Assays used are ExplominTM assays reported by SGS (XRF Borate Fusion complete digestion). The upper dunite is distinguished from lower dunite using a cut off of 5.3% Fe. Definitions for sub-dividing incipiently, partially and the two completely serpentinized facies are based on olivine and Fe-serpentine content defined in: Incipiently and partially serpentinized, OL>5%, Completely serpentinized Fe-serpentine facies, OL<=5% and FESP>=5%, Completely serpentinized Mg-serpentine sacies , OL<=5% and FESP<5%.

The density of dunite at Dumont ranges from 2700 to 3000 kgm-3 in partially serpentinized rocks (containing max 50% relict olivine) and 2400 to 2700 kgm-3 in completely serpentinized rocks. Completely serpentinized rocks (as defined by the absence of olivine), exhibit two groups, one with moderate densities (the Mg-serpentine facies) whose density drop is retarded by magnetite growth (2600 to 2800 kgm-3) and the other (the Fe-serpentine facies) with the lowest densities observed at Dumont, whose host for iron is serpentine (2300 to 2600 kgm-3) (Fig. 15F). An isocheimal serpentinization process implies little to no mass loss in the system and must therefore imply a volume expansion that can account for the decrease in density. Evidence for significant volume expansion may be seen in feature commonly preserved in incipiently to partially serpentinized dunites, exemplified by sample EXP_394 (Fig. 5a). This is among the freshest samples observed at Dumont, containing 42.3% olivine within a single 1.5 m ExplominTM sample. Pentlandite within the illustrated intercumulus bleb displays dilatational fracturing along cleavage planes, filled by magnetite. The notion that magnetite within the sulfide blebs occurs as dilatant fracture fillings is consistent with the results of the Ni mass balance exercise conducted above, where it was inferred that the original volume of the blebs was currently accounted for by the volume presently occupied by sulfide and metallic phases, excluding magnetite (Fig. 13).

An upper limit estimate on the volume expansion in completely serpentinized dunites can be assessed by comparing the lowest values for specific gravity, measured in Dumont dunite completely serpentinized Fe-serpentine facies, to the specific gravity prior to serpentinization. In un-mineralized samples from the upper dunite zone, where there is little to no intercumulus chromite or magnetite, the primary whole rock specific gravity will have been close to that of forsterite (3300 kgm-3; Fig. 15f). This would imply a volume expansion of approximately 40% during the transition from fresh dunite to completely serpentinized Dumont dunites of the Fe- serpentine facies.

Serpentine composition and the redox state of Fe The Dumont dunite was initially and extreme adcumulate containing up to 100% olivine with an average of Fo 91, yet completely serpentinized Mg-serpentine facies serpentines have Mg # up to 99. The increase in Mg# in secondary products requires the precipitation of additional Fe-rich phases. In incipiently serpentinized dunite, the main reservoir for Fe is as Fe2+

in olivine. Completely serpentinized Fe-serpentine facies contains Fe-serpentine and brucite as well as moderate amounts of magnetite. Brucite hosts Fe as Fe2+, whereas serpentine can host Fe in both 2+ and 3+ states (O’Hanley and Dyar, 1993; Evans, 2008). There are several possible serpentine end member components that can serve as hosts for Fe. Divalent Fe is accommodated by the lizardite-greenalite solid solution. As suggested by O’Hanley and Dyar, (1993) and Evans (2008), Fe3+ may enter the serpentine structure as a cronstedtite substitution (Fig. 16a ), wherein Fe3+ in the tetrahedral site is charge-balanced by Fe3+ in the M-site. Alternatively, a pair of Fe3+ ions may appear in the M sites coupled with an M-site vacancy (Wicks and Plant, 1979) in a component referred to by Evans (2008) as ferrikaolinite. Serpentine compositions determined by EMPA have been recast in an idealized 5- cation stoichiometry to address the behaviour of Fe and H2O during the formation of serpentine and plotted in Figure 16A. The calculation is done assuming that Al is evenly distributed between the tetrahedral and octahedral sites according to a Tschermak's type substitution. Any remaining deficiency in the tetrahedral site is assumed to be filled with Fe3+, which is charge- balanced by Fe3+ in the M site. Iron remaining after the assignment of Fe3+ is assumed to be divalent and is assigned to the M site. If there is insufficient Fe to fill it then a vacancy is assumed to exist in the tetrahedral site. A perfect cronstedtite substitution would form a slope of

-0.5 in Figure 16A. A trend line in Fig. 16b contaminated by brucite (XFe=0.3) would have a slope of -1.4, and contamination with less rich Fe-brucite would have a slope of -2.8. Fig. 16A shows a steeper negative slope than a perfect cronstedtite substitution (-0.7 to -0.9). Fe serpentine at Dumont has a large deficit in the tetrahedral site even if all Fe is assumed to be Fe3+ and split evenly between tetrahedral and octahedral sites. As Fe in serpentine increases, the anhydrous totals decrease. One way to accommodate this observation is to consider the possibility that a hydrogrossular-like substitution may replace tetrahedral Si4+ ions + with four H ions when serpentine forms at extremely low SiO2 activity. This would imply a large excess of H2O, above what could fit in a lizardite-greenalite stoichiometry and even exceeds what could be in the Fe-kaolinite type of M-site vacancy substitution (Wicks and Plant, 1979). A tetrahedral deficit in Fe-serpentine would go some way toward accounting for the low T-site occupancy inferred from EMPA and whole rock LOI data. Alternatively, what is referred to as Fe-serpentine in this study may be poorly crystalline “proto-serpentine” or fine Fe-rich amorphous phase associated with olivine dissolution (Rucklidge and Patterson, 1977; King et al,

2010; Daval. et al, 2011; Gouze et al. 2011) but since it does not appear isotropic in transmitted light it will henceforth be continued to be referred to as Fe-serpentine. To further test for the presence of Fe3+ in serpentine, metavanadate FeO titration was performed on 20 pulps from partially serpentinized and completely serpentinized Fe-serpentine facies containing less than 1.8% magnetite and up to 38% Fe-serpentine as reported by ExplominTM. Figure 16B shows that 30 to 73% of the total Fe contained in the whole rock occurs as Fe3+ and that completely serpentinized Fe-serpentine facies samples contain more Fe3+ and have higher LOI contain less magnetite compared to partially serpentinized samples.(Fig. 14d). BSE images of partially serpentinized and completely serpentinized Fe-serpentine facies samples do indicate the presence of some fine, microscale magnetite inclusions however there is far more Fe3+ in whole rock than could be accounted for by magnetite even if fine magnetite results in under-reporting reporting by ExplominTM. In the completely serpentinized Mg serpentine facies the tetrahedral site in serpentine is apparently filled with Si and Al/2 without requiring Fe to fill in, suggesting that the Fe in serpentine is dominantly Fe2+, and that magnetite is the main reservoir of all iron, accounting for both Fe2+ and Fe3+.

Nickel remobilization As shown in Figures 10 and 11, the original nickel content of olivine varied between 0.2 and 0.4 wt% Ni. The secondary products serpentine and magnetite, which make up > 90 % of the serpentinized rock mass, contain significantly less Ni than olivine (Table 8 &9, Figure 10), yet the serpentinization process has been isochemical with respect to nickel (as discussed above). A comparison of weakly, near completely and completely serpentinized rocks (Table 8) displays differences in the nickel content for serpentine and magnetite through various serpentinization states for sulfide bearing and low-sulfide samples. Both sulfide and low-sulfide samples show a decreasing trend for nickel content in serpentine with serpentinization (Table 8), consistent with the trend toward lower Ni in serpentine with higher Mg# in Fig. 10. The decrease in nickel in serpentine differs for sulfide and non-sulfide samples, the latter showing almost double in completely serpentinized dunites(Table 8). Magnetite follows similar trends whereby sulfide-bearing samples contain little to no nickel in magnetite, while low -sulfide samples contain more, however the nickel increase in magnetite through serpentinization is only observed in non-sulfide samples.

Misra and Fleet (1973) and Harris and Nickel (1972) found that the Ni content of pentlandite should increase with that of the bulk composition of the sulfide, however the pentlandite compositional distribution does not follow parallel primary features (enrichment with bulk sulfide enrichment) of the sill suggesting the compositional variations are not solely the result of low temperature equilibration. Pentlandite samples co-existing with heazlewoodite in

2.4

2.2 A y = -0.89x + 2.0714 R² = 0.785

2.0

1.8

IV Si+Al IV 1.6 y = -0.7047x + 1.9995 1.4 R² = 0.7683 1.2

1.0 0 0.05 0.1 0.15 0.2 0.25 0.3 IV Fe apfu

80% 100% 70% 90%

B C

60% 80%

50% 70%

40% SP % 60% 3+

% oftotal %FE 30% 50%

Fe 3+ 3+

Fe 20% 40% 10% 30% 0% 20% 5 10 15 20 20% 45% 70% 95%

3+ LOI % Fe % whole rock

Figure 16 A) Total Fe a.p.f.u. per 5 cation formula, versus tetrahedral Si + Al for serpentine. Stoichiometry was calculated as described in the text. Red squares represent data collected at McGill and XPS. Blue diamonds represent data collected at University of Toronto. B) Fe3+ versus LOI in whole rocks. Whole rock data were determined by metavanadate titration and XRF(SGS). Note that the indicated LOI underestimates the amount of 2+ H2O bound up in the rock because oxidation of Fe during ignition adds O2 to the rock, lessening the apparent loss on ignition. C) Calculated Fe3+ in serpentine vs whole rock (open symbols) was based on reported ExplominTM mineralogy for olivine, chromite, magnetite and serpentine (assumed olivine composition of Fo91, assumed composition of chromite and serpentine based on microprobe averages). Blue and green correspond to partially serpentinized and Fe-serpentine facies samples respectively.

completely serpentinized Mg-serpentine facies samples have higher nickel contents than those co-existing with magnetite ± awaruite assemblage which are restricted to the completely serpentinized fe-serpentine facies assemblages (Table 3, Table 8). The lower values of Ni in serpentine and magnetite in sulfide samples combined with the increase in pentlandite nickel content with serpentinization may suggest that metallic nickel phases in intercumulus blebs have taken up some of the nickel release by olivine during serpentinization. The increase in Ni in serpentine and magnetite, along with the presence of nickel enriched awaruite in low sulfide, completely serpentinized Mg-serpentine facies samples may reflect the lack of an intercumulus sulfide phase to preferentially remove uptake. The original adcumulate upper dunite contained olivine and sulfides. The average sulfide content in the 29 freshest ExplominTM samples (containing more than 15% olivine and more than 0.1% sulfur) is 0.74 wt % (dominantly pentlandite). If we consider the average Ni content of the primary olivine and sulfide blebs at 0.3 wt% Ni (Table 10) and 40wt% Ni (Fig. 13) respectively, then the serpentinization of this adcumulate dunite containing a maximum of 99% olivine could release 0.28 g of Ni per 100g of rock. Based on the electron microprobe data for completely serpentinized Mg-serpentine facies sulfide-bearing samples, the secondary resulting minerals contain 0.1% and 0.05% Ni in serpentine and magnetite (Table 8). If we assume the nickel released by olivine were to be taken up by the sulfide bleb in the serpentinized rock containing the average of 89.5%, 4.4% and 2.2% serpentine, magnetite and brucite respectively (Table 8) then the Ni content of the blebs should be approximately 75% wt Ni, which is very close to the observed heazlewoodite Ni content in completely serpentinized samples. This suggests that the blebs on average can account for the Ni realised from olivine. Similar calculation for low-sulfide samples, where the original rock contained 100% olivine with 0.25% Ni would produce a metallic nickel assemblage of 0.75% wt Ni in alloy, approximately equal to awaruite at the observed average modal abundances of awaruite and heazlewoodite in completely serpentinized, Mg-serpentine facies non-sulfide samples. This suggests that the modal abundances of observed secondary minerals can account for the Ni released by primary olivine during serpentinization and that the upgrading of the blebs during the progressive serpentinization of Dumont dunite likely occurs upon the decomposition of iron serpentine as observed with the increase in Ni in pentlandite and the decrease in Ni in serpentine toward the Mg-serpentine facies.

Chlorine has been noted in partially serpentinized ultramafic associated with either fine grained Fe-serpentine or a Fe-hydroxide phase (the literature is inconclusive) from various ultramafic bodies (Janecky and Seyfried, 1986; Dahlberg and Saini-Eidukat, 1991; Springer, 1998; Anselmi et al., 2000). Older studies (Edel'shtein 1965; Faust 1965; Naldrett 1966a, Bogolepov,1969; Ashley, 1975) have suggested that Fe and Ni are immobile during serpentinization however more recent evidence (Keays and Kirkland, 1972; Keays and Jowitt, 2013), including the present, study suggest that this may not be the case. Rucklidge (1972a) described mobility of Ni associated with chlorine in partially serpentinized dunites, and described a larger scale of mobility for Ni as a chloride. The introduction of seawater via serpentinizing fluid could provide a source of Cl to the alteration process which lowers the activation barrier energy required to dissolve Ni into solution (Miura et al, 1981). More recent experimental studies by Liu et al., 2011 determined Ni solubility and speciation in hydrothermal systems and confirmed Ni mobility as chloride species and that solubility and speciation of the chloride phase was dependent on fO2, providing a mechanism for Ni transport from primary phases. Jambor and Smith, (1975) and Rucklidge and Patterson (1977), as well as this study, have reported Cl associated with fine grained Fe-enriched lizardite in partially serpentinized dunites, although it is not certain entirely in which phase the Cl resides. The presence of a Cl enriched, Fe-serpentine/hydroxide phase at Dumont which is also associated with higher measured Ni contents, provides evidence that this mechanism may have operated during serpentinization of Dumont dunite and that Ni chloride complexes may have increased the range of Ni mobility on the scale of a grain or two of olivine.

Fluid Chemistry and Thermodynamics

Phase relations

The serpentinization process is commonly thought to occur by diffusion of H2O into the rock in the absence of an aqueous phase. Although it is convenient to discuss the process in terms of the composition of a "fluid" phase, there is no requirement for the existence of a distinct fluid phase, and indeed the extremely low activities of H2O commonly inferred to obtain during serpentinization demand that if such a phase was present it must have been dominated by H2 or

CH4 rather than H2O or CO2 (Evans, 2008; Frost, 1985). At such low fluid/rock ratios, the mineral assemblage can impose on the fluid its values of local chemical potentials of SiO2, H2O,

O2 or H2 and the exchange potential FeMg-1(Evans, 2008). Figure 17a is a phase diagram

calculated at 100 MPa, 200 C (after Fig. 4a of Frost and Beard, 2007) that illustrates the extremely low activities of SiO2 and H2O that will be obtained in a system at the invariant point where olivine, serpentine and brucite coexist. With cooling above the thermal stability limit of serpentine (Fig. 17b), olivine will react with H2O to form serpentine by reaction (4), consuming silica until the silica activity (aSiO2) in the fluid has fallen to the invariant point represented by reaction 5 where brucite coexists with serpentine and olivine. If olivine persists metastably due to the limited supply of H2O at low water activities at temperatures below its thermal stability limit (i.e., the temperature of reaction

(6) then the aSiO2 will fall along the metastable continuation of invariant reaction 5 indicated by the dashed line in Figure 17b. On the other hand, if an excess of H2O is present then olivine will be consumed immediately after T falls below that of reactions (5) and (6), and the stable assemblage brucite + serpentine will exist at equilibrium with a free aqueous phase. Because in reality serpentinization is governed by access of water to the rock in systems characterized by low water-rock ratios and isochemical serpentinization like Dumont, reaction (5) dictates the reaction as the water activity required for (6) to proceed falls with temperature.

3Mg2SiO4 + SiO2 + 4H2O = 2Mg3Si2O5(OH)4 (4)

Mg2SiO4 + 4Mg(OH)2 + 3SiO2 = 2Mg3Si2O5(OH)5 (5)

2Mg2SiO4 + 3H2O = Mg3Si2O5(OH)4 + Mg(OH)2 (6)

Mg2SiO4 + H2O = 2Mg(OH)2 + SiO2 (7)

Mg3Si2O5(OH)4 + H2O = 3Mg(OH)2 + 2SiO2 (8)

To understand the change in fO2 with aSiO2 with the often reduced assemblages of metal alloys observed in serpentinized systems we need to consider the Fe component in olivine. In an idealized isobaric equilibrium model of decreasing temperature, the progressive serpentinization of Fayalite (Fa9-10) should show a temperature interval where (in order of increasing XFe) serpentine, olivine and brucite co-exist with Fe-Mg partitioning between the three phases governed by equilibrium distribution coefficients. However in nature at the low temperatures at which serpentinization proceeds the kinetics of diffusion within the phases are too slow to permit re-equilibration of the olivine composition (Evans 2008, Kunugiza 1982); instead, sluggish kinetics in low temperature serpentinization allows olivine to retain its original composition

(Evans, 2010). Despite the slow diffusion kinetics, the compositions of the coexisting serpentine and brucite are still constrained by the equilibrium KD for Fe-Mg exchange, with the result that extremely Mg-rich serpentine and brucite tend to form, leaving an excess of Fe2+ in the system which must somehow be accommodated (Evans, 2008). A commonly cited means of disposing 2+ of the excess Fe involves oxidation of iron and reduction of H from H2O to produce H2 and magnetite as shown in reaction (9), which describes the comportment of the fayalite component of olivine at the metastable reaction (5) in Figure 17. The formation of brucite due to reaction 5 is inhibited by reaction (9), which drives the formation of extra Mg-serpentine and magnetite while removing Fe from silicates to form magnetite. Under these circumstances fO2 must fall below that of the decomposition of water to H2.

6Fe2SiO4 + H2O + 9Mg(OH)2 = 4Fe3O4 + 4H2 + 3Mg3Si2O5(OH)4 (9)

Although in some serpentinized ultramafic suites magnetite occurs as the dominant Fe- bearing phase along with Mg-rich serpentine, in others little to no magnetite occurs, and instead Fe-enriched serpentine is observed, particularly early in the serpentinization process when silica and water activities are very low and olivine persists as a metastable phase (Toft et al., 1990; O’Hanley and Dyar, 1993; Oufi and Cannat, 2002; Bach et al., 2006; Frost and Beard, 2007; Evans et al., 2009; this study). During low temperature serpentinization (<300°C), lizardite is the stable serpentine mineral (Fig. 17, Grauby et al 1998, Evans et al 2013). Early suggestions that Fe-bearing lizardite (as recognized by Page 1967) tends to be enriched in Fe3+ was confirmed by Mössbauer spectroscopy (O’Hanley and Dyar, 1993). The hydration of the fayalite component of olivine to produce Fe3+ hosted in lizardite as a cronstedtite component in the presence of brucite can be expressed as:

2+ 3+ 3+ 4Fe2SiO4 + 5H2O + 3Mg(OH)2 = 2(Fe2 Fe )(Fe Si)O5(OH)4 + Mg3Si2O5(OH)4 + 2H2 (10)

Reaction 10 shows that in the presence of brucite, the fayalite component of metastable olivine can react to enter lizardite as a cronstedtite component without requiring the formation of magnetite. This reaction would evolve H2 while remaining at conditions close to the very low silica and H2O activities implied by reaction (5). At such low silica activities it might not be

surprising also to encounter the type of hydrogrossular substitution that was suggested above on the basis of observed deficiencies in tetrahedral cations in EMPA data for serpentine at Dumont; however no thermodynamic or experimental data exist to assess whether such increases in H2O content of serpentine would be possible at the very low aH2O that obtains during reaction (5). 3+ The substitution of Fe as a cronstedtite component offers a mechanism for H2 production at very low fO2 without the formation of magnetite as would otherwise be indicated by reaction (9). This proposition is consistent with the results of experiments performed by Seyfried et al, 2007 (at 200°C) and Marcaillou et al 2011 (at 300°C), where seawater was reacted with peridotite at low temperature and low water-to-rock ratios, forming iron-bearing lizardite and little to no magnetite. Moody (1976a,b) and Janecky and Seyfried, (1986) performed similar experiments at similar temperatures (200-300°C), at higher water-rock ratios which resulted in the production of magnetite and Mg-lizardite. Observations of natural serpentinites (ocean floor massifs) generally indicate that Fe- serpentine forms early with little to no magnetite, during the onset of serpentinization, followed by later replacement by the assemblage Mg-serpentine + magnetite in completely serpentinized (but unsteatitized) peridotite or dunite. (Ikin and Harmon, 1983; Toft et al., 1990; Oufi and Cannat, 2002; Bach et al., 2006). While reaction (10) is not directly dependent on aSiO2, the 3+ ability for Fe to substitute for Si in the tetrahedral site, must place an upper limit on the aSiO2 of the fluid. Therefore, the presence of magnetite Mg-lizardite vs Fe-lizardite as cronstedtite may be a sensitive function of temperature, water to rock ratio and aSiO2 (Evans, 2008).

In higher temperature serpentinization (>320°C) magnesian antigorite becomes the stable serpentine (Evans et al., 2013). Above 400°C (for example in fore arc basins) olivine is stable and reaction/diffusion kinetics may be rapid enough to permit Fe-Mg exchange equilibrium between olivine and serpentine to be attained. In such cases olivine can accommodate Fe into its 2+ 3+ composition, no longer requiring the oxidation of Fe to Fe , and as a result no H2 is produced (McCollum and Bach, 2009; Evans at al., 2013). Conversely during serpentinization at very low 2+ temperatures 100-200°C, Fe is accommodated in brucite and there is consequently no H2 production due to oxidation of the iron from the fayalite component (Moody 1976b, Seyfried et al. 2007).

2+ 3+ Where conditions are such that the oxidation of Fe to Fe produces high H2 concentrations approaching the di-hydrogen saturation of water, pre-existing primary sulfide minerals may be destabilized, while Fe-Ni alloys are stabilized (Eckstrand, 1975; Frost 1985; Seyfried et al., 2007; Klein and Bach, 2009). In the sulfur-free system the Ni-Fe alloy can be produced by the Ni and Fe made available from the decomposition of olivine, as inferred by its

Figure 17 A) Log aSiO2 vs log aH2O after Frost and Beard (2007). Pressure and temperature fixed at 1kbar and 200oC. At the onset of hydration of olivine the silica activity will fall until it is defined by the invariant equilibrium between olivine, serpentine, and brucite (5) which buffers water activity and silica activity far below unity. Once olivine has been depleted from the rock silica activity will rise to values determined by the serpentine-brucite buffer

(8) B)Temperature vs log aSiO2. Drawn after Frost and Beard (2007). Symbols: TLC=Talc, SP=Serpentine, BRU=brucite, OL=olivine, EN=Enstatite. Numbers 4 to 9 refer to reactions in text (after Frost and Beard, 2007). Letters A to H represent assemblages in Figure 6 and correspond to red arrows which suggest the path of changing silica activity followed during serpentinization of Dumont dunite.

presence in partially serpentinized sulfur-free rocks, and is stable in the assemblage brucite + serpentine + olivine at fO2 in the range of 4 to 5 log units below FMQ (Frost, 1985). Where sulfides are present, awaruite has been observed to have formed by the desulfidation by (13) and (14) (Nickel, 1959; Kanehira et al., 1964; Chamberlain et al., 1965; Beard and Hopkinson, 2000; Bach and Klein, 2009; Çina, 2010). Under the reducing conditions caused by the generation of copious amounts of H2 by reactions 10 and 11, the high fH2 will tend to diminish both fO2 and fS2 low and push S out of primary sulfides into the fluid phase by forming H2S (reactions 11, 12).

S2g + 2H2,aq = 2H2S,aq (11)

S2 + 2H2O = 2H2S + O2 (12)

Reactions (13) to (16) show how the high fH2 and low fO2 accompanying serpentinization causes desulfidation reactions to produce awaruite and heazlewoodite at equilibrium with magnetite from primary pentlandite. Reactions (17) to (18) are written without explicitly involving MgO-bearing minerals to illustrate how iron contained in serpentine and brucite convert to magnetite, increasing aH2O and SiO2.

2Fe4.5Ni4.5S8 + 16H2 = 2Fe3O4 + 3Ni3Fe + 16H2S (13)

2Fe4.5Ni4.5S8 + 12H2O + 4H2 = 3Ni3S2 + 3Fe3O4 + 16H2S (14)

3Ni3Fe + 6H2S + 5O2 = 3Ni3S2 + Fe3O4 + 6H2O (15)

3NiTaenite+2H2S+O2=Ni3S2+2H2O (16) 2+ 3+ 3+ 3(Fe2 Fe )(Fe Si)O5(OH)4 = 4Fe3O4 + 3SiO2 + 5H2O + H2 (17) 2+ +3 +2 3Fe (OH)2 + 0.5O2 = (Fe 2Fe )O4 + 3H2O (18)

Stability of Ni-Fe Alloys Equations (13) and (15) express the stability of awaruite in the presence of sulfides, however, because of the relationship between fO2, fS2, and aH2S expressed by equations (11) and

(12), Ni-Fe alloys are stable over wider ranges of fO2 under low fS2 conditions representative of low-sulfide assemblages (Fig. 18). The fields near the lower left of Figure 18 are enlarged in the inset to illustrate the variation from pure native Fe at the lowest fO2 through kamacite

(disordered Ni-poor Fe) and awaruite (ordered Ni3Fe) to taenite, which is a disordered  phase more Ni-rich than awaruite and can reach pure Ni composition (e.g., Cacciamani et al., 2010). as shown by Figure 18, the lower the ΣS, the larger the range of stability of alloys until Hz becomes stable phase by (15) at increased fO2’s. Note that the upper limit of fO2 for stability of taenite in

Figure 18 is well below that where carbonate minerals replace graphite, above the millerite boundary. As a result, carbonated ultramafic rocks do not contain Fe-Ni alloy phases.

Serpentinization of the Dumont Sill The freshest Dumont dunites contain olivine with some trace Fe-lizardite, trace magnetite along Mg-serpentine rims and little to no brucite (Fig. 8a). Serpentinization therefore likely began along reaction (4) to produce the Fe-analogue of the idealized Mg-serpentine in Figure 17

o Figure 18: ΣS2 vs fO2 for Ni assemblages of interest at 300 C and 2 kbar (altered after Frost, 1985). Numbers refer to reactions in the text. Letter A to O refer to assemblages observed at Dumont as illustrated in Figure 6. Symbols: Py = pyrite field, Po = pyrrhotite field, Mt = magnetite field, Hm = hematite field Fe = iron field. IM = iron – magnetite buffer, Aw = awaruite, Ta = taenite, Hz = heazlewoodite, Ml = millerite. as well as minor quantities of magnetite (Fig. 14d). Primary magmatic sulfide blebs observed (Fig. 6A) are those most indicative of the original sulfide liquid composition during magmatic

emplacement (dominantly pentlandite) ± some early addition of magnetite due reaction 9 and emplaced within the blebs due to initial volume expansion in the silicate framework surrounding the blebs. As suggested by the consumption of brucite in reactions (9) and (10), incipiently and partially serpentinized Dumont dunites contain minimal brucite, Mg-lizardite along pseudomorphed olivine boundaries and Fe-lizardite aureoles around relict olivine grains, with minor magnetite (Fig. 6B, Fig. 14 d to h). The substitution of, and increase in iron into the 3+ serpentine structure is the result of a decrease in aSiO2 and Fe substitution as a partial cronstedtite component (Fig.16a; reaction 10). This would require the aSiO2 to continue to drop steeply along (4), until (5) prevails (Fig. 17b, path a to c and the first appearance of brucite is observed (Fig. 6c) as brucite centers after olivine. Although the serpentinization reaction at the invariant point represented by reaction (5) should produce brucite at the expense of olivine in a nominally Fe-free system, the reactions involving the fayalite component (i.e., reactions 9 and 10) both diminish brucite production. Formation of the cronstedtite component of serpentine via reaction (10) diminishes brucite production less than does the direct formation of magnetite via reaction ( 9). The aSiO2 of the fluid phase in partially serpentinized samples (Fig. 6b) must be very low to promote substitution of Fe3+ into the cronstedtite component via (10). As shown in

Figure 6c, brucite occurs as centers after olivine suggesting that a drop in aSiO2 after the initial serpentinization has permitted brucite to stabilize (7), suggesting a change in condition, namely a decrease in aH2O (Fig. 17a), caused by the consumption of water to form H2 in a system where new additions of H2O are limited by diffusion.

The effect of increasing H2 on primary sulfide, by diminishing the activity of H2S and thereby destabilizing sulfide minerals, is seen in Figure 6b and Figure 5b, where the coarsest awaruite grains observed at Dumont occur as replacements of primary sulfides (pentlandite) with little to no brucite in the silicate matrix and Fe-Lizardite around relict olivine grains, although the occurrence of these large awaruite grains are not restricted to brucite free rocks (Fig. 6). The effect of this very reducing environment in partially serpentinized Dumont dunites can also be seen within the olivine and Fe-lizardite matrix surrounding the sulfide blebs or in low sulfide assemblages. Micrometer to sub-m sized grains of Ni-Fe alloy and Co-Fe alloys ± magnetite are observed within the silicate matrix (Fig. 5d, j-l, Fig. 18) as inclusions within very fine grained Fe-lizardite, characteristic of the Fe-serpentine facies. However these particles are less prevalent in completely serpentinized Mg-serpentine facies samples (Fig. 5e). In completely

serpentinized Mg-serpentine facies samples, awaruite occurs more abundantly as coarser grains, therefore represented as higher average modal abundances (Table 12) whereas the low sulfide samples in Figure 11 which plot in the high-chrome, “barren” range, contain abundant awaruite. As serpentinization increases to Mg-serpentine facies, Dumont nickeliferous assemblages are marked by increasing portions of Mg-lizardite and magnetite (Fig. 14d) at the expense of Fe- lizardite and Fe-brucite (Fig. 6 d to e and Fig. 14 e,f) until only Mg serpentine, magnetite and Mg-brucite are present (Fig. 6e and onwards). Increasing silica activities after total consumption of olivine and consequent departure from the invariant reaction 4 may consume previously formed brucite and convert it to serpentine through reaction 8 (Fig. 14j). The conversion of Fe in cronstedtite and brucite to magnetite is given by (17) and (18). The required trigger to cause (17) to proceed to the right may be a change in local conditions; for example once olivine is exhausted from the rock reaction (8) in Figure 17 will prevail, which requires water and results in an increase the aSiO2 of the fluid phase. The exhaustion of olivine also limits the amount of

H2 production via (11) and conditions may increase slightly in fO2, and aSiO2 as to drive (17) to the right.

The increase in fO2 of the system due to the exhaustion of olivine (and oxidation of Fe in 2+ serpentine) may be limited by further H2 production by the oxidation of any remaining Fe in the Fe-serpentine facies (Fig. 16b) contained within in serpentine and brucite through (18) . However according to Fig. 16b the majority of the whole rock Fe would have been oxidized in the Fe-serpentine facies stage. Further, reaction (17) which convert Fe3+ in serpentine to magnetite releases H2O , consistent with the drop in LOI observed in the change in Fe-serpentine to Mg-serpentine facies (Fig. 14c).

Evidence for a slight increase in fO2 in the conversion between Fe-serpentine facies, to Mg-serpentine facies is seen in the destabilization of awaruite in the sulfide assemblage (B-D Fig. 6) to more oxidizing assemblages as (13) reverses to produce the assemblages heazlewoodite and pentlandite (e &f Fig. 6 and Fig.18) or heazlewoodite+ magnetite ( Fig. 6g and Fig. 18) in primary intercumulus sulfide blebs and Fe3+ in serpentine is converted abundant to magnetite in completely serpentinized Mg-serpentine facies rocks (Fig. 14 d to e) . The most oxidizing conditions within the Dumont ore body are represented by the heazlewoodite- magnetite assemblage in Fig. 6g. The magnetite content of completely serpentinized Mg- serpentine facies rock (Fig. 6 e to g) is on the order of double of those that are incipiently,

partially and completely serpentinized Fe-serpentine facies (Fig. 6 a to d). Although trace millerite was observed in 2 of 1420 ExplominTM samples, these were restricted to the lower peridotite zone where aSiO2 and fO2 conditions were higher or near large regional fault zones where fluid flux was probably high, postdating the preliminary serpentinization hydrothermal event. Similarly in low sulfide zones, the previously discussed reactions involving the silicate and hydroxide matrix prevail, and as such buffer the fluid phase, the major difference being that these reactions occur at lower ΣS2 due to the lack of primary sulfides. Numerous previous publications which comment on the stability of Ni-Fe alloys have noted a tendency for the occurrence of Ni-Fe alloys found in peridotites at the early stages of serpentinization, and their general absence in rocks that have been completely serpentinized (Frost, 1985; Ashley, 1975; Eckstrand, 1975). At Dumont awaruite does occur in completely serpentinized rocks, defined by samples bearing no relict olivine, however these samples are either in non-sulfide, non-carbonate bearing Mg-serpentine facies, or in sulfide bearing, Fe-serpentine facies. According to the classification of Ni-Fe alloys (Cacciamani et al., 2010) Dumont nickel-enriched “awaruite” may plot in the “taenite” field between 200-300°C and occurs with relatively nickel enriched magnetite; this is consistent with the experimental work of Filippidis (1985), who found that serpentinization in the sulfur-free system produced an opaque mineral assemblage of Ni-rich magnetite + Ni-rich “awaruite”. Ni rich awaruite or native nickel has been reported at various localities in association with heazlewoodite (Ahmed and Bevan, 1981; Auge et al., 1999) and Ni- rich magnetite or “trevorite” (Hudson and Travis, 1981). Awaruite formed after primary sulfides displays a consistent compositional range of approximately 70-75% Ni at Dumont and in the literature (Chamberlain, 1965; Chamberlain et al., 1965; Radhakrishna et al., 1985; Beard, 2000; Bach and Klein, 2009; Evans 2009), whereas Fe-Ni alloys associated with serpentine, olivine and magnetite in non-sulfide or low sulfur zones at Dumont and elsewhere show much more compositional variability of between 60% Ni to that of native Ni (This study, Chamberlain 1965; Dick, 1974; Bird and Weathers, 1979; Hudson and Travis, 1981; Fillidipis, 1985; Auge, 1999).

Table 12: Average modal abundance for minerals of interest as reported by ExplominTM particle scan

NI ppm S % Upper Dunite AW PN HZ MT BRU NISI SP OL N (XRF) (ICP-41) Partially serpentinized 0.08 0.04 0.03 1.45 1.89 68.4 2869 0.019 80.7 14.7 51 Low sulfide Completely serpentinized Fe-serpentine facies 0.12 0.02 0.03 2.48 2.74 59.3 2690 0.018 90.1 0.9 74 Completely serpentinized Mg-serpentine facies 0.13 0.01 0.06 4.66 2.98 50 2701 0.022 90.3 0.1 143 Partially serpentinized 0.08 0.73 0.14 1.74 2.98 31.4 5079 0.303 72.1 17.2 27 Completely serpentinized Fe-serpentine facies 0.09 0.99 0.06 1.95 3.6 22.4 4862 0.34 81.4 0.5 31 sulfide Completely serpentinized Mg-serpentine facies PN 0.08 0.73 0.15 4.28 3.7 12.6 4910 0.276 84.5 0.08 30 Completely serpentinized Mg-serpentine facies HZ 0.1 0.1 0.45 4.43 2.21 14 3938 0.161 89.5 0.1 111 NI ppm S% Lower dunite AW PN HZ MT BRU NISI SP OL N (XRF) (ICP-41) Partially serpentinized 0.06 0.04 0 2.62 1.04 72.9 2334 0.0197 85 18.6 43 Low-sulfide Completely serpentinized Fe-serpentine facies 0.1 0.02 0 4.83 1.48 58.3 2433 0.0207 90.1 0.89 130 Completely serpentinized Mg-serpentine facies 0.12 0.01 0 7.12 2.42 45.2 2390 0.1212 87.3 0.09 279 Partially serpentinized 0.03 0.79 0.09 2.2 2.42 38.6 4028 0.0209 74.7 18.8 19 Completely serpentinized Fe-serpentine facies 0.04 0.93 0.07 3.95 2.94 17.6 3888 0.2878 86.2 1.5 61 sulfide Completely serpentinized Mg-serpentine facies PN 0.07 1.04 0.17 8.33 2.1 12.6 4470 0.1662 84.7 0.06 96 Completely serpentinized Mg-serpentine facies HZ 0.08 0.06 0.3 6.6 1.35 22.5 3271 0.1214 80.9 0.09 14 Ni in silicates (NISI) is a calculated value based on the modal abundances of pentlandite (Pn), heazlewoodite (Hz) and awaruite (Aw) in the ExplominTM sample. % Ni in silicates = [(Nickel Assay - Metallic Nickel)/Nickel Assay], where the metallic nickel = % Modal abundance of Pn * %Ni in Pn + % Modal abundance of Hz * %Ni in Hz + % Modal abundance of Aw * %Ni in Aw. Where heazlewoodite modal abundance <0.1%, the average value of 27.3% Ni in Pn from electron microprobe data was used, for heazlewoodite modal abundance >=0.1, 32% Nickel was used for pentlandite. 73% and 72% Ni was used for Aw and Hz respectively across all domains. “Non-sulfide” is considered to be samples with sulfur <0.05%. AW=awaruite, PN=Pentlandite, Hz=heazlewoodite, MT=magnetite, BRU=brucite, SP=Total Mg and Fe Serpentine, OL=Olivine, N is the number of ExplominTM samples. Minerals reported in %,

Filippidis’ (1985 and 1982) experimental work has shown that awaruite compositional variability is controlled by fS2, and fO2, Ni-rich compositions being produced at lower fS2. In sulfide bearing assemblages at Dumont desulfurization of pentlandite will have occurred within a small range of fO2 and fS2, resulting in the restricted range of Ni compositions observed, whereas in the low fS2 case where primary sulfide minerals were absent, there is larger range of fO2 over which the Ni-Fe alloys are stable, resulting in more variable nickel contents. At Dumont, awaruite associated with sulfide contains consistently ~73wt% Ni whereas awaruite associated with low sulfide serpentinite assemblages exhibits a wider range of compositions which includes the more Ni-rich population that may actually represent the disordered  phase approaching pure Ni. If the compositions of serpentinizing fluids are governed by the composition of the protolith due to very low fluid-rock ratios (i.e., diffusion-controlled ingress of H2O), olivine- 2+ 3+ serpentine partitioning of Fe is likely to govern the fO2 via the oxidation of Fe into Fe . Equations (19) and (21) demonstrate that the greater the activity of Fe serpentine or Fe brucite in their respective host phases, the lower the fO2. It follows that dunites or peridotites that contain higher iron content may be capable of producing more reducing environments trending towards more Fe-rich alloys, whereas those with higher forsterite contents may produce less reducing conditions. Dumont awaruite exhibits this tendency. Awaruite contains lower nickel contents are found in the lower dunite, which contained more fayalitic olivine, whereas awaruite from the more forsteritic upper dunite tends to be more nickel rich (Table 8) It must also be noted that Mg rich olivine also typically contains higher Ni contents (This study; Barnes et al., 2011b; Sobolov et al., 2007), and that the availability of nickel from primary olivine to produce Ni-Fe alloy may play a part in terms of the abundance and composition of awaruite (Table 12, Fig. A-6). While average abundances of awaruite are similar for upper dunite and lower dunite at Dumont, the upper dunite shows more low-sulfide samples containing higher abundance of awaruite. The experiments work of Fillipidis (1985) showed a similar result where the olivine with higher nickel contents produced relatively increased modal abundances of awaruite. Insufficient information on olivine compositions and awaruite (sulfide vs non sulfide associations, olivine compositions, modal abundances etc.) makes it difficult to compare Dumont to the existing literature on other alloy occurrences.

Thus it is argued that the metallic and sulfide nickel assemblages observed at Dumont represent the changes in fO2 and ΣS2 by rock-buffered serpentinizing fluids in a low temperature, retrograde environment The sequence of serpentinization began in the partially serpentinized facies, followed by the Fe-serpentine facies and lastly followed by the Mg-serpentine facies.

Evolution of H2 and thus fO2 captured by Dumont assemblages is consistent with the recent experimental work of Marcaillou et al., (2011) where peridotite was reacted with seawater at

300°C and 300 bar to assess the mineralogical products and H2 production. They found the hydrogen production occurred in three stages: (1) initial H2 production by preliminary magnetite formation; (2) an intermediate stage when Fe3+ is incorporated into serpentine and (3) a third stage during which the amount of magnetite doubles and the serpentine becomes increasingly Mg-rich by the destruction of Fe serpentine. They found that the amount of Fe3+ in serpentine or 3+ magnetite was directly related to the H2 production and that the incorporation of Fe in serpentine accounted for up to 50% of the H2 produced. The arrow in Figure 6 is drawn to express changes in metallic and silicate/hydroxide mineralogy to represent the progressive low temperature serpentinization as represented by Dumont assemblages.

Mass Transfer to Intercumulus Sulfide Blebs and Nickel Remobilization Figure 5a illustrates the transfer of Fe between the silicate matrix and the sulfide blebs during progressive serpentinization at Dumont. As olivine is consumed it is replaced first by Fe- rich serpentine close to the reaction front, and later by Mg-rich serpentine farther from the serpentinization front. The total serpentine content in the sample from Figure 5a is 52.7% (as reported by ExplominTM), 15.7% of which is Fe-serpentine. This suggests some of the iron originally contained in the olivine, has been transferred to the sulfide bleb to create magnetite during expansion of intercumulus sulfide blebs, where it appears as dilatant fillings along cleavage planes in pentlandite. This is consistent with Duke's (1986) observation that the Fe/Ni ratio of intercumulus sulfide blebs increases with serpentinization at Dumont. Magnetite can also be added to blebs during the desulfurization of pentlandite to produce awaruite and heazlewoodite via reactions (15) and (16). The volume expansion of blebs along with several modes of magnetite production in intercumulus sulfides blebs renders it challenging to perform mass balance calculations using 2D ExplominTM field stitches to assess whether Ni has also been released from silicate matrices to upgrade the nickel content of intercumulus

sulfide blebs on a sample per sample basis. We look to the variability and distribution of the metallic nickel and silicate phases observed at Dumont and the variability of such with the progress of serpentinization to gain an understanding of nickel remobilization in a macroscopically isochemical system.

Nickel Distribution and Controls Serpentinization has remobilized nickel contained in primary silicates to intercumulus blebs, resulting in an increase in the tenor of the nickel bearing phases in the intercumulus blebs (e.g., Figs. 6, 8 13; Duke, 1986). As a result, in zones where serpentinization is incipient, partial Fig. 19, Domain 4) or near complete but remaining within the Fe-serpentine facies, the percentage of nickel in silicates existing within the silicates structure or as microscopic inclusions of alloy and sulfide is generally higher (Table 12). This incomplete remobilization of nickel to the intercumulus blebs has resulted in the population of lower tenor pentlandite (Fig. 9 A-4a) associated with incomplete serpentinization. As serpentinization continues into the Mg serpentine facies and olivine is consumed, magnetite is produced at the expense of Fe-lizardite and Fe-brucite, while the remaining serpentine and brucite become increasingly magnesium rich (Fig. 6d to g). At this point, the stability fields for serpentine and brucite expand, fO2 rises, and both awaruite and pentlandite are destabilized to produce heazlewoodite via reactions 15 and 16 (Fig. 6 transition from D to E; Fig. 18). This process coincides with a decrease nickel concentration in silicates (Table 12). In Mg-serpentine zones where serpentinization is complete (Fig. 8a), intercumulus blebs contain abundant magnetite ± pentlandite ± heazlewoodite with little to no awaruite (Fig. 6, e to h). Generally, heazlewoodite and awaruite exhibit negative correlation on a zone scale. Where heazlewoodite content is high, awaruite is low. Where sulfides are not present, awaruite exists as finely disseminated grains associated with magnetite or brucite mesh rims. However, on the scale of a thin section, heazlewoodite and awaruite can occur together in the same bleb. Nickel remobilization to intercumulus spaces has been completed by late-stage serpentinization, thus the percentage of nickel hosted in silicates is generally lower (Table 12 ) in Zones 1, 2 and 3b in Figure 19 and the nickel tenor of pentlandite is higher where it is present. This is represented by the population of higher Ni tenor of 30% to 35% in (Fig. A-4A).

In late-stage reactions beyond Mg-serpentine facies with heazlewoodite, steatitization can occur, accompanied by replacement of magnetite by sulfur-rich nickel sulfides such as millerite (Fig. 6G). These transitions indicate increasing oxygen and sulfur fugacities (Eckstrand, 1975; Frost 1985) where serpentine and brucite are replaced by talc and carbonate (Bach and Klein, 2008, Frost and Beard 2007). This is observed rarely locally within the Dumont dunite around major structures, but regularly at the basal contact of the intrusion (well outside of resource envelope).

Figure 19. Distribution of sulfide minerals and Fe serpentine facies alteration in the block model (modified after Sciortino et al., 2013). Serpentinization zones which are analogous to metallurgical domains (Royal Nickel, 2013) correspond to assemblages A to G in Fig 6.. Zone 1 (Fig. 6G); heazlewoodite dominant, fully serpentinized +/- awaruite (metallurgical domain heazlewoodite dominant, Hz/Pn>=5, SPFE<14. Zone 2 (Fig. 6 E to G); low iron serpentine, mixed sulfide, pentlandite and heazlewoodite +/- awaruite (metallurgical domain: mixedsulfide,114. Note that assemblage H is not displayed because it does not corresponds to broad zones but is restricted locally to large fault zones and the basal contact which is found outside of the mineralization envelope. (Note: HZPN is the heazlewoodite to pentlandite ratio, SPFE is high iron serpentine).

The silicate or matrix phase assemblages observed for low sulfide layers display the same range of serpentinization as those for sulfide bearing assemblages, except that they contain trace to no sulfides. Where serpentinization is incipient, the nickel in silicates is higher which in turn is associated with low values of awaruite modal abundances (Table 12). In low grade, non-sulfide zones where serpentinization is complete Mg-serpentine facies, the nickel in silicates values are lower and the modal abundances of awaruite is highest (Table 12).This evidence suggests that where serpentinization is incipient or partial, and the remobilization of nickel is not complete, more nickel is hosted in silicates as opposed to creating awaruite, as suggested by the higher reported probe values for nickel in serpentine and magnetite in non-sulfide samples (Table 8). Where re-mobilization is complete (in Mg-serpentine facies) larger amounts of nickel have gone into forming awaruite, leaving less nickel in silicates. Mg-serpentines are Fe-poor up to Mg # 97, and the Fe-enriched lizardite which is so prevalent in initially, partially and completely (Fe- serpentine facies) serpentinized samples show a trend for increasing nickel content with iron . Therefore, to consider the nickel component in incipient Fe-lizardite, we propose iron serpentine converts to magnetite via reaction (17) and as a result produces more awaruite instead of reaction (5) given in Evans 2008. This would account for the increase in awaruite (above the resolution of ExplominTM) and decrease in Ni in serpentine.

Implications Weakly serpentinized dunite contains the highest nickel in silicates, whereas those that are completely serpentinized contain the least and in addition contain higher nickel tenor minerals. It then follows that better recoveries as well as higher concentrate grades may be expected from completely serpentinized Mg-serpentine facies samples, which is what is observed in the metallurgical tests results of the Dumont dunite (Royal Nickel Corporation, 2013). The bulk isochemical nature of the serpentinization process, combined with the grain-scale redistribution of elements including Ni to secondary phases, has allowed Royal Nickel to subdivide the Dumont dunite into mineralogical domains as shown in Figure 19 to predict Ni recovery (Sciortino et al. 2013). Understanding the process of serpentinization and the distribution of resulting secondary minerals is the key to unlocking the values in the Dumont dunite and other similarly hosted, low grade ultramafic bodies which were similarly serpentinized.

Conclusions and Exploration Implications Low temperature serpentinization occurs under non-equilibrium conditions due to sluggish diffusion kinetics, however local equilibrium during reactions involving metastable phases allows reliable predictions of the observed assemblages based on a thermodynamic approach to the phase relations both at Dumont and at other localities described in published literature. Slow diffusive fluxes of H2O to the reaction site permit the mineral assemblage in the 2+ rock to impose on the fluid its values of the chemical potentials of H2O, H2, SiO2 and Fe Mg-1.

Rates of reaction (consumption of H2O and release of H2) and delivery of H2O will be very slow, classic of retrograde serpentinization of massifs (Evans, 2008). Nickel release from primary silicates during the serpentinization of ultramafic bodies in low temperature isochemical system may lead to the formation of nickel-iron alloys and upgrading of primary sulfides leading to and increase available recoverable nickel in secondary metallic phases. The production of Ni-Fe alloys from primary silicates (or after primary sulfides) requires reducing conditions, requires local changes in fO2 which are driven by increased H2 potentials due to the oxidation of Fe2+ in primary silicates to Fe3+ in secondary Fe-lizardite and magnetite. Such reducing conditions are only observed during serpentinization over a limited range of temperatures (~200-350°C) at low water to rock ratios where the evolved H2 may be concentrated in the fluid phase The amount of nickel available for upgrading and production of Ni-Fe alloys may be maximized in systems that are dominanted by olivine and that contain abundant nickel in primary olivine. Sobolev et al., (2007) compared the compositions of olivine from various magmatic systems and showed that the greatest Ni contents in terrestrial magmas are found in highly forsteritic olivine associated with komatiitic magmas. Thus the most favorable rock type for Ni enrichment through serpentinization is Dumont type, komatiitic extreme adcumulate dunites, followed by komatiitic meso and ortho cumulate dunites, komatiitic peridotites and mid ocean ridge peridotites. Nickel upgrading during serpentinization of ultramafic rocks does not occur universally, because a specific set of conditions must be met during the serpentinization process. The key to maintaining the exceptionally low fO2 required for the processes documented at Dumont is for serpentinization occur at the serpentine-brucite-olivine invariant reaction point by diffusive transport of H2O in the absence of a free aqueous phase. Therefore in cases where the ultramafic

rock has undergone extensive deformation and brittle fracturing, potentially producing open system conditions and allowing water to penetrate along fractures as might be expected in alpine scenarios, any formerly produced Ni-Fe alloys may be destabilized (however there is plenty of evidence for alloys occurring in alpine ophiolites). This may be indicated by the occurrence of extensive fractures, chrysotile vein formation, and potentially several generations of overlapping serpentine growth. Serpentinization at high temperature and pressure typical of fore arc basins and subduction-related serpentinization in general may permit diffusive re-equilibration of olivine with serpentine, eliminating the driver to extremely high Mg# in serpentine and the 2+ resulting release of Fe into the fluid. The result is little to no magnetite and H2 production, generating Mg-antigorite in an oxidizing environment in which Ni-alloys may not be expected..

Late stage serpentinization is marked by increase aSiO2 and increased fO2, marked by the presence of talc and talc carbonate alteration. In an open systems specifically including contact zones with changes in lithology, where there is no mechanism to retain produced H2, fO2 may also rise with an accompanying increase in aSiO2 thus formatting talc, destabilizing any previously formed alloys. Therefore talc-carbonates in may be an indicator for destabilization of Ni-Fe alloys, however under such conditions also may result in S-rich assemblages containing nickel originally hosted in silicates (Groves et al., 1974; Groves and Keays, 1979;

Dundonaldson, 1981). The presence of localized talc in serpentinite is evidence of higher aSiO2 for local alteration either as a result of abundance of pyroxenes available during serpentinization or as a result of proximity to a silica source, namely cross cutting dykes. The higher fO2 and fS2 associated with this result is higher sulfide assemblages whereby Ni-Fe alloys are not stable. Serpentinization of the Dumont dunite occurred at low temperatures and low rock-water ratios. The resulting mineral assemblage records the changing fO2 and fS2 as result of the progressive serpentinization. Early serpentinization was marked by reducing assemblages 2+ including Fe-Ni alloys as a result of the H2 production via the oxidation of Fe in the olivine. Late serpentinization assemblages mark slightly more oxidizing conditions where Ni rich pentlandite and heazlewoodite occur with Mg-serpentine and abundant magnetite. During the process, nickel originally contained in primary silicates (olivine) has been re-mobilized to secondary products to produce awaruite or enriched Ni-sulfides. As a result the degree of serpentinization can be used to infer the distribution of nickel between metallic and non-metallic phases and thus economic potential.

References

Abrajano, T.A., and Pasteris, J.D., 1989, Zambales Ophiolite, Philippines: Sulfide petrology of the critical zone of the Acoje Massif: Contributions from Mineralogy and Petrology, v. 103, p. 64-77.

Ahmed, Z., and Bevan, J.C., 1981, Awaruite, iridian awaruite, and a new Ru-Os-Ir-Ni-Fe alloy from the Sakhakot-Qila complex, Malakand Agency, Pakistan: Mineralogical Magazine, v. 44, p. 225-230.

Anselmi, B., Mellini, M., Viti, C., 2000, Chlorine in the Elba, Monti Livornesi and Murlo serpentinites: evidence for sea water interaction: European Journal of Mineralogy, v. 12, p.137- 146.

Arndt, N.T., 1986, Differentiation of komatiite flows: Journal of Petrology, v. 27, p. 279–301.

Arndt, N.T., Teixeria, N.A., and White, W.M., 1989, Bizarre geochemistry of komatiites from the Crixás greenstone belt, Brazil: Contributions from Mineralogy and Petrology, p. 101, v. 187– 197.

Ashley, P.M., 1975, Opaque mineral assemblages formed during serpentinization in the Coolac ultramafic belt, New South Wales: Journal of the Geological. Society Of Australia, v.22, p. 91- 102.

Auge, T., Cabri, L.J., Legendre, O., McManhon, G., and Cocherie, A., 1999, PGE distribution in base-metal alloys and sulfides of the New Caledonia Ophiolite: The Canadian Mineralogist, v. 37, p. 1147-1161.

Bach, W., Paulick, H., Garrido, C.J., Ildefonse, B., Meurer, W.P., and Humphris, S.E, 2006, Unraveling the sequence of serpentinization reactions: petrography, mineral chemistry, and petrophysics of serpentinites from MAR 15°N (ODP Leg 209, Site 1274): Geophysical Research Letters, v. 33, p. 1-4.

Ballhaus, C., Berry R.F., Green, D.H., 1991, High Pressure experimental calibration of the olivine-orthoyroxene-spinel oxygen geobarometer: implications for the oxidation state of the upper mantle: Contributions to Mineral and Petrology, v. 107, p. 27-40.

Barnes, S.J., and Brand, N.W., 1999, The distribution of Cr, Ni and chromite in komatiites, and applications to exploration for komatiites-hosted nickel sulfide deposits: ECONOMIC GEOLOGY, v. 97, p. 129-132.

Barnes, S.J., Godel, B.M., Locmelis, M., Fiorentini, M.L., Ryan, C.G., 2011b, Extremely rich Fe-Ni sulfide assemblages in komatiitc dunite at Bethano, Western Australia: results from synchrotron x-ray fluorescence mapping: Australian Journal of Earth Science, v. 88, p. 691-709.

Beard, J.S., 2000, A fossil, serpentinization-related hydrothermal vent, Ocean Drilling Program Leg 173, Site 1068 (Iberia Abyssal Plain): Some aspects of mineral and fluid chemistry: Journal of Geochemical Research, v. 105, no. B7, p. 16527-165239.

Bird, J.M., and Weathers, M.S, 1979, Origin of josephinite: Geochemical Journal, v. 13, p. 41- 55.

Bogolepov, G., 1969, Problem of serpentinization of ultrabasic rocks: International Geology Review., v. 12, p. 421-482

Brenan, J.M., and Caciagli, N.C., 2000, Fe-Ni exchange between olivine and sulfide liquid: Implications for oxygen barometry in sulfide-saturated magmas: Geochimica Cosmochimica Acta, v. 64, p. 307–320.

Cacciamani, G., Dinsdale, A., Palumbo, M., Pastruel, A., 2010, The Fe-Ni system: Thermodynamic modelling assisted by atomistic calculations: Intermetallics, v. 18, p. 1148- 1162.

Daval, D., Sissmann, O., Menguy, N., Saldi, G.D., Guyot, F., Martinez, I., Corvisier, J., Garcia, B., Machouk, I., Knauss, K.G., and Hellmann, R., 2011, Influence of amorphous silica layer formation on the dissolution rate of olivine at 90 °C and elevated pCO2: Chemical Geology, v. 284, p. 193–209,

Chamberlain, J.A., 1966, Heazlewoodite and awaruite in serpentinites of the Eastern Township, Quebec: Canadian Mineralogist, v. 8, p. 519-522.

Chamberlain, J.A., McLeod, C.R., Traill, R.J., and Lachance, G.R., 1965, Native Metals in the Muskox Intrusion: Canadian Journal of Earth Sciences, v. 2, p. 188-215.

Charlou, J.L., Fouquet, Y., Bougault, H., Donval, J.P., Etoubleau, J., Jean-Baptiste, P., Dapoigny, A., Appriou, P., and Rona, P.A., 1998, Intense CH4 plumes generated by serpentinization of ultramafic rocks at the intersection of the 15820’N fracture zone and the Mid- Atlantic Ridge: Geochimica et Cosmochimica Acta, v. 62, p. 2323-2333.

Çina, A., 2010, Pentlandite and mineralization related to Albanian ophiolites.,Scientific Annals, School of Geology, Aristotle University of Thessaloniki Proceedings of the XIX CBGA Congress, Thessaloniki, Greece Special v. 100, p. 317-323.

Chown, E.H., Daigneault, R., Mueller, W., Mortensen, J.K., 1992,Tectonic Evolution of the Northern Volcanic Zone, Abitibi belt, Quebec, Canadian Journal of Earth Sciences, v.29(10), 2211-2225.

Cross, W., Iddings, J.P., Pirsson, L.V., and Washington, H.S., 1903, Quantitative classification of igneous rocks, University of Chicago Press, London, 286 p.

Dahlberg, E.H., and Saini-Eidukat, B., 1991, A chlorine-bearing phase in the drill core of serpentinized troctolitic rocks of the Duluth Complex, Minnesota: Canadian Mineralogist, v. 29, p. 239-244.

Dick, H.J.B., 1974, Terrestrial nickel-iron from the Josephine peridotite, Its Geologic occurrence, associations, and origin: Earth and Planetary Science Letters, v. 24, p. 291-298.

Donaldson, M.J., 1981, Redistribution ore elements during serpentinization and talc–carbonate alteration of some Archaean dunites: Western Australia ECONOMIC GEOLOGY, v. 76, p. 1698–1715.

Duke, J.M., 1986, Petrology and economic geology of the Dumont Sill: An Archean intrusion of komatiitic affinity in northwestern Quebec: Geological Survey of Canada, no. 35, p. 1-56.

Eckstrand, O.R., 1975, The Dumont serpentinite: A model for control of nickeliferous Opaque Mineral Assemblages by Alteration Reactions in Ultramafic Rocks: ECONOMIC GEOLOGY, v. 70, p. 183-201.

Edel'shteinI,. I., 1968, Petrology and nickel content of ultrabasic intrusion in the Tobol-Buryktal area of the southern Urals: Magmatizm, Metamorfizm, Metallogeniya Urala, Akad. Nauk SSSR, Ural'sk Filial, Gorn.-Geol. Inst. Tr. Pervogo Ural'sk. Petrogr.Soveshch.S, verdlovsk1, 961, p. 19- 28

2+ Evans, B.W., 2008, Control of the products of serpentinization by the Fe Mg-1 exchange potential of olivine and orthopyroxene: Journal of Petrology, v. 49, no. 10, p. 1873-1887.

Evans, B.W., 2010, Lizardite versus antigorite serpentinite: Magnetite, hydrogen, and life : Geology, v. 38, no. 10, p. 879-882.

Evans, B.W., Kuehner, S.M., and Chopelas, A., 2009, Magnetite-free, yellow lizardite serpentinization of olivine websterite, Canyon Mountain complex, N.E. Oregon: American Mineralogist, v. 94, p. 1731-1734.

Evans, B.W., Hattori, K., Baronnet, A., 2013, Serpentinite: What, Why, Where?: Elements, v.9, p. 99-106.

Faust, G.T., 1963, Minor elements in serpentine—additional data: Geochimica et Cosmochimica Acta, v. 27, p. 665-668.

Filippidis, A., 1982, Experimental Study on the Serpentinization of Mg-Fe-Ni Olivine in the Presence of sulfur: Canadian Mineralogist, v.20, p.567-574

Filippidis, A., 1985, Formation of Awaruite in the System Ni-Fe-Mg-Si-O-H-S and Olivine Hydration with NaOH Solution, An Experimental Study: ECONOMIC GEOLOGY, v. 80, p. 1974-1980.

Fleet, M.E., and MacRae N.D., 1988, Partition of Ni between olivine and sulfide: Equilibria with sulfide-oxide liquids: Contributions from Mineralogy and Petrology, v. 100, p. 462–469.

Frost, R.B., 1985, On the Stability of Sulfides, Oxides, and Native Metals in Serpentinite: Journal of Petrology, v. 26, Part 1, p. 31-63.

Frost, R.B., and Beard J.S., 2007, On Silica Activity and Serpentinization: Journal of Petrology, v. 48, no. 7, p. 1351-1368.

Gaetani, G.A., and Grove, T.L., 1997, Partitioning of moderately siderophile elements among olivine, silicate melt and sulfide melt: Constraints on core formation in the Earth and Mars: Geochimica et Cosmochimica Acta, v. 61, p. 1829–1846.

Godwin, A.M., and Ridler, R.H., 1970, Abitibi orogenic belt: Geological Survey of Canada, paper 70-40, p. 1-24.

Gouze, P., Luquot, L., Andreani, M., Godard, M., and Gibert, B., 2011, Serpentinization of olivine by seawater: A fl ow-through experiment: American Geophysical Union, Fall Meeting 2011, abstract V41C-2506.

Grauby, O., Baronnet, A., Devouard, B., Schoumacker, K., and Demirdjian, l., 1998, The chrysotile-polygonal serpentine suite synthesized from a 3MgO-2SiO2-excess H2O gel , EPMG, 7th,Orleans, Terra Nova Abstract Supplement , p.24.

Groves, D.I., and Keays, R.R., 1979, Mobilization of ore-forming elements during alteration of dunites, Mt Keith-Betheno, Western Australia: Canadian Mineralogy, v. 17, p. 373–389.

Groves, D.I., Hudson, D.R., and Hack, T.B.C., 1974, Modification of iron–nickel sulfides during serpentinization and talc–carbonate alteration at Black Swan, Western Australia: ECONOMIC GEOLOGY, v. 69, p. 1265–1281.

Harris, D.C. and Nickel, E.H., Pentlandite compositions and associations in some mineral deposits: Canadian Mineralogist, v.11, p.861-878.

Hudson, D.R., and Travis, G.A., 1981, A native nickel-heazlewoodite-ferroan trevorite assemblage from Mount Clifford, Western Australia: ECONOMIC GEOLOGY, v. 76, p. 1686- 1697.

Ikin, N.P., and Harmon, R.S., 1983, Mineralogy and petrology of the highland Border Suite serpentinites: Mineralogical Magazine, v. 47, p. 301-310.

Jambor, J.L., and Smith, C.H., 1964, Olivine composition determination with small diameter X- ray powder camera: Mineralogical Magazine, v. 33, p. 730-741.

Janecky, D. R. & Seyfried,W. E., Jr ,1986, Hydrothermal serpentinization of peridotite within the ocean crust: Experimental investigations of mineralogy and major element chemistry: Geochimica et Cosmochimica Acta, v.50, p. 1357-1378.

Kanehira, K., Banno, S., and Hashimoto, M., 1964, Notes on rock forming minerals (28) finding of awaruite (native nickel-iron) from serpentinite near the city of Koti, Sikoku: The Journal of the Geological Society Japan, v. 70, no.824, p. 272-277.

Keays, R.R, and Jowitt, S., 2013, The Averbury Ni deposit, Tasmania: A case study of an unconventional nickel deposit: Ore Geology Reviews, v. 52, p. 4-17.

Keays, R.R., and Kirkland, M.C., 1972, Hydrothermal mobilization of gold from copper–nickel sulfides and ore genesis at Thomson River Copper Mine, Victoria, Australia: ECONOMIC GEOLOGY, v. 67, p. 1263–1275.

Kelley, D.S., Karson, J.A., Blackman, D.K., Fru«h-Green, G.L., Butterfield, D.A., Lilley, M.D., Olson, E.J., Schrenk, M.O., Roe, K.K., Lebon, G.T., and Rivizzigno, P., 2001, An off-axis 17 hydrothermal vent field near the Mid-Atlantic Ridge at 30: Nature, v. 412, p. 145-149.

King, H.E., Plümper, O., and Putnis, A., 2010, Effect of secondary phase formation on the carbonation of olivine: Environmental Science & Technology,v. 44, p. 6503–6509,

Kitakaze, A., Sugaki, A., Itoh, H., and Komatsu, R., 2011, A Revision of Phase Relations in the System Fe-Ni-S from 650 to 450°C: The Canadian Mineralogist, v. 49, p. 1687-1710, doi: 10.3749/canmin.49.6.1687.

Klein, F., and Bach, W., 2009, Fe-Ni-Co-O-S Phase Relations in Peridotite-Seawater Interactions: Journal of Petrology, v. 50, no. 1, p. 37-59.

Kunugiza, K., 1982, Formation of zoning of olivine with progressive metamorphism of serpentinite: an example from the Ryumon peridotite body of the Sanbagawa metamorphic belt, Kii peninsula: Journal of the Japanese Association of Mineralogy, Petrology and ECONOMIC GEOLOGY, v. 77, p. 157-170.

Lahaye, Y., Arndt, N.T., Byerley, G., Chauvel, C., Fourcade, S., and Gruau, G., 1995, The influence of alteration on the trace-element and Nd isotopic compositions of komatiites: Chemical Geology, v. 126, p. 43–64.

Liu, W., Migdisov, A., Williams-Jones, A., 2012, The stability of aqueous nickel chloride complexes in hydrothermal solution: results of UV-Visible spectroscopic experiments: Geochemica et Cosmochemica Acta, v. 94, p. 276-290.

Marcaillou, C., Muñoz, M., Vidal, O., Parra, T., and Harfouche, M., 2011, Mineralogical evidence for H₂ degassing during serpentinization at 300°C/300 bar: Earth and Planetary Science Letters, v. 303, p. 281-290.

McCollom, T.M., Bach, W., 2009, Thermodynamic constraints on hydrogen generation during serpentinization of ultramafic rocks: Geochimica et Cosmochimica Acta, v. 73, p. 856–875.

Misra, K.C., and Fleet, M.E., 1973, The Chemical Compositions of synthetic and natural pentlandite assemblages: ECONOMIC GEOLOGY, v. 68, p. 518-539.

Miura, Y., Rucklidge, J., and Nord G.L., 1981, The Occurrence of Cl in Serpentine –Minerals: Contributions to Mineralogy and Petrology, v. 76, p. 17-23.

Moody, J.B., 1976a, An experimental study of the serpentinization of iron-bearing olivines: Canadian Mineralogist, v. 14, p. 462-478.

Moody, J.B., 1976b, Serpentinization: a review: Lithos, v. 9, p. 125-138.

Moretti, R. and Baker, D. R., 2008, Modeling the interplay of fO2 and fS2 along the FeS-1083 silicate melt equilibrium. Chem. Geol. 256, 286-298.

Mungall, J.E., and Brenan, J.M., 2014, Partioning of platinum-group elements and Au between sulfide liquid abd basalt and the origins of mantle rust fractionation of the chalcophile elements: Geochimica et Cosmochimica Acta, in prep

Mumpton, F.A. and Thompson C.S., 1966, The stability of brucite in the weathering zone of the New ldria serpentinite, Clay and Clay Minerals National Conference, 4th, New York, Proceedings, p. 249-257.

Naldrett, A.J., 1966a, Talc-carbonate alteration of some serpentinized ultramarine rocks south of Timmons, Ontario: Journal of Petrology, v. 7, p. 489-499.

Nickel, E.H., 1959, The occurrence of native nickel-iron in the serpentine rock of the Eastern Townships of Quebec Province: Canadian Mineralogist, v. 6, p. 307-319.

O’Hanley, D.S., 1996, Serpentinites: Records of tectonic and petrologic history: Oxford University Press, New York.

O’Hanley, D.S., and Dyar, M.D., 1993, The composition of lizardite 1T and the formation of magnetite in serpentinites: American Mineralogist, v. 78, p. 391-404.

Oufi, O., and Cannat, M., 2002, Magnetic properties of variably serpentinized abyssal peridotites: Journal of Physical Research, v. 107, no. B5, p.209-2113.

Page, N. J., 1967. Serpentinization at Burro Mountain, California: Contributions to Mineralogy and Petrology, v. 14, p. 321-342.

Radhakrishna, T., Rao, D., Rao J., 1982, Occurrence and significance of awaruite in Dras ultramafics, Kashmir Himalaya India: Mineralogical Magazine, v. 46, p.483-484.

Roeder and Emslie, 1970, Olivine-liquid equilibrium: Contributions to Mineralogy and Petrology, v.29, no.4, p. 275-289.

Royal Nickel Corporation, 2013, Technical report on the Dumont Ni Project, Launay and Trecesson Townships, Quebec, Canada.( http://www.sedar.com/)

Rucklidge, J.C., 1972, Chlorine in partially serpentinized dunite: ECONOMIC GEOLOGY, v. 67, p. 38-40.

Rucklidge, J.C., and Patterson, G.C., 1977, The role of chlorine in serpentinization: Contributions to Mineralogy and Petrology, v. 65, p. 39-44.

Seyfried, W.E., Foustoukos, D.I., and Fu, Q., 2007, Redox evolution and mass transfer during serpentinization: An experimental and theoretical study at 200°C, 500 bar with implications for ultramafic-hosted hydrothermal systems at Mid-Ocean Ridges: Geochimica et Cosmochimica Acta, v.71, p. 3872-3886.

Sciortino, M., Muinonen, J. Korczak, J. and St-Jean, A. Geometallurgical Modelling of the Dumont Deposit. AusIMM International Geometallurgy Conference, 2nd, Brisbane, Sept 30-Oct 2nd, p.93-100.

SGS, Exploration Mineralogy: ExplominTM 2013. (http://www.sgs.com/en/Mining/Exploration- Services/Geometallurgy/EXPLOMIN.aspx)

Shervais, J.W., Kolesar, P., and Andreasen, K., 2005, A field and chemical study of serpentinization Stonyford, California: chemical flux and mass balance: International Geology Review, v. 47, p. 1-28.

Sleep, N.H., Meibom, A., Fridriksson, T.H, Coleman, R.G., and Bird, D.K., 2004, H2-rich fluids from serpentinization: geochemical and biotic implications: Proceedings of the National Academy of Science, v. 101, p. 12818–12823.

Sobolev, A., Hoffman, A., Kuzmin, D., Yaxley, G., Arndt, N., Chung, S.L., Danyushevsky, L., Elliott, T., Frey, F., Garcia, M., Gurenko, A., Kamenetsky, V., Kerr, A., Krivolutskaya, N., Matvienkov, V., Nikogosian, I., Rocholl, A., Sigurdson, I., Sushchevskaya, N., and Teklay, M., 2007, The amount of recycled crust in sources of mantle-derived melts: Science, v. 316, p. 412−417.

Sproule, R.A., Lesher, C.M., Houle, C.M., and Keays, R.R., 2005, Chalcophile element geochmistry and metallogenisis of komatiitic Rocks in the Abitibi Greenstone Belt, Canada. ECONOMIC GEOLOGY, v. 100 p. 1169-1190.

Thompson, J.F.H., Barnes, S.J., and Duke, J.M., 1984, The distribution of nickel and iron between olivine and the magmatic sulfides in some natural assemblages: Canadian Mineralogist, v. 22, p. 55-66.

Toft, P.B, Arkani-Hamed, J., and Haggerty, S.E., 1990, The effects of serpentinization on density and magnetic susceptibility: petrophysical model: Physics of the Earth and Planetary Interiors, v. 65, p. 137-157.

Wicks, F. J. and Plant, A. G. ,1979,. Electron microprobe and X-ray microbeam studies of serpentine textures. Canadian Mineralogist v. 17, p.785-830.

Wood, B.J., 1990, An experimental test of the spinel peridotite oxygen barometer: Journal of Geophysical Research, v. 95b, p. 15845-15851.

Appendix A: Methods Additional Information

A-1: Electron Micro Probe Measurement Conditions by Lab Lab/Device: McGill/JEOL 8900 Beam Correcti Curren Beam on Mineral Acc. Voltage t Size Method Si Mg Fe Cr Ni Al Mn Ti Co Zn V Ca Na P K Cu As S Count Time 20 20 20 20 20 20 20 20 20 20 20 20 20 20 20 Standar Kya Spessa Ruti Met Wille Vana Diop Alb Apa Ortho ds nite CR HM CR NiO CR rtine le al mite dinite side ite tite clase Detectio n Limits 0.03 0.02 0.02 0.03 0.06 0.03 0.02 0.0 0.03 Magnetite 20 kV 30 nA 3 um ZAF (wt%) 8 6 0.040 7 0.029 0 0.027 3 7 0.043 0.048 4 25 0 0.019 Count Time 20 20 20 60 20 20 20 20 20 Standar Pyro Spessa Ruti Met Diop Alb ds OL OL OL CR NiO pe rtine le al side ite Detectio n Limits 0.02 0.02 0.02 0.01 0.02 0.02 0.0 Serpentine 20 kV 30 nA 10 um ZAF (wt%) 2 2 0.020 9 0.017 8 0.026 6 2 0.017 18 Count 20.0 20. Time 20 20 20 20 20.000 00 000 Co Co Standar NiA NiA ds PN PN s SPH CuS s PN Detectio n Limits 0.02 0.05 0.0 Pentlandite 20 kV 30 nA 3 um PRZ (wt%) 0.021 0.039 9 0.040 0.033 3 19 Count 20.0 20. Time 20 20 20 20 20.000 00 000 Co Co Standar NiA NiA ds PN Metal s SPH CuS s PN Detectio Heazlewood n Limits 0.03 0.05 0.0 ite 20 kV 30 nA 3 um PRZ (wt%) 0.021 0.040 1 0.040 0.034 2 18 Count 20.0 20. Time 20 20 20 20 20.000 00 000 Standar Co Co Awaruite 20 kV 30 nA 3 um PRZ ds PN Metal NiA SPH CuS NiA PN

s s Detectio n Limits 0.03 0.05 0.0 (wt %) 0.024 0.044 5 0.040 0.037 3 17 Lab/Device: University of Toronto /Cameca SX-50 Count Time 40 20 20 60 40 60 30 40 Co CR_ Busta TiO NiA Standards Fo93 MT PS97 PN mite 2 s SPH Detection Limits (wt 0.01 0.02 0.02 0.02 0.03 Magnetite 20 kV 35 nA 1 um ZAF %) 5 7 0.040 8 0.038 0.020 4 2 0.039 Count Time 20 30 30 30 30 30 30 20 kaer kaer kaerss kaer pxA kaers Standards ssx1 ssx1 x1 CR PN ssx1 lsx1 sx1 Detection Limits (wt 0.02 0.01 0.03 0.01 0.04 0.02 Serpentine 15kV 35 nA 15 ZAF %) 4 8 0.038 8 0.049 5 0 7 Count Time 20 20 30 30 40 40 30 Co Co NiA Chalco NiA Standards PN PN s SPH pyrite s PN Detection Limits (wt 0.02 0.08 0.0 Pentlandite 25kv 20nA 1 ZAF %) 0.027 5 0.048 0.021 0 28 Count Time 20 20 40 40 40 40 SRM- INVA Ni-Cr Cu Standards R Alloy SX1 Brass Brass S Detection Limits (wt 0.03 0.0 Awaruite 20 kV 20nA 1 um ZAF %) 0.040 0.076 1 0.041 0.061 23 Count Time 40 40 20 20 40 40 40 40 40 40 Co Fo9 Busta Ruti NiA Vana Standards 3 CR CR CR PN CR mite le s SPH dinite Detection Limits (wt 0.00 0.00 0.02 0.00 0.01 0.01 Chromite 20 kV 50nA 1um ZAF %) 7 9 0.017 6 0.022 8 0.017 4 8 0.019 0.010 Count Olivine 20 kV 30nA 1um ZAF Time 20 30 30 40 40 40 40

Fo8 pxTi Standards 5 Fo85 Fo85 CR PN Al TiO2 Detection Limits (wt 0.01 0.00 0.01 0.00 0.01 %) 1 9 0.015 6 0.013 9 8 Lab/Device: Xstrata Processing Support: Cameca SX-100 Count Time 20 20 20 20 20 20 20 20 20 20 20 20 20 20 20 Kya Spessa Ruti Met Wille Vana Diop Alb Apa Ortho Standards nite CR HM CR NiO CR rtine le al mite dinite side ite tite clase Detection ZAF- Limits (wt 0.01 0.02 0.01 0.01 0.01 0.02 0.01 0.0 0.01 Magnetite 20 kV 30 nA 0 um PAP %) 3 4 0.027 4 0.017 4 0.023 4 8 0.047 0.006 3 26 5 0.015 Count Time 20 20 20 20 20 20 20 Co Co NiA NiA Standards PN Metal s SPH CuS s PN Detection ZAF- Limits (wt 0.02 0.0 Pentlandite 20 kV 30 nA 0 um PAP %) 0.022 0.028 3 0.034 20 Count Time 20 20 20 20 20 20 20 Co Co NiA NiA Standards PN Metal s SPH CuS s PN Detection ZAF- Limits (wt 0.02 0.0 Awaruite 20 kV 30 nA 0 um PAP %) 0.022 0.028 3 0.034 20 Count Time 20 20 20 20 20 20 20 Co Co NiA NiA Standards PN Metal s SPH CuS s PN Detection ZAF- Limits (wt 0.02 0.0 Heazlewoodite 20 kV 30 nA 0 um PAP %) 0.022 0.028 3 0.034 20 CR= Chromite, HM=Hematite, OL=Olivine, MT=magnetite, PN=Pentlandite, CuS=Chalcopyrite

Appendix A: Figures Figure A-1: Explomin Data Density and Distribution for Royal Nickel Corporation’s 2013 Technical report.(Figure provided by Royal Nickel Corporation)

Figure A-2: Drill hole data density and distribution for Royal Nickel Corporation’s 2013 Technical report. (Figure provided by Royal Nickel Corporation).

Figure A-3: Electron microprobe data density and distribution for Royal Nickel Corporation’s 2013 Technical report. (Figure provided by Royal Nickel Corporation)

Figure A-4. a,b,c) Distribution of Ni contents in pentlandite, awaruite, and magnetite, respectively as measured by electron microprobe. Average values for individual samples are displayed within the Feasibility pit outline.

Figure A-5. plot of fO2 vs T for the olivine spinel pairs.

-12 600 700 800 900 1000 1100 1200 1300 -13

-14

-15

-16 fO2

-17

-18

-19

-20 T (C)

Figure A-6. Frequency Distribution of Aw for low-sulfide samples. While the mean values are similar the upper dunite (UDZ) displays populations of samples which contain higher amounts of awaruite.”LDZ”: lower dunite

LDZ Partially LDZ Fe-Serpentine Serpentinized Facies Facies 20 100% 90% 20 100% 90% 80% 18 15 16 80% 70% 14 70% 60% 12 60% 10 50% 10 50% 40% 8 40% 30% 6 30% 5 20% 4 20% 10% 10% 2 0 0%

0 0%

0.3 0.1 0.2 0.4

0.25 0.05 0.15 0.35

0.1 0.2 0.3 0.4

0.025 0.075 0.125 0.175 0.225 0.275 0.325 0.375

0.05 0.15 0.25 0.35

0.075 0.125 0.175 0.225 0.275 0.325 0.375 0.025 Aw% Aw%

LDZ Mg Serpentine UDZ Partially Facies Serpentinized 25 100% 14 100% 90% 12 90% 20 80% 80% 70% 10 70% 15 60% 8 60% 50% 50% 10 40% 6 40% 30% 4 30% 20% 5 20% 2 10% 10%

0 0% 0 0%

0.1 0.2 0.3 0.4

0.25 0.05 0.15 0.35

0.05 0.15 0.25 0.35

0.025 0.075 0.125 0.175 0.225 0.275 0.325 0.375

0.025 0.075 0.125 0.175 0.225 0.275 0.325 0.375 Aw% Aw %

UDZ Fe-Serpentine UDZ Mg- Serpentine Facies Facies

35 100% 40 100% 30 90% 35 90% 80% 80% 25 70% 30 70% 20 60% 25 60% 50% 20 50% 15 40% 40% 30% 15 10 30% 20% 10 20% 5 10% 5 10% 0 0%

0 0%

0.3 0.1 0.2 0.4

0.05 0.15 0.25 0.35

0.1 0.2 0.3 0.4

0.175 0.025 0.075 0.125 0.225 0.275 0.325 0.375

0.05 0.15 0.25 0.35

0.175 0.025 0.075 0.125 0.225 0.275 0.325 0.375 Aw% Aw%

Appendix B: Reported Explomin Mineralogy and Electron Microprobe Data for Figure 6 Avg Ni ICP-41 Assay ExplominTM Field Stitch Mineralogy (wt %) Ni wt% in Phase by EMP wt% of sulfide SAMPLE Ni (ppm) S % SG SP OL MT BRU CG CR PN AW HZ ML SPFE OL Fo SP MT PN HZ AW and Aw1 A EXP_394 4090 0.37 2.97 49.46 42.52 3.44 0.82 0.09 0.03 3.1 0 0 0 14.6 0.25 0.91 0.27 0.07 30.6 n/a n/a 30.6 B EXP_275*2 5630 0.28 2.66 68.78 23.88 1.42 1.26 0.32 0.1 3.31 0.76 0.02 0 28.36 0.38 0.89 0.28 0.07 30.3 n/a 74.9 38.4 RNC- C 4240 0.2 2.46 86.85 2.62 1.34 3.86 0.87 2.15 1.88 0.19 0 0 41.79 N/A N/A 0.19 0.02 28.2 n/a 72.9 216A-EXP 32.3 RNC- D 8380 0.83 2.39 90.79 0.05 1.02 5.79 0.29 0.15 1.7 0.16 0.01 0 18.75 N/A N/A 0.1 0.01 28.2 n/a 72.4 197D-EXP 32.2 E EXP_052 2600 0.1 2.56 85.66 0 6.37 6.34 0.15 0.02 1.21 0.16 0.06 0 1.07 N/A N/A N/A 0.03 33.1 n/a n/a 39.2 F EXP_159 4580 0.3 2.59 94.96 0 2.56 1.48 0.15 0.07 0.6 0 0.15 0 0.41 N/A N/A 0.1 0.01 32.6 71.8 n/a 40.4 G EXP_531 2660 0.05 2.61 90.68 0 4.64 3.67 0.21 0.1 0.02 0.07 0.57 0 0.49 N/A N/A 0.19 0.04 n/a 72.7 n/a 71.6 H EXP_187 2670 0.05 2.61 90.12 0.15 7.16 1.45 0.11 0.04 0.02 0.03 0.85 0.28 2.31 N/A N/A N/A N/A n/a n/a n/a 70.2 I EXP_650 2820 0.04 3.19 29.92 68.73 0.91 0.04 0 0.09 0.02 0.01 0.06 0 6.06 N/A N/A N/A N/A n/a n/a n/a J EXP_397 2990 0.02 2.8 73.47 24.02 0.95 0.43 0.1 0.2 0.05 0.2 0 0 31.17 0.35 0.9 0.32 0.09 n/a n/a n/a K EXP_164 3140 0.02 2.6 96.57 0 1.38 1.32 0.09 0.37 0.02 0.18 0 0 22.43 N/A N/A 0.21 0.01 n/a n/a 73.4 L EXP_383 2870 0.1 2.61 96.15 0.04 2.21 1.1 0.18 0 0 0.28 0 0 2.38 N/A N/A 0.22 0 n/a n/a 72.6 M EXP_150 3380 0.03 2.59 95.4 0 1.82 2.08 0.18 0.28 0 0.08 0.11 0 7.39 N/A N/A 0.16 0.65 n/a 72.3 88.4 RNC- N 2630 0.06 2.53 92.39 0.31 2.28 4.55 0.11 0.02 0.02 0.04 0.1 0 0.3 N/A N/A N/A 0.03 n/a 73.3 n/a 224A-EXP Specific Gravity, SP=total serpentine, OL=olivine, MT=magnetite, BRU=brucite, CG=Coalingite (indicator for Fe bearing hydroxide phase), CR=chromite, PN=pentlandite, Aw=awaruite, HZ=Heazlewoodite, SPFE=Iron Serpentine, ML=Millerite

1 Calculated based on reported mineralogy where the average ni content in metallic nickel phases (Aw, Pn, Hz ) = (modal abundance Aw * Nickel in Aw + modal abundance Pn * Ni in Pn+ Modal Abundance Hz*Nickel in Hz)/sum modal abundance (Hz, Aw, Pn). Millerite was included where applicable. Where probe values were not measured for Aw, Hz or Ml, the average of all probe points was used.