generation of high-ni sulfide and alloy phases during ... · ii generation of high-ni sulfide and...
TRANSCRIPT
Generation of high-Ni sulfide and alloy phases during serpentinization of the Dumont Dunite,
Quebec
By
MICHELLE SCIORTINO
A Thesis submitted in conformity with the requirements for the degree of MASc Graduate Department of
Geology, University of Toronto
© Copyright Michelle Sciortino 2014
ii
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
iii
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.
iv
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
v
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
vi
Table of Abbreviations
aH2O activity of H2O
aSiO2 activity of SiO2
AW Awaruite
BRU Brucite
CNisulfide
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
vii
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).
Figure 18: ΣS2 vs fO2 for Ni assemblages of interest at 300oC 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
viii
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 ExplominTM particle scan 54
1
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 Fe
3+. The result is the
production of a ferric iron bearing phase, most commonly magnetite [(Fe23+
Fe2+
)O4] or ferric-
iron bearing serpentine (cronstedtite [(Fe2+2
Fe3+
)(Fe3+
Si)O5(OH)4]). The oxidation of Fe2+
in
olivine to Fe3+
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 Fe
3+ occurs in the early stage of serpentinization by via the
2
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.
3
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
4
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
5
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.
To determine the proportions of whole-rock FeO vs Fe2O3, 20 ExplominTM
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
6
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
7
this distinction is purely based on composition; the phase recognized as coalingite
(Mg10Fe+3
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
8
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).
9
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,
10
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).
11
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
12
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
13
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
14
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
15
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.
16
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.
17
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.
18
Table 1: Phase assemblages and textures in serpentinized dunite at Dumont.
Silicate Matrix Other Properties
Serpentinizatio
n State Olivine Serpentine Brucite Magnetite Other Ni Phases
Proportion
of Dunite*
Density
(kg*m-3
)
Magnetic
Susceptibility
Incipient
Present
slightly
fractured
(OL>20%)
Dominantly
Fe SP
trace, spots
within grains or
along
boundaries
trace, spots
within grains
or along
boundaries
4 >2800 <40
Partial
Some to
none
(20%<OL>
5%)
Dominantly
FESP,
variable
MGSP
some both Fe
and Mg as
minor rims and
centers
minor spots
within grains
or along
boundaries
small
grains<5µm of
PN or/and AW
within
pseudomorphed
grain
26 2600-2800 <40
Complete Fe-
Serpentine
Facies
None
(OL<5%,
FESP>5%)
Variable Fe
and MGSP
weak brucite
rings around
pseudomorphed
grains, variable
Fe and mg
minor spots
within grains
or along
boundaries
small
grains<5um of
PN or/and AW
within
pseudomorphed
grain
29 2300-2600 10-25
Complete Mg
Serpentine
Facies
None
(OL<5%,
FESP<5%)
MGSP
Mg brucite rings
around
pseudomorphed
olivine grains
within brucite
rings
surrounding
pseudomorphe
d olivine
grains
trace secondary
Aw or Hz grains
near edges of
pseumorhped
grains.
41 2600-2800 40-80
*Dunite inside Feasibility pit shell and within mineral wireframe. “FESP=Iron Serpentine”, “MGSP=Magnesium Serpentine, “OL”=olivine”
19
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
20
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.
21
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)
22
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
23
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).
24
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).
25
Table 8: Comparison of Ni content for phases of interest in sulfide-bearing and sulfide-poor
samples.
AW PN HZ SP MT
Low-
sulfide
Incipient and
partially
serpentinized
73.1(7) N/A N/A 0.18(6) 0.11(7)
Low-
sulfide
Completely
serpentinized Fe-
serpentine facies
74.4(16) N/A N/A 0.2(12) 0.11(16)
Low-
sulfide-
Completely
serpentinized Mg-
serpentine facies
74.2(24) N/A N/A 0.1(6) 0.14(27)
sufide-
bearing
Incipient and
partially
serpentinized
73.2(4) 29.3(8) 72.3(1) 0.23(6)
sulfide-
bearing
Completely
serpentinized Fe-
serpentine facies
72.8(13) 29.0(18) 71.9(7) 0.14(13) 0.02(18)
sulfide-
bearing
Completely
serpentinized Mg-
serpentine facies
HZ/PN>1
72.8(8) 33.2(14) 72.5(17) 0.1(1) 0.06(17)
sulifide-
bearing
Completely
serpentinized Mg-
serpentine facies
HZ/PN<1
72.5(17) 32.7(15) 71.3(13) 0.04(2) 0.02(15)
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).
26
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
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.
86
88
90
92
94
96
98
100
0 0.1 0.2 0.3 0.4 0.5 0.6
Mg
#
Ni wt %
27
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.
28
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.
29
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
500
5000
1000 10000
LOG
[Cr
(pp
m)
XR
F]
LOG[Ni (ppm) XRF]
Non-Mineralized
Low Sulfide
Sulphide-Pn Dominant
Sulphide-Hz Dominant
30
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
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.
-4.00
-3.00
-2.00
-1.00
0.00
1.00
88 89 90 91 92 93
Δ F
MQ
Mg # Olivine
EXP_041
EXP_397
EXP_438
271-EXP008
EXP_128
EXP_275
31
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
on fS2, fO2, and the Ni content of the sulfide melt CNisul
:
8.70log0.11312log7.342
2
2
2
fS
fO
fS
fOCK sulfide
NiD 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).
In the present formulation KD has a linear dependence on CNisul
rather than the (CNisul
)3
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)
cubic dependence of KD on CNisul
leads to unphysical results like CNisul
> 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
32
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).
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.
20
30
40
50
20 30 40 50
Ni (
wt%
, KD)
Ni (wt%, ExplominTM)
33
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
34
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-
40
50
60
70
80
90
100
0 20 40 60
SP (
TOTA
L)
OL
Fe-SerpentineFaciesMg-SerpentineFaciesPartiallySerpentinized
A
50
60
70
80
90
100
10 15 20 25
MG
SP
LOI %
B
0
10
20
30
40
50
10 15 20 25
FE S
P
LOI
C
0
2
4
6
8
10
10 15 20 25
MT
LOI
D
0
10
20
30
40
50
0 2 4 6 8 10
FE S
P
MT
E
0
0.5
1
1.5
2
0 2 4 6 8 10
CG
MT
F
35
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).
50
60
70
80
90
100
0 2 4 6 8 10
SP (
TOTA
L)
MT
G
0
2
4
6
8
10
12
14
0 2 4 6 8 10
Fe2
O3
MT
H
0
1
2
3
4
5
6
7
8
10 15 20 25
BR
U
LOI
0
1
2
3
4
5
6
7
8
80 85 90 95 100
BR
U
SP TOTAL
I J
36
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
37
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
0.1
0.2
0.3
0.4
0.5
0.6
50 60 70 80 90 100
NI %
SP (total)
A
4
6
8
10
12
50 60 70 80 90 100
Fe2
O3
%
Sp (total)
42
44
46
48
50
52
50 60 70 80 90 100
MgO
%
SP(total)
C
38
39
40
41
42
43
44
50 60 70 80 90 100
SiO
2%
SP (total)
D
B
38
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%
Ni
(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%
Ni
(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%.
0
0.2
0.4
0.6
0.8
1
50 60 70 80 90 100
CaO
%
SP (total)
E
2.2
2.4
2.6
2.8
3
50 60 70 80 90 100
SG
SP (total)
F
39
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+
40
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,
41
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.
42
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
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
H2O bound up in the rock because oxidation of Fe2+
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.
y = -0.89x + 2.0714 R² = 0.785
y = -0.7047x + 1.9995 R² = 0.7683
1.0
1.2
1.4
1.6
1.8
2.0
2.2
2.4
0 0.05 0.1 0.15 0.2 0.25 0.3
IV S
i+A
l
IV Fe apfu
A
0%
10%
20%
30%
40%
50%
60%
70%
80%
5 10 15 20
Fe3
+ %
of
tota
l FE
LOI %
B
20%
30%
40%
50%
60%
70%
80%
90%
100%
20% 45% 70% 95%
Fe3
+ % S
P
Fe3+ % whole rock
C
43
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.
44
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
45
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
46
(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
of the excess Fe2+
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:
4Fe2SiO4 + 5H2O + 3Mg(OH)2 = 2(Fe22+
Fe3+
)(Fe3+
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
47
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).
The substitution of Fe3+
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
ability for Fe3+
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
composition, no longer requiring the oxidation of Fe2+
to Fe3+
, and as a result no H2 is produced
(McCollum and Bach, 2009; Evans at al., 2013). Conversely during serpentinization at very low
temperatures 100-200°C, Fe2+
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).
48
Where conditions are such that the oxidation of Fe2+
to Fe3+
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.
49
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)
3(Fe22+
Fe3+
)(Fe3+
Si)O5(OH)4 = 4Fe3O4 + 3SiO2 + 5H2O + H2 (17)
3Fe2+
(OH)2 + 0.5O2 = (Fe+3
2Fe+2
)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
50
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
Figure 18: ΣS2 vs fO2 for Ni assemblages of interest at 300oC 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
51
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
serpentine structure is the result of a decrease in aSiO2 and Fe3+
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
52
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
serpentine) may be limited by further H2 production by the oxidation of any remaining Fe2+
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,
53
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).
54
Table 12: Average modal abundance for minerals of interest as reported by ExplominTM
particle scan
Upper Dunite AW PN HZ MT BRU NISI NI ppm
(XRF)
S %
(ICP-41) SP OL N
Low sulfide
Partially serpentinized 0.08 0.04 0.03 1.45 1.89 68.4 2869 0.019 80.7 14.7 51
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
sulfide
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
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
Lower dunite AW PN HZ MT BRU NISI NI ppm
(XRF)
S%
(ICP-41) SP OL N
Low-sulfide
Partially serpentinized 0.06 0.04 0 2.62 1.04 72.9 2334 0.0197 85 18.6 43
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
sulfide
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
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 %,
55
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-
serpentine partitioning of Fe is likely to govern the fO2 via the oxidation of Fe2+
into Fe3+
.
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.
56
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
magnetite was directly related to the H2 production and that the incorporation of Fe3+
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
57
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).
58
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,1<HZ/PN<5, SPFE<14. Zone 3a (Fig. 6 C-D); low –moderate iron serpentine, pentlandite dominate
commonly with coarse awaruite. Zone 3b (Fig. 6 D-E); low iron serpentine, pentlandite dominate +/- awaruite
(metallurgical domain: Pentlandite Dominant HZPN<=1, SPFE<14. (4)A&B pentlandite dominate with high iron
serpentine +/- relict olivine (Metallurgical domain high iron serpentine SPFE>14. 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).
59
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.
60
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
rock to impose on the fluid its values of the chemical potentials of H2O, H2, SiO2 and Fe2+
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
61
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
resulting release of Fe2+
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
including Fe-Ni alloys as a result of the H2 production via the oxidation of Fe2+
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.
62
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.
63
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.
64
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
Evans, B.W., 2008, Control of the products of serpentinization by the Fe2+
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.
65
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.
66
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.
67
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.
68
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.
69
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.
70
Appendix A: Methods Additional Information
A-1: Electron Micro Probe Measurement Conditions by Lab
Lab/Device: McGill/JEOL 8900
Mineral Acc. Voltage
Beam
Curren
t
Beam
Size
Correcti
on
Method Si
Mg Fe Cr Ni Al Mn Ti
Co Zn V Ca
Na P K Cu
As S
Magnetite 20 kV 30 nA 3 um ZAF
Count
Time 20 20 20 20 20 20 20 20 20 20 20 20 20 20 20
Standar
ds
Kya
nite CR HM CR NiO CR
Spessa
rtine
Ruti
le
Met
al
Wille
mite
Vana
dinite
Diop
side
Alb
ite
Apa
tite
Ortho
clase
Detectio
n Limits
(wt%)
0.03
8
0.02
6 0.040
0.02
7 0.029
0.03
0 0.027
0.06
3
0.03
7 0.043 0.048
0.02
4
0.0
25
0.03
0 0.019
Serpentine 20 kV 30 nA 10 um ZAF
Count
Time 20 20 20 60 20 20 20 20 20
Standar
ds OL OL OL CR NiO
Pyro
pe
Spessa
rtine
Ruti
le
Met
al
Diop
side
Alb
ite
Detectio
n Limits
(wt%)
0.02
2
0.02
2 0.020
0.02
9 0.017
0.01
8 0.026
0.02
6
0.02
2 0.017
0.0
18
Pentlandite 20 kV 30 nA 3 um PRZ
Count
Time 20 20 20 20 20.000
20.0
00
20.
000
Standar
ds PN PN
Co
NiA
s SPH CuS
Co
NiA
s PN
Detectio
n Limits
(wt%) 0.021 0.039
0.02
9 0.040 0.033
0.05
3
0.0
19
Heazlewood
ite 20 kV 30 nA 3 um PRZ
Count
Time 20 20 20 20 20.000
20.0
00
20.
000
Standar
ds PN Metal
Co
NiA
s SPH CuS
Co
NiA
s PN
Detectio
n Limits
(wt%) 0.021 0.040
0.03
1 0.040 0.034
0.05
2
0.0
18
Awaruite 20 kV 30 nA 3 um PRZ
Count
Time 20 20 20 20 20.000
20.0
00
20.
000
Standar
ds PN Metal
Co
NiA SPH CuS
Co
NiA PN
71
s s
Detectio
n Limits
(wt %) 0.024 0.044
0.03
5 0.040 0.037
0.05
3
0.0
17
Lab/Device: University of Toronto /Cameca SX-50
Magnetite 20 kV 35 nA 1 um ZAF
Count
Time 40 20 20 60 40 60 30 40
Standards Fo93 MT
CR_
PS97 PN
Busta
mite
TiO
2
Co
NiA
s SPH
Detection
Limits (wt
%)
0.01
5
0.02
7 0.040
0.02
8 0.038 0.020
0.02
4
0.03
2 0.039
Serpentine 15kV 35 nA 15 ZAF
Count
Time 20 30 30 30 30 30 30 20
Standards
kaer
ssx1
kaer
ssx1
kaerss
x1 CR PN
kaer
ssx1
pxA
lsx1
kaers
sx1
Detection
Limits (wt
%)
0.02
4
0.01
8 0.038
0.03
8 0.049
0.01
5
0.04
0
0.02
7
Pentlandite 25kv 20nA 1 ZAF
Count
Time 20 20 30 30 40 40 30
Standards PN PN
Co
NiA
s SPH
Chalco
pyrite
Co
NiA
s PN
Detection
Limits (wt
%) 0.027
0.02
5 0.048 0.021
0.08
0
0.0
28
Awaruite 20 kV 20nA 1 um ZAF
Count
Time 20 20 40 40 40 40
Standards
SRM-
INVA
R
Ni-Cr
Alloy SX1 Brass Brass
Cu
S
Detection
Limits (wt
%) 0.040 0.076
0.03
1 0.041 0.061
0.0
23
Chromite 20 kV 50nA 1um ZAF
Count
Time 40 40 20 20 40 40 40 40 40 40
Standards
Fo9
3 CR CR CR PN CR
Busta
mite
Ruti
le
Co
NiA
s SPH
Vana
dinite
Detection
Limits (wt
%)
0.00
7
0.00
9 0.017
0.02
6 0.022
0.00
8 0.017
0.01
4
0.01
8 0.019 0.010
Olivine 20 kV 30nA 1um ZAF
Count
Time 20 30 30 40 40 40 40
72
Standards
Fo8
5 Fo85 Fo85 CR PN
pxTi
Al TiO2
Detection
Limits (wt
%)
0.01
1
0.00
9 0.015
0.01
6 0.013
0.00
9
0.01
8
Lab/Device:
Xstrata
Processing
Support:
Cameca SX-100
Magnetite 20 kV 30 nA 0 um
ZAF-
PAP
Count
Time 20 20 20 20 20 20 20 20 20 20 20 20 20 20 20
Standards
Kya
nite CR HM CR NiO CR
Spessa
rtine
Ruti
le
Met
al
Wille
mite
Vana
dinite
Diop
side
Alb
ite
Apa
tite
Ortho
clase
Detection
Limits (wt
%)
0.01
3
0.02
4 0.027
0.01
4 0.017
0.01
4 0.023
0.01
4
0.02
8 0.047 0.006
0.01
3
0.0
26
0.01
5 0.015
Pentlandite 20 kV 30 nA 0 um
ZAF-
PAP
Count
Time 20 20 20 20 20 20 20
Standards PN Metal
Co
NiA
s SPH CuS
Co
NiA
s PN
Detection
Limits (wt
%) 0.022 0.028
0.02
3 0.034
0.0
20
Awaruite 20 kV 30 nA 0 um
ZAF-
PAP
Count
Time 20 20 20 20 20 20 20
Standards PN Metal
Co
NiA
s SPH CuS
Co
NiA
s PN
Detection
Limits (wt
%) 0.022 0.028
0.02
3 0.034
0.0
20
Heazlewoodite 20 kV 30 nA 0 um
ZAF-
PAP
Count
Time 20 20 20 20 20 20 20
Standards PN Metal
Co
NiA
s SPH CuS
Co
NiA
s PN
Detection
Limits (wt
%) 0.022 0.028
0.02
3 0.034
0.0
20
CR= Chromite, HM=Hematite, OL=Olivine, MT=magnetite, PN=Pentlandite, CuS=Chalcopyrite
73
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).
74
Figure A-3: Electron microprobe data density and distribution for Royal Nickel Corporation’s 2013
Technical report. (Figure provided by Royal Nickel Corporation)
75
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.
76
Figure A-5. plot of fO2 vs T for the olivine spinel pairs.
-20
-19
-18
-17
-16
-15
-14
-13
-12
600 700 800 900 1000 1100 1200 1300
fO2
T (C)
77
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
0%
10%
20%
30%
40%
50%
60%
70%
80%
90%
100%
0
5
10
15
20
0.0
25
0.0
50
.07
50
.10
.12
50
.15
0.1
75
0.2
0.2
25
0.2
50
.27
50
.30
.32
50
.35
0.3
75
0.4
Mo
re
Aw%
LDZ Partially Serpentinized Facies
0%
10%
20%
30%
40%
50%
60%
70%
80%
90%
100%
02468
101214161820
0.0
25
0.0
50
.07
50
.10
.12
50
.15
0.1
75
0.2
0.2
25
0.2
50
.27
50
.30
.32
50
.35
0.3
75
0.4
Mo
re
Aw%
LDZ Fe-Serpentine Facies
0%
10%
20%
30%
40%
50%
60%
70%
80%
90%
100%
0
5
10
15
20
25
0.0
25
0.0
50
.07
50
.10
.12
50
.15
0.1
75
0.2
0.2
25
0.2
50
.27
50
.30
.32
50
.35
0.3
75
0.4
Mo
re
Aw%
LDZ Mg Serpentine Facies
0%
10%
20%
30%
40%
50%
60%
70%
80%
90%
100%
0
2
4
6
8
10
12
14
0.0
25
0.0
50
.07
50
.10
.12
50
.15
0.1
75
0.2
0.2
25
0.2
50
.27
50
.30
.32
50
.35
0.3
75
0.4
Mo
re
Aw %
UDZ Partially Serpentinized
78
.
0%
10%
20%
30%
40%
50%
60%
70%
80%
90%
100%
0
5
10
15
20
25
30
35
0.0
25
0.0
50
.07
50
.10
.12
50
.15
0.1
75
0.2
0.2
25
0.2
50
.27
50
.30
.32
50
.35
0.3
75
0.4
Mo
re
Aw%
UDZ Fe-Serpentine Facies
0%
10%
20%
30%
40%
50%
60%
70%
80%
90%
100%
0
5
10
15
20
25
30
35
40
0.0
25
0.0
50
.07
50
.10
.12
50
.15
0.1
75
0.2
0.2
25
0.2
50
.27
50
.30
.32
50
.35
0.3
75
0.4
Mo
re
Aw%
UDZ Mg- Serpentine Facies
79
Appendix B: Reported Explomin Mineralogy and Electron Microprobe Data for Figure 6
ICP-41 Assay ExplominTM
Field Stitch Mineralogy (wt %) Ni wt% in Phase by EMP Avg Ni
wt% of
sulfide
and Aw1
SAMPLE Ni (ppm) S % SG SP OL MT BRU CG CR PN AW HZ ML SPFE OL Fo SP MT PN HZ AW
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
C RNC-
216A-EXP 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
32.3
D RNC-
197D-EXP 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
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
N RNC-
224A-EXP 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
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.