The Application of Platinum Group Elements€¦ · prospective to host mineralization. The potential use of the chalcophile elements, specifically the platinum group elements (PGE:

The Application of Platinum Group Elements

in Komatiite-Hosted

Nickel Sulfide Exploration

Geoffrey John Heggie, B.Sc., M.Sc.

This thesis is presented for the degree

of Doctor of Philosophy in Geology

of the University of Western Australia

December 2010.

ii

THE APPLICATION OF PLATINUM GROUP ELEMENTS IN KOMATIITE-HOSTED NICKEL SULFIDE EXPLORATION

Summary

Exploration for komatiite-hosted nickel (Ni) deposits is a continued challenge, due

to small target size, discontinous nature, and lack of an alteration halo associated

with the ore forming process. New komatiite-hosted Nil deposits are becoming

increasingly difficult to locate, as the remaining prospective areas are typically

under cover and at greater depths. Lithogeochemistry has the capacity to increase the

target size beyond the physical mineralization and indicate whether the system is

prospective to host mineralization. The potential use of the chalcophile elements,

specifically the platinum group elements (PGE: iridium, rhodium, ruthenium,

platinum, and palladium), as lithogeochemical prospectivity indictors is widely

recognized, since the chalcophile elements are intimately associated with Ni ore

formation process. The ore formation in komatiites is a consequence of sulfur

saturation and the strong partitioning of the chalcophile elements from the silicate

magma into the sulfide phase. During these processes, two mineralization signatures

develop: chalcophile element enrichment and chalcophile element depletion. As

such, chalcophile elements are used as prospectivity indicators. Previous

applications utilizing the chalcophile elements as prospectivity (mineralization)

indicators were limited, as the size of the ore forming system remained

unconstrained. This limitation prevented the transformation of these prospectivity

indicators into lithogeochemical vectors to Ni mineralization. This research

identified two scale (size) components critical to the development of a Ni

mineralization vector: 1) chalcophile element signature magnitude (relative

enrichment and depletion): and 2) the spatial correlation (distance) between

chalcophile element signatures and known Ni mineralization.

These scale components of mineralized komatiite Ni systems are addressed through

the investigation of three komatiite-hosted Ni sulphide areas. Two areas represent

deposit case studies: Long-Victor mine (Kambalda Dome) and Maggie Hays mine

(Lake Johnston Greenstone Belt), both in Western Australia; and the third area

comprises scattered ultramafic outcrop in northern Finland and Norway (Karelian

Craton). The Long-Victor, Maggie Hays, and the Karelian Craton areas, provide

diversity in: age (2.7, 2.9, 2.0 Ga, respectively), geochemistry (Munro-, Barberton-,

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Karasjok-type komatiites, respectively), and style of mineralizing system (extrusive,

intrusive, poorly constrained, respectively). This diversity enables two scale

components of mineralized komatiite systems to be quantified.

The relative magnitude of chalcophile element signatures allows for the

identification and classification of a residual anomaly from an established

background abundance. Background chalcophile element abundances are calculated

as a function of MgO content of the respective samples. These calculations are a

product of linear regressed best-fit lines, derived from iteratively filtered

geochemical data sets for the individual deposits. This methodology reproduces

previously reported initial liquid chalcophile element abundances, and characterizes

the background chalcophile element content of samples, from >10 to <50 wt% MgO

in sulfur undersaturated conditions. These background values enables the systematic

quantification of residual anomalies within the deposit data sets (mineralization

signatures: positive [enrichment] and negative [depletion]). Chalcophile element

enrichment and depletion signatures are the result of the Ni ore forming process.

Mineralization signatures are not equally apparent for all chalcophile elements.

Enrichment signatures are apparent with most chalcophile elements (Ni, Cu, PGE:

Ir, Rh, Ru, Pt, Pd); however, depletion signatures are primarily identified through

select PGE (Pt, Pd, Rh). Nickel appears relatively insensitive as a depletion

indicator, and several other chalcophile elements (Cu, Ir, Ru) are limited in use.

These limitations are due to several factors: mobility, compatibility in additional

phases, or occurring at too low of an abundance to produce meaningful

quantification with the current analytical techniques.

This research demonstrates that quantified mineralization signatures from select

PGE (depletion and enrichment) exhibit a spatial correlation to known Ni

mineralization. Both case study deposits (Long-Victor and Maggie Hays) are

characterized by enrichment signatures that increase in magnitude with decreasing

distance to mineralization. This geochemical gradient is interpreted as a primary

disseminated mineralization halo. However, this halo is not necessarily visually

mineralized, as samples with < 0.25 wt% sulfur exhibit chalcophile element

enrichment, a result of sulfur loss during alteration and metamorphism.

iv

The spatial correlation between depletion signatures and mineralization differs

within the two case study deposits, a product of contrasting mineralization setting

(extrusive versus intrusive). However, depletion signatures in both deposits reflect

the influence of recharge in the magmatic system.

The Long-Victor Ni deposit is hosted in an extrusive komatiite system where the

mineralization is associated within linear bodies of thickened olivine cumulates, the

product of sustained channelized magma flow. Magma transport in the system is

both linear (within channel) and lateral, due to the development of adjacent flank

facies by channel splays and over-bank magma flooding. Consequently,

mineralization depletion signatures are preserved in the flanking environment and

exhibit decreasing depletion gradients with proximity to the mineralized channel

environment.

The Maggie Hays Ni deposit is hosted within a sub-volcanic intrusion that acted as a

feeder to overlying extrusive komatiites. As such, magma flow within the feeder

system is purely linear. Mineralization hosted within the intrusion is the result of

sulfur saturation induced by the assimilation of a sulfidic stratigraphic unit overlying

the intrusion, forming a point source. Mineralization depletion signatures occur

proximal to the sulfur point source, and exhibit an increasing and subsequent

decreasing magnitude of depletion signatures with decreasing distance to

mineralization.

These two case study Ni deposits provide an ideal environment to identify, quantify

and constrain the spatial correlation between chalcophile element mineralization

signatures and known mineralization. In summary, mineralization signatures in Ni

mineralized systems are observed in approximately half of the sample population.

Of these, 80% are characterized as enriched, and 20% as depleted. The

understanding gained from the Long-Victor and Maggie Hays deposits was applied

to the Karelian Craton (northern Finland and Norway) to assess the practical

application of chalcophile element signatures in complex terranes with sparse

outcrop and limited volcanological interpretation.

The Karelian Craton, comprising both Archean Munro-type and Proterozoic

Karasjok-type komatiitic rocks, was sampled in three locations: the Karasjok belt,

v

the Pulju belt, and the Enontekiö area, with the latter two hosting known Ni

mineralization. Sampling and limited mapping by the author in these areas identified

ultramafic rocks comprising thin flows and unconstrained cumulate bodies. Major

element whole-rock geochemistry was used to further classify the ultramafic rocks;

where the majority of the cumulate bodies were classified as lava lakes and ponded

flows, rather than the more prospective dunitic bodies. Chalcophile element

abundances were characterized based on fields defined by Barberton- and Munro-

type systems. Consequently, all three locations within the Karelian Craton exhibit

mineralization signatures and have high prospectivity. The Karasjok Belt, despite

being dominated by low prospectivity thin and pillowed flows, contains

mineralization signatures and warrants further research targeting higher volume flow

conduits. The high prospectivity indicators for sample locations within Pulju Belt,

and Enontekiö area, both known to host Ni mineralization, validate the application

of lithogeochemistry and chalcophile element mineralization signatures.

Lithogeochemistry, is a vital tool for Ni sulfide mineralization targeting. The

combination of major and chalcophile elements provides a number of Ni

prospectivity and mineralization vectoring tools. Major elements allow for the

interpretation and discrimination of volcanic facies; whereas, the chalcophile

elements are associated with mineralization. Chalcophile elements, specifically the

platinum group elements (PGE) exhibit quantifiable mineralization signatures and

spatial correlations to known mineralization, resulting in practical and applicable

PGE based vectors to target komatiite hosted Ni sulfide mineralization.

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Table of Contents

Page

Title Page i

Summary iii

Table of Contents vii

List of Figures xv

List of Tables xxxi

Acknowledgments xxxv

1.0 Purpose and Scope: The Application of Platinum Group Elements in Komatiite-Hosted Nickel Sulfide Exploration

1.1. Introduction 1

1.2. World Nickel Use and Discovery 1

1.3. Nickel Prospectivity and Purpose 2

1.4. Chalcophile Element Mineralization Signatures in Komatiites 3

1.5. Research Scope 4

1.6. Thesis Overview 5

1.7. References 9

2.0 Komatiites and Orthomagmatic Nickel

2.1. Introduction 11

2.2. Komatiite Geochemistry and Volcanic Processes 11

a. Classification 12

b. Geochemistry 12

i. Melt generation 13

ii. Chalcophile elements 14

iii. Crystallization 16

iv. Contamination 17

c. Transport and eruption 18

d. Volcanic textures 19

i. Spinifex 19

ii. Cumulates 22

iii. Harrisite 26

iv. Breccia-volcaniclastic 26

v. Vesicles 28

e. Volcanic flow field 28

vii

i. Propagation and field development 29

ii. Flow thickness 32

iii. Channel and Trough 33

iv. Flank 35

v. Scale 36

2.3. Orthomagmatic Mineralization Model 37

a. Sulfur in orthomagmatic nickel systems 39

b. Nickel sulfide distribution 40

c. Metal tenor and distribution in sulfide ores 41

2.4. Mineralization Indicators 43

a. Major elements - whole rock geochemistry 43

b. Trace elements - whole rock geochemistry 45

c. Mineralization 46

d. Chalcophile elements - whole rock geochemistry 47

i. Chalcophile element partitioning 48

ii. R-factor 48

iii. Chalcophile element mineralization signatures 50

iv. Examples of chalcophile element signatures 52

e. Minerals and mineral separates 54

f. Spatial distribution and size of mineralized systems 55

2.5. Conclusion, Implications and the Way Forward 56

a. Komatiite generation 57

b. Tectonic setting 57

c. Mineralization processes 57

d. Considerations 58

2.6. References 59

3. 0 The Kambalda Dome


3.2. Regional Geology and Tectonics 72

a. Stratigraphic sequences 74

i. Lower Kambalda sequence 74

ii. Middle Kalgoorlie sequence 74

iii. Upper Kurrawang and Merougil sequences 75

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b. Geodynamic setting of the Kambalda Domain 75

3.3. Lower Kambalda Sequence Stratigraphy 76

a. Basement 76

b. Lunnon Basalt Formation 77

c. Metasedimentary rocks 79

i. Sediment provenance 80

d. Kambalda Komatiite Formation 81

i. Silver Lake Member 81

ii. Tripod Hill Member 89

e. Devon Consuls Basalt, Kapai Slates, and Paringa Basalt Formations 90

i. Devon Consols Basalt Formation 90

ii. Kapai Slate Formation 91

iii. Paringa Basalt Formation 91

f. Intrusions 92

3.4. Structural Evolution 93

3.5. Alteration and Metamorphism 95

3.6. Summary 99

3.7. References 101

4. 0 The Size of Nickel Mineralized Systems: Examination of Platinum Group Element Distribution in the Long-Victor system, Kambalda Dome, W.A.


4.2. Kambalda Dome 113

a. Geological setting 113

b. Structural modification 116

4.3. Chalcophile Element Abundance 117

4.4. Materials and Methods 118

a. Sample selection 118

b. Distance to mineralization 120

c. Analytical techniques 121

4.5. Results 122

a. Major and trace element geochemistry 123

b. Chalcophile element geochemistry 127

i. Sulfur-bearing 130

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ii. Sulfur-poor 131

4.6. Discussion 132

a. Flow field 132

b. Chalcophile element abundance 136

i. Background chalcophile element values 136

ii. Chalcophile element enrichment 138

iii. Chalcophile element depletion 142

c. Spatial correlation of chalcophile element values 144

d. Timing of komatiite spinifex growth and relation to ore formation 151

e. Volcanological control on spatial distribution of chalcophile element values 153

4.7. Conclusion 156

4.8. References 159

5. 0 Stratigraphic Control on the Style of Komatiite Emplacement in the 2.9 Ga Lake Johnston Greenstone Belt, Yilgarn Craton, Western Australia.


5.2. Regional Geology 170


5.4. Stratigraphy and Geochemistry 174

a. Felsic volcanic unit 176

i. Interpretation of the felsic volcanic unit 180

b. Transition zone unit 181

i. Interpretation of the TZU 183

c. Banded iron formation unit 184

i. Interpretation of the BIF unit 185

d. Sedimentary unit 186

i. Interpretation of the sedimentary unit 187

e. Ultramafic units 187

i. Interpretation of the WUU and CUU 194

5.5. Discussion 197

a. Structural modification 197

b. Tectonic setting and deposition of the Honman Formation 199

c. Stratigraphic control on emplacement of ultramafic magmas 200

x

5.6. Conclusions 204

5.7. References 206

6. 0 Nickel Mineralization Signatures in an Intrusive Komatiite Sequence: Examination of the Spatial Distribution of PGE in the Maggie Hays Ni system, Lake Johnston Greenstone Belt, Western Australia.


6.2. Geological Setting 217

a. Regional stratigraphy 217

i. Central ultramafic unit 220

ii. Maggie Hays Ni deposit 222

b. Metamorphism and structural modification 224


a. 3D model 225

b. Sample selection 226

c. Analytical techniques 227

6.4. Results 228

a. Major and trace element geochemistry 228

b. Chalcophile element geochemistry 232

6.5. Discussion 236

a. Whole-rock geochemistry 236

i. Western ultramafic unit 236

ii. Central ultramafic unit 236

b. Chalcophile element abundance 237

c. Chalcophile element enrichment 240

i. Sulfide-bearing samples 240

ii. Sulfide-poor samples 241

d. Chalcophile element depletion 242

e. Spatial correlation of ore forming signatures 244

6.6. Genetic Model for Ore Formation and the Spatial Distribution of Ore Forming Signatures 246


6.8. References 257

7. 0 Application of Lithogeochemical Prospectivity for Komatiite-Hosted Nickel Sulfide Mineralization, Northern Finland and Norway.

xi


a. Volcanic facies 262

b. Mineralization indicators 263

c. Test area 263

7.2. Regional Setting 264

a. Central Karelian Craton 264

i. Archean komatiites (2.9-2.7 Ga) 266

ii. Paleoproterozoic komatiites (2.0-1.9 Ga) 266

7.3. Sampling and Physical Volcanology 268

a. Archean komatiites (Enontekiö area) 269

b. Paleoproterozoic komatiites (Pulju and Karasjok Greenstone Belts) 269


7.5. Whole-Rock Geochemistry Results 271

a. Archean komatiites (Enontekiö area) 271

b. Paleoproterozoic komatiites (Karasjok and Pulju Greenstone Belts) 272

7.6. Lithogeochemical Prospectivity Indicators 274

a. Petrogenetic classification and initial chalcophile content 274

b. Volcanic facies 277

c. Chalcophile element mineralization signatures 279


7.8. References 285

Appendix Table 7.1A 288

8. 0 Conclusions: Application of Platinum Group Elements in Komatiite-Hosted Nickel Exploration.


8.2. References 303

Appendix A. Sample locations and Summary Descriptions

Long-Victor, Kambalda Dome, Western Australia A1

Maggie Hays, Lake Johnston Greenstone Belt, Western Australia A7

Karelian Craton, northern Finland and Norway A12

Appendix B. Geochemical Analyses

Long-Victor B1

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Maggie Hays B20

Karelian Craton B44

Appendix C. Data Quality

a. Sampling techniques C1

b. Chemical Analysis C1

c. Error in Data C5

d. Quality Assurance and Control C6

e. References C12

Appendix D. Chalcophile Elements as a Function of MgO.

a. Purpose D1

b. Assumptions D1

c. Procedure D2

d. Results D5

e. References D7

xiii

xiv

List of Figures Page

Figure 2.1. World map showing distribution of major orthomagmatic

deposits, Ni mineralization districts and geographical locations referenced in

this thesis. Komatiite-hosted deposits comprise: Mt. Keith, Perseverance,

Black Swan, and Kambalda deposits of Western Australia; Reliance deposit of

Africa, and Abitibi Greenstone Belt of Canada. Komatiitic basalt-hosted

deposits comprise the Thompson Ni-belt and Raglan Ni-belt of Canada. High

MgO basalt deposits are characterized by Noril’sk-Talnakh of Russia,

Jinchuan deposit of China, and Kabanga deposit of Tanzania. Ferro-picrite is

associated with the Pechenga deposit of Russia. Troctolite is associated with

the Voisey’s Bay deposit of Canada. Meteorite impact related deposits are

characterized with the Sudbury region of Canada. Large layered intrusions,

hosting reef-type platinum group element mineralization, are characterized by

the Stillwater Complex of the United States of America, and Bushveld

Complex of South Africa. The Karelian Craton of Finland and Norway is

included for reference to Karasjok-type komatiites. 15

Figure 2.2. Diagram illustrating fully differentiated komatiite flow with upper

A-zone spinifex and lower B-zone olivine cumulates. Modified from Pyke et

al. (1973) and Arndt et al. (1977). 21

Figure 2.3. Komatiite cooling units matrix with increasing olivine

accumulation on left and increasing differentiation along the bottom axis. UN

= undifferentiated non-cumulate (massive, pillowed or volcaniclastic), DN =

differentiated non-cumulate, UC = undifferentiated cumulate, DC =

differentiated cumulate. Modified from Lesher and Keays (2002). 25

Figure 2.4. Komatiite flow field model as proposed by Hill (2001) showing

the transition from massive sheet flow to channelized flow. Modified from

Arndt et al. (2008). 30

Figure 2.5. Komatiite flow field model as proposed by Hill (2001) showing

lobe development at the advancing front and lateral development. Modified

from Arndt et al. (2008). 31

xv

Figure 2.6. Idealized schematic cross-section showing both channel and flank

facies with associated sediments and Ni-sulfide mineralization as observed at

the Kambalda Dome. Modified from Cowden and Roberts (1990). 34

Figure 2.7. Blind persons and the elephant. Cartoon based on poem by John

Godfrey Saxe (1816-1887). Modified from Yeh and Rousseau (2000). 46

Figure 3.1. Regional map of the Yilgarn Craton showing the South West and

Youanmi Terranes and Eastern Goldfields Superterrane. Kalgoorlie, Kurnalpi

and Burtville Terranes shown, and domains within each terrane shown in red.

Nickel deposits hosted within the Yilgarn Craton shown as red squares.

Modified from Cassidy et al. (2006). 72

Figure 3.2. Stratigraphic column within the Kalgoorlie Terrane, with

lithostratigraphic divisions shown on left. Modified from Lesher and Arndt

(1995); Beresford et al. (2002); Krapez and Hand (2008). Stratigraphy adapted

from Gresham and Loftus-Hills (1981); Cowden and Roberts (1990); Swager

et al. (1992); Krapez (1997). Ages U/Pb SHRIMP from Claoue-Long et al.

(1988); Krapez et al. (2000); Kositcin et al. (2008). 73

Figure 3.3. Block model showing distribution of contact sediments within the

channel and flank facies. Modified from Gresham and Loftus-Hills (1981) and

Stone and Masterman (1998). 82

Figure 3.4. Geological map of the Kambalda Dome area with mineralized Ni

ore shoots projected to surface. Major ore shoots are labeled. Map projection

UTM zone 16 with WGS84 datum. 87

Figure 4.1. Generalized geological map of the Kambalda Dome with nickel

sulfide ore shoots shown in plan projection with major faults and fold axis

shown. Area of the Long-Victor Ni deposit shown by dashed outline.

Modified after Ross and Hopkins (1975) and Stone et al. (2005). 114

Figure 4.2. Local Kambalda Dome mine stratigraphy in an idealized cross-

section showing the Lunnon Basalt Formation (footwall), and Kambalda

Komatiite Formation comprising the Silver Lake and Tripod Hill Members.

The Silver Lake Member exhibits thickened channel facies, thin flank facies,

interflow metasedimentary rocks and Ni sulfide mineralization within a trough

feature. Modified from Lesher and Groves (1984). 116

xvi

Figure 4.3. 3D model of the Lunnon Basalt surface (shown in green) and

0.4% Ni grade shell (shown in red) as modeled with Leapfrog®. Victor trough

and Long trough interpretations shown with dashed lines, with select ore

shoots labeled (Gibb, Victor, McCleay, Long and Moran). Grey shading

delineates approximate flank facies distribution. View looking west. 119

Figure 4.4. Plot of distance (m) and azimuth of samples from nickel

mineralization > 0.4 wt% Ni. Each data point is an average of the closest three

distances and azimuths. Rose diagram showing distribution of azimuths with

general trend (335°) of the Long-Victor channels shown by grey arrow, as

observed in Figure 4.1. 121

Figure 4.5. Plot of FeOtot versus MgO wt% for the basal flow within the

Long-Victor area. Oxides as wt% and recalculated to anhydrous. Samples are

characterized as channel facies B-zone (Ch Bz) and spinifex (Ch Spfx) and

flank facies B-zone (Fl Bz) and spinifex (Fl Spfx). Volcanic flow facies fields

from Barnes (2006). Modelled olivine compositions (Fo) in pure adcumulate

shown on right hand side. Magma liquids in equilibrium calculated olivine

compositions (Fo) shown on left hand side and along top. 126

Figure 4.6. Plot of Al2O3 and TiO2 versus MgO for the basal flow within the

Long-Victor area. Oxides as wt% and recalculated to anhydrous. Samples are

characterized as channel facies B-zone (Ch Bz) and spinifex (Ch Spfx) and

flank facies B-zone (Fl Bz) and spinifex (Fl Spfx). 126

Figure 4.7. Median primitive mantle normalized trace element plots of the

samples from the basal flow in the Long-Victor area. Samples divided into

channel and flank facies, and spinifex textured and B-zone cumulates. 127

Figure 4.8. Primitive mantle normalized chalcophile element metal diagrams

for the basal flow within the Long-Victor area. Spinifex textured samples

shown in black and B-zone cumulate samples in black. Normalizing values

from McDonough and Sun (1995). 128

Figure 4.9. MgO wt% versus chalcophile element for all samples from the

basal flow. Visual trends shown by dashed lines. 129

xvii

Figure 4.10. PGE/Tipmn versus MgO wt% for all samples from the basal flow.

Samples with S > 0.25 wt% on the left hand side and samples with S < 0.25

wt% on the right hand side. Samples are subdivided based on flow facies

(channel = Ch, and flank = Fl) and komatiite flow facies (B-zone cumulates =

Bz, and spinifex textured = Spfx). 130

Figure 4.11. Inter-chalcophile element relationships for samples from the

Long-Victor basal flow with S>0.25wt%. 131

Figure 4.12. Platinum (ppb) versus sulfur (S wt%), and sulfur (S wt%) versus

MgO (wt%) for sulfur-poor (S<0.25 wt%) Long-Victor basal flow samples. 132

Figure 4.13. Major and trace element abundances plotted as a function of

distance from known mineralization (Ni >0.4%) which characterizes the

channel (c.f. Fig. 4.3). Samples are classified as channel (Ch) and flank (Fl),

as interpreted from constructed cross-sections. Samples are further subdivided

based on texture: B-zone (Bz) and spinifex (Spfx). Median values for B-zones

(solid line) and spinifex (dashed line) for channel and flank environments are

shown. Calculated best fit lines for flank B-zones (blue) and spinifex (red) are

shown, with R2 values for spinifex. Channel and flank subdivision at a

distance of 100 m is based on data distribution. 135

Figure 4.14. A. Ni/Tipmn versus MgO for the Long-Victor system, basal flow

samples shown in red diamonds. Calculated Ni normalized to actual Tipmn

plotted as black triangles. Ni/Ti trend line based on a derived equation. B.

Pt/Tipmn versus MgO for Long-Victor, basal flow samples shown in red

diamonds. Calculated Pt normalized to actual Tipmn plotted as black triangles.

Trend line of Pt/Ti represents perfectly incompatible elements at a determined

constant ratio of 0.67. 138

Figure 4.15. Plots of PGE/Tipmn versus MgO (wt%) for Long-Victor samples

exhibiting chalcophile element enrichment based on Pt and Pd abundances.

Samples are plotted as analytical data in red and calculated chalcophile

element abundance in grey (Table 4.4). Samples with sulfur greater than 0.25

wt% are shown on the left hand side and samples with sulfur less than 0.25

wt% on the right hand side. Blue lines define the analytical uncertainly field

around the numerically modelled background values (see Appendix C). 139

xviii

Figure 4.16. Plots of Pt correlations to incompatible elements (TiO2 and S)

and chalcophile elements (Pd, Ni) for the Long-Victor basal flow samples

with low sulfide abundance (< 0.25 wt%) and a chalcophile element

enrichment signature. 141

Figure 4.17. Ni/Tipmn versus MgO (wt%) and Pt/Tipmn versus MgO (wt%) for

Long-Victor basal flow samples, filtered to remove enrichment signature

(Pt/Ti pmn <0.88 and Pd/Ti pmn < 1.65). Samples are plotted as analytical data in

red and calculated chalcophile element abundance in grey with lines

delineating ± 500 ppm uncertainty for Ni, and ±2 ppb uncertainty for Pt. 143

Figure 4.18. Change in chalcophile element abundance from calculated

background values (Δ) for sample from Long-Victor basal flow. Samples

exhibiting enrichment signatures are removed. A. Calculated Pt (ppb)

depletion, with modeled depletion lines of 100%, 75%, 50% and 0% shown.

Dark grey shading delineates fields of uncertainty. B. Calculated Pt depletion

versus Pd depletion with ±2 ppb uncertainty applied to both. C. Calculated Rh

depletion versus Pt depletion with uncertainty shown by grey bars. D. Nickel

depletion versus Pt depletion with uncertainty shown by grey bar. 143

Figure 4.19. Leapfrog 3D-model of chalcophile element (PGE) mineralization

signatures within the basal flow of the Long-Victor channels. A. Lunnon

Basalt surface with 0.4% Ni grade shell shown. B. Modeled surface of the

basal flow spinifex with colour gradients representing ore forming signatures

observed in the spinifex; green = background, blue = depletion, and red =

enrichment. C. Mineralization signatures observed in the B-zone cumulate,

projected to the modeled surface of the basal flow spinifex. 146

Figure 4.20. Plot of distance (m) versus Ni grade (%) for all samples from the

basal flow of the Long-Victor system. Distances are an average of the three

closest Ni occurrences to each sample. Ni grade (%) represents the average Ni

abundance for those three occurrences. 147

Figure 4.21. A. Pt/Ti pmn and B. Pd/Ti pmn versus distance (m) to nickel

mineralization. Samples are classified as Enriched (mineralized based on Pt/Ti

and Pd/Ti ratios), Enriched LS (mineralized based on Pt/Ti and Pd/Ti ratios

with S<0.25%), Depleted (chalcophile element depleted samples as

xix

determined from previous section) and Background (samples which exhibit no

indication of chalcophile element enrichment or depletion). Plots are

domained into three spatial regions A, B, and C based on predominant ore

forming signatures at the respective distances. 148

Figure 4.22. Pt/Tipmn and Pd/Tipmn versus distance (m) to Ni mineralization,

focusing on samples within 80 m of known mineralization. Enriched

(mineralized based on Pt/Ti and Pd/Ti ratios), Enriched LS (mineralized based

on Pt/Ti and Pd/Ti ratios with S<0.25%), Depleted (chalcophile element

depleted samples as determined from previous section) and Background

(samples which exhibit no indication of chalcophile element enrichment or

depletion). 149

Figure 4.23. Pt, Pd and Rh for each chalcophile element depleted sample

from Area C shown as % depletion versus distance from mineralization ≥

0.4% Ni. 150

Figure 4.24. Chalcophile element depletion (left) and enrichment (right) as a

percentage change from the calculated background for each chalcophile

element of the basal flow, from drill hole KD6024. No chalcophile element

uncertainty was applied to the interpreted mineralization signatures. Samples

195.7, 196.0, 196.9, and 209 m from Lesher and Arndt (1995) and Lesher et

al. (2001). 152

Figure 4.25. Schematic cross-section through interpreted paleo-volcanic

setting of Victor and Long channels showing relative locations of flank

environments. Chalcophile element enrichment zones shown in red dots,

chalcophile element depletion shown in blue shading and areas of recharge

(background) in grey. 154

Figure 4.26. Time sequence block model for the progressive emplacement,

mineralization and preservation of chalcophile element ore forming

signatures. Komatiite flows colour coded for chalcophile signature: green =

background, blue = depleted, red = enriched. 156

Figure 5.1. Yilgarn Craton showing subdivision of the South West Terrane,

Youanmi Terrane and Eastern Goldfields Superterrane. Youanmi Terrane

granite-greenstone belts (dark grey) include: Lake Johnston (LJGB),

xx

Ravensthorpe (RGB), Forrestania (FGB) and Southern Cross (SCGB)

greenstone belts. Eastern Goldfields Superterrane granite-greenstone belts

(medium grey) include: Norseman (NGB) and Kalgoorlie (KGB). Lake

Johnston Greenstone Belt nickel mines include: EA (Emily Anne deposit) and

MH (Maggie Hays deposit). Modified from Department of Industry and

Resources (2008). 171

Figure 5.2. Generalized stratigraphic column for the Lake Johnston

Greenstone Belt; modified from Gower and Bunting (1976). * U-Pb age

determinations from Wang et al. (1996). 172

Figure 5.3. Geological plan map of the study area within the Lake Johnston

Greenstone Belt, showing the Honman and Maggie Hays Formations.

Honman Formation is subdivided into lithologic units. Strong deformation at

the northern end and along basal contact of the CUU in proximity to

remobilized Ni sulfide mineralization shown as wavy lines. All diamond drill

holes examined are shown, and key drill holes referenced in the paper labeled. 173

Figure 5.4. Composite stratigraphic column for the Honman Formation as

observed from diamond drill cores (LJD0126, LJD0048, LJD0011,

LJD0054A, LJD0087A, LJD003A, LJD0039, LJD0038, LJD0049, LJD0074,

LJD0055W2, LJD0092). Approximate intrusive level of the Central

Ultramafic Unit and narrow intrusive sills (banded iron formation-hosted sills)

shown along the left hand side. 175

Figure 5.5. Oblique Leapfrog® model view looking down and north-east

towards the local Maggie Hays nickel-deposit stratigraphy. Stratigraphy from

left to right consists of the Banded Iron Formation Unit, Transition Zone Unit,

Central Ultramafic Unit and Felsic Volcanic Unit. Scale bar in metres.

Western ultramafic unit not shown for clarity, but occurs to the left of the

Banded Iron formation. 176

Figure 5.6. Jensen cation plot from the Felsic Volcanic Unit and ultramafic

units from the Lake Johnston Greenstone Belt: felsic volcanic rocks, Central

Ultramafic Unit (CUU) pyroxenites and olivine cumulates, and Western

Ultramafic Unit (WUU) komatiites. H-Fe th as (high-Fe tholeiitic andesite),

H-Mg th ba (high-Mg tholeiitic basalt). 178

xxi

Figure 5.7. Primitive mantle-normalized trace element patterns for the Felsic

Volcanic Unit shown as black lines. Data fields for TTG/TTD type (Black

Flag Formation: Morris and Witt, 1997) and Arc-type felsic volcanism from

Eastern Goldfields Superterrane (EGS: Morris and Witt, 1997; Messenger,

2000; Barley et al., 2008). Normalizing values from McDonough and Sun

(1995). 180

Figure 5.8. Drill core photos and photomicrographs of representative Honman

Formation units. A. Part of the Transition Zone (TZ) Unit from LJD0038.

Felsic Volcanic Unit lithology with minor garnet on left, garnetite in middle

(magnified in B.), and chert with minor sulfide on right. B. Garnetite lithology

(LJD0038). C. Banded Iron Formation Unit (LJD0011). D. Iron-poor Fe-

formation. E. Spinifex texture from the Western-UU (LJD0011). F. Flow top

breccia texture from the Western-UU (LJD0126). G. Polarized light

photomicrograph of garnetite (LJD0038) amp = amphibole, bio = biotite, grt =

garnet. H. Reflected light photomicrograph of quartz-arenite (quartz with

trace pyrite), exhibiting graded bedding (LJD0011). 182

Figure 5.9. Plot of FeO versus MgO (anhydrous wt.%) for WUU (spinifex

and B-zone cumulates) and CUU, subdivided into spatial zones as shown in

Figure 5.10 (zone 0 = gabbroic, zone 1 = amphibolite, zone 2, 3, 4 = olivine

cumulates). Modified from Barnes et al., (2004). 188

Figure 5.10. Cross-section from line 6430470mN through the Honman

Formation stratigraphy, showing stratigraphic succession (Felsic Volcanic

Unit, Transition Zone Unit, BIF Unit, WUU) and conformal setting of the

Central Ultramafic Unit (CUU) and smaller banded iron formation-hosted

ultramafic sub-unit (BIF-hosted intrusions). Spatial geochemical zones shown

in within the CUU (as used in Figs. 5.9), zone 0 = gabbroic; zone 1=

pyroxenite; zone 2 = mixture of adcumulates to orthocumulates with lower

forsterite olivine; zone 3 = dominant adcumulates with moderate forsterite

olivine (Fo90-92); zone 4 = olivine adcumulates with highest forsterite content

(Fo93-94). 190

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Figure 5.11. Primitive mantle normalized trace element patterns of select

samples from the CUU (blue lines), WUU (grey lines) and mean FVU (red

line). Data from Chapter 6 and Appendix B. Normalizing values of Sun and

McDonough (1989). 191

Figure 5.12. Drill core photos and photomicrograph of the Central Ultramafic

Unit. A. Top contact between the BIF Unit and the CUU. Note the low-angle

bedding in the banded iron formation (i.e. parallel to core axis) and

conformable contact between CUU and BIF Unit (LJD0054A). B. Small

siliceous xenolith, with felsic xeno-melt on top-left side, hosted in the CUU

proximal to the footwall contact. C. Cross-polarized photomicrograph of

weakly altered olivine cumulate within the CUU (LJD003A). 192

Figure 5.13. Bi-variant plot of TiO2 and Al2O3 for all samples from the

Maggie Hays system data from this volume (Chapter 6). WUU spinifex

textured samples (spfx WUU). CUU; pyroxenite lithology (Border), olivine

cumulate lithology (CUU Ol), gabbroic lithology (gabbro). Felsic Volcanic

Unit (felsic) with calculated averages for contaminant 1 and 2 shown.

Barberton-type komatiite trend line shown for comparison with two

component mixing lines between Barberton-type liquid and both potential

felsic contaminants shown. Effects of olivine accumulation shown as %

trapped liquid lines below the Barberton-type liquid origin. 194

Figure 5.14. Schematic graphic model of the emplacement of the CUU,

showing the dominant role that stratigraphy plays in controlling the intrusions

morphology. A. Two layer stratigraphy BIF with density of 3.2 overlying

felsic volcanic with density of 2.4. Upward propagation of ultramafic magma

through the felsic volcanic shown. B. Upward propagation is inhibited at the

boundary between BIF and felsic volcanic, causing the lateral spreading of the

ultramafic magma. C. Continual magma injection results in over-pressuring of

the magma chamber (CUU) and eventual breach of the BIF occurs. Ultramafic

magma progresses to the surface and develops into an extrusive komatiite

flow field (WUU). 203

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Figure 6.1. Southwestern region of Western Australia, with Yilgarn Craton

and the three constituent subdivisions: South West Terrane, Youanmi Terrane

and Eastern Goldfields Superterrane shown (Cassidy et al., 2006). Greenstone

belts shown as light grey within the Eastern Goldfields Superterrane, with

Kalgoorlie (K) and Norseman (N) areas labeled. Greenstone belts within

Youanmi Terrane shown as dark grey, with Lake Johnston Greenstone Belt

(LJGB), Southern Cross (SCGB), Forrestania (FGB), and Ravensthorpe

(RGB) shown. Nickel mines Maggie Hays (MH) and Emily Ann (EA) shown. 218

Figure 6.2. Stratigraphic sequence of the Lake Johnston Greenstone Belt.

Modified from Gower and Bunting (1972; 1976); (see Chapter 5). 219

Figure 6.3. Geological plan map of the of the Maggie Hays Ni deposit

stratigraphy, comprising Maggie Hays, Honman and Glasse Formations. The

Honman Formation is divided into five lithological units: felsic volcanic,

transition zone unit (TZU), banded iron formation (BIF unit), sedimentary

unit, Western ultramafic unit (WUU), Central ultramafic unit (CUU) and

Eastern ultramafic unit (EUU). Strong deformation at the northern end and

along the basal contact of the CUU in proximity to mobilized Ni sulfide

mineralization shown by wavy lines. Diamond drill holes examined and

sampled in this study shown by the drill hole trace, and key drill holes referred

to in this work are labeled with the collar identification. 221

Figure 6.4. Cross-section on line 6430610mN through the Maggie Hays

deposit stratigraphy (Honman Formation: Felsic Volcanic, TZU, BIF-unit, and

WUU) with crosscutting CUU. Major lithological divisions of the CUU

shown. Facing direction as determined from spinifex texture within the WUU

and graded bedding within the quartz arenite shown by black arrow. Two drill

holes logged and sampled are labeled and shown in black (LJD0003A,

LJD00011). 222

Figure 6.5. 3D computer generated lithological model of the northern portion

of the CUU (purple), with point of view from the NE looking to the SW (see

Fig. 6.3). Stratigraphy dips towards the east at 60°, as shown by the Transition

Zone unit. Maggie Hays and North Shoot mineralized zones shown in red

(0.4% Ni grade shell). 224

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Figure 6.6. Bi-variant whole-rock geochemistry plots of major and trace

elements for samples from the CUU (diamonds) and the WUU (triangles).

Major elements are recalculated to anhydrous abundances. Chromite liquid

trends from Barnes (2006). 230

Figure 6.7. Median primitive mantle normalized trace element patterns for the

CUU (amphibolite samples), WUU (spinifex textured samples) and felsic

volcanic rocks. Median Barberton Formation komatiites (Barberton-type

komatiites) and median Silver Lake Formation komatiites from Kambalda

Dome (Munro-type komatiite) shown for comparison (Chapter 4). Primitive

mantle normalizing values from McDonough and Sun (1995). Barberton data

from Blichert-Toft et al. (2004) and Chavagnac (2004). 232

Figure 6.8. Bi-variant whole-rock geochemistry plots of chalcophile elements

and sulfur from the CUU (diamonds) and WUU (squares). Samples filtered

for S <1% to remove strong enrichment resulting from accumulated sulfide

liquid. 233

Figure 6.9. PGE/Tipmn versus MgO (wt%) for samples from the WUU

(squares) and CUU (diamonds). Dashed line of constant PGE/Tipmn are

median values of low-sulfur samples of both CUU and WUU. 235

Figure 6.10. Plot of FeO versus MgO (anhydrous wt.%) for WUU (spinifex

and B-zone cumulates) and CUU, subdivided into spatial zones as shown in

Figure 5.10 (zone 0 = gabbroic, zone 1 = amphibolite, zone 2, 3, 4 = olivine

cumulates). Calculated olivine compositions (Fo) for pure olivine adcumulates

are shown on the right hand side of the figure. Calculated olivine

compositions (Fo) in equilibrium with magma liquid compositions are shown

on left and along top of the figure. Modified from Barnes et al. (2004). 237

Figure 6.11. Plots of titanium normalized chalcophile elements versus MgO

for the Maggie Hays system. Geochemical assay data plotted as grey

diamonds, with equivalent calculated values shown as (+). Calculated

background lines shown as solid black lines with error lines light grey (Ni ±

500 ppm; Pt, Pd ± 2 ppb; Rh ± 1 ppb). 239

xxv

Figure 6.12. All whole-rock samples that are chalcophile element enriched

samples from the Maggie Hays system with S >0.25 wt% (A) and S<0.25

wt% (B). Raw data plotted as diamonds (WR data), calculated background for

each sample shown as (x: Pt/Ti n calc). Ideal calculated background shown as

constant solid line with ± 2 ppb error bars shown as dashed lines. 241

Figure 6.13. Chalcophile element depleted samples. A. Pd/Tipmn versus

Pd/Tipmn for all samples with background and depleted signatures. Lines at

0.63 Pt/Tipmn and 0.85 Pd/Tipmn define median background ratios. B.

Calculated Pd and Pt depletion as ppb with ± 2 ppb uncertainty (grey shading)

shown. C. Calculated Pt depletion as ppb with modeled lines of percent

depletion (50, 75 and 100%) with ± 2 ppb uncertainty shown by grey shading.

D. Calculated depletion for Ru and Pt (ppb). E. Calculated Ir depletion (ppb)

versus Pt (ppb) depletion. F. Calculated Ni (ppm) depletion versus Pt (ppb)

depletion. 243

Figure 6.14. Pt/Tipmn versus distance (metres) for all samples from within the

CUU. Samples are classified as background, and chalcophile element enriched

and depleted. The following Figure 6.15 represents samples within 350 m of

mineralization. 245

Figure 6.15. Pt/Tipmn and Pd/Tipmn versus distance for samples within 350 m

of mineralization within the CUU (close up of Fig. 6.14). Samples are

classified as background, and chalcophile element enriched and depleted.

Arrows show visual trends of increasing and decreasing magnitude of the

chalcophile element depletion signature. 245

Figure 6.16. Cartoon long section of the Lake Johnston Greenstone Belt

stratigraphy showing the CUU conduit system and overlying WUU. A.

Emplacement model. B. Ore forming process, through assimilation of the

overlying sulfur-rich contaminant, with small inset cross-section shown. C.

Ore forming process with areas hosting mineralization signatures indicated. D.

Final stage of the conduit system and the spatial distribution of ore forming,

and background chalcophile element abundances shown. 248

Figure 6.17. 3D computer generated lithological model of the northern

portion of the CUU with point of view from the NW looking to the SE (see

xxvi

Fig. 6.3) showing the areas of intersection between the CUU (purple) and the

modeled TZU surface (light grey). Lithological drill intersections utilized in

TZU modeling shown as black circles. 250

Figure 7.1. Map of northern Sweden, Norway, Finland and northwestern

Russia showing the distribution of the Paleoproterozoic Central Lapland

Greenstone Belt (green), and associated komatiite and picritic rocks (black).

Sampling areas are delineated by boxes comprising the: Archean Enontekiö

Area, and Paleoproterozoic Pulju and Karasjok Greenstone Belts. Inset map of

Norway, Sweden and Finland showing major tectonic divisions of the Baltic

Shield. Modified from Hanski et al. (2001). 265

Figure 7.2. Paleoproterozoic stratigraphic sequences and correlations within

the Central Lapland Greenstone Belt, comprising the Karasjok, Pulju and

Kittilä Greenstone Belts; with arrows indicating formations sampled within

the Karasjok and Pulju belts. Formations and Groups are identified with

characteristic lithologies summarized: mf. vol. = mafic volcanic, amp. =

amphibolite, vol. clast. = volcaniclastic, kom. = komatiite, psam. = psammite,

thole. vol. = tholeiitic volcanic, cong. = conglomerate, fels. vol. = felsic

volcanic, suf. sed. = sulfidic sediment, qutz. = quartzite, BIF = banded iron

formation. Complied from Braathen and Davidson (2000); Papunen (1998);

Lehtonen et al. (1998). Age determinations from Pihiaja and Manninen

(1988), Hanski et al. (1997). 268

Figure 7.3. Bivariant plots of major elements for the ultramafic units from the

three areas within the central Karelian Craton, as determined by XRF and

ICP-MS. Komatiites from the Archean Enontekiö area (Sarvisoaivi), and

Paleoproterozoic areas: Karasjok (Karasjok Greenstone Belt), and Nilivaara

and Hotinvaara (Pulju Greenstone Belt). 271

Figure 7.4. Bivariant plots of chalcophile and major elements for the

ultramafic units from the three areas within the central Karelian Craton, as

determined by fire-assay ICP-MS. Komatiites from the Archean Enontekiö

area (Sarvisoaivi) and Paleoproterozoic areas: Karasjok (Karasjok Greenstone

Belt), and Nilivaara and Hotinvaara (Pulju Greenstone Belt). 273

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Figure 7.5. [Al2O3] versus [TiO2] high-MgO volcanic discrimination diagram

of Hanski et al. (2001). Where [Al2O3] and [TiO2] are normalized mole

proportions using the equations [Al2O3] = Al2O3/(2/3-MgO-FeO) and [TiO2]

= TiO2/(2/3-MgO-FeO): (see Hanksi, 1992). 275

Figure 7.6. FeO wt% versus MgO wt% recalculated to volatile free for

ultramafic samples from Central Karelian Craton. Olivine compositions in

equilibrium with liquid shown as solid lines (Fo91-94) and olivine compositions

in adcumulates (pure olivine) shown as diamonds (Fo95-85), with volcanic

facies discrimination fields as determined by Barnes (2006a). 278

Figure 7.7. Pt/Alpmn versus Pd/Alpmn diagram for classifying chalcophile

element mineralization signatures within komatiitic systems. Fields derived

from mineralized Munro- (Long-Victor deposit, Kambalda Dome) and

Barberton-type (Maggie Hays deposit, Lake Johnston Greenstone Belt)

komatiite systems in Western Australia (Chapters 4 and 6). 280

Figure C.1. Duplicate analyses plots for the platinum group elements.

Coefficient of Variation (CV) vs. Duplicate Mean, with Relative Error shown

(RE). Half absolute relative difference (HARD) vs. Ranked Percentile with

vertical line demarking 95th percentile and horizontal lines 2 standard

deviations (2s). C5

Figure C.2. A graphical representation of the estimation of precision by the

regression of duplicate analyses using analytical data determined by FA-ICP-

MS by Geolabs. C7

Figure C.3. Calculated precision as a function of concentration for platinum

determined by FA-ICP-MS by Geolabs. C8

Figure C.4. Plots of calculated Pt (see Appendix D) versus MgO and

calculated Pt/Tipmn versus MgO, with total maximum sample uncertainty

shown by dashed lines. Total uncertainty includes precision estimates of MgO

and TiO2 as derived from Ultratrace duplicate analyses and Pt as derived from

Geolabs. C9

Figure C.5. Comparison between ICP-MS and ICP-OES, and ICP-MS and

XRF analyses for elements that were analyzed by both analytical methods.

xxviii

Linear regressions and r2 correlation coefficients shown for each data set from

Ultratrace and Geolabs. C11

Figure D.1. A. Plot of Pt (ppb) versus MgO (wt%) for all Kambalda data with

sulfur < 0.25 wt%, showing general negative correlation with MgO with

potential Pt depletion (D) and enrichment (E) overprinting trend as shown by

arrows. B. TiO2 versus MgO (wt%) showing strong negative correlation

between the two elements. D1

Figure D.2. Plots Pt/Tipmn versus MgO wt% and Pd/Tipmn versus MgO wt% of

all Kambalda data with sulfur < 0.3 wt%, showing constant value with

varying MgO content. Deviation from a constant value shown as D (depletion)

and E (enrichment). D2

Figure D.3. Plot of Pd/Ti pmn versus Pt/Ti pmn for all Kambalda samples with

S<0.3wt%. Trend lines shown for low sulfur Pt and Pd

enrichment/mineralization (Pt+Pd En), Pt and Pd depletion (Pt+Pd De) and

enrichment or depletion of either Pt or Pd from a constant value. D3

Figure D.4. Final data set (n=75) from Kambalda which falls within ± 2 ppb

of calculated Pt/Tipmn and Pd/Tipmn ratios. D4

Figure D.5. Primitive mantle normalized noble metal plot of select samples. D4

Figure D.6. Ni (ppm) versus MgO and Ir/Ti pmn versus MgO for Kambalda

samples with linear regressions and R2 values. D4

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List of Tables Page

Table 2.1. Greenstone belts containing volcaniclastic textured ultramafic

lithologies. Barberton-type komatiite (B-type), Munro-type komatiite

(M-type), Karasjok-type komatiite (K-type). 27

Table 2.2. Case study intrusions that have chalcophile element ratios utilized to

identify orthomagmatic mineralization. 51

Table 4.1. Summary of geochemistry for the basal flow at Long-Victor: Median

(Med), Maximum (Max), Minimum (Min), Number of samples (N).

Data filtered for S<0.25 wt%. Oxides are recalculated to anhydrous

conditions and reported in wt%, metals and trace elements are reported

as ppm unless denoted * then ppb. 124

Table 4.2. Average (n=19) chalcophile element abundances, MgO and TiO2

content of spinifex textured samples from the Long-Victor area. (TiO2

and MgO as wt%, Ni, Cu, Co, Cr, Zr, Gd as ppm, and Ir, Ru, Rh, Pt, Pd,

Au as ppb). 128

Table 4.3. Comparison of geochemical and physical attributes of channel and

flank facies. Compiled from Gresham and Loftus-Hills (1981); Lesher et

al. (1984); Lesher (1989); Lesher and Arndt (1995); Lesher et al. (2001);

Barnes (2006). 134

Table 4.4. Equations derived and used to calculate background abundances of

Ni, Pt, Pd, and Rh as a function of the MgO content of the sample.

Complete list of equations provided in Appendix D. 137

Table 5.1. Whole rock geochemistry analyses of representative units from the

Honman Formation. With drill collar, sample depth, lithological unit

(FVU = Felsic Volcanic Unit; WUU = Western-UU; CUU = Central-

UU) and lithology (Rhy-dac = rhyolite-dacite; Spfx = spinifex; OC =

olivine cumulate, Pyr = pyroxenite) in header. Trace element ratios

La/Sm*, Th/Sm*, Nb/Th* and Gd/Yb* primitive mantle normalized.

Normalization values from McDonough and Sun, (1995). 179

xxxi

Table 6.1. Median values of major and trace elements for WUU (B-zone

cumulates, Spinifex textured samples) and CUU (amphibolite and

olivine cumulate) with data from Kambalda Dome Long-Victor system.

(Channel B-zone, Flank B-zone, Channel Spinifex and Flank Spinifex).

All data filtered S<0.25%. Trace elements and chalcophile elements in

ppm unless marked * indicating ppb. 231

Table 6.2. Correlation matrix for select major elements and chalcophile

elements from Maggie Hays Samples. Filtered for S <1%. 234

Table 6.3. Equations derived and utilized to calculate background abundances

of Ni, Pt, Pd, and Rh as a function of the MgO content of the samples

within the Maggie Hays system. 239

Table 7.1A. Whole-rock geochemistry of ultramafic rocks from Karasjok and

Pulju Greenstone Belts and Enontekiö area. Major elements analyzed by

XRF and given in wt% oxide and chalcophile elements by ICP-MS from

NiS fire assay pre-concentration with PGE concentrations in ppb and Ni,

Cu in ppm. Morphology as determined from outcrop mapping: TF =

Thin flow, MF = Massive flow, PF = Pillowed flow, FR = Fragmental

textured, Flt = Flow top. Sample location given as decimal degrees

latitude (Lat) and longitude (Long) with WGS84 datum. LOI = loss on

ignition, n.d. = not determined. 279

Table 8.1. Partition coefficients for the chalcophile elements between silicate

liquid and sulfide liquid. 1. Francis (1990); 2. Sattari et al. (2002); 3.

Gaetani and Grove (1997); 4. Peach et al. (1990); 5. Jana and Walker

(1997); 6. Rajamani and Naldrett (1978); 7. Stone et al. (1990); 8.

Bezmen et al. (1994); 9. Fleet et al. (1999); 10. Peach et al. (1994); 11.

Helz and Rait (1988). 295

Table 8.2. Mineralization signature characteristics of the chalcophile elements 295

Table C.1. Lowest Level of Detection (LLD) reported by both analytical labs

for each analytical technique. C3

xxxii

Table C.2. Summary of Precision as determined for major and chalcophile

elements through duplicate analyses. S0 and K are Y-intercept and

slope, respectively, from linear regressed duplicate analyses described

previously. MDL = method detection limit. Precision (%) is a median

value over the compositional range given. Range in wt% for oxides, ppm

for Cr, Ni, and Cu, and ppb for the PGE. C8

Table C.3. Calculated total maximum uncertainty for the chalcophile elements.

Values are median values covering the range of compositions observed

in the Long-Victor system (9-48 wt% MgO). C9

Table D.1. Step results of iteratively filtered Kambalda Dome data set. D3

Table D.2. Chalcophile elements as a function of MgO as derived for the

Kambalda Dome system (2.7 Ga Munro-type) with calculated R2 values D5

Table D.3. Calculated chalcophile content of a theoretical Kambalda primitive

magma (24 wt% MgO) compared with median spinifex textured samples

(n=15: filtered to remove mineralizing signatures) from Kambalda

Dome. D5

Table D.4. Chalcophile elements as a function of MgO as derived for the

Maggie Hays System (2.9 Ga Barberton-type) with calculated R2 values D6

Table D.5. Calculated chalcophile content of a theoretical Maggie Hays

primitive magma (26.8 wt% MgO) compared with median spinifex

textured samples (n=7: filtered to remove mineralizing signatures) from

Western Ultramafic Unit. D6

xxxiii

xxxiv

Acknowledgements

I would like to acknowledge the support and contributions from numerous

individuals, companies, organizations, and institutions that made this thesis possible.

Shannon Johns who never hesitated at the opportunity for an adventure in Australia,

only to endure four years of komatiite and nickel discussion. This thesis would not

have the same polish or sparkle without her continuous support, encouragement and

reviews.

Marco Fiorentini, Mark Barley and Steve Barnes, my supervisors who diligently

worked to put the project together, recruit me sight unseen, and persevered to

supervise the completion of the thesis.

This study would not have been possible without the assistance and contributions

from the nickel industry: BHP-Billiton, Independence Group NL., and Noril’sk

Nickel (formerly LionOre Pty.), with their continued support to the AMIRA P710A

project after the completion of P710. Exploration/project managers: Steve Beresford

(BHP-Biliton), Paull Parker (Independence Group), Chris Stott and Ian Gregory

(Noril’sk Nickel) provided access to mine sites and drill core, essential feedback and

discussion during the course of the P710A project and during the completion of this

thesis. Additional thanks to the geology groups at the Long-Victor (Somely Shepard

and Ricky Gordon), and Maggie Hays Mines (Alex Johnston and Nesbert Nyama)

who provided access to mines, drill core and digital data. Your assistance is greatly

appreciated.

Thesis work conducted in Finland and Norway was made possible through a student

grant from the Hickok-Radford Memorial Fund of the Society of Economic

Geologists Foundation Inc. to support field-based research. The field experience was

greatly enhanced by Juka Jokela who gave an excellent field trip of komatiites in

northern Finland and Norway.

xxxv

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Chapter 1. Purpose and Scope

Chapter 1. Purpose and Scope: The Application of Platinum Group Elements in Komatiite-Hosted Nickel Sulfide Exploration

1.1. Introduction

Exploration for komatiite-hosted nickel deposits is a continued challenge, due to

small target size and lack of alteration haloes associated with mineralization.

Currently, the discovery rate of komatiite-hosted nickel deposits is decreasing

(Hronsky and Schodde, 2006), as new deposits are typically located under cover and

at greater depths. Recent advances in targeting techniques, such as geophysics and

lithogeochemistry, have provided limited discovery success. However,

lithogeochemical targeting has the capacity to enhance the target size beyond

mineralization. Lithogeochemical indicators that have shown potential for use in

targeting are the platinum group elements (PGE), nickel (Ni), and copper (Cu), due

to their chalcophile nature (Lesher et al., 1981; Keays, 1982; Barnes et al., 1985;

1987; 1995; Barnes, 1990; Maier et al., 1998; Lesher et al., 2001). The chalcophile

nature of the PGE, Ni, and Cu results in the generation of predictable and

recognizable ore forming signatures. To date, the use of these chalcophile elements

as mineralization indicators is limited, as the size of ore-related anomalies remains

unconstrained.

This thesis will quantify the relative magnitude of chalcophile element (PGE, Ni,

Cu) ore forming signatures, and test the spatial correlation between these signatures

and known nickel mineralization in komatiite systems. By doing so, the size of ore

forming systems will be constrained, and a prospectivity indicator can be translated

into a nickel mineralization vector for application in the resource industry.

1.2. World Nickel Use and Discovery

Nickel is primarily used by the iron foundry industry in the generation of steel and in

the production of specialized Ni alloys (Reck et al., 2008). Nickel imparts corrosion

resistance and enhances tensile properties. Large scale use of Ni commenced in the

19th century, with substantial production increases during the Industrial Revolution

of the middle to late 1800s. More recent Ni consumption increases have occurred

due to the proliferation of Ni metal hydride batteries.

1


With the evolving use of Ni, production has fluctuated with global growth and

demand. Nevertheless, nickel demand has exhibited a steady increase of

approximately 4% per annum over the last 20 years (Mudd, 2009). Despite this

steady increase in demand, there has been no increase in the rate of discovery of new

deposits. In Australia, the bulk of the current Ni resource was identified prior to

1973 (Jaques et al., 2005; Hoatson et al., 2006). This decreasing rate of discovery in

Australia is skewed even more if Ni laterite systems are excluded from the

calculation.

1.3. Nickel Prospectivity and Purpose

Nickel deposits occur in a number of tectonic settings (rifts, plumes, etc.) and host

lithologies (ultramafic rocks, high-MgO rocks, etc.). Some Ni deposits and

magmatic systems are better understood and constrained than others, but a general

understanding of the search space for Ni mineralization is well documented in the

literature. Previous Ni deposit research has identified prospective lithologies and

tectonic settings, applied Ni mineralization models, and constrained the higher

prospectivity areas within the systems (Naldrett, 1997; Barnes and Lightfoot, 2005;

Naldrett, 2005; Hoatson et al., 2006; Eckstrand and Hulbert, 2007). Much of the

prospective Ni mineralization search space is covered by active mining or mineral

exploration claims and tenements. Additional Ni discoveries will be made in these

areas, although the discovery rate will be low and at a higher cost.

By constraining the relative magnitude of chalcophile element ore forming

signatures and the spatial correlation between these signatures and Ni

mineralization, it is here postulated that: (1) exploration programs can become more

effective by maximizing the information obtained from each iterative step in the

exploration process; (2) gross prospectivity of a greenstone belt, stratigraphic

sequence, or lithologic unit can be assessed, prior to the outlay for intensive regional

airborne surveys and drill testing; and (3) lithogeochemical vectoring of Ni

mineralization within komatiite systems is possible.

2


1.4. Chalcophile Element Mineralization Signatures in Komatiites

Chalcophile element ore forming signatures in komatiites are the product of a sulfide

phase that becomes saturated within a magmatic system (often referred to as sulfur

saturation: Lesher et al., 1981; Barnes et al., 1985; 1987a; b; 1995; Barnes, 1990;

Maier et al., 1998; Lesher et al., 2001; Fiorentini et al., 2010). Identification of

chalcophile element ore forming signatures is dependent on comparisons with

background geochemical conditions. The background chalcophile element budget

reflects the composition of the source, and is defined when a sulfur undersaturated

komatiite magma leaves the source area. This background geochemical condition

can only be used if the initially sulfur undersaturated magma erupts, differentiates,

fractionates, and crystallizes without sulfur saturation occurring. Under these

conditions, the chalcophile element background is a function of the current

constituent phases in the rock, and the partitioning of the chalcophile elements into

these phases (e.g. olivine, pyroxene, oxides, trapped liquid, glass: Barnes and Maier,

1999). The chalcophile element background is also defined by komatiite

geochemical type (Keays, 1982; Arndt et al., 2005), and PGE content dependent on

age (Maier et al., 2009; Fiorentini et al., in press).

If the sulfur undersaturated background magma erupts and becomes sulfur saturated

through changes in either composition (assimilation, crystallization: Lesher et al.,

1984; Naldrett, 2005; Li et al., 2009) or physical characteristics (temperature,

oxygen fugacity, pressure: MacLean, 1969; Haughton et al., 1974; Mavrogenes and

O’Neil, 1999), then a chalcophile element ore forming signature will result from the

segregation of the immiscible sulfide liquid. Due to the chalcophile nature of the

PGE, Ni, and Cu, these elements strongly partition out of the silicate liquid into the

sulfide liquid (Ragamani and Naldrett, 1979). As a consequence of sulfur saturation

and sulfide liquid segregation, two chalcophile element signatures are produced: (1)

chalcophile element enrichment in the sulfide liquid (mineralization), and (2)

chalcophile element depletion of the silicate magma.

Chalcophile element enrichment signatures are readily seen in the form of

accumulated Ni sulfide, ranging from massive accumulations to fine, disseminated

interstitial sulfides. Chalcophile element depletion signatures are more subtle and are

3


only identified by geochemical analysis of the silicate rock. Chalcophile element

(PGE, Ni, Cu) enrichment and depletion signatures are present within mineralized

systems, although both signatures are not always identified. Incidentally,

mineralization signatures commonly characterize a small fraction of the total volume

of the magmatic system, with the majority of the volume characterized as baseline

abundances (Fiorentini et al., 2010). Background chalcophile element abundances

are ubiquitous within Ni mineralized systems, and are a function of discontinuous

sulfur saturation and extensive recharge within the komatiite systems (Lesher et al.,

1984; Lesher and Arndt, 1995).

The recognition of chalcophile element mineralization signatures represents a viable

solution to increase the discovery rate of komatiite-hosted Ni deposits. Chalcophile

elements are intimately associated with mineralization, and chalcophile element

mineralization signatures are identifiable within mineralized systems (Lesher et al.,

2001; Fiorentini et al., 2010). However, chalcophile element mineralization

signatures do not represent vectors to mineralization, due to the current lack of

distance and size components.

1.5. Research Scope

The effective utilization of chalcophile element (PGE, Ni, Cu) signatures as

lithogeochemical mineralization vectors requires the definition of a quantifiable size

for the ore forming system. This thesis aims to translate chalcophile element

mineralization signatures into chalcophile element based mineralization vectors in

two ways:

1. Quantifying the magnitude of enrichment and depletion of the chalcophile

element signatures associated with Ni mineralization from background

abundances; and

2. Quantifying the spatial distribution and spatial correlation of the chalcophile

element mineralization signatures to Ni mineralization within komatiite

systems.

When these two components are constrained it is possible to interpret the variation

of chalcophile element mineralization signatures as vectors to Ni mineralization.

4


Both factors are addressed by the application of spatially constrained

lithogeochemical variation. Whole-rock chalcophile element abundances (PGE:

platinum [Pt], palladium [Pd], rhodium [Rh], ruthenium [Ru], iridium [Ir]; nickel

[Ni]; copper [Cu]) are analyzed and modelled from two orthomagmatic nickel

systems in two case study deposits within the Yilgarn Craton of Western Australia.

The first case study deposit (Chapter 4) is the Long-Victor Ni mine located on the

eastern side of the Kambalda Dome, which is hosted in 2.7 Ga Munro-type

komatiites of the Eastern Goldfields Terrane. The second case study deposit

(Chapter 6) is the Maggie Hays Ni mine, associated with 2.9 Ga Barberton-type

komatiites of the Lake Johnston Greenstone Belt, Youanmi Terrane. These two

deposits were selected to provide a range in geochemical type, age, and volcanic

setting. The candidate conducted preparatory work of 3-dimensional modelling and

spatial targeting of sample areas with Leapfrog® prior to field work. Field-based

research was carried out by the candidate at the Long-Victor and Maggie Hays Ni

mines between 2006-2009. This work consisted of core logging and lithological

sampling.

The understanding of ore forming signatures and size of Ni systems, obtained from

the two case studies, is applied to ultramafic units within the Karelian Craton in

northern Norway and Finland to assess the Ni prospectivity of the ultramafic rocks

within select greenstone belts (Chapter 7). The Karelian Craton was selected for

two reasons. Firstly, the craton hosts Proterozoic Karasjok-type komatiites (high Fe-

Ti-komatiites), which display similar volcanological and geochemical properties to

both Munro- and Barberton-type systems, yet contrast them in other aspects.

Secondly, to date, no economic Ni sulfide mineralization has been identified within

the northern portion of the craton, despite extensive work carried out by the Finnish

Geological Survey (GTK) and industry. Field-based research in Norway and Finland

was conducted by the candidate in 2007, and involved field mapping and lithological

sampling.

1.6. Thesis Overview

The thesis consists of eight chapters (including this Introduction), and is arranged in the sequence listed below to provide the necessary background information relevant to the topic and study areas. As the topic is pertinent to Ni exploration, this

5


thesis is prepared in manuscript form for future publication. The thesis comprises four manuscripts that have been submitted for publication (Chapters 4, 5, 6, 7), and linking chapters (1, 2, 3 and 8) that will not be published.

Chapter 2. Komatiites and Orthomagmatic Nickel

Chapter 2 provides a summary and overview of komatiites, the ore forming process, and mineralization indicators. This chapter provides a review and necessary background discussion to support the following chapters. It should be noted that Chapters 4 to 7 are written as manuscripts and are consequently less descriptive in “komatiite explanation”. This is due to the focus on specific aspects of the magmatic systems or the spatial correlation of the chalcophile element mineralization signatures. Chapter 2 is divided into four topical sections. The first section examines the main geochemical types of komatiites and contributing factors leading to geochemical variation within komatiites, igneous textures characterizing komatiites, and the current komatiite flow field development model. The next section examines the processes leading to ore formation within komatiite systems and the characteristics of mineralization. The final section examines the current mineralization indicators for Ni sulfide exploration, with examples and applications. Additional summaries are also provided in the last section, regarding chalcophile elements in intrusion-hosted PGE exploration and within the Noril’sk Ni system (high-MgO).

Chapter 3. The Kambalda Dome

Chapter 3 introduces the geological setting of the Long-Victor Ni mine, which comprises the first case study deposit in Chapter 4. The Long-Victor mine is located on the eastern flank of the Kambalda Dome, an area that has been extensively studied over the past 40 years. This chapter summarizes the tectonic setting, stratigraphy, deformation, alteration and metamorphism within the Kambalda Dome.

Chapter 4. The Size of Nickel Mineralization Systems: Examination of the PGE Distribution in the Long-Victor System, Kambalda Dome, Western Australia

Chapter 4 provides the first case study deposit for the use of chalcophile element signatures (specifically the PGE) as lithogeochemical mineralization vectors. The Long-Victor Ni deposit is characteristic of the majority of mineralized komatiites

6


identified globally. The deposit is hosted in 2.7 Ga Munro-type komatiites, the system is extrusive, and mineralization is hosted within a channelized flow environment. Additionally, the Long-Victor system is well defined by diamond drilling, with both flank and channel environments identified over a strike length of approximately 2 km. These characteristics make the Long-Victor deposit an ideal setting to investigate the size and geometry of the spatial and genetic correlation between localization of Ni sulfide mineralization and the variability of PGE abundance.

Chapter 5. Stratigraphy and Stratigraphic Control on the Style of Komatiite Emplacement in the 2.9 Ga Lake Johnston Greenstone Belt, Yilgarn Craton, Western Australia

Chapter 5 provides a detailed overview of the stratigraphic and tectonic setting for the Maggie Hays Ni deposit within the Lake Johnston Greenstone Belt. Relative to the Long-Victor deposit, the Maggie Hays Ni deposit is relatively unstudied in terms of the deposit or greenstone belt. This chapter examines the local mine stratigraphy of the Maggie Hays deposit and the nature of ultramafic magmatism hosting the Ni mineralization. This information provides a geological framework for the chalcophile element mineralization signatures examined in Chapter 6.

Chapter 6. Platinum Group Element Signatures and Spatial Distribution in an Intrusive Komatiite System: Examination of the Maggie Hays System, Lake Johnston Greenstone Belt, Western Australia

Chapter 6 provides the second case study deposit for the use of chalcophile element signatures (specifically the PGE) as lithogeochemical mineralization vectors. The Maggie Hays Ni deposit contrasts the Long-Victor system, as it is hosted within a 2.9 Ga, Barberton-type komatiite system. Additionally, the Ni mineralization is hosted within the sub-volcanic feeder, rather than within the extrusive component. Thus, the Maggie Hays deposit provides essential age, geochemistry, and deposit style diversity in the understanding of the spatial distribution of chalcophile element mineralization signatures.

Chapter 7. Application of Lithogeochemical Prospectivity for Komatiite-Hosted Nickel Sulfide Mineralization, Northern Finland and Norway

7


Chapter 7 provides an opportunity to assess the application of chalcophile element based mineralization indicators and vectors in other prospective systems. The understanding of chalcophile element mineralization signatures obtained from the contrasting Long-Victor and Maggie Hays mineralization systems is applied in a Ni prospectivity assessment of ultramafic rocks (Karasjok- and Munro-type komatiites) from the Karelian Craton of northern Norway and Finland. This field area provides a unique opportunity to test the application of lithogeochemical prospectivity in terranes with complex geology, sparse outcrop, and limited research and exploration activities to date.

Chapter 8. Application of Platinum Group Elements in Komatiite-Hosted Nickel Sulfide Exploration

Chapter 8 forms the conclusion to the thesis. The application of chalcophile elements in the exploration for komatiite-hosted Ni deposits was previously limited to regional prospectivity, due to the unconstrained size of ore forming systems. This thesis has outlined a systematic approach to: (1) identifying initial background chalcophile element concentrations in komatiite-hosted Ni systems; (2) quantifying deviations from the background, in the form of chalcophile element (PGE) enrichment and depletion signatures; and has (3) determined the spatial correlation between mineralization signatures and known mineralization within two differing Ni mineralized systems (Long-Victor and Maggie Hays). Therefore, the size of Ni mineralization systems is now constrained and the use of chalcophile elements (PGE) as mineralization vectors is possible.

8


1.7. References Arndt, N.T., Lesher, C.M., Czamanske, G., 2005. Mantle-derived magmas and magmatic Ni-Cu-

(PGE) deposits: Economic Geology, v. 100, p. 5-23.

Barnes, S-J., 1990. The use of metal ratios in prospecting for platinum-group element deposits in mafic and ultramafic intrusions: Journal of Geochemical Exploration, v. 37, p. 91-99.

Barnes, S-J., Boyd, R., Korneliussen, A., Nilsson, L.P., Often, M., Pedersen, R.B., Robins, B., 1987b. The use of Mantle normalization and metal ratios in discriminating between the effects of partial melting, crystal fractionation and sulfide segregation on platinum-group elements, gold, nickel and copper: Examples from Norway: Geo-Platinum 87, p. 113-143.

Barnes, S-J., Lightfoot, P.C., 2005. Formation of magmatic nickel sulfide ore deposits and processes affecting their copper and platinum-group element contents. In: Hedenquist, J.W., Thompson, J.F.H., Goldfarb, R.J., Richards, J.P. (eds.), Economic Geology 100th Anniversary Volume, p. 179-213.

Barnes, S-J., Maier, W.D., 1999. The fractionation of Ni, Cu and the noble metals in silicate and sulphide liquids. In: Keays, R.R., Lesher, C.M., Lightfoot, P.C., Farrow, C.E.G., (eds.), Dynamic processes in magmatic ore deposits and their application to mineral exploration: Geological Association of Canada, Short Course Notes, v. 13, p. 69-106.

Barnes, S-J., Naldrett, A.J., 1987. Fractionation of the Platinum-Group Elements and gold in some komatiites of the Abitibi Greenstone Belt, Northern Ontario: Economic Geology, v. 82, p. 165-183.

Barnes, S-J., Naldrett, A.J., Gorton, M.P., 1985. The origin of the fractionation of Platinum-group elements in terrestrial magmas: Chemical Geology, v. 53, p. 303-323.

Barnes, S.J., Lesher, M.C., Keays, R.R., 1995. Geochemistry of mineralised and barren komatiites from the Perseverance nickel deposit, Western Australia: Lithos, v. 34, p. 209-234.

Eckstrand, O.R., Hulbert, L.J., 2007. Magmatic nickel-copper-platinum group element deposits. In: Goodfellow, W.D., (ed.), Mineral deposits of Canada: A synthesis of major deposit types, district metallogeny, the evolution of geological provinces, and exploration methods: Geological Association of Canada, Mineral Deposit Division, Special Publication No. 5, p. 205-222.

Fiorentini, M.L., Barnes, S.J., Lesher, C.M., Heggie, G.J., Keays, R.R., Burnham, M.O., 2010. Platinum-group element geochemistry of mineralized and non-mineralized komatiites and basalts: Economic Geology, v. 105, p. 795-823.

Fiorentini, M.L., Barnes, S.J., Maier, W.D., Burnham, M., Heggie, G.J., in press. Global variability in the platinum-group element contents of komatiites: Journal of Petrology.

Haughton, D.R., Roeder, P.L., and Skinner, B.J., 1974. Solubility of sulfur in mafic magmas: Economic Geology, v. 69, p. 451-467.

Hoatson, D.M., Jaireth, S., Jaques, A.L., 2006. Nickel sulfide deposits in Australia: characteristics, resources, and potential: Ore Geology Reviews, v. 29, p. 177-241.

Hronsky, J.M.A., Schoddle, R.C., 2006. Nickel exploration history of the Yilgarn Craton: From the nickel boom to today. In: Barnes, S.J., (ed.), Nickel Deposits of the Yilgarn Craton: Geology, Geochemistry, and Geophysics Applied to Exploration: Society of Economic Geologists, Special Publication, No. 13, p. 1-12.

Jaques, A.L., Huleatt, M.B., Ratajkoski, M., Towner, R.R., 2005. Exploration and discovery of Australia’s copper, nickel, lead and zinc resources 1976-2005: Resources Policy, v. 30, p. 168-185.

Keays, R.R., 1982. Palladium and iridium in komatiites and associated rocks: application to petrogentic problems. In: Arndt, N.T., and Nisbet, E.G. (eds.), Komatiites: George Allen and Unwin, London. p. 436-457.

Li, C., Ripley, E.M., Naldrett, A.J., 2009. A new genetic model for the giant Ni-Cu-PGE sulfide deposits associated with the Siberian Flood Basalts. Economic Geology, v. 104, p. 291-301.

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Lesher, C.M., Arndt, N.T. Groves, D.I. 1984. Genesis of komatiite-associated nickel sulfide deposits at Kambalda, Western Australia: a distal volcanic model. In: Buchanan, D.L., and Jones, M.J., (eds.), Sulfide deposits in mafic and ultramafic rocks.

Lesher, C.M., Lee, R.F., Groves, D.I., Bickle, M.J. Donaldson, M.J., 1981. Geochemistry of komatiites from Kambalda, Western Australia: I. Chalcophile element depletion- a consequence of sulfide liquid separation from komatiitic magmas: Economic Geology, v. 76, p. 1714-1728.

Lesher, C.M., Burnham, O.M., Keays, R.R., Barnes, S.J., Hulbert, L., 2001. Geochemical discrimination of barren and mineralized komatiites associated with magmatic Ni-Cu-(PGE) sulfide deposits. Canadian Mineralogist, v. 39, p. 673-696.

Maier, W.D., Barnes, S.J., Campbell, I.H., Fiorentini, M.L., Peltonen, P., Barnes, S-J., Smithies, R.H., 2009. Progressive mixing of meteoritic veneer into the early Earth's deep mantle: Nature, v. 460, p. 620-623.

Maier, W.D., Barnes, S-J., de Waal, S.A., 1998. Exploration for magmatic Ni-Cu-PGE sulphide deposits: a review of recent advances in the use of geochemical tools, and their application to some South African ores: South African Journal of Geology, v. 101, p. 237-253.

Mavrogenes, J.A., O’Neil, H.St.C., 1999. The relative effects of pressure, temperature and oxygen fugacity on the solubility of sulfide in mafic magmas: Geochimica et Cosmochimica Acta, v. 63, p. 1173-1180.

McLean, W.H., 1969. Liquidus phase relations in the FeS-FeO-Fe3O4-SiO2 system, and their application in geology: Economic Geology, v. 64, p. 865-994.

Mudd, G.M., 2009, Nickel sulfide versus laterite: the hard sustainability challenge remains. Proceedings 48th Annual Conference of Metallurgists, Canadian Metallurgical Society, Sudbury, Ontario, Canada, August, 2009.

Naldrett, A.J., 1997. Key factors in the genesis of Noril’sk, Sudbury, Jinchuan, Voisey’s Bay and other world class Ni-Cu-PGE deposits: implications for exploration: Australian Journal of Earth Sciences, v. 44, p. 283-315.

Naldrett, A.J., 2005. A history of out understanding of magmatic Ni-Cu sulfide deposits: The Canadian Mineralogist, v. 42, p. 2069-2098.

Ragamani, V., Naldrett, A.J. 1979. Partitioning of Fe, Co, Ni, and Cu between sulfide liquid and basaltic melts and the composition of Ni-Cu sulfide deposits: Economic Geology, v. 73, p. 82-93.

Reck, B.K., Muller, D.B., Rostkowski, K., Graedel, T.E., 2008. Anthropogenic nickel cycle: insights into use, trade, and recycling: Environmental Science and Technology, v. 42, p. 3394-3400

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11

Contents

1.1. Introduction ........................................................................................... 1 1.2. World Nickel Use and Discovery .......................................................... 1 1.3. Nickel Prospectivity and Purpose .......................................................... 2 1.4. Chalcophile Element Mineralization Signatures in Komatiites ............ 3 1.5. Research Scope ...................................................................................... 4 1.6. Thesis Overview .................................................................................... 5 1.7. References ............................................................................................. 9

Chapter 2. Komatiites and Orthomagmatic Nickel Deposits


2.1. Introduction

Nickel sulfide mineralization associated with komatiites was discovered in 1966 at

Kambalda Dome in Western Australia (Woodall and Travis, 1970). This discovery

drastically changed the view of orthomagmatic nickel deposits and prospectivity

within komatiite systems. Prior to this discovery, there was no clear link between

komatiitic rocks and nickel mineralization. Nickel sulfides had been identified in

mafic systems typified by Noril’sk (Russia) and Sudbury (Canada), however,

mineralization hosted within the komatiite-associated systems of Raglan (Canada)

and Pechenga (Russia) was incorrectly characterized. Since 1966 subsequent work

on komatiite-hosted nickel deposits has resulted in a robust mineralization model.

This chapter critically evaluates and discusses the current knowledge of komatiites,

komatiite-hosted nickel deposits, and geochemical targeting techniques derived from

research at Kambalda Dome and other localities. The purpose of this chapter is to

provides necessary background information as a basis for understanding the

subsequent chapters of this thesis. The chapter is subdivided into three sections: (1)

komatiite geochemistry, (2) orthomagmatic mineralization model, and (3)

mineralization indicators. The discussion at the end of this chapter identified

knowledge gaps which will be addressed in this thesis, and provides an initial

framework for more efficient targeting of nickel sulfide mineralization.

2.2. Komatiite Geochemistry and Volcanic Processes

This section on komatiite geochemistry and volcanic processes draws upon both

local Australian examples (e.g. Kambalda Dome) and international locations, in

order to present a comprehensive understanding of komatiitic rocks. Komatiites are

defined and discussed based on the current understanding of: (1) geochemistry (melt

generation, chalcophile elements, crystallization, and contamination); (2) transport

and eruption; (3) volcanic textures (spinifex, cumulates, harrisite, breccia-

volcaniclastic, and vesicles; and (4) volcanic flow field (propagation and field

development, flow thickness, channel and trough, flank, and scale).

11


a. Classification

Komatiites are defined as ultramafic rocks with liquid compositions >18 wt% MgO,

which can be divided into volcanic or sub-volcanic settings based on field

relationships (Arndt and Nesbit, 1982; Arndt et al., 2008). Dendritic olivine and

pyroxene crystal habits referred to as “spinifex texture” are commonly identified

within komatiites; however, these crystal structures are not ubiquitous in all units.

Komatiites are divided into two groups based on MgO content and other

characteristics. The first group includes komatiites characterized by olivine spinifex

texture, with cumulate olivine compositions ranging from Fo89 to Fo94, and liquidus

contents of >18 wt% MgO (Arndt et al., 1977; Arndt and Brooks, 1980). The second

group comprises komatiitic basalts, characterized by liquidus MgO content of <18

wt%, and are dominated by clinopyroxene spinifex texture. Further subdivision of

komatiite types is possible based on major and trace element geochemistry, which is

largely a function of melt generation (discussed below in the Geochemistry Section

2.2b).

b. Geochemistry

Komatiites can be subdivided into three main geochemical groups based on major

and trace element abundances. These geochemical groups consist of Munro-type

(Al-undepleted) and Barberton-type (Al-depleted) as summarized by Kerrich and

Wyman (1996), Lesher and Stone (1996), Arndt et al. (2008); and Karasjok-type

(high Fe-Ti-komatiites) described by Barnes and Often (1990) and Barley et al.

(2000).

The three komatiite types exhibit differing major and rare earth element abundances,

which are mainly controlled by five factors: (1) source region composition, (2)

degree of partial melting, (3) mantle residual phases, (4) crystallizing phases, and (5)

contamination (Lesher and Stone, 1996; Arndt et al., 2008; Fiorentini et al., 2010).

The first three factors are intrinsic to the melt generation stage; whereas, the latter

two (crystallization and contamination) are commonly associated with the final stage

of emplacement.

12


i. Melt generation

Research indicates that komatiites were derived from a mantle source. High MgO

contents (20-30 wt%) measured in fine-grained random spinifex and aphanitic

chilled margins indicate that a mantle-derived magmatic liquid was generated.

Additionally, high forsterite values (Fo90 - Fo94) measured in relict olivine represent

equilibrium conditions with the residues of mantle melting (Lesher et al., 1981;

Redman and Keays, 1985; Arndt and Jenner, 1986). Further evidence for partial

melting of a mantle source is demonstrated by isotopic signatures and the abundance

of rare earth elements (Sun and Nesbitt, 1978; Nesbitt et al., 1979).

Geochemical modelling of Munro-type (Al-undepleted) komatiites indicates partial

melting (25-60% volume) of the source area, with olivine representing the only

residual mineral phase (Sun and Nesbitt, 1978; Nesbitt et al., 1979; Arndt and

Nesbitt, 1982; Lesher and Stone, 1996). This high-degree partial melting and the

presence of residual olivine generated a high MgO melt with chondritic ratios of the

lithophile elements (Al, Ca, Ti, Zr, Y, Hf). Conversely, Barberton-type (Al-depleted)

komatiites are the product of lower degrees of partial melting and formation at

greater depths. Evidence for greater depths of melting is shown in the depletion of:

Al, Sc, Y, and heavy rare earth elements (HREE) relative to Ti, middle rare earth

elements (MREE), light rare earth element (LREE), and large ion lithophile

elements (LILE). The depletion in Al, Sc, Y, and HREE is attributed to stabilization

of majorite garnet at greater pressures, and retention of these elements in the garnet

structure and source area, which resulted in melts with an Al and HREE depleted

signature (Sun and Nesbitt, 1978; Xie et al., 1993; Chavagnac, 2001). A similar

HREE depleted signature is observed in Karasjok-type komatiites which also

reflects a garnet influence in the source area during melt generation (Barnes and

Often, 1990; Barley et al., 2000). Melt generation for komatiitic basalts is estimated

to be lower (15-25% volume) than for Munro-, Barberton-, and Karasjok-type

komatiites. However, these estimates are dependent upon the composition of the

source area, degree of previous melt extraction, and metasomatism (Keays 1982,

1995; Peach et al., 1990; Lesher and Stone, 1996).

Compositionally, mantle source areas vary both spatially and temporally, leading to

geochemical variability in the resultant melts (Zhang et al., 2008; Maier et al.,

13


14

2009). Source areas that have undergone prior melt extraction are depleted in highly

incompatible elements relative to moderately incompatible elements (Th/Smmn <1

and La/Smmn<1; mn subscript denotes mantle normalized: Lesher et al., 2001). The

degree of depletion in the highly incompatible elements is proportional to the

amount of initial melt extracted from the source area. Conversely, source areas that

are enriched in highly incompatible elements from subducted oceanic crust and

associated sedimentary component exhibit an enrichment in highly incompatible

elements (Th/Smmn>1 and La/Smmn>1) relative to moderately incompatible

elements. Consequently, the disparity in LREE contents observed between

Barberton- and Munro-type komatiites is a result of prior low degree partial melting

associated with the Munro-type. Munro-type komatiites exhibit LREE depletion, a

geochemical characteristic generated by previous low degrees of partial melting in

the source area (interpreted to be the extraction of the crust), which is not observed

in Barberton-type komatiites.

ii. Chalcophile elements

Significant nickel mineralization is associated with komatiite systems in Western

Australia, Canada, and Africa (Fig. 2.1: Prendergast, 2003; Sproule et al., 2005;

Barnes et al., 2006b; Arndt et al., 2008). In these regions nickel-copper-platinum

group element (PGE) mineralization (orthomagmatic mineralization) is the product

of fertile melt generation and the effective concentration of the metals upon

emplacement (discussed in Section 2.3).

The chalcophile elements (Ni, Cu, Co, PGE) are strongly partitioned into sulfide

mineral phases, and in the mantle source area the chalcophile elements reside within

the interstitial sulfides (Mitchell and Keays, 1981). The mantle is estimated to

contain 250 ppm sulfur (Sun and McDonough, 1989) and to liberate the maximum

abundance of chalcophile elements, extensive melting (>25%) of the mantle source

area is required generating a sulfur (S)-undersaturated magma. Lower degrees of

partial melting are potentially S-saturated and residual sulfides in the mantle retain

the chalcophile elements and the melt is chalcophile element depleted (Keays, 1982;

1985).


Figure 2.1. World map showing distribution of major orthomagmatic deposits, Ni mineralization districts and geographical locations referenced in this thesis. Komatiite-hosted deposits comprise: Mt. Keith, Perseverance, Black Swan, and Kambalda deposits of Western Australia; Reliance deposit of Africa, and Abitibi Greenstone Belt of Canada. Komatiitic basalt-hosted deposits comprise the Thompson Ni-belt and Raglan Ni-belt of Canada. High MgO basalt deposits are characterized by Noril’sk-Talnakh of Russia, Jinchuan deposit of China, and Kabanga deposit of Tanzania. Ferro-picrite is associated with the Pechenga deposit of Russia. Troctolite is associated with the Voisey’s Bay deposit of Canada. Meteorite impact related deposits are characterized with the Sudbury region of Canada. Large layered intrusions, hosting reef-type platinum group element mineralization, are characterized by the Stillwater Complex of the United States of America, and Bushveld Complex of South Africa. The Karelian Craton of Finland and Norway is included for reference to Karasjok-type komatiites.

15


The chalcophile element abundance in komatiites is controlled by the petrological

history of the source area. Original melt generation models utilized a homogenous

mantle source area. However, current hypothesis of mantle source area compositions

are examining the possibility of a more heterogeneous sources (Sproule et al., 2005;

Zhang et al., 2008; Begg et al., 2009). A heterogeneous source, both spatially and

temporally is the product of previous melt extraction and subsequent secondary

enrichment events. These events lead to potential variability in the chalcophile

element budget as the mineralogy and composition of the source area changes prior

to a melt extraction event (Zhang et al., 2008; Maier et al., 2009). There has been

limited research comparing the chalcophile element contents of komatiitic liquids in

different deposits, greenstone belts, komatiite types, or eruptive ages. Current

research by Maier et al. (2009) indicates that komatiites older than >3.0 Ga exhibit

age dependent PGE abundances, but contain lower PGE contents than komatiites

from 2.9 and 2.7 Ga systems. Although eruptive age (< 2.9 Ga) appears to influence

the PGE budget, Fiorentini et al. (2010) have demonstrated komatiite type and

greenstone belt location do not affect the chalcophile element budget.

iii. Crystallization

The crystallizing phases in ultramafic to mafic magmas exert a strong control on the

distribution of major and trace elements (Lesher and Stone, 1996; Barnes et al.,

2004a; 2006; 2007; Arndt et al., 2008). Olivine provides the dominant control on the

geochemistry in magmas with MgO contents of 20 wt% to 30 wt%. Conversely,

olivine and chromite are the controlling phases in systems with MgO contents of 10

to 20 wt%. In systems with liquid compositions less than 10 wt% MgO, pyroxene,

chromite and plagioclase control the major and trace element distribution. In

komatiite systems (MgO = 20-30 wt%), strong positive correlations are observed

between MgO and compatible elements in olivine (Ni, Co); whereas, strong linear

negative correlations are observed between MgO and incompatible elements in

olivine (Ti, Al, REE, HFSE, Cu, Pd, Pt: Lesher and Stone, 1996; Barnes et al.,

2004a; 2006; 2007; Arndt et al., 2008) .

Crystal fractionation and accumulation also control the REE distribution in

komatiites. Rare earth element abundances are dependent upon the crystallization

and accumulation of olivine, pyroxene, feldspar, chromite, and sulfide, and the

16


partition coefficients of the REE into these mineral phases. Olivine, chromite, and

sulfide represent the dominant crystallizing phases in most komatiite systems; where

chromite and sulfide have a limited effect on the total REE abundance in a rock.

Olivine/liquid partition coefficients for the REEs are quite low (<<0.1), with HREE

having slightly higher partition coefficients than the LREE (Arndt and Lesher,

1992). Consequently, olivine cumulates exhibit considerable variation in

incompatible trace element concentrations, as the abundance of these elements is a

function of the composition and the volume of the trapped interstitial liquid. Even

though the total abundance of incompatible trace elements varies with the proportion

of trapped liquid, the relative concentrations between elements vary only slightly

(Arndt, 1986; Lesher and Arndt, 1995).

Due to the variability in proportions of accumulated crystallized phases in a rock,

liquid magma compositions can only be measured from quenched textures (fine to

very fine-grained spinifex texture and flow top breccias). Quenched textures seldom

contain phenocryst phases, and therefore are the best approximation of liquid

compositions. Advanced spinifex (A2 and A3 as discussed in section 2.2.d.i) are not

preferable as representations of liquid compositions, as these textured rocks contain

a component of accumulated olivine and trapped liquids, resulting in deviation from

the liquid composition (Barnes et al., 1983).

iv. Contamination

Crustal contamination is not observed in all komatiite systems and varies from

minor to extensive with dependence upon physical environmental factors (e.g.

substrate rheology, composition, emplacement dynamics: Lesher and Stone, 1996;

Lesher et al., 2001). Although variable in extent, crustal contamination can greatly

affect the trace element abundances in komatiites. Continental crust is enriched in

highly incompatible lithophile elements (Cs, U, Th, Nb, Ta, LREE) relative to

moderately incompatible lithophile elements (MREE, Y, Zr, Hf, HREE).

Continental crust also exhibits negative Ta and Nb anomalies relative to Th, and

lower Ti enrichment relative to MREE. These anomalies are attributed to retention

in oxide phases during crustal recycling (Weaver and Tarney 1981, Rudnick et al.

1998). This complementary signature between continental crust (LREE enriched)

and komatiites (LREE depleted) results in contaminated komatiites exhibit variable

17


degrees of LREE enrichment with negative Ta and Nb anomalies (Lesher et al.,

2001). Crustal contamination signatures are also characterized by an increase in

La/Sm, La/Yb, and Zr/Y ratios, and a decrease in Nb/La and Nb/Th ratios. Crustal

contamination in komatiite systems (e.g. Black Swan, Kambalda) is physically

recognized by the presence of xenoliths, both macro-(lithic fragments) and micro-

scale (inherited zircons: Barnes et al., 2004b; Compston et al., 1986).

The unique chemical composition of Karasjok-type (high-Fe-Ti) komatiites is

argued to result from contamination and source area composition. Karasjok-type

komatiites are characterized by subchondritic Al2O3/TiO2 ratios with LREE and

HREE depletion, but are enriched in middle rare earth elements (MREE) and high

field strength elements (HFSE), as documented by Barnes and Often (1990);

Lehtonen et al. (1998); Barley et al. (2000); Hanski et al. (2001); and Gangopadhyay

et al. (2006). The subchondritic Al2O3/TiO2 ratios and HREE depletion indicate

residual garnet in the source area and melting at high pressures (Barnes and Often,

1990). Conversely, enrichment of MREE and HFSE result from either the

interaction of the ascending melt with metasomatised or subduction-modified mantle

lithosphere (Barley et al., 2000; Gangopadhyay et al., 2006; Fiorentini et al., 2008a).

The MREE and HFSE enrichments may also result from eclogite contamination

(LREE depleted oceanic crust) of the source area prior to melting (Hanski et al.,

2001; Gangopadhyay et al., 2006).

c. Transport and eruption

Research indicates that komatiites erupted at temperatures up to 1650°C, with high

heat contents (ca. 200cal/g), low viscosities (0.01 Pa/s), and a large temperature

interval between the liquidus temperature and solidus (400°C: Williams et al., 1998;

2001). Numerical modelling of magma flow in komatiites indicates that turbulent

flow occurs at high velocities within confined channels (Williams et al., 1998;

2001). Laminar flow occurs at lower velocities within propagating flow fronts and

thinner flows (Cas et al., 1999). These physical attributes allow for rapid eruption of

large volumes of highly mobile lava, resulting in extensive flow fields (Hill et al.,

1995). Within volcanic flow fields, positive feedback mechanisms control

development. Turbulent flow promotes the thermal-mechanical erosion of the

substrate (Lesher et al., 1984; Arndt, 1986; Groves et al., 1986; Greenley et al.,

18


1998; Williams et al., 1998; 2001) which leads to the progressive development of

channelized flow and increasing distances of sustained turbulent flow.

d. Volcanic textures

Volcanic textures within extrusive volcanic rocks are a function of: (1) initial

magma composition, (2) flow dynamics, (3) cooling rates and thermal gradients, and

(4) availability of nucleation sites. These variables control the mineralogy,

morphology, and abundances of minerals (via accumulation and fractionation)

observed in ultramafic rocks (Arndt, 1986). As discussed previously in section 2.2.a,

classification of the two main types of komatiites is based on measured MgO

contents. The highest MgO content lavas (MgO > 18 wt%: komatiites) are

characterized by olivine spinifex texture, with interstitial clinopyroxene and

cumulate lithologies dominated by equant olivine. Lower MgO content lavas (MgO

< 18 wt%: komatiitic basalts) are characterized by equant or platy skeletal crystals

of olivine in a groundmass of devitrified glass or skeletal clinopyroxene without the

presence of plagioclase. The following discussion of volcanic textures is restricted to

komatiites (MgO >18%). Similar textures are observed in komatiitic basalts (MgO

>18%) with lower temperature mineral phases (e.g. pyroxene) replacing skeletal

olivine in the spinifex zones (Arndt et al., 1977; 1979). Both systems display a wide

range of igneous textures, from quenched liquids to mineral cumulates. Despite this

range in textures the primary mineralogy is simple in komatiite systems, with olivine

± chromite representing the dominant crystallizing and accumulating mineral phases

(Arndt, 1986).

i. Spinifex

Spinifex texture is characterized by bladed or acicular dendritic olivine or pyroxene

crystals, which were first described from the Barberton komatiites of South Africa

(Viljoen and Viljoen, 1969). Spinifex texture is commonly observed at the top of

extrusive ultramafic lavas flows, and more rarely within intrusive bodies as spinifex

textured intervals (Donaldson, 1974; Arndt et al., 2004). Since the first complete

textural descriptions of spinifex texture (Pyke et al., 1973), extensive research has

classified, interpreted and modelled the formation of spinifex texture. Spinifex

texture is currently subdivided into flow top breccia and three zones: A1, A2, and

19


A3 (Fig. 2.2). Theses subdivisions are a function of variable thermal gradients

within a cooling ultramafic magma.

Flow top breccia is characterized by well-sorted, subangular to amoeboid clasts

ranging in size from 0.5 to 1 cm. Texturally, the fragments range from

microporphyritic with olivine phenocrysts (1-3%) to micro-spinifex with glass.

Olivine microlites and small amygdules (1-2%) are periodically observed (Arndt,

1986). Common alteration minerals include: chlorite, magnetite, serpentine and

tremolite.

A1-spinifex, representing the chilled margin, is characterized by fine-grained

devitrified glass, and commonly exhibits polyhedral joint sets with rare olivine

phenocrysts (Pyke et al., 1973). A1-spinifex commonly forms a thin zone from 1 to

10 cm thick (Fig. 2.2). Petrographically small (0.1-0.5 mm) and sparse olivine

phenocrysts (1-2%) are observed within the fine-grained groundmass along with

small hopper olivine crystals. Compositionally, olivine exhibits a restricted range of

Fo94.1 with more fractionated rims (Fo88), as documented in the Munro Township

flows (Arndt et al., 1977).

A2-spinifex, known as random spinifex, consists of fine (~2 mm) skeletal and

dendritic olivine crystals which are randomly orientated in a matrix of fine-grained

skeletal pyroxene, skeletal and equant chromite, and devitrified glass (Pyke et al.,

1973; Arndt, 1986). A2-spinifex commonly forms a zone 5 to 50 cm in thickness

(Fig. 2.2). Compositionally, olivine varies from Fo93.5 to Fo93 in the cores and

exhibits continual zonation with Fe-enrichment toward the margins (Fo87: Arndt,

1986). The formation of A2-spinifex occurs after the generation of a thin aphanitic

crust or chilled margin (A1-spinifex) during progressive cooling under a steep

thermal gradient (Faure et al., 2006).

20


Figure 2.2. Diagram illustrating fully differentiated komatiite flow with upper A-zone spinifex and lower B-zone olivine cumulates. Modified from Pyke et al. (1973) and Arndt et al. (1977).

A3-spinifex, represented by orientated, bladed, and chevron spinifex, is

characterized by coarse-grained platy olivine. The olivine crystals are dominantly

elongate along the C-axis and are orientated perpendicular to the flow top (Fig. 2.2).

The transition from overlying A2-spinifex to A3-spinifex is commonly sharp, but

gradational. A3-spinifex can form zones up to 3m thick, with individual crystals up

to 1m in length (Pyke et al., 1973). Olivine crystals are commonly skeletal and range

in thickness from 0.02 to 2.0 mm, forming booklets of 0.3 to 15 cm in thickness.

A3-spinifex contains approximately 60 to 65% olivine, with inter-blade areas

consisting of discrete and elongate skeletal crystals of clinopyroxene and devitrified

glass (Pyke et al., 1973). Compositionally, olivine has cores of Fo93.9 and

fractionated margins of < Fo86 (Arndt, 1986). The formation of A3-spinifex is a

continuation of A2 olivine crystallization, preferentially favouring the growth of

21


olivine crystals perpendicular to the cooling surface, parallel to the direction of heat

loss (Huppert and Sparks, 1985; Shore and Fowler, 1999; Faure et al., 2006).

Hypotheses related to the formation of spinifex texture in ultramafic magmas and

lavas has undergone continual revision from initial ideas of super cooling (Drever

and Johnston, 1957), directional super cooling (Donaldson, 1982; Huppert et al.,

1984 and Arndt 1986), superheated magma (Aitken and Echeverria, 1984 and Arndt,

1994), magma degassing (Donaldson, 1974), and constitutional supercooling due to

degassing of originally high water contents (Parman et al., 1997). Current models for

the formation of spinifex textures rely on constant thermal gradients and under-

cooling (Shore and Fowler, 1999; Faure et al., 2006).

Komatiite systems differ from mafic systems in terms of liquidus and solidus

temperatures. Temperature differences up to 500°C between the liquidus and the

solidus have been inferred in komatiite systems. This large temperature interval of

crystal free magma is not observed in mafic systems (Arndt, 1994). Initial cooling

upon emplacement is rapid in komatiites, leading to the generation of A1-spinifex

through direct quenching of the ultramafic magma with seawater. This sustained

rapid heat loss from quenching (>50C°/hour) continues until the crystallizing crust

has reached a thickness >1 m, leading to the formation of A2-spinifex or random

spinifex. Upon reaching a crust thickness ≥1 m, the cooling rate within the komatiite

flow drops substantially, as heat loss is primarily due to conduction through the crust

(Faure et al., 2006). Therefore, the thermal gradient established from conductive

cooling within the flow is critical for the formation of A3-spinifex. The formation of

A3-spinifex was as modelled by Faure et al. (2006), where orientated bladed olivine

crystals were generated using thermal gradients of from 25°C/cm to <10°C/cm with

low cooling rates of 2°C/hr to 5°C/hr.

ii. Cumulates

Cumulate lithologies characterize the lower portion of differentiated komatiite flows

(e.g. upper spinifex and lower cumulate: Fig. 2.2) and occur in areas of sustained

flow (e.g. magma channels). Cumulate mineralogy exhibits a range in composition

that is dependent upon the initial magma. Komatiite systems are dominated by

olivine, komatiitic basalts are characterized by olivine and pyroxene, and basaltic

22


differentiates contain olivine, pyroxene, and feldspar. Cumulates are indicative of

sustained magma flow in both open and closed systems and form through two

processes. The first process is gravitational crystal accumulation, as proposed by

Wagner et al. (1960). Crystal accumulation results from the slow settling (up or

down) of crystallizing mineral phases, due to density contrasts between the mineral

and the magma. The second process is in-situ nucleation and crystal growth

(Campbell, 1978; McBirney and Noyes, 1979; McBirney and Hunter, 1995; Barnes

and Hill, 1995). Cumulates developed by this method grow in-situ at the interface

between the crystal pile and the magma, and are continuously exposed to flowing

magma, which facilitates chemical exchange.

Olivine cumulate textures display a range of crystal morphologies from equigranular

polyhedral olivine to irregular hopper morphologies, and a wide distribution in

crystal size. Cumulates exhibit a continuum of crystal packing densities based on the

varying abundance of crystals and the presence of intercumulus liquid. Cumulate

lithologies are divided into three groups: adcumulate, mesocumulate, and

orthocumulate. These cumulate divisions are based on the proportion of cumulate

crystals relative to the proportion of trapped liquid. Adcumulates have little or no

intercumulate liquid and are commonly mono-mineralic. Adcumulates are also

thought to represent prolonged periods of turbulent lava flow, where crystallization

of olivine occurred close to the liquidus temperature at the top of the cumulate pile

(Barnes and Hill, 1995). Mesocumulates occur between the adcumulate and

orthocumulate end members with extensive mutual crystal boundaries and minor

intercumulate liquid. Orthocumulates exhibit a high proportion of trapped

intercumulate liquid. Both mesocumulate and orthocumulate lithologies, containing

high proportions of intercumulus liquid, are hypothesized to form during two

processes. The first process is in-situ crystallization under laminar flow with

increasing degrees of super cooling (Hill et al., 1995). During this process,

crystallization occurs at the top of the crystal pile, sealing off and trapping the

intercumulus liquid before it can be removed (Hill, 2001). The second process is

gravitational sedimentation of olivine crystals from flowing and ponded magma, as

proposed by Wagner et al. (1960) and extended to komatiite systems by Hill et al.

(1995) and Hill (2001).

23


Cumulate lithologies in extrusive komatiite flow systems are divided into four

groups based on igneous texture. Paralleling the spinifex zones, cumulates comprise

B1-, B2-, B3- and B4-zones (Fig. 2.2). However, not all zones are observed within

one flow, and variability along strike within the same flow is common (Pyke et al.,

1973; Arndt et al., 1977). Changes in cumulate mineral abundance results in a

continuum of possible volcanic facies (Fig. 2.3), as observed and outlined by Lesher

et al. (1984), Lesher (1989), Lesher et al. (1999), and Lesher and Keays (2002).

The B1-zone (foliated skeletal olivine) is characterized by tabular olivine crystals

with a more apparent skeletal habit to hopper-bladed crystals that are orientated

parallel to the top of the flow and parallel to the plane of flow (Pyke et al., 1973;

Arndt, 1986). A maximum thickness B1 cumulates is 30 cm as documented in

Munro Township komatiites (Arndt et al., 1977). The contact with the overlying A-

spinifex zone is characterized as abrupt, irregular, and being the most distinctive

internal contact in a komatiite flow (Pyke et al., 1973). The lower contact with the

B2-zone is gradational and rapid (Pyke et al., 1973). The B1-zone is not always

observed within a differentiated flow and can change rapidly along strike in a flow

(Fig. 2.2).

The B2-zone, composed of peridotite is characterized by equant olivine crystals,

with a matrix of skeletal clinopyroxene and cruciform, dendritic or euhedral

chromite. Elongate, partially skeletal olivine grains are minor and commonly occur

sub-parallel to flow direction (Arndt et al., 1977). Foliation within the B2-zone is

more developed towards the top of the zone. The B2-zones also exhibits a decrease

in crystal size towards to the basal contact. Compositionally, this zone is composed

of 70-75% olivine with lesser interstitial clinopyroxene and minor chromite.

24


Figure 2.3. Komatiite cooling units matrix with increasing olivine accumulation on left and increasing differentiation along the bottom axis. UN = undifferentiated non-cumulate (massive, pillowed or volcaniclastic), DN = differentiated non-cumulate, UC = undifferentiated cumulate, DC = differentiated cumulate. Modified from Lesher and Keays (2002).

The B3-zone (knobby peridotite) ranges in thickness from 15 to 35 cm and occurs in

the central-lower portion of the B-zone (Fig. 2.2: Pyke et al., 1973). The B3-zone is

well defined in continuous flow units and poorly defined, patchy or absent in other

flow units. The B3-zone exhibits a gradational contact with the adjacent B2- and B4-

zones and is defined by the presence of 10-20% small (~2 mm) semi-round

protuberances. Mineralogically, the B3-zone typically consists of ~65% olivine and

5% clinopyroxene within a fine-grained matrix.

The B4-zone (basal peridotite: Fig. 2.2), has similar moderate foliation as the B2-

zone. One notable difference between the B2- and B4-zones is the presence of a

narrow chilled margin (~ 1 cm wide), at the contact between the B4-zone underlying

komatiite flow unit (Pyke et al., 1973).

25


Large ultramafic komatiite bodies (>25m) dominated by cumulate lithologies (e.g.

Mt. Keith: Rosengren et al., 2007) are not subdivided into B-zones. These thick

komatiite units are typically described based on the proportions of olivine and

interstitial liquid (adcumulate, mesocumulate, orthocumulate) and other

distinguishing textural and mineralogical attributes.

iii. Harrisite

Harrisitic texture is characterized by exceptionally large olivine crystals with

branching crystal morphologies and parallel growth habits (see Donaldson [1974]

for a comprehensive harristic texture description from the Rhum Intrusion). In

komatiites, harristic texture is typically found within the adcumulate sequences and

is thought to represent nucleation and rapid crystal growth due to directional cooling

in a supersaturated liquid (Hill et al., 1995; Hill, 2001). Harrisitic texture is

commonly identified overlying olivine adcumulates. Based on this association, it is

thought that harrisite texture marks the transition from continuous turbulent flow in

a system, to laminar and stagnating flow (Hill et al., 1995).

iv. Breccia-volcaniclastic

Breccia and volcaniclastic rocks are scarce in komatiite systems. Breccia textures

commonly are either flow top breccias or hyaloclastite as described on thick flows in

Munro Township (Arndt et al., 1977). Volcaniclastic komatiite rocks are even more

rare and interpreted to form under unique conditions. It is argued that the low

volatile content, and low viscosity of ultramafic magmas limits the abundance of

explosive lava eruptions in komatiites (Arndt, 2008). Furthermore, the deep

submarine environment proposed for the eruption of most komatiites is thought to

inhibit phreatomagmatic brecciation (McPhie et al., 1993).

Ultramafic volcaniclastic lithologies (e.g. lapilli tuff, accrectionary lapilli tuffs,

komatiitic volcanic breccia) have been identified in a limited number of greenstone

belts (see Table 2.1). Ultramafic volcaniclastic rocks are documented in the

Barberton Greenstone Belt (South Africa); Quetico Subprovince and Abitibi

Greenstone Belt (Canada); Ruth Well, Scotia Ni-deposit and Meekatharra-Wydgee

Belt (Australia); and Karasjok and Kittilä Greenstone Belts (Finland-Norway). The

26


latter three settings (Meekatharra-Wydgee, Kittilä and Karasjok) and several

occurrences in Quetico Subprovince are Karasjok-type komatiites (Fe-Ti-enriched).

Volcaniclastic textured units are characteristic of Karasjok-type komatiites (Barnes

and Often, 1990; Barley et al., 2000; Gangopadhyay et al., 2005; Goldstein and

Francis, 2008).

Table 2.1. Greenstone belts containing volcaniclastic textured ultramafic lithologies. Barberton-type komatiite (B-type), Munro-type komatiite (M-type), Karasjok-type komatiite (K-type). BIF = banded iron formation, Int. Vol. = intermediate volcanics, Metased. = metasedimentary rocks.

Belt/Area Age (Ga) Komatiite type

Strat. Associations

Reference:

Barberton 3.0-2.9 B-type Chert, Evaporite

Stiegler et al., 2008

Quetico 2.78 B-type

K-type

Carbonates, BIF, Int. Vol.

Schaefer and Morton, 1991; Fralick et al. 2008

Abitibi 2.7 Not reported Not reported Mueller et al., 2006; Gelinas et al., 1977.

North Spirit 3.0 Not reported Metased., komatiites, BIF

Houlé et al., 2008.

Ruth Well 3.5 M-type Komatiite Nisbet and Chinner, 1981

Meekatharra-Wydgee belt

3.0-2.9 K-type Komatiite Barley et al., 2000

Scotia Ni-deposit 2.7 M-type Komatiite Page and Schmulian, 1981; Stolz and Nesbitt, 1981

Karasjok-Kittilä 2.0-1.9 K-type Komatiite Barnes and Often, 1990; Saverikko, 1985; Gangopadhyay et al., 2006

It should be noted that volcaniclastic units within the Scotia Ni-deposit are suspect,

since a broad spatial relationship between fragmentals and thick olivine cumulates

was documented by Page and Schmulian (1981). In addition, intercalated

metasedimentary rocks are missing from local stratigraphy, yet present in the rest of

the sequence. Based on these observations, it is speculated that these fragmental

textures represent flow auto-brecciation related to episodic-fluctuating and pulsating

flow regimes.

27


v. Vesicles

Vesicles are not ubiquitous to komatiites, but also not uncommon. Komatiite units

hosting vesicles are identified within the Barberton Greenstone Belt, South Africa

(Dann, 2001); Murphy Well, Australia (Lewis and Williams, 1973); Kambalda

Dome, Australia (Beresford et al., 2000; 2002; 2005); Black Swan, Australia (Hill et

al., 2004); Scotia Ni-deposit, Australia (Stolz and Nesbitt, 1981); Lake Johnston

Greenstone Belt, Australia (Heggie et al., 2007); Abitibi Greenstone Belt, Canada

(Stone et al., 1996); and Dismal Ashrock, Canada (Schaefer and Morton, 1991).

The presence of vesicles in komatiites is controversial, as they indicate the presence

of volatiles within the magma. Initial volatile contents within ultramafic magmas is

interpreted to be low, due to the hot and anhydrous nature of the proposed source

area (Arndt et al., 1998). Arguably, ultramafic magmas can gain volatiles through

assimilation of hydrated contaminants (Black Swan: Hill et al., 2004; Freds Flow,

Abitibi Greenstone belt: Stone et al., 1996). Recent research on hydrous minerals

(amphibole) in both komatiites and more fractionated systems supports the presence

of some volatile content in primitive magmas (Stone et al, 1997; Fiorentini et al.,

2008a). Although, a low volatile content (<0.5%) can produce high volumes of

vesicles, komatiite systems typically contain <2% vesicles. The discrepancy

between volatile content and the abundance of vesicles is interpreted to be controlled

by the depth of eruption and confining pressure.

e. Volcanic flow field

Komatiite flow fields are diverse and complex volcanic settings. There are no

modern analogs of lava flows with similar extreme eruption temperatures (~1600C)

and low viscocities. However, modern day ocean island basaltic eruptions (e.g.

Hawaii) have been used to provide insight into komatiite flow field development

(channelized flow, inflation, lobe budding: Hill, 2001). These observations are

supported by geochemistry, rock types and interpreted flow facies identified in

komatiite flow fields (Gresham and Loftus-Hills, 1981; Lesher, 1983; Hill et al.,

1995).

28


i. Propagation and field development

The variations in komatiite volcanic facies and flow morphologies form a continuum

that is dependent upon the eruption rate, host environment, and proximity to source.

Komatiite flow can be divided into three zones based on the inferred proximity to

the source: proximal, medial, and distal. Numerical modelling identifies proximal

flow as dominantly turbulent (Huppert et al., 1984; Huppert and Sparks, 1985). Non-

channelized flow in medial to distal areas is dominantly laminar (Cas et al., 1999).

Conversely, channelized flow in medial to distal areas is turbulent, based on Ni

mineralization models and the necessary thermal-mechanical erosion of footwall

contaminants (Lesher, 1983).

Hill et al. (1995) and Hill (2001) developed a flow field model based on the

assumption that continuous flowing lava will tend to form cumulate dominated

lithologies; whereas, episodic flow will lead to the generation of differentiated

flows. The time sequence volcanic model by Hill (2001), identifies initial volcanic

activity in the form of continuous unconstrained eruption, that results in the

formation of proximal sheet flows and the continuous switching of developing

channels (Fig. 2.4). Once the direction of preferred lava flow is established,

dependent upon slope and pre-existing topography, sustained lava channels develop

within the flow field. Concurrent with channel development, flanking environments

develop through progressive channel breakouts and inflationary advances. This

results in a complex stratigraphy consisting of thin to thick, differentiated and

undifferentiated flows (Fig. 2.5).

The preceding komatiite flow field model concludes that at very high eruption rates,

sheet and channelized flow would dominate, forming an extensive compound flow

system. This system would be characterized by thick mesocumulate to adcumulate

bodies within channels, and a combination of thick undifferentiated and thin

differentiated flows in more distal areas. At lower eruption rates, komatiite flow

fields would be characterized by thin differentiated flows with episodic development

of channelized flows and adcumulate bodies.

29


Figure 2.4. Komatiite flow field model as proposed by Hill (2001) showing the transition from massive sheet flow to channelized flow. Modified from Arndt et al. (2008).

30


Figure 2.5. Komatiite flow field model as proposed by Hill (2001) showing lobe development at the advancing front and lateral development. Modified from Arndt et al. (2008).

31


Structure has some control on the development of komatiite flow fields. This is

based on observed lithological distribution and identification of major structures

within flow fields. Major N-NW faults in the Kambalda region of Western Australia

were interpreted to have been active during the Archean and influenced sedimentary

basin development and the emplacement of magmas (Horwitz and Sofoulis, 1965;

Williams, 1970; Gee, 1979 and O’Driscoll, 1981). Many Ni deposits in the Eastern

Goldfields Superterrane of Western Australia occur west of major faults, as

documented in the Kambalda area where the Kambalda Dome mineralization occurs

west of the Lefroy fault (Gresham and Loftus-Hills, 1981).

ii. Flow thickness

Komatiite flows are divided into thee types based on flow thickness: thin, thick, and

very thick flows (Hill et al., 1995). Flow thickness is associated with flow lithology

and textural variability (as shown in Fig. 2.3) and used to identify flow facies. Flow

facies divisions are controlled by: (1) ponding and differentiation, and (2) flow

through and olivine accumulation (Barnes, 2006).

Thin flow facies are characterized by flows < 25m thick, and are more commonly

0.5m to 10m in thickness, ranging in width from 10s to 100s of metres. Thin flows

can be either differentiated or undifferentiated. Differentiated flows contain well

developed A and B zones, reflecting rapid cooling and crystallization under stagnant

conditions (Hill et al., 1995). Undifferentiated flows are entirely composed of either

spinifex texture or B-zone cumulates (Fig. 2.3). Undifferentiated spinifex flows are

thought to occur in distal environments, whereas undifferentiated B-zone flows

occur in crystal laden lava tubes (Hill et al., 1995; Barnes, 2006).

Thin flow facies can also form as a complex lava tube system, with continual

budding of new flows and lobes during system propagation (Barnes, 1985). Budding

flow lobes from the Munro Township are characterized by large 2 to 6m wide lobes

that are roughly cylindrical with a convex upper surface and a concave lower

surface, with small cusps where they have conformed to underlying flows. These

lava lobes exhibit both spinifex texture and spinifex devoid margins with basal

accumulations of equant olivine crystals (Arndt et al., 1977).

32


Thick flows represent flows that range in thickness from 25 to 150m, and are up to

several kilometers in width. Thick flows are characterized as pathway subfacies and

differentiated cumulate subfacies, depending upon the extent of textural

differentiation (Barnes, 2006). Differentiated cumulate subfacies are the product of

ponded komatiite lava undergoing in-situ differentiation. This results in the

commonly observed sequence of basal olivine accumulation, overlain by pyroxenite

and gabbroic cumulates, with spinifex textures occurring at the top (Hill et al., 1995;

Barnes, 2006). Pathway subfacies are thick undifferentiated flows dominated by

cumulate olivine. The composition ranges from orthocumulate to mesocumulate,

with thin spinifex zones occurring at the top of the sequence. The development of

pathway subfacies is attributed to continued channelized magma flow in a

developing komatiitic flow field (Lesher et al., 1984).Within the Silver Lake

Member of the Kambalda Komatiite Formation, pathway subfacies characterize the

“channel” portion of the basal flows as described by Lesher (1983) and Lesher et al.

(1984). The Kambalda channel facies comprise olivine cumulates up to 150m thick

and 200m wide, forming a linear sinuous body. Thin flow facies define the margins

of the channel and form the flanking environment (Lesher et al., 1984).

Very thick flows include ultramafic units that range in thickness from 150 to 1000m,

and are characterized by a preponderance of olivine cumulates with olivine-chromite

adcumulates forming the core of many igneous bodies (Hill et al., 1995). Very thick

flows form dunite lens or sheet subfacies and dunitic differentiated cumulate

subfacies depending upon the degree of differentiation (Barnes, 2006). Recent

research on these very thick flow units has led to their reclassification as intrusive

bodies, rather than extrusive (Rosengren et al., 2005).

iii. Channel and Trough

Komatiite magma flow channels, equivalent to modern day lava tubes and feeders

are identified at the base of the komatiite sequence at the Kambalda Dome and

within other Archean komatiite flow fields. Channels are characterized as thickened

linearly continuous cumulate bodies (Fig. 2.6). Channels are characteristically

thicker (up to 150m) than the adjacent flanking flows (< 25m) and dominated by

olivine cumulate rocks. Both the greater thickness and olivine accumulation are the

result of sustained magma flow through in the channel. The mechanism for the

33


development of thickened channels is still debated, but, two end-member hypotheses

for the formation of channels are proposed. The first end-member model involves

pre-existing topography to control the development of channels and channelized

flows (Lesher et al., 1984). The second model states that channels represent a

positive feedback from thermo-mechanical erosion of the substrate, thus

perpetuating the development of channels (Huppert et al., 1984). Regardless of the

chosen model for formation of thickened channels, channelized flow provides the

sustained transport of primitive lava through the channel system to the advancing

flow front.

Figure 2.6. Idealized schematic cross-section showing both channel and flank facies with associated sediments and Ni-sulfide mineralization as observed at the Kambalda Dome. Modified from Cowden and Roberts (1990).

Sustained magma transport results in channels having a greater degree of olivine

accumulation, as flow through provides a steady influx of fresh magma crystallizing

and accumulating olivine. Sustained magma flow through also results in more

34


primitive mineral and whole rock compositions being hosted within the channels.

Primitive lavas with >20% MgO are chromite undersaturated and only crystallize

olivine, whereas, more fractionated lavas with <20% MgO are chromite saturated

and will crystallize both olivine and chromite at cotectic proportions of 50:1 (Muruk

and Campbell, 1986). Consequently, channels are relatively depleted in Cr with

lower Cr/Mg and Cr/Ni ratios than more fractionated flanking environments (Barnes

and Brand, 1999; Lesher and Groves, 1984; Lesher and Arndt, 1995).

Although a number of similarities are observed in the function, morphology and

chemical relationships of Archean and modern lava tube systems, the formation of

the channels may have differed slightly (Hill et al., 1995). Modern lava tube systems

form by the gradual roofing through inward solidification, inflation, and downward

growth of crust on the centralized flow. The width that roofing occurs is limited due

to the weight of the roofing material. Consequently, it is proposed that in komatiitic

channels, lava would have flowed in direct contact with seawater, as any crust

forming would have been dense and sunk back into the turbulent lava (Barnes et al.,

1983; Hill, 2001).

iv. Flank

Flank environments vary in lithology from thick undifferentiated flows, to thin well-

differentiated flows, and differentiated ponded lava lakes (Figs. 2.3 and 2.6). Flows

in the flanks are thin (< 25m thick), contain interflow metasedimentary rocks, and

lack Ni-sulfide mineralization (e.g. lower flows of the Kambalda Dome). Flanks are

also characterized by lower MgO content, and higher Cr, Ti, Al, Fe, and Zn contents

than the channel environments. This distinct geochemistry is likely a result of the

more evolved lava compositions (Lesher, 1989).

Flank flow development occurs through two processes: (1) initial sheet flood flow,

and (2) lateral breakouts (Figs. 2.4 and 2.5: Hill et al., 1995; Hill, 2001). Initial

sheet flood flow is the first outpouring of magma and results in the generation of a

thin laterally continuous basal flow. The establishment of channelized flow results in

channel inflation leading to the second process of flank flow development; lateral

breakouts. Lateral breakouts occur along the channelized flow and form small lava

35


lobes that either coalesce and form a continuous sheet or remain as discrete

individual lobes.

Detailed work by Beresford et al. (2002), within the Long-Victor channel of the

Kambalda Dome correlated channel and flank stratigraphy within the basal flow

units of the Silver Lake Member. Above the basal unit, flows are the result of

breakouts and cannot be correlated, as supported by the presence of thickened

interflow sediments.

v. Scale

The relative size (scale) of komatiite magmatic systems is largely unconstrained in

the rock record. Distal lava fronts are poorly outlined, and eruptive centres with

proximal facies and feeder systems have rarely been identified within the sequences.

Small lava lobes intercalated with sedimentary-volcaniclastic material are argued to

define terminal environments or low flow rates (Cas et al., 1999; Arndt et al., 2008).

However, this facies association is not observed in the Kambalda Dome komatiites,

or in any other high MgO-system. The most “distal” facies observed in the Eastern

Goldfields Superterrane are thin differentiated flows, with or without intercalated

sedimentary-volcaniclastic material, as observed in the Tripod Hill Member of the

Kambalda Komatiite Formation. However, rather than a distal setting, declining

eruption rates are proposed for the flow facies, as extensive proximal channelized

flows are identified stratigraphically below in the Silver Lake Member.

Proximal facies are characterized by heterogeneous pyroclastic deposits consisting

of heterolithic fragments ranging from course- to fine-grained, lava lakes, and

magma feeder systems (Arndt et al., 2008). The only known komatiite example of

heterogeneous heterolithic pyroclastic deposits is the Dismal Ashrock of the Steep

Rock-Lumby Lake greenstone belt in Canada (Table 2.1: Schaefer and Morton,

1991). Yet, there are no komatiite flows identified within the Steep Rock sequence.

Proximal facies are identified within the Scotia Ni-deposit of Western Australia,

where thick breccia-textured units are associated with thickened olivine cumulates

(Page and Schmulian, 1981; Stolz and Nesbitt, 1981). A proximal setting is also

represented by evidence for simultaneous komatiite and dacite eruptions within the

Boorara Domain, Western Australia (Trofimovs et al., 2004). A possible intrusive

36


feeder system is also identified within the Reliance komatiites of Zimbabwe

(Prendergast, 2003).

Thickened olivine cumulate bodies are the product of sustained high-volume flow

rates, and can occur both proximally and distally. Therefore the presence of

thickened olivine cumulate bodies is ambiguous in terms of system scale. Within the

Eastern Goldfields Terrane of Australia it is unknown if the thickened olivine

adcumulate bodies documented between Agnew and Norseman, Australia represent

multiple eruptive sites along the 2.7 Ga system, or if these adcumulate bodies are the

product of a single point source within a laterally and linearly extensive komatiite

flow field containing channelized flow.

The scale of komatiite flow fields have been estimated by several researchers. Hill et

al. (1995) proposed a fractal approach to estimate the size of komatiite flow fields.

Individual flow lobes are identified on the scale of 10m, compound flows potentially

kilometers in size and perhaps >100 km for single cooling units, as suggested by the

laterally continuous Walter Williams Formation (Hill et al., 1995). Barnes et al.

(2007) hypothesized that channelized facies within the Scotia-Kambalda-St. Ives-

Widgiemooltha areas of Western Australia, represent the distal equivalents to the

more proximal facies (intrusive bodies) within the Agnew-Wiluna area. The scale of

the identified proximal and distal facies resulted in a flow field that exceeds 500 km

N-S and 150 km E-W. Work by Prendergast (2003) suggests that extrusive

komatiitic-basalts of the Reliance flow sequence of Zimbabwe extend linearly for 85

km, with the greatest width (2500m) observed in the central portion of the sequence.

2.3. Orthomagmatic Mineralization Model

Nickel sulfide deposits hosted within komatiites are an important resource for world

Ni supply. Consequently, exploration for additional Ni deposits is extensive. A

conceptual model for the generation of komatiite-hosted Ni deposits is critical in the

exploration process. The following mineralization model is derived from the

Kambalda Dome area of Western Australia and other Ni mineralized systems around

the world. This model is presented in a generalized and process orientated fashion to

allow for application in a variety of settings that may deviate from the “Kambalda-

type”.

37


The ore forming process for komatiite-hosted orthomagmatic Ni-Cu-PGE

mineralization is solely dependent upon sulfur saturation within the system, and the

development of an immiscible sulfide liquid (Campbell and Naldrett, 1979; Naldrett,

1981; Campbell and Barnes, 1984). Processes leading to sulfur saturation in Ni

deposits are still under debate amongst researchers (Stone et al., 1996; Lesher et al.,

2001; Fiorentini et al., 2006; Barnes, 2006b). The generally accepted model involves

the assimilation of a crustal contaminant, either sediments or felsic volcanics

underlying an area that contains a variable abundance of sulfur. Within the

Kambalda Dome setting, research indicates that sediments were originally located

beneath the flow channels, where thermal-mechanical erosion at the base of the

channels incorporated sulfidic sediments, causing localized sulfur saturation (Lesher

et al., 1984; Lesher et al., 2001). Mass independent S-isotope fractionation analyses

have linked exhalative S with mineralization, thus supporting the theory of

underlying sediment assimilation (Bekker et al., 2009). Similarly, most

orthomagmatic deposits globally and temporally involve the assimilation of crustal

S. An exception to this model is the Nebo-Babel deposit in the Musgrave block of

central Australia, where a crustal sulfur source is not observed, but rather a mantle

source (Seat et al., 2008).

Once an immiscible sulfide is present within the magmatic system, the chalcophile

elements (Ni, Cu, and PGE) preferentially partition into an immiscible sulfide phase

over the silicate phase. The strong partitioning of the chalcophile elements into the

sulfide phase depletes the interacting silicate magma in chalcophile elements.

Continued interaction between a metal-bearing silicate liquid and an immiscible

sulfide liquid causes the latter to become progressively more enriched in chalcophile

elements, as described by the R-factor model (Campbell and Naldrett, 1979;

Campbell and Barnes, 1984). Sulfide-silicate liquid interaction continues until the

sulfide liquid is isolated from the lava through gravitational settling and the silicate

liquid is solidified through crystallization.

Channelized flow is also an important component in the generation of

mineralization, as it provides both turbulent and sustained flow, which results in

thermal-mechanical erosion of the sulfidic substrate (Fig 2.4: Groves et al., 1979;

1986; Huppert et al., 1984; Lesher et al., 1984; Huppert and Sparks, 1985; Frost and

38


Groves, 1989; Lesher, 1989; Williams et al., 1998; 1999; Lesher and Burnham,

1999), and effective mixing between the silicate and sulphide liquids.

Primary economic Ni mineralization in komatiite systems is divided into two

morphological types: Type-1 massive sulfide accumulations at the base of

channelized flow conduits (Fig. 2.6: komatiite flows or sub-volcanic feeders), and

Type-2 disseminated sulfide associated with large dunite bodies (Lesher, 1989;

Lesher and Keays, 2002). Although distinct in morphology, the two deposit types

may form a continuum of mineralization types (Barnes et al., 2007).

Type-1 Ni deposits are sulfide-rich and are dominated by massive ore (75-100%

sulfide), with lesser matrix/net-textured ore (20-70% sulfide) grading laterally into

minor disseminated sulfide. The type example of this style of mineralization is the

Kambalda Dome, as summarized by Barnes (2006b). Type-1 ores commonly have

high Ni tenors and low Cu tenors, as defined by concentration of the metal

normalized to 100% sulfides. These ores also exhibit extensive grade variability

between ore bodies and deposits. Nickel grade variability between ore shoots is

largely interpreted to represent primary mineralization signatures (Ross and Keays,

1979; Woolrich et al., 1981; Keays et al., 1981; Barnes, 2004b). Type-1 ores form

early in the development of the volcanic field, during localized sulfur saturation

within the channelized flow conduit, followed by sulfide transport and deposition

within the channel.

Type-2 deposits are characterized by homogenous low-grade disseminated sulfides

hosted within thickened linear olivine cumulate bodies (Lesher, 1989). The type

examples of this mineralization style are the Mt. Keith and Yakabindie deposits of

Western Australia (Grguric et al., 2006). Nickel grades are typically homogeneous,

dominantly <1%, commonly ~ 0.6%, and are associated with 1-5% sulfide. The

relatively high abundance and interstitial position of the sulfide precludes formation

by trapped liquid and interstitial precipitation of Ni-sulfides. Barnes (2007) proposed

a model of non-cotectic precipitation and physical transport of pre-existing sulfide.

39


a. Sulfur in orthomagmatic nickel systems

Sulfur plays a crucial role in the development of orthomagmatic Ni deposits, as

described previously. The S-solubility of a magmatic system is affected by

temperature, magma composition (aFeO, aSiO2, aNa2O), oxygen and sulfur

fugacity, pressure, and the presence of water, as summarized by Li and Ripley

(2005). Sulfur solubility in a magma decreases with decreasing temperature, aFeO,

sulfur-fugacity and pressure; whereas, increases in oxygen-fugacity, aSiO2 and

aNa2O will also cause decreases in S-solubility (Haughton et al. 1974; Shima and

Naldrett, 1975; Wallace and Carmicheal, 1992; Mavrogenes and O'Neil, 1999).

Although these factors contribute to S-solubility within a magma, most occurrences

of economic komatiite-hosted Ni result from excess S and over-saturation. Sulfur

over-saturation is attained through the assimilation of a local S-rich contaminant,

which is commonly but not exclusively exhalative in origin (Bekker et al., 2009).

b. Nickel sulfide distribution

Nickel sulfide distribution in extrusive orthomagmatic systems is divided into two

groups: primary mineralization, and secondary, or (re)mobilized mineralization.

Primary Ni sulfide mineralization is also subdivided based on sulfide location and

sulfide abundance (Lesher and Keays, 2002). Primary Ni sulfide mineralization

consists of basal contact, strata-bound, and stratiform mineralization types. Basal

contact mineralization is restricted to the footwall contact of the basal flow unit,

whereas hanging wall mineralization occurs at the base of a flow unit on a higher

stratigraphic level (Fig. 2.6: e.g. Lunnon, Hunt and McMahon ore shoots of the

Kambalda Dome: Gresham and Loftus-Hills, 1981). Both basal and strata-bound

mineralization occur in the form of disseminated (cloudy) sulfide (<20% sulfide),

matrix (net-textured) sulfide (20-60% sulfide), and massive sulfide (>60% sulfide).

Stratiform or reef mineralization is restricted within the disseminated and matrix

sulfides, and is typically contained within the differentiated cumulate units

(Fiorentini et al., 2007).

Secondary or mobilized mineralization consists of two sub-classes: metamorphic

and tectonic. Metamorphic mineralization is restricted to the mobilization of Ni from

massive ore into adjacent sulfidic metasediments, and the development of

40


mineralized veins in wall rock adjacent to primary mineralization with a quartz and

or carbonate vein association (Lesher and Keays, 2002). Metamorphic

mineralization within the Avebury Ni deposit of Tasmania, Australia, consists of

mineralization that is intimately associated with granite intrusion (Hoatson et al.,

2006; Keays and Jowitt, 2009). Tectonic mineralization is the mechanical

mobilization of primary mineralization into faults, shear zones, and fold hinges.

Tectonically mobilized deposits include those with complete detachment of the ore

from the primary magmatic host rock (e.g. Thompson Nickel Belt deposits: Layton-

Matthews et al., 2007), and partial detachment and mobilization of the

mineralization into adjacent lithologies (e.g. Kambalda Dome Ni deposits: Stone and

Archibald, 2004).

c. Metal tenor and distribution in sulfide ores

Nickel tenor is the Ni content of the sulfide assemblage, based on recalculation to

100% sulfide. Nickel tenor determines the relative abundance of pentlandite,

pyrrhotite and pyrite in the ore. Mineralization at Kambalda Dome exhibits a range

in Ni tenor between different ore shoots, and within different mineralization styles

within individual ore shoots (Gresham and Loftus-Hills, 1981). Nickel tenor within

the basal contact ore shoots from the Kambalda Dome are classified as either high

tenor (S:Ni > 2.5: Otter, Durkin, Gibb, Victor and Ken ore shoots) or low tenor

(S:Ni < 2.5: Long, Lunnon, Hunt and Gellatly ore shoots: Gresham and Loftus-Hills,

1981). Within the Fisher and Jaun ore shoots of the Kambalda Dome, nickel tenor is

consistent within an ore shoot, but adjacent ore shoots may have contrasting

compositions (Gresham and Loftus-Hills, 1981).

Variation in nickel tenor between adjacent ore shoots is attributed to differing R-

factors between the ore shoots. R-factor is a ratio of silicate to sulfide liquid, and is a

means to quantify mixing between silicate and sulfide liquids within the magma

(Campbell and Naldrett, 1979: see Section 2.4d on Chalcophile elements and

whole rock geochemistry). The occurrence of differing R-factors between ore

shoots is supported by positive correlations between S:Ni and the tenor of Pd, Pt,

and Ir (Ross and Keays, 1979; Cowden and Woolrich, 1987; Lesher and Campbell,

1983; Barnes, 2004b).

41


In addition to tenor differences, metal zonation within the mineralization is observed

on both the ore body and ore shoot scale. Matrix ores typically contain more Pd, Au,

Cu, and Ni, than the associated massive ores that are relatively enriched in Ir (Keays

et al., 1981). In detailed vertical ore profiles from the Kambalda Dome, a systematic

increase in Pd, Au, and Cu, and decrease in Ir is observed upward from the base of

the ore zone (Keays et al., 1981). A similar distribution is also observed in the Alexo

deposit within the Abitibi Greenstone belt (Barnes and Naldrett, 1985). A reversed

pattern with decrease in Pd, Au, and Cu, and increase in Ir in the Silver Swan ore

body, was attributed to an inverted cooling regime (Barnes, 2004).

Narrow sulfide stringers represent an extreme end member of metal zonation within

sulfide ores. Sulfide stringers commonly crosscut footwall lithologies and form a

minor ore component in many deposits (e.g. Kambalda Dome, Silver Swan, and

Widgiemooltha Dome of Western Australia). Sulfide stringers typically contain

elevated Cu, Au, and Pd concentrations relative to the adjacent massive ore, and are

interpreted as either the final stage in the generation of a zoned ore body, or a

product of hydrothermal alteration and remobilization of select ore elements.

The formation of metal zonation within ore bodies is attributed to differential

crystallization of a monosulfide solid solution (MSS: Distler et al., 1977; Naldrett et

al., 1994; Li et al., 1996; Barnes et al., 1997; Beswick, 2002). The cooling of a

sulfide liquid results in the generation of a Fe-rich MSS, where Os, Ir, Ru and Rh

represent compatible elements. Elements incompatible with a MSS form a fractioned

liquid (intermediate solid solution: ISS) that is enriched in Cu, Pt and Pd.

Fractionated ore bodies displaying metal zonation are well-documented at the

Voisey’s Bay Ovoid deposit of Canada (Huminicki and Sylvester, 2007), the

Noril’sk-Talnakh ores of Russia (Naldrett et al., 1994), and the offset ores of the

Sudbury Igneous Complex, Canada (Li et al., 1993; Beswick, 2002).

42


2.4. Mineralization Indicators

The identification of prospective magmatic systems to host mineralization is straight

forward when known deposits are delineated within a brownfield terrane (e.g.

komatiites in both Eastern Goldfields Terrane of Australia and Abitibi Greenstone

Belt of Canada). However, targeting potential mineralization hosted within a

prospective greenfields terrane is more difficult, due to the volume of associated

igneous lithologies. Consequently, a number of mineralization indicators and

mineralized systems characteristics have been identified to aid in the process of

targeting.

Mineralization indicators are divided into two groups: (1) magmatic process

indicators, and (2) mineralization process indicators (Barnes et al., 2004a).

Magmatic process indicators identify environments, settings and processes that are

conducive to the formation of orthomagmatic sulfide deposits. These consist of the

identification of channelized flow, continuous fluxing of primitive magma,

assimilation and contamination of the magma. Magmatic process indicators are

addressed through the use of (a) major and (b) trace element abundances.

Mineralization process indicators utilize the chalcophile nature of Ni, Cu, Co, Zn

and PGE to directly assess the S-saturation history of the melt through either (c)

mineralization or (d) chalcophile element depletion. Due to limited work on

chalcophile element mineralization indicators in Ni-mineralization systems, several

brief case studies are included to further discuss chalcophile element signatures of

ore formation (Persevearnce, Rocky’s Reward, and Mt. Keith of Western Australia,

and Noril’sk-Talnkh of Russia). Mineralization process indicators are also identified

in (e) minerals and mineral separates. The last section, (f) spatial distribution and

scale of mineralized systems, examines the known spatial relationships between ore

and mineralization indicators.

a. Major elements - whole rock geochemistry

Major element analyses of whole rock samples commonly include: Mg, Fe, Na, K,

Ca, Ti, Al, Mn, Si, P, Cr and Zn. As a result, there is a large amount of major

element geochemical data for komatiites in areas of significant mineralization (e.g.

Kambalda Dome and Mt. Keith of Western Australia), or in areas of readily

43


accessible outcrop (Abitibi Greenstone Belt of Canada). As Munro-type komatiites

host the majority of identified nickel mineralization most geochemical data is from

Munro-type komatiites. Munro-type komatiites are similar to Barberton- and

Karasjok-type in major and chalcophile elements and differ only in incompatible and

trace element abundances. Similarly, the ore forming processes and mineralization

setting are the same in the three types of komatiites. Consequently, the major

elements reflect the volcanological setting, regardless of geochemical classification

(Barberton-, Munro- or Karasjok-type).

Initial work by Lesher and Groves (1984) identified major element associations with

mineralization, that were later interpreted as the product of komatiite flow field

development (Hill et al., 1995; Hill, 2001). Mineralization is typically hosted

channels, representing areas of channelized and sustained magma transport. Olivine

cumulate rocks characterize channel environments and exhibit high MgO contents,

with limited trapped liquid and low abundances of the more incompatible elements

Ti and Al, (Hill and Gole, 1990; Lesher, 1989; Brand, 1999; Barnes and Brand,

1999; Hill, 2001; Lesher et al., 2001; Barnes et al., 2004a; Barnes et al., 2007). This

link between major element contents and volcanic facies is shown by a series of

density contoured bivariant major element plots (MgO-FeO, SiO2-MgO, Cr-Ni, Ni-

Ti) produced by Barnes et al. (2004a; 2007). These plots use a large

lithogeochemical data set which includes komatiites from the Eastern Goldfields

Superterrane of Western Australia and the Abitibi Greenstone Belt of Canada.

Volcanic facies probability fields are used to compare olivine abundance, olivine

fractionation, and trapped liquid, with volcanic facies, as reflected in the major

element abundances. Major elements (e.g. Cr-Ni) have been applied to map out

volcanic facies in both metamorphosed and weathered lithologies at the Kambalda

Dome (c.f. Fig. 5 Brand, 1999).

The link between Ni mineralization and major element abundance was also

examined by Barnes et al. (2007) on a geographically broader data set. Overall, the

conclusions of this study support the association of Ni mineralization with the

highest MgO content rocks (e.g. olivine cumulates). A difference in the initial liquid

MgO content (as determined from spinifex textured samples) was identified between

the nickel-rich Eastern Goldfields Superterrane and the relatively nickel-poor Abitibi

44


Greenstone Belt, with the latter exhibiting an overall lower MgO content during

eruption. The most significant difference between these two areas was the higher

proportion of olivine-rich cumulates in the Eastern Goldfields Superterrane the

result of larger magmatic systems.

b. Trace elements - whole rock geochemistry

Trace elements analyses of whole rock samples commonly include: large ion

lithophile elements (LILE: Cs, Rb, Ba), high field strength elements (HFSE: Th, Nb,

Hf, Zr), light rare-earth elements (LREE: La, Ce, Nd, Sr), middle rare-earth

elements (MREE: Eu, Gd, Tb, Dy, Y), heavy rare-earth elements (HREE: Ho, Er,

Tm, Yb, Lu), and transition metals (Sc, V). The highly incompatible lithophile

elements (Cs, U, Th, Nb, Ta and LREE), that are enriched in the crust relative to the

mantle are typically used to identify magmatic processes, including crustal

contamination (see review by Lesher et al., 2001). In general high-degree partial

melts from a depleted mantle source will be depleted in the highly incompatible

elements relative to the MREE and HREE. Deviations from this initial melt

composition (e.g. LREE enrichment, negative Nb and Ta anomalies) are attributed

to assimilation of crustal material (Perring et al., 1996; Lesher et al., 2001). On this

basis, a series of “assimilation-sensitive” ratios has been outlined by Lesher et al.

(2001) and Barnes et al. (2004b). These ratios include La/Smn, Th/Ybn and Zr/Tin

(where n denotes mantle normalized), and are used to assess the interaction between

komatiite magmas and crustal contaminants.

Crustal contamination can be an important process in the development

mineralization if the magmas become S-saturated during contamination (Barnes,

2006b; Arndt et al., 2008). However, there is no direct correlation between

contamination and mineralization. The Kambalda Dome system is characterized by

variable contamination signatures. The flanking environments show positive

contamination indicators (LREE enrichment), whereas the channel environment

does not show any evidence for contamination. This is though to be related to

system flushing and recharge within the channelized environments (Lesher and

Arndt, 1995).

45


c. Mineralization

Chalcophile element enrichment, as discussed previously, represents the best

indicator of Ni sulfide mineralization (Barnes et al., 2004). However, the application

of this indicator is not simple or straight forward. For example, the significance of

disseminated mineralization or a narrow zone of mobilized Ni sulfide, in term of the

larger system scale is unknown. Likewise, how doe you get from the tail, trunk or

ear of the elephant to the main body (Fig. 2.7)?

Figure 2.7. Blind persons and the elephant. Cartoon based on poem by John Godfrey Saxe (1816-1887). Modified from Yeh and Rousseau (2000).

Effective targeting of mineralization is based on the identification of geochemical

haloes (both positive and negative anomalies: Goldberg et al., 2003). However, the

current belief in Ni targeting is that mineralization haloes are of limited spatial

extent or are non-existent. This assumption is based on the relative timing between

mineralization and the adjacent host-rock, where crystallization of the host rock

often postdates mineralization (Lesher and Campbell, 1993; Lesher and Arndt,

1995). Consequently, geochemical haloes associated with massive Ni sulfides

typically comprise only of the adjacent disseminated sulfide zone. Recent research

on mineralization haloes has focused on the identification of chalcophile element

(Ni, Cu, PGE) enrichment beyond the zone of visually identifiable mineralization

(Fiorentini et al., 2010). Chalcophile element enrichment is observed; however it is

unknown if this enrichment is related to primary or secondary processes. Examples

of questionable enrichment include “unsupported PGE enrichment,” as described by

46


Fiorentini et al. (2010) and metal enriched hydrous alteration minerals described by

Hanley and Bray (2009). Metal enriched hydrous alteration minerals are discussed

further in Section 2.5E.

Unsupported PGE enrichment occurs when whole rock samples contain PGE

abundances above the normal background level, without elevated sulfide values.

Fiorentini et al. (2007) identified unsupported PGE enrichment in sulfide-poor

whole rock samples within the Albion Downs mineralized komatiite system of

Western Australia (Jericho and Jordan deposits). Mineralization within the belt is

both massive (type-1: Jordan deposit) and disseminated (type-2: Jericho deposit).

Unsupported enrichment signatures have been identified within 10m of the

disseminated type-1 mineralization and within 100m of the massive type-2

mineralization. Samples taken at a greater distance from known mineralization (~2

km) did not show the same enrichment pattern. Unsupported enrichment was also

documented in the Black Swan deposit of Western Australia (Barnes, 2004b; Barnes

et al., 2004b; Hill et al., 2004). Barnes et al. (2004b) identified two samples from

within the channel facies (unknown distance to mineralization) that exhibited

elevated PGE abundances. Flanking units to the mineralized channel exhibited

normal background values. These anomalous PGE values within the channel were

attributed to sulfur loss from pre-existing disseminated mineralization.

d. Chalcophile elements - whole rock geochemistry

With current analytical techniques, chalcophile elements comprising Ni, Cu, Co, and

the platinum group elements (PGE: Pt, Pd, Rh, Ru, Ir, Os) are detectable at normal

background concentrations in mafic and ultramafic whole-rock samples (Barnes and

Fiorentini, 2008b). The use of chalcophile elements as mineralization indicators

differs from the previous two mineralization indicators (major and trace elements).

During the ore forming process, chalcophile elements are extracted from the silicate

melt and concentrated in the sulfide melt (mineralization). Consequently,

chalcophile element abundance is a direct indicator of Ni mineralization, and is

divided into two end-members: (1) enrichment, and (2) depletion. Enrichment

(mineralization) as discussed in Section 2.5C, is the most effective method to

identify mineralized systems. In contrast, the use of chalcophile element depletion is

limited, as its practical application is tied closely to advances in analytical

47


techniques, precision, and the understanding of orthomagmatic mineralization

systems. Initial research and numerical models working with Ni, Co, and Zn

predicted the existence of silicate melts in ore forming systems that are chalcophile

element depleted due to the high sulfide partition coefficients of the chalcophile

elements (MacLean and Shimazaki, 1976; Duke and Naldrett, 1978; Rajamani and

Naldrett, 1978; Duke, 1979; Campbell and Naldrett, 1979).

i. Chalcophile element partitioning

Chalcophile element depletion is the result of the chalcophile elements having

partition coefficients > 1 for the sulfide liquid phase. The chalcophile elements as a

group have a range of partition coefficients where Co is the lowest at ~30, followed

by Ni at 100-200, Cu: 300-1000, and PGE with the highest range of 10000 to

>100000 (see reviews of Barnes and Maier, 1999; Mathez, 1999). Absolute partition

coefficient values are not yet determined and substantial single element variation is

well documented in the literature. This variation is inferred to be dependent upon

melt composition, oxygen fugacity, sulfur fugacity, and temperature.

The initial application of chalcophile element depletion as a mineralization

indicators was largely restricted to the use of Ni, Cu, Co and Zn as these elements

were routinely analyzed with good precision. However, the low partition coefficients

of Ni, Cu, Co, and Zn resulted in ambiguous depletion signatures (Lesher and

Groves, 1984; Lesher 1989, 1993; Lesher and Arndt, 1995). Platinum group

elements have partition coefficients two orders of magnitude higher than Ni, Cu, Co,

and Zn and consequently are more sensitive indicators of sulfur saturation. However,

the lack of analyses precise at the ppb level limited the initial application of the

PGEs. Commercially available PGE analyses at applicable levels (<ppb) have only

been available since the 1990s, thus providing a new opportunity to develop a more

accurate assessment of chalcophile element depletion in mineralized komatiites.

ii. R-factor

One critical factor that controls the magnitude of chalcophile element depletion

signatures within the silicate magma is the effective mixing between the silicate and

sulfide phase. The enrichment of chalcophile elements within the sulfide phase is

48


defined as the R-factor, a ratio of silicate to sulfide liquid (Campbell and Naldrett,

1979):

CS = CLD(R+1)/(R+D)

Rearranged to

R=CSD/(CLD-CL)

Where CS is the concentration of metal in sulfide, CL is the concentration of metal in the initial silicate liquid, D is the partition coefficient D=Dsul/sil , and R is the ratio of silicate to sulfide liquid (Campbell and Naldrett, 1979; Campbell and Barnes, 1984; Lesher and Stone, 1996).

A high R-factor implies a low sulfide abundance interacting with a large volume of

silicate magma, causing effective removal of the high partition coefficient

chalcophile elements (PGE) from the silicate magma, relative to the moderately

chalcophile elements Cu and Ni. A low R-factor indicates a greater abundance of

sulfide interacting with a volume of magma, and the effective removal of all

chalcophile elements from the silicate magma, resulting in lower grade sulfides.

If R is greater than 10 times D, the enrichment factor (CS/CL) in the sulfide liquid

approaches D; whereas, if R is less than 10 times D, the enrichment factor is

approximately equal to R. Consequently, high PGE-enrichment requires R values of

greater than 10000. The effects of varying R-factors on mineralization systems (both

enrichment and depletion) are summarized by Lesher and Campbell (1993) and

Lesher et al. (1999; 2001), and listed below as high, moderate, and low R-factors:

1. Magmas that equilibrate at high R-factor values may not exhibit

significant depletion in any of the chalcophile elements, and the sulfides

generated may be enriched in all chalcophile elements.

2. Magmas that equilibrate at moderate R-factor values may exhibit

chalcophile element depletion for elements where R <10*D, and the sulfides

may be relatively enriched in Co>Ni>Cu>PGE.

3. Magmas that equilibrate at low R-factor values may exhibit significant

depletion in all the chalcophile elements, and the generated sulfides may

only be slightly enriched in the chalcophile elements.

49


Calculated R-factor values for komatiite Ni deposits are variable, and values

obtained from the Kambalda Dome range from 100-500 (Campbell and Barnes,

1984; Lesher and Campbell, 1993).

iii. Chalcophile element mineralization signatures

The identification of chalcophile element mineralization signatures is typically by

plotting the chalcophile elements as either a function of MgO (Duke and Naldrett,

1979), or as a chalcophile element ratio (e.g. Cu/Pd, Ir/Pd: Campbell and Barnes,

1984; Barnes et al., 1985; Barnes et al., 1988; Barnes, 1990; Li et al., 2001). MgO is

useful for most ultramafic systems, as the majority of systems exhibit strong olivine

control and MgO functions as an olivine accumulation index. Nickel is compatible

in olivine and as a result strong positive correlations are observed between MgO and

Ni. Most of the other chalcophile elements (Pt, Pd, Rh, and Cu) are incompatible in

olivine, and exhibit strong negative correlations with MgO. Initial chalcophile

element depletion signatures were originally based on this relationship and

deviations from theoretical olivine control lines (Duke and Naldrett, 1978; Duke,

1979). Chalcophile element depletion was also identified through comparative

geochemical analysis of barren and mineralized komatiites. The mineralized

komatiites exhibited lower chalcophile element abundance than the barren

komatiites. This was interpreted as a result of S-saturation and chalcophile element

depletion (Lesher and Groves, 1984; Barnes et al., 2004a).

The same methodology utilizing MgO as a fractionation index was also applied to

the PGE elements (Barnes et al., 1985; Lesher et al., 2001; Barnes et al., 2004a) with

significant scatter in PGE concentrations observed in the binary plots. Normalization

of the PGE abundance to Ti was able to remove the effects of olivine accumulation

and fractionation within the system, and reduced the observed data scatter (Barnes et

al., 2007). Utilizing this methodology komatiites in a terrane-scale comparisons

between the Eastern Goldfields Superterrane, Western Australia and the Abitibi

Greenstone Belt, Canada clearly displayed a strong depletion signature in a small

fraction of the dataset from Kambalda.

Chalcophile element ore forming signatures have traditionally been used for the

understanding and exploration of PGE-dominated mineralization (reef-type

50


mineralization hosted within large layered intrusions), as well as conduit type

systems (komatiites, flood basalts). The application of chalcophile element

signatures in large intrusions is commonly based on the stratigraphic examination of

a chalcophile element ratio. Ideally, the chalcophile element ratio indicates a

stratigraphic interval that is depleted in the highly chalcophile elements (Pt, Pd)

relative to a less chalcophile element (Cu). The simplest interpretation of this

setting, is that the mineralization event occurred stratigraphically below a change in

the chalcophile element ratio, and is preserved in the igneous stratigraphy (Maier et

al., 1998; Maier and Barnes, 2005). There are numerous applications for this

stratigraphic technique, with examples summarized in Table 2.2.

Table 2.2. Case studies that have utilized chalcophile element ratios to identify orthomagmatic mineralization in intrusions.

Intrusion Reference

Bushveld Complex Maier and Barnes, 1999; Barnes et al., 2004

Tete Complex Maier et al., 2001

Uitkomst Complex Maier et al., 2004

Stella Intrusion Maier et al., 2003

Munni Munni Hoatson and Keays, 1989; Barnes et al., 1993

Panton Hoatson and Blake, 2000

Duluth Complex Theriault et al., 2001; Miller et al. 2002

Mordor Complex Barnes et al. 2008

Stillwater Complex Godel and Barnes, 2008

Honngge Intrusion Zhoug et al., 2002

Baimazhai Intrusion Wang et al., 2006

Open systems with extensive magma flow-through that host mineralization are the

product of metal extraction from a silicate melt. However, chalcophile element

depletion signatures within these systems (e.g. komatiite) are scarce and poorly

constrained, due to the continuous recharging of the system and the extensive

development of flow fields (Lesher and Arndt, 1995). Whole-rock chalcophile

element depletion signatures, attributed to S-saturation in mineralized komatiite

systems (PGE-based), are limited to the Kambalda Dome (Lesher et al., 2001),

Perseverance (Barnes et al., 1995; 2004) and Mt. Keith deposits (Barnes et al.,

2004). Depletion signatures are also documented within the Noril’sk deposit of

51


Russia; however, this has been highly debated (Naldrett, 1992; Naldrett et al., 1992;

Brügmann et al., 1993; Latypov, 2002; 2007). Chalcophile element depletion

signatures are observed within the West Greenland flood basalts and the Deccan

Traps; however, the depletion is not related to known mineralization (Momme et al.,

2002; Keays and Lighfoot, 2007; 2010).

iv. Examples of chalcophile element signatures

The Kambalda Dome Ni komatiites of Western Australia are characterized by a

number of ore shoots (as described in Chapter 3) hosted within the Silver Lake

Member. The most studied ore shoot is the Long-Victor shoot, located on the eastern

flank of the dome (Fig. 3.4). Timing of mineralization is poorly constrained within

the system, but is assumed to be quite early in the flow field development. The

greatest abundance of depleted samples occurs in the non-cumulate (spinifex

textured) lithologies of the flanking sheet flow facies within the Silver Lake

Peridotite Member (Lesher et al., 2001; Hill, 2001). Channelized environments that

host the majority of nickel mineralization at Kambalda Dome do not display a

depletion signature. This was addressed by Lesher and Campbell (1993) and Lesher

and Arndt (1995), using trace elements to examine the spatial distribution of

contamination signatures. The research indicated that channel facies were refreshed

with new magma (lacking a contamination signature), and the expected

mineralization signature was subsequently flushed from the channel environment.

The Tripod Hill Member of the Kambalda Dome (continuous with, and overlying

the Silver Lake Member) hosts no known Ni mineralization. Whole-rock

geochemistry indicates normal undepleted chalcophile element abundances for the

Tripod Hill member (Keays et al., 1981; Keays, 1982; Lesher et al., 2001),

suggesting that the magmas were not sulfur saturated at any time during

emplacement.

Bavinton and Keays (1978) examined the possibility of using precious metal (Au,

Pd, Ir, Ag) abundances in metasedimentary rocks from the Silver Lake Member to

identify proximity to nickel ores. Despite having a laterally extensive sample set (44

samples over 35-40 km2), the authors did not identify any systematic indicator of

proximity to nickel mineralization.

52


The Perseverance and Rocky’s Reward Ni system of Western Australia is hosted

in extrusive komatiites within a felsic volcanic sequence (Barnes et al., 1995). These

deposits formed during the initial outpouring of komatiitic lava, with sulfur

saturation induced by the assimilation of local weakly-sulfidic felsic volcanics. On

the deposit scale, a persistent Ni depletion in the dataset was identified by Barnes et

al. (1988; 2004a). However, whole-rock analysis of sulfide-poor samples (S<0.2%)

did not display a similar depletion in PGE. The PGE were actually enriched,

resulting in ambiguous signatures (Barnes et al., 2004a).

The Mt. Keith Ni-system of Western Australia is characterized as a large intrusion

dominated by olivine cumulate with disseminated sulfide (Grguric et al., 2006;

Rosengren et al., 2007). This intrusion is hosted within felsic volcanic rocks

(Rosengren et al., 2005; 2008). Mineralization within the intrusion is interpreted to

have formed upstream of the current location and was subsequently transported as

sulfide droplets to its current location (Barnes, 2007). Deposit scale analyses

identified a persistent Ni depletion in samples from the Eastern and Central units

that contain the Mt. Keith and Cliffs Ni deposits, respectively (Barnes et al., 2004a).

PGE analyses from the Eastern and Central and unmineralized Western ultramafic

unit, exhibited depletion signatures at a deposit scale resolution (Barnes et al.

2004a).

The Noril’sk-Talnakh Ni-system of Russia is associated with the Permo-Triassic

(248-250 Ma: Renne and Basu, 1991; Campbell et al., 1992) Siberian Trap flood

basalts and represents one of the largest Ni-Cu-PGE deposits in the world. The

Noril’sk-Talnakh system is perhaps both the best and worst example of the use of

chalcophile elements as vectors. Controversy still exists around the relationships

between mineralization, mineralized host rocks, and volcanic stratigraphy within the

systsem.

Nickel mineralization within the Noril’sk-Talnakh system is hosted by small,

differentiated mafic-ultramafic intrusions emplaced within Permian sedimentary

rocks which are overlain by flood basalts (Naldrett et al., 1992). Initial research

examining the stratigraphic sequence of eruptive flood basalts identified strong

chalcophile element depletion and LREE enrichment within the Nadezhdinsky

Formation. Other basalt formations stratigraphically below and above this formation

53


did not contain this signature (Naldrett, 1992; Naldrett et al., 1992; Brügmann et al.,

1993). It was interpreted that the heavily mineralized intrusions within the Noril’sk-

Talnakh deposits were co-magmatic with a portion of the extrusive stratigraphy

(Nadezhdinsky Formation). These intrusions may have acted as S-saturated

subvolcanic feeders, that concentrated the chalcophile elements from a large volume

of magma that had flowed through the system (Naldrett et al., 1992; Brugmann et

al., 1993; Naldrett, 1997).

Open dynamic magma conduits are one of the most effective settings for extraction

of chalcophile element from a silicate magma under S-saturated conditions, and

provided a plausible mineralization model for the Noril’sk-Talnakh system.

However, contention with this original Noril’sk model is in the genetic link between

intrusions hosting mineralization and the flood basalts. Disparity between the two is

identified in the geochemistry (isotopes, and major and trace elements), and

mineralogy of the intrusions relative to the extrusive basalts (Latypov, 2002; 2007).

Further refinements to the mineralization model of the Noril’sk-Talnakh system

were proposed by Lightfoot and Keays (2005). It was suggested that magma mixing

and S-saturation occurred within a deep-seated staging chamber. This was followed

by the dissolution of early formed sulfides and re-precipitation of an immiscible

sulfide liquid within the intrusions following emplacement. Re-precipitation of an

immiscible liquid was attributed to the assimilation of local sulfate-rich

metasedimentary rocks (Li et al., 2009).

e. Minerals and mineral separates

The use of individual minerals and mineral separates in the identification of

orthomagmatic mineralization signatures is limited. Traditionally unaltered olivine,

a common mineral phase in komatiite systems, has been used as Ni readily partitions

into the olivine crystal structure. Theoretical models indicate that Ni depletion

occurs when olivine crystallizes from a sulfide saturated magma (Naldrett and Duke,

1978; Duke, 1979). These theoretical models have been tested with varying success

in mineralized and unmineralized settings. Relict olivine crystals within the

mineralized channel facies at Kambalda are not depleted in Ni, but occur proximal

to mineralization (Lesher, 1989). The discrepancy between the normal Ni abundance

54


in the olivine and the presence of massive Ni sulfide within the channel was

attributed to recharge occurring within the channel. Olivine temporally associated

with the generation of mineralization is removed from within the channel by

recharge (Lesher and Campbell, 1993; Lesher and Arndt, 1995). The Perseverance

system (Agnew deposit) exhibits a range of olivine nickel abundances, which are

attributed to be a result of variable mixing between an immiscible sulfide and

silicate liquid (Barnes et al., 1988). Olivine from the Scotia deposit also exhibited

mixed patterns; however, Ni depletion in some olivine was attributed to alteration

rather than sulfide extraction (Stolz and Nesbitt, 1981).

Recent research has focused on chromite, and the relative abundance of Ru in

chromite that crystallizes within a mineralized system, versus a barren komatiite

system (Fiorentini et al., 2008b). Chromite from komatiite systems with Ni

mineralization exhibit a lower Ru concentration than unmineralized systems.

Continuing work on the application of chromite and Ru as mineralization indicators,

is being undertaken by Locmelis et al. (2009).

Research examining alteration veins associated with mineralization in the Sudbury

Igneous Complex of Canada (Ames and Farrow, 2007), has identified a spatial

correlation between the metal content of amphiboles and proximity to mineralization

(Hanley and Bray, 2009). Amphibole-bearing (actinolite to actinolitic-hornblende)

alteration veins that postdate the bolide impact, but predate the emplacement of

mineralization, were analyzed for chalcophile element abundances. Metal content

within the amphiboles varies from <100 ppm Ni at ~3 km from known

mineralization, to >1% Ni within 1 m of mineralization (Hanley and Bray, 2009). A

distance of 700m was identified as a probable range for a proximity indicator. A

similar spatial relationship was observed in Cu and Sn, but on a much smaller scale

(<5m). Several chalcophile elements (Pb, Co, and Zn) did not exhibit a spatial

correlation with mineralization. Nickel enrichment in the amphiboles was ultimately

attributed to hydrothermal metasomatism by a saline fluid enriched in Ni via

leaching of contact style ores (Hanley and Bray, 2009).

55


f. Spatial distribution and size of mineralized systems

The spatial distribution of ore-forming signatures (chalcophile elements) has been

discussed briefly in the previous sections. Overall there has been limited research on

the spatial distribution of ore forming signatures. The Kambalda Dome area contains

the only empirical evidence of a depletion signature within spinifex from the

flanking environments (Lesher et al., 2001). In general, the size of Ni ore forming

systems is poorly constrained, with vague descriptors such as “distal” usually

applied to this aspect of the mineralized systems. The distribution of chalcophile

element enrichment proximal to known mineralization was approached by Fiorentini

et al. (2007), at the Jericho deposit in Western Australia. Elevated PGE abundances

were identified in low-sulfide samples within 10 to 100m of known mineralization.

However, the systematic spatial work examining ore forming signatures has not

been done within the Kambalda system.

Preservation of chalcophile element depletion within a mineralized system is

variable due to the dynamic and open nature of komatiite systems. Ore forming

signatures are moderately preserved in orthomagmatic systems, characterized by

large layered intrusions and PGE deposits (e.g. Bushveld Complex of South Africa,

and Stillwater Complex of the United States of America). This moderate

preservation of ore forming signatures is due to the “closed” nature of these systems

(Li et al., 2001). In orthomagmatic systems that host massive Ni-sulfide, the

evidence for chalcophile element depletion is scarce (Lesher and Campbell, 1993;

Lesher and Arndt, 1995; Lesher et al., 2001). This lack of preservation of

chalcophile element depletion signatures is interpreted as a result of sustained

recharge. However, the timing, duration, and localization of sulfur saturation are

also critical factors.

2.5. Conclusion, Implications and the Way Forward

Komatiites provide a unique window for the examination of mantle and crustal

evolution prior to modern plate tectonics. World peak komatiite production occurred

between 2.7 and 2.9 Ga, with the majority of komatiite-hosted Ni deposits restricted

to lithological units of these ages. Consequently, there are no modern analogs.

56


Archean komatiite systems underwent multiple episodes of deformation, alteration,

and metamorphism, and are currently exposed as discontinuous sequences. The

available material from Archean komatiite sequences does not provide a complete

picture in terms of komatiite generation, tectonic setting, or mineralization

processes; all of which remain unresolved or controversial.

a. Komatiite generation

Komatiite melt generation is argued result from anhydrous melting of a mantle

plume (Arndt et al., 2008). However, recent research has identified increasing

proportions of water within plume-related mafic and ultramafic systems (Stone et

al., 1997; Beresford et al., 2000; Dann, 2001; Wilson et al., 2003; Fiorentini et al.,

2006; 2008b; Barr et al., 2009). The influence of hydrous phases in the generation of

komatiite melts is unknown.

b. Tectonic setting

Ultramafic magmatism is believed to be the product of mantle plume melting, and is

supported by both the high-degree partial melting required to generate high-MgO

lavas, and the observed lithostratigraphic sequences (e.g. Eastern Goldfields

Terrane: Campbell et al., 1989). Although mantle plumes can occur in any tectonic

setting (oceanic, rift, continental, etc.), the reoccurring association between felsic

volcanic rocks (volcanic arc-type) and komatiites warrants further research. Is it

possible to generate komatiite and komatiitic magmas in other tectonic settings (e.g.

subduction zones), as proposed by Parman et al. (2001; 2004), Parman and Grove

(2004), Smithies et al. (2004).

Komatiites form extensive flow fields comprising numerous facies (Hill et al., 1995;

Hill, 2001; Barnes, 2006) and extend for 100s of kilometres. Yet, komatiite vent

sites are not conclusively identified in the rock record. It is unknown if komatiite

systems develop from a point source, multiple point sources, linear fractures, distal

source, or a proximal source. All of these locations have been proposed throughout

the literature, but definitive supporting evidence is required (Lesher et al., 1984; Hill

et al., 1995; Hill, 2001; Prendergast, 2001; 2003).

57


c. Mineralization processes

External sulfur is identified as a critical component in the development of komatiite

hosted Ni deposits. However, the process of incorporating sulfur into the magmas

remains problematic, as outlined by Naldrett (2005). It is unknown if sulfur is

incorporated into the system through complete assimilation of a sulfidic

contaminant, or if it is possible for S-diffusion to play a role in S-saturation.

Mineralization hosted within komatiite systems is usually well-constrained. Mass

balance calculations between the mineralization and the metal abundance in the

initial magma have indicated that a large volume of magma has interacted with the

immiscible sulfide liquid. Although a large volume of magma is involved with the

ore-forming process, the spatial extent or scale of most ore-forming systems is

unknown.

d. Considerations

Knowledge gaps are present in the understanding of komatiite systems and this

summary is by no means complete. Continued research will resolve some of these

questions, and generate new questions during the process. Some knowledge gaps

may never be answered due to a lack of suitable material or exposure.

Two knowledge gaps are addressed in this thesis. The first is the scale of ore

forming systems. The scale of mineralized komatiite systems is intrinsic to the

targeting of Ni sulfide mineralization, where knowledge of the system volume can

indicate that mineralization is present, without actually identifying the

mineralization. The scale of mineralized systems is the main focus, as presented in

Chapters 4 and 6. These chapters constrain the size of two komatiite systems

(Long-Victor and Maggie Hays of Western Australia) that interacted directly with a

sulfide liquid during the mineralization process. Chapter 7 applies the concepts

developed in Chapters 4 and 6 to a greenfields exploration scenario for komatiite

hosted nickel sulfide mineralization in northern Finland and Norway.

The second knowledge gap addressed in this thesis is the morphology of a komatiite

complex. Chapter 5 presents the stratigraphy and stratigraphic control on the

58


development of an ultramafic complex, containing both intrusive subvolcanic feeder

and overlying extrusive flows (Maggie Hays deposit of Western Australia).

The next chapter in this thesis, Chapter 3, provides an introduction into the geology

and mineralization within the Kambalda Dome area of Western Australia; which

includes the Long-Victor deposit.

59


2.6. References

Aitken, B. C., Echeverria, L. M. 1984. Petrology and geochemistry of komatiites and tholeiites from Gorgona Island, Colombia: Contributions to Mineralogy and Petrology v. 86, p. 94-105.

Ames, D.E., Farrow, C.E.G., 2007. Metallogeny of the Sudbury mining camp, Ontario: In: Goodfellow, W.D., (ed.), Mineral Deposits of Canada: A Synthesis of Major Deposit-types, District Metallogeny, the evolution of Geological Provinces, and Exploration Methods: Geological Association of Canada, Mineral Deposit Division, Special Publication No. 5, p. 329-350.

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

Arndt, N., 1994. Archean komatiites: In: Condie, K. C. (ed.), Archean Crustal Evolution. Amsterdam: Elsevier, p. 11-44.

Arndt, N., Brooks, C., 1980. Komatiites: Penrose conference report: Geology, v. 8, p. 155-156.

Arndt, N.T., Jenner, G.A. 1986. Crustally contaminated komatiites and basalts from Kambalda, Western Australia: Chemical Geology, v. 56, p. 229-255

Arndt, N.T., Lesher, C.M., 1992. Fractionation of REE's by olivine and the origin of Kambalda komatiites, Western Australia: Geochimica et Cosmochimica Acta, v. 56, p. 4191-4204.

Arndt, N.T., Nesbitt, R.W., 1982. Geochemistry of Munro Township basalts. In: Arndt, N.T., Nisbet, E.G. (eds.), Komatiites: Allen and Unwin, London, p. 309-329.

Arndt, N.T., Francis, D., Hynes, A.J., 1979. The field characteristics and petrology of Archean and Proterozoic komatiites: Canadian Mineralogist, v. 17, p. 147-163.

Arndt, N., Ginibre, C., Chauvel, C., Albarede, F., Cheadle, M., Herzberg, C., Jener, G., Lahaye, Y., 1998. Were komatiites Wet?: Geology, v. 26, p. 739-742.

Arndt, N.T., Lesher, C.M., Barnes, S.J., 2008. Komatiite: Cambridge University Press, 467p.

Arndt, N.T., Lesher, C.M., Houle, M.G., Lewin, E., Lacaze, Y., 2004. Intrusion and crystallization of a spinifex-textured komatiite sill in Dundonald Township, Ontario: Journal of Petrology, v. 45, p. 2555-2571.

Arndt, N.T., Naldrett, A.J., Pyke, D.R., 1977. Komatiitic and Iron-rich Tholeiitic Lavas of Munro Township, Northeast Ontario: Journal of Petrology, v. 18, p. 319-369.


Barnes, S-J., Boyd, R., Korneliussen, A., Nilsson, L.P., Pedersen, R.B. Robins, B., 1988. The use of mantle normalization and metal ratios in discriminating between the effects of partial melting, crystal fractionation and sulfide segregation on platinum-group elements, gold, nickel and copper: Examples from Norway. In: Geo-Platinum 87, p. 113-143.

Barnes, S-J., Couture, J.-F., Sawyer, E.W., Bouchaib, C., 1993. Nickel-copper occurrences in the Belleterre-Angliers belt of the Pontiac Subprovince and the use of Cu-Pd ratios in interpreting platinum-group element distribution: Economic Geology, v. 88, p. 1402-1419.

Barnes, S-J., Gorton, M.P., Naldrett, A.J., 1983. A comparative study of olivine and clinopyroxene spinifex flows from Alexo, Abitibi Greenstone Belt, Ontario, Canada: Contributions to Mineralogy and Petrology, v. 83, p. 293-308.

Barnes, S-J., Makovicky, E., Makovicky, M., Rose-Hansen, J., Karup-Moller, S., 1997. Partition coefficients for Ni, Cu, Pd, Pt, Rh and Ir between monosulfide solid solution and sulfide liquid and the formation of compositionally zoned Ni-Cu sulfide bodies by fractional crystallization of sulfide liquid: Canadian Journal of Earth Sciences, v. 34, p. 366-374.

Barnes, S-J., Maier, W.D., 1999. The fractionation of Ni, Cu and the noble metals in silicate and sulfide liquids: In: Keays, R.R., Lesher, C.M., Lightfoot, P.C., Farrow, C.E.G., (eds.), Dynamic processes in magmatic ore deposits and their application in mineral exploration, Geological Association of Canada, Short Course, v. 13, p. 69-106.

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Barnes, S-J., Maier, W.D., Ashwal, L.D., 2004. Platinum-group element distribution in the Main Zone and Upper Zone of the Bushveld Complex, South Africa: Chemical Geology, v. 208, p. 293-317.

Barnes, S-J., Naldrett, A.J., 1986. Variations in platinum group element concentrations in the Alexo mine komatiite, Abitibi greenstone belt, northern Ontario: Geological Magazine, v. 123, p. 515-524.


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Index

2.1. Introduction ..................................................................................................... 11 2.2. Komatiite Geochemistry and Volcanic Processes ........................................... 11

a. Classification ............................................................................................. 12 b. Geochemistry ............................................................................................. 12

i. Melt generation ................................................................................... 13 ii. Chalcophile elements ........................................................................... 14 iii. Crystallization ..................................................................................... 16 iv. Contamination ..................................................................................... 17

c. Transport and eruption .............................................................................. 18 d. Volcanic textures ....................................................................................... 19

i. Spinifex ................................................................................................ 19 ii. Cumulates ............................................................................................ 22 iii. Harrisite .............................................................................................. 26 iv. Breccia-volcaniclastic ......................................................................... 26 v. Vesicles ................................................................................................ 28

e. Volcanic flow field .................................................................................... 28 i. Propagation and field development ..................................................... 29 ii. Flow thickness ..................................................................................... 32 iii. Channel and Trough ............................................................................ 33 iv. Flank .................................................................................................... 35 v. Scale .................................................................................................... 36

2.3. Orthomagmatic Mineralization Model ............................................................ 37 a. Sulfur in orthomagmatic nickel systems ................................................... 40 b. Nickel sulfide distribution ......................................................................... 40 c. Metal tenor and distribution in sulfide ores ............................................... 41

2.4. Mineralization Indicators ................................................................................. 43 a. Major elements - whole rock geochemistry ............................................. 43 b. Trace elements - whole rock geochemistry ............................................... 45 c. Mineralization ............................................................................................ 46 d. Chalcophile elements - whole rock geochemistry ..................................... 47

i. Chalcophile element partitioning ........................................................ 48 ii. R-factor ................................................................................................ 48 iii. Chalcophile element mineralization signatures .................................. 50 iv. Examples of chalcophile element signatures ....................................... 52

e. Minerals and mineral separates ................................................................. 54 f. Spatial distribution and size of mineralized systems ................................. 56

2.5. Conclusion, Implications and the Way Forward ............................................. 56 a. Komatiite generation ................................................................................. 57 b. Tectonic setting ......................................................................................... 57 c. Mineralization processes ........................................................................... 58 d. Considerations ........................................................................................... 58

2.6. References ....................................................................................................... 60

List of Figures

Figure 2.1. World map showing distribution of major orthomagmatic deposits, Ni mineralization districts and geographical locations referenced in this thesis. Komatiite-hosted deposits comprise: Mt. Keith, Perseverance, Black Swan, and Kambalda deposits of Western Australia; Reliance deposit of Africa, and

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Abitibi Greenstone Belt of Canada. Komatiitic basalt-hosted deposits comprise the Thompson Ni-belt and Raglan Ni-belt of Canada. High MgO basalt deposits are characterized by Noril’sk-Talnakh of Russia, Jinchuan deposit of China, and Kabanga deposit of Tanzania. Ferro-picrite is associated with the Pechenga deposit of Russia. Troctolite is associated with the Voisey’s Bay deposit of Canada. Meteorite impact related deposits are characterized with the Sudbury region of Canada. Large layered intrusions, hosting reef-type platinum group element mineralization, are characterized by the Stillwater Complex of the United States of America, and Bushveld Complex of South Africa. The Karelian Craton of Finland and Norway is included for reference to Karasjok-type komatiites. .................................................................................................... 15

Figure 2.2. Diagram illustrating fully differentiated komatiite flow with upper A-zone spinifex and lower B-zone olivine cumulates. Modified from Pyke et al. (1973) and Arndt et al. (1977). ............................................................................ 21

Figure 2.3. Komatiite cooling units matrix with increasing olivine accumulation on left and increasing differentiation along the bottom axis. UN = undifferentiated non-cumulate (massive, pillowed or volcaniclastic), DN = differentiated non-cumulate, UC = undifferentiated cumulate, DC = differentiated cumulate. Modified from Lesher and Keays (2002)............................................................. 25

Figure 2.4. Komatiite flow field model as proposed by Hill (2001) showing the transition from massive sheet flow to channelized flow. Modified from Arndt et al. (2008). ............................................................................................................. 30

Figure 2.5. Komatiite flow field model as proposed by Hill (2001) showing lobe development at the advancing front and lateral development. Modified from Arndt et al. (2008). ............................................................................................... 31

Figure 2.6. Idealized schematic cross-section showing both channel and flank facies with associated sediments and Ni-sulfide mineralization as observed at the Kambalda Dome. Modified from Cowden and Roberts (1990). ......................... 34

Figure 2.7. Blind persons and the elephant. Cartoon based on poem by John Godfrey Saxe (1816-1887). Modified from Yeh and Rousseau (2000). ............. 46

List of Tables

Table 2.1. Greenstone belts containing volcaniclastic textured ultramafic lithologies. Barberton-type komatiite (B-type), Munro-type komatiite (M-type), Karasjok-type komatiite (K-type). ....................................................................................... 27

Table 2.2. Case study intrusions that have chalcophile element ratios utilized to identify orthomagmatic mineralization. ............................................................... 51


Chapter 3. The Kambalda Dome.

3.1. Introduction

The Yilgarn Craton of Western Australia (Fig. 3.1) is well known for its wealth of

mineral resources. Gold was first targeted by early prospectors and miners in the

region, with many of the original 1890s discoveries still in operation within the

Eastern Goldfields Superterrane (e.g. Kalgoorlie: Golden Mile-Super Pit). The

Kalgoorlie Terrane contains prolific gold deposits, but also hosts substantial nickel

(Ni) sulfide deposits. Komatiite-hosted Ni sulfide was discovered in 1966,

approximately 55 km south of the town Kalgoorlie, at Kambalda, Western Australia.

Subsequent Ni discoveries are focused along the 550 km extent of the Kalgoorlie

Terrane (Fig 3.1), as summarized by Barnes (2006).

Komatiite-hosted Ni mineralization at Kambalda occurs around a doubly plunging

anticline cored by granitic intrusions (Kambalda Dome) that post-date ore formation.

Nickel mineralization is identified within the volcanic host rock stratigraphy that is

exposed along the flanks of the dome. The Ni mineralization does not occur as a

single body, but as discontinous to semi-continuous lenticular bodies of

mineralization, termed “ore shoots”. Ten to twenty-four ore shoots, depending upon

interpretation (Gresham and Loftus-Hills, 1981) define nickel mineralization hosted

within the Kambalda Dome. These ore shoots do not represent the biggest Ni deposit

in the Kalgoorlie terrane based on total contained tonnes; however, the Kambalda

Dome contains high-grade, well-defined ore bodies that have been the focus of

active Ni sulfide mining since the 1970s. Significant research has been carried out

on the Kambalda Dome because it was the first komatiite-hosted Ni mineralized

system identified in Western Australia (Woodall and Travis, 1966).

This chapter summarizes previous research studies on the Kambalda Dome

mineralization and stratigraphy and is divided into four principle components: (1)

regional geology and tectonics, (2) Kambalda Sequence stratigraphy, (3) structural

evolution, and (4) alteration and metamorphism. These four components are not

mutually exclusive and all are inherently required for a comprehensive

understanding of komatiites and mineralization. This summary of the Kambalda

Dome will provide a framework for more in-depth work examining the chalcophile

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elements within ore forming systems, and the spatial correlation between

mineralization and ore forming signatures (see. Chapter 4. The scale of nickel

mineralized systems: Examination of platinum group element distribution in

the Long-Victor system, Kambalda Dome, W.A.).

Figure 3.1. Regional map of the Yilgarn Craton showing the South West and Youanmi Terranes and Eastern Goldfields Superterrane. Kalgoorlie, Kurnalpi and Burtville Terranes shown, and domains within each terrane shown in red. Nickel deposits hosted within the Yilgarn Craton shown as red squares. Modified from Cassidy et al. (2006).

3.2. Regional Geology and Tectonics

The Eastern Goldfields Superterrane (Myers, 1997) comprises elongate belts of

deformed and metamorphosed volcanic and sedimentary rocks, intruded by

extensive granitoid intrusions. These large granitoid-greenstone belts define the

Eastern Goldfields Superterrane. This superterrane is divided into three

chronological and fault bounded tectono-stratigraphic terranes (from west to east):

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Kalgoorlie, Kurnalpi, and Burtville (Fig. 3.1: Cassidy et al., 2006; Kositcin et al.,

2008).

The Kalgoorlie Terrane (formerly Norseman-Wiluna greenstone belt) extends in a

north-northwest direction for approximately 800 km, with a maximum width of 200

km. The terrane is bounded to the west by the regional eastward-dipping Ida Fault

and to east by the east-dipping Mount Monger Fault (Swager et al., 1997). The

Kalgoorlie Terrane is further subdivided into 10 structural-stratigraphic domains that

are separated by structural discontinuances: Coolgardie, Ora Banda, Kambalda,

Parker, Boorara, Moilers, Jundee, Wiluna, Depot, and Norseman Domains (Fig. 3.1:

Swager et al., 1992; Kositcin et al., 2008).

Figure 3.2. Stratigraphic column within the Kalgoorlie Terrane, with lithostratigraphic divisions shown on left. Modified from Lesher and Arndt (1995); Beresford et al. (2002); Krapez and Hand (2008). Stratigraphy adapted from Gresham and Loftus-Hills (1981); Cowden and Roberts (1990); Swager et al. (1992); Krapez (1997). Ages U/Pb SHRIMP from Claoue-Long et al. (1988); Krapez et al. (2000); Kositcin et al. (2008).

Each domain preserves a characteristic volcanic-sedimentary sequence. Correlation

between domains from the Kalgoorlie Terrane is possible due to the identification of

three lithostratigraphic divisions within the volcanic-sedimentary sequence

(Gresham and Loftus-Hills, 1981). These lithostratigraphic divisions are

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unconformity bounded sequences, and consist of: the Lower Kambalda Sequence,

Middle Kalgoorlie Sequence, and Upper Kurrawang and Merougil Sequences

(Krapez et al., 2000).

a. Stratigraphic sequences

i. Lower Kambalda sequence

The Lower Kambalda Sequence is characterized by basalt and komatiite unit, as

observed in the lithostratigraphic section at the Kambalda Dome, within the

Kambalda Domain. The Lower Kambalda Sequence comprises the: Lunnon Basalt

Formation, Kambalda Komatiite Formation (Silver Lake Member and Tripod Hill

Member), Devon Consols Basalt Formation, Kapai Slate Formation, and Paringa

Basalt Formation. The komatiites within the Kambalda Komatiite Formation are

constrained by a Re-Os isotope isochron age of 2706 ± 36 Ma (Foster et al., 1996).

A similar age of 2707 ± 4 Ma was obtained from dacite flows (Kositcin et al., 2008;

Claoue-Long et al., 1988; Nelson, 1995; 1997; 1998), interpreted to be

contemporaneous with the komatiites in the Boorara Domain (Trofimovs et al.,

2004). A detritial zircon age of 2692 ± 4 Ma was obtained from the Kapai Slate

Formation, representing an upper age limit for the komatiites within the Lower

Kambalda Sequence (Claoue-Long et al., 1988). The Lower Kambalda Sequence is

unconformably overlain by the Middle Kalgoorlie Sequence. The Lower Kambalda

Sequence is further expanded on as it hosts all known Ni mineralization in the

Kalgoorlie Domain.

ii. Middle Kalgoorlie sequence

The Middle Kalgoorlie Sequence (Fig. 3.1), also known as the Black Flag Group

comprises four unconformably bound sequences; and is characterized by andesitic,

dacitic and rhyolitic volcaniclastic and epiclastic rocks, with minor mafic flow units

and sedimentary rocks (Woodall, 1965; Travis et al., 1971; Hunter, 1993; Hand,

1998; Krapez et al., 2000; Krapez and Hand, 2008). The deposition of the Middle

Kalgoorlie Sequence is constrained by zircon U-Pb age determinations between

2686 ± 3 Ma and 2658 ± 3 Ma (Krapez et al., 2000). The Middle Kalgoorlie

Sequence represents deposition in a series of deep marine intra-arc basins, within an

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extensional to trans-tensional tectonic environment (Hand, 1998; Brown et al., 2001;

Krapez and Hand, 2008).

iii. Upper Kurrawang and Merougil sequences

The Upper Kurrawang Sequence comprises an upwards-fining succession of

conglomerate, sandstone and mudstone units, interpreted as high-density coarse-

grained to low-density fine-grained turbidites (Krapez et al., 2000). The Merougil

Sequence also consists of upward-fining successions of conglomerate and sandstone

units, but is interpreted as fluvial bar and channel systems (Krapez et al., 2000).

Both the Upper Kurrawang and Merougil Sequences formed through sediment

deposition within remnant ocean basins (Krapez et al., 2000).

b. Geodynamic setting of the Kambalda Domain

Overall, the three lithostratigraphic sequences of the Kalgoorlie Terrane (Lower

Kambalda, Middle Kalgoorlie, and Upper Kurrawang and Merougil Sequences)

represent a progression of crustal development initiated by plume related rifting,

followed by accretion and formation of a late basin. The Lower Kambalda Sequence

represents a regressive lava sequence that formed during the emplacement of a

mantle plume (Lesher and Arndt, 1995). Mantle plume emplacement beneath the

lithospheric crust (basement) resulted in primary melting of the plume head, and

generation of voluminous tholeiitic basalt flows (Lunnon Basalt: Fig. 3.2: Campbell

et al., 1989). Melting of the plume head is typically followed by hotter, deeper

melting of the plume tail, and generation of ultramafic komatiite magmas

(Kambalda Komatiite: Campbell et al., 1989). Progressive contamination,

fractionation and differential source melting generated the overlying basaltic

sequence at the top of the Lower Kambalda Sequence (Devon Consols and Paringa

Basalt). This initial volcanic-sedimentary sequence, was subsequently intruded by

late felsic intrusions. These late intrusions represent secondary crustal melts

generated during the thermal transfer of heat from the plume to the overlying crust

(Campbell et al., 1989).

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3.3. Lower Kambalda Sequence Stratigraphy

The Lower Kambalda Sequence, exposed in the Kambalda Dome area, has

represented a focal point of exploration and discovery for mining companies and

academics over the past 50 years. A continuous evolution of ideas and research has

been published on the stratigraphy and associated nickel mineralization of the Lower

Kambalda Sequence. Nickel mineralization in the Kambalda Dome area is confined

to the Lower Kambalda Sequence; consequently, further discussion of the

stratigraphy in this chapter is limited to units within the Lower Kambalda Sequence

(Lunnon Basalt Formation, Kambalda Komatiite Formation, Devon Consols

Formation, Kapai Slate Formation, and Paringa Basalt Formation: Fig. 3.2) and

basement. The overlying sequences will not be discussed further in this chapter

(Middle Kalgoorlie and Upper Kurrawang and Merougil Sequences).

a. Basement

The presence of basement beneath the Kambalda Dome stratigraphy is debatable due

to the lack of outcropping basement rocks. However, several lines of indirect

evidence (e.g. zircons, trace elements, isotopes) support the presence of an older

(>2.7 Ga) basement.

Xenocrystic zircons from the Lunnon Basalt, Devon Consuls, and Paringa Basalt

Formations display a range of ages (Compston et al., 1986). Zircon cores contain

measured ages of >3.4 Ga, with metamorphic overgrowths occurring between 3.2 to

3.1 Ga, and final overgrowths at 2.7 Ga (Compston et al., 1986). Zircon chemistry

and morphology indicate the xenocrystic zircons crystallized from felsic magma

(Compston et al., 1986). Zircons containing 3.4 to 3.2 Ga remnant cores were also

identified in felsic intrusions from the southeast Yilgarn Craton. These older zircon

cores may have formed from existing crust of intermediate composition, with crustal

reworking at 2.7 Ga (Oversby, 1975; Hill et al., 1989).

Trace element data from the Lower Kambalda Sequence (Lunnon Basalt, Kambalda

Komatiite, Denvon Consols, and Paringa Basalt Formations) contains varaible

abundances of light rare earth elements (LREE), a product of crustal contamination

and fractional crystallization (Arndt and Jenner, 1986). The Lunnon Basalt

Formation is characterized by Nb/Th and Nb/U ratios that indicate either minor

76


crustal contamination or the presence of a heterogeneous mantle source (Sylvester et

al., 1997). The Kambalda Komatiite Formation exhibits the lowest degree of

contamination, with flat to depleted LREE patterns. The overlying Devon Consols

and Paringa Basalt Formations exhibit higher degrees of contamination from a felsic

source, and are enriched in LREE (Arndt and Jenner, 1986).

The Sm-Nd isotopic system was initially used on whole-rock samples to determine

the eruption age of the Lower Kambalda Sequence. These Sm-Nd isotopic analyses

provided a wide range of ages from 3.2 Ga to 2.7 Ga (Claoué-Long et al., 1984;

Chauvel et al., 1985), with the latter age supported by U-Pb zircon work. The

disparity between the Sm-Nd isotopic age and zircon was attributed to a mixing

isochron between mantle melts and an older crustal contaminant (Chauvel et al.,

1985). Lead isotopic data obtained from late granititic intrusions supports the

presence of older re-worked crust (Oversby, 1975). However, this data indicates

differing crustal histories for the Kalgoorlie and Norseman Domains. The Kalgoorlie

Domain preserves a 2700 Ma and younger signature, whereas the Norseman Domain

preserves evidence for crustal development between 3300 and 2600 Ma.

The Norseman Domain occurs to the south of the Kalgoorlie and Kambalda

Domains and contains older, pre-2700 Ma rocks. Within the Norseman Domain, the

Penneshaw Formation (minor tholeiitic basalt, felsic lithic tuffs, minor greywacke

and shales), the overlying Noganyer Group (well-defined beds of banded iron

formation, conglomerate, sandstone, graphitic slate, biotite-andualusite schists), and

the Woolyeeryer Formation (basalt), all predate the Lower Kambalda Sequence

(Hall and Bekker, 1965; Doepel, 1973; Krapez et al., 2000). The felsic volcanic

rocks of the Penneshaw Formation contain zircon with ages of ~2930 Ma,

interpreted to represent an eruption age (Campbell and Hill, 1988). This age and

litho-stratigraphic succession is similar to that identified within the Lake Johnston,

Ravensthorpe and Forrestania Greenstone Belts in the Youanmi Terrane (Wong et

al., 1996; c.f. Chapter 5).

b. Lunnon Basalt Formation

The Lunnon Basalt Formation forms the footwall to the ultramafic Kambalda

Komatiite Formation, and is at least 2000 m in thickness. The Lunnon Basalt

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Formation is dominated by 2 to 30 m thick flow units, with a minimum lateral extent

of 500 km2 (Squire et al., 1998). Stratigraphically equivalent basalts to the Lunnon

Basalt Formation are observed throughout the Eastern Goldfields Superterrane, and

potentially represent up to 1.5 million km3 of erupted basalt flows (Lesher and

Arndt, 1995).

Four lithological facies are identified within the Lunnon Basalt Formation and

comprise pillowed basalt, massive basalt, basalt breccia, and sulfidic

metasedimentary rocks (Squire et al., 1998). Pillowed basalt flows comprise

approximately 45% of the Lunnon Basalt Formation stratigraphy, and commonly

exhibit well-defined pillow rims with radial and sub-concentric perlitic fractures and

associated periodic flow top breccia (Gresham and Loftus-Hills, 1981; Squire et al.,

1998). Pillowed flow intervals range in thickness from 3 to 15 m, with pillows

ranging from 30 cm to 5 m in size.

Massive basalt comprise approximately 45% of the Lunnon Basalt Formation and

are dominated by fine- to medium-grained basalt flows ranging in thickness from 10

to 140 m.

Based on volcanology, the Lunnon Basalts erupted in an aqueous environment with

at least 700 m of water depth (Squire et al., 1998). The eruption was generally

passive with magma transport through lava tubes on an average slope of <10°, with

paleo-flow towards the west, ranging from southwest to north-northwest (Squire et

al., 1998). Xenocrystic zircons within the basalts contain age ranges of 3422 ± 7 Ma

to 2667 ± 18 Ma, with the latter representing the youngest possible eruption age

(Compston et al., 1986).

The Lunnon Basalt Formation is chemically characterized as tholeiite, with

moderately high MgO contents, high Ni and Cr abundances, low incompatible

element concentrations, and minor LREE depletion (La/Smpmn (primitive mantle normalized) =

0.76 to 0.85: Redman and Keays, 1985). The Lunnon Basalt Formation is

subdivided into an upper and lower sequence, separated by a thin unit of interflow

sedimentary rocks. The lower sequence (high-Mg series basalts: HMSB) is slightly

less evolved (0.69% TiO2, 8.3% MgO) and contains olivine as phenocrysts (Redman

and Keays, 1985). The upper sequence (low-Mg series basalt: LMSB) is more

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evolved (0.91% TiO2, 7.8% MgO) and olivine phenocrysts are not observed

(Redman and Keays, 1985). Additionally, vesicles and amygdules are observed in

the lower sequence, but are absent from the upper sequence (Squire et al., 1998).

The Lunnon Basalt Formation formed during decompression melting of a mantle

plume in the subcontinental lithospheric mantle (Redman and Keays, 1985;

Campbell et al., 1989). Geochemical and isotopic studies indicate that source area

was depleted in LREE by a previous small-degree partial melt extraction; both

depleted mantle and primitive mantle sources were involved in the generation of the

basalts (Lesher and Arndt, 1995). Minor crustal contamination during basalt

emplacement is identified by Nb/Th ratios (Sylvester et al., 1997).

c. Metasedimentary rocks

Metasedimentary rocks occur throughout the entire Lower Kambalda Sequence.

Within the Lunnon Basalt Formation sedimentary rock units are commonly thin and

discontinous, representing accumulated interflow sedimentation. Rare sedimentary

structures (low-angle cross lamination, small scale scours and scour truncations) are

observed within the sedimentary rocks and indicate a very low energy subaqueous

environment, comprising either deep or quiet shallow conditions (Squire et al.,

1998). A thin semi-continuous horizon of sedimentary rocks is documented

approximately 100-200 m below the ultramafic contact of the Kambalda Komatiite

Foramtion. This horizon represents a stratigraphic marker that divides the less

evolved mafic lava flows from slightly more evolved lava flows within the Lunnon

Basalt Formation (Gresham and Loftus-Hills, 1981; Redman and Keays, 1985).

Sedimentary rock abundance increases towards the top of the Lunnon Basalt

Formation, where the unconformity between the Lunnon Basalt Formation and the

Kambalda Komatiite Formation is marked by a thin (≤ 5 m) sedimentary rock unit

(contact sediments of Bavinton, 1981).

Within the Silver Lake Member interflow sedimentary rocks (internal sediments of

Bavinton, 1981) are intercalated with komatiite flows, defining the boundary

between successive flows lobes. Age determinations of xenocrystic zircons from the

sedimentary rocks indicate an age of 2702± 4 Ma (Claoue-Long et al., 1988).

Sedimentary rocks within the Silver Lake Member are dominantly restricted to the

79


flanking environments, and are ubiquitously absent from the channel facies (ore

prism) and a 100-300 m wide zone flanking the channel (Bavinton, 1981). Interflow

sedimentary rocks have limited lateral continuity (200-500 m), with highly variable

thicknesses (Bavinton, 1981). A cumulative maximum sediment thickness is

attained at a distance of approximately 500 m from the channel facies, thinning

towards the channel (Bavinton, 1981)

The absence of metasedimentary rocks at the Lunnon Basalt Formation and

Kambalda Komatiite Formation contact represents a well-documented indicator of

the Ni ore environment within the Kambalda Dome area. Limited occurrences of

sedimentary rocks within the ore prism at Kambalda Dome are documented by

Bavinton (1979; 1981), but are restricted in spatial distribution. Interflow

metasedimentary rocks become rare in units overlying the Silver Lake Member, with

only a few thin discontinous intervals reported in the Tripod Hill Member

(Bavinton, 1979; Gresham and Loftus-Hills, 1981).

Three main types of metasedimentary rocks are documented within the Kambalda

Dome area, and are described in detail by Bavinton (1979; 1981). In order of

decreasing abundance, they consist of: (1) light grey to white siliceous chert, (2)

dark grey to black carbon-bearing slate, and (3) dark green chlorite and amphibolite-

rich non-siliceous sedimentary rock. The metasedimentary rocks typically contain

20-25 wt% iron sulfide in the form of pyrrhotite. The pyrrhotite occurs in thin (5-15

mm) layers and small trains of spherical sulfide nodules parallel to the apparent

layering. The total sulfide content increases up through the stratigraphy.

i. Sediment provenance

Metasedimentary rocks within the Lower Kambalda Sequence accumulated in an

unstable volcanically active and rifting basin (Bavinton, 1979). Bavinton (1979) and

Bavinton and Taylor (1980) identified several sources contributing to sedimentation.

These included two detrital components, an exhalative component, carbonaceous

material, and silica precipitation.

The first detrital component was extra-basinal and felsic in origin, as identified

through the presence of zircon, apatite, rutile, thorianite, monazite and baddelyite.

The first detrital component had an interpreted transport distance of 100s of

80


kilometers (Bavinton, 1979). The second detrital component was intra-basinal, and

is dominated by mafic and ultramafic fragments. These fragments were interpreted

as fragments of devitrified glass.

Cyclical metal grading is observed within each interflow sediment horizon.

Increases in S, Zn, Pb, Cu and Fe abundance are observed up through the sediment

horizon. Chromium, Ti, V, Ga, and Rb (± Ba, Sr) exhibit increasing abundances

down stratigraphy. Constant element abundances are observed for Co, Mn, Th, Nb,

Zr, Ce, La, Y, Ni/S, Ni/Zn. The cyclical variations were the result of changes in sea

floor exhalations (Bavinton, 1978; 1981; Bavinton and Keays, 1978; Bekker et al.,

2009).

Carbonaceous material within the metasedimentary rocks was interpreted as organic,

and sourced from primitive organisms within the water column (Bavinton, 1979).

Chemical precipitation of silica as amorphous silica gel is also proposed as some

sedimentary units contain up to 70% SiO2.

Bavinton (1981) identified a number of differences between what are termed as

“contact sediments” and “internal sediments”. Contact metasedimentary rock units

occur at lithologic contacts and are thinner than internal sedimentary rock units.

These contact metasedimentary rocks are enriched in Mg, Fe, and Mn, and are

depleted in Si, Ti, Ga, Na, Y, Cr, Pb, Zn, S, REE, Th, U, and Zr, relative to the

internal metasedimentary rocks. These differences were attributed to a shorter

accumulation period for the contact sediments relative to internal sediments.

d. Kambalda Komatiite Formation

i. Silver Lake Member Stratigraphy and volcanology

The Silver Lake Member consists of one or more laterally continuous komatiite flow

units, characterized by thick adcumulate channels, and thinner flanking

environments (Fig. 3.3: Hill et al., 1995; Lesher and Arndt, 1995; Beresford et al.,

2002).The Silver Lake Member varies in thickness from 50 to 200 m and comprises

approximately 1/3 of the Kambalda Komatiite Formation. Nickel sulfide

mineralization is associated with thickened (> 30 m) channels, whereas the flanking

facies are commonly barren. The basal flow channel is interpreted to have developed

81


along a shallow pre-existing linear topographic feature, that was later modified by

thermal-mechanical erosion, and deformation (Gemuts and Theron, 1975; Lesher et

al., 1984; Groves et al., 1986; Lesher, 1983; 1989; Stone and Archibald 2004; Stone

et al., 2005; Williams et al., 1998).

Figure 3.3. Block model showing distribution of contact sediments within the channel and flank facies. Modified from Gresham and Loftus-Hills (1981) and Stone and Masterman (1998).

Channel facies within the Silver Lake Member are up to 100 m in thickness, and are

characterized by olivine orthocumulate to mesocumulate compositions. The thick

olivine cumulate units are the product of sustained lava flow and continuous olivine

accumulation with variable abundance of chromite (Hill et al., 1995). Consequently,

these channel facies are interpreted as a highly dynamic portion of the komatiite

system. Olivine cumulates within the channel facies commonly represent the

occurrence of multiple composite cooling units. Spinifex layers (<1-10 m thick)

form at the channel flow top when flow velocity decreases. Therefore, spinifex

texture can post-date olivine accumulation in a portion of the channel facies (Hill et

al., 1995; Lesher and Arndt, 1995; Arndt et al., 2008).

Flanking komatiite flow facies within the Silver Lake Member typically have a

constant flow thickness of 15 to 35 m. Flanks exhibit a well-differentiated sequence

of A-zone spinifex and B-zone cumulates, dominated by orthocumulates. Thin

interflow metasedimentary rocks are also common in the flanking facies and are <1

to 10 m in thickness.

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The relationship between channel facies and flanking facies in the Silver Lake

Member was addressed by Beresford et al. (2002). Examination of the volcanic

facies identified thinning of the basal flow unit from east to west across the Victor

ore shoot (Fig. 3.4). The research also indicated that correlations are only possible

between the channel and flank facies in the basal unit. Flow units overlying the basal

flow could not be correlated, due to the transition from laminar sheet flow in the

basal flow to random channel breakouts, and lensing of successive flows in the

evolving komatiite volcanic pile.

Geochemistry

The Kambalda Komatiite Formation (Silver Lake and Tripod Hill Members) is

composed of Munro-type komatiites, with initial liquid compositions of up to 30%

MgO (Lesher et al., 1984; Lesher, 1989; Lesher and Arndt, 1995). Olivine in

equilibrium with the initial liquid, would have an approximate composition of Fo94,

which is similar to that observed within the channel facies olivine cumulate zones

(Ross and Hopkins, 1975; Lesher, 1989). The Kambalda Komatiite Formation

exhibits major and trace element variations consistent with the fractionation and

accumulation of olivine and minor chromite, which is akin to other Munro-type

komatiites (Barnes et al., 2004; 2007). Accumulation of pyroxene is not observed in

the channel facies; however, pyroxene (metamorphosed to amphibole) is prevalent

in the more fractionated flanking facies of the Silver Lake Member. Pyroxene is

interpreted to reflect a more fractionated magma composition in the flanking facies

(Lesher and Arndt, 1995).

Channel facies of the Silver Lake Member are characterized by > 35% MgO, and

inferred olivine compositions of Fo90-94 (Ross and Hopkins, 1975; Lesher, 1989;

Barnes et al., 2007). Spinifex within the channel facies ranges in composition from

16-31% MgO, 0.31-0.53% TiO2, 385-1610 ppm Ni, 1280-3670 ppm Cr, with REE

abundances characterized by La/Smcn ratios from 0.4-0.7 with slight LREE depletion

over HREE (Lesher and Arndt, 1995).

Flanking facies of the Silver Lake Member are characterized by lower MgO contents

(35-40% MgO), and olivine compositions of Fo89-91. Spinifex textured samples from

the flanks are characterized by 12-21% MgO, 0.41-0.55% TiO2, 424-1810 ppm Cr

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and 71-410 ppm Ni, with LREE enrichment relative to the channel spinifex (Lesher

and Arndt, 1995).

Overall, the Silver Lake Member exhibits LREE depletion (La/Smcn of 0.6-0.7),

chondritic ratios of MREE and HFSE ratios (Al2O3/TiO2 ~ 20, Ti/Zr 97, Gd/Ybcn 1;

Arndt and Jenner, 1986). However, crustal contamination is variable within the

lower units of the Silver Lake Member (Lesher and Arndt, 1995; Lesher et al. 2001).

Footwall metasediments are absent from beneath the channel facies, and may have

been assimilated by turbulent flowing magma and thermal-mechanical erosion

(Groves et al., 1986; Williams et al., 1998; Bekker et al., 2009). It is belieived that

these initial assimilation and contamination processes were followed by extensive

recharge within the channel, which effectively removed and diluted any geochemical

signature of the initial interaction. The flanking environments of the Silver Lake

Member exhibit limited thermal-mechanical erosion of the underlying

metasediments, yet record a crustal contamination signature of LREE enrichment.

This contamination likely occurred within the channel facies upstream of the

preservation site within the flank (Lesher and Arndt, 1995). Limited visible physical

evidence of contamination is present within the Kambalda Dome komatiites. Felsic

ocelli are identified along channel margins and are argued to be derived from

sediment assimilation (Frost, 1985; McNaughton et al., 1988; Frost and Groves,

1989; Frost, 1992).

Metasedimentary rock-ore association

Metasedimentary rocks are not commonly associated with Ni sulfide ore zones;

therefore, it is interpreted that sediment assimilation usually occurred within the

channel environment (Fig. 3.3). Turbulent magma within the channel environment

effectively scoured the sediment away and exposed the basalt footwall (Lesher,

1983, Lesher et al., 1984). Most Ni ore zones are characterized by a trough-like

feature hosting mineralization, with an abrupt transition to a barren contact,

commonly containing a 5-30 cm thick chlorite zone. Laterally, this basal chlorite

zone grades into planar metasediments which are dominantly cherty in appearance

(Bavinton, 1979). Sediment distribution in the Kambalda Dome area is best

summarized by Gresham and Lofuts-Hills (1981), with the statement

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“approximately 60-70% of the ultramafic-basalt contact at the Kambalda Dome is

sediment bearing, and the majority of sediment-free contact areas contain ore”.

Orthomagmatic mineralization

Nickel sulfide mineralization identified at the Kambalda Dome, Widgiemooltha

Dome, St. Ives, Tramways, Blue Bush is all hosted with the Silver Lake Member of

the Kambalda Komatiite Formation (Gresham and Loftus-Hills, 1981; Marston et

al., 1981; Barnes, 2006). Mineralization within the Silver Lake Member is spatially

restricted to the basal flow and within trough-like structures in the basal footwall

Lunnon Basalts (Fig. 3.3). Occurrences of Ni mineralization higher within the Silver

Lake Member stratigraphically (directly above the basal contact mineralization) are

documented in several ore shoots (e.g. Lunnon Shoot: Gemuts and Theron, 1975).

However, this “hanging-wall” mineralization comprises only a small fraction of the

total observed mineralization.

Nickel sulfide mineralization identified around the Kambalda Dome includes at least

24 separate ore shoots, with each ore shoot contributing to more than 350 ore

surfaces (Gresham and Loftus-Hills, 1981). The eastern flank of the Kambalda

Dome contains the Gibb, Long, and Victor ore shoots (Fig. 3.4). These ore shoots

are characterized by dominant basal contact mineralization with a strong structural

control on trough development or modification. Metasedimentary rocks are absent

from within the ore environment of the shoots. However, contact metasedimentary

rocks are observed in the flanking positions to the troughs. Hanging-wall

metasedimentary rocks are also observed in the flanks, and can stratigraphically

overlap the trough structures.

The Gibb ore shoot, up-dip of the Long ore shoot, is 1300 m in length and attains a

maximum width of 150 m (Fig. 3.4). The Gibb shoot is arc-like, plunging shallowly

to the north and south and terminated at the northern end by extensive felsic

intrusions. Ni sulfide mineralization is hosted within the basal komatiite flow, which

attains a maximum thickness of 50 m. The mineralization resides in a complex

trough structure dominated by basalt-basalt pinch-outs (Gresham and Loftus-Hills,

1981).

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The Long ore shoot (Fig. 3.4) occurs down dip of the Gibb ore shoot, has a known

plunge length of 2300 m, and remains open both up- and down-plunge. The Long

shoot attains a maximum width of 300 m, and is characterized by steep to sub-

vertical dips, but appears to shallow as it plunges to the south. Mineralization is

contained within a low-relief trough structure within the basal komatiite flow, which

attains a thickness of ~100 m.

The Victor ore shoot (Fig. 3.4) represents the down-plunge extension of the Gibb

ore shoot, separated by extensive felsic intrusions. The basal flow unit to the Victor

shoot attains thicknesses of >75 m within the trough, has a defined mineralized

plunge length of 850 m, and another 700 m of unmineralized extension. Nickel

mineralization within the Victor shoot occurs in a well-defined trough structure

~200 m in width, and is defined by high-angle normal faulting up-dip and low angle

reverse faulting down-dip.

Mineralization

The ore shoots within the Kambalda Dome comprise three ore settings: 1) basal

contact mineralization, 2) hanging-wall mineralization, and 3) structurally mobilized

mineralization. The mineralization is dominated by the basal contact type, with

lesser hanging-wall mineralization.

Basal contact mineralization occurs at the contact between the footwall basalts and

the overlying ultramafic flows. Basal contact ore surfaces typically occur within

embayments or depressions in the top of the footwall basalts, termed troughs or

channels (see Fig. 3.3: Lesher, 1983). Troughs within the Kambalda Dome area vary

in size, but are commonly narrow (<300 m) and elongate with lengths up to 2300 m

(Gresham and Loftus-Hills, 1981). Mineralization within major troughs is

dominantly continuous, but occurs as small (20-130 m) elliptical ore bodies in minor

troughs.

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Figure 3.4. Geological map of the Kambalda Dome area with mineralized Ni ore shoots projected to surface. Major ore shoots are labeled. Major structures identified shown in black. Map projection UTM zone 16 with WGS84 datum.

The formation of these trough or embayment features is contentious, with two

existing hypotheses. The first hypothesis infers that thermal-mechanical erosion by

the turbulent flowing ultramafic lavas was responsible for the down-cutting and

entrenchment of the ultramafic flow into the Lunnon Basalt (Lesher, 1983; 1989;

Beresford et al., 2005). The second hypothesis suggests a structural control on the

development of troughs. Troughs are formed either through pre-existing faults with

syn-eruption graben development, or during subsequent deformation of the

greenstone belt (Stone and Archibald, 2004; Stone et al., 2005). Evidence for each

87


hypothesis is extensively documented. The current ore surface configuration is likely

due to the combination of both pre-existing structures and topography, and the

erosive action of the ultramafic magma, followed by regional deformation.

Hanging-wall mineralization occurs stratigraphically higher, but usually within 100

m of the ultramafic-basalt contact (Gresham and Loftus-Hills, 1981). Mineralization

is documented in the third flow unit and higher (Gresham and Loftus-Hills, 1981).

However, this strata-bound mineralization exhibits a strong spatial relationship to

basal contact mineralization, and commonly grades laterally into stratigraphically

equivalent interflow metasedimentary units. Hanging-wall mineralization within the

Lunnon, Hunt, and McMahon ore shoots (Fig. 3.4) occurs at the contact of the basal

ultramafic flow unit and the second komatiite flow unit; where the Ni sulfide

mineralization resides on the A-zone spinifex unit of the basal flow (Groves et al.,

1986).

Structurally mobilized mineralization is characterized by the mechanical

transportation of sulfide into areas of dilation and lower tectonic pressure (e.g. fold

hinges, fault dilation zones, shear zones), away from the primary accumulation site

(Lesher and Keays, 2002). Mobilized mineralization is restricted to massive sulfides,

as massive sulfides are more ductile than disseminated sulfide and moved easier

(McQueen, 1981; 1987).

Style of mineralization

Basal and hanging-wall ore zones are commonly 1 to 3 m thick, with larger intervals

up to 10 m thick. Mineralization typically consists of a massive sulfide layer (<1 m

thick) overlain by a zone of matrix mineralization. Massive sulfide ore is defined as

>80% sulfide, and comprises pyrrhotite, pentlandite, pyrite and chalcopyrite, with

minor spinels concentrated at the basal or top contacts of the sulfide interval (Groves

et al., 1977). Massive ore is often banded, with alternating layers of pyrrhotite- and

pentlandite-rich bands. These bands form during recrystallization under directed

stress, and are generally parallel to the adjacent wall rock contacts (Ewers and

Hudson, 1972).

Matrix mineralization is defined as mineralization with 40 to 80% sulfide

abundance, with the remainder comprising serpentine or talc (pseudomorphs after

88


original olivine) within a continuous matrix of sulfide. In the Kambalda Dome area,

matrix mineralization exhibits a gradation of sulfide abundance from 40 to 60% at

the top of the matrix mineralization, to 60-80% sulfide at the base of the unit, and

ranges in thickness from 1 to 3 m (Gresham and Loftus-Hills, 1981; Keays et al.,

1981). Matrix mineralization exhibits a greater lateral continuity than the massive

sulfide mineralization.

Disseminated mineralization is characterized by 1 to 33% sulfide, but commonly 5%

interstitial sulfide within an ultramafic host. Disseminated mineralization is rarely

documented within the Kambalda Dome area, but is observed in both basal contact

and hanging-wall settings.

ii. Tripod Hill Member

The Tripod Hill Member ranges in thickness from 100-1000 m and comprises

approximately 2/3 of the Kambalda Komatiite Formation. The Tripod Hill Member

is thickest on the northern and western flanks of the Kambalda Dome, and thins on

the eastern flank of the Kambalda Dome and in the St. Ives, Tramways and

Bluebush areas to the south (Lesher and Arndt, 1995). The Tripod Hill Member is

composed of thin (1 to10 m) well-differentiated komatiite flow units. Flows exhibit

well-developed flow top breccia, thick spinifex zones, and well-developed B-zone

cumulates (Gresham and Loftus-Hills, 1981). Whole-rock geochemistry of spinifex

textured samples ranges from 15-32% MgO, 0.4-0.5% TiO2, 440-920 ppm Ni, and

2500-4020 ppm Cr (Lesher and Arndt, 1995). Overall, the flows are characterized by

a lower MgO content than in the underlying Silver Lake Member, the result of a

lower proportion of cumulate olivine. A trend of decreasing MgO content up-

sequence is also observed in the Tripod Hill Member rocks (Gresham and Loftus-

Hills, 1981). Metasedimentary rocks are generally absent from this member. The

well differentiated flow units, thinner flows, lower MgO content, and lack of

interflow metasedimentary rocks are interpreted to represent continuous outpouring

of compound lava flows at lower discharge rates than in the Silver Lake Member

(Lesher, 1989; Lesher and Arndt, 1995).

The Tripod Hill Member exhibits LREE enrichment relative to the Silver Lake

Member. Numerical modelling indicates that this enrichment is the product of minor

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(~ 5%) crustal contamination (Lesher and Arndt, 1995). The Tripod Hill Member

does not host any Ni mineralization within the sequence and displays normal

chalcophile element contents (Lesher et al., 2001), indicating that the magma was

not sulfur saturated during ascent or emplacement (Keays, 1982).

e. Devon Consuls Basalt, Kapai Slates, and Paringa Basalt Formations

A sequence of mafic volcanics and intrusive bodies overlies the Kambalda

Komatiite Formation. The transition from the underlying ultramafic units (Kambalda

Komatiite Formation) to the mafic volcanics is sharp within the Kambalda Dome

area, but interfingering transitions are observed elsewhere (St. Ives, Tramways:

Gresham and Loftus-Hills, 1981). The overlying mafic basalts are siliceous high

magnesium series basalts (SHMSB), comprising two members: the Devon Consuls

Basalt Formation (lower member) and the Paringa Basalt Formation (upper member:

Redman and Keays, 1985). The two members are separated by the Kapai Slate

Formation a thin (1-10 m) metasedimentary rock unit. Both the lower and upper

members contain abundant (up to 30%) phenocryst phases of olivine, pyroxene and

feldspar (Redman and Keays, 1985).

i. Devon Consols Basalt Formation

The Devon Consols Basalt Formation (lower member) has a total thickness of 60 to

100 m and is characterized by two lithologies: pillowed flows with felsic ocelli, and

massive komatiitic-basalt with minor pillowed phases (Ferguson and Currie, 1972).

The basalts are further classified into two geochemical groups: 1) high-Si, low-Mg

basalt characterized by 52-60% SiO2, 4-6% MgO, 6.7-7.4% FeOt (total), 0.71-0.83%

TiO2, 742-896 ppm Cr, 231-278 ppm Ni; and 2) low-Si, high-Mg basalt

characterized by 47-52% SiO2, 9-16% MgO, 9.8-12% FeOt, 0.64-0.77% TiO2, 576-

1173 ppm Cr, and 152-393 ppm Ni (Redman and Keays, 1985; Arndt and Jenner,

1986). Trace element data from the Devon Consuls Basalt exhibits flat HREE

primitive mantle normalized patterns with moderate LREE enrichment and no

apparent Nb depletion (Arndt and Jenner, 1986; Bateman et al., 2001). Chalcophile

element abundances within the basalt are constant and were not S-saturated

(Redman and Keays, 1985; Lesher et al., 2001).

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SHRIMP and U-Pb age determinations of xenocrystic zircons within the basalt

exhibit a range of ages from 3450 ± 3 Ma to 2652 ± 12 Ma (Compston et al., 1986).

Two geochrons are identified, where the oldest (3385 ± 10 Ma) represents the

crystallization age of the basement, and the younger (2693 ± 50 Ma) represents the

age of the basaltic volcanism (Compston et al., 1986).

ii. Kapai Slate Formation

The Kapai Slate Formation is characterized into two facies assemblages: a lower

carbonaceous shale, and upper incised turbidites and carbonaceous shales (Krapez et

al., 2000). Lithologically, the Kapai Slate Formation is composed of carbonaceous

shale, with minor pale chert and felsic volcaniclastic rocks (Bavinton, 1979;

Bateman et al., 2001). Xenocrystic zircon age determinations recorded minimum

ages of 2692 ± 4 Ma, and contain grains as old as 3441 ± 18 Ma (Claoue-Long et al.,

1988).

iii. Paringa Basalt Formation

The Paringa Basalt Formation (upper member) exceeds 500 m in thickness, and is

dominated by massive or pillowed mafic flows. Massive units are interpreted as

either massive sheet flows or intrusive units, and commonly contain medium- to

coarse-grained differentiated portions in the central sections (Gresham and Loftus-

Hills, 1981; Said and Kerrich, 2009). The Paringa Basalt rocks are characterized by

~ 10.6 wt% MgO, 10.7 wt% FeO, 13.0 wt% Al2O3 1070-2020 ppm Cr, and 280-470

ppm Ni, with strong LREE enrichment (Arndt and Jenner, 1986; Lesher and Arndt,

1995).

The Paringa Basalt Formation is geochemically subdivided into a lower enriched

basalt characterized as komatiitic-basalt to high-magnesium tholeiitic basalt

(HMTB), and an upper depleted basalt characterized as HMTB (Said and Kerrich,

2009). The lower enriched basalt is characterized by Mg# from 53 to 76, with LREE

enriched primitive mantle normalized patterns (Bateman et al., 2001; Said and

Kerrich, 2009). The upper depleted basalt exhibits a narrow compositional range

(Mg# 61-75), and a flat primitive mantle normalized pattern with slight LREE

depletion. The disparity between the lower enriched basalt and the upper depleted

basalt units was attributed to a mantle plume interacting with an asthenospheric

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mantle that had a component of older crustal recycled back into it (Said and Kerrich,

2009). Despite the complex petrogenetic history, the Paringa Basalts are not S-

saturated and preserve normal chalcophile element concentrations (Redman and

Keays, 1985; Lesher et al., 2001).

f. Intrusions

A complex sequence of intrusions post-date and cross-cut the mafic and ultramafic

units of the Lower Kambalda sequence stratigraphy. Geochronology work indicates

that the majority of these granitoid intrusions were emplaced continuously between

2.7 to 2.63 Ga, both coeval and post-dating extrusive felsic magmatism in the

overlying Middle Kalgoorlie Sequence (Brown et al., 2001). Consequently, granitoid

intrusions exhibit a range of deformation, from intense foliation and lineation to

non-foliated. Intrusion lithologies vary from biotite monzogranite, to granodiorite

and trondhjemite (Witt and Swager, 1989; Champion and Sheraton, 1993; 1997;

Witt and Davy, 1997). These intrusions are divided into five geochemical groups:

high-Ca (granodiorite, trondhjemite, monzogranite); low-Ca (granodiorite,

monzogranite, syenogranite); high-HFSE (granite); mafic (diorite, tonalite,

trondhjemite, granodiorite, granite); and syenitic (syenite, quartz syenite, monzonite:

Champion and Sheraton, 1997).

Granitoid petrogenesis is based on the trace element and isotopic geochemistry of

each of the identified geochemical groups. High-Ca granitoids are similar to other

tonalite-trondjhemite-granodiorite (TTG) systems, and may have been derived from

a garnet-stable, plagioclase-unstable mafic source (Martin, 1994). Probable melt

generation models involve partial melting of a thickened crust, or melting of a

subducted oceanic slab (Martin, 1994). The presence of older inherited zircons

within the high-Ca granitoids supports the melting of a pre-existing thickened

crust(Hill et al., 1992; Nelson, 1997). The low-Ca granitoids represent crustal

reworking and partial melting of tonalitic rocks (Brown et al., 2001). Similarly, the

high-HFSE granitoids are also derived from crustal melting (Champion and

Sheraton, 1997). The source for the mafic and syenitic groups is more ambiguous,

with possible contributions from both crustal and mantle sources (Champion and

Sheraton, 1997).

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3.4. Structural Evolution

The structural history of the Eastern Goldfields Superterrane (including the

Kambalda Dome area) has undergone significant examination. Deformation within

the Eastern Goldfields Superterrane comprises four main shortening events, D1

through D4 (Swager, 1989; Swager and Griffin, 1990; Swager et al., 1997); with

possible extension predating both the D1 and D2 shortening events (Weinberg et al.,

2003).

D1e (extension) is an important early deformation phase within Kalgoorlie

Terrane and is responsible for the emplacement of basalts, komatiites and

felsic volcanics of the Kambalda Dome (Williams and Whitaker, 1993).

However, D1e is largely obscured by subsequent progressive regional

deformation (Weinberg et al., 2003).

D1 is the earliest phase of folding defined within the Kalgoorlie Terrane. D1 is

characterized as localized tight to isoclinal, recumbent and napped folds,

with south over north thrust stacking that resulted in large-scale stratigraphic

repetition (Cowden and Roberts, 1990; Swager et al., 1997). This early phase

of deformation is identified within some sedimentary layers and at the basal

contact between the Lunnon Basalt Formation and Silver Lake Member of

the Kambalda Komatiite Formation. A S1 (schistosity) fabric is weakly

developed within these units, and trends N-NE and N-NW.

D2e represents an extensional collapse following D1 thrust stacking (Witt,

1994), and is responsible for the generation of shallow basins and late

sedimentation (e.g. Kalgoorlie and Kurrawang Sequences: Swager, 1997).

D2 is transpressive and commonly involves N-NW striking, upright to gently

plunging fold axes, that refold D1 structures. D2 deformation resulted in the

generation of open to tight, recumbent to inclined folds, that are commonly

dislocated by a shear zone sub-parallel to the axial plane. D2 deformation is

most evident in the re-entrant trough structures along the contact between

Lunnon Basalt Formation and Silver Lake Member. These D2 structures

occur at a low angle along the contact and are responsible for the dislocation

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and repetition of mineralization. Granitoid intrusion was synchronous with

D2 deformation (Nelson, 1997), along with the initial development of

regional subvertical foliation that was further intensified during D3.

D3 comprises transcurrent faulting and associated en echelon folding. D3

structures are characterized by the Boulder Fault and Kunanalling Shear

(Swager et al., 1990). Faults vary in width, and comprise anatomizing zones

of intensely foliated rock with pods of less deformed rock.

D4 is characterized by continued shortening with brittle dextral shearing and

regional scale development of fracture zones (Weinberg et al., 2003).

Structural deformation in the Eastern Goldfields Superterrane is prevalent in the

Kalgoorlie Terrane volcanic sequence. The sulfide mineralization associated with

the komatiites (e.g. Silver Lake Member) is particularly susceptible to mobilization

during deformation, alteration, and metamorphic events (McQueen, 1981; 1987;

Mason et al., 2003; Seat et al., 2004; Stone et al., 2004).

Potential mechanisms for sulfide mobilization in Ni sulfide deposits consist of both

mechanical and chemical transport. Overall, mechanical mobilization is brittle below

200-250°C and ductile above this temperature (McQueen, 1987), resulting in sulfide

mobility towards areas of lower pressure (fold hinges, faults: Marshall and Gilligan,

1993; Marshall et al., 2000). Mechanical mobilization readily moves massive

sulfides relative to the adjacent host rocks; as large contrasts in stress and strain rates

develop between the soft ductile massive sulfides and the hard brittle silicate

lithologies as pressure and temperature increase through green schist to amphibolite

grade metamorphic facies. Chemical transport of sulfides can also occur through

grain boundary diffusion, and fluid-aided dissolution and precipitation (Marshall and

Gilligan, 1987; Gilligan and Marshall, 1987; Plimer, 1987; McQueen, 1987).

Evidence for sulfide mobilization during deformation is found in the majority of

deposits around the Kambalda Dome. Structural deformation of the sulfide

mineralization and associated host rocks is evident in both macrostructures and

textures. The trough-like ore shoots identified at the Kambalda Dome represent

components of both the primary morphology (pre-existing topography and thermal-

mechanical erosion of a channel: Lesher, 1983), and the products of regional

94


deformation (Stone and Archibald, 2004). This regional deformation is observed in

the elongation of textures parallel to regional and parasitic fold plunge directions,

around the Kambalda and Widgiemooltha Domes (McQueen, 1987; Cowden, 1988;

Stone and Archibald, 2004; Stone et al., 2005). Deformation fabrics are also

observed in the sulfide ores, divergent veins, footwall stringers, sulfide-filled

fractures, breccia ores, sulfide layering, and in areas of extensive sulfide

recrystallization (Cowden and Archibald, 1987; McQueen, 1987; Cowden and

Roberts, 1990).

3.5. Alteration and Metamorphism

Alteration and metamorphism studies on the Kambalda Dome rocks were completed

by: Ross (1974), Barrett et al. (1977), Bavinton (1979), Marston and Kay (1980),

Gresham and Loftus-Hills (1981), Arndt and Jenner (1986), Barley and Groves

(1987), and summarized by Swager et al. (1990) and Lesher et al. (2001).

Komatiite units in the Kambalda Dome underwent seafloor hydrothermal alteration,

with alteration intensity (serpentine) increasing towards the top of the volcanic

succession (Barley and Groves, 1987). Later regional metamorphism, comprising

further serpentine and talc-carbonate alteration, preserved and overprinted the

primary seafloor alteration assemblage.

Regional metamorphism in the Kalgoorlie Superterrane (e.g. Kambalda Dome) is

dominated by upper greenschist facies, but variation from prehnite-pumpellyite to

lower amphibolite facies also present in some areas. Metamorphism occurred at low

to intermediate pressures (2.5 ± 1 kb to > 4.5 kb), and temperatures of 500 to 600°C,

with peak metamorphism during late D2 to D3 (Binns et al., 1976; Bavinton, 1979;

McQueen, 1981; Bickel and Archibald, 1984; Wong, 1986). Low-grade

metamorphism is associated with the central cores of the greenstone belts; whereas,

higher grades are observed along the periphery (Brown et al., 2001).

Complete replacement of the primary mineralogy by alteration assemblages has

occurred. However, many primary igneous textures and features are still visible

within lithological units around the Kambalda Dome. The limited development of

penetrative deformation fabrics also aided in the preservation of primary igneous

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textures. Progressive mineralogical change through hydrothermal alteration was

observed in the ultramafic lithologies (Gresham and Loftus-Hills 1981; Cowden and

Roberts 1990). Glass and pyroxene were progressively hydrated to form tremolite

and chlorite, whereas olivine was altered to serpentine. Antigorite was identified as

the dominant serpentine mineral, and either formed via direct serpentinization of

olivine at peak metamorphic conditions, or was the result of prograde

metamorphism of a lizardite assemblage. Progressive carbonation of the

serpentinites resulted in destruction of tremolite and antigorite, and the formation of

talc-dolomite and talc-magnesite assemblages. The only relic igneous minerals

present are chromite and rare portions of cumulate olivine within the Durkin and

Victor shoots of the Kambalda Dome (Gresham and Loftus-Hills, 1981, Lesher,

1983).

The effects of alteration and metamorphism on sulfide mineralization are variable

and highly dependent upon the metals involved, abundance of sulfide, and alteration

and metamorphic conditions. Research on the Mt. Keith disseminated ore body

demonstrated the progressive upgrading and Ni enrichment of the sulfides with

alteration intensity (Donaldson, 1981; Grguric et al., 2006). In contrast, progressive

alteration at the Black Swan deposit and select ore shoots at the Kambalda Dome

have shown limited effects on the composition of mineralization (Lesher and

Campbell, 1993; Barnes, 2004; Barnes et al., 2009).

Rocks with high MgO contents (komatiites) are typically reactive and susceptible to

element redistribution during low-grade metamorphism. Overall, komatiite systems

exhibit high loss on ignition (LOI) values, which are attributed to the addition of

volatiles to the system. More advanced alteration and metamorphism has variable

effects on the geochemistry and mobility of elements within a komatiite sequence

(Arndt and Jenner, 1986; Lahaye et al., 1995; Lesher and Arndt, 1995; Kerrich and

Wyman, 1996; Lesher and Stone, 1996). Large ion lithophile elements (LILE: Cs,

Rb, K, Na, Ba, Sr, Ca, Eu+2), with large ionic radii and low charge, are highly

susceptibility to mobilization during alteration events (Xie et al., 1993; Lesher and

Arndt, 1995; Lahaye et al., 1995). LILE mobility varies from local remobilization to

complete removal or enrichment within the system during low temperature seafloor

96


alteration and higher temperature hydrothermal alteration, and regional

metamorphism.

Rare-earth element and high field strength element mobility is dependent upon fluid

composition, with CO2-rich fluids having a stronger influence on element mobility

than H2O-rich fluids (Lahaye et al., 1995). Light-rare earth elements (LREE: La, Ce,

P, Nd,) are relatively immobile, yet may become mobile in the presence of CO2-rich

fluids. Limited mobility is observed at low fluid/rock ratios for the high field

strength elements (U+4, Th, Ta, Nb, Zr, Y, HREE), aluminum, the first period

transition elements (Sc, Ti, V, Cr, Mn, Co, Ni) and the highly siderophile elements

(Fe, PGE).

Evidence for S mobility within the ore forming systems of the Kambalda Dome was

initially proposed by Marston and Kay (1980), Seccombe et al. (1981), and

McQueen (1987). However, subsequent research by Keays et al. (1981), on ore

tenors and chalcophile elements in the silicate host rocks, did not show a strong

relationship between sulfur and metal abundance in lower sulfur samples (S <

0.2%). This lack of correlation between sulfur and metal abundance was attributed

to metamorphic and metasomatic redistribution of S (± Au and Cu). Work by

Seccombe et al. (1981) and Stone et al. (2004) identified S-loss by oxidation during

prograde metamorphism; where disseminated sulfides were more susceptible to S-

loss than net-textured and massive sulfides. Lesher and Campbell (1993) identified

post-crystallization mobilization of sulfur, with no systematic correlation between

degree of S-mobility and change in chalcophile element abundance. The chalcophile

elements (S, Au, Cu, Zn, Pb) are commonly mobile during complexation with

hydrothermal and metamorphic fluids.

The effects of alteration on PGE abundances in whole-rock samples are difficult to

identify, which leads to the inference of limited PGE mobility during alteration.

However, PGE-enriched hydrothermal ore deposits are identified (Lac des Iles, New

Rambler, Salt Chuck Intrusion: Hanley, 2006; Wilde, 2005). Additionally, platinum

group element mobility and fractionation during weathering is documented in a

number of locations (Cameron and Hattori, 2005; Traoré et al., 2008). Furthermore

the transport of PGE in aqueous solutions at very high salinities and oxidation states

is documented in laboratory experiments (Wood and Normand, 2008). These

97


observations support the mobility of PGE in S-poor rocks under certain conditions

(oxidizing high salinity fluids). Conversely, PGE immobility associated with

pervasive talc-carbonate alteration is documented in the Black Swan Ni system

(Barnes et al., 2004).

Platinum and palladium are incompatible in olivine and pyroxene and increase in

abundance in the residual melt with igneous fractionation. As a result of the

incompatibility a strong positive correlation with igneous fractionation is observed.

Both Pt and Pd are equally chalcophile, and the extraction of these elements from

magma by a sulfide liquid will not cause fractionation between the two elements.

However, Pt and Pd are argued to be fractionated by hydrothermal processes, with

Pd more easily mobilized. Consequently, deviations of inter-element ratios from

igneous ratios are attributed to fractionation of Pt and Pd during alteration and

metamorphism.

Mobilization of sulfide mineralization can result in fractionation of the chalcophile

elements, as observed in Cu-enriched divergent veins in the Kambalda Dome

deposits (McQueen, 1987). However, mobilization of sulfides can be isochemical, as

observed in the Perseverance (Agnew) Ni deposit (Barnes et al., 1988). Within the

Perseverance deposit, the similarity in composition between mobilized and in-situ

sulfides was the result of mechanical mobilization; as fluid phase transport would

result in significant chemical fractionation of the PGE (e.g. enrichment in Pd, Pt, Au

and Cu: Barnes et al., 1988).

Alteration and metamorphism is prevalent in all greenstone belts, but the presence of

alteration and metamorphism does not preclude the use of komatiite geochemistry.

A number of broad studies which focused on element mobility within high MgO

systems have identified elements that are susceptible to redistribution and those that

are not, allowing for quantitative research into lithogeochemistry (Lahaye et al.,

1995, Kerrich and Wyman, 1996 and Lesher and Stone, 1996).

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3.6. Summary

The Kambalda Dome is a doubly plunging anticline cored with a granitic intrusion,

located in the Yilgarn Craton of Western Australia. The Kambalda Dome has been a

focal point For Ni exploration, mining and research since 1970. This dome exposes

an Archean 2.7 Ga plume related extrusive volcanic and sedimentary sequence.

These volcanic and sedimentary rocks are divided into three sequences: Lower

Kambalda Sequence, Middle Kalgoorlie Sequence, and Upper Kurrawang and

Merougil Sequences. Komatiites are only identified in the Lower Kambalda

Sequence, comprising: basal tholeiitic basalts (Lunnon Basalt Formation), and

komatiites of the Kambalda Komatiite Formation (Silver Lake and Tripod Hill

Members). This first phase of magmatism in the Lower Kambalda Sequence is

interpreted as the initial plume melting beneath the crust. The overlying Middle

Kalgoorlie sequence reflects progressive melting and the transfer of heat from the

plume to the crust, as reflected in the evolving magma compositions. Post-plume

emplacement is dominated by late stage subsidence and basin development, as

recorded in the Upper Kurrawang and Merougil Sequences. Although the complete

sequence reflects plume-related volcanism and tectonics, economic Ni

mineralization is only hosted within the Silver Lake Member.

The Silver Lake Member comprises a sequence of komatiite flows. Volcanology and

lithogeochemistry have divided the komatiite flows into two volcanic facies: flank

and channel facies. Flank facies are characterized by thin flows (< 30 m) that are

well differentiated (A-zone spinifex and B-zone olivine cumulates). These flows are

the result of a single magma pulse with limited magma flow-through. Based on

geochemistry, flank facies are slightly more fractionated and more crustally

contaminated than the channel facies. Channel facies are characterized by thickened

(> 30 m) narrow sinuous bodies dominated by olivine cumulate rocks. Channel

facies exhibit limited differentiation, with only thin A-zone spinifex textured rocks

overlying massive olivine cumulates. Channel facies are interpreted as long lived

magma conduits transporting ultramafic magma through the system. This

interpretation is supported by the observed primitive lithologies and

lithogeochemistry. The dynamic setting of channel facies with continuous magma

99


transport is a critical factor in the generation of Ni sulfide mineralization hosted

along the base of these channels.

Nickel mineralization within the basal komatiite flows is hosted by shallow

depressions in the Lunnon Basalt footwall at the base of the channel facies.

Metasedimentary rocks, commonly located along the contact between the Lunnon

Basalt and the Silver Lake Member, are ubiquitously missing from the shallow

footwall depressions. The missing metasedimentary rocks are argued to have been

assimilated by magma flowing through the channel and caused the saturation of an

immiscible sulfide melt. The sulfide melt strongly partitioned the chalcophile

elements and accumulated on the channel floor due to density contrasts between

sulfide and silicate liquid. Accumulated sulfide occurs as both massive and semi-

massive ore bodies with lesser disseminated sulfides. It is these Ni sulfide

accumulations at the base of the Silver Lake Member that have been the focus of

exploration, active mining, and research since 1970.

Research in the Kambalda Dome area has been broad, covering: tectonics, plume

magmatism, volcanology, geochemistry, stratigraphy, orthomagmatic

mineralization, alteration, metamorphism, and structural deformation, just to name a

few. As such, the Kambalda Dome area provides an ideal setting and location to

further expand our research knowledge.

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3.7. References

Arndt, N.T., Jenner, G.A. 1986. Crustally contaminated komatiites and basalts from Kambalda, Western Australia: Chemical Geology, v. 56, p. 229-255

Arndt, N.T., Lesher, C.M., Barnes, S.J., 2008. Komatiite, Cambridge University Press, 467p.

Barley M.E., Groves, D.I., 1987. Hydrothermal alteration of Archaean supracrustal sequences in the central Norseman-Wiluna belt, Western Australia, In: Ho, S.E., and Groves, D.I., (eds.), Recent advances in understanding Precambrian Gold deposits: University of Western Australia, Geology Department and University Extension, Publication 11, p. 51-66.

Barley, M.E., Krapez, B. Brown, S.J.A., Hand, J., Cas, R.A.F., 1998. Mineralized volcanic and sedimentary successions in the Eastern Goldfields Province, Western Australia. In: Aust. Miner. Ind. Res. Associ. Final Report, 437.

Barnes, S.J., 2004. Komatiites and nickel sulfide ores of the Black Swan area, Yilgarn Block, Western Australia. 4. Platinum group element distribution in the ores, and genetic implications: Mineralium Deposita, v. 39, p. 752-765.

Barnes, S.J., 2006. Komatiite-hosted nickel sulfide deposits: Geology, Geochemistry, and Genesis. In Barnes, S.J. (ed.), Nickel Deposits of the Yilgarn Craton: Geology, Geochemistry, and Geophysics applied to Exploration. Society of Economic Geologists. Special Publication No. 13. p. 50-97.

Barnes, S.J., Gole, M.J., Hill, R.E.T., 1988. The Agnew nickel deposit, Western Australia: Part II. Sulfide geochemistry, with emphasis on the platinum group elements: Economic Geology, v. 83, p. 537-550.

Barnes, S.J., Hill, R.E.T., Evans, N.J., 2004. Komatiites and nickel sulfide ores of the Black Swan area, Yilgarn Craton, Western Australia. 3: komatiite geochemistry and implications for ore forming processes: Mineralium Deposita, v. 39, p. 729-751.

Barnes, S.J., Hill, R.E.T., Perring, C.S., and Dowling, S.E., 2004. Lithogeochemical exploration for komatiite-associated Ni-sulfide deposits: strategies and limitations: Mineralogy and Petrology, v. 82, p. 259-293.

Barnes, S.J., Lesher, C.M., Sproule, R.A., 2007. Geochemistry of komatiites in the Eastern Goldfields Superterrane, Western Australia and the Abitibi Greenstone Belt, Canada, and implications for the distribution of associated Ni-Cu-PGE deposits: Applied Earth Science, v. 116, p. 167-187.

Barnes, S.J., Wells, M.A., Verrall, M.R., 2009. Effects of magmatic processes, serpentinization, and talc-carbonate alteration on sulfide mineralogy and ore textures in the Black Swan disseminated nickel sulfide deposit, Yilgarn Craton: Economic Geology, v. 104, p. 539-562.

Barrett, F.M., Binns, R.A., Groves, D.I., Marston, R.J., McQueen, K.G., 1977. Structural history and metamorphic modification of Archean Volcanic-Type Nickel Deposits, Yilgarn Block, Western Australia: Economic Geology, v. 72, p. 1195-1223.

Batemann, R., Costa, S., Swe, T., Lambert, D., 2001. Archaean mafic magmatism in the Kalgoorlie area of the Yilgarn Craton, Western Australia: a geochemical and Nd isotopic study of the petrogenetic and tectonic evolution of a greenstone belt: Precambrian Research, v. 108, p. 75-112.

Bavinton, O.A., 1979. Interflow sedimentary rocks from the Kambalda ultramafic sequence. Unpublished PhD Thesis. Australia National University, Canberra.

Bavinton, O.A., 1981. The nature of sulfidic metasediments at Kambalda and their broad relationships with associated ultramafic rocks and nickel ore: Economic Geology, v. 78, p. 1606-1628.

Bavinton, O.A., Keays, R.R., 1978. Precious metal values from interflow sedimentary rocks from the komatiite sequences at Kambalda, Western Australia: Geochimica et Cosmochimica Acta, v. 42, p. 1151-1163.

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Table of Contents

3.1. Introduction ..................................................................................................... 71 3.2. Regional Geology and Tectonics ..................................................................... 72

a. Stratigraphic sequences ............................................................................. 74 i. Lower Kambalda sequence .................................................................. 74 ii. Middle Kalgoorlie sequence ................................................................ 74 iii. Upper Kurrawang and Merougil sequences ....................................... 75

b. Geodynamic setting of the Kambalda Domain ......................................... 75 3.3. Lower Kambalda Sequence Stratigraphy ........................................................ 76

a. Basement ................................................................................................... 76 b. Lunnon Basalt Formation .......................................................................... 77 c. Metasedimentary rocks .............................................................................. 79

i. Sediment provenance ........................................................................... 80 d. Kambalda Komatiite Formation ................................................................ 81

i. Silver Lake Member ............................................................................. 81 ii. Tripod Hill Member ............................................................................. 89

e. Devon Consuls Basalt, Kapai Slates, and Paringa Basalt Formations ...... 90 i. Devon Consols Basalt Formation ....................................................... 90 ii. Kapai Slate Formation ........................................................................ 91 iii. Paringa Basalt Formation ................................................................... 91

f. Intrusions ................................................................................................... 92 3.4. Structural Evolution ......................................................................................... 93 3.5. Alteration and Metamorphism ......................................................................... 95 3.6. Summary .......................................................................................................... 99 3.7. References ..................................................................................................... 101

List of Figures

Figure 3.1. Regional map of the Yilgarn Craton showing the South West and Youanmi Terranes and Eastern Goldfields Superterrane. Kalgoorlie, Kurnalpi and Burtville Terranes shown, and domains within each terrane shown in red. Nickel deposits hosted within the Yilgarn Craton shown as red squares. Modified from Cassidy et al. (2006). ................................................................ 72

Figure 3.2. Stratigraphic column within the Kalgoorlie Terrane, with lithostratigraphic divisions shown on left. Modified from Lesher and Arndt (1995); Beresford et al. (2002); Krapez and Hand (2008). Stratigraphy adapted from Gresham and Loftus-Hills (1981); Cowden and Roberts (1990); Swager et al. (1992); Krapez (1997). Ages U/Pb SHRIMP from Claoue-Long et al. (1988); Krapez et al. (2000); Kositcin et al. (2008). ......................................... 73

Figure 3.3. Block model showing distribution of contact sediments within the channel and flank facies. Modified from Gresham and Loftus-Hills (1981) and Stone and Masterman (1998). ............................................................................ 82

Figure 3.4. Geological map of the Kambalda Dome area with mineralized Ni ore shoots projected to surface. Major ore shoots are labeled. Map projection UTM zone 16 with WGS84 datum. ............................................................................. 87

Chapter 4. PGE Signatures in the Long-Victor system.

Chapter 4. The Size of Nickel Mineralized Systems: Examination of Platinum Group Element Distribution in the Long-Victor System, Kambalda Dome, W.A. Abstract

The komatiite-hosted Long-Victor nickel deposit, located on the eastern flank of the Kambalda Dome in Western Australia, was selected to investigate the size and geometry of the spatial and genetic correlation between localization of nickel-sulfide mineralization and the variability of chalcophile element (specifically the platinum group elements: PGE) abundance. The Long-Victor deposit is hosted in 2.7 Ga, Munro-type extrusive komatiites. Nickel mineralization is associated with the initial flows in the komatiite stratigraphy. Sampling was restricted to the basal flow unit, in order to isolate the ore forming signatures

Chalcophile element variability within a nickel mineralized komatiite system is the product of magmatic sulfur saturation. The resultant immiscible sulfide is enriched in PGE, and conversely the silicate melt is depleted of PGE. Enrichment and depletion are quantified relative to calculated normal background PGE abundances as a function of sample MgO content. The resultant data set (133 samples) exhibits variability in the whole-rock concentrations of the chalcophile elements. The majority of samples (47%) exhibit background values, enriched samples are the second most common with 39%, whereas, depletion is only recognized in approximately 14% of the dataset.

Within the basal flow, a spatial correlation is observed between chalcophile element (PGE) enrichment, depletion, and known nickel mineralization. Komatiite channel facies rocks hosting the mineralization exhibit both enrichment and background values; whereas the flank facies exhibits a complex distribution of background, depleted and enriched values. Chalcophile element enrichment occurs proximal to the channel mineralization, and displays a positive correlation between increasing enrichment and decreasing distance to mineralization. Chalcophile element depletion occurs in the flank facies, with maximum depletion occurring 340 metres from channel mineralization. This flank depletion displays a negative correlation between decreasing magnitude and increasing proximity to mineralization.

The spatial correlation between nickel mineralization and chalcophile element ore forming signatures (PGE enrichment and depletion) provides the first framework for development of a chalcophile element based vectoring tool for Ni sulfide exploration. The examination of the spatial distribution of chalcophile element abundances in mineralized systems is a crucial step for a better understanding of the nature of dynamic orthomagmatic nickel systems.

Keywords: komatiite, Munro-type, PGE, Ni, mineralization vector, Yilgarn,

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4.1. Introduction

Nickel sulfide deposits are becoming more difficult to find and the discovery rate of

new deposits is decreasing (Hronsky and Schodde, 2006). We propose that by

understanding the spatial correlation between chalcophile element ore forming

signatures and nickel mineralization hosted within a komatiite system, the size of the

mineralized system can be constrained. This information will allow the development

of vectors towards nickel sulfide mineralization for the application in both

greenfields and brownfields komatiite nickel exploration.

In Australia komatiite-hosted nickel sulfide deposits have been mined and explored

over the last 40 years providing a significant proportion of past nickel (Ni)

production and future reserves. Komatiite-hosted Ni mineralization was initially

identified at Kambalda, Western Australia in 1966 (Woodall and Travis, 1970;

Marston et al., 1981; Hronsky and Schodde, 2006). Since the discovery, the

Kambalda area (see Fig. 3.1) has become the type locality for komatiite-hosted

nickel deposits, with the geological and mineralization setting used for genetic

komatiite ore deposit models (Ross and Hopkins, 1979; Lesher et al., 1982; Arndt et

al., 2008). Subsequent research throughout the Eastern Goldfields Superterrane of

Western Australia (Fig. 3.1) and terranes hosting both mineralized and

unmineralized komatiites world wide, have provided a number of permutations of

komatiite systems, resulting in a comprehensive picture of mantle melting, tectonic

setting, flow field development, and mineralization processes as summarized by

Arndt et al. (2008).

Our understanding of the processes leading to ore formation in komatiite systems is

evolving, but the rate of discovery of economic Ni mineralization has decreased due

to the remaining prospective search spaces being under cover and at greater depths.

Recent advances in Ni targeting efficiency have occurred through the development

of new tools and evolution of existing tools, ranging from geophysical applications

(gravity and electromagnetic: Peters and Buck, 2000; Wolfgram and Golden, 2001;

Peters, 2006), volcanology (Hill et al., 1995; Hill, 2001) and lithogeochemistry

(major, trace and chalcophile elements: Lesher et al., 2001; Barnes et al., 2004;

2007).

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Commonly, Ni exploration methodologies use geophysical targeting, with follow-up

drill testing of conductive and magnetic anomalies. The presence of even minor

amounts of low-grade mineralization indicates that the preserved magmatic system

attained sulfur saturation, and identifies the potential for substantial Ni sulfide

accumulations within the system. Since komatiite-hosted nickel deposits are

typically small and lack an alteration halo, an unmineralized indicates there is no

mineralization within the sample but does not indicate the potential for the system to

host Ni mineralization.

Sulfur saturation is the key requirement for the generation of Ni sulfide

mineralization. A magma that attains sulfur saturation either through contamination

by the addition of S, Fe, or Si, fractionation, cooling, or a change in oxygen fugacity

develops an immiscible sulfide liquid phase within the silicate magma (MacLean,

1969; Haughton et al., 1974; Lesher et al., 1984; Mavrogenes and O’Neil, 1999; Li

and Ripley, 2005; Naldrett, 2005; Li et al., 2009). The chalcophile elements; nickel

(Ni), copper (Cu), platinum (Pt), palladium (Pd), iridium (Ir), osmium (Os),

ruthenium (Ru), and rhodium (Rh) strongly partition into the immiscible sulfide

liquid phase, becoming concentrated in the sulfide liquid. The progressive

accumulation of the immiscible sulfide liquid forms the Ni sulfide mineralization

(Ragamani and Naldrett, 1979; Campbell and Naldrett, 1979).

During the ore forming process and the partitioning of the chalcophile elements into

the sulfide phase, a reciprocal depletion of these metals in the silicate magma occurs,

resulting in both positive (enriched) and negative (depleted) signatures. Preserved

silicate magmas with chalcophile element depletion have been documented in

komatiites (Lesher et al., 2001). However, the magnitude of chalcophile element

depletion is less than predicted in many cases (Fiorentini et al., in press), and the

spatial correlation between chalcophile element depletion and mineralization within

komatiite systems is unconstrained.

We hypothesize that by understanding the spatial correlation between chalcophile

element mineralization signatures in the preserved silicate magma and Ni

mineralization hosted within a komatiite system, the magnitude of the signature

associated with the ore forming process can be constrained. Ultimately,

understanding the physical and chemical size of the mineralized systems will lead to

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the application of chalcophile element vectors for targeting Ni sulfide

mineralization.

To test this spatial correlation hypothesis, we have selected the Long-Victor Ni

deposit located on the eastern flank of the Kambalda Dome in Western Australia

(Fig. 4.1). The Long-Victor deposit was discovered in the early 1970s and since then

has seen almost continuous active mining, development and exploration in the area,

resulting in an extensive spatial drill hole data set. Additionally, the Long-Victor

deposit has been the focus of active research since discovery, resulting in a

significant body of literature, data, samples, and hypotheses which describe all

aspects of mineralization, geochemistry, volcanology, stratigraphy, structure, and

metamorphism (see Ross, 1974; Keays and Davidson, 1976; Ross and Hopkins,

1979; Gresham and Loftus-Hills, 1981; Keays et al., 1981; Keays, 1982; Lesher,

1983; Lesher et al., 1984; Redman and Keays, 1985; Arndt, 1986; Lesher, 1989;

Lesher and Arndt, 1995; Lesher and Stone, 1996; Moore et al., 2000; Lesher et al.,

2001; Beresford et al., 2002; Stone and Archibald, 2004; Stone et al., 2005).

The Long-Victor Ni deposit occurs within the basal komatiite flow unit, and

comprises two sub-parallel mineralized channels plunging shallowly to the north and

south, with an identified strike length of approximately 3 km. The local stratigraphy

is well-constrained and dips moderately to the east. Komatiite volcanic

environments, both proximal (channel) and distal (flank) to mineralization are

identified in the Long-Victor deposit and are shown in Figure 4.2. These volcanic

environments are interpreted to represent both passive (flank) and dynamic

(channel) components of a lava system that has extensive magma recharge and

magma flux through parts of the system (Donaldson et al., 1986; Hill et al., 1995;

Lesher and Stone, 1996; Barnes et al., 1999; Barnes, 2006). As a result, magmas that

were the source of the chalcophile elements (i.e. now depleted in chalcophile

elements) have flowed away (both linearly and laterally) from the site of ore

formation, which was subsequently occupied by the crystallization products of

magma unrelated to the ore forming process.

In this study, we present chalcophile element (PGE) data obtained from high-

precision analysis of sulfide-poor whole-rock samples from flank and channel

environments within the basal flow of the Long-Victor Ni deposit. The data are

numerically classified as either chalcophile element enriched or depleted, as

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observed on the basis of deviations from a normal background value. This

background value is a function of MgO content and deposit-specific derived

equations. Enrichment, depletion and background values are interpreted within a

komatiite volcanology framework (Hill et al., 1995; Hill, 2001; Lesher and Arndt,

1995; Lesher et al., 2001) and the spatial correlation between ore forming signatures

and Ni mineralization is examined.

4.2. Kambalda Dome

a. Geological setting Komatiite-hosted nickel sulfide deposits are identified throughout the greenstone

belts of the Yilgarn Craton in Western Australia (Fig. 3.1), specifically the 2.7 Ga

Kalgoorlie Terrane comprising the Wiluna, Agnew, Mt Clifford, Coolgardie, Ora

Banda, Boorara, Norseman and Kambalda Domains (Swager et al., 1995; Cassidy et

al., 2006; Kositicin et al., 2008). Within the Kalgoorlie Terrane, Ni deposits are

located discontinuously along the 600 km strike length. The northern portion of the

terrane is characterized by both large low-grade and smaller high-grade deposits

hosted within felsic volcanics (Mt Keith and Perseverance deposits: Barnes et al.,

1995; Grguric et al., 2006; Fiorentini et al., 2007). The central portion of the terrane

is characterized by extrusive komatiite systems hosted within intermediate to felsic

volcanics (Black Swan: Barnes et al., 2004; Dowling et al., 2004; Hill et al., 2004).

In the southern extent of the terrane, the Ni deposits are characterized by extrusive

komatiites with mafic footwalls (Widgiemootha and Kambalda Domes: Gresham

and Loftus-Hills, 1981; Lesher, 1983; Marston, 1984). This research focuses on the

Kambalda Dome area.

The Kambalda Dome within the Kalgoorlie Terrane (Fig. 3.1, Fig. 4.1), forms a

doubly plunging anticline cored by a series of felsic intrusions. Pioneering research

in the 1970s and 1980s has led to the Kambalda Dome becoming the type area for

komatiite-hosted orthomagmatic Ni deposits (Ross, 1974; Bavinton, 1979; Gresham

and Loftus-Hills, 1981; Keays et al., 1981; Lesher, 1983; Redman and Keays, 1985).

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Figure 4.1. Generalized geological map of the Kambalda Dome with nickel sulfide ore shoots shown in plan projection with major faults and fold axis shown. Area of the Long-Victor Ni deposit shown by dashed outline. Modified after Ross and Hopkins (1975) and Stone et al. (2005).

Since the initial discovery of the Lunnon ore shoot in 1966, on the south east flank

of the Kambalda Dome, at least 12 Ni mines have operated around the Dome in the

last 40 years; Ni ore was extracted from at least 24 ore shoots (Fig. 4.1: Gresham

and Loftus-Hills, 1981). The Long and Victor mines on the eastern flank of the

Dome commenced operation in the early 1970s, and have been mined almost

continuously.

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Within the Kambalda Dome area, several stratigraphic units are closely associated

with Ni mineralization. These units comprise the basal Lunnon Basalt Formation,

overlying sulfidic metasedimentary rocks, and the extrusive Kambalda Komatiite

Formation (specifically the lower Silver Lake Member: Fig. 4.2), as summarized by

Gresham and Loftus-Hills (1981). The Lunnon Basalt Formation is characterized by

a thick sequence of mafic volcanic flows, both pillowed and massive. These mafic

flow units commonly form the direct footwall to much of the Ni mineralization

(Gresham and Loftus-Hills, 1981; Redman and Keays, 1985).

The Kambalda Komatiite Formation overlies the Lunnon Basalt Formation and

comprises extrusive komatiite flows divided into two members: the lower Silver

Lake Member and upper Tripod Hill Member. The Silver Lake Member hosts all

known Ni mineralization in the Kambalda Dome area (Fig. 4.2), and consists of a

varying number of flows (from three to more than twenty: Stone and Archibald,

2004). Komatiite flows of the Silver Lake Member exhibit well-developed channel

and flank facies, a product of the flow field development and channelized magma

transport (Lesher et al., 1984). The upper Tripod Hill Member ranges in thickness

from 100 to 200 m and is dominated by numerous thin (1 to 10 m), well-

differentiated komatiite flows. No known significant Ni mineralization is hosted in

the Tripod Hill Member.

Sulfidic metasedimentary rocks are observed throughout the igneous stratigraphy,

but are commonly less than 10 m in thickness, and are discontinous laterally and

along strike (Bavinton and Keays, 1978; Bavinton, 1979; Gresham and Loftus-Hills,

1981). Metasedimentary rocks occur along the contact between the Lunnon Basalt

Formation and Kambalda Komatiite Formation, and are intercalated with flank

facies komatiite flows of the Silver Lake Member. However, metasedimentary rocks

are rarely observed at the base of the channels or stratigraphically above the

channels.

The lack of metasedimentary rocks within the channel facies of the Silver Lake

Member is argued to be related to the ore forming process and the generation of Ni

mineralization (Lesher et al., 1984). Nickel sulfide mineralization is restricted to the

channel facies of the Silver Lake Member, and is dominantly found at the basal

contact within shallow depressions (trough: Fig. 4.2). The mineralized contact is

mostly sediment free, suggesting that mineralization formed early in the flow history

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due to assimilation of sulfidic sediments originally hosted within the trough

(Huppert et al., 1984; Arndt, 1986; Williams et al., 1998; 2001). Local assimilation

of a sulfur-bearing contaminant induced the saturation of an immiscible sulfide

phase within the system and acted as a chalcophile element collector.

Figure 4.2. Local Kambalda Dome mine stratigraphy in an idealized cross-section showing the Lunnon Basalt Formation (footwall), and Kambalda Komatiite Formation comprising the Silver Lake and Tripod Hill Members. The Silver Lake Member exhibits thickened channel facies, thin flank facies, interflow metasedimentary rocks and Ni sulfide mineralization within a trough feature. Modified from Lesher and Groves (1984).

Mineralization within the Long-Victor deposit comprises three main ore shoots: the

Long, Victor and Gibb shoots. Recent development in the mine has identified down

plunge extensions of the Long (Moran shoot) and Victor (Victor South and McLeay

shoot). Although a series of mineralized shoots are identified, the shoots can be

grouped together into two sub-parallel discontinuously mineralized troughs (Fig.

4.3). The Victor trough comprises the Gibb, Victor, Victor South and McLeay ore

shoots, whereas the Long trough comprises the Long and Moran ore shoots (Figs.

4.1 and 4.3).

b. Structural modification

The Kambalda Domain has undergone significant structural modification with

complex folding and faulting observed in a number of deposits. Four main phases of

deformation are identified within the Kambalda Dome (Cowden and Archibald,

1987; Cowden and Roberts, 1990; Stone and Archibald, 2004; Stone et al., 2005).

Within the Long-Victor mine, the effects of deformation on the stratigraphy are

visible, with the most qualitative being the trust duplication of the Lunnon Basalt

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Formation within the overlying Silver Lake Member in the flanking environments.

Deformation within the channels is identified (Stone and Archibald, 2004); however,

the extent of structural complexity and duplication within the channels is difficult to

assess due to lithological ambiguity. Although the trough-like features within the

channels host the majority of nickel sulfide mineralization in the Kambalda Dome

area, the formation of the trough-like features is controversial. Trough formation has

been attributed to pre-existing topography, syn-volcanic faulting, thermal-

mechanical erosion of channels, or post volcanic deformation, with evidence

supporting aspects of all models (Barnes, 2006). Although structural complexity is

observed within the flank environments of the sequence (e.g. trust faulting, folding,

duplication of units), the volcanic stratigraphy is intact and specific stratigraphic

units are easily identified (i.e. the basal flow of the Silver Lake Member).

4.3. Chalcophile Element Abundance

In magma that has segregated an immiscible sulfide liquid producing sulfur

saturation the strong partitioning of the chalcophile elements (PGE: Pt, Pd, Ir, Ru,

Rh; Ni; Cu) into the sulfide phase depletes the silicate magma of these elements

(Ragamani and Naldrett, 1978; Campbell and Naldrett, 1979; Naldrett and

Campbell, 1982). If the sulfide liquid is completely removed from the silicate liquid

and the silicate liquid is isolated from further interaction with the magmatic system,

the resultant crystallization products will be depleted in the chalcophile elements.

However, if the immiscible sulfide phase accumulates within an igneous unit during

silicate crystallization, the unit will exhibit chalcophile element enrichment.

Chalcophile element depletion and enrichment form ore forming signatures. The

presence of ore forming signatures in a komatiite sequence indicates that

mineralization is present beyond the physical limits of mineralization, thus

increasing the potential volume that can be targeted to identify mineralization.

Volumetrically, the majority of komatiites exhibit normal background or baseline

abundances, due to extensive magma recharge that is unrelated to the ore forming

process (Lesher and Arndt, 1995; Lesher et al., 2001; Fiorentini et al., in press).

The identification of ore forming signatures is possible through the use of

normalized geochemical plots (e.g. chondrite, primitive mantle: Barnes et al., 1985;

Barnes and Naldrett, 1987), chalcophile element ratios (e.g. Pd/Ir, Pt/Ir, Ni/Cu,

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Cu/Pd: Barnes, 1990; Maier et al., 1998), or chalcophile-incompatible ratios (e.g.

Cu/Zr, Pd/Zr, PGE/Ti: Lightfoot and Keays, 2005; Fiorentini et al., in press). The

current application utilizes PGE/Tipmn ratios (where pmn is primitive mantle

normalized: McDonough and Sun, 1995) to remove the effects of magmatic

fractionation and olivine accumulation (Barnes et al., 2004; 2007; Fiorentini et al., in

press). This approach is based on the assumption that Pt, Pd, Rh, and the lithophile

incompatible trace elements are not fractionated from each other during olivine

fractionation and accumulation, and will exhibit a constant value (Pearce and Norry,

1977; Fiorentini et al., 2010).

The use of PGE/Tipmn ratios allows for the quantification and comparison of

mineralization signatures within a wide range of komatiite lithologies, from spinifex

to adcumulates, and covering the composition range of ~10 to 45 wt% MgO. At

concentrations >45 wt% MgO, the use of PGE/Tipmn ratios is limited by

compounding analytical uncertainties in the precision of both PGE and Ti at low

whole-rock abundances. PGE/Tipmn ratios provide systematic way to examine the

abundance and distribution of normal background, enriched and depleted

chalcophile element values from within the Long-Victor Ni mineralization system.

4.4. Materials and Methods

a. Sample selection A total of 118 samples were collected from drill core within the Long-Victor Ni

systems and associated flanks, channels, and ore shoots. In order to constrain the

spatial correlation between ore forming signatures and mineralization, a sampling

strategy was developed to utilize existing samples and maximize the extent of new

samples collected within the basal flow unit. A three dimensional (3D) geological

model was generated by the author to examine the spatial distribution of all existing

diamond drill hole data from the Long-Victor system. This computer generated

model was created using the commercial software package Leapfrog®. The main

surfaces and shells that were generated comprise: (1) the mafic footwall basalts

(Lunnon Basalt Formation), (2) Ni mineralization (massive sulfide, disseminated

sulfide, 0.4%, 1% and 3% Ni grade shells), and (3) an ultramafic shell (Fig. 4.3).

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Figure 4.3. 3D model of the Lunnon Basalt surface (shown in green) and 0.4% Ni grade shell (shown in red) as modeled with Leapfrog®. Victor trough and Long trough interpretations shown with dashed lines, with select ore shoots labeled (Gibb, Victor, McCleay, Long and Moran). Grey shading delineates approximate flank facies distribution. View looking west.

Modelling of individual komatiite flows (i.e. basal flow) was not possible, due to

inconsistencies and variable logging codes that were used over the exploration

history. Therefore, cross sections generated from an unconstrained Leapfrog®

model were compared with previously published sections generated by Lesher

(1983) and Beresford et al. (2002), and found to be consistent with previous

interpretations in areas where extensive structural modification had not occurred

(e.g. flanks).

Sampling of drill core from the Long-Victor deposit was conducted in 2006 and

2007. Initial sampling covered the complete strike length of the Long-Victor system

(~ 3 km), and examined both channel facies and flank facies from the Long and

Victor ore shoots (Fig. 4.3). The second round of sampling was focused on the

Victor channel and flanking environments, with the objective of data infill for the

initial sampling and historic samples (i.e. Keays, 1982; Lesher and Arndt, 1995;

Lesher et al., 2001).

Sampling throughout the Long-Victor system was largely restricted to the basal

flow, with additional samples from 2nd and 3rd overlying stratigraphic flows for

comparison. The basal flow was identified in the flanking environments by a distinct

contrast in lithology between the footwall Lunnon Basalt Formation, sulfidic

metasedimentary rocks, and the basal komatiite flow. The channel facies were

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characterized by homogenously textured adcumulates and mesocumulates.

Consequently, it was difficult to quantify displacement along shears, faults and

lithologic contacts, and difficult to identify duplication and thickening of lithologic

units. As such, the accuracy of stratigraphic position decreases with increasing

distance from the footwall contact. Both spinifex textured and B-zone cumulates

(see Pyke et al., 1973; Arndt et al., 1977) from the basal flow were identified in drill

core and sampled, thus providing profiles through the komatiite flow. Local infill

sampling complemented previous work by Keays (1982), Lesher and Arndt (1995),

Lesher et al. (2001). Infill samples also provided additional resolution of fine detail

within chalcophile element fractionation processes, due to crystallization within

spinifex horizons. Selected samples were visually sulfide free (low-sulfide),

carbonate unaltered and distal to cross-cutting felsic intrusive bodies, in order to

minimize possible contact metamorphic and regional alteration effects.

Samples were split with a diamond saw and a representative slab was retained for

documentation and further examination. Samples selected for geochemical analysis

were cleaned and cut to remove weathering effects accumulated during storage.

Samples were then coarse crushed at the University of Western Australia using a jaw

crusher, which was flushed with quartz, cleaned with a wire brush and acetone, and

blown dry with compressed air after each sample. The samples were packaged in

clear locking plastic bags and sent to Geoscience Laboratories (Sudbury, Canada)

for further milling and geochemical analysis.

b. Distance to mineralization The 3D computer generated geological model was utilized to examine the spatial

distribution of samples and geochemistry within the system. The model was also

used in the calculation of distances from geochemical samples to known

mineralization. An ore shell with a cut-off grade of 0.4% Ni defines mineralization

hosted within the channel environments (Fig. 4.3). This ore shell was used as a

proxy for channel environment in the calculation of vector distances, and orientation

from the channel environment. Distance vectors and orientation were calculated

using Euclidean norm from the three dimensional Cartesian coordinates of the

geochemical sample and the 2 m assay composites filtered for Ni >0.4%. For each

geochemical sample, an average of the closest three Ni occurrences > 0.4% was

used as the vector distance and orientation value. The resulting distances exhibit a

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range from a minimum distance of 1.5 m to a maximum of 509 m, with orientations

dominantly perpendicular (60-90° and 210-240°) to the current trend of the

mineralized channels (Fig. 4.4), with a lesser number of samples characterizing an

up and down stream component (150-180°).

Figure 4.4. Plot of distance (m) and azimuth of samples from nickel mineralization > 0.4 wt% Ni. Each data point is an average of the closest three distances and azimuths. Rose diagram showing distribution of azimuths with general trend (335°) of the Long-Victor channels shown by grey arrow, as observed in Figure 4.1.

c. Analytical techniques Samples were analysed at Geoscience Laboratories (Geolabs) in Sudbury, Ontario,

Canada in two batches. Major elements (Al2O3, CaO, Fe2O3, K2O, MgO, MnO,

Na2O, P2O5, SiO2, TiO2) were analyzed by wavelength dispersive X-Ray

fluorescence (XRF) on a 4 g sample which was fused to a glass bead with a borate

flux. Minor elements and some major elements (Al, Sb, Ba, Be, Bi, Cd, Ca, Ce, Cs,

Cr, Co, Cu, Dy, Er, Eu, Gd, Ga, Hf, Ho, Fe, La, Pb, Li, Lu, Mg, Mn, Mo, Nd, Ni,

Nb, P, K, Pr, Rb, Sm, Sc, Na, Sr, S, Ta, Tb, Tl, Tm, Sn, Ti, W, U, V, Yb, Y, Zn, Zr)

were analyzed by ICP-MS following a four acid (hydrofluoric, hydrochloric,

perchloric, and nitric) closed beaker digestion of 0.5 g sample. Additional and

duplicate analyses of select trace elements (As, Ba, Cr, Cu, Ni, Rb, Sc, Sr, V, Y, Zr)

were analysed by XRF on a 10 g sample pressed into a 40 mm pellet excited by a Rh

target. Total sulfur was measured by infrared adsorption during the combustion the

0.5 g sample in an oxygen-rich environment.

Platinum group elements (Pt, Pd, Rh, Ru and Ir) were analyzed by ICP-MS

following a nickel sulfide fire assay pre-concentration step, aqua regia dissolution of

the sulfide button and co-precipitation of the PGE with tellurium from a 15 g

sample.

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The precision of the analytical methods were evaluated through the use of internal

standards, blanks and duplicate analyses. Analytical precision was assessed with

duplicate analyses by the method outlined by Thompson and Howarth (1976). Major

elements exhibited median errors of <1% for the concentrations observed.

Chalcophile elements exhibited median errors of 8% Ir, 19% Ru, 13% Rh, 11% Pt,

and 7% Pd over normal unmineralized range of abundances, as summarized in

Barnes and Fiorentini (2008), and shown in Appendix C.

Multiple techniques are available and have been previously used for PGE analysis:

fire assay (Barnes and Fiorentini, 2008; Maier et al., 2009; Fiorentini et al., 2010; in

press), Carius tube isotope dilution (Puchtel and Humayun, 2001; Fiorentini et al.,

2004) and instrumental neutron activation analysis (Maier et al., 2004; Maier et al.,

2007). Detection limits and precision vary between the three methodologies. Current

studies commonly use fire assay due to lower cost and shorter preparation time.

Although, the Carius tube isotope dilution method provides better instrumental

precision, duplicate analysis by fire assay ICP-MS produces analytical results

reproducible within 5% (Barnes and Fiorentini 2008). Additional data from

published and unpublished work are also utilized in the study as summarized and

shown in Appendix A. This additional data was derived from similar, but not

identical analytical techniques, as described in the respective documents; therefore,

some discrepancies may exist. However, all data was carefully assessed and only

used if analytical methodologies were equivalent or superior to fire assay ICP-MS.

4.5. Results

Since discovery, a total of 133 publicly available sulfide-poor samples have been

analyzed for PGE from the basal flow of the Long-Victor system. Of these samples,

118 were generated in this study, with the remaining 15 sample analyses derived

from previous research (Keays, 1982; Lesher and Arndt, 1995; Lesher et al., 2001).

The new samples were dominantly collected from the basal flow in the Victor area

(Victor, Victor South and McLeay: Fig. 4.3), and comprise up-dip flank, channel

and down-dip flank areas, as interpreted from core logging (Fig. 4.3). Complete

sample coverage represents approximately one square kilometre of surface area,

with a mineralized strike length of 1300 m. Within the Victor channel setting, the

up-dip flank is better characterized, with sampling extending up to 870 m from the

channel mineralization. Conversely, the down-dip flank has yet to be tested to the

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same extent by exploration drilling. Accordingly, sampling attains a maximum

distance of 450 m from channel hosted mineralization on the down dip flank.

Previous studies on komatiite-hosted Ni deposits have rigorously filtered the

geochemical data sets to remove samples that are enriched in Ni, or have sulfur

contents greater than 0.25%. Samples containing more than 0.25% S are typically

excluded, as this is the estimated capacity of S based on the solubility of S in high-

MgO komatiite magmas. Consequently, values >0.25% S are believed to contain an

immiscible sulfide component (Fiorentini et al., in press). This research study does

not exclude data based on a mineralization filter (elevated Ni) or on a S-filter. The

reasoning behind this decision is based on two principles: firstly chalcophile element

enrichment is an ore forming signature, and secondly S-mobility within sulfide ores

is documented in other deposits (Marston and Kay, 1980; Seccombe et al., 1981;

McQueen, 1987; Stone et al., 2004b). The data for this research study have been

divided into low-S (<0.25 wt%) and high-S (>0.25 wt%) for comparison purposes.

Samples are grouped as either flank or channel facies, and as spinifex textured (spfx)

or B-zone cumulates (Bz). Channel and flank discrimination is based on the spatial

distribution of known mineralization, interflow metasediments, flow thickness, and

relative thickness of the A-zone spinifex and B-zone cumulates, as outlined by

Beresford et al. (2002).

a. Major and trace element geochemistry Major and trace element abundances from the basal flow are summarized in Table

4.1 as maximum, minimum and median values of major, trace and chalcophile

elements from flank (spinifex and B-zones) and channel facies (spinifex and B-

zones). The complete whole rock data set generated in this thesis is in Appendix B.

These major and trace element abundances exhibit a strong olivine control on

distribution. Both flank (n=18) and channel (n=17) are characterized by a median

spinifex composition of 25 wt% MgO and 11 wt% FeOtot, with maximum MgO

content measured in the B-zone cumulates of 47 wt% (flank) and 49 wt% (channel).

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Table 4.1. Summary of geochemistry for the basal flow at Long-Victor: Median (Med), Maximum (Max), Minimum (Min), Number of samples (N). Data filtered for S<0.25 wt%. Oxides are recalculated to anhydrous conditions and reported in wt%, metals and trace elements are reported as ppm unless denoted * then ppb.

Channel B-Zone Flank B-zone Channel Spinifex Flank Spinifex Wt% Med Max Min N Med Max Min N Med Max Min N Med Max Min N

SiO2 44.1 49.7 40.5 39 44.7 49.8 42.3 21 46.6 51.4 44.4 11 46.9 47.7 45.2 8

TiO2 0.14 0.29 0.04 45 0.17 0.31 0.08 23 0.36 0.54 0.31 11 0.45 0.53 0.31 8

Al2O3 2.55 5.67 1.68 39 3.14 6.36 1.77 21 7.66 11.2 6.29 11 9.57 10.9 6.56 8

FeO 7.33 10.6 5.95 39 8.14 9.79 6.01 21 9.59 10.8 8.97 11 11.6 12.1 9.8 8

Fe2O3 0.22 0.8 0.03 39 0.39 0.89 0.08 21 1.11 1.33 0.86 11 1.43 1.5 1.05 8 FeO tot 7.52 11.3 6.06 39 8.48 10.6 6.48 21 10.6 12 9.97 11 12.9 13.5 10.7 8

MnO 0.15 0.2 0.08 39 0.17 0.24 0.12 21 0.19 0.22 0.12 11 0.25 0.29 0.22 8

MgO 43 49.1 32.6 45 39.5 47.5 27.3 23 26.6 30.9 13.7 11 18.8 27.8 17.8 8

CaO 0.98 7.58 0.11 39 1.84 8.63 0.1 21 7.36 11.7 3.33 11 7.83 8.96 6.27 8

Na2O 0.03 0.13 0.01 38 0.05 0.13 0.01 21 0.14 2.02 0.06 11 0.35 0.52 0.1 8

K2O 0.01 0.34 0.01 38 0.01 0.05 0 21 0.99 4.29 0.01 11 2.75 4.32 0.02 8

Cr2O3 0.27 1.1 0.23 39 0.29 0.39 0.16 21 0.36 0.5 0.13 11 0.26 0.46 0.16 7

P2O5 0.01 0.02 0 39 0.01 0.02 0.01 21 0.02 0.06 0.02 11 0.04 0.08 0.02 8

S* 0.17 0.3 0 45 0.18 0.3 0 23 0.2 0.27 0.01 11 0.05 0.27 0.02 8

Channel B-Zone Flank B-zone Channel Spinifex Flank Spinifex ppm Med Max Min N Med Max Min N Med Max Min N Med Max Min N

Ni 2694 3730 580 41 2302 3117 292 23 966 1416 109 11 516 1197 276 7

Cu 19 70 1 43 33 99 1 21 35 169 2 11 24 81 1 8

Co 101 142 0 39 0 119 0 21 85 93 0 11 0 84 0 8

Cr 1867 7524 1562 39 1989 2652 1106 21 2453 3433 866 11 1811 3155 1108 7

Zn 52 231 0 39 0 161 0 21 72 113 0 11 0 230 0 8

Ir* 4.7 14.1 1.19 45 2.54 8.58 0.13 23 0.95 1.83 0.01 11 0.32 0.99 0.18 8

Ru* 3.48 32.2 0.26 45 3.44 21.1 0.16 23 3.96 5.85 0 11 0.48 3.3 0.33 8

Rh* 0.68 9.51 0.03 45 0.62 6.15 0.01 23 1.2 2.72 0 11 0.19 1.43 0.07 8

Pt* 3.45 48.5 0.7 44 4.2 38 0.36 23 8.48 14.4 0 11 2.78 12.1 1.72 8

Pd* 4.56 70.7 0.23 45 4.48 52.9 0.19 23 8.46 15.1 0.14 11 2.16 11.6 0.67 8

Au* 9.04 87.8 0.82 39 8.45 23.5 1.32 22 3.23 36 0.6 8 1.04 30.8 0.59 7

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Table 4.1 continued.

Channel B-Zone Flank B-zone Channel Spinifex Flank Spinifex ppm Med Max Min N Med Max Min N Med Max Min N Med Max Min N

Th 0 0.09 0 16 0 0 0 15 0.04 0.16 0 5 0.11 0.24 0 7

Nb 0.24 0.82 0.13 29 0.24 0.45 0.17 19 0.62 0.82 0.44 9 0.71 2 0.45 8

La 0.24 0.66 0.16 37 0.37 0.71 0.19 21 0.74 1.42 0.44 9 0.93 4.99 0.53 8

Ce 0.7 1.63 0.42 37 0.92 1.84 0.53 21 1.94 3.39 1.19 9 2.32 9.17 1.6 8

Pr 0.11 0.26 0.07 37 0.16 0.29 0.08 21 0.34 0.54 0.24 9 0.41 1.15 0.29 8

Nd 0.62 1.33 0.37 37 0.88 1.5 0.34 21 1.98 2.98 1.37 9 2.35 5.34 1.73 8

Hf 0.29 0.45 0.12 22 0.29 0.43 0.22 6 0.56 0.65 0.52 5 0.69 0.69 0.69 1

Zr 9.35 16.4 5.79 37 10 16.6 6.91 21 20.4 27.7 16.8 9 24.5 32.3 16 8

Sm 0.24 0.48 0.14 37 0.32 0.55 0.13 21 0.75 1.13 0.53 9 0.9 1.37 0.68 8

Eu 0.09 0.18 0.05 37 0.11 0.19 0.04 21 0.25 0.55 0.17 9 0.37 1.22 0.17 8

Gd 0.34 0.65 0.21 37 0.44 0.83 0.19 21 0.97 1.44 0.72 9 1.28 1.77 0.91 8

Tb 0.06 0.12 0.04 37 0.08 0.15 0.04 21 0.19 0.28 0.14 9 0.26 0.32 0.17 8

Dy 0.44 0.81 0.25 37 0.57 0.99 0.25 21 1.29 1.95 0.97 9 1.77 2.08 1.17 8

Ho 0.1 0.18 0.05 37 0.13 0.22 0.06 21 0.29 0.41 0.21 9 0.39 0.65 0.26 8

Y 4.39 5.82 3.43 15 4.66 7.73 4.45 15 9.67 12.3 7.26 5 10.6 13.5 8.53 7

Er 0.27 0.51 0.16 37 0.36 0.65 0.18 21 0.85 1.18 0.64 9 1.1 1.27 0.75 8

Tm 0.04 0.08 0.02 37 0.05 0.1 0.03 21 0.13 0.18 0.09 9 0.17 0.2 0.12 8

Yb 0.29 0.51 0.18 37 0.35 0.63 0.19 21 0.86 1.2 0.64 9 1.08 1.3 0.74 8

Lu 0.04 0.08 0.03 37 0.05 0.1 0.03 21 0.14 0.19 0.1 9 0.17 0.2 0.11 8

The similarities and differences between channel and flank facies and spinifex

textured and B-zone cumulates are visible in Table 4.1 and Figure 4.5 FeO wt%

versus MgO wt%. Significant overlap of channel and flank B-zone cumulates is

observed. Similarly, spinifex textured samples from the channel and flank exhibit

overlapping distributions.

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Figure 4.5. Plot of FeOtot versus MgO wt% for the basal flow within the Long-Victor area. Oxides as wt% and recalculated to anhydrous. Samples are characterized as channel facies B-zone (Ch Bz) and spinifex (Ch Spfx) and flank facies B-zone (Fl Bz) and spinifex (Fl Spfx). Volcanic flow facies fields from Barnes (2006). Modelled olivine compositions (Fo) in pure adcumulate shown on right hand side. Magma liquids in equilibrium calculated olivine compositions (Fo) shown on left hand side and along top.

Negative correlations are observed between MgO and TiO2, Al2O3 (Fig. 4.6). Large

ion lithophile elements exhibit negative correlations with MgO, with moderate

scatter attributed to secondary mobility. The trace elements exhibit flat primitive

mantle normalized patterns with minor light rare earth depletion (Fig. 4.7). Trace

element total abundances exhibit a negative correlation with MgO, reflecting both

the proportion of trapped liquid and fractionation.

Figure 4.6. Plot of Al2O3 and TiO2 versus MgO for the basal flow within the Long-Victor area. Oxides as wt% and recalculated to anhydrous. Samples are characterized as channel facies B-zone (Ch Bz) and spinifex (Ch Spfx) and flank facies B-zone (Fl Bz) and spinifex (Fl Spfx).

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Figure 4.7. Median primitive mantle normalized trace element plots of the samples from the basal flow in the Long-Victor area. Samples divided into channel and flank facies, and spinifex textured and B-zone cumulates.

Major and trace element geochemistry characteristics within the basal flow are

similar to other published results from the Kambalda Dome area (Keays et al., 1981;

Lesher et al., 1981; Lesher, 1983; Redman and Keays, 1985; Arndt and Jenner,

1986; Lesher and Arndt, 1995; Lesher and Stone, 1996; Lesher et al., 2001) and

other komatiite systems (Barnes et al., 2004; Barnes, 2006; Barnes et al., 2007).

b. Chalcophile element geochemistry Chalcophile elements (Ni, Cu, Co, Pt, Pd, Ir, Ru, Rh) were measured in 133 samples

from the basal flow of the Long-Victor area, with summary analytical results

presented in Table 4.1, and complete analyses in Appendix B. Within the data set,

the chalcophile elements exhibit a wide range of abundances from below detection

limits for individual PGE (Pt, Pd, Ir, Rh, Ru), to anomalously high >260 ppb for

total PGE. Primitive mantle normalized chalcophile element abundances (Fig 4.8)

exhibit a wide range of patterns in both channel and flank settings and within

spinifex textured samples and B-zone cumulates. Characteristic concave up and

concave down normalized chalcophile element patterns are observed, indicative of

sulfide segregation (Maier et al., 1998).

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Figure 4.8. Primitive mantle normalized chalcophile element metal diagrams for the basal flow within the Long-Victor area. Spinifex textured samples shown in black and B-zone cumulate samples in black. Normalizing values from McDonough and Sun (1995).

Spinifex textured samples from the channel and flank areas display similar ratios of

the chalcophile elements normalized to primitive mantle (Ni/Cupmn 0.38 and 1.2,

Ni/Pdpmn 598 and 980 Ir/Pdpmn 0.15 and 0.35, Ru/Pdpmn 0.28 and 0.29, respectively).

These spinifex textured flank and channel samples also exhibit PGE abundances that

are similar to other documented 2.7 Ga Munro-type komatiites (Table 4.2:

Fiorentini et al., 2010; Maier et al., 2009).

Table 4.2. Average (n=19) chalcophile element abundances, MgO and TiO2 content of spinifex textured samples from the Long-Victor area. (TiO2 and MgO as wt%, Ni, Cu, Co, Cr, Zr, Gd as ppm, and Ir, Ru, Rh, Pt, Pd, Au as ppb).

TiO2 MgO Ni Cu Co Cr Ir Ru Rh Pt Pd Au Zr Gd

0.41 22.3 747 45 87 2144 0.76 2.29 0.85 6.22 5.96 7.43 22.62 1.18

Overall Ni and Ir exhibit a positive correlation with MgO, whereas Pt, Pd and Rh

exhibit strong to moderate negative correlations (Fig. 4.9). Ruthenium exhibits a

complex pattern with a positive correlation observed between 10-25 wt% MgO and

no apparent correlation with MgO above 25 wt%. Copper exhibits extensive scatter,

with the general trend of lowest abundances at high MgO and increasing abundances

at lower MgO contents. Gold does not exhibit any systematic relationship between

abundance and MgO content.

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Figure 4.9. MgO wt% versus chalcophile element for all samples from the basal flow. Visual trends shown by dashed lines.

Titanium normalized PGE (Pt, Pd, Rh) diagrams exhibit a central cluster of data

points along constant values for Pt and Pd, and a slightly decreasing value for Rh

(Fig. 4.9). In Figure 4.10 samples that plot above and below the central values and

contain MgO contents >25 wt% are dominated by B-zone cumulates from both the

channel and flank facies. Samples plotting below the central values and contain <25

wt% MgO are largely restricted to the flank facies spinifex textured samples.

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Figure 4.10. PGE/Tipmn versus MgO wt% for all samples from the basal flow. Samples with S > 0.25 wt% on the left hand side and samples with S < 0.25 wt% on the right hand side. Samples are subdivided based on flow facies (channel = Ch, and flank = Fl) and komatiite flow facies (B-zone cumulates = Bz, and spinifex textured = Spfx).

The chalcophile element abundance of a sample can be controlled by the

accumulation of sulfide. Sulfur abundance is generally imprecise as a mineralization

filter (e.g. S-loss or S-gain). Although sulfur abundance is a poor indicator of

mineralization the samples are divided into two groups (sulfide-poor and sulfide-

bearing) utilizing a S<0.25 wt% cut off to compare and contrast the two sub-sets of

data (Fig. 4.10). Approximately half of the samples (67 of 133) contain S<0.25 wt%,

with the remaining (66 of 133) containing S>0.25 wt%.

i. Sulfur-bearing The S-bearing samples (S>0.25 wt%) range from 0.25 to 4.0 wt% S with a median

value of 0.27 wt%. Sulfide-bearing samples generally exhibit the same trends

relative to MgO as those observed in the comprehensive data set, with the exception

of a minor portion of the data that exhibit elevated metal abundances and plot well

above the crystal fractionation and accumulation trend lines (Fig. 4.9). A general

positive correlation is observed between S and chalcophile element abundance.

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However, four samples exhibit increased chalcophile element abundance at S <0.5

wt% as shown in Fig. 4.11 with Pt versus S.

Inter-chalcophile element relationships exhibit a number of trends (Fig. 4.11).

Nickel and Ir exhibit a positive correlation. Nickel and the other PGE (Rh, Ru, Pt,

Pd) exhibit two contrasting trends. A negative correlation between Ni and the PGE

represents the dominant trend observed in Figure 4.11. A second trend with a

positive correlation between Ni and PGE is superimposed on the first trend and is

characterized by highly elevated metal abundances (e.g. Pt >15 ppb versus Ni). The

PGE (excluding Ir) exhibit strong positive inter-element correlations, with Ir

exhibiting a similar trend as observed between Ni and Rh, Ru, Pt, Pd.

Figure 4.11. Inter-chalcophile element relationships for samples from the Long-Victor basal flow with S>0.25wt%.

ii. Sulfur-poor Sulfur-poor samples (S<0.25 wt%) exhibit the same strong mineral control trends as

the comprehensive data set (Fig. 4.9). The one exception is the lack of samples with

highly elevated chalcophile abundances no longer being present. Additionally, sulfur

does not correlate with MgO, and there appears to be no correlation between S and

any of the chalcophile elements (Ni, Cu, Co, Pt, Pd, Ir, Ru, Rh, Au), as shown with

Pt versus S (Fig. 4.12).

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Figure 4.12. Platinum (ppb) versus sulfur (S wt%), and sulfur (S wt%) versus MgO (wt%) for sulfur-poor (S<0.25 wt%) Long-Victor basal flow samples.

4.6. Discussion

Quantifying the physical and chemical size of Ni mineralized systems using

chalcophile element (PGE) ore forming signatures requires the understanding of two

system aspects: 1) flow field development and volcanology; and 2) the significance

and meaning of varying chalcophile element abundances. These two aspects are

used to determine the spatial correlation of chalcophile element values with the basal

flow of the Long-Victor system. The timing of komatiite crustal growth and its

relation to ore formation, and volcanological controls on the spatial distribution of

chalcophile element values, will be discussed for the purpose of developing PGE

vectors for targeting Ni sulfide mineralization.

a. Flow field Extensive work has examined the physical volcanology and theoretical development

of komatiite volcanic fields based on analogs to modern volcanic flow systems

(Arndt et al., 1977: Gresham and Loftus-Hills, 1981; Lesher, 1983, Barnes et al.,

1983; Hill et al., 1995; Lesher and Arndt, 1995; Williams et al., 1998; Moore et al.,

2000; Hill, 2001, Lesher et al., 2001; Beresford et al., 2002; Barnes, 2006; Barnes

and Lesher, 2008). A unified emplacement (flow field) model that incorporated both

field observations, geochemistry, and volcanology constraints from modern systems

was proposed by Lesher et al. (1984), and further refined by Lesher (1989) and

Arndt et al. (2008).

Within the komatiite flow field model, initial volcanic activity occurs as continuous

unconstrained lava eruption, resulting in the formation of a vent proximal sheet

flow, and the continuous changing of areas of channelized flow within the sheet

flow. Once the direction of the preferred lava flow is established (dependent upon

slope and pre-existing topography), sustained lava channel(s) will develop.

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Concurrent to channel development, the flanking environments develop through

progressive channel breakouts and inflationary advances, resulting in a complex

stratigraphy consisting of thin to thick differentiated and undifferentiated komatiite

flows.

A number of salient points in the komatiite flow field model and ore forming

process are not apparent in the above summary, but are relevant to understanding the

spatial correlation between Ni mineralization and ore forming signatures: (1) ore

formation occurs early in the flow field development, and consequently the genetic

relevance between channel-hosted mineralization and successive komatiite flows

decreases rapidly up stratigraphy (Lesher, 1989); (2) stratigraphic correlation

between channel and flank facies is only possible within the basal flow, and

successive stratigraphic flows in the flank environment have an unknown temporal

relationship to mineralization (e.g. large time gaps between successive flows

represented by accumulations of thick interflow sediments: Beresford et al., 2002);

(3) olivine cumulate rocks within the channels are the product of sustained magma

transport, and are dominantly unrelated to the mineralization event (e.g. magma

recharge of the system: Lesher and Arndt, 1995; Lesher et al., 2001); and (4)

quenched crusts and spinifex textured lithologies located stratigraphically above

channelized environments are transient, changing and developing with time, and as

such are probably coeval with ore generation (Hill et al., 1995; Beresford et al.,

2002; Barnes and Lesher, 2008).

The development of channelized flow is a critical step in the process leading to

economic Ni mineralization. Thus, the identification of channel facies, or the

proximity to channel facies is significant in terms of successful exploration.

Previous work has identified geochemical differences between channel and flank

facies within the flow field model (as summarized in Table 4.3). Geochemical

differences between facies are controlled by the sustained flow of slightly more

primitive magma in the channelized environments, resulting in the accumulation of

more primitive mineralogy (olivine ± chromite).

The spatially constrained geochemical data set from the Long-Victor system made it

possible to examine the geochemical gradients from mineralized channel facies to

unmineralized flank facies (Fig. 4.13). Three elements (MgO, Cr2O3 and Zn) were

identified as good discriminators between channel and flank facies (Lesher et al.,

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1984; Lesher, 1989; Lesher and Arndt 1995; Lesher et al., 2001) and are shown in

Figure 4.12. The select elements exhibit differences in median abundances between

the flank and channel. However, MgO, Cr2O3 and Zn exhibit weak positive or

negative chemical gradients with increasing distance from the channel. These trends

reflect increased fractionation, and contamination with increasing distance from the

channel. All other major elements associated with komatiites (Al2O3, TiO2) exhibit

flat patterns with a large spread in the data, and no systematic chemical gradients in

either the spinifex or B-zone textured samples.

Table 4.3. Comparison of geochemical and physical attributes of channel and flank facies. Compiled from Gresham and Loftus-Hills (1981); Lesher et al. (1984); Lesher (1989); Lesher and Arndt (1995); Lesher et al. (2001); Barnes (2006).

Attribute Channel Flank

Physical Setting Thickness > 30 up to 100 metres < 30 metres

Length/width < 200 m wide, extending kms Laterally extensive

Topography Hosted within a shallow depression

Flat setting

Lithology Olivine mesocumulate Meso to orthocumulates

Texture Cumulate dominant

(thin spinifex)

Spinifex (thickened) and B-zone differentiated

Sediments

Contact Devoid Present

Intercalated Devoid Common

Ni-mineralization Present Devoid

Geochemistry

Major > Fo olivine, lower Cr/Mg and Cr/Ni ratios

< Fo olivine, higher Cr, Ti, Al, Fe and Zn

Trace Uncontaminated Contaminated

Chalcophile Enriched and background Depleted and background

The abundance of MgO within spinifex textured samples displays a trend of

decreasing MgO with increasing distance from the channel, reflecting fractionation

away from the channel. Conversely, Cr2O3 abundances within spinifex textured

samples decrease with increasing distance from the channel. However, the B-zone

cumulates, exhibit constant MgO values and slightly increasing Cr2O3

concentrations. Zinc shows the largest contrast in median values between the

channel and flank in both B-zone cumulates and spinifex textures samples.

However, a significant spread in the data results in a low R2 value (Fig. 4.13).

Previous work by Lesher and Groves (1984) and Brand (1999) have identified

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anomalous Zn concentrations in distal environments interpreted to be related to

assimilation of Zn-rich sediments (e.g. exhalatives). La/Smpmn ratios were used as a

crustal contamination index on the assumption that LREE enrichment was due to

contamination. A weak trend of increasing crustal contamination with increasing

distance from mineralization is observed in the La/Smpmn ratios (Fig. 4.13).

However, the highest La/Smpmn values often occur within 150 m of mineralization.

Figure 4.13. Major and trace element abundances plotted as a function of distance from known mineralization (Ni >0.4%) which characterizes the channel (c.f. Fig. 4.3). Samples are classified as channel (Ch) and flank (Fl), as interpreted from constructed cross-sections. Samples are further subdivided based on texture: B-zone (Bz) and spinifex (Spfx). Median values for B-zones (solid line) and spinifex (dashed line) for channel and flank environments are shown. Calculated best fit lines for flank B-zones (blue) and spinifex (red) are shown, with R2 values for spinifex. Channel and flank subdivision at a distance of 100 m is based on data distribution.

135


The discrepancy between the spinifex and cumulate trends (Cr and MgO) are

interpreted to be the result of physical decoupling between the spinifex and

cumulates zones, a function of emplacement dynamics. Spinifex forms early in the

flank environment due to quenching and directional crystallization; whereas the

underlying B-zone cumulates are generated by sustained magma flow-through, and

are potentially unrelated to the overlying spinifex crust (Lesher, 1983; Hill et al.,

1995; Hill, 2001). Similarly, chalcophile element ore forming signatures can differ

within the spinifex and B-zones. One textural unit can be depleted, whereas the other

is characterized by background or enriched chalcophile element abundances.

Additionally, the progressive crystallization of spinifex can record chemical

variation in the magma flow-through with time.

b. Chalcophile element abundance Samples from the basal flow of the Long-Victor system exhibit a wide range of

chalcophile element (PGE) abundances. Chalcophile element depletion and

enrichment are observed in both the chalcophile element and the PGE/Tipmn versus

MgO graphs as they plot well above and below the general trends (Fig. 4.10). Strong

PGE depletion (Rh, Pt, Pd) is observed in 8 samples, whereas >20 samples exhibit

moderate to strong enrichment (Fig. 4.9; 4.10). To quantify the chalcophile element

enrichment or depletion signatures beyond strong, moderate and low, a background

value that characterizes the sample under non-ore forming conditions is required.

With the use of background values, chalcophile element ore forming signatures

represent the magnitude of the residual anomaly: positive and negative (enrichment

and depletion, respectively).

i. Background chalcophile element values Previous volcanology and geochemistry research on komatiite systems indicates that the systems are very dynamic and subject to substantial magma recharge and flow-through (Hill et al., 1995; Lesher and Arndt, 1995; Hill, 2001; Barnes et al., 2004; 2007). Consequently, large volumes of the komatiite system contain normal background chalcophile element abundances, due to the continuous supply of non-ore related magma and the flushing out of ore forming signatures from within the system. As such, the vast majority of samples collected from komatiites exhibit either enrichment or normal background chalcophile values (Lesher et al., 1981;

136


Lesher and Groves, 1984; Lesher et al., 2001; Barnes et al., 2004; 2007; Fiorentini et al., in press).

A methodology of iteratively filtering and removing enriched and depleted samples from the data set was used to generate equations from best fit lines that describe the abundance of chalcophile elements in a sample as a function of MgO content (Table 4.4; Fig. 4.13; Appendix D). Arguably, numerical fractionation and crystal accumulation models could generate similar trends and equations. However, utilizing real data sets to generate the equations bypasses the unresolved partition coefficients of the chalcophile elements into the crystallizing mineral phases (e.g. olivine, chromite, pyroxene).

Table 4.4. Equations derived and used to calculate background abundances of Ni, Pt, Pd, and Rh as a function of the MgO content of the sample. Complete list of equations provided in Appendix D. Ni Fn(MgO) = 90.04(MgO)-1175 R2 = 0.92

Pt Fn(MgO) = -0.369(MgO)+17.99 R2 = 0.77

Pd Fn(MgO) = -0.36(MgO)+18.0 R2 = 0.75

Rh Fn(MgO) = -0.0366(MgO)+2.1166 R2 = 0.57

One limitation of this methodology becomes apparent in Figure 4.13, as the

modeled trends curve at low chalcophile element abundances. This is an artifact of

increasing analytical uncertainties at low element abundances (e.g. Ni at low MgO;

Ti, Pt, Pd and Rh at high MgO). Consequently, the equations under-estimate the

abundance of the chalcophile elements at high or low MgO contents dependent upon

the incompatibility of the element. However, when the total sampling, preparation

and analytical errors are taken into consideration (± 500 ppm Ni, ± 2 ppb Pt and Pd,

and ± 1 ppb for Rh; Appendix C; Figure 4.14), the equations permit an accurate

estimate of background chalcophile element abundances that are expected within the

sample.

Calculated Ni and Ru (not shown) plot along straight lines with a slope determined

by the olivine and chromite partition coefficient for Ni and Ru. The incompatible

chalcophile elements (Pt, Pd, Rh, and Cu) plot as constants. Within this context,

deviations from background abundances are apparent with the PGE and less so with

Ni (Fig. 4.14). Given a sample population with 10 to 40 wt% MgO, it is possible to

accurately calculate a background chalcophile element budget for the sample that

would represent the sulfide-free crystallization of the sample (e.g. contains no ore

forming signature either depleted or enriched).

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Figure 4.14. A. Ni/Tipmn versus MgO for the Long-Victor system, basal flow samples shown in red diamonds. Calculated Ni normalized to actual Tipmn plotted as black triangles. Ni/Ti trend line based on a derived equation. B. Pt/Tipmn versus MgO for Long-Victor, basal flow samples shown in red diamonds. Calculated Pt normalized to actual Tipmn plotted as black triangles. Trend line of Pt/Ti represents perfectly incompatible elements at a determined constant ratio of 0.67.

The derived equations (Table 4.4, Appendix D) also permit the calculation of

chalcophile element abundances at common MgO contents for direct comparison

with other komatiite systems. Assuming a parental liquid MgO content of 24 wt%,

the liquid background chalcophile element content for the Long-Victor system is:

Ni: 1001 ppm, Cu: 48 ppm, Pt: 9.7 ppb, Pd: 9.3 ppb, Ru: 3.7 ppb, Rh: 1.3 ppb, Ir:

1.1 ppb. These values are similar to those reported by Fiorentini et al. (in press) in a

global komatiite comparison and previous work from the Kambalda Dome (Keays et

al., 1981; Keays, 1982).

ii. Chalcophile element enrichment Chalcophile element enrichment (> background values: mineralization) is the result

of sulfur saturation within the system and accumulation of immiscible sulfides

within the sample. Chalcophile element enrichment in the Long-Victor system is

defined as Pt/Ti pmn values greater than 0.88 and Pd/Ti pmn values greater than 1.65.

These values are derived from the background median value of each ratio, with a ±2

ppb error added to accommodate for sample heterogeneity and analytical

uncertainty.

For the purpose of discussion, chalcophile element enrichment has been divided into

two groups based on sulfur content: sulfide-bearing (S>0.25 wt%: Fig. 4.15), and

sulfide-poor (S<0.25 wt%: Fig. 4.15). Sulfide-bearing samples (S>0.25 wt%) are

enriched relative to the calculated background, and commonly contain fine

disseminated sulfides. Sulfide-poor samples contain less than 0.25 wt% S, yet have

chalcophile element abundances that higher than the calculated background

abundance for the sample.

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Figure 4.15. Plots of PGE/Tipmn versus MgO (wt%) for Long-Victor samples exhibiting chalcophile element enrichment based on Pt and Pd abundances. Samples are plotted as analytical data in red and calculated chalcophile element abundance in grey (Table 4.4). Samples with sulfur greater than 0.25 wt% are shown on the left hand side and samples with sulfur less than 0.25 wt% on the right hand side. Blue lines define the analytical uncertainly field around the numerically modelled background values (see Appendix C).

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Sulfur-bearing samples

Sulfur-bearing samples that exhibit chalcophile element enrichment (Pt/Ti pmn >

0.88, Pd/Ti pmn > 1.65 and S>0.25 wt%) have a range of sulfide content from 0.26 to

4 wt%. Texturally, these samples consist of both spinifex textured and B-zone

cumulates, and are dominantly from within the channel environment. Chalcophile

element enriched samples are also identified from within the flank environment. At

high sulfur contents, the sulfur bearing samples exhibit a strong correlation between

chalcophile element content and sulfur abundance, whereas at lower sulfur

abundances (S<1 wt%) the correlation becomes less apparent (Fig. 4.11; 4.15).

Considerable scatter is observed in the data set at very low sulfur contents.

Primitive mantle-normalized noble metal plots exhibit element profiles that are

greater than the mantle, and display a convex-up shaped pattern, characteristic of

metal accumulation. A small negative Pt anomaly is observed in most samples. All

samples exhibit positive Pt, Pd, Rh element enrichment when compared to

calculated values based on the MgO content. Two of the 18 samples exhibit Ru

negative depletion values, and Ir exhibits depletion in half of the samples relative to

the calculated values.

The sulfur-bearing chalcophile element enriched samples from the Long-Victor

system are interpreted to represent orthomagmatic mineralization. Four lines of

evidence support this interpretation: 1) the presence of disseminated sulfides, 2)

good correlation between S content and chalcophile element abundance, 3)

occurrence of samples exhibiting enriched primitive mantle normalized abundances,

and 4) the presence of samples showing PGE/Tipmn enrichment relative to

background values (Fig. 4.14).

Sulfur-poor samples

Sulfur-poor chalcophile element enriched samples (Pt/Ti pmn > 0.88, Pd/Ti pmn > 1.65

and S<0.25 wt%) make up a minor portion of the data set, constituting only 13

samples out of 133 from the basal flow. These enriched samples include B-zone

cumulates from both the channel and flank environments. The enriched samples do

not exhibit correlations between PGE abundance and fractionation indexes (MgO,

TiO2, Al2O3), or visible correlation between the chalcophile elements and S (Fig.

4.16). Positive inter-element correlations are observed between the PGE, but, not

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with Ni. The one exception to this is Ir, which exhibits a positive correlation with

Ni. When compared with calculated background values, all sulfur-poor chalcophile

enriched samples exhibit positive Pt, Pd and Rh enrichment. However, Ru is

depleted in 5 samples, Ir is depleted in 7, and Ni depleted in 6 of the 13 samples,

relative to the calculated values.

Figure 4.16. Plots of Pt correlations to incompatible elements (TiO2 and S) and chalcophile elements (Pd, Ni) for the Long-Victor basal flow samples with low sulfide abundance (< 0.25 wt%) and a chalcophile element enrichment signature.

The elevated PGE values (chalcophile element enrichment) could be an artifact of an

orthomagmatic mineralization signature that has undergone sulfur loss. However,

this enrichment may also be unrelated to Ni mineralization and a function of

alteration, analytical error, or a primary feature of magmas anomalously rich in the

PGE. Sulfur-loss due to metamorphism, alteration and oxidation is the most direct

explanation of enrichment, as S-mobility is well documented in other orthomagmatic

settings (Seccombe et al., 1981; Stone et al., 2004b). A sample originally containing

disseminated orthomagmatic mineralization that undergoes S-loss would retain the

elevated chalcophile element signature of a mineralized sample, but contain a low-

sulfur content.

Alteration and element mobility are arguably able to generate an enriched PGE

signature. The removal of MgO or TiO2 from the system would result in the over

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estimation of the PGE (mineralization signature), whereas the addition of MgO or

TiO2 would result in underestimation (depletion signature). Alteration processes

could also liberate the PGE from within the silicate or sulfide mineralogy and

redistribute them within the system, leading to complementary alteration induced

enrichment and depletion signatures. However, if the samples are plotted as

calculated PGE normalized to TiO2, all the data points fall within the background

field. This would not be anticipated if TiO2 or MgO were added or removed from

the sample. Additionally, Al2O3 and TiO2 exhibit a strong positive correlation (Fig.

4.6) and alteration does not seem to play a crucial role in the in the redistribution of

these elements.

The possible mobility of the PGE within the system is difficult to constrain. The

current S-poor PGE enriched samples have mantle normalized noble metal patterns

similar to those of the disseminated mineralization. Additionally the samples do not

exhibit significant PGE fractionation if plotted on PGE binary element plots (Fig.

4.16). Consequently, the S-poor chalcophile element enriched samples are

interpreted to represent orthomagmatic mineralization that has undergone S-loss.

Enriched chalcophile element signatures are readily identifiable on the basis of

deviations in PGE/Tipmn values from calculated background values (Fig. 4.15).

Enrichment is associated with both S-bearing and S-poor samples highlighting the

limited functionality of S as a filter to identify mineralization, and the possibility of

enriched mineralization signatures present in samples not appearing to be

mineralized (e.g. trace sulfide, S<0.25 wt %).

iii. Chalcophile element depletion Chalcophile element depletion in a sample occurs as a result of sulfur saturation

within the system and the removal of a sulfide liquid from the silicate liquid. The

sulfur saturation and segregation process leads to the removal of the chalcophile

elements from the silicate magma prior to crystallization. To quantify depletion,

analytical values are compared with calculated values and presented as a ppm or ppb

deviation from a background value. Uncertainties were incorporated to

accommodate sample heterogeneity and analytical precision (Pt and Pd ±2 ppb, Rh

±0.5 ppb, 500 ppm for Ni: Fig. 4.17). Chalcophile element depletion is apparent in

the Pt/Tipmn and Pd/Tipmn ratios for 18 samples (11 with S<0.25 wt% and 7 with

S>0.25 wt%: Fig. 4.16).

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Figure 4.17. Ni/Tipmn versus MgO (wt%) and Pt/Tipmn versus MgO (wt%) for Long-Victor basal flow samples, filtered to remove enrichment signature (Pt/Ti pmn <0.88 and Pd/Ti pmn < 1.65). Samples are plotted as analytical data in red and calculated chalcophile element abundance in grey with lines delineating ± 500 ppm uncertainty for Ni, and ±2 ppb uncertainty for Pt.

Chalcophile element depletion signatures for Pt, Pd, and Rh are shown in Figure

4.18 as a calculated deviation from the background valued (ΔPGE). Platinum, Pd,

and Rh exhibit strong positive inter-element correlations (ΔPt versus ΔPd: R2 =

0.94), as shown in Figure 4.18B and C. This relationship is maintained even when

the ±2 ppb uncertainty is removed, indicating the uncertainly could be lower. The

other chalcophile elements (Ni, Cu, Ru, Ir) exhibit variable correlations relative to Pt

and Pd depletion. Ruthenium appears to correlate moderately with Pt depletion.

Nickel and iridium depletion (not shown) exhibit a scatter of data, rather than the

expected positive correlation with Pt enrichment or depletion (Fig. 4.18D).

Figure 4.18. Change in chalcophile element abundance from calculated background values (Δ) for sample from Long-Victor basal flow. Samples exhibiting enrichment signatures are removed. A. Calculated Pt (ppb) depletion, with modeled depletion lines of 100%, 75%, 50% and 0% shown. Dark grey shading delineates fields of uncertainty. B. Calculated Pt depletion versus Pd depletion with ±2 ppb uncertainty applied to both. C. Calculated Rh depletion versus Pt depletion with uncertainty shown by grey bars. D. Nickel depletion versus Pt depletion with uncertainty shown by grey bar.

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Chalcophile element depletion is best identified using Pt, Pd and Rh, which exhibit

the strongest incompatibility with olivine, as observed by strong negative

correlations with MgO (Fig. 4.9). This strong incompatibility results in constant or

near constant PGE/Tipmn values (Fig. 4.10). Platinum and Pd occur at the highest

relative abundances that even with conservative uncertainties of ±2 ppb, the

quantified ore forming signatures are meaningful. Conversely, Ru and Ir are similar

to Ni, and exhibit positive correlations with MgO (Fig. 4.9). The positive correlation

with MgO is interpreted to be independently controlled by IPGE rich liquidus alloy

phases, resulting in changing PGE/Tipmn values with MgO content (Fiorentini et al.,

2008; Barnes and Fiorentini, 2008; Locmelis et al., 2009). This characteristic along

with the lower total abundance of Ir and Ru complicates the interpretation of ore

forming signatures.

Nickel appears relatively insensitive to sulfur saturation, relative to the PGE in

komatiite systems. Despite being a chalcophile element, Ni has experimentally

determined partition coefficients for the sulfide phase that range from 300 to 1000,

which are orders of magnitude smaller than the partition coefficients measured for

the PGE (>10000: Fleet and MacRae, 1983; Peach et al., 19990; Stone et al., 1990;

Fleet et al., 1991; Fleet et al., 1996; Barnes and Maier, 1999). Consequently, the

PGE are more susceptible to extraction from the silicate melt once sulfur saturation

is attained. This is reflected in the data set, where Pt, Pd and Rh exhibit strong

depletion, yet there is no apparent correlation with the Ni abundance.

c. Spatial correlation of chalcophile element values Understanding the spatial correlation between the chalcophile element (PGE) ore

forming signatures and Ni mineralization is essential to the development of a

working mineralization vectoring tool. Ore forming signatures in the form of strong

chalcophile element depletion were predicted to be hosted within komatiites of the

Silver Lake Member of the Kambalda Dome (Keays et al. 1981; Keays, 1982), and

were documented in select samples from the flanking environment (Lesher et al.,

1981; Lesher and Groves, 1984; Lesher et al., 2001; Lesher and Keays, 2002).

Within the current data set, strongly depleted samples contribute only 10% of the

data, which is considerably lower than anticipated. Eighteen samples from 9

different drill holes, out of a total of 133 samples from the basal flow in the Long-

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Victor area, exhibit quantifiable depletion on the basis of the above calculations and

methodologies (Appendix D). These 18 samples consist of both channel (5 samples)

and flank environments (13 samples). Within the flank environment, both spinifex

(11 samples) and B-zone cumulates (2 samples) exhibit depletion. In the channel

environment, depletion is only observed in one drill core in the spinifex textured

flow top, and none from the B-zone cumulates. Enriched samples within the data set

comprise 52 samples and represented all facies of the komatiite system. The

remaining 63 samples within the 133 sample data set display background values

which are found throughout the system. Figure 4.19 shows the relationship between

known channel hosted mineralization and the modelled intensity of enrichment,

depletion and background signatures as projected to the interpreted basal flow top.

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Figure 4.19. Leapfrog 3D-model of chalcophile element (PGE) mineralization signatures within the basal flow of the Long-Victor channels. A. Lunnon Basalt surface with 0.4% Ni grade shell shown. B. Modeled surface of the basal flow spinifex with colour gradients representing ore forming signatures observed in the spinifex; green = background, blue = depletion, and red = enrichment. C. Mineralization signatures observed in the B-zone cumulate, projected to the modeled surface of the basal flow spinifex.

The Long-Victor exploration and development drilling database was utilized to

assess the spatial correlation between the analyzed samples and known

mineralization. Distances and orientation were calculated from a sample to known

mineralization greater than 0.4% Ni (as shown in Fig. 4.4 and Fig. 4.20). The

resulting distances exhibit a range from 1.5 m to 509 m with Ni grades ranging from

0.43 to 6.5% Ni. Two trends are apparent in the distance versus Ni grade data set

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(Fig. 4.20). The first trend of increasing Ni grade with proximity to mineralization is

associated with high-grade (>2% Ni) mineralization. The second trend observed in

the data set is a constant Ni grade of 0.4% (cutoff grade) at all distances from

mineralization, with no apparent spatial correlation between grade and distance.

Figure 4.20. Plot of distance (m) versus Ni grade (%) for all samples from the basal flow of the Long-Victor system. Distances are an average of the three closest Ni occurrences to each sample. Ni grade (%) represents the average Ni abundance for those three occurrences.

In order to display the effects of both strong enrichment and strong depletion

mineralization signatures are plotted as log scaled Ti normalized values (Fig. 4.21).

The data are also placed into four classes based on the methodologies outlined in the

previous sections: (1) background values, (2) chalcophile element depleted, (3)

sulfide-bearing chalcophile element enriched, and (4) sulfide-poor chalcophile

element enriched. The resulting plot indicates background values are found

throughout the system at all distances from known mineralization (Fig. 4.21A, B).

However, three areas (A, B, C: highlighted in Fig. 4.21A, B) are identified where

there appears to be a distance control on the chalcophile element abundance in the

samples.

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Figure 4.21. A. Pt/Ti pmn and B. Pd/Ti pmn versus distance (m) to nickel mineralization. Samples are classified as Enriched (mineralized based on Pt/Ti and Pd/Ti ratios), Enriched LS (mineralized based on Pt/Ti and Pd/Ti ratios with S<0.25%), Depleted (chalcophile element depleted samples as determined from previous section) and Background (samples which exhibit no indication of chalcophile element enrichment or depletion). Plots are domained into three spatial regions A, B, and C based on predominant ore forming signatures at the respective distances.

Area A (Fig. 4.21A, B), is dominated by samples that contain enriched sulfide-

bearing and sulfide-poor ore forming signatures. The main grouping of data occurs

within 80 m of known mineralization. Additional enriched samples are observed at a

distance of 200 m with three strongly enriched samples occurring at 120 to 180 m.

Within Area A, the main grouping of data within 80 m of mineralization is

characterized by a relatively constant enriched signature over the interval from 30 to

80 m (as shown with Pt/Tipmn and Pd/Tipmn in Fig. 4.22). From 0 to 30 m, the

enriched signature exhibits a general trend of increasing PGE abundance with

proximity to mineralization, which is interpreted to be a mineralization halo. This

trend is observed in both sulfur-bearing and sulfide-poor enriched samples,

supporting the S-loss model for the generation of sulfide-poor chalcophile element

enriched samples.

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Figure 4.22. Pt/Tipmn and Pd/Tipmn versus distance (m) to Ni mineralization, focusing on samples within 80 m of known mineralization. Enriched (mineralized based on Pt/Ti and Pd/Ti ratios), Enriched LS (mineralized based on Pt/Ti and Pd/Ti ratios with S<0.25%), Depleted (chalcophile element depleted samples as determined from previous section) and Background (samples which exhibit no indication of chalcophile element enrichment or depletion).

The three sulfur-bearing samples that exhibit chalcophile element enrichment at a

distance >100 m from mineralization (Area A, Fig. 4.21) are from three different

drill holes: one spinifex textured and two B-zone cumulate samples. Although from

three different holes, the samples are all located in a similar proximal position to the

channel, which is best classified as a transition zone from the channel environment

to flank.

Area B (Fig. 4.21) is characterized by sulfur-bearing and sulfur-poor enriched

samples. Overall, samples from this area appear similar in Pt and Pd abundance to

those in Area A, but occur at a distance of 300 to 450 m away from known

mineralization. Similarly, these samples are interpreted to represent orthomagmatic

mineralization that has undergone variable S-loss. The origin of the enriched

samples in Area B is due to two possible events. The first is halo mineralization

proximal to more substantial mineralization, akin to that observed in Area A. The

second interpretation is that chalcophile element enrichment is the result of localized

sulfur saturation leading to the development of patchy orthomagmatic disseminated

sulfide in the flanking environment. To date, there is insufficient exploration drilling

within the area to sample and definitively resolve the two hypotheses. However, as

discussed later in this section, the distance of 300 to 450 m correlates to a potential

additional channel, as identified by thickened olivine cumulates and a sediment free

basal contact. The presence of an additional channel supports the idea that

enrichment in Area B is a halo to more substantial mineralization.

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Two working hypotheses are inferred for the creation of chalcophile element

enriched halos. The first is a cloud of disseminated sulfide, which is a primary

feature of orthomagmatic mineralization that has undergone partial sulfur loss. The

second model infers the mobilization of chalcophile elements out of the primary

mineralization and into the surrounding host rock during post-crystallization

chemical dispersion. There is little evidence within the current samples and data set

to support one model over the other. The mix of sulfur-poor and sulfur-bearing

chalcophile element enriched samples that occur within the same distance interval

and do not exhibit contrasting PGE fractionation trends, supports the hypothesis that

enrichment is a single magmatic process. Alternatively, PGE fractionation would be

anticipated if the process involved secondary hydromagmatic mobilization. An

orthomagmatic mineralization event in the form of a disseminated cloud of

mineralization, that has undergone variable sulfur mobility and loss, is the preferred

explanation of the enriched signature in Area B.

Area C (Fig. 4.21) is dominated by chalcophile element depleted samples. Two

depleted samples occur closer to mineralization (45 and 165 m) and are located on

the margin of the Victor channel on the down-dip side. The remaining depleted

samples occur on the up-dip flank of the Victor channel and at an average distance

of 340 m from known mineralization. A general trend of increasing magnitude of

depletion with increasing distance from mineralization is observed with Pt, Pd and

Rh in the data set (Fig. 4.23). The depleted samples also show increased depletion in

the more fractionated samples, which spatially parallels the observation of a weak

linear relationship between the MgO content of the spinifex and distance from

mineralization.

Figure 4.23. Pt, Pd and Rh for each chalcophile element depleted sample from Area C shown as % depletion versus distance from mineralization ≥ 0.4% Ni.

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Chalcophile element depletion, specifically in the PGE, has been numerically

modeled by fractional segregation and batch equilibrium segregation (Barnes et al.

1988; Barnes et al., 1995; Lesher and Stone, 1996; Fiorentini et al., 2010). Both

numerical models can be applied to orthomagmatic systems and identify the

extraction of chalcophile elements from the silicate magma by the sulfide phase.

However, the rate at which this occurs differs between the two models. Fractional

segregation modeling indicates a rapid and complete extraction (Lesher and Stone,

1996); whereas batch equilibrium shows a slower rate of depletion, which occurs

once sulfur saturation is attained.

Within the Long-Victor system, the majority of chalcophile element depleted

samples exhibit relatively small depletions (Fig. 4.18A). The small depletions are

the product of either: (1) the ore forming system being dominated by batch

equilibrium segregation, or (2) the result of fractional segregation followed by

mixing and dilution of the strongly PGE depleted magma (responsible for the Ni

mineralization) with a recharging magma that was not sulfur-saturated or PGE

depleted. Both numerical models have been applied to natural systems and produce

ambiguous results. A hybrid equilibrium-fractional segregation model is suggested

for dynamic channel settings (Fiorentini et al., 2010). Within this dynamic

environment, chalcophile element depleted magmas are progressively replaced with

undepleted recharging magma resulting in the dilution of the mineralization

signature.

d. Timing of komatiite spinifex growth and relation to ore formation

The progressive replacement of chalcophile element (PGE) depleted magma with

undepleted magmas is observed within the flanking environment of the Long-Victor

system. Drill hole KD6024 (Fig. 4.24) intersected a 17 m thick flow with well-

developed A-zone spinifex and B-zone cumulates, with sediments on the basal flow

contact. Seven samples have been collected from the 17 m thick flow, where 4

samples are from the upper spinifex textured portion, and 3 are from the underlying

cumulates. The spinifex samples exhibit a trend of decreasing PGE depletion down-

hole over 1.2 m. Conversely the B-zone cumulates exhibit enrichment in the middle

of the flow, and background values at the basal contact.

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The decreasing depletion signature with depth in the spinifex is argued to represent a

progressive change in the magma composition with time. Faure et al. (2006)

concluded that a constant temperature gradient is the critical component in the

progressive downward growth of oriented spinifex into a komatiite flow. As the

magma composition changes due to progressive removal of the depleted magma

with time, the continuous growth of orientated spinifex records and preserves this

change in the chalcophile element abundance.

Figure 4.24. Chalcophile element depletion (left) and enrichment (right) as a percentage change from the calculated background for each chalcophile element of the basal flow, from drill hole KD6024. No chalcophile element uncertainty was applied to the interpreted mineralization signatures. Samples 195.7, 196.0, 196.9, and 209 m from Lesher and Arndt (1995) and Lesher et al. (2001).

The observed chalcophile element trends in KD6024 are markedly different than

those observed in other closed non-mineralized komatiite systems consisting of thick

and thin komatiite flows and lava lakes, as identified by Keays et al. (1981); Keays,

(1982); Dowling and Hill, (1992); Zhou, (1994); Puchtel and Humayun, (2001). In

these non-mineralized systems, (Vetreny Belt, Mt. Clifford, Belingwe), the

chalcophile element profiles through flows are homogenous and exhibit relatively

constant Pt/Tipmn and Pd/Tipmn ratios. Conversely, mineralized komatiite systems

(Mt. Keith and Lunnon Shoot) exhibit variability in Pt/Tipmn and Pd/Tipmn through

the flow units associated with mineralization (e.g. basal flow of Silver Lake

Member), and can be utilized as a Ni prospectivity indicator. The observed

chalcophile element profile (Fig. 4.24) indicates that the sulfur saturation event

leading to ore formation occurred prior to an extensive flow-through event in the

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flanking environment, in addition to the flow-through already identified within the

channel (Lesher and Arndt, 1995).

e. Volcanological control on spatial distribution of chalcophile element values

This section discusses the relation between the spatial distribution of chalcophile

element mineralization signatures and the dynamic model for komatiite

emplacement, taking into account primary syn-volcanic fault controlled topography.

This relationship is displayed in a schematic cross-section through both the Long

and Victor channels and adjacent flanking environments (Fig. 4.25). Three flank

environments are identified based on the current structural setting: up-dip flank

(closest to the surface), inter-channel flank (occurring between the Victor and Long

channels), and the down dip flank (below the Long channel: Fig. 4.25). The

interpreted cross-section through the channels and flanks is also divided into

enriched, depleted and recharge zones (e.g. background) based on PGE signature

and volcanology.

The enriched zone is found in two areas. The first area is proximal to mineralization

within the lower portion of the channel, as anticipated in a Ni mineralized system

(Fig. 4.25). The enriched signatures are preserved in the lower portion of the

channel as finely disseminated and possibly sulfide-poor mineralization due to

alteration and S-loss. This ore forming signature became isolated from the active

magmatic system by progressive channel infill, a result of continuous olivine

crystallization from overlying undepleted magma recharging the system (Fig. 4.26).

Figure 4.26 represents a time sequence block model for the progressive

emplacement, mineralization and preservation of chalcophile element (PGE) ore

forming signatures in the Long-Victor system.

The second area where an enriched signature is prevalent, but not inclusive, is the

up-dip flank in the B-zone cumulates (Fig. 4.25). Intuitively, an enriched signature

would not be expected in the flanking environment, as the flanks are interpreted to

be barren. On the assumption that flank environments are barren, the enriched

signatures in both channel and flank environments are difficult to interpret, unless

the enrichment occurring within the up-dip flank of the Victor channel is related to

the development of an additional subsidiary channel. Evidence supporting the

presence of an additional “channel-like” environment exists in an adjacent drill hole

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(KD6012). Drill hole KD6012 exhibits a thickened basal flow unit and no basal

sediments, akin to the majority of ore shoots (excluding the presence of massive

sulfide) identified within the Kambalda Dome. Consequently, the presence of

enriched mineralization signatures in the up-dip flank area indicates potential for Ni

mineralization down plunge of the developing channel.

Figure 4.25. Schematic cross-section through interpreted paleo-volcanic setting of Victor and Long channels showing relative locations of flank environments. Chalcophile element enrichment zones shown in red dots, chalcophile element depletion shown in blue shading and areas of recharge (background) in grey.

The depleted zone (blue shading in Fig. 4.25) occurs in the flank and is preserved

within the spinifex flow top, approximately 340 m from channel mineralization. The

spinifex zone records both complete chalcophile element depletion, and a trend of

decreasing signature magnitude with time (down spinifex: Fig. 4.24). This

diminishing trend is the result of magma recharge and dilution of the chalcophile

element depletion signature with time (Fig. 4.26).

The depletion preserved in the up-dip flank spinifex can arguably be related to two

mineralization events: (1) main channel, and (2) flank. The main channel

mineralization event hosted within the Victor channel is the preferred correlation, as

it is the only significant ore body identified in a proximal location. However, the

development of a “channel-like” environment with chalcophile element enrichment

(KD6024: Fig. 4.24) within the same area as depletion (Fig. 4.21), implies the

enrichement could be a very local (10s to <100 m) correlation between

mineralization and chalcophile element depletion.

Two lines of evidence support a more distal (Victor channel) rather than a close

correlation between mineralization and chalcophile element depletion. The first is

the lack of meaningful sulfide accumulation yet identified in the vicinity of the drill

hole KD6024. The second indicator for a distal correlation is the occurrence of a

chalcophile element depleted sample in the spinifex portion of the second flow, at

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the same distance from the channel. This would indicate that the depletion is

unrelated to the “channel-like” environment located stratigraphically below. The

occurrence of multiple samples exhibiting depletion at similar distances from the

channel supports the premise that a distance of 340 m from the channel is critical.

Additionally, the depleted sample within the flank, stratigraphically above the basal

flow, that preserves magma recharge in the form of PGE depletion dilution (Fig.

4.24) supports an early prolonged sulfur saturation event spanning successive

recharge and flushing events during the emplacement of the Silver Lake member.

Recharge zones (grey shading in Fig. 4.25) are characterized by dominant olivine

cumulate lithologies and samples that exhibit background chalcophile element

contents. Both these characteristics support sustained magma flow-through.

Recharge areas are identified both above and beside the mineralized zone. The

observed spatial pattern is the result of early ore formation within the channel,

followed by magma recharge above the accumulated sulfides within the channel.

Recharge is not restricted to the channel, but also extends into the flanks. Flank

recharge is observed at a distance greater than 200 m from the channel, as preserved

within spinifex as a decreasing chalcophile element depletion trend at a fixed

distance from the channel (Fig. 4.24).

Additionally, if all the spinifex formed at the same time, spinifex from basal flow on

the up-dip flank should systematically record the same ore forming signature.

However, this is not observed in the flank spinifex. At a distance of 380 m from the

channel, the spinifex exhibits strong chalcophile element depletion; whereas closer

to the channel, no depletion signature is recorded within the spinifex. The lack of a

depletion signature closer to the channel may support the idea of transient crusts

(Hill et al., 1995: Fig. 4.26). The lava pulse (chalcophile element depleted) that

formed the up-dip basal flow at the time of emplacement instantly formed a

quenched surface. Subsequently, spinifex began to grow, thus preserving the

depletion signature. However, closer to the channel, the flow velocity is higher and

the initial crust may have been continuously formed and destroyed (transient crust),

until the flow velocity decreased enough for a flow top to be preserved.

Consequently, the spinifex above the channel records a lava composition which was

no longer chalcophile element depleted.

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Figure 4.26. Time sequence block model for the progressive emplacement, mineralization and preservation of chalcophile element ore forming signatures. Komatiite flows colour coded for chalcophile signature: green = background, blue = depleted, red = enriched.

4.7. Conclusion

The Long-Victor system is dominated by normal background chalcophile element

abundances, typical of those found in Munro-type komatiites <3.0 Ga worldwide.

The presence of normal background chalcophile element abundances allows for the

identification of positive (enrichment) and negative (depletion) deviations from this

background. Enrichment signatures are evident with all the chalcophile elements

(Ni, Pt, Pd, Ir, Ru, Rh). However, depletion is most evident with the PGE,

specifically Pt and Pd, and is not at all discernable utilizing Ni.

Within the Long-Victor system, enrichment is considerably more common than

depletion. This work was initially carried out to identify chalcophile element

depletion, and targeted low-sulfide samples (visibly S-free). Yet, the resulting data

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set (133 samples) is dominated by samples that exhibit enriched signatures (39%),

whereas, depletion constitutes only 14% of the samples.

Depletion signatures identified within the Long-Victor system are restricted to

spinifex textured samples in the flanking environment within the basal flow, and at

least the overlying second flow unit. Based on the current sample density in the

Victor area, depletion is only recognized within the up-dip flank, a function of the

local volcanological setting and flow dynamics. The strongest depletion signatures

are preserved in the uppermost portions of the spinifex zone and decrease in

magnitude with increasing depth from the top of the spinifex, a result of progressive

flushing by recharging lava.

Depletion signatures are preserved at an average distance of 340 m from channel

hosted mineralization, and systematically decrease in the magnitude of depletion

closer to the channel: the result of recharging undepleted magma. Extensive

sampling has not been carried out beyond the depletion zone (340-400 m from the

channel). However, limited sampling completed beyond this distance would indicate

an enriched mineralization signature is once again present, suggesting another

channelized environment up-dip.

Enrichment signatures are largely restricted to cumulate lithologies, with the

distribution falling into two groups, proximal to known mineralization and distal to

known mineralization. Proximal enrichment signatures occur less than 80 m from

mineralization. These signatures have a general trend of increasing enrichment

magnitude with proximity, forming a halo around the mineralization which extends

at least twice as far as the detectable extent of anomalous Ni concentrations. This

trend is observed in both sulfide-bearing samples (S>0.25%) and sulfide-poor

samples (S<0.25%), with the latter representing sulfur loss. Development of the

enriched halo around mineralization is suspected to result from primary fine

disseminated mineralization that developed at the same time as the more massive

accumulations. Distal enrichment represents sulfur saturation occurring away from

the known mineralization. This enrichment is interpreted as either localized sulfur

saturation forming disseminated sulfides, or a halo proximal to more substantial

mineralization not yet identified.

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Although, this study was carried out in a well-constrained komatiite system, the

results and conclusions are applicable to all levels of komatiite-hosted Ni sulfide

exploration (brownfields to greenfields), as key volcanological areas are identified

that preserve mineralization signatures. In summary, within the basal flow, the flank

environment hosts PGE depletion in the uppermost spinifex zone, whereas,

chalcophile element enrichment is observed in the lower to middle portions of the B-

zone cumulates, if prospective. Thickened mineralized channelized environments

host chalcophile element enrichment in the lower portions of the B-zone cumulates.

Within both environments, S is a poor indicator of the mineralization processes.

Enriched samples may have no visible sulfides, whereas depleted and background

samples may contain abundant secondary S, unrelated to a mineralization process.

This study provides an outline for the use of PGE-based vectors in exploration for

massive Ni sulfides in komatiite systems. Both enrichment and depletion signatures

exhibit a recognizable spatial correlation to mineralization, providing viable vectors

to Ni sulfide mineralization. This study is the first detailed work relating the spatial

distribution of quantifiable mineralization signatures in the form of chalcophile

element enrichment and PGE depletion to the dynamic physical volcanology of an

ore forming environment.

158


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Williams, D.A., Kerr, R.C., Lesher, C.M., Barnes, S.J., 2001. Analytical/numerical modeling of komatiite lava emplacement and thermal erosion at Perseverance, Western Australia: Journal of Volcanology and Geothermal Research, v. 110, p. 27-55.

Wolfgram, P., Golden, H., 2001. Airborne EM applied to sulfide nickel: Examples and Analysis: Exploration Geophysics, v. 32, p. 136-140.

Zhou, M. F., 1994. PGE Distribution in 2.7-Ga layered komatiite flows from the Belingwe greenstone belt, Zimbabwe: Chemical Geology, v. 118, p. 155-172.

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Contents 4.1. Introduction ................................................................................................... 110 4.2. Kambalda Dome ............................................................................................ 113

a. Geological setting .................................................................................... 113

b. Structural modification ............................................................................ 116

4.3. Chalcophile Element Abundance .................................................................. 117 4.4. Materials and Methods .................................................................................. 118

a. Sample selection ...................................................................................... 118

b. Distance to mineralization ....................................................................... 120

c. Analytical techniques .............................................................................. 121

4.5. Results ........................................................................................................... 122 a. Major and trace element geochemistry .................................................... 123

b. Chalcophile element geochemistry .......................................................... 127

i. Sulfur-bearing ............................................................................................ 130

ii. Sulfur-poor ................................................................................................ 131

4.6. Discussion ...................................................................................................... 132 a. Flow field ................................................................................................. 132

b. Chalcophile element abundance .............................................................. 136

i. Background chalcophile element values ................................................... 136

ii. Chalcophile element enrichment ............................................................... 138

iii. Chalcophile element depletion .................................................................. 142

c. Spatial correlation of chalcophile element values ................................... 144

d. Timing of komatiite spinifex growth and relation to ore formation ........ 151

e. Volcanological control on spatial distribution of chalcophile element

values 153

4.7. Conclusion ..................................................................................................... 156 4.8. References ..................................................................................................... 159

List of Figures Figure 4.1. Generalized geological map of the Kambalda Dome with nickel sulfide ore shoots shown

in plan projection with major faults and fold axis shown. Area of the Long-Victor Ni deposit shown by dashed outline. Modified after Ross and Hopkins (1975) and Stone et al. (2005). . 114

Figure 4.2. Local Kambalda Dome mine stratigraphy in an idealized cross-section showing the Lunnon Basalt Formation (footwall), and Kambalda Komatiite Formation comprising the Silver Lake and Tripod Hill Members. The Silver Lake Member exhibits thickened channel facies, thin flank facies, interflow metasedimentary rocks and Ni sulfide mineralization within a trough feature. Modified from Lesher and Groves (1984). ................................................................. 116

Figure 4.3. 3D model of the Lunnon Basalt surface (shown in green) and 0.4% Ni grade shell (shown in red) as modeled with Leapfrog®. Victor trough and Long trough interpretations shown with dashed lines, with select ore shoots labeled (Gibb, Victor, McCleay, Long and Moran). Grey shading delineates approximate flank facies distribution. View looking west. ........................ 119

Figure 4.4. Plot of distance (m) and azimuth of samples from nickel mineralization > 0.4 wt% Ni. Each data point is an average of the closest three distances and azimuths. Rose diagram

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showing distribution of azimuths with general trend (335°) of the Long-Victor channels shown by grey arrow, as observed in Figure 4.1. ............................................................................... 121

Figure 4.5. Plot of FeOtot versus MgO wt% for the basal flow within the Long-Victor area. Oxides as wt% and recalculated to anhydrous. Samples are characterized as channel facies B-zone (Ch Bz) and spinifex (Ch Spfx) and flank facies B-zone (Fl Bz) and spinifex (Fl Spfx). Volcanic flow facies fields from Barnes (2006). Modelled olivine compositions (Fo) in pure adcumulate shown on right hand side. Magma liquids in equilibrium calculated olivine compositions (Fo) shown on left hand side and along top. .................................................................................... 126

Figure 4.6. Plot of Al2O3 and TiO2 versus MgO for the basal flow within the Long-Victor area. Oxides as wt% and recalculated to anhydrous. Samples are characterized as channel facies B-zone (Ch Bz) and spinifex (Ch Spfx) and flank facies B-zone (Fl Bz) and spinifex (Fl Spfx). 126

Figure 4.7. Median primitive mantle normalized trace element plots of the samples from the basal flow in the Long-Victor area. Samples divided into channel and flank facies, and spinifex textured and B-zone cumulates. ............................................................................................... 127

Figure 4.8. Primitive mantle normalized chalcophile element metal diagrams for the basal flow within the Long-Victor area. Spinifex textured samples shown in black and B-zone cumulate samples in black. Normalizing values from McDonough and Sun (1995). .............................. 128

Figure 4.9. MgO wt% versus chalcophile element for all samples from the basal flow. Visual trends shown by dashed lines. ............................................................................................................. 129

Figure 4.10. PGE/Tipmn versus MgO wt% for all samples from the basal flow. Samples with S > 0.25 wt% on the left hand side and samples with S < 0.25 wt% on the right hand side. Samples are subdivided based on flow facies (channel = Ch, and flank = Fl) and komatiite flow facies (B-zone cumulates = Bz, and spinifex textured = Spfx). ............................................................... 130

Figure 4.11. Inter-chalcophile element relationships for samples from the Long-Victor basal flow with S>0.25wt%. ...................................................................................................................... 131

Figure 4.12. Platinum (ppb) versus sulfur (S wt%), and sulfur (S wt%) versus MgO (wt%) for sulfur-poor (S<0.25 wt%) Long-Victor basal flow samples. .............................................................. 132

Figure 4.13. Major and trace element abundances plotted as a function of distance from known mineralization (Ni >0.4%) which characterizes the channel (c.f. Fig. 4.3). Samples are classified as channel (Ch) and flank (Fl), as interpreted from constructed cross-sections. Samples are further subdivided based on texture: B-zone (Bz) and spinifex (Spfx). Median values for B-zones (solid line) and spinifex (dashed line) for channel and flank environments are shown. Calculated best fit lines for flank B-zones (blue) and spinifex (red) are shown, with R2 values for spinifex. Channel and flank subdivision at a distance of 100 m is based on data distribution. .............................................................................................................................. 135

Figure 4.14. A. Ni/Tipmn versus MgO for the Long-Victor system, basal flow samples shown in red diamonds. Calculated Ni normalized to actual Tipmn plotted as black triangles. Ni/Ti trend line based on a derived equation. B. Pt/Tipmn versus MgO for Long-Victor, basal flow samples shown in red diamonds. Calculated Pt normalized to actual Tipmn plotted as black triangles. Trend line of Pt/Ti represents perfectly incompatible elements at a determined constant ratio of 0.67. ......................................................................................................................................... 138

Figure 4.15. Plots of PGE/Tipmn versus MgO (wt%) for Long-Victor samples exhibiting chalcophile element enrichment based on Pt and Pd abundances. Samples are plotted as analytical data in red and calculated chalcophile element abundance in grey (Table 4.4). Samples with sulfur greater than 0.25 wt% are shown on the left hand side and samples with sulfur less than 0.25 wt% on the right hand side. Blue lines define the analytical uncertainly field around the numerically modelled background values (see Appendix C). ................................................. 139

Figure 4.16. Plots of Pt correlations to incompatible elements (TiO2 and S) and chalcophile elements (Pd, Ni) for the Long-Victor basal flow samples with low sulfide abundance (< 0.25 wt%) and a chalcophile element enrichment signature. ........................................................................... 141

Figure 4.17. Ni/Tipmn versus MgO (wt%) and Pt/Tipmn versus MgO (wt%) for Long-Victor basal flow samples, filtered to remove enrichment signature (Pt/Ti pmn <0.88 and Pd/Ti pmn < 1.65). Samples are plotted as analytical data in red and calculated chalcophile element abundance in grey with lines delineating ± 500 ppm uncertainty for Ni, and ±2 ppb uncertainty for Pt. ...... 143

Figure 4.18. Change in chalcophile element abundance from calculated background values (Δ) for sample from Long-Victor basal flow. Samples exhibiting enrichment signatures are removed. A. Calculated Pt (ppb) depletion, with modeled depletion lines of 100%, 75%, 50% and 0% shown. Dark grey shading delineates fields of uncertainty. B. Calculated Pt depletion versus Pd depletion with ±2 ppb uncertainty applied to both. C. Calculated Rh depletion versus Pt

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169

depletion with uncertainty shown by grey bars. D. Nickel depletion versus Pt depletion with uncertainty shown by grey bar. ................................................................................................ 143

Figure 4.19. Leapfrog 3D-model of chalcophile element (PGE) mineralization signatures within the basal flow of the Long-Victor channels. A. Lunnon Basalt surface with 0.4% Ni grade shell shown. B. Modeled surface of the basal flow spinifex with colour gradients representing ore forming signatures observed in the spinifex; green = background, blue = depletion, and red = enrichment. C. Mineralization signatures observed in the B-zone cumulate, projected to the modeled surface of the basal flow spinifex. ............................................................................. 146

Figure 4.20. Plot of distance (m) versus Ni grade (%) for all samples from the basal flow of the Long-Victor system. Distances are an average of the three closest Ni occurrences to each sample. Ni grade (%) represents the average Ni abundance for those three occurrences. ....... 147

Figure 4.21. A. Pt/Ti pmn and B. Pd/Ti pmn versus distance (m) to nickel mineralization. Samples are classified as Enriched (mineralized based on Pt/Ti and Pd/Ti ratios), Enriched LS (mineralized based on Pt/Ti and Pd/Ti ratios with S<0.25%), Depleted (chalcophile element depleted samples as determined from previous section) and Background (samples which exhibit no indication of chalcophile element enrichment or depletion). Plots are domained into three spatial regions A, B, and C based on predominant ore forming signatures at the respective distances. .................................................................................................................................. 148

Figure 4.22. Pt/Tipmn and Pd/Tipmn versus distance (m) to Ni mineralization, focusing on samples within 80 m of known mineralization. Enriched (mineralized based on Pt/Ti and Pd/Ti ratios), Enriched LS (mineralized based on Pt/Ti and Pd/Ti ratios with S<0.25%), Depleted (chalcophile element depleted samples as determined from previous section) and Background (samples which exhibit no indication of chalcophile element enrichment or depletion). ........ 149

Figure 4.23. Pt, Pd and Rh for each chalcophile element depleted sample from Area C shown as % depletion versus distance from mineralization ≥ 0.4% Ni. ...................................................... 150

Figure 4.24. Chalcophile element depletion (left) and enrichment (right) as a percentage change from the calculated background for each chalcophile element of the basal flow, from drill hole KD6024. No chalcophile element uncertainty was applied to the interpreted mineralization signatures. Samples 195.7, 196.0, 196.9, and 209 m from Lesher and Arndt (1995) and Lesher et al. (2001). ............................................................................................................................. 152

Figure 4.25. Schematic cross-section through interpreted paleo-volcanic setting of Victor and Long channels showing relative locations of flank environments. Chalcophile element enrichment zones shown in red dots, chalcophile element depletion shown in blue shading and areas of recharge (background) in grey. ................................................................................................ 154

Figure 4.26. Time sequence block model for the progressive emplacement, mineralization and preservation of chalcophile element ore forming signatures. Komatiite flows colour coded for chalcophile signature: green = background, blue = depleted, red = enriched. ......................... 156

List of Tables Table 4.1. Summary of geochemistry for the basal flow at Long-Victor: Median (Med), Maximum

(Max), Minimum (Min), Number of samples (N). Data filtered for S<0.25 wt%. Oxides are recalculated to anhydrous conditions and reported in wt%, metals and trace elements are reported as ppm unless denoted * then ppb............................................................................. 124

Table 4.2. Average (n=19) chalcophile element abundances, MgO and TiO2 content of spinifex textured samples from the Long-Victor area. (TiO2 and MgO as wt%, Ni, Cu, Co, Cr, Zr, Gd as ppm, and Ir, Ru, Rh, Pt, Pd, Au as ppb). .................................................................................. 128

Table 4.3. Comparison of geochemical and physical attributes of channel and flank facies. Compiled from Gresham and Loftus-Hills (1981); Lesher et al. (1984); Lesher (1989); Lesher and Arndt (1995); Lesher et al. (2001); Barnes (2006). ............................................................................ 134

Table 4.4. Equations derived and used to calculate background abundances of Ni, Pt, Pd, and Rh as a function of the MgO content of the sample. Complete list of equations provided in Appendix D. ............................................................................................................................................. 137

Chapter 5. Stratigraphic Control on the Maggie Hays deposit

Chapter 5. Stratigraphic Control on the Style of Komatiite Emplacement in the 2.9 Ga Lake Johnston Greenstone Belt, Yilgarn Craton, Western Australia

Abstract

Komatiites occur in many Archean greenstone belts and host significant nickel-

sulfide ore deposits. Establishing the stratigraphy and the control that stratigraphy

has on the emplacement and morphology of ultramafic magmatism is crucial to

understanding Archean geodynamic environments and the targeting of nickel sulfide

mineralization within these environments.

The 2.9 Ga Lake Johnston Greenstone Belt, in the southern portion of the Youanmi

Terrane of Western Australia contains komatiite flows and related subvolcanic

intrusions, mafic volcanic rocks, felsic volcanic rocks, banded iron formation and

sedimentary rocks. The stratigraphic sequence is intact, preserving original

sedimentary and igneous textures and contact relationships, despite being overturned

and variably deformed.

This study proposes that the lithostratigraphic succession and ultramafic intrusions

identified within the Lake Johnston Greenstone Belt record a transition from arc-

dominated to plume-dominated magmatism, accompanied by the establishment of a

banded iron formation-dominated sedimentary basin.

It is proposed that the rheological contrast between the felsic volcanic unit and

overlying banded iron formation acted as a stratigraphic barrier and magma trap for

ascending ultramafic magmas. The stratigraphic barrier inhibited the upward ascent

of ultramafic magma causing the development of a sub-volcanic magma chamber.

Magma trapped beneath the banded iron formation progressively inflated and spread

out along the contact, until over-pressuring breached the banded iron formation and

magma escaped, forming the overlying extrusive komatiites. Both the geodynamic

and lithological transitions gave rise to favourable substrate lithologies and an ideal

tectonic setting for formation of komatiite-hosted nickel sulfide ores.

Keywords: komatiite, Barberton-type, Archaean, Yilgarn Craton, Youanmi Terrane,

stratigraphy, magma trap

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5.1. Introduction

Komatiites are identified in many Archean greenstone belts throughout the world

and host significant nickel sulfide ore deposits, producing approximately 10% of the

world’s annual nickel production (Arndt et al., 2008). Understanding of the host

stratigraphy within these belts is crucial for the interpretation of the geodynamic

environment of komatiite emplacement (Dostal and Mueller, 1997; Mueller and

Corcoran, 2001; Mueller et al., 2005; Trofimovs et al., 2006; Houlé et al., 2008).

Three main settings for komatiite volcanism are identified on the basis of the

associated stratigraphy, and summarized by Arndt (2008) as: 1) deep-water “mafic

plain” oceanic setting with the association of tholeiite flood basalts and komatiites

(e.g. Kambalda; Gresham and Loftus-Hills, 1981); 2) convergent margins where

komatiites are associated with arc-type volcanism (Dostal and Mueller, 1997;

Hollings and Wyman, 1999; Hollings et al., 1999; Rosengren et al., 2008); and 3)

submerged continental platforms characterized by sequences of komatiites and

shallow continental sedimentation (quartz-arenite, conglomerates, banded iron

formation: Donaldson and de Kemp, 1998; Bleeker et al., 2000; Mueller et al.,

2005).

The deep-water mafic plain environment comprises a coherent substrate (derived

from the cooling and solidification of molten lava or magma: McPhie et al., 1993) of

mafic pillow basalt overlain by komatiite flows with a thin interleaving veneer of

sedimentary rock (e.g. Kambalda Dome area: Bavinton and Keays, 1978; Bavinton,

1981; Gresham and Loftus-Hills, 1981; Redman and Keays, 1985). This system is

dominated by komatiite lava flows, and no subvolcanic ultramafic intrusions are

observed within the lower basalt sequences (Houlé et al., 2008). Conversely,

convergent margin environments that are characterized by variably incoherent

substrates (composed of disaggregated particles of sedimentary and volcaniclastic

accumulations: Fisher, 1961; McPhie et al., 1993) host extensive intrusive komatiitic

systems (e.g. Mt. Keith ultramafic complex: Rosengren et al., 2005; 2007; Fiorentini

et al., 2007; Houlé et al., 2008) and associated komatiite lava flows. Submerged

continental platform environment comprises both coherent substrate (crystalline

basement) and incoherent overlying sediments. Ultramafic magmatism within this

environment occurs also as komatiitic sills and lava flows intercalated with the

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overlying sediments as observed in the Slave Craton, Canada (Bleeker and Hall,

2007).

Both the mafic plain and convergent margin ultramafic environments consist of

simple, single unit stratigraphic sequences prior to komatiite magmatism (coherent

basalt or incoherent volcaniclastic). The submerged continental platform

environment is a more complex, multiple unit stratigraphic sequence prior to

komatiitic magmatism. However, to date no thickened olivine cumulate bodies

(intrusive or extrusive), indicative of sustained high volume magma transport are

identified in this type of environment.

Limited work has been carried out examining the relationship between the pre-

existing stratigraphy of simple, single unit sequences and the morphology of

komatiite volcanism (Prendergast, 2003; Houlé et al., 2008; 2009). However, the

physical effects of more complex stratigraphic sequences, consisting of multiple

lithological units with contrasting coherent and incoherent properties and large

density contrasts between adjacent units, are completely unconstrained.

The Honman Formation from the Maggie Hays area within the 2.9 Ga Lake

Johnston Greenstone Belt of Western Australia (Fig. 5.1), provides an opportunity to

examine, in detail, a complex stratigraphic setting that hosts komatiite bodies

ranging from thin differentiated flow lobes to thick lenticular dunite bodies. These

komatiites are associated with regionally extensive banded iron formation, and

hence could be regarded as having formed in the submerged continental platform

environment. However, the presence of underlying felsic and mafic volcanic

sequences implies a more complex setting. Diamond drill cores generated during the

exploration and resource-delineation stages of the Maggie Hays and Emily Ann

nickel sulfide deposits provides excellent continuous, and unique three-dimensional

exposure for this area. This research complements the recent regional structural

study of the same area by Joly et al. (2008; 2009).

The purpose of the paper is to: 1) document the lithological succession in the

Honman Formation within the 2.9 Ga Lake Johnston Greenstone Belt and propose a

coherent stratigraphy through the use of volcanic and sedimentary petrology and

whole-rock geochemistry; 2) interpret a tectonic setting and tectonic progression

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model leading to komatiite magmatism based on the geochemistry and stratigraphic

succession, and 3) propose a stratigraphically controlled emplacement model for a

sub-volcanic intrusion. Constraining the physical stratigraphic controls on magma

emplacement provides an important tool for: the reconstruction of Archean terranes,

mineral exploration within these terranes (e.g. Prendergast, 2003; Fiorentini et al.,

2007; Finamore et al., 2008; Houlé et al., 2008), and the study of modern systems

that host intrusive bodies.

5.2. Regional Geology

The Lake Johnston Greenstone Belt is located in the south-eastern portion of the

Youanmi Terrane of the Archean Yilgarn Craton, Western Australia (Swager, 1997;

Cassidy et al., 2006). Two other greenstone belts; the Forrestania, and Ravensthorpe

Greenstone Belts are also located in the south-western and south-central portion of

the Youanmi Terrane (Fig. 5.1). The Forrestania, Ravensthorpe and Lake Johnston

Greenstone Belts are believed to be correlative, based on their similar stratigraphy

and the presence of Barberton-type komatiites associated with banded iron

formation (Perring et al., 1995; 1996; Swager, 1997; Barnes, 2006). The Lake

Johnston Greenstone Belt trends NNW-SSE and is approximately 100 km in length,

varying in width from 20 km to less than 6 km. The belt is bounded to the east and

west by Archean granitic batholiths and migmatitic gneisses. Upper greenschist to

amphibolite facies are present within the central portion of the greenstone belt with

peak pressure of 5-7 ± 2.1 kbars and temperatures of 596-678 ± 65°C (Joly et al.,

2008).

Deformation within the belt varies from zones of intense shearing and boudinage, as

observed within the felsic volcanic rocks, to undeformed igneous and sedimentary

textures observed within the komatiite and sedimentary rocks. Four deformation

phases are identified (Joly et al., 2008; 2009). The first phase (D1) is NNE-SSW

shortening resulting in the generation of large fold-nappes. This is followed by static

prograde metamorphism to amphibolite facies during emplacement of granitoid

intrusions. D2 is recognized as shortening due to dextral shearing in NNW-SSE to

NW-SE direction under peak metamorphic conditions. The D3 event is E-W

shortening, apparent from the development of crenulation cleavages. The final

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deformation event (D4) occurs under brittle conditions and is characterized by

steeply dipping N-NE trending dextral faults.

Figure 5.1. Yilgarn Craton showing subdivision of the South West Terrane, Youanmi Terrane and Eastern Goldfields Superterrane. Youanmi Terrane granite-greenstone belts (dark grey) include: Lake Johnston (LJGB), Ravensthorpe (RGB), Forrestania (FGB) and Southern Cross (SCGB) greenstone belts. Eastern Goldfields Superterrane granite-greenstone belts (medium grey) include: Norseman (NGB) and Kalgoorlie (KGB). Lake Johnston Greenstone Belt nickel mines include: EA (Emily Anne deposit) and MH (Maggie Hays deposit). Modified from Department of Industry and Resources (2008).

The Lake Johnston greenstone belt is divided into three Formations, from east to

west: the Maggie Hays, Honman and Glasse formations (Gower and Bunting 1972;

1976). Outcrop is limited and mining activities are restricted to the Maggie Hays and

Emily Ann Ni mines (Fig. 5.1). Age determinations (U-Pb) on zircon from the felsic

volcanic rocks of the Honman Formation (Fig. 5.2) indicate ages of 2921 ±4 Ma to

2903 ±5 Ma for the greenstone belt (Wang et al., 1996).

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Figure 5.2. Generalized stratigraphic column for the Lake Johnston Greenstone Belt; modified from Gower and Bunting (1976). * U-Pb age determinations from Wang et al. (1996).


The exploration drilling database for the area around the Maggie Hays Ni-mine (Fig.

5.3) was provided by Noril’sk Nickel Pty. Ltd. (formerly LionOre Ltd.). Leapfrog®,

a 3D numerical modeling program was used to generate lithological block models

from the diamond drilling database. The lithological block model was subsequently

used to select 47 key diamond drill holes for detailed examination and sampling of

the entire Honman Formation lithostratigraphic sequence.

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Figure 5.3. Geological plan map of the study area within the Lake Johnston Greenstone Belt, showing the Honman and Maggie Hays Formations. Honman Formation is subdivided into lithologic units. Strong deformation at the northern end and along basal contact of the Central-UU in proximity to remobilized Ni sulfide mineralization shown as wavy lines. All diamond drill holes examined are shown, and key drill holes referenced in the paper labeled.

Of the initial 294 diamond core samples, a subset of 157 of the least altered samples

were selected for whole rock geochemical analysis. Samples averaging 1 kg were

coarse crushed at the University of Western Australia using a mechanical jaw

crusher, which was flushed with quartz and cleaned with wire brush and compressed

air between samples. Further sample preparation was carried out at Ultratrace

Laboratories in Perth, Western Australia, and consisted of pulverization in a

tungsten carbide mill. Loss on ignition (LOI) was determined gravimetrically

between 105-1000°C. A sample split was fused to form a glass bead for X-Ray

fluorescence spectrometry (XRF) for analyses of Al2O3, TiO2, MgO, Fe2O3, MgO,

Na2O, K2O, CaO, BaO, Cr2O3, P2O5, V2O5, ZrO2, SO3, Cu, Ni, Rb and Sr. Rare-

173


earth elements and Th, Nb and Se were analysed by inductively coupled plasma

mass spectrometry (ICP-MS) following a four-acid dissolution in closed beakers.

5.4. Stratigraphy and Geochemistry

The Honman, Maggie Hays and Glasse Formations within the Lake Johnston

Greenstone Belt (Fig. 5.2) were interpreted by Gower and Bunting (1972; 1976) as a

conformable sequence based on outcrop mapping. The Honman Formation, the only

formation of the three to contain komatiites, comprises five laterally continuous

lithological units (Fig. 5.4); from oldest to youngest: 1) Felsic Volcanic Unit (FVU),

2) interbedded felsic volcanic rocks, sediments and iron formation, termed the

Transition Zone Unit (TZU), 3) Banded Iron Formation (BIF) Unit, 4) quartz-rich

sedimentary rocks (consisting of quartz-arenite and massive sulfide) of the

Sedimentary Unit, and 5) extrusive komatiites of the Western Ultramafic Unit

(WUU: Fig. 5.4). In addition, a series of ultramafic intrusions of the Central

Ultramafic Unit (CUU) and Eastern Ultramafic Unit (EUU) cross-cut parts of the

Honman Formation. The five laterally continuous units strike north-northwest and

south-southeast and dip approximately 60-70° to the east with younging direction to

the west, indicating that the formation has been overturned. The majority of the

diamond drilling has been carried out from the east and is west-directed

(approximately perpendicular to strike and dip) and up-stratigraphy (Fig. 5.3; 5.5).

Figure 5.5 is a computer generated block model of the Honman Stratigraphy as

modelled from the diamond drill hole database. The block model shows the spatial

distribution and the lateral continuity of the lithological units with depth and along

strike.

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Figure 5.4.Composite stratigraphic column for the Honman Formation as observed from diamond drill cores (LJD0126, LJD0048, LJD0011, LJD0054A, LJD0087A, LJD003A, LJD0039, LJD0038, LJD0049, LJD0074, LJD0055W2, LJD0092). Approximate intrusive level of the Central Ultramafic Unit and narrow intrusive sills (banded iron formation-hosted sills) shown along the left hand side.

Textural preservation is variable within the greenstone belt. Most of the primary

igneous textures that occur stratigraphically below the BIF (Fig. 5.4) have been

obliterated, as observed in the CUU, and variably deformed as observed in the felsic

volcanic rocks. Units above the BIF are remarkably well-preserved and contain

175


abundant pseudomorphed igneous and primary sedimentary textures. Therefore,

lithological classification is based on preserved textures where present and current

metamorphic mineral assemblages or whole-rock geochemistry in extensively

altered lithologies. The following descriptions use igneous and sedimentary

terminology to refer to the protolith rather than the observed metamorphosed

equivalent. The prefix “meta-” has been omitted from rock names but can be

universally assumed.

Figure 5.5. Oblique Leapfrog® model view looking down and north-east towards the local Maggie Hays nickel-deposit stratigraphy. Stratigraphy from left to right consists of the Banded Iron Formation Unit, Transition Zone Unit, Central Ultramafic Unit and Felsic Volcanic Unit. Scale bar in metres. Western ultramafic unit not shown for clarity, but occurs to the left of the Banded Iron formation.

a. Felsic volcanic unit

The Felsic volcanic unit (FVU) occurs at the base of the Honman Formation (Fig.

5.4), and has a minimum thickness of 600 m. Uranium-lead age determination from

zircons extracted from a feldspar porphyritic rock range from 2921 ±4 Ma to 2903

±5 Ma (Wang et al., 1996). Volcanic textures are commonly very coarse, as

recrystallization generally obliterates the finer more delicate igneous textures. As a

result, three lithologies of the Felsic Volcanic Unit are recognized and comprise: 1)

felsic volcaniclastic rocks, 2) feldspar-porphyritic rocks, and 3) quartz-porphyritic

rocks.

The felsic volcaniclastic lithology comprises approximately 60% of the volume of

the felsic volcanic unit. It is white to light grey in colour and dominated by pebble-

176


sized fragments (commonly less than 10 cm) hosted within a sand-sized (< 2 mm)

matrix. A minor volcaniclastic sub-unit is defined by homogenous sand-sized grains,

and forms layers 1-2 m in thickness. Breccia fragments preserve some internal

structure and range from massive, amygdaloidal to porphyritic in texture. They

exhibit length to width ratios of approximately 3:1 parallel to the local foliation.

Thickness of individual volcaniclastic units varies from approximately 10 cm to 10

m.

The feldspar porphyritic lithology represents approximately 20% volume of the

FVU and is grey-white in colour and massive in appearance. Feldspar phenocrysts

are observed throughout the lithology ranging in size from 2-5 mm in diameter and

commonly between 20-30% in volume. The individual units vary in thickness from

~ 1 to 10 m.

The quartz porphyritic lithology also represents approximately 20% volume of the

unit and increases in abundance towards the top of the FVU. The lithology is

visually homogenous and white in colour with minor dark amphiboles defining a

moderate foliation. The lithology is fine-grained with 10-20% fine to medium (1-3

mm) phenocrysts of quartz occurring throughout. Small garnet porphyroblasts (< 3

mm) are dispersed (<< 1%) within the unit. Discontinuous intercalations of

sedimentary lithologies are observed throughout the felsic volcanic unit and increase

in abundance towards the top of the unit.

Intercalations are thin (<10 cm to 1 m) and either pelitic (mudstone) or semi-pelitic

(sandstone) in composition with the later dominated by sand sized grains (2-5 mm).

Mudstone lithologies appear massive and dark grey-green in colour, dominated by

Fe-Mg-silicate mineralogy consisting of biotite, actinolite, and grunerite. Locally,

high abundances (> 60%) of coarse idioblastic and porphyroblastic garnet are

present with minor fine disseminated or narrow stringer sulfide (< 1 vol.%).

Sandstone lithologies dominated by quartz and feldspar with minor mafic minerals

present. Intervals commonly appear massive, with rare thin sedimentary laminations.

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Figure 5.6. Jensen cation plot from the Felsic Volcanic Unit and ultramafic units from the Lake Johnston Greenstone Belt: felsic volcanic rocks, Central Ultramafic Unit (CUU) pyroxenites and olivine cumulates, and Western Ultramafic Unit (WUU) komatiites. H-Fe th as (high-Fe tholeiitic andesite), H-Mg th ba (high-Mg tholeiitic basalt).

The felsic volcanic rocks are characterized by SiO2 contents from 62-77 wt% (Table

5.1), and range from dacite to rhyolite in composition (Fig. 5.6). Major oxides FeO,

MgO, TiO2, Al2O3, CaO, and K2O, and trace elements Sr, Zr, Ni and Cr, decrease in

abundance with decreasing SiO2. A positive correlation is observed between SiO2

and Na2O, and most of the rare-earth elements. Multi-element primitive mantle

normalized plots indicated all of the felsic volcanic rocks are enriched in light rare-

earth elements (LREE) relative to heavy rare-earth elements (HREE: Fig. 5.7).

Additionally, the HREE patterns are flat to slightly concave downward. The felsic

volcanic rocks have pronounced negative Nb, Sr and Ti anomalies, but no apparent

negative Eu anomaly.

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Table 5.1. Whole rock geochemistry analyses of representative units from the Honman Formation. With drill collar, sample depth, lithological unit (FVU = Felsic Volcanic Unit; WUU = Western-UU; CUU = Central-UU) and lithology (Rhy-dac = rhyolite-dacite; Spfx = spinifex; OC = olivine cumulate, Pyr = pyroxenite) in header. Trace element ratios La/Sm*, Th/Sm*, Nb/Th* and Gd/Yb* primitive mantle normalized. Normalization values from McDonough and Sun, (1995).

LJD0018-

265.60 LJD0048-

164.05 LJD0126 313.10

LJD003A-524

LJD107-445

LJD0069- 238.00

LJD0077- 345.10

FVU WUU WUU CUU CUU CUU CUU (wt%) Rhy-dac Spfx Spfx OC OC Pyr Pyr

SiO2 73.70 44.90 44.40 39.46 40.09 47.70 42.60TiO2 0.35 0.45 0.45 0.09 0.06 0.36 0.36

Al2O3 14.30 4.86 3.91 1.21 0.72 4.31 4.25FeOt 0.95 11.25 11.34 8.17 7.67 8.91 9.45MnO n.d. n.d. n.d. n.d. n.d. n.d. n.d.MgO 0.77 24.30 27.40 38.10 46.20 24.00 23.30CaO 1.40 7.63 6.20 0.35 0.07 9.04 12.10

Na2O 4.83 0.08 0.16 0.06 0.10 0.14 0.15K2O 2.87 n.d. 0.02 0.16 n.d. 0.05 0.05

P2O5 0.09 0.04 0.04 0.01 0.01 0.03 0.04LOI 0.38 4.78 4.04 11.40 4.35 4.19 5.97

Total 99.6 98.3 98.0 99.0 99.3 98.7 98.3

Al/Ti 40.9 10.8 8.7 14.1 11.4 12.0 11.8(ppm)

Ni 30 1080 1590 4520 3020 1680 1230Cu 5 60 70 50 20 80 50Cr 21 2662 2600 2270 1920 2210 2128

Rb 20 n.d. n.d. 50 40 n.d. n.d.Sr 90 30 30 30 50 30 30V 50 151 140 60 40 129 123Y 6 9 8 2 2 8 8Zr 148 n.d. n.d. 30 35 n.d. n.d.Th 6.65 0.65 0.25 0.20 0.10 0.55 0.50Nb 1.00 0.60 0.70 0.20 0.20 0.70 0.60Hf 2.00 0.30 0.20 0.10 n.d. 0.60 0.50Ta 0.10 n.d. n.d. n.d. n.d. n.d. n.d.La 7.00 1.50 1.00 2.70 0.95 3.50 3.00Ce 13.00 4.00 3.50 5.40 2.80 7.00 6.50Pr 1.40 0.60 0.60 0.55 0.35 1.00 0.80

Nd 6.00 3.50 3.00 2.00 1.85 4.50 3.50Sm 1.00 1.00 1.00 0.40 0.40 1.00 1.00Eu 0.40 0.40 0.40 0.10 0.05 0.40 0.40Gd n.d. n.d. n.d. 0.35 0.50 n.d. n.d.Dy 1.00 2.00 1.50 0.35 0.40 1.50 1.50Tb n.d. 0.20 0.20 0.05 0.05 0.20 0.20Ho 0.20 0.40 0.40 0.05 0.05 0.40 0.40Er 0.50 1.00 1.00 0.20 0.20 1.00 1.00

Tm n.d. n.d. n.d. n.d. n.d. n.d. n.d.Yb 0.50 1.00 0.50 0.20 0.15 1.00 1.00Lu n.d. n.d. n.d. n.d. n.d. n.d. n.d.

La/Sm* 4.39 0.94 0.63 4.23 1.49 2.19 1.88Th/Sm* 33.96 3.32 1.28 2.55 1.28 2.81 2.55Nb/Th* 0.02 0.11 0.34 0.12 0.24 0.15 0.14Gd/Yb* 1.42 2.70

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i. Interpretation of the felsic volcanic unit

The presence of thin intercalated sedimentary rocks within the felsic volcanics

indicates volcanism was in a sub-aqueous environment. Volcaniclastic lithologies

with mixed fragment assemblages support the hypothesis that the unit was deposited

as either a primary pyroclastic deposits or re-sedimentation of pyroclastic and

autoclastic felsic material. Massive units (quartz and/or feldspar pheric) occur as

either intrusions or coherent lava flows throughout the unit.

Major element variation defines a calc-alkaline trend (Fig. 5.6). Negative Nb and Ti

anomalies with LREE enrichment over HREE (La/Ybpm ~10), and HREE patterns

that are generally flat to slightly concave downward (Fig. 5.7) are indicative of both

modern subduction related, arc-type volcanism (Pearce, 1982; Pearce and Peate,

1995) and Archean tronjemite-tonalite-granodiorite (TTG) series volcanism (Martin,

1994). However, TTG systems typically exhibit strong HREE depletion (La/Ybpm ~

34: Morris and Witt, 1997; Brown et al., 2001), whereas arc-type systems within the

Yilgarn Craton exhibit moderate HREE depletion (La/Ybpm ~ 10: Messenger, 2000;

Brown et al., 2002; Barley et al., 2008), which is more akin to that observed within

the Honman Formation felsic volcanic unit (La/Ybpm of 10).

Figure 5.7. Primitive mantle-normalized trace element patterns for the Felsic Volcanic Unit shown as black lines. Data fields for TTG/TTD type (Black Flag Formation: Morris and Witt, 1997) and Arc-type felsic volcanism from Eastern Goldfields Superterrane (EGS: Morris and Witt, 1997; Messenger, 2000; Barley et al., 2008). Normalizing values from McDonough and Sun (1995).

A subduction-related arc setting appears to be the best interpretation of the Honman

Formation felsic volcanic rocks. Subduction zones occur below either a pre-existing

sialic basement or on juvenile oceanic crust. Subduction occurring beneath a pre-

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existing sialic basement commonly contains older inherited zircons, as documented

in the Kalgoorlie Terrane; interpreted as evidence for the presence of an older

ensialic crust, and a potential source of felsic magmas (Campbell and Hill, 1988;

Compston et al., 1986; Chauvel et al., 1985). Age determinations carried out on

zircons extracted from the Felsic Volcanic Unit by Wang et al. (1996) did not

identify an older inherited zircon population, supporting subduction associated with

juvenile oceanic crust.

b. Transition zone unit

The transition zone unit (TZU) had not been identified previously within the

Honman Formation stratigraphy. Since its identification, the interpretation of the

unit has been contentious as to whether it represents a stratigraphic unit, or a

structural zone (altered shear zone/complexly folded altered shear zone). The TZU is

a heterogeneous sequence approximately 55 m thick, occurring between the

underlying felsic volcanic unit and overlying BIF Unit (Fig. 5.4). The TZU is

identified in 83 drill holes to a maximum depth of approximately 900 m below

surface. It maintains the same stratigraphic position at both the mine scale

(approximately 3 km strike length within the Maggie Hays mine area: Fig. 5.5 and

regionally, as observed in drill core 8 km north of the Maggie Hays mine.

Lithologically, the base of the TZU is defined by a garnet-grunerite to garnetite layer

varying in thickness from 3 m in drill hole LJD0038 to approximately 10 m in

LJD0039 (Fig. 5.3). The transition between the FVU and TZU is gradual, as an

increase in garnet abundance is observed over the preceding 10 m, from rare

porphyroblastic garnet (<< 1 vol.%) to abundant porphyroblasts (20 vol.%) in the

underlying fine-grained felsic volcaniclastic rock as the basal TZU contact is

approached (Fig. 5.8A). The garnet-grunerite to garnetite is followed by fine-grained

massive siliceous rock with narrow sulfide (pyrrhotite) stringers and fine

disseminated sulfides ranging from 5-10 vol.% in abundance (Fig. 5.4 & 5.8A). The

siliceous rock has an observed thickness of approximately 7 m in drill hole

LJD0087A and exhibits a decrease in sulfide abundance up-stratigraphy.

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Figure 5.8.Drill core photos and photomicrographs of representative Honman Formation units. A. Part of the Transition Zone (TZ) Unit from LJD0038. Felsic Volcanic Unit lithology with minor garnet on left, garnetite in middle (magnified in B.), and chert with minor sulfide on right. B. Garnetite lithology (LJD0038). C. Banded Iron Formation Unit (LJD0011). D. Iron-poor Fe-formation. E. Spinifex texture from the Western-UU (LJD0011). F. Flow top breccia texture from the Western-UU (LJD0126). G. Polarized light photomicrograph of garnetite (LJD0038) amp = amphibole, bio = biotite, grt = garnet. H. Reflected light photomicrograph of quartz-arenite (quartz with trace pyrite), exhibiting graded bedding (LJD0011).

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A sharp transition occurs to a thin (0.9 m) amphibole-rich lithology that contains

abundant porphyroblastic garnet at the top and bottom contacts. A massive fine-

grained siliceous lithology (14 m thick) overlies the amphibole-rich lithology and

exhibits an increase in iron (magnetite) content down the drill hole (interpreted up-

stratigraphy) with minor (< 2 vol.%) sulfide (pyrite and pyrrhotite) observed

throughout (Fig. 5.4). A narrow (1.5 m thick) dark green amphibole-rich lithology

with pyrrhotite stringers and variable abundance of garnet bands is observed.

Overlying this lithology is a 30 m thick sequence of fine-grained white quartz

porphyritic felsic volcanic rocks containing trace porphyroblastic garnet, intercalated

with minor thin (0.3-4 m) quartz-rich sandstones (light grey in colour) with thin

laminar graded bedding, indicating a younging direction to the west and massive

dark green amphibole-rich lithologies. A massive fine-grained siliceous lithology

with minor sulfide (< 2 vol.%), minor magnetite, and sparse intervals of garnet rich

bands (5-10 cm thick) occurs above for approximately 4 m, and marks the contact

between the TZU and the overlying BIF Unit.

A weak to moderate foliation permeates the TZU. Primary sedimentary structures

(graded bedding, planar bedding) are preserved and visible where lithologies are not

massive. Bedding observed through the TZU in drill core LJD0087A dips at a

constant angle over the full 75 m thickness (Fig. 5.3).

i. Interpretation of the TZU

Several factors support the contention that the TZU is a stratigraphic unit, rather

than a structural zone: the gradational contact relationship between the underlying

FVU, the presence of sedimentary structures (graded bedding) along with the

consistent stratigraphic position of the TZU, the lack of apparent shear zones, and a

consistent younging direction. The garnet-gunerite, garnetite, amphibolite-rich

lithologies and the massive fine-grained silicified lithologies (Fig. 5.8B, G) are

interpreted to be of sedimentary origin. Garnet and amphibole-rich lithologies are

derived from a metamorphosed iron and clay-rich protoliths (Klein, 2005), and

silicified lithologies derived from a high silica protolith, rather than alteration. This

interpretation fits with the proposed depositional environment for the Honman

Formation, as a transition between felsic volcanism and BIF accumulation. Iron and

clay-rich protoliths are derived from the accumulation of iron from exhalative

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sources (Klein and Beukes, 1989; Frei and Polat, 2007; Ohmoto et al., 2006) and

clay from hemipelagic rainout (Klein, 2005) as documented in other silicate facies

iron formations (Spry et al., 2000; Heggie, 2002; Slack et al., 2009). The increasing

abundance of metamorphic garnet in the lower part of the unit reflects an increase in

the abundance of exhalative iron accumulating with the underlying felsic

volcaniclastic unit during the waning stages of volcanism.

Exhalative venting appears to be low temperature (< 200°C) as other higher

temperature phases (i.e. copper, zinc: Ohmoto et al., 2006) are not visually apparent

within the system. Oxygen activity during the deposition of the TZU was variable,

as the unit contains both Fe-oxides and Fe-sulfides. The presence of both oxides and

sulfides indicate conditions were oxic to moderately reducing during iron

precipitation and anoxic to very-reduced euxinic during the formation of sulfide

phases (Klein, 2005).

The presence of thinly-graded sedimentary beds within the TZU indicates a minor

detrital component was still being contributed to the depositional basin during the

accumulation of this unit. These sediments represent the product of small periodic

debris flows from the shallower margins of the basin, accumulated from the erosion

of the exposed basin margin. Hemipelagic rainout and detrital input ceases during

the transition to the overlying BIF Unit.

c. Banded iron formation unit

The banded iron formation (BIF) Unit forms the second thickest unit in the Honman

Formation (Fig. 5.4, 5.5) with maximum thickness of more than 190 m and average

thickness of 120 m. The lateral continuity of the BIF Unit is well constrained

throughout the greenstone belt by drilling and airborne magnetic surveys (Geol.

Survey WA, 2005). The unit is locally deformed with tight isoclinal folding. It is

characterized by alternating macrobands (5-10 mm thick) of silica and magnetite

(James, 1954) in the central portion (Fig. 5.4, 5.8C). Although classically banded in

the centre of the unit, the BIF exhibits variability in the abundance of magnetite

towards the top and bottom of the unit. A trend of increasing magnetite content is

observed upward from the TZ Unit where siliceous and aluminum silicate minerals

are dominant, and the reverse trend of decreasing magnetite content is observed in

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the upper 32 m of the unit, which becomes increasingly dominated by siliceous and

aluminum silicate mineralogy (Fig. 5.8D) towards the top contact with the quartz-

arenite unit.

i. Interpretation of the BIF unit

The BIF unit is interpreted to represent the continued accumulation of iron within

the basin from hydrothermal activity initially identified within the TZU. However,

the co-accumulation of volcaniclastic and detrital material into the basin ceases,

resulting in the deposition of a thick (120 m) homogenous hydrothermal-

sedimentary sequence of alternating magnetite and chert layers forming classic

oxide-facies banded iron formation. The end of the BIF deposition was not abrupt,

rather the sequence exhibits a gradual decrease in iron (magnetite) and increase in

aluminum silicate abundance at the top. The observed progression implies a

decrease in both hydrothermal activity and temperature of the venting fluids with an

increase in the deposition of aluminum-enriched pelagic and hemipelagic sediment

into the basin.

Banded iron formation deposition is interpreted to occur during volcanic hiatuses,

forming repeating cyclical sequences of volcanic activity-BIF accumulation, which

has been documented in both the Abitibi Greenstone Belt (Thurston et al., 2008) and

the Brockman Supersequence of the Hamersley Ranges (Krapež et al., 2003). In the

Abitibi Greenstone Belt, BIF accumulation occurs at multiple stratigraphic levels

prior to komatiite magmatism in numerous volcanic episodes, as well documented

between the 2734-2724 Ma Deloro Assemblage (mafic to felsic volcanic rocks with

several horizons of iron formation at the top) and the 2710-2704 Ma Tisdale

Asssemblage (mafic to ultramafic rocks: Thurston et al., 2008). The Brockman

Supersequence cyclical sequences described by Krapež et al. (2003) are

characterized by mudrock defining the base of the sequences, representing a period

of lowstand deposition in the basin. A gradual transition upward to overlying BIF is

observed resulted from a transgression from lowstand to highstand conditions within

the basin. Deposition during basin highstand conditions is restricted, and results in a

condensed section/depositional gap as observed in the homogenous sequences of

BIF. The top of the highstand sequences are marked by a gradual decrease in iron

content, with continued chert deposition, interpreted by Krapež et al. (2003) to

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indicate that BIF deposition within the basin ceased prior to the start of the next

sequence and deposition of lowstand mudrock. This detailed stratigraphic work led

Krapež et al. (2003) to hypothesize a causal relationship between hydrothermal

activity-volcanic activity and rising and falling sea level.

Similar to the cycles within the Abitibi and Brockman sequences, the BIF Unit

within the Honman Formation follows felsic volcanic activity, contains a transitional

unit (TZU) and exhibits a decreasing iron content up-stratigraphy from the middle

homogenous portion of the unit. However, within the Honman Formation, the top of

the sequence and the start of another sequence is not marked by lowstand mudrock,

but rather by quartz-arenite and massive sulfide pre-staging renewed volcanic

activity.

d. Sedimentary unit

The sedimentary unit identified within the Honman Formation (Fig. 5.4) is the

thinnest identified stratigraphic unit and consists of a quartz-arenite and massive

sulfide sub-units.

Quartz-arenite sub-unit: Quartz-arenite sub-unit is only observed in a single drill

hole (LJD0011: Fig. 5.3) during the field work. However, this sub-unit is identified

in the regional mapping by Gower and Bunting (1976).The quartz-arenite sub-unit is

approximately 3.2 m thick, and contains minor narrow (< 10 cm) bands of garnet-

grunerite and sulfide-rich intercalations. The quartz-arenite lithology is characterized

by 90-95% sutured quartz grains varying from 0.1 to 2 mm in diameter. Trace

minerals amphibole, biotite, muscovite and garnet define a weak to moderate

foliation throughout the sub-unit. Primary laminar graded bedding is visible

throughout (Fig. 5.8H) and indicates a younging direction to the west. Minor

disruptions to the bedding by small faults with visible displacement of centimeters

are observed. Transition from the underlying BIF is sharp without any visible

intercalation of underlying highly siliceous rocks into the quartz-arenite.

Massive sulfide sub-unit: The massive sulfide sub-unit is only identified in four

drill holes (LJD0011, LJD0048, LJD0050, LJD0049) as the contact is rarely drilled.

However, the four drill holes represent 2.5 km of strike length, thus providing a

minimum strike length for the unit (Fig. 5.3).The massive sulfide sub-unit marks the

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top of the sedimentary package of the Honman Formation. This sulfide-bearing unit

is approximately 2 m thick and directly overlies the quartz-arenite sub-unit. The

transition is gradational, with minor narrow stringers of sulfide observed in the

underlying quartz-arenite. The massive sulfide sub-unit is dominated by massive

nodular pyrite, narrow stringers of pyrrhotite and a fine matrix of quartz.

i. Interpretation of the sedimentary unit

The abrupt transition observed between the BIF Unit and overlying Sedimentary

Unit marks a substantial change in the depositional system. The quartz-arenite sub-

unit dominantly consists of thin planar graded beds interpreted to represent the

depositional product of debris flows. Debris flows of quartz-rich detrital material

require the extensive sub-aerial weathering of a protolith (Thurston and Kozhevniko,

2000). Within the Honman Formation stratigraphy, the most plausible source is the

felsic volcanic rocks, sub-aerially exposed in the periphery of the basin undergoing

weathering, erosion and accumulation.

Thin layers enriched in iron-silicates similar to lithologies observed within the TZU

are intercalated within the debris flows, implying that hydrothermal-Fe and Al-rich

detrital material (clay) was also accumulating. The top of the quartz-arenite sub-unit

is marked by a gradual increase in sulfide content in the form of iron sulfides and

minor copper sulfides, indicating either a change in oxidation state, sulfur

availability, or the onset of more proximal hydrothermal venting in the area.

Substantial changes in the sedimentary unit in both the material accumulating (BIF

to detrital sedimentation) and the facies of iron formation (oxide to sulfide) are a

prelude to the eruption of the Western Ultramafic Unit komatiites (WUU) at the top

of the Honman Formation sequence.

e. Ultramafic units The Honman Formation contains three ultramafic units: Eastern Ultramafic Unit

(EUU), Central Ultramafic Unit (CUU), and Western Ultramafic Unit (WUU). The

three ultramafic units occur at distinct stratigraphic settings (Fig. 5.3) and preserve

igneous relationships and textures allowing for interpretation. The EUU occurs on

the eastern side of the Honman Formation, and has been defined by diamond

drilling. Although a significant igneous body, no mineralization has been identified

within the EUU, and no research has been carried out on its relationship to the CUU

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and WUU. The Central Ultramafic Unit (CUU) and the Western Ultramafic Unit

(WUU) occur on the western side of the Honman Formation and are intimately

associated with Ni mineralization as the Maggie Hays deposit is hosted within the

CUU. The CUU and WUU are stratigraphically within 200 m of each other, with the

WUU overlying the CUU.

Central ultramafic unit

The CUU has good exposure through diamond drilling, and two ultramafic sub-units

are identified (Heggie et al., 2007). The first sub-unit is a volumetrically minor set of

ultramafic bodies hosted within the BIF unit (BIF-hosted ultramafic). The second

sub-unit (CUU proper) is the major volumetric ultramafic body that lies to the east

of the BIF unit and hosts the Maggie Hays nickel sulfide mineralization (Fig. 5.10).

Figure 5.9. Plot of FeO versus MgO (anhydrous wt.%) for WUU (spinifex and B-zone cumulates) and CUU, subdivided into spatial zones as shown in Figure 5.10 (zone 0 = gabbroic, zone 1 = amphibolite, zone 2, 3, 4 = olivine cumulates). Modified from Barnes et al., (2004).

BIF-hosted ultramafic sub-unit: The BIF-hosted ultramafic sub-unit of the CUU

are observed in five drill holes and characterized thin (< 20m) ultramafic bodies

hosted within the BIF unit. Ultramafic bodies have sharp top and bottom contacts

with light grey-green amphibole-rich reaction zones with coarse magnetite crystals.

The central portions of the ultramafic bodies are uniformly fine-grained, dark green

to black in colour and comprise amphibole-chlorite rock. The size, lateral continuity

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and total number of the BIF-hosted ultramafic bodies are not known due to limited

exposure. Drill hole LJD0011 intersects five BIF-hosted ultramafic bodies ranging

in thickness from 2-17 m with a total thickness of approximately 43 m, whereas,

LJD0120 intersects three ultramafic bodies ranging in thickness from 3-16 m with a

total thickness of 26 m (Fig. 5.10).

Main ultramafic sub-unit: The main ultramafic sub-unit (CUU proper) of the CUU

(Fig. 5.3), strikes parallel to sub-parallel to the regional magnetic trend as defined by

the BIF Unit for approximately 3 km, reaching a maximum stratigraphic thickness of

350-400 m at the northern extent as identified from diamond drilling. The main

ultramafic sub-unit hosts the Maggie Hays nickel sulfide deposit (Barnes, 2006) and

is dominated by mutually gradational olivine cumulate lithologies (mesocumulate to

adcumulate: Fig. 5.12C), flanked by lesser pyroxenite and minor gabbro-troctolite.

Pyroxenite is identified along all ultramafic-host rock contacts and varies in

thickness from 1 m up to 10 m in thickness. Pyroxenite is also observed as a

transitional phase between olivine cumulates in the central core and gabbro-

troctolite lithology occurring along the western side of the CUU (Fig. 5.10). Gabbro-

troctolite is observed in two drill holes (LJD003A and LJD051) but is not

constrained in true thickness.

Contact relationships observed at the top and bottom contacts in numerous drill

holes are sharp and un-deformed (Fig. 5.12A), and characterized by the ubiquitous

occurrence of a border pyroxenite lithology adjacent to the host-rock contact. The

border pyroxenite lithology grades into more olivine-rich lithologies farther from the

contact, similar to that observed at Mt. Keith (Rosengren et al., 2005). Small

xenoliths and xenomelts (Fig. 5.12B) are observed sporadically in proximity to both

the eastern (paleo-base) and western (paleo- top) contacts of the CUU. The CUU

transgresses the stratigraphy and progressively increases in width and thickness from

south to north. At the southern extent of body, the CUU is hosted within the felsic

volcanic unit. At the northern extent the top contact is against BIF unit and the basal

contact is against the felsic volcanic unit (Fig. 5.3).

As a whole, the CUU has undergone retrograde alteration to serpentine, talc, and

chlorite. Consequently, the majority of primary and prograde metamorphic minerals

have been obliterated. Minor primary mineralogy is preserved in relict olivine

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fragments in the olivine adcumulate and mesocumulate lithologies, and relict

clinopyroxene observed in the more orthocumulate portions of the CUU. Randomly

orientated anthophyllite is observed in wide bands (10-100s m) throughout the CUU

and represents a prograde metamorphic mineral assemblage formed by inward

migration of water-rich fluids during dehydration of early-formed marginal

serpentinite, as observed in the Perseverance (Agnew) ultramafic body (Gole et al.,

1987). The margins of the CUU (20-60 m thick) have lower MgO contents (Table

5.1.), which dominantly consist of tremolite and chlorite, with or without

metamorphic olivine, and preserve contact relationships in low-strain areas.

Figure 5.10. Cross-section from line 6430470mN through the Honman Formation stratigraphy, showing stratigraphic succession (Felsic Volcanic Unit, Transition Zone Unit, BIF Unit, WUU) and conformal setting of the Central Ultramafic Unit (CUU) and smaller banded iron formation-hosted ultramafic sub-unit (BIF-hosted intrusions). Spatial geochemical zones shown in within the CUU (as used in Figs. 5.9), zone 0 = gabbroic; zone 1= pyroxenite; zone 2 = mixture of adcumulates to orthocumulates with lower forsterite olivine; zone 3 = dominant adcumulates with moderate forsterite olivine (Fo90-92); zone 4 = olivine adcumulates with highest forsterite content (Fo93-94).

Overall, major and trace element abundances from the CUU exhibit control by

varying degrees of accumulation of olivine. The CUU is characterized by strong

negative correlations between TiO2, Al2O3, FeOtot and CaO with MgO, with an

average Al2O3/TiO2 ratio of approximately 14, intermediate between Barberton

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(Al2O3/TiO2 ~10) and Munro-type (Al2O3/TiO2 ~20) komatiites (Lesher and Stone,

1996: Fig. 5.13). Trace element abundances for the CUU exhibit flat primitive

mantle normalized patterns (Fig. 5.11) with slight HREE depletion and strong

positive La anomalies relative to Munro- and Barberton-type komatiites.

Figure 5.11. Primitive mantle normalized trace element patterns of select samples from the CUU (blue lines), WUU (grey lines) and mean FVU (red line). Data from Chapter 6 and Appendix B. Normalizing values of Sun and McDonough (1989).

The main portion of the CUU body comprises olivine cumulates that are

characterized by median values of 42.2 wt% MgO, 0.1 wt% TiO2, 1.6 wt% Al2O3,

8.4 wt% FeOtot, 0.2 wt% Cr2O3 and 2800 ppm Ni (Table 5.1). Whole-rock FeO and

MgO exhibit a range from 4-14 wt% and 12-50 wt% respectively, and represent a

mixture of olivine cumulates (orthocumulates to adcumulates) and variable amounts

of fractionated and locally contaminated trapped liquid. Olivine adcumulate

lithologies exhibit a range in inferred olivine composition from Fo85 to Fo94, with the

majority of the samples falling between Fo90 and Fo94 (Fig. 5.9).

The CUU exhibits a systematic zonation in geochemistry and lithology and is

divided into 5 litho-geochemical zones as shown in Figure 5.10. The core of the

CUU (Zone 4) is composed of olivine adcumulates with the highest forsterite

content (Fig. 5.9) and lowest Al2O3/TiO2 ratios. Enclosing Zone 4 are Zones 3 and

2, which outwards from Zone 4 progressively exhibit a higher abundance of an

interstitial liquid component and lower forsterite content of the olivine (Fig. 5.10).

Zone 1 occurs along the margins of the body and is characterized by a pyroxenite

lithology exhibiting lower total olivine content (Fig. 5.9). Litho-geochemical Zone 0

dominantly occurs along the western contact (Fig. 5.10), where it is characterized as

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a gabbro-troctolite. Occurrences of Zone 0 (identified on the basis of geochemistry)

also occur along the eastern contact of the Central-UU and felsic volcanic unit.

Based on geochemistry, Zone 0 is more fractionated and displays lower MgO

contents, moderate FeO (Fig. 5.9) with higher Al2O3 and TiO2 contents. Zone 0

samples plot along a mixing line between ultramafic liquid compositions observed

in the Western-UU and a felsic volcanic unit composition (Heggie, 2007: Fig. 5.13).

Figure 5.12. Drill core photos and photomicrograph of the Central Ultramafic Unit. A. Top contact between the BIF Unit and the CUU. Note the low-angle bedding in the banded iron formation (i.e. parallel to core axis) and conformable contact between CUU and BIF Unit (LJD0054A). B. Small siliceous xenolith, with felsic xeno-melt on top-left side, hosted in the CUU proximal to the footwall contact. C. Cross-polarized photomicrograph of weakly altered olivine cumulate within the CUU (LJD003A).

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Mineralization within the Maggie Hays nickel sulfide deposit comprises both

primary (massive and disseminated sulfides) and secondary mobilized sulfides.

Primary orthomagmatic mineralization occurs dominantly along the eastern contact,

as disseminated sulfide hosted in olivine cumulates and as massive sulfide along the

north-eastern felsic volcanic contact (Buck et al., 1998; Barnes, 2006: Fig. 5.3).

Mobilized secondary mineralization is present at the northern termination of the

ultramafic body, which is hosted within shear zones of the felsic volcanic unit (Joly

et al., 2008).

Western ultramafic unit

The Western ultramafic unit (WUU) represents the uppermost and youngest unit

within the Honman Formation. The WUU directly overlies the massive sulfide sub-

unit of the Sedimentary Unit as observed in drill holes LJD0011 and LJD0048 and

documented in the exploration drill logs for LJD0050 and LJD0049. The WUU

contains well-preserved pseudomorphed igneous textures. A-zone spinifex (flow top

breccia, A1, A2: Fig. 5.8E), and B-zone cumulate textures (Fig. 5.8F) characteristic

of extrusive komatiites, are observed repeatedly in successive flow units (see Arndt

et al., 2008 for a review of the igneous textures in layered komatiitic flows).

The komatiite flows observed in LJD0011 are classified as thin differentiated flow

lobes using the terminology of Barnes (2006). Komatiite flows are dominantly thin

(approximately 1 m ), with a range in thickness from 30 cm to 10 m. Flows are

differentiated and exhibit well-developed B-zone cumulate layers, and A-zone

spinifex layers with flow top breccias best defined in drill hole LJD0126. Flow top

breccias in LJD0126 are characterized by angular to subrounded, centimeter sized

fragments of both fine-grained and homogenous komatiite. Breccia fragments

containing spinifex texture are also locally observed. Hole LJD0126 exhibits a

transition from komatiite flows with spinifex tops dominating the lower stratigraphy

to flows characterized by thick (up to 40 m) flow top breccia zones (Fig. 5.8F)

higher in the stratigraphy. Flow top breccia zones throughout the WUU exhibit

minor shearing evident as flattening of fragments and a weak foliation throughout

the extrusive komatiite flow unit. A minimum thickness of 280 m is observed in drill

hole LJD0126 (Fig. 5.3), but the total thickness of the WUU is not known.

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The WUU flows are aluminum-depleted (Barberton-type) komatiites with an

average Al2O3/TiO2 ratio of 10.5 (range of 9 to 12: Fig. 5.13). Spinifex-textured and

flow top breccia samples have an average composition of 25 wt% MgO and appear

in equilibrium with olivine from Fo92 to Fo94 (Table 5.1 and Fig. 5.9). Primitive

mantle normalized trace element patterns are generally flat to slightly LREE-

depleted (La/Ybpm ~0.75: Fig. 5.11).

Figure 5.13. Bi-variant plot of TiO2 and Al2O3 for all samples from the Maggie Hays system data from this volume (Chapter 6). WUU spinifex textured samples (spfx WUU). CUU; pyroxenite lithology (Border), olivine cumulate lithology (CUU Ol), gabbroic lithology (gabbro). Felsic Volcanic Unit (felsic) with calculated averages for contaminant 1 and 2 shown. Barberton-type komatiite trend line shown for comparison with two component mixing lines between Barberton-type liquid and both potential felsic contaminants shown. Effects of olivine accumulation shown as % trapped liquid lines below the Barberton-type liquid origin.

i. Interpretation of the WUU and CUU

Physical volcanology

Komatiite volcanism is interpreted to be the result of high degree (> 20%) partial

melting of a mantle plume based on their geochemistry (Arndt et al., 2008). The

WUU flows are characterized as Barberton-type komatiites on the basis of

Al2O3/TiO2 ratio and flat to slightly depleted LREE. These characteristics of

Barberton-type komatiites indicates that garnet was stable during melting and that

melting was initiated at a pressure greater than 7 GPa (> 200 km depth) as

summarized by Arndt et al. (2008). Thin differentiated flows and sheets are

indicative of a mid to distal proximity to the volcanic vent, or moderate to low flow

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rates (Hill et al., 1995; Barnes, 2006). The presence of thick flow top breccia on

subsequent flows in drill hole LJD0126 is an uncommon but not a unique feature.

The presence of spinifex textured fragments within the flow top implies the

fragmentation or foundering of solid crust during inflation and break-out formation

(Arndt et al., 2008).

Both the BIF-hosted ultramafic and main ultramafic CUU sub-units are interpreted

as intrusive ultramafic sills post-dating the deposition of the FVU, BIF, TZU and

Sedimentary Units. The main ultramafic CUU was initially interpreted to be

intrusive (Marston, 1984) and later re-interpreted to be a fault duplication of the

WUU extrusive komatiite and associated stratigraphy (Buck et al., 1998). We here

argue that the CUU represents an intrusive body, based on: 1) the cross-cutting

relationship between the CUU and the Honman Formation stratigraphy, 2) the

concentric zonation in litho-geochemical units observed within the CUU, 3) the

presence of a chilled zone (pyroxenite) between the main ultramafic sub-unit of the

CUU and wall rocks at both upper and lower contacts, and 4) the presence of felsic

xenoliths and xenomelts along the upper and lower contacts. Although spinifex

texture is not limited to komatiite flows and cannot be used as an absolute

discriminator between intrusive and extrusive ultramafics. Spinifex textures or flow

top breccia were not observed anywhere within the CUU, whereas they were in the

WUU.

Definitive intrusive features such as continuous ultramafic apophyses from the main

body into the host rock, as observed at Mt. Keith (Rosengren et al., 2005), have not

been observed along the upper margin of the Central-UU. It is distinctly possible

that the BIF-hosted ultramafic bodies represent apophyses, but it has not been

possible to demonstrate this purely from drill core evidence.

Contamination

Both the Central-UU and Western-UU are of Barberton-type komatiite composition

on the basis of whole-rock Al2O3/TiO2 ratios (Fig. 5.13). Rare-earth element

patterns observed within the Central-UU and Western-UU are similar to each other,

yet deviate from the expected Barberton-type pattern in that both ultramafic units

exhibit predominant Th and La enrichment with negative Nb (Fig. 5.11). The

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geochemical trends in the ultramafic units are similar to those observed in the felsic

volcanic unit (Fig. 5.6) suggesting that the differences in chemical composition

between the two units may be the result of in-situ crustal contamination.

Felsic contamination of both the ultramafic units was numerically modeled by

simple two component mixing utilizing TiO2 and Al2O3, and assimilation-

fractionation-crystallization (AFC) modeling with the trace elements. All ultramafic

units within the Lake Johnston greenstone belt exhibit Th, Nb and La anomalies

deviating from “normal Barberton-type” and are interpreted to be all crustally

contaminated to some degree. Therefore, an average composition for Barberton

komatiites from published data (Blichert-Toft et al., 2004 and Chavagnac, 2004) was

utilized as a starting composition for the modeling, with a local felsic volcanic

lithology as the contaminant. Two types of felsic contaminants are identified within

the geochemical data set (Fig. 5.13). One characterized as low Al2O3 and TiO2

(~14% and 0.4% respectively: contaminant 1) and the other as high Al2O3 and TiO2

(~18% and 0.8% respectively: contaminant 2).

Simple two component mixing utilizing TiO2 and Al2O3 between a Barberton-type

komatiite liquid and a common local felsic volcanic lithology (contaminant 1) was

carried out (Fig. 5.13). Contamination of the initial liquid by 10-15% followed by

various proportions of contaminated liquid trapped in the olivine cumulates

reproduces the transitional Al2O3/TiO2 ratios discussed above (Fig. 5.13). AFC

numerical modeling carried out with the parameters of only olivine and chromite as

crystallizing phases and fractionation within the system limited to 5% prior to

recharge, reproduced the trace element patterns of the WUU with ~ 30%

contamination.

Similarly, the CUU was modeled as a product of olivine accumulation with a minor

component of trapped liquid from the initial contaminated liquid. Utilizing these

parameters the resulting trace element pattern of modeled extrusive magma does not

exactly replicate the observed patterns in the WUU due to the large negative Nb

anomaly in the WUU (Fig. 5.11). In the AFC model for the CUU, felsic volcanic

contamination greater than 30% starts to generate a similar positive Th anomaly

with equivalent MREE and HREE abundances. However, the Nb concentration in

the felsic volcanic is higher than the WUU, resulting in an incremental increase in

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the abundance of Nb in the WUU with increasing contamination. Without another

phase crystallizing that strongly partitions Nb out of the melt, it is impossible to

generate the negative Nb anomaly observed in the WUU through AFC within the

CUU, with the current contaminant and starting composition.

Although, modeling does not generate an exact match of the trace element

geochemistry of both units, similar patterns can be modeled with the WUU

representing a liquid contaminated by 10-20% of a felsic contaminant, and the CUU

being a product of olivine accumulation from the WUU magma. The CUU overall is

more primitive (median 42% MgO) than the WUU, a result of olivine accumulation,

the pyroxenite lithology of the CUU found along the intrusive contact is very similar

in major (and trace) element abundance to the spinifex textured samples from the

WUU.

Although there is no physical correlation (cross-cutting relationships) between the

WUU and the CUU the geochemical observations support a direct petrogenetic

relationship between the two units. The resulting volcanic system is characterized by

the CUU acting as a sub-volcanic feeder to the overlying extrusive WUU.

Supporting the proximal interpretation of the primitive thin differentiated flows

observed at the base of the WUU.

5.5. Discussion

Extensive drilling has provided a unique three-dimensional dataset and an

opportunity to address three aspects of Archean greenstone development: 1)

structural modification and preservation of stratigraphy from the effects of regional

strain on lithological units due to contrasting rheological properties (e.g. felsic,

ultramafic, BIF, sulfide); 2) establishment of the tectonic and depositional setting

within a 2.9 Ga basin prior to the emplacement of komatiite magmas; and 3)

stratigraphic control and the link between tectono-stratigraphic architecture of

greenstone belts and style of subsequent komatiite volcanism.

a. Structural modification

The Lake Johnston Greenstone Belt has undergone significant deformation as

identified by research carried out proximal to nickel mineralization at the Maggie

Hays and Emily Ann mines (Mason et al., 2003; Joly et al., 2008; 2010). Extensive

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secondary mobilization of massive sulfide ore bodies and deformation of adjacent

lithologies is observed in the strongly sheared felsic volcanic rocks along the eastern

contact of the CUU. Although the greenstone belt is deformed, it is interpreted that

the local stratigraphy in the Honman Formation is intact for the following reasons:

1) Primary sedimentary structures (graded bedding: as observed in the both the TZU

and Sedimentary Units) and igneous textures within the WUU (spinifex, flow top

breccias), 2) Gradational sedimentary transitions are preserved (FVU-TZU-BIF-

Sedimentary Unit), 3) Intrusive igneous contacts are observed.

The local Honman Formation stratigraphy is preserved as deformation within the

greenstone belt as a whole is not uniform, rather partitioned between the various

lithological units dependent upon the rheology. Deformed rocks, terranes and ore

deposits commonly exhibit partitioned and heterogeneous strain (Lister and

Williams, 1983; Maiden et al., 1986; Ramsay and Lisle, 2000). Numerical and

physical modeling of rock deformation (Ramsay and Graham, 1970; Ramsay and

Huber, 1987; Treagus, 1988; Treagus, 1993: Jaing, 1994; Jaing, 1994b; Jiang and

White, 1995; Goodwin and Tikoff, 2002; Tanaka et al., 2004) identifies contrasting

rheological properties, competency contrasts, boundary conditions, or boundary

discontinuities between adjacent rock masses as controlling factors on the

heterogeneous distribution of strain.

Within the Honman Formation, the FVU exhibits the highest degree of deformation,

and would have deformed plastically under amphibolite-facies metamorphism (550-

650°C: Shelton and Tullis, 1981). Moderate elongation of phenocrysts and volcanic

fragments are observed throughout the unit, with the intensity of deformation

increasing with proximity to nickel mineralization. Conversely, lithological units

occurring stratigraphically above the BIF Unit exhibit only minor deformation

textures, as is observed in the preserved spinifex textures and minor elongation of

fragments within the flow-top breccia in the WUU (Fig. 5.8F).

The CUU, does not exhibit a pervasive tectonic fabric, and is interpreted to have

undergone deformation restricted to discrete, early formed structural discontinuities.

At peak metamorphic conditions of 596-678°C ± 65°C and 5-7 kbars ± 2.1 kbars

(Joly et al., 2008; 2010), olivine is stable, either as relic grains or as neoblastic

metamorphic crystals formed by dehydration of early-formed serpentinites and less

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susceptible to deformation (Chopra and Paterson, 1984). Consequently, zones within

the CUU dominated by chlorite with minor relic and neoblastic olivine were

susceptible to deformation (Escartín et al., 1997), and partitioned the strain during

D1 and D2, prior to the growth of random anthophyllite within these zones under

prograde metamorphism conditions (Joly et al., 2008; 2010). Anthophyllite zones

within the CUU are located in the central portion of the ultramafic body and trend

parallel to the regional strike. The linear extension of these zones to the north

beyond the ultramafic unit corresponds to the highly deformed felsic volcanic rocks

and remobilized massive sulfide mineralization.

Consequently, deformation within the greenstone belt is heterogeneous varying from

high-, homogenous strain within the most incompetent lithologies, to low-strain or

discrete zones of strain within the competent lithologies, rather than uniform

homogenous deformation throughout the greenstone belt. As a result, preserved

primary contact relationships are observed between the CUU and adjacent

lithological units, and the local stratigraphy within the Honman Formation is

continuous and stratigraphically intact.

b. Tectonic setting and deposition of the Honman Formation

The Honman Formation consists of a sequence of felsic volcanics, a transition zone,

banded iron formation, sedimentary rock and komatiite. Interpretation of individual

units (lithology, volcanic and sedimentary facies, textures, and geochemistry)

identifies indicators of tectonic setting, which together constrain the geodynamic

setting of the Honman Formation.

The felsic volcanic rocks have arc-type geochemical signatures and no evidence for

an older crust component, implying juvenile crust and an active subduction zone. As

felsic volcanism wanes, a transgression from lowstand to highstand results in the

deposition of the transition zone. Following the cease of felsic volcanism and

highstand basin conditions, a period of quiescence and limited deposition gives rise

to banded iron formation deposition in a basin with restricted clastic input; which is

similar to depositional gaps observed in the Abitibi Greenstone Belt (Thurston et al.,

2008), and highstand-condensed sedimentation within the Brockman Supersequence

of the Hamersley Province (Krapež et al., 2003). Hydrothermal activity within the

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basin progressively diminishes, as evident from the decreasing Fe-content of the

accumulating sediments up stratigraphy. The top of the Honman Formation BIF Unit

is chert-rich, and similar to the chert-rich intervals observed within the Brockman

Supersequence, which mark the top of each sequence and the start of the next.

Unlike the Brockman Supersequence, where mud-rich sedimentary rocks mark the

start of a new cycle, clastic sedimentary rocks (detrital quartz-arenite) occur at the

top of the BIF sequence in the Honman Formation. The occurrence of detrital

quartz-arenite at this stratigraphic interval results from tectonic activity destabilizing

accumulated sediments along the margin of the basin and transporting them to

deeper, distal portions. Tectonic activity precedes a substantial change in sulfur

availability within the basin as the massive exhalative sulfide overlies the quartz-

arenite.

These two dramatic changes in tectonic activity and sulfur availability within the

basin precede the eruption of komatiitic lavas on the paleo-basin floor, marking the

start of magmatism sourced from a mantle plume. Arndt et al. (2008) argued that

mantle plumes can impinge upon any pre-existing tectonic setting, accounting for

the wide range of settings within which komatiites are found. The occurrence of

plume-derived komatiites at the top of an arc-volcanic sequence results from the

emplacement of a mantle plume beneath the subduction zone and shuts down the

active margin, similar to that documented within the greenstone belts of the Superior

Craton (Dostal and Mueller, 1997; Hollings and Wyman, 1999; Hollings et al.,

1999). Coincidently, the Ravensthorpe greenstone belt adjacent to Lake Johnston

greenstone belt within the Youanmi Terrane contains arc-type felsic volcanics

(average La/Ybpm of 11, excluding one data point of 226: Witt, 1999), banded iron

formation and komatiite units, supporting the initial correlation (Swager, 1997;

Barnes, 2006) and a similar tectonic setting for the two greenstone belts.

c. Stratigraphic control on emplacement of ultramafic magmas

Sub-volcanic ultramafic intrusions are well documented within the Agnew-Wiluna

Greenstone Belt, Western Australia (Rosengren et al., 2005) and in the Dundonald

Beach komatiite complex (Houlé et al., 2008) and Shaw Dome (Stone and Stone,

2000; Houlé et al., 2010), Abitibi greenstone belt, Canada, (Houlé et al., 2008).

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Examination of other greenstone stratigraphic sections that contain felsic volcanic

rocks and extrusive komatiites reveals a number of strata-bound ultramafic bodies

occurring stratigraphically beneath the komatiite lava flows as observed at Windarra,

Western Australia, with the “Corridor ultramafic” and “Inter-BIF” (Schmulian,

1984; Marston, 1984), within the Forrestania Greenstone Belt, Western Australia at

Liquid Acrobat and Cosmic Boy Ni deposits (Marston, 1984), and within the

Gweru-Midlands Greenstone Belt, Zimbabwe, as a peridotite sill within the Kwe

Kwe Felsitic Formation (Prendergast, 2001; 2003). The overall observation from

these greenstone belts is: shallow level ultramafic intrusions are more frequently

documented in settings where the substrate comprises incoherent material

(volcaniclastic/sediment). Conversely, komatiite systems with a coherent substrate

(e.g. Kambalda Dome: basalt footwall) ultramafic sub-volcanic bodies are not

identified. In both substrate settings (coherent and incoherent) the footwall

lithologies to the ultramafic magmas largely control the morphology of the resulting

volcanism (Houlé et al., 2008).

The Honman Formation sequence which was deposited prior to the WUU consists

wholly of incoherent volcaniclastics and sediments of unknown induration.

Consequently, the identification of the CUU as sub-volcanic intrusion fits the

empirical observation of intrusions associated with incoherent footwall lithologies.

However, the emplaced of the CUU at a depth of <200 m (thickness of BIF Unit)

below the paleo-basin floor where komatiites (WUU) are extruded is an unusual

setting for the development of a sub-volcanic magma chamber. Sub-volcanic

intrusions commonly occur at greater depths >1 km, or are shallow and form an

internal part of the komatiite complex (Houlé et al., 2008).

The CUU crosscuts the felsic volcanic and the TZU, but not the BIF Unit. The

modeled 3D morphology of the CUU (Figs. 5.4, 5.5) indicates the intrusion is

flattened and enlarged in size at the northern end, where it is in contact with the BIF

Unit, relative to the southern extent. The CUU intrusion morphology and

stratigraphic location imply that the BIF Unit controlled the ascent of the ultramafic

magma and the development of a concordant magma staging chamber.

The presence of concordant intrusions within sedimentary sequences has been

documented in numerous studies (Mudge, 1968; Johnson and Pollard, 1973; Hogan

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et al., 1998; Burchardt, 2008), leading to the inference of a physical control between

concordant intrusions and their distribution in thick sedimentary sequences. The

stratigraphic level of emplacement within sedimentary basins was proposed to be a

function of depth, with three contributing factors: 1) lithostatic pressure, 2) bedding

plane weaknesses and 3) fluid barrier (Mudge, 1968).

Lithostatic pressure and level of neutral buoyancy of a magma (i.e. density controls)

was further examined by Walker (1989), who concluded that magmas that are less

dense than the surrounding country rocks will ascend to higher levels until they are

either erupted on surface or “trapped” in host lithologies of equal density. However,

this model is unable to explain the abundance of dense Fe-rich tholeiitic and Mg-rich

komatiitic magmas that have erupted through less dense crust. This setting is

typified by komatiitic magmas within the Kalgoorlie-Wiluna greenstone belt that

contain older inherited zircons which are sourced from an andesitic-like crust

(Campbell and Hill, 1988). Emplacement and eruption of dense magmas through

less dense (i.e. andesitic-like) crust is thought to be a function of magma driving

pressure (Baer and Reches, 1991; Hogan et al., 1998). Magma driving pressure can

exceed lithostatic pressure, as shown in experimental modeling studies (Galland et

al., 2009), thus negating the effects of buoyancy in magma emplacement.

Although magma driving pressure can bring dense magmas to the surface, the

formation of concordant intrusions requires a significant subhorizontal strength

anisotropy. This strength anisotropy is a combination of Mudges, (1968) bedding

plane weaknesses, and fluid barriers in conjunction with mechanical obstacles

described by Hogan et al. (1998). The importance of subhorizontal strength

anisotropy in the development of concordant intrusions has been identified

numerous times in analog experimental modeling (Roman-Berdiel et al., 1995;

Galland et al., 2009). The analog experimental work indicates a subhorizontal

strength anisotropy separating an upper rigid layer from a lower weaker media acts

as a barrier to vertical propagation. This process is essential for the formation of

concordant intrusions, as shown experimentally by Kavanagh et al. (2006) and

summarized by Menand (2008). Burchard (2008), documented a field example of

this rheological setting, with mafic rocks overlying felsic rocks controlling the

emplacement of sills along the contact in Iceland.

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A similar rheological setting of a rigid lithological unit overlying a weak lithological

unit is identified within the Honman Formation. It is argued that the BIF Unit forms

the upper rigid layer, whereas the underlying TZU and FVU are weaker (Figs. 5.10

and 5.14). The BIF Unit inhibits the upward propagation of the magma, forcing the

magma to spread out laterally and inflate, until the confining pressures and strengths

are overcome.

Figure 5.14. Schematic graphic model of the emplacement of the CUU, showing the dominant role that stratigraphy plays in controlling the intrusions morphology. A. Two layer stratigraphy BIF with density of 3.2 overlying felsic volcanic with density of 2.4. Upward propagation of ultramafic magma through the felsic volcanic shown. B. Upward propagation is inhibited at the boundary between BIF and felsic volcanic, causing the lateral spreading of the ultramafic magma. C. Continual magma injection results in over-pressuring of the magma chamber (CUU) and eventual breach of the BIF occurs. Ultramafic magma progresses to the surface and develops into an extrusive komatiite flow field (WUU).

All the analog models for laccolith emplacement indicate that with continued

inflation and expansion, rupture and breaching to higher stratigraphic levels will

occur (Roman-Berdiel et al., 1995; Galland et al., 2009). It is argued that this

process occurred during the emplacement of the CUU, leading to the eruption of the

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WUU. A portion of the CUU intrusion is interpreted to have been removed by

erosion, resulting in minimum estimates of the intrusion original width and

thickness. However, if a symmetry axis is placed at the intrusions current maximum

thickness, a pre-erosional intrusion width of approximately 500-700 m is attained.

This width is similar in magnitude to that estimated by the analog models. Analog

models indicate lateral spreading of three times (3x) the overburden thickness prior

to breaching the overlying sediment (Pollard and Johnson, 1973 and Kavanagh et al.,

2006). The BIF unit overlying the CUU has a maximum thickness of approximately

200 m. With the analog model’s 3x thickness, this would equate to approximately

600 m of lateral spreading for the CUU prior to breaching of the BIF and outpouring

of ultramafic magma (WUU) at the surface.

5.6. Conclusions

The 2.9 Ga Honman Formation of the Lake Johnston Greenstone Belt contains a

conformable basin stratigraphy that records the transition from felsic volcanism

through exhalative iron-formation to the intrusion and eruption of komatiitic

magmas. Basin stratigraphy played an important role in controlling both the

temporal, spatial and the volcanic architecture of komatiite magmatism.

Felsic volcanic rocks occurring at the base of the formation are similar in trace

element geochemistry to modern subduction-related volcanism (arc-type), implying

a subduction component. It is proposed that the cessation of felsic volcanism

resulted from the arrival of a mantle plume beneath the subduction zone. Waning

felsic volcanism and the concurrent increase in hydrothermal activity within the

basin resulted in the deposition of a transitional unit, comprising silicate iron

formation intercalated with chert and felsic volcanic rocks. Felsic volcanism ceased

and homogenous oxide-facies iron formation was deposited during a basin high-

stand. Hydrothermal activity was not constant, and a resultant decrease in iron

content is observed up-sequence. A thin quartz-rich clastic sediment overlaying the

oxide-facies iron formation marks a sharp change in the sedimentation sequence in

the basin, from dominant pelagic to an interval of detrital sedimentation. Although

hydrothermal activity continued after the detrital sedimentation, the conditions were

considerably more anoxic, as preserved in the sulfide-facies iron formation

overlying the clastic sedimentary rocks. These two significant changes within the

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basin precede the eruption of komatiite lavas sourced from the mantle plume on the

paleo-basin floor.

The eruption of the extrusive komatiites to the surface was temporally delayed by

the development of a subvolcanic magma chamber (CUU) beneath the BIF. The

rheological contrast between the BIF Unit and underlying felsic volcanic unit acted

as a mechanical barrier inhibiting the ascent of the ultramafic magma. Magma that

was unable to breach the BIF Unit collected along the rheological boundary between

the two units, spread out and inflated from the progressive injection of magma (Fig.

5.14). The emplacement of the intrusion contributed to changes in both the tectonic

(clastic sedimentation) and hydrothermal activity within the area. During the

injection-inflation period, the intrusion (acting as a local heat source) drastically

changed in sulphur content in the iron formation, as observed in the change from

oxide-facies iron formation (BIF Unit) to sulfide-facies (exhalative sulfide-unit)

accumulating in the basin. Continued magma injection into the chamber caused

overpressure and breached the BIF Unit, resulting in magma erupted on the basin

floor directly on top of the sulfide iron formation.

Komatiites within the Lake Johnston Greenstone Belt record a dramatic change in

volcanism from arc-related felsic activity to emplacement of plume-related magmas

as flow-intrusion complexes. The transition is marked by a period of quiescence and

establishment of a sedimentary basin with limited detrital input, fluctuating

oxidation state and episodic development of exhalative sedimentary sulfide. This

transition gave rise to favorable substrate lithologies and an ideal tectonic setting for

formation of komatiite-hosted nickel sulfide ores. Transitions from arc volcanism to

BIF basins may be indicative of plume-arc interactions, and constitute favorable

exploration targets for komatiite associated Fe-Ni-Cu sulfide mineralization.

Acknowledgements

The greenstone belt stratigraphy described in this paper was examined as part of the AMIRA P710A project. Funding and access to sites was provided by BHP-Billiton, Independence Group and Noril’sk Nickel (formerly Lionore Australia). The authors would like to thank C. Stott and staff of the Maggie Hays Ni-Mine for their generous contribution of support and time while in the field and office.

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5.7. References Arndt, N.T., Barnes, S.J., Lesher, C.M., 2008. Komatiite: Cambridge, Cambridge University Press,

467p.

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Contents

5.1. Introduction ................................................................................................... 168 5.2. Regional Geology .......................................................................................... 170 5.3. Materials and Methods .................................................................................. 172 5.4. Stratigraphy and Geochemistry ..................................................................... 174

a. Felsic volcanic unit .................................................................................. 176 i. Interpretation of the felsic volcanic unit ................................................. 180

b. Transition zone unit ................................................................................. 181 i. Interpretation of the TZU ........................................................................ 183

c. Banded iron formation unit ..................................................................... 184 i. Interpretation of the BIF unit .................................................................. 185

d. Sedimentary unit ...................................................................................... 186 i. Interpretation of the sedimentary unit ..................................................... 187

e. Ultramafic units ....................................................................................... 187 i. Interpretation of the WUU and CUU ...................................................... 194

5.5. Discussion ...................................................................................................... 197 a. Structural modification ............................................................................ 197 b. Tectonic setting and deposition of the Honman Formation .................... 199 c. Stratigraphic control on emplacement of ultramafic magmas ................. 200

5.6. Conclusions ................................................................................................... 204 5.7. References ..................................................................................................... 206

List of Figures

Figure 5.1. Yilgarn Craton showing subdivision of the South West Terrane, Youanmi Terrane and Eastern Goldfields Superterrane. Youanmi Terrane granite-greenstone belts (dark grey) include: Lake Johnston (LJGB), Ravensthorpe (RGB), Forrestania (FGB) and Southern Cross (SCGB) greenstone belts. Eastern Goldfields Superterrane granite-greenstone belts (medium grey) include: Norseman (NGB) and Kalgoorlie (KGB). Lake Johnston Greenstone Belt nickel mines include: EA (Emily Anne deposit) and MH (Maggie Hays deposit). Modified from Department of Industry and Resources (2008). ............................................................................................ 171

Figure 5.2. Generalized stratigraphic column for the Lake Johnston Greenstone Belt; modified from Gower and Bunting (1976). * U-Pb age determinations from Wang et al. (1996). ................................................................................. 172

Figure 5.3. Geological plan map of the study area within the Lake Johnston Greenstone Belt, showing the Honman and Maggie Hays Formations. Honman Formation is subdivided into lithologic units. Strong deformation at the northern end and along basal contact of the Central-UU in proximity to remobilized Ni sulfide mineralization shown as wavy lines. All diamond drill holes examined are shown, and key drill holes referenced in the paper labeled. ......................................................................................................................... 173

Figure 5.4. Composite stratigraphic column for the Honman Formation as observed from diamond drill cores (LJD0126, LJD0048, LJD0011, LJD0054A, LJD0087A, LJD003A, LJD0039, LJD0038, LJD0049, LJD0074, LJD0055W2, LJD0092). Approximate intrusive level of the Central Ultramafic Unit and narrow intrusive sills (banded iron formation-hosted sills) shown along the left hand side. ......................................................................................................... 175

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Figure 5.5. Oblique Leapfrog® model view looking down and north-east towards the local Maggie Hays nickel-deposit stratigraphy. Stratigraphy from left to right consists of the Banded Iron Formation Unit, Transition Zone Unit, Central Ultramafic Unit and Felsic Volcanic Unit. Scale bar in metres. Western ultramafic unit not shown for clarity, but occurs to the left of the Banded Iron formation. ........................................................................................................ 176

Figure 5.6. Jensen cation plot from the Felsic Volcanic Unit and ultramafic units from the Lake Johnston Greenstone Belt: felsic volcanic rocks, Central Ultramafic Unit (CUU) pyroxenites and olivine cumulates, and Western Ultramafic Unit (WUU) komatiites. H-Fe th as (high-Fe tholeiitic andesite), H-Mg th ba (high-Mg tholeiitic basalt). .............................................................. 178

Figure 5.7. Primitive mantle-normalized trace element patterns for the Felsic Volcanic Unit shown as black lines. Data fields for TTG/TTD type (Black Flag Formation: Morris and Witt, 1997) and Arc-type felsic volcanism from Eastern Goldfields Superterrane (EGS: Morris and Witt, 1997; Messenger, 2000; Barley et al., 2008). Normalizing values from McDonough and Sun (1995). 180

Figure 5.8. Drill core photos and photomicrographs of representative Honman Formation units. A. Part of the Transition Zone (TZ) Unit from LJD0038. Felsic Volcanic Unit lithology with minor garnet on left, garnetite in middle (magnified in B.), and chert with minor sulfide on right. B. Garnetite lithology (LJD0038). C. Banded Iron Formation Unit (LJD0011). D. Iron-poor Fe-formation. E. Spinifex texture from the Western-UU (LJD0011). F. Flow top breccia texture from the Western-UU (LJD0126). G. Polarized light photomicrograph of garnetite (LJD0038) amp = amphibole, bio = biotite, grt = garnet. H. Reflected light photomicrograph of quartz-arenite (quartz with trace pyrite), exhibiting graded bedding (LJD0011). .............................................. 182

Figure 5.9. Plot of FeO versus MgO (anhydrous wt.%) for WUU (spinifex and B-zone cumulates) and CUU, subdivided into spatial zones as shown in Figure 5.10 (zone 0 = gabbroic, zone 1 = amphibolite, zone 2, 3, 4 = olivine cumulates). Modified from Barnes et al., (2004). ........................................... 188

Figure 5.10. Cross-section from line 6430470mN through the Honman Formation stratigraphy, showing stratigraphic succession (Felsic Volcanic Unit, Transition Zone Unit, BIF Unit, WUU) and conformal setting of the Central Ultramafic Unit (CUU) and smaller banded iron formation-hosted ultramafic sub-unit (BIF-hosted intrusions). Spatial geochemical zones shown in within the CUU (as used in Figs. 5.9), zone 0 = gabbroic; zone 1= pyroxenite; zone 2 = mixture of adcumulates to orthocumulates with lower forsterite olivine; zone 3 = dominant adcumulates with moderate forsterite olivine (Fo90-92); zone 4 = olivine adcumulates with highest forsterite content (Fo93-94). ..................... 190

Figure 5.11. Primitive mantle normalized trace element patterns of select samples from the CUU (blue lines), WUU (grey lines) and mean FVU (red line). Data from Chapter 6 and Appendix B. Normalizing values of Sun and McDonough (1989). ............................................................................................................. 191

Figure 5.12. Drill core photos and photomicrograph of the Central Ultramafic Unit. A. Top contact between the BIF Unit and the CUU. Note the low-angle bedding in the banded iron formation (i.e. parallel to core axis) and conformable contact between CUU and BIF Unit (LJD0054A). B. Small siliceous xenolith, with felsic xeno-melt on top-left side, hosted in the CUU proximal to the footwall contact. C. Cross-polarized photomicrograph of weakly altered olivine cumulate within the CUU (LJD003A). ........................................................... 192

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Figure 5.13. Bi-variant plot of TiO2 and Al2O3 for all samples from the Maggie Hays system data from this volume (Chapter 6). WUU spinifex textured samples (spfx WUU). CUU; pyroxenite lithology (Border), olivine cumulate lithology (CUU Ol), gabbroic lithology (gabbro). Felsic Volcanic Unit (felsic) with calculated averages for contaminant 1 and 2 shown. Barberton-type komatiite trend line shown for comparison with two component mixing lines between Barberton-type liquid and both potential felsic contaminants shown. Effects of olivine accumulation shown as % trapped liquid lines below the Barberton-type liquid origin. ........................................................................... 194

Figure 5.14. Schematic graphic model of the emplacement of the CUU, showing the dominant role that stratigraphy plays in controlling the intrusions morphology. A. Two layer stratigraphy BIF with density of 3.2 overlying felsic volcanic with density of 2.4. Upward propagation of ultramafic magma through the felsic volcanic shown. B. Upward propagation is inhibited at the boundary between BIF and felsic volcanic, causing the lateral spreading of the ultramafic magma. C. Continual magma injection results in over-pressuring of the magma chamber (CUU) and eventual breach of the BIF occurs. Ultramafic magma progresses to the surface and develops into an extrusive komatiite flow field (WUU). ............................................................................................................ 203

List of Tables

Table 5.1. Whole rock geochemistry analyses of representative units from the Honman Formation. With drill collar, sample depth, lithological unit (FVU = Felsic Volcanic Unit; WUU = Western-UU; CUU = Central-UU) and lithology (Rhy-dac = rhyolite-dacite; Spfx = spinifex; OC = olivine cumulate, Pyr = pyroxenite) in header. Trace element ratios La/Sm*, Th/Sm*, Nb/Th* and Gd/Yb* primitive mantle normalized. Normalization values from McDonough and Sun, (1995). .............................................................................................. 179

Chapter 6. PGE signatures, Maggie Hays Ni deposit.

Chapter 6. Nickel Mineralization Signatures in an Intrusive

Komatiite Sequence: Examination of the Spatial Distribution of

PGE in the Maggie Hays Ni System, Lake Johnston Greenstone

Belt, Western Australia

Abstract

Enrichment and depletion of the highly chalcophile platinum group elements,

relative to abundances expected in sulfide-undersaturated mantle-derived magmas, is

a potentially powerful exploration tool. Constraining the spatial distribution of

enrichment and depletion signatures in the context of a robust volcanology and

mineralization model makes it possible to quantify the size of nickel sulfide forming

systems, and ultimately target nickel (Ni) sulfide mineralization within komatiite

sequences.

The Maggie Hays Ni deposit within the Lake Johnston Greenstone Belt of Western

Australia is hosted within a komatiite complex consisting of both extrusive

komatiites and an ultramafic intrusive sub-volcanic feeder conduit, with

mineralization hosted in the feeder conduit. Ore formation is attributed to the

assimilation of a local sulfur rich sedimentary unit, located above the sub-volcanic

feeder. Assimilation of this unit when intersected by the sub-volcanic feeder,

induced sulfur saturation within the sub-volcanic feeder magmas. Sulfur saturation

within the system generated enriched and depleted chalcophile element ore forming

signatures. Ore forming signatures are quantified as deviations from calculated

background abundances. The spatial distributions of these signatures are examined

relative to known Ni mineralization.

Platinum group element (PGE) depletion and enrichment signatures occur at a

distance of approximately 320 m upstream from mineralization. This area is a site of

intersection between the ultramafic intrusive magma and the sulfur rich sedimentary

unit, and is interpreted to mark the point of sulfur saturation within the system. The

magnitude of PGE enrichment displays a progressive increase with proximity to

mineralization; whereas depletion signatures exhibit a more complex V-shaped

pattern attributed to progressive mixing between sulfide liquid, depleted silicate

magma and undepleted recharging magma. Ore forming signature preservation

within the system is controlled by the volcanology and timing of sulfur saturation.

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Initial magmas are sulfur undersaturated, and preserved as a marginal phase along

the intrusive contact and basal flows within the extrusive komatiites. Enrichment

and depletion signatures associated with accumulation and fractional extraction of

sulfide liquid, respectively, are preserved within the intrusive sub-volcanic feeder,

and depletion is identified within the extrusive sequence. The final stage of ore

system formation involved the influx of sulfur undersaturated magmas within the

central portion of the sub-volcanic feeder, and emplacement of stratigraphically

higher flows.

The presence of a spatial relationship between Ni mineralization and chalcophile

element depletion and enrichment signatures within the Maggie Hays system

provides the basis for the development of a PGE-based vector towards Ni ores in

komatiite systems.

Keywords: Barberton-type; Yilgarn Craton; Ni-Cu-PGE; vector; 2.9 Ga; Archean; platinum group element; chalcophile

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6.1. Introduction

Archean komatiite-hosted Fe-Ni-Cu sulfide deposits are an important source of Ni

worldwide, and were identified in the mid 1960s with the discovery of Ni

mineralization at Kambalda, Western Australia. The genetic linkage between

komatiitic rocks and Ni mineralization resulted in a subsequent boom in mineral

exploration, and the discovery of numerous outcropping Ni systems within the

Yilgarn Craton of Western Australia. However, the rate of discovery has

substantially decreased in the post-boom period (1972-present: Hronsky and

Schoddle, 2006), where most new deposits are identified by electromagnetic

methods (Peters, 2006), or through follow-up on extensions of known

mineralization.

Alternative exploration models (not geophysical-based) have mainly targeted the

physical aspects of Ni systems (e.g. komatiite volcanology, sulfur source, evidence

for contamination, and mineralization: Lesher, 1989; Barnes, 2006), with a lesser

focus on geochemical targeting (Lesher et al., 2001; Barnes et al., 2007). Limited

research has investigated the practical application of chalcophile elements (platinum

group elements [PGE]: Pt, Pd, Ru, Rh, Ir; Ni and Cu) in komatiite-hosted Fe-Ni-Cu

sulfide exploration. The chalcophile elements differ from other mineralization

indicators (e.g. rare earth elements, major elements: Lesher et al., 2001; Barnes et

al., 2007) as they are both physically and chemically linked to the ore forming

process.

During Ni ore formation, the chalcophile elements strongly partition into the sulfide

phase in the presence of an immiscible sulfide liquid (Campbell and Naldrett, 1979;

Naldrett, 1979; 1981; Naldrett and Campbell, 1982; Campbell and Barnes, 1984).

This process results in strong chalcophile element enrichment in the sulfide liquid

and chalcophile element depletion in the ore forming silicate melt. Both enrichment

and depletion signatures may be present in the komatiite system, and represent a

powerful targeting tool for regional lithogeochemical-based exploration.

Chalcophile element targeting of Fe-Ni-Cu sulfide hosted in both extrusive and

intrusive komatiite settings is challenging, as these magmatic systems have high

recharge rates with dynamic and turbulent flow. This dynamic setting produces

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rocks adjacent to mineralization that are temporally and chemically unrelated to the

ore forming process (Lesher and Keays, 1995; Lehser et al., 2001). However, the

dynamic setting of these types of Ni deposits is essential to ore formation, as the

immiscible sulfide liquid must equilibrate with a large volume of the silicate magma

(R-Factor: Campbell and Naldrett, 1979). The contained chalcophile element

abundance in the sulfide ores bodies (e.g. Ni tenor) typically indicates that the

sulfide liquid equilibrated with hundreds to thousands of times its own volume of

silicate magma (Campbell and Naldrett, 1979). It is predicted that low tenor

mineralized systems should exhibit strong chalcophile element depletion (Campbell

and Naldrett, 1979), as the deposits formed at lower silicate:sulfide ratios (R-

Factor), resulting in the most extensively depleted silicate host rocks.

Extensively chalcophile element depleted rocks, at hundreds to thousands of times

the volume of the Ni ore body, represent a powerful lithogeochemical prospectivity

tool. Quantifying the magnitude of chalcophile element depletion, constraining the

spatial distribution of the chalcophile element depleted rocks, and understanding the

spatial correlation between this signature and Ni mineralization, transforms a

prospectivity tool into a mineralization vector within a komatiite system.

The Maggie Hays Fe-Ni-Cu sulfide system is a low-tenor nickel sulfide forming

system (Barnes, 2006). As such, the Maggie Hays Ni deposit, hosted in the 2.9 Ga

Lake Johnston Greenstone Belt, within the Yilgarn Craton, of Western Australia

(Fig. 6.1), is an ideal natural laboratory to assess the spatial relationship between

chalcophile element depletion, enrichment and background abundances (ore forming

signatures) and Ni mineralization.

The local stratigraphy and ore deposit geometry of the Maggie Hays Ni system are

well-defined by extensive resource evaluation drilling (see Chapter 5). The deposit

has a number of distinctive features which contrast the more common Kambalda-

type setting (see Chapters 2, 3, 4). The Kambalda-type setting occurs within 2.7 Ga

extrusive Munro-type komatiites (Lesher and Keays, 2002), whereas the Maggie

Hays deposit is hosted within 2.9 Ga Barberton-type (aluminum depleted)

komatiites. In addition, the Maggie Hays system contains the juxtaposition of

extrusive and intrusive components, with the intrusive, sub-volcanic feeder conduit

hosting the Ni mineralization. The Maggie Hays system provides insight on the

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mechanisms, spatial distribution, and relative timing of sulfur saturation and sulfide

liquid accumulation within intrusive komatiite systems.

This research examines the ore forming process that led to the generation of the

Maggie Hays Ni deposit. Within the ore forming framework, chalcophile element

mineralization signatures are identified and quantified. The spatial relationship

between these chalcophile element mineralization signatures and Ni mineralization

is also examined in a dynamic conduit setting. The result is an enhanced

understanding of the size of ore forming systems, and the use of chalcophile

elements as lithogeochemical vectors to Ni mineralization.

6.2. Geological Setting

a. Regional stratigraphy

The Maggie Hays Ni deposit is hosted within the Archean Lake Johnston

Greenstone Belt (LJGB), located in the south eastern portion of the Youanmi

Terrane of the Archean Yilgarn Craton, Western Australia (Fig. 6.1). The LJGB is

located east of the Forrestania Greenstone Belt, northeast of the Ravensthorpe

Greenstone Belt, and to the west of the Norseman Greenstone Belt of the Eastern

Goldfields Superterrane (Fig. 6.1: Swager, 1997; Cassidy et al., 2006). The LJGB

trends NW-SE, has a strike length of approximately 100 km, and varies in width

from less than 6 km to 20 km. Age determinations of the felsic volcanic unit which

underlies the Maggie Hays deposit indicate the greenstone to be at least 2921±4 Ma

and emplacement of the ultramafic units occurred after 2903±5 Ma (Wang et al.

1996) contrasting with the 2700 Ma age of the komatiite volcanism in the Eastern

Goldfields Superterrane. The Forrestania, Ravensthorpe and Lake Johnston

greenstone belts can be correlated, based on their similar stratigraphy and the

presence of komatiites associated with banded iron formation (Swager, 1997;

Barnes, 2006; see Chapter 5).

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Figure 6.1. Southwestern region of Western Australia, with Yilgarn Craton and the three constituent subdivisions: South West Terrane, Youanmi Terrane and Eastern Goldfields Superterrane shown (Cassidy et al., 2006). Greenstone belts shown as light grey within the Eastern Goldfields Superterrane, with Kalgoorlie (K) and Norseman (N) areas labeled. Greenstone belts within Youanmi Terrane shown as dark grey, with Lake Johnston Greenstone Belt (LJGB), Southern Cross (SCGB), Forrestania (FGB), and Ravensthorpe (RGB) shown. Nickel mines Maggie Hays (MH) and Emily Ann (EA) shown.

The LJGB comprises three formations: Maggie Hays, Honman, and Glasse

Formations, as described by Gower and Bunting (1972, 1976: Fig. 6.2). Sulfide

nickel mineralization is associated exclusively with ultramafic lithologies within the

middle Honman Formation, forming the Maggie Hays and Emily Ann deposits. The

Honman Formation comprises five distinct litho-stratigraphic units that are variably

deformed and overturned, with stratigraphy dipping to the east at approximately 60°

and younging to the west (see Chapter 5). The litho-stratigraphic units from oldest

to youngest are: felsic volcanic unit, transition zone unit, banded iron formation unit,

sedimentary unit, and komatiite unit (Fig. 6.2).

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Figure 6.2. Stratigraphic sequence of the Lake Johnston Greenstone Belt. Modified from Gower and Bunting (1972; 1976); (see Chapter 5).

The oldest unit is a thick sequence of fragmental felsic volcanic lithologies. Two

samples have been dated at 2921±4 Ma and 2903±5 Ma utilizing U-Pb SHRIMP

(Wang et al., 1996). Overlying and intercalated with the felsic volcanic unit is a

laterally extensive sequence of sulfidic volcanic and sedimentary lithologies termed

the transition zone unit (TZU). The TZU represents the transition from felsic

volcanic to banded iron formation and is approximately 50-75 m thick, dominated

by iron-rich silicates, abundant garnet, thin chert units, and exhalative sulfides (both

disseminated and stringer). The top of the TZU exhibits a gradational increase in

iron (magnetite) and silica (chert) content over 5-10 m. Overlying the TZU, are well

defined alternating bands of magnetite and chert, and form an approximately 120 m

thick banded iron formation (BIF) unit. Overlying the BIF unit is a thin (< 15 m)

sequence of sedimentary units (quartz arenite sub-unit overlain by a massive

exhalative sulfide sub-unit). Extrusive komatiites termed the Western Ultramafic

Unit (WUU) are emplaced on top of the volcano-sedimentary sequence and

represent the stratigraphically youngest unit of the Honman Formation. The WUU is

219


characterized by thin flows (< 20 m), with well-developed A-zone spinifex textures

and B-zone cumulates. These cumulates have similar textures to type examples

within the Munro Township komatiites of the Abitibi Greenstone Belt, Canada

(Pyke et al., 1973; Arndt et al., 1977.

The Honman Formation volcano-stratigraphic succession is interpreted to represent

the cessation of subduction along an active margin by the emplacement of a mantle

plume (see Chapter 5). The stratigraphic succession accumulated prior to komatiite

volcanism, specifically the contact between the felsic volcanic unit and BIF unit, is

interpreted to have acted as a magma trap and controlled the emplacement of

intrusive sub-volcanic feeders to the komatiite flows (see Chapter 5). The Central

Ultramafic Unit (CUU) is a sub-volcanic intrusion and hosts the Maggie Hays Ni

deposit (Fig. 6.2).

i. Central ultramafic unit

The CUU does not outcrop and is delineated entirely through drill core intersections

and geophysical response (Fig. 6.3). It is irregular in geometry and forms a sub-

horizontal tube-like body. The CUU intrusion cross-cuts a portion of the Honman

Formation stratigraphy and decreases in size as it extends southward approximately

3.5 km from the northern end, where Ni sulfide mineralization is located in the

Maggie Hays deposit (Fig. 6.3; 6.4). The intrusion reaches maximum dimension

(300-400 m thick, > 400 m in width) in the vicinity of the Ni mineralization. The

CUU consists of olivine cumulates with volumetrically minor amphibolite and

gabbroic to felsic-pyroxenitic lithologies. Amphibolite lithologies occur as a thin (<

10 m) wide margin along the intrusion-wall rock contact (Fig. 6.4). The lithological

contact between the wall-rock and amphibolite is sharp, with preserved igneous

contacts observed locally, whereas the remainder exhibits sheared and deformed

contacts. The amphibolite locally contains small xenoliths and crystallized felsic

melts. The transition from amphibolite to olivine cumulate appears gradational but

occurs over a narrow interval (5-10 m).

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Figure 6.3. Geological plan map of the of the Maggie Hays Ni deposit stratigraphy, comprising Maggie Hays, Honman and Glasse Formations. The Honman Formation is divided into five lithological units: felsic volcanic, transition zone unit (TZU), banded iron formation (BIF unit), sedimentary unit, Western ultramafic unit (WUU), Central ultramafic unit (CUU) and Eastern ultramafic unit (EUU). Strong deformation at the northern end and along the basal contact of the CUU in proximity to mobilized Ni sulfide mineralization shown by wavy lines. Diamond drill holes examined and sampled in this study shown by the drill hole trace, and key drill holes referred to in this work are labeled with the collar identification.

Olivine cumulates internal to the amphibolite form a homogenous sequence

consisting of olivine mesocumulates to adcumulates. The igneous olivine cumulates

are replaced by metamorphic assemblages. These assemblages formed through

hydration and dehydration during prograde metamorphism, resulting in mineral

assemblages of metamorphic olivine, talc and anthophyllite. In the central portion of

the CUU, all primary igneous textures are obscured by metamorphic olivine and

zones of random anthophyllite, that formed during static prograde metamorphism

(Joly et al., 2010). Consequently, no textural variability or mineralogical layering

(olivine, pyroxene, feldspar, oxides) is documented within the olivine cumulates.

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Figure 6.4. Cross-section on line 6430610mN through the Maggie Hays deposit stratigraphy (Honman Formation: Felsic Volcanic, TZU, BIF-unit, and WUU) with crosscutting CUU. Major lithological divisions of the CUU shown. Facing direction as determined from spinifex texture within the WUU and graded bedding within the quartz arenite shown by black arrow. Two drill holes logged and sampled are labeled and shown in black (LJD0003A, LJD00011).

Within the upper portions of the intrusion, a transition is observed from olivine

dominant cumulates, through increasing pyroxene content, to gabbroic rocks with

increasing in feldspathic components (see Chapter 5). These gabbroic to felsic

pyroxenitic lithologies are significant in revealing internal differentiation, as well as

pre-deformation facing indicators. The presence of these lithologies on one side of

the Maggie Hays intrusive body is taken as corroborative evidence that the body

represents a single complete intrusion, and is not an isoclinally folded tectonic slice,

as previously hypothesized based on external geometry.

ii. Maggie Hays Ni deposit

Exploration for Ni sulfide within the LJGB was initiated in 1966. However, Ni

sulfide mineralization was not discovered at Maggie Hays until 1971. Anomalous Ni

concentrations were encountered over the next ten years of exploration, and in 1981

drilling intersected significant Fe-Ni-Cu sulfide. Due to structural complexity and

low metal price, limited exploration occurred over the next ten years. Exploration

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drilling in 1991 and 1993 intersected the main disseminated mineralization at

Maggie Hays; however, low grades and limited understanding of the deposit

geology prevented the deposit from being brought into production until 2007.

Concurrently, regional exploration identified the Emily Ann Ni deposit 5 km to the

north of Maggie Hays deposit in 1997 (Peters, 2006), leading to production

commencing in 2001 and terminating in 2007 when the identified resource was

exhausted.

The Emily Ann Ni deposit (1507 kt at 3.8% Ni: Barnes, 2006) is a highly tectonized

deposit that has been detached from its original ultramafic host, and is now almost

entirely hosted within felsic volcanics (Mason et al., 2003; Barnes, 2006). The

Maggie Hays deposit (12284 kt at 1.5% Ni) is intimately associated with the CUU

intrusive body described above. Mineralization occurs in two forms: primary

orthomagmatic mineralization, and secondary mobilized mineralization (Barnes,

2006).

Primary mineralization within the Maggie Hays deposit occurs as both massive

sulfide along the basal contact and as a large disseminated and locally matrix-

textured zone. Massive sulfides are variable in thickness but commonly less than 7

m in thickness, and are found at the northern end of the CUU along the contact with

the felsic volcanic footwall. Two zones of disseminated mineralization are observed

within the CUU. The larger zone occurs above the massive ore within the olivine

cumulates with a thickness of ~ 50 m as shown in Figure 6.5 (Maggie Hays Ni-S

ore body). The second unnamed minor zone of disseminated mineralization occurs

at the southern end of the main CUU intrusion, and lacks massive sulfide (Fig. 6.5).

Secondary mobilized mineralization is observed within the Maggie Hays deposit and

described as the North Shoot (Fig. 6.5), due to its location relative to the primary

mineralization. The North Shoot mineralization is characterized by sub-parallel

massive sulfide zones hosted in strongly sheared felsic volcanics (Joly and Miller,

2008). This mineralization is commonly less than 9 m in width and extends up to

800 m north of the CUU.

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Figure 6.5. 3D computer generated lithological model of the northern portion of the CUU (purple), with point of view from the NE looking to the SW (see Fig. 6.3). Stratigraphy dips towards the east at 60°, as shown by the Transition Zone unit. Maggie Hays and North Shoot mineralized zones shown in red (0.4% Ni grade shell).

The North Shoot mineralization is interpreted as the result of remobilization of

primary massive mineralization from the basal contact into shear zones. These shear

zones extend off of the northern end of the CUU and into the felsic host rock.

Remobilized massive sulfide is intimately associated with quartz veining and

contains numerous rounded quartz fragments, the result of continuous brecciation

and milling of syn-deformational quartz veins. Research by Joly and Miller (2008)

mapped out the mineralization and shear zones within the North Shoot and

discovered that mineralization is not continuous, but occurs as en echelon shears

zones that step laterally.

b. Metamorphism and structural modification

Metamorphic facies and structural deformation within the LJGB are variable. Upper

greenschist to amphibolite facies are present within the central portion of the

greenstone belt with a calculated peak pressure of 5-7 ± 2.1 kbars and temperatures

of 596-678 ± 65°C (Joly et al., 2008; 2010). As the greenstone belt has undergone

metamorphism, the prefix ‘meta’ has been omitted from rock descriptions for

simplification.

Four phases of deformation are identified within the LJGB (Joly et al., 2010). The

first phase of deformation (D1) is represented by NNE-SSW shortening, resulting in

the generation of large fold nappes. This was followed by elevated temperatures and

224


static prograde metamorphism to amphibolite facies, during the emplacement of

granitoid intrusions. D2 is recognized as shortening due to dextral shearing in NNW-

SSE to NW-SE direction under peak metamorphic conditions. The D3 event is

characterized by E-W shortening, apparent from the development of crenulation

cleavages. The final deformation event (D4) occurs under brittle conditions and is

characterized by steeply dipping N-NE trending dextral faults.

Although four phases of deformation are identified within the LJGB, deformation is

not evenly distributed and is partitioned into discrete zones and ductile lithologies

(see Chapter 5). Intense deformation is observed within the felsic volcanic unit with

pervasive shearing, localized mylonites and boudinage. Strong deformation is also

associated with the mobilization of massive Ni sulfide along the basal contact into

the North Shoot by dextral shearing (Joly and Miller, 2008b). Deformation within

the komatiites and sedimentary lithologies is weak to absent, or highly partitioned

into narrow shear zones. These shear zones separate blocks of undeformed rock

which preserve both igneous (spinifex and volcanic breccias) and sedimentary

textures (laminar and graded bedding: Chapter 5).


a. 3D model

A 3D lithological and structural model was generated utilizing the exploration and

resource delineation drill-hole data from the Maggie Hays system provided by

Noril’sk Nickel Pty. Ltd. (formerly LionOre Ltd.). This model was used to aid in the

understanding of the intrusion morphology, stratigraphic relationships, and the

spatial distribution of samples and geochemistry within the mineralized intrusion.

The computer generated models were created using a commercial software package,

Leapfrog®. Modeling was carried out as an iterative process, with field observations

refining the litho-stratigraphic units and spatial distribution. All litho-stratigraphic

units were modeled, thus providing a visual interpretation of cross-cutting

relationships and distribution of Ni mineralization. Block models from the

Leapfrog® model were utilized to generate successive cross-sections through the

CUU intrusion (e.g. Fig. 6.4). The CUU intrusion block model, Ni ore grade shell,

and the TZU were the main focus of modelling and are shown in Figure 6.5. The

225


computer model was also used to: 1) visualize the spatial distribution of

geochemical samples and geochemical models within the volcano-stratigraphic

framework; and 2) extract Cartesian coordinates of mineralization and geochemical

samples for the calculation of distances and azimuths between mineralization and

geochemistry samples through the Euclidian norm.

b. Sample selection

Sample collection from the Maggie Hays Fe-Ni-Cu system was undertaken on drill

core drilled between 1991 and the 2007. Sampling was conducted in two campaigns

in 2007 aided by the 3D lithological model to maximize the spatial coverage.

Exploration drilling was dominantly carried out along east-west drill sections and

provided excellent vertical and horizontal drill coverage. The initial sampling

campaign was broad and covered the drilled strike length of the Maggie Hays

system (~ 1.6 km), focusing on both the CUU and WUU. The second sampling

campaign focused on the northern mineralized portion of the Maggie Hays intrusion,

in order to infill data gaps from the initial sampling and previous research. Previous

research by Perring et al. (1994) and Perring (1995) on the CUU focused on the

mineralization and adjacent areas within 50 m of mineralization. Samples collected

for the current research focused on regions away from mineralization, with a limited

number of samples proximal to mineralization along the north-eastern margin of the

CUU. Samples within the CUU were selected to maximize the spatial coverage of

the intrusion, addressing both proximity to the host-rock contact and central portions

of the intrusion. Sampling of the WUU was limited, as the majority of drill holes

were terminated prior to reaching this stratigraphic unit. Consequently, sampling

within WUU was mainly restricted to the lowest four komatiite flows in the

stratigraphy.

A total of 47 diamond drill holes throughout the intrusion and local stratigraphy

(Fig. 6.3), with 294 samples selected for further study. Samples were selected to be

visually sulfide-free, carbonate unaltered, and distal to cross-cutting felsic intrusive

bodies. Samples were also selected to avoid the patchy, but locally advanced

anthophyllite replacement. Textural identification of lithologies within the CUU was

difficult due to prograde and retrograde metamorphic assemblages. Therefore,

general lithological core-logging was carried out in the field. The WUU exhibits less

226


metamorphic overprinting and preserves more primary textures, allowing for further

interpretation (spinifex and B-zone textures). Other stratigraphic units are visually

distinct (felsic volcanics, banded iron formation) and provide a good contrast to the

ultramafic lithologies.

Samples were split with a diamond saw and a representative slab retained for


were cleaned and cut to remove weathering effects accumulated during storage.

Samples were coarse crushed at the University of Western Australia using a jaw

crusher, which was flushed with quartz, cleaned with a wire brush, acetone and

blown dry with compressed air after each sample. Samples were subsequently

packed in clear locking plastic bags and sent to the geochemical lab for further

milling and geochemical analysis.

c. Analytical techniques

Samples up to 1 kg in size were analyzed in two batches at Ultratrace Analytical

Laboratories, Perth, Western Australia. Major elements and several trace elements

(Al2O3, CaO, Fe2O3, K2O, MgO, MnO, Na2O, P2O5, SiO2, TiO2, Cr2O3, SO3, Ni, Cu,

Ba, Rb, Sr, V, Zr,) were analyzed by wavelength dispersive X-Ray fluorescence

(XRF) on a 0.66 g sample fused to a glass bead. Minor elements (Y, Th, Nb, Hf, Ta,

La, Ce, Pr, Nd, Sm, Eu, Gd, Dy, Tb, Ho, Er, Tm, Tb, Lu, Te, Se) were analyzed by

ICP-MS following four acid (hydrofluoric, hydrochloric, perchloric, and nitric)

digestion of a 0.3 g aliquot. Platinum group elements (Au, Pt, Pd, Rh, Ru, Ir) were

analyzed by ICP-MS following a nickel sulfide fire pre-concentration, aqua regia

dissolution of the sulfide button, and co-precipitation of the PGE with tellurium

from a 25 g aliquot. Total sulfur was measured by infrared adsorption during the

combustion of the pulped sample in an oxygen-rich environment.

The precision of the analytical methods was evaluated through the use of internal

standards, blanks and duplicate analyses. Analytical precision was assessed with

duplicate analyses by the method outlined by Thompson and Howarth (1976). Major

elements exhibited median errors of <1% for the observed concentrations.

Chalcophile elements exhibited median errors of 8% Ir, 19% Ru, 13% Rh, 11% Pt,

and 7% Pd over a normal unmineralized range of abundances.

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Multiple techniques are available for PGE analysis and all are utilized in research.

These techniques include: fire assay ICP-MS (Barnes and Fiorentini, 2008; Maier et

al., 2009; Fiorentini et al., 2010; in press), Carius tube isotope dilution (Puchtel and

Humayun, 2001; Fiorentini et al., 2004), and instrumental neutron activation

analysis (Maier et al., 2004; Maier et al., 2007). Detection limits and precision vary

between the three methodologies. Current applications commonly utilize fire assay,

due to lower cost and shorter preparation time. Although the Carius tube isotope

dilution method provides better instrumental precision, duplicate analysis by fire

assay ICP-MS produces analytical results reproducible within 5% (Barnes and

Fiorentini, 2008). Additional analytical data from published and unpublished work

are also utilized in the study, as summarized and shown in Appendix A. This data

was derived from similar, but not identical analytical techniques as described in the

respective documents, therefore some discrepancies may exist. However, all

additional data were carefully assessed and only used if analytical methodologies

were equivalent or superior to the fire assay ICP-MS method.

6.4. Results

The geochemical analytical results are grouped in two ways: major and trace

element geochemistry samples are grouped based on stratigraphic location (WUU

and CUU); whereas chalcophile element distribution data are further subdivided into

sample lithology for the CUU, as these elements are sensitive to time variable sulfur

saturation events.

a. Major and trace element geochemistry

Major element, trace element and PGE analyses from collected samples and

previous research data on the Maggie Hays system is presented in Table 6.1 and

Appendix B. Major and trace element abundances from the CUU and the WUU are

presented in Figure 6.6 as binary element plots with MgO (wt%) as a fractionation

index. In conjunction with previous research a total of 205 whole-rock analyses were

from the CUU, 20 from the WUU (komatiites), and 20 characterizing the remaining

stratigraphic units. Overall, the samples from the CUU and WUU exhibit common

olivine fractionation trends. The CUU and WUU are characterized by strong

negative correlations between TiO2, Al2O3, FeOtot, and CaO with MgO (Fig. 6.6).

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Both units are also characterized by constant, and similar average Al2O3/TiO2 ratios

(WUU ~10 with σ = 0.1, CUU ~14 with σ = 3.0: Fig. 6.6). Samples from the CUU

are both chromite saturated and unsaturated, plotting along both the olivine liquid

trend and chromite saturated liquid trends as defined by Barnes (1998, 2006).

However, a large proportion of the samples plot as chromite-undersaturated on the

olivine liquid trend, with only a few samples defining the chromite-saturated and

accumulated chromite fields, which is anomalous for rocks with this high of MgO

content (Barnes and Fiorentini, in press).

Amphibolite samples from the intrusive contact of the Central-UU and spinifex

textured samples from the WUU have similar major element abundances. The

amphibolite lithologies and WUU spinifex are characterized by median values of

25.5 and 26.7 wt% MgO, 0.4 and 0.5 wt% TiO2, 4.6 and 5.0 wt% Al2O3, 10.1 and

11.9 wt% FeOtot, 0.3 and 0.4 wt% Cr2O3 and 1330 and 1309 ppm Ni, respectively

(see Table 6.1). Amphibolite samples exhibit a range of 14.3 to 28.8 wt% MgO,

whereas a range of 23.8 to 26.8 wt% MgO is measured from WUU spinifex. The

complementary B-zone cumulates in the WUU are characterized by median MgO

contents of 30.0 wt% (with a range from 24 to 31 wt% MgO), 0.3 wt% TiO2, 2.9

wt% Al2O3, 9.3 wt% FeOtot and 0.3 wt% Cr2O3.

The high field strength elements (HFSE: Th, Nb, Hf, Zr) for the CUU and WUU

exhibit negative correlations with MgO with minor scatter in the data (Fig. 6.6).

Rare earth elements (REE), consisting of light rare earths (LREE: La, Ce, Pr, Nd),

medium rare earths (MREE: Sm, Eu, Gd, Tb) and heavy rare earths (HREE: Dy, Ho,

Y, Er, Tm, Yb, Lu), uniformly exhibit negative correlations with MgO, with minor

scatter of the data for both units.

Rare earth element patterns for the amphibolite intrusive contact samples and WUU

are similar (Fig. 6.7). Both units exhibit elevated total REE abundance with LREE

and MREE enrichment over HREE, and negative Nb and positive Th anomalies. The

amphibolite lithology has generally flat patterns with minor LREE and MREE

enrichment (La/Smpmn of 1.43 and La/Ybpmn of 1.2), with Nb/Nb* anomaly of 0.26

(Fig. 6.7), (Nb/Nb* is defined as Nb/10^(Log La)+(Log La)-Log Ce). The WUU

exhibits lower LREE enrichment with similar MREE and HREE abundances

(La/Smpmn of 0.79 and La/Ybpmn of 0.67), with a smaller Nb/Nb* anomaly of 0.47

(Fig. 6.7).

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Figure 6.6. Bi-variant whole-rock geochemistry plots of major and trace elements for samples from the CUU (diamonds) and the WUU (triangles). Major elements are recalculated to anhydrous abundances. Chromite liquid trends from Barnes (2006).

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Table 6.1.Median values of major and trace elements for WUU (B-zone cumulates, Spinifex textured samples) and CUU (amphibolite and olivine cumulate) with data from Kambalda Dome Long-Victor system. (Channel B-zone, Flank B-zone, Channel Spinifex and Flank Spinifex). All data filtered S<0.25%. Trace elements and chalcophile elements in ppm unless marked * indicating ppb. WUU CUU Long-Victor (wt%) B-zone Spfx Amph Olivine C Ch Bz Flk Bz Ch Spfx Flk Spfx SiO2 50.33 47.31 48.65 44.72 44.13 44.70 46.58 46.85 TiO2 0.30 0.48 0.39 0.12 0.14 0.17 0.36 0.45 Al2O3 3.39 5.05 4.66 1.61 2.55 3.14 7.66 9.57 FeO 8.59 10.86 9.18 8.33 7.33 8.14 9.59 11.57 Fe2O3 0.94 1.22 1.04 0.37 0.22 0.39 1.11 1.43 FeO tot 9.43 11.93 10.12 8.39 7.52 8.48 10.58 12.86 MnO 0.16 0.16 0.19 0.22 0.15 0.17 0.19 0.25 MgO 26.95 26.78 25.57 42.37 42.99 39.52 26.59 18.82 CaO 9.53 7.40 9.49 1.67 0.98 1.84 7.36 7.83 Na2O 0.19 0.14 0.10 0.08 0.03 0.05 0.14 0.35 K2O 0.02 0.01 0.05 0.01 0.01 0.01 0.99 2.75 Cr2O3 0.30 0.41 0.34 0.25 0.27 0.29 0.36 0.26 P2O5 0.04 0.04 0.04 0.01 0.01 0.01 0.02 0.04 S 0.02 0.01 0.02 0.12 0.17 0.18 0.20 0.05 (ppm) Ni 1566 1309 1330 2826 2694 2302 966 516 Cu 69 70 54 22 19 33 35 24 Co n.d. n.d. n.d. 87 101 n.d. 85 n.d. Cr 2031 2789 2317 1679 1867 1989 2453 1811 Zn n.d. n.d. n.d. 90 52 n.d. 72 n.d. Ir* 2.02 2.17 1.81 2.54 4.70 2.54 0.95 0.32 Ru* 4.78 6.27 5.74 5.30 3.48 3.44 3.96 0.48 Rh* 0.93 1.49 1.17 0.79 0.68 0.62 1.20 0.19 Pt* 6.40 11.38 8.16 3.47 3.45 4.20 8.48 2.78 Pd* 4.52 7.92 5.87 2.25 4.56 4.48 8.46 2.16 Au* 0.24 9.04 8.45 3.23 1.04 (ppm) Th 0.60 0.65 0.35 0.20 0.04 0.11 Nb 0.38 0.64 0.73 0.22 0.24 0.24 0.62 0.71 La 1.49 1.34 2.62 0.81 0.24 0.37 0.74 0.93 Ce 3.80 3.71 6.91 1.86 0.70 0.92 1.94 2.32 Pr 0.62 0.64 0.86 0.23 0.11 0.16 0.34 0.41 Nd 2.89 3.19 3.76 1.19 0.62 0.88 1.98 2.35 Hf 0.21 0.32 0.56 0.16 0.29 0.29 0.56 0.69 Zr 22.11 n.d. 39.51 20.45 9.35 10.04 20.41 24.52 Sm 1.04 1.06 1.08 0.34 0.24 0.32 0.75 0.90 Eu 0.24 0.42 0.30 0.11 0.09 0.11 0.25 0.37 Gd 1.20 n.d. 1.60 0.40 0.34 0.44 0.97 1.28 Dy 1.23 1.60 1.57 0.54 0.06 0.08 0.19 0.26 Tb 0.21 0.21 0.22 0.06 0.44 0.57 1.29 1.77 Ho 0.22 0.42 0.32 0.11 0.10 0.13 0.29 0.39 Y 9.48 n.d. n.d. 7.62 4.39 4.66 9.67 10.55 Er 0.67 1.06 0.91 0.28 0.27 0.36 0.85 1.10 Tm 0.12 n.d. 0.11 n.d. 0.04 0.05 0.13 0.17 Yb 0.54 1.05 0.86 0.23 0.29 0.35 0.86 1.08 Lu 0.11 n.d. 0.11 n.d. 0.04 0.05 0.14 0.17

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Figure 6.7. Median primitive mantle normalized trace element patterns for the CUU (amphibolite samples), WUU (spinifex textured samples) and felsic volcanic rocks. Median Barberton Formation komatiites (Barberton-type komatiites) and median Silver Lake Formation komatiites from Kambalda Dome (Munro-type komatiite) shown for comparison (Chapter 4). Primitive mantle normalizing values from McDonough and Sun (1995). Barberton data from Blichert-Toft et al. (2004) and Chavagnac (2004).

The CUU olivine cumulates (excluding the amphibolite) are characterized by

median values of 42.2 wt% MgO (with a range of 2.5 to 51.1 wt% MgO), 0.1 wt%

TiO2, 1.6 wt% Al2O3, 8.4 wt% FeOtot, 0.2 wt% Cr2O3 and 2826 ppm Ni. Trace

element patterns for the CUU olivine cumulate lithologies are similar to the

amphibolite lithologies but exhibit a wide range of abundances, generally lower in

total trace element content relative to the amphibolite. Central ultramafic unit olivine

cumulate lithologies are also characterized by LREE and MREE enrichment over

HREE with a negative Nb and positive Th anomaly (La/Smpmn of 1.0 and La/Ybpnm

of 15.7), flat HREE patterns (Gd/Ybpmn of 1.46) and a negative Nb/Nb* anomaly

(0.4)(Fig. 6.7).

b. Chalcophile element geochemistry

A total of 138 samples were analyzed for chalcophile element (Ni, Cu, Pt, Pd, Rh,

Ru, Ir) abundances from the CUU and include from samples proximal to

mineralization (<1 m) to more distal (>1000 m). The samples from the CUU exhibit

a range in chalcophile element abundances that range from below detection limits

(<1 ppb) to highly elevated (>100 ppb total PGE: Fig. 6.8). Eighteen samples from

the WUU were analyzed for chalcophile element abundances and exhibit more

uniform abundances, commonly clustering on the graphs (Fig. 6.8).

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Figure 6.8. Bi-variant whole-rock geochemistry plots of chalcophile elements and sulfur from the CUU (diamonds) and WUU (squares). Samples filtered for S <1% to remove strong enrichment resulting from accumulated sulfide liquid.

Figure 6.8 presents the relationships between ore forming phases (chalcophile

elements, sulfur) and olivine fractionation for sample from the CUU and WUU. As

observed in Figure 6.8, S< 1% does not correlate with MgO and exhibits a wide

scatter at all MgO contents in both the CUU and WUU. A similar pattern is

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observed in the relationship between Pt and S, with a random distribution of the

analyses. Although Pt does not correlate with S content, there is a strong positive

correlation among all the PGE (see Fig. 6.8 Pt versus Rh and Table 6.2). This

relationship does not exist for all chalcophile element inter-relationships, as PGE

and Ni, or PGE and Cu do not exhibit a similar correlation.

Nickel exhibits a strong positive correlation with MgO (Fig. 6.8, Table 6.2).

However, a few samples deviate and exhibit either elevated or depleted Ni values.

Platinum, palladium, and rhodium display a general negative correlation with MgO

(Table 6.2). Although, numerous samples contain anomalously high or low values

resulting in extensive observed scatter relative to MgO content (Fig. 6.8).

Iridium displays a positive correlation, whereas, Ru exhibits no visual correlation

with MgO (Table 6.2). Several outlying samples occur above and below the general

observed trends between Ru, Ir and MgO (Fig. 6.8). Copper exhibits extensive

scatter with the MgO distribution and does not correlate with the other chalcophile

elements.

Table 6.2. Correlation matrix for select major elements and chalcophile elements from Maggie Hays Samples. Filtered for S <1%.

TiO2 MgO Cr2O3 S Ni Cu Ir Ru Rh Pt Pd TiO2 1 MgO -0.95 1 Cr2O3 -0.03 -0.02 1 S -0.07 0.09 0.27 1 Ni -0.87 0.89 -0.13 0.20 1 Cu 0.16 -0.25 -0.04 -0.04 -0.09 1 Ir -0.50 0.49 -0.03 0.14 0.56 0.01 1 Ru -0.32 0.29 0.28 0.19 0.33 -0.02 0.86 1 Rh -0.21 0.14 0.12 0.22 0.24 0.12 0.77 0.87 1 Pt -0.03 -0.01 0.19 0.09 0.08 0.16 0.66 0.79 0.82 1 Pd -0.12 0.07 0.18 0.24 0.21 0.20 0.64 0.77 0.76 0.90 1

Titanium normalized PGE (Pt, Pd, Rh) diagrams are used to remove the effects of

magmatic fractionation and olivine accumulation (Barnes et al., 2004; 2007;

Fiorentini et al., 2010; in press). This approach is based on the assumption that Pt,

Pd, and Rh and the lithophile incompatible trace elements are not fractionated from

each other during olivine fractionation and accumulation and will exhibit a constant

value (Fiorentini et al., 2010; in press).

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Constant PGE/Tipmn values represent the background ratios of the elements in the

absence of an ore forming signature. Titanium normalized PGE exhibit clustering of

data along constant values for Pt/Tipmn and Pd/Tipmn, and a slightly decreasing value

for Rh/Tipmn (Fig. 6.9). Central ultramafic unit samples plot both above and below

the central values with a large scatter of data observed above 40 wt% MgO,

representing the olivine cumulates of the CUU. Depleted CUU samples occur at

lower MgO contents. Samples from the WUU plot dominantly along a constant

value, with two samples occurring below.

Figure 6.9. PGE/Tipmn versus MgO (wt%) for samples from the WUU (squares) and CUU (diamonds). Dashed line of constant PGE/Tipmn are median values of low-sulfur samples of both CUU and WUU.

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6.5. Discussion

a. Whole-rock geochemistry

Major and trace element distribution within the ultramafic units (CUU and WUU) of

the LJGB exhibit negative correlations with MgO content, reflecting the fractional

crystallization of olivine within both magmatic units. Despite the high grade

metamorphism (upper amphibolite), the large ion lithophile elements (Ca, Na)

continue to display negative correlations with minor data scatter (Fig 6.6). These

negative correlations indicate metamorphism was isochemical and the metamorphic

assemblages are the product of a low fluid to rock ratio.

i. Western ultramafic unit

The WUU comprises < 10 m thick differentiated flows with a median spinifex

composition of 26 wt% MgO, reflecting the approximate composition of the

primitive magma. The komatiite B-zones exhibit olivine enrichment and plot as thin

differentiated flow lobes to channelized sheet flow facies (Barnes et al., 2004: Fig.

6.10). Olivine in equilibrium with the magma exhibits a range in composition from

Fo92-94, based on whole-rock FeO and MgO contents. These olivine compositions

are similar to those observed in the adcumulates (zone 3 and 4) from the CUU (Fo91-

93: Fig. 6.10), thus supporting a genetic link between the two ultramafic units.

ii. Central ultramafic unit

The CUU is dominated by mutually gradational olivine cumulate lithologies, flanked

by lesser marginal amphibolite and minor felsic pyroxenite. The central portion of

the CUU (zone 3 and 4: Fig. 6.10) comprises olivine cumulates with median values

of 42.2 wt% MgO, 0.1 wt% TiO2, 1.6 wt% Al2O3, 8.4 wt% FeOtot, 0.2 wt% Cr2O3

and 2800 ppm Ni. The negative correlations among TiO2, Al2O3, FeOtot and CaO

with MgO indicate a strong olivine control on the CUU geochemical trends. Whole-

rock FeO and MgO exhibit a range from 4-14 wt% and 12-50 wt%, respectively

(Fig. 6.10), and represent a mixture of cumulus olivine with variable amounts of

fractionated and locally contaminated trapped liquid, indicating significant magma

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flow-through and subsequent olivine accumulation in the magma chamber (see

Chapter 5).

Figure 6.10. Plot of FeO versus MgO (anhydrous wt.%) for WUU (spinifex and B-zone cumulates) and CUU, subdivided into spatial zones as shown in Figure 5.10 (zone 0 = gabbroic, zone 1 = amphibolite, zone 2, 3, 4 = olivine cumulates). Calculated olivine compositions (Fo) for pure olivine adcumulates are shown on the right hand side of the figure. Calculated olivine compositions (Fo) in equilibrium with magma liquid compositions are shown on left and along top of the fiugre. Modified from Barnes et al. (2004).

Deviations from normal Barberton-type komatiite geochemistry are observed within

the major and trace element data sets (Fig. 6.7). The LJGB ultramafic units exhibit

Al2O3/TiO2 values that are transitional between Barberton- and Munro-type (Fig.

6.6), with strong positive Th and La, and negative Nb anomalies in addition to

enrichment of LREE over the MREE and HREE. These deviations are attributed to

crustal contamination of the ultramafic magma through the assimilation of local

felsic volcanics (see Chapter 5). It is difficult to quantify any contribution from the

BIF unit due to the lack of any distinct trace element abundance in the BIF unit.

b. Chalcophile element abundance

Samples from the CUU have a wide range of chalcophile element abundances, even

within a restricted set of samples containing < 0.25 wt % S (Figs. 6.8; 6.9). The

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PGE correlate well with each other but poorly with Ni and Cu (Table 6.2). The lack

of correlation between Ni, Cu and the PGE is attributed to the Maggie Hays Ni-

system being a low R-factor system (Barnes, 2006).

Once a magma has reached sulfur saturation, the strong partitioning of the

chalcophile elements into the sulfide phase depletes the silicate magma in these

elements (Lesher et al., 1981; Campbell and Barnes, 1984; Barnes, 1990). If the

sulfide liquid is removed from the silicate liquid and the silicate liquid is isolated

from further interaction with the magmatic system, the resulting crystallization

products of the silicate liquid will be chalcophile element depleted. However, if a

sample contains a component of accumulated immiscible sulfide liquid, it will

exhibit chalcophile enrichment relative to a normal background abundance. A

normal background is defined as the chalcophile element abundance that would be

observed within a sample if it were representative of sulfide-free fractional

crystallization.

Fiorentini et al. (2010) used PPGE:Ti ratios (PPGE = Pt, Pd and Rh) as indices of

sulfide-free background abundances, on the grounds that these elements and Ti are

incompatible in olivine, the dominant phase involved in silicate fractionation of

komatiites. Within the Maggie Hays samples set, normalizing to Ti does not appear

to remove all the background variance and the influence of sulfide liquid is

suspected to control the observed variability. Therefore, this research derives the

expected values of background sulfide-free chalcophile element abundance as a

function of MgO. This allows for the quantification of the degree of chalcophile

element enrichment and depletion in each sample.

PPGE:Ti ratios were used to identify samples with background chalcophile element

abundances. Using Pt/Tipmn and Pd/Tipmn, the data set was iteratively filtered, thus

removing outlying samples from the median PPGE:Ti ratios. Outlying samples were

assumed to be enriched or depleted due to an ore forming process. The resultant

filtered data set was in-turn utilized to generate best-fit lines through linear

regressions. These regressions describe the predicted background abundance of

chalcophile elements (not Ti-normalized) in a sample as a function of MgO content

(Table 6.3; Fig. 6.11; Appendix D).

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Table 6.3. Equations derived and utilized to calculate background abundances of Ni, Pt, Pd, and Rh as a function of the MgO content of the samples within the Maggie Hays system.

Ni Fn(MgO) = 83.51(MgO)-823 r2 = 0.85 Pt Fn(MgO) = -0.338(MgO)+17.75 r2 = 0.70

Pd Fn(MgO) = -0.230(MgO)+12.16 r2 = 0.67

Rh Fn(MgO) = -0.0269(MgO)+1.857 r2 = 0.49

One limitation of this methodology becomes apparent in Figure 6.11, as the

modeled trends curve at low chalcophile element abundances. This is an artifact of

increasing analytical uncertainties at low element abundances (e.g. Ni at low MgO;

Ti, Pt, Pd and Rh at high MgO). Consequently, the equations under-estimate the

abundance of the chalcophile elements at high or low MgO contents dependent upon

the elements incompatibility. However, when taking into consideration the total

uncertainty (sampling, preparation, and analytical: Appendix C) of ± 500 ppm Ni, ±

2 ppb Pt and Pd, and ± 1 ppb for Rh (as shown in Figure 6.11), the equations permit

an accurate estimate of background chalcophile elements abundance that are

expected within each sample.

Figure 6.11. Plots of titanium normalized chalcophile elements versus MgO for the Maggie Hays system. Geochemical assay data plotted as grey diamonds, with equivalent calculated values shown as (+). Calculated background lines shown as solid black lines with error lines light grey (Ni ± 500 ppm; Pt, Pd ± 2 ppb; Rh ± 1 ppb).

Calculated background values for Ni, Rh, and Ru (not shown) plot along lines with a

slope determined by the olivine and chromite partition coefficient for Ni, Rh and Ru;

whereas, the incompatible chalcophile elements Pt, Pd, Cu plot as constant values

with fractionation (Fig. 6.11). Within this context, deviations from the background

are apparent with the highly chalcophile elements (PGE) and less so with Ni (Fig.

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6.11). The derived background equations also allow for the calculation of

chalcophile element abundances at any MgO content for comparison with other

komatiite systems. Assuming a liquidus MgO content of 25%, the initial chalcophile

element content for the Maggie Hays system is: 1265 ppm Ni, 46 ppm Cu, 9.3 ppb

Pt, 6.4 ppb Pd, 5.5 ppb Ru, 1.2 ppb Rh, 1.9 ppb Ir, similar to both Barberton- and

Munro-type systems, reported by Fiorentini et al. (2010).

c. Chalcophile element enrichment

Chalcophile element enrichment (above background) in a sample is the result of

sulfide liquid accumulation. Chalcophile element enrichment within individual

samples is identified as Pt/Tipmn > 0.8 and Pd/Tipmn > 1.2. These values are derived

from the background median value of each ratio, derived from the iteratively filtered

data set with a +2 ppb uncertainty. Chalcophile element enrichment is recognized in

both sulfur-bearing (S > 0.25 wt%: Fig. 6.12A) and sulfur-poor samples (S < 0.25

wt%: Fig. 6.12B). It is assumed that the sulfur bearing phases are sulfides rather

than sulfates within the current data set.

i. Sulfide-bearing samples

Within the dataset, 19 samples are identified as mineralized sulfide-bearing samples

on the basis of Pt/Ti pmn > 0.8, Pd/Ti pmn > 1.2 and sulfur > 0.25 wt%. All samples

characterized as enriched based on Pt/Tipmn and Pd/Tipmn also exhibit enrichment in

Ni, Ru, Rh and Ir relative to background abundances. Enrichment varies from 2-55

ppb for Pt, with an average enrichment of 19 ppb relative to calculated background

values. The sulfide-bearing samples have a range of sulfur content from 0.25 wt% to

17 wt%. Samples with > 2 wt% S are from previous work by Perring et al. (1994).

Sulfide-bearing mineralized samples represent olivine cumulates from the CUU and

exhibit a strong correlation between chalcophile element content and sulfur

abundance. Primitive mantle-normalized noble metal plots exhibit element

concentrations that are greater than mantle, and display a convex-up pattern,

characteristic of metal accumulation. Small negative Ir and Pt anomalies are

observed in most samples.

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Figure 6.12. All whole-rock samples that are chalcophile element enriched samples from the Maggie Hays system with S >0.25 wt% (A) and S<0.25 wt% (B). Raw data plotted as diamonds (WR data), calculated background for each sample shown as (x: Pt/Ti n calc). Ideal calculated background shown as constant solid line with ± 2 ppb error bars shown as dashed lines.

Within the data set the calculated chalcophile element abundances (Pt/Ti n calc: a

function of MgO) for nine of the samples plot above the predicated background

abundance (Fig. 6.12A), which implies either TiO2 loss or MgO gain to the sample,

resulting in the data points plotting above the background. Seven of these samples

also plot along the general TiO2 versus Al2O3 data trend (Fig. 6.6). However, all 9

samples plot below the general trend of Al2O3 versus MgO (Fig. 6.6), thus

indicating possible MgO loss from the samples, resulting in an over estimation of

the background Pt and Pd abundance for the sample. The lower MgO abundances

are either a result of alteration or differing analytical precision, as 7 of the 9 samples

are from previous work by Perring (1994; 1995).

ii. Sulfide-poor samples

Sulfide-poor high-PGE samples (Pt/Ti pmn > 0.8, Pd/Ti pmn > 1.2, and S < 0.25 wt%),

comprise a large proportion of the data set and constitute 35 samples out of 138 from

all ultramafic units within the LJGB (Fig. 6.12B). These high-PGE low-S samples

are dominantly from olivine cumulates within the CUU, with two samples from the

amphibolite border phase, and one from the felsic pyroxenite. The samples exhibit a

good correlation between Pt, Pd, and TiO2 although there is no visible correlation

between Pt, Pd and MgO or S.

All sulfide-poor samples exhibit positive Pt and Pd enrichment (Pt enriched 2 to 25

ppb) when compared with calculated values based on the MgO content. However,

six of the 35 samples exhibit minor Rh depletion. Ruthenium appears depleted in 13

samples, and 15 samples exhibit Ir depletion.

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The sulfide-poor enriched signature (elevated PGE values) could be a function of:

low TiO2 and MgO, analytical error or alteration, or sulfur loss from the samples.

When calculated background PGE abundances (as a function of MgO) for the

enriched samples is plotted on the PGE/Tipmn plots, the majority of the samples plot

along the background trend (Fig. 6.12B), indicating that the observed enriched

signature is not due to the loss of MgO or TiO2 by alteration or analytical error.

Consequently, sulfur loss from pre-existing orthomagmatic sulfide is hypothesized

to be the cause of the enriched chalcophile element signature in sulfide-poor

samples. Sulfur loss from orthomagmatic mineralization is difficult to quantify.

Crystallization modeling and mass balance research on mineralization hosted within

the Skaergaard Intrusion of Greenland identified sulfur mobility (Andersen, 2006).

Additionally, disseminated sulfides are identified as susceptible to sulfur loss

through oxidation from prograde metamorphism (Seccombe et al., 1981; Stone et al.,

2004), thus supporting the hypothesis that the enriched signatures in the sulfide-poor

samples are a relict orthomagmatic sulfide signal.

All chalcophile element enriched samples (both sulfide-poor and sulfide-bearing)

exhibit positive inter-element correlations among the PGE. Strong positive

correlations are observed between Pt, Pd and Rh, strong correlations are also

observed between Ir-Ni, Rh-Ru and Ir-Ru. Moderate correlations are observed

between Pt-(Ni, Ir, Ru), Pd-(Ni, Ir, Ru), and Rh-Ni.

d. Chalcophile element depletion

Chalcophile element depleted samples are defined as Pt/Tipmn < 0.54 and Pd/Tipmn <

0.4. Chalcophile element depletion is visually identifiable with PGE/Tipmn ratios in

14 samples from the CUU and WUU (Fig. 6.11). However, depletion is less evident

with Ni/Tipmn ratios (Fig. 6.11). The difference in signal magnitude between the

differing chalcophile elements is interpreted to be a result in the differences in

partitioning coefficients between the chalcophile and highly chalcophile elements

(Ni ~ 300 versus Pt > 10000). The difference is also attributed to a much higher

background value of Ni in the silicate component relative to the PGE.

Samples with background and depleted abundances are shown in Figure 6.13.

Background chalcophile element abundances are plotted to give spatial context to

242


the observed depletion trends. Uncertainty in the values are shown as grey bars in

some of the plots with all data points occurring within these zones plotting as

undepleted (0 value).Data are presented as both PGE/Tipmn plots and calculated

element depletion as Δ ppb or ppm. Calculated element depletion were quantified as

negative residual PGE:Ti anomalies from the background abundances as shown in

Figure 6.13B,C, & D as ppm and ppb deviations of Ni, Pt, and Pd.

Figure 6.13. Chalcophile element depleted samples. A. Pd/Tipmn versus Pd/Tipmn for all samples with background and depleted signatures. Lines at 0.63 Pt/Tipmn and 0.85 Pd/Tipmn define median background ratios. B. Calculated Pd and Pt depletion as ppb with ± 2 ppb uncertainty (grey shading) shown. C. Calculated Pt depletion as ppb with modeled lines of percent depletion (50, 75 and 100%) with ± 2 ppb uncertainty shown by grey shading. D. Calculated depletion for Ru and Pt (ppb). E. Calculated Ir depletion (ppb) versus Pt (ppb) depletion. F. Calculated Ni (ppm) depletion versus Pt (ppb) depletion.

The raw data plot of Pt/Tipmn versus Pd/Tipmn (Fig. 6.13A) displays a depletion trend

from the cluster of samples with background abundances to samples with increasing

degrees of depletion towards the lower left hand corner of the plot of. Similar

patterns are observed between both the whole-rock assay data of Pt/Tipmn and

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Pd/Tipmn (Fig. 6.13A) and the calculated negative residual anomaly for Pt and Pd

(Fig. 6.13B). Samples plotted as calculated Pt (Δ ppb) versus MgO (wt%) with

maximum calculated depletion lines (50%, 75%, 100%) indicated the majority of

samples exhibit strong depletion between 75 and 100% as shown in Figure 6.13C.

The other chalcophile elements generally exhibit the same positive correlation

relative to Pt and Pd, but with a higher degree of variability. Rhodium depletion

shows a good correlation relative to Pt and Pd depletion (not shown), but only when

the ±1 ppb uncertainty is filter removed, as the Rh budget for Maggie Hays system

at 25% MgO is 1.2 ppb. Ruthenium depletion correlates moderately well with Pt and

Pd depletion (Fig. 6.13D). Conversely, Ir which correlates well with Pt and Pd

enrichment displays a relatively constant trend with increasing Pt and Pd depletion

(Fig. 6.13E). Nickel depletion does not exhibit any correlation with Pt or Pd

depletion (Fig. 6.13F).

e. Spatial correlation of ore forming signatures

The spatial correlation between known Ni sulfide mineralization and ore forming

signatures (enrichment or depletion) is expressed in terms of a calculated average

distance between the sample and the closest three occurrences of mineralization with

grades greater than 0.4% Ni. The resulting distances range from a minimum of 1.4 m

to a maximum of 1486 m within the CUU, with Ni grades ranging from 0.43 to

2.8% Ni.

Ore forming signatures are plotted as log scaled, Ti normalized values to

accommodate both strong enrichment and strong depletion on the same graph. The

samples are sorted into three classes based on the previously described

methodologies: background, chalcophile element enriched and chalcophile element

depleted as shown in Figure 6.14.

Within the CUU intrusion, normal background values (sulfur undersaturated) occur

at distances > 320 m from mineralization; whereas enrichment and depletion ore

forming signatures and background signatures occur at distances < 320 m. Within

320 m of mineralization, enrichment and depletion exhibit differing geochemical

trends. Enrichment exhibits an increasing magnitude of enrichment with decreasing

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distance to mineralization (Fig. 6.15). The observed increasing enrichment trend

comprises both low-sulfide and sulfide-bearing samples.

Figure 6.14. Pt/Tipmn versus distance (metres) for all samples from within the CUU. Samples are classified as background, and chalcophile element enriched and depleted. The following Figure 6.15 represents samples within 350 m of mineralization.

Figure 6.15. Pt/Tipmn and Pd/Tipmn versus distance for samples within 350 m of mineralization within the CUU (close up of Fig. 6.14). Samples are classified as background, and chalcophile element enriched and depleted. Arrows show visual trends of increasing and decreasing magnitude of the chalcophile element depletion signature.

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The spatial correlation between mineralization and samples characterized by

depletion signatures displays a complex association. The depletion spatial

correlation exhibits a ‘V’-shaped pattern with proximity to mineralization (Fig.

6.15). This pattern is attributed to progressive interaction between three liquid

phases (sulfide liquid, interacting magma, and recharging magma), as the magma is

transported through the conduit prior to removal from the conduit and extrusion at

surface.

6.6. Genetic Model for Ore Formation and the Spatial Distribution of Ore Forming Signatures

Nickel mineralization hosted within the CUU intrusive conduit comprises: massive,

matrix and disseminated sulfide. Nickel mineralization hosted within the intrusive

conduit formed as a result of sulfur saturation, the development of an immiscible

sulfide liquid, and the accumulation of the immiscible sulfide. Sulfur saturation can

be reached through a number of processes, as summarized by Barnes and Maier

(1999), Naldrett (1997, 1999), and Barnes and Lightfoot (2005). However, the most

direct process for sulfur saturation is the assimilation of a sulfide bearing

contaminant causing the sulfur solubility of the magma to be exceeded. Based on

quantitative numerical modeling of major and trace element abundances, the CUU

has assimilated up to 20% volume from a felsic volcanic contaminant. However,

felsic contamination is not the sole cause of sulfur saturation. Samples from the

southern extent of CUU intrusive conduit are contaminated, yet massive sulfide

mineralization is not observed until the northern terminus of the conduit.

Examination of the felsic volcanic unit in drill core also reveals a general lack of

sulfides within the unit (<<1%). Two other lithological units, the sedimentary unit

and TZ unit (TZU) occur within the local deposit stratigraphy (Fig. 6.2) and contain

considerably more sulfur in the form of disseminated, stringer and massive sulfides

(pyrite and pyrrhotite). Both units are laterally continuous along strike and down dip

of the deposit as defined by diamond drilling (Chapter 5).

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It is hypothesized that the emplacement of the CUU was controlled by the existing

stratigraphy. The banded iron formation (BIF unit) restricted the upward propagation

of the emplacing magma (Fig. 6.16A), forcing the magma to pond beneath the iron

formation until over-pressuring and rupture of the BIF unit occurred with venting

lava forming the WUU (Fig. 6.16B). Ultramafic magma pooling beneath the BIF

unit makes the TZU stratigraphically more accessible to the CUU than the

sedimentary unit overlying the BIF-unit (Fig. 6.2). Additionally, 3D-modeling of the

TZU and the CUU identified a large area at the northern end of the CUU where the

ultramafic rocks have assimilated the TZU (Fig. 6.17), and come into contact with

the BIF unit.

Considering the lithological control of magma ponding within the CUU, Fe-Ni-Cu

mineralization is inferred to be the result of localized sulfur saturation occurring at

the northern end of the intrusion. Sulfur saturation was induced by the assimilation

of the TZU from both above the intrusion and locally at the base of the intrusion

through, thermal-mechanical erosion by magma moving through the sub-volcanic

feeder conduit (Fig. 6.16B). Once an immiscible sulfide liquid was formed, it began

to settle due to the large density contrast between sulfide and silicate magma.

Continued magma flow-through within the sub-volcanic feeder conduit intrusion

progressively transported the sulfide forward in the CUU. The sulfide settled on the

floor of the intrusion forming massive sulfide (Fig. 6.16B) and progressive

accumulation of olivine led to the development of both matrix textured ores and

disseminated sulfide zones.

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Figure 6.16. Cartoon long section of the Lake Johnston Greenstone Belt stratigraphy showing the CUU conduit system and overlying WUU. A. Emplacement model. B. Ore forming process, through assimilation of the overlying sulfur-rich contaminant, with small inset cross-section shown. C. Ore forming process with areas hosting mineralization signatures indicated. D. Final stage of the conduit system and the spatial distribution of ore forming, and background chalcophile element abundances shown.

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This ore forming model shown in Figure 6.16, infers a very localized sulfur source

with limited transport (200-400 m) of the sulfide liquid in the magma. The limited

transport distance is supported by R-factor values; where the R-factor is a measure

of the relative volumes of silicate liquid and sulfide liquid which equilibrate with

one another, as described by Campbell and Naldrett (1979):

R=CSD/(CLD-CL)

Where CS is the concentration of metal in sulfide, CL is the concentration of metal in the initial silicate liquid, D is the partition coefficient D=Dsul/sil and R is the mass ratio of silicate to sulfide liquid involved in the reaction.

R-factors calculated for the Maggie Hays mineralization (based on Ni) range from 5

to 19 with an average of approximately 7. These R-factors are relatively low

compared to other komatiite-hosted Ni deposits (Kambalda Dome 100-500:

Campbell and Barnes, 1984; Lesher and Campbell, 1993). This indicates that there

was limited interaction between sulfide liquid and silicate liquid within the Maggie

Hays system; thus supporting the hypothesis of local interaction of a large volume of

assimilated sulfide with a restricted volume of silicate magma, and limited transport

distance of the sulfide from the site of assimilation.

The CUU, acting as a feeder and magma conduit represents a simple linear tube-like

magmatic system. In comparison, complex extrusive komatiite systems typically

comprise both flank and channel facies, as recognized at Kambalda in Western

Australia (see Chapter 3 and 4: Gresham and Loftus-Hills, 1981). Consequently,

due to this apparent simplicity observed within the CUU, any changes in sulfur

saturation that occurred in the magma at a specific place within the conduit would be

recorded in the crystallization products further along the flow path. Therefore any

ore forming signatures will be contained within the conduit after the point of sulfur

saturation.

The preservation of chalcophile element signatures requires magma to become

isolated from the continued influx of new magma. Therefore, the preservation of a

mineralization signature is potentially limited within a conduit system. Similar to

channel facies in extrusive komatiite systems (Lesher and Arndt, 1995), sulfur

saturation and the development of a mineralization signature may occur early in the

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system development. Yet, the infill of the conduit may occur later, under waning

recharge conditions, and is unrelated to the ore forming event.

Assuming magma flow was unidirectional through the CUU conduit, several

features indicated that magma flow was from right to left within Figure 6.16 (SE to

NW in the field: Fig. 6.3). These features include: the presence of mineralization at

the northern terminus of the CUU; the intrusion progressively cross-cutting

stratigraphy to the north; the morphology of the CUU influenced by the BIF-unit;

the identified proximal sulfur source; and the lack of sulfide signatures in the

southern extent of the CUU.

Within this conduit model, with flow from southeast to northwest, the distribution of

depletion and enrichment signatures indicates that a distance of approximately 320

m from mineralization was the point at which the system initially attained sulfur

saturation. This distance physically corresponds to the computer modeled location of

the intersection between the CUU and the sulfur rich contaminant (TZU: Fig. 6.17).

Figure 6.17. 3D computer generated lithological model of the northern portion of the CUU with point of view from the NW looking to the SE (see Fig. 6.3) showing the areas of intersection between the CUU (purple) and the modeled TZU surface (light grey). Lithological drill intersections utilized in TZU modeling shown as black circles.

The distance of 320 m from mineralization demarks the start of mixed

mineralization signatures (enriched and depleted). Depletion of chalcophile elements

from a silicate melt is not instantaneous once sulfur saturation occurs. Rather,

chalcophile element depletion occurs gradationally as mixing between the

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immiscible sulfide liquid and the silicate liquid occurs. Chalcophile element

depletion continues until either a chemical equilibrium is established between

sulfide and silicate liquid, or the sulfide is removed from the silicate liquid

(accumulation by gravity). The observed V-shaped depletion pattern (Fig. 6.15) is

interpreted to reflect the progressive mixing of three components: 1) immiscible

sulfide melt, 2) interacting silicate melt, and 3) recharging silicate melt, resulting in

both increasing and decreasing the depletion signature. Maximum chalcophile

element depletion occurs 80-150 m further downstream than the initial site of sulfur

saturation (Fig. 6.15), where the 80-150 m distance is interpreted to represent

progressive mixing between sulfide and silicate phases. However, maximum

depletion is not observed through the remaining distance to mineralization (< 80 m);

rather, a progressive decrease in the amount of depletion is observed in this interval.

It should be noted that the depleted silicate liquid is not isolated from the system.

The conduit system has sustained flow-through of variably depleted to undepleted

silicate liquid, and the gradational decrease in the depletion signature is interpreted

to represent progressive mixing between a chalcophile element depleted magma and

a recharging un-depleted magma.

Ore forming signatures occur within the system, but are viewed as transient. The

replacement of depleted magma by an undepleted magma subsequently flushes and

dilutes the depletion signature of the combined magmas prior to crystallization

within the system, or removal from the conduit. Within the CUU, the observed ore

forming signatures (enrichment and depletion) display a spatial relationship to

mineralization and exhibit a volcanological control, as observed in the distribution

patterns of the mineralization signatures.

The narrow portion of the CUU that extends southeast from the main mineralized

body (Fig. 6.3) is characterized by normal background values. This area is thought

to have been undersaturated for most, or all of the intrusive history, as it contains

only two samples that exhibit enrichment. Normal background values are observed

within the main body and form concentric ring patterns (Fig. 6.16C, D). The

concentric ringed pattern is interpreted to represent sulfur undersaturated magmas

during both the initial emplacement of the intrusion and during final stages of

magmatic activity. Initial emplacement of the intrusion and the formation of the

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amphibolite border phase occurred under sulfur undersaturated conditions, as

observed by the preserved background chalcophile element values identified within

the amphibolite unit along the igneous contacts. Final stages of magmatic activity

are restricted to the central portion of the intrusion, by progressive infilling and

plating of the conduit from crystal accumulation during the intrusive history. This

central portion of the intrusion preserves sulfur undersaturated conditions, as

identified by samples with background values from this area (Fig. 6.16D).

Progressive infilling of the conduit also limits the duration of sulfur saturation

induced by a local contaminant. The TZU, representing the sulfur source, is located

stratigraphically above the sub-volcanic feeder and would have become isolated

from the magma with minor amounts of crystal plating along the roof of the

chamber or a decrease in recharge rate. As a result, the enriched signature is

restricted to the lower, down-stream portion of the intrusion (Fig. 6.16C, D).

Primary Ni mineralization is hosted along the basal contact of the intrusion and

extends up the west paleo-intrusion wall. The enriched signature appears to envelope

mineralization from the southern-most mineralization to the northern terminus of the

intrusion and up along the west paleo-intrusion wall (Fig. 6.16C, D).

The depletion signature makes up the smallest volume of the three signatures and

occurs in restricted areas in close proximity to the TZU (Fig. 6.16C, D). Two areas

of depletion are observed; with the largest area located at the top of the intrusion

overlying the enriched zone and partial enclosed by background values (Fig. 6.16C,

D). This zone forms the top intrusive contact with TZU. The second depleted area is

only recognized by one sample, and found along the basal contact approximately 10

m from Ni mineralization (0.78% Ni).

The presence of Ni mineralization within the sub-volcanic CUU indicates that sulfur

saturation was attained within the system for a duration of time. The interval in

which the system was sulfur saturated would arguably be recorded in the

stratigraphy of the WUU, as the CUU functioned as a sub-volcanic feeder to the

WUU. The WUU is characterized by thin komatiite flows, with limited magma

flow-through, resulting in less potential mixing and dilution of a mineralization

signature (depletion and/or enrichment) prior to crystallization. Within the WUU

stratigraphy, the lowest flows lack ore forming signatures, and indicate that the

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system was not sulfur saturated at the time of eruption, supporting the observed

distribution of mineralization signatures within the CUU (e.g. normal background

along the intrusive contacts). Within the WUU, one spinifex textured sample is

observed to have chalcophile element depletion. This sample is not from within

initial lava flow units; however, the sample location within the stratigraphy is

unknown, as the drill hole begins stratigraphically above the known marker units

(BIF unit, sediment unit). Yet, samples stratigraphically above the depleted one do

not exhibit a mineralization signature, supporting the argument that a specific

interval in the stratigraphy of the WUU records the CUU mineralization event. Even

though the spatial distribution of the chalcophile element depletion signature is not

constrained within the WUU by this one depleted sample, it is interesting to note

that this sample occurs 3400 m along strike from the Maggie Hays Ni deposit. In

summary the identification of depletion within the WUU supports the link between

the CUU and WUU units. It also presents the possibility of the WUU to host

transported Ni mineralization (yet unidentified), arguably at a higher tenor than the

Maggie Hays deposit.

6.7. Conclusions

The Maggie Hays Ni deposit is hosted in a 2.9 Ga Barberton-type komatiite complex

consisting of both intrusive and extrusive units (see Chapter 5 for further

description). Although metamorphism and structural deformation are documented in

the stratigraphic sequence (Joly et al., 2010), a cohesive mineralization model, that

is consistent with a geochemical architecture largely controlled by primary

volcanological and magmatic parameters, is generated for the CUU. This

mineralization model is also consistent with the previously proposed emplacement

model (Chapter 5; Fig. 6.16). Mineralization is hosted within an intrusive sub-

volcanic feeder conduit (CUU), and is the result of assimilation of sulfur rich

contaminant. Two local sulfur sources are identified in the mine stratigraphy and

consist of the Transition Zone Unit (TZU) and a massive exhalative sulfide sub-unit

within the sedimentary unit (e.g. stratigraphically above the BIF unit; Fig. 6.16; see

Chapter 5). 3D computer-generated models of the stratigraphically lower TZU and

the CUU reveals probable areas of intersection between the units and zones where

the assimilation of the TZU has occurred. The effects of interaction between the

253


TZU and the CUU are observed in two forms. Firstly, the effects of interaction are

observed in three drill holes in the form of disseminated and stringer sulfide

occurring in the ultramafic unit proximal to the TZU contacts. Secondly, the TZU is

stratigraphically thinned or absent in the areas of interaction. Both observations

support the hypothesis that sulfur saturation occurred at the top of the CUU

intrusion, a result of interaction between the sub-volcanic feeder and the overlying

sulfur rich sediments (TZU).

A mineralization model that has sulfur contamination occurring above the feeder and

conduit system provides a point source for sulfur saturation, and a unique

opportunity to examine the spatial distribution of chalcophile element mineralization

signatures within a conduit system (CUU) that has a direct extrusive (WUU)

component. Chalcophile element ore forming signatures occur as depletion and

enrichment signatures. Both mineralization signatures are identified as deviations

from a calculated background condition. Chalcophile element depletion, enrichment

and background signatures are observed within the CUU.

Chalcophile element depletion characterizes approximately 10% of the samples

within the dataset (14 out of 138), and is observed in Pt, Pd, Rh, and Ru.

Chalcophile element depletion is most recognizable with Pt and Pd, as these

elements occur in the highest abundances and are strongly incompatible with

olivine. Rhodium and Ru correlate well with Pt and Pd depletion, although these

elements exhibit slight compatibility with olivine, and occur at abundances nearing

the analytical detection limits. The remaining chalcophile elements (Ni, Cu, Ir)

correlate poorly with depletion observed in the other PGE. This poor correlation is a

result of several factors, including; higher partition coefficients into olivine and

lower partition coefficients into the sulfide phase for Ni, hydrothermal element

mobility for Cu, and temperature dependent saturation phases for Ir (Barnes and

Fiorentini, 2008). These factors act to decouple the respective chalcophile element

from orthomagmatic mineralization and the other PGE.

The spatial distribution of mineralization signatures within the Maggie Hays system

is consistent with the emplacement model (Fig. 6.16) indicating magma flow from

the southeast to the northwest. Magma moving through the system interacted

extensively with the felsic volcanics, as observed in both the major elements and

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trace element patterns of both the CUU and WUU. Despite interacting with and

assimilating the felsic volcanics, the initially emplaced magma was sulfur

undersaturated. This magma did not attain sulfur saturation until the assimilation of

the overlying sulfur-rich TZU, at a point marginally upstream of the current

mineralization at the northern terminus of the CUU intrusion. The point of sulfur

saturation within the intrusion is identified where both depletion and enrichment

signatures occur concurrently, approximately 320 m upstream from mineralization.

Depletion and enrichment signatures are not observed before this distance in the

southeasterm part of the intrusion.

Chalcophile element depletion is limited in spatial distribution, and is controlled by

the volcanological setting and magma flow dynamics (velocity, turbulence,

viscosity, volume). At the time of sulfur saturation within the CUU, the preservation

of this chalcophile element depletion signature appears restricted to an area at the

top of the intrusion, down-stream from the point of sulfur saturation. Enrichment

signatures are more commonly observed within the intrusion and are restricted to the

lower portion of the intrusion. This is a combined effect of magma flow and density

separation of the denser immiscible sulfides to the base and lower portions of the

intrusion.

Ore forming signatures within the CUU display two different patterns with

proximity to Ni mineralization. Enrichment signatures exhibit an increasing degree

of enrichment with proximity to mineralization. Conversely, depletion signatures

exhibit a V-shaped pattern with proximity (Fig. 6.15). The pattern of progressive

increase in depletion signature to a maximum followed by progressive lessening of

the signature with proximity to mineralization, is attributed to mixing between three

liquid phases: 1) sulfide liquid extracting the chalcophile elements; 2) silicate liquid

undergoing depletion; and 3) undepleted recharging magma. The increasing

depletion signature is a result of liquids 1 and 2 continuing to interact; whereas the

decreasing depletion signature is a result dilution by fluids 2 and 3 mixing, prior to

removal of the resultant magma from the CUU. This magma removal subsequently

contributes to the developing WUU extrusive flow field.

Stratigraphically, the WUU reflects these three ore forming processes within the

CUU. The WUU basal flows do not exhibit any enrichment or depletion signature

255


(e.g. background), but higher in the stratigraphy chalcophile element depletion is

observed, which is overlain once again by undepleted magmas. This pattern of

background-depleted-background supports the proposed mineralization model in

which sulfur saturation is of a limited duration, occurring part way through the

evolution of the system.

The Maggie Hays system consists of the intrusive CUU and the extrusive WUU,

where preserved chalcophile element signatures occur within the sub-volcanic

intrusive feeder conduit (CUU) and stratigraphically within the extrusive komatiites

(WUU). The spatial distribution of the chalcophile element signatures associated

with Ni mineralization within the CUU provides valuable insight into both

magmatic and ore forming processes. Chalcophile element signatures are a viable

way to constrain Ni ore forming systems. Chalcophile element ore forming

signatures identify a point source of sulfur saturation, the development of an

immiscible sulfide liquid, and the sulfide and silicate liquid mixing. Additionally,

chalcophile element ore forming signatures constrain the site of sulfide deposition

and accumulation, and indicate post ore formation magmatic activity within the

conduit with the removal of the depleted magma. Quantifying the magnitude of

chalcophile element ore forming signatures and constraining the spatial distribution

of the signatures within dynamic volcanic settings establishes the functionality of

chalcophile element based vectors.

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Index

6.1. Introduction ................................................................................................. 215 6.2. Geological Setting ....................................................................................... 217

a. Regional stratigraphy ................................................................................... 217 i. Central ultramafic unit ............................................................................ 220 ii. Maggie Hays Ni deposit .......................................................................... 222

b. Metamorphism and structural modification ................................................ 224 6.3. Materials and Methods ................................................................................ 225

a. 3D model ..................................................................................................... 225 b. Sample selection .......................................................................................... 226 c. Analytical techniques .................................................................................. 227

6.4. Results ......................................................................................................... 228 a. Major and trace element geochemistry ........................................................ 228 b. Chalcophile element geochemistry .............................................................. 232

6.5. Discussion .................................................................................................... 236 a. Whole-rock geochemistry ........................................................................... 236

i. Western ultramafic unit ........................................................................... 236 ii. Central ultramafic unit ............................................................................ 236

b. Chalcophile element abundance .................................................................. 237 c. Chalcophile element enrichment ................................................................. 240

i. Sulfide-bearing samples .......................................................................... 240 ii. Sulfide-poor samples ............................................................................... 241

d. Chalcophile element depletion .................................................................... 242 e. Spatial correlation of ore forming signatures .............................................. 244

6.6. Genetic Model for Ore Formation and the Spatial Distribution of Ore Forming Signatures ................................................................................................. 246 6.7. Conclusions ................................................................................................. 253 6.8. References ................................................................................................... 257

List of Figures

Figure 6.1. Southwestern region of Western Australia, with Yilgarn Craton and the three constituent subdivisions: South West Terrane, Youanmi Terrane and Eastern Goldfields Superterrane shown (Cassidy et al., 2006). Greenstone belts shown as light grey within the Eastern Goldfields Superterrane, with Kalgoorlie (K) and Norseman (N) areas labeled. Greenstone belts within Youanmi Terrane shown as dark grey, with Lake Johnston Greenstone Belt (LJGB), Southern Cross (SCGB), Forrestania (FGB), and Ravensthorpe (RGB) shown. Nickel mines Maggie Hays (MH) and Emily Ann (EA) shown. ................................ 218

Figure 6.2. Stratigraphic sequence of the Lake Johnston Greenstone Belt. Modified from Gower and Bunting (1972; 1976); (see Chapter 5). .............................. 219

Figure 6.3. Geological plan map of the of the Maggie Hays Ni deposit stratigraphy, comprising Maggie Hays, Honman and Glasse Formations. The Honman Formation is divided into five lithological units: felsic volcanic, transition zone unit (TZU), banded iron formation (BIF unit), sedimentary unit, Western ultramafic unit (WUU), Central ultramafic unit (CUU) and Eastern ultramafic unit (EUU). Strong deformation at the northern end and along the basal contact of the CUU in proximity to mobilized Ni sulfide mineralization shown by wavy

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lines. Diamond drill holes examined and sampled in this study shown by the drill hole trace, and key drill holes referred to in this work are labeled with the collar identification. ........................................................................................ 221

Figure 6.4. Cross-section on line 6430610mN through the Maggie Hays deposit stratigraphy (Honman Formation: Felsic Volcanic, TZU, BIF-unit, and WUU) with crosscutting CUU. Major lithological divisions of the CUU shown. Facing direction as determined from spinifex texture within the WUU and graded bedding within the quartz arenite shown by black arrow. Two drill holes logged and sampled are labeled and shown in black (LJD0003A, LJD00011). ......... 222

Figure 6.5. 3D computer generated lithological model of the northern portion of the CUU (purple), with point of view from the NE looking to the SW (see Fig. 6.3). Stratigraphy dips towards the east at 60°, as shown by the Transition Zone unit. Maggie Hays and North Shoot mineralized zones shown in red (0.4% Ni grade shell). ..................................................................................................... 224

Figure 6.6. Bi-variant whole-rock geochemistry plots of major and trace elements for samples from the CUU (diamonds) and the WUU (triangles). Major elements are recalculated to anhydrous abundances. Chromite liquid trends from Barnes (2006). ........................................................................................ 230

Figure 6.7. Median primitive mantle normalized trace element patterns for the CUU (amphibolite samples), WUU (spinifex textured samples) and felsic volcanic rocks. Median Barberton Formation komatiites (Barberton-type komatiites) and median Silver Lake Formation komatiites from Kambalda Dome (Munro-type komatiite) shown for comparison (Chapter 4). Primitive mantle normalizing values from McDonough and Sun (1995). Barberton data from Blichert-Toft et al. (2004) and Chavagnac (2004). ................................................................... 232

Figure 6.8. Bi-variant whole-rock geochemistry plots of chalcophile elements and sulfur from the CUU (diamonds) and WUU (squares). Samples filtered for S <1% to remove strong enrichment resulting from accumulated sulfide liquid.......................................................................................................................... 233

Figure 6.9. PGE/Tipmn versus MgO (wt%) for samples from the WUU (squares) and CUU (diamonds). Dashed line of constant PGE/Tipmn are median values of low-sulfur samples of both CUU and WUU. ......................................................... 235

Figure 6.10. Plot of FeO versus MgO (anhydrous wt.%) for WUU (spinifex and B-zone cumulates) and CUU, subdivided into spatial zones as shown in Figure 5.10 (zone 0 = gabbroic, zone 1 = amphibolite, zone 2, 3, 4 = olivine cumulates). Calculated olivine compositions (Fo) for pure olivine adcumulates are shown on the right hand side of the figure. Calculated olivine compositions (Fo) in equilibrium with magma liquid compositions are shown on left and along top of the fiugre. Modified from Barnes et al. (2004). .......................... 237

Figure 6.11. Plots of titanium normalized chalcophile elements versus MgO for the Maggie Hays system. Geochemical assay data plotted as grey diamonds, with equivalent calculated values shown as (+). Calculated background lines shown as solid black lines with error lines light grey (Ni ± 500 ppm; Pt, Pd ± 2 ppb; Rh ± 1 ppb). .................................................................................................... 239

Figure 6.12. All whole-rock samples that are chalcophile element enriched samples from the Maggie Hays system with S >0.25 wt% (A) and S<0.25 wt% (B). Raw data plotted as diamonds (WR data), calculated background for each sample shown as (x: Pt/Ti n calc). Ideal calculated background shown as constant solid line with ± 2 ppb error bars shown as dashed lines. ....................................... 241

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263

Figure 6.13. Chalcophile element depleted samples. A. Pd/Tipmn versus Pd/Tipmn for all samples with background and depleted signatures. Lines at 0.63 Pt/Tipmn and 0.85 Pd/Tipmn define median background ratios. B. Calculated Pd and Pt depletion as ppb with ± 2 ppb uncertainty (grey shading) shown. C. Calculated Pt depletion as ppb with modeled lines of percent depletion (50, 75 and 100%) with ± 2 ppb uncertainty shown by grey shading. D. Calculated depletion for Ru and Pt (ppb). E. Calculated Ir depletion (ppb) versus Pt (ppb) depletion. F. Calculated Ni (ppm) depletion versus Pt (ppb) depletion. .............................. 243

Figure 6.14. Pt/Tipmn versus distance (metres) for all samples from within the CUU. Samples are classified as background, and chalcophile element enriched and depleted. The following Figure 6.15 represents samples within 350 m of mineralization. ................................................................................................. 245

Figure 6.15. Pt/Tipmn and Pd/Tipmn versus distance for samples within 350 m of mineralization within the CUU (close up of Fig. 6.14). Samples are classified as background, and chalcophile element enriched and depleted. Arrows show visual trends of increasing and decreasing magnitude of the chalcophile element depletion signature. ............................................................................ 245

Figure 6.16. Cartoon long section of the Lake Johnston Greenstone Belt stratigraphy showing the CUU conduit system and overlying WUU. A. Emplacement model. B. Ore forming process, through assimilation of the overlying sulfur-rich contaminant, with small inset cross-section shown. C. Ore forming process with areas hosting mineralization signatures indicated. D. Final stage of the conduit system and the spatial distribution of ore forming, and background chalcophile element abundances shown. ......................................................... 248

Figure 6.17. 3D computer generated lithological model of the northern portion of the CUU with point of view from the NW looking to the SE (see Fig. 6.3) showing the areas of intersection between the CUU (purple) and the modeled TZU surface (light grey). Lithological drill intersections utilized in TZU modeling shown as black circles. .................................................................... 250

List of Tables

Table 6.1. Median values of major and trace elements for WUU (B-zone cumulates, Spinifex textured samples) and CUU (amphibolite and olivine cumulate) with data from Kambalda Dome Long-Victor system. (Channel B-zone, Flank B-zone, Channel Spinifex and Flank Spinifex). All data filtered S<0.25%. Trace elements and chalcophile elements in ppm unless marked * indicating ppb. . 231

Table 6.2. Correlation matrix for select major elements and chalcophile elements from Maggie Hays Samples. Filtered for S <1%. ........................................... 234

Table 6.3. Equations derived and utilized to calculate background abundances of Ni, Pt, Pd, and Rh as a function of the MgO content of the samples within the Maggie Hays system. ....................................................................................... 239

Chapter 7. Finland and Norway Ni Prospectivity

Chapter 7. Application of Lithogeochemical Prospectivity for Komatiite-Hosted Nickel Sulfide Mineralization, Northern Finland and Norway.

Submitted as: Heggie, G.J., Fiorentini, M.L., Barnes, S.J., and Barley, M.E., Application of lithogeochemical prospectivity for komatiite-hosted nickel sulfide mineralization, northern Finland and Norway, Economic Geology (in revision).

Abstract

The application of major and chalcophile elements (platinum group elements [PGE],

Ni, Cu) as key indicators of Fe-Ni-Cu sulfide prospectivity in komatiites is tested in

the poorly exposed, geologically complex terranes of the central Karelian Craton of

northern Finland and Norway. Major element abundances are indicative of volcanic

processes, allowing for the detection of prospective volcanic facies. Conversely, the

highly chalcophile PGE are intimately associated with Fe-Ni-Cu sulfide

mineralization and record mineralization signatures in the form of PGE depletion

and enrichment. Mineralization signatures are identified by Pt/Alpmn (primitive

mantle normalized) and Pd/Alpmn ratios, removing effects of low-pressure olivine

crystallization. PGE-based mineralization signatures (Ni prospectivity indicators)

have been characterized from mineralized Munro- and Barberton-type komatiites,

and are tested on ultramafic units of the central Karelian Craton.

Lithogeochemistry identifies Paleoproterozoic Karasjok-type (high Fe-Ti)

komatiites and picritic rocks in the Karasjok and Pulju Greenstone Belts, and

Archean Munro-type komatiites in the Enontekiö area. Major element komatiite

flow facies characterization identifies higher prospectivity channelized sheet flows

and ponded lava lakes in the Pulju Greenstone Belt and Enontekiö area, which are

known to host Fe-Ni-Cu sulfide mineralization. The Karasjok Greenstone Belt is

characterized by low prospectivity thin flow facies. All three areas contain PGE-

based mineralization signatures of enrichment and depletion, indicating a high Ni

prospectivity for the magmatic systems to host Fe-Ni-Cu sulfide mineralization; thus

supporting the application of major element and PGE lithogeochemistry as an

effective Ni prospectivity indicator in terranes with complex deformation histories,

limited outcrop, alteration and metamorphism.

Keywords: komatiite, platinum group elements, chalcophile elements, Karasjok-

type.

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7.1. Introduction

Komatiite-hosted nickel (Ni) deposits continue to provide an exploration challenge,

due to small target size and absence of an alteration halo. Currently, the discovery

rate of new komatiite-hosted Ni deposits is decreasing, where the remaining deposits

represent more challenging targets under cover and at greater depths (Hronsky and

Schodde, 2006). Advances in targeting techniques have led to increased discovery

success, specifically geophysics (magnetic and electromagnetic: Peters, 2006) and

the use of lithogeochemistry.

Lithogeochemical targeting has the capacity to increase the target size beyond that

of physical mineralization. Two areas have shown potential as lithogeochemical

targeting tools: 1) major elements are used to identify prospective volcanic facies

(Barnes et al., 2004; 2007); and 2) platinum group elements (PGE) along with Ni

and copper (Cu), are used in the form of mineralization indictors, due to their

chalcophile nature and intimate association with the ore forming process (Keays,

1982; Barnes et al., 1985; Lesher et al., 2001; Fiorentini et al., 2010; Chapters 4

and 6).

a. Volcanic facies

Komatiite-hosted Ni deposits are typically associated with areas of sustained magma

flow (Lesher et al., 1984; Lesher and Keays, 2002; Barnes, 2006a; b; Barnes et al.,

2004; 2007; Arndt et al., 2008). Within the Yilgarn Craton of Western Australia, Fe-

Ni-Cu sulfide mineralization is hosted within narrow linear units of thick olivine

cumulates (>30 m) that grade laterally into thin (<30 m) well-differentiated

komatiite flows (Lesher et al., 1984; Barnes 2006a, b). These volcanic komatiite

facies are the product of sustained magma flow in a channelized environment based

on the accumulation of extensive olivine, with periodic breakouts into the flank

environment generating thinner well-differentiated flows (Hill et al., 1995; Barnes,

2006a). Nickel prospectivity based on volcanology involves the physical

identification of thickened olivine cumulates, and is consequently dependent upon

exposure or geophysical interpretation. Major element geochemistry can provide an

indirect assessment of the volcanology, if limited physical identification is possible

(e.g. poor outcrop). This assessment is useful, as channelized environments display

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major element geochemical indicators of sustained flow (e.g. high MgO: Barnes,

2006a; Barnes et al., 2004; 2007).

b. Mineralization indicators

Mineralization indicators are based on several chalcophile elements: Ni, Cu, and the

PGE (platinum [Pt], palladium [Pd], iridium [Ir], rhodium [Rh], ruthenium [Rh]).

The chalcophile nature of the PGE, Ni and Cu generates predicable and recognizable

mineralization signatures in ore forming systems that attain sulfur (S) saturation and

segregate an immiscible sulfide liquid (Barnes et al., 1985; Lesher et al., 2001;

Fiorentini et al., 2010). Two PGE mineralization indicators are identified:

enrichment and depletion, representing the positive and negative residual anomalies

from a background baseline. This background baseline represents the concentrations

of chalcophile elements in a sample if it had crystallized without sulfide

accumulation or removal. Baselines are a product of initial magma composition,

subsequent fractionation and crystal accumulation. Background baseline abundances

and observed enrichment and depletion ranges of the mineralization signatures are

identified for both Barberton- and Munro-type komatiites hosting nickel sulfide

mineralization (Fiorentini et al., in press b; Chapters 4 and 6). Although Barberton-

and Munro-type komatiites differ in geochemistry, a function of petrogenetic history

(Arndt et al., 2008), they display similar mineralization signatures and background

baselines (Chapters 4 and 6). These established baselines provide guidelines for the

interpretation of PGE whole-rock geochemistry to assess the prospectivity of

komatiites to host Fe-Ni-Cu sulfide mineralization.

In this study, major and chalcophile (PGE, Ni, Cu) element lithogeochemical Ni

prospectivity indicators are applied to ultramafic units within the central Karelian

Craton of northern Finland and Norway, to assess the application of Ni prospectivity

indicators in terranes with complex tectonic histories, limited exposure and differing

komatiite types.

c. Test area

The central Karelian Craton of northern Finland and Norway contains both Archean

and Paleoproterozoic komatiites (Fig. 7.1). The identified Archean komatiites are

Munro-type (Slabunov et al., 2006), whereas Paleoproterozoic units are Karasjok-

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type (high Fe-Ti) to picritic in composition (Barnes and Often, 1990; Hanski et al.,

2001). Nickel sulfide mineralization is associated with both age groups and

geochemical types (Kuriki and Papunen, 1985; Saltikoff et al., 2006; Makkonen et

al., 2009), thus providing geochemical and age diversity to the field test of

lithogeochemical prospectivity indicators in areas of limited outcrop, but with

identified Fe-Ni-Cu sulfide mineralization.

Although diversity in komatiite type and age are observed, the ore forming process

is common. This study assesses the relevance of applying Munro- and Barberton-

type background abundances to Karasjok-type komatiites. The chalcophile element

budgets are similar between all three komatiite types (Fiorentini et al., in press b); as

such the resulting chalcophile element signatures can be applied to Karasjok-type

komatiites. However, the differing petrogenetic history of the Karasjok-type

komatiites (Barnes and Often, 1990; Barley et al., 2000; Hanski et al., 2001), may

influence the resulting background and ranges of observable mineralization

signatures. The Ni prospectivity of select komatiite units from the central Karelian

Craton are assessed using: 1) PGE mineralization indicators (Fiorentini et al., 2010;

Chapters 4 and 6); 2) physical volcanology obtained from outcrop interpretation;

and 3) major element concentrations to assist in the interpretation of volcanology.

This research presents new PGE and major element geochemical data on Karasjok-,

and Munro-type komatiites from the central Karelian Craton of northern Finland and

Norway, and discusses the application of PGE-based Ni prospectivity indicators in

terranes with complex tectonic histories, variable alteration and metamorphism, and

limited outcrop.

7.2. Regional Setting

a. Central Karelian Craton

The Baltic or Fennoscandian Shield of Sweden, Finland, Norway and northwestern

Russia (Fig. 7.1) comprises three major crustal domains. From SW-NE the domains

include: the Paleoproterozoic Svecofennian Province exposed in the SW of Sweden

and Finland, the Archean Karelian Craton occupying northeastern Finland and

western Russia, and the Kola-Lapland Province covering the Kola Peninsula and

northern-most Finland and Norway (Fig. 7.1 inset).

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During the Paleoproterozoic, the Archean Karelian Craton acted as both a stable

continental platform forming basement to the Paleoproterozoic 2.5-1.9 Ga Central

Lapland Greenstone Belt along the northeastern margin, and as a core for subsequent

accrectionary tectonics (Lehtonen et al., 1998; Hanski et al., 2001; Slabunov et al.,

2006). Accrectionary tectonism commenced at 1.9 Ga with the accretion of the

Kola-Lapland Province along the northern margin of the Karelian Craton, followed

by the Paleoproterozoic Svecofennian Province between 1.97-1.86 Ga along the

southwestern margin (Gaál and Gorbatschev, 1987; Weihed et al., 2005).

Figure 7.1. Map of northern Sweden, Norway, Finland and northwestern Russia showing the distribution of the Paleoproterozoic Central Lapland Greenstone Belt (green), and associated komatiite and picritic rocks (black). Sampling areas are delineated by boxes comprising the: Archean Enontekiö Area, and Paleoproterozoic Pulju and Karasjok Greenstone Belts. Inset map of Norway, Sweden and Finland showing major tectonic divisions of the Baltic Shield. Modified from Hanski et al. (2001).

The central Karelian Craton comprises lithological units as old as 3.1 Ga, but is

dominated by younger 2.9-2.7 Ga granitoids and gneissic domains that intrude

greenstone belts of similar age (Lobach-Zhuchenko et al., 1993; Vaasjoki et al.,

1993; Slabunov et al., 2006). The central Karelian Craton also forms basement to

younger Paleoproterozoic (2.0-1.9 Ga) greenstone sequences of the Central Lapland

265


Greenstone Belt (Fig. 7.1). More than 20 Archean and Paleoproterozoic greenstone

belts are recognized within the central Karelian Craton, with belts exhibiting a linear

alignment of NNW-NW and N-S. These greenstone belts are characterized by

multiple deformation events, which resulted in greenschist facies in the central

portions, grading into amphibolite facies along the margins of the belts. Komatiite

lithological units of the central Karelian Craton are identified within the older

Archean 2.9-2.7 Ga remnant greenstone fragments and metamorphic complexes

associated with supracrustal sedimentary and volcanic units (Papunen et al., 1977;

Kröner et al., 1981; Slabunov et al., 2006), and within the Paleoproterozoic Central

Lapland Greenstone Belt (Fig. 7.1: Hanski et al., 2001).

i. Archean komatiites (2.9-2.7 Ga)

Komatiites occurring within Archean greenstone fragments of the central Karelian

Craton exhibit diverse litho-stratigraphic associations, ranging from komatiite-

tholeiitic, basalt-calc-alkaline volcanic rocks and sedimentary sequences, to the

more dominant komatiite with intercalated felsic volcanic rocks, basalt, tuff and

graphitic schist (Papunen et al., 1977; Kayryak and Morozov, 1985; Slabunov et al.,

2006). Currently, documented komatiites are restricted to the Munro-type (Slabunov

et al., 2006). Nickel mineralization is identified within komatiites of the Sumozero-

Kenozero and Kuhmo-Suomussalmi-Tipasjärvi Greenstone Belts, and ultramafic

(amphibolite) units within the Lieksa Complex and Enontekiö area (Papunen et al.,

1977; Saltikoff et al., 2006; Slabunov et al., 2006; Makkonen et al., 2009).

ii. Paleoproterozoic komatiites (2.0-1.9 Ga)

Paleoproterozoic komatiites are hosted within the Central Lapland Greenstone Belt

(Fig. 7.1). The belt extends from northern Norway into central Finland, and

comprises three sections: the Karasjok Greenstone Belt in the north (Norway), the

Kittilä Greenstone Belt to the south (Finland), and the Pulju Greenstone Belt

occurring in between (Fig. 7.1). These belts can be extended and correlated with the

adjacent Vetreny Greenstone Belt in Russia, based on similar stratigraphic position,

lithology, and geochemistry (Hanski et al., 2001).

The volcanic-sedimentary succession documented in the Central Lapland

Greenstone Belt is interpreted to represent rifting of the Archean Karelian Craton

266


(Lehtonen et al., 1998). Rift-associated ultramafic volcanism and sedimentation

within intracratonic basins occurred for approximately 500 Ma, until compression

and inversion of the basins and rift system occurred at ca. 1.9-1.8 Ga, a result of

continent-continent collision during the Svecokarelian Orogeny (Lehtonen et al.,

1998; Braathen and Davidson, 2000; Weihed et al., 2005).

The volcanic-sedimentary succession observed in the Paleoproterozoic greenstone

belts of the Central Lapland Greenstone Belt is variable from north to south. The

stratigraphy is best documented within the Kittilä Greenstone Belt (Lehtonen et al.,

1998); whereas, differing komatiite lithological units are identified within the rift

sequences of the Karasjok and Pulju Greenstone Belts (Fig. 7.2: Papunen, 1998;

Braathen and Davidson, 2000). The main stratigraphy of the Kittilä Greenstone Belt

is subdivided into the upper Kumpu formation and lower Lapponian schists, which

are separated by an unconformity (Fig. 7.2). The Kumpu formation is dominated by

quartzite and conglomerate with lesser felsic volcanic rocks. The Lapponian schists

comprise sedimentary and volcanic rocks, and are further subdivided into five

lithostratigraphic groups: Salla, Onkamo, Sodankylä, Sovukoski, and Kittilä Groups

(Fig. 7.2: Räsänen et al., 1995; Lehtonen et al., 1998; Hanski et al., 2001).

Within all of the greenstone belt rift sequences of the Central Lapland Greenstone

Belt, komatiite to komatiite-picrite units are identified within the lower portions of

the stratigraphy (Fig. 7.2). Within the Kittilä Greenstone Belt, two geochemical

units are identified within the lower Lappoinian Schists (lower and upper: Hanski et

al., 2001). The lower ultramafic geochemical subdivision within the Onkamo Group

comprises a komatiite-tholeiite sequence (approximately 250 m thick), which

erupted upon both older intermediate-felsic volcanics of the Salla Group and

Archean basement (Lehton et al., 1998). The upper geochemical ultramafic unit

extruded upon deeper water sediments of the Savukoski Group (Hanski et al., 2001),

and comprises Karasjok-type komatiites and picrites (Hanski et al., 2001; Barnes

and Often, 1990). Komatiitic units within the Kittilä Greenstone Belt are

characterized by high MgO contents, variable light rare-earth element enrichment or

depletion, heavy rare-earth depletion and middle rare-earth and high field strength

element enrichment (Hanski et al., 2001).

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Figure 7.2. Paleoproterozoic stratigraphic sequences and correlations within the Central Lapland Greenstone Belt, comprising the Karasjok, Pulju and Kittilä Greenstone Belts; with arrows indicating formations sampled within the Karasjok and Pulju belts. Formations and Groups are identified with characteristic lithologies summarized: mf. vol. = mafic volcanic, amp. = amphibolite, vol. clast. = volcaniclastic, kom. = komatiite, psam. = psammite, thole. vol. = tholeiitic volcanic, cong. = conglomerate, fels. vol. = felsic volcanic, suf. sed. = sulfidic sediment, qutz. = quartzite, BIF = banded iron formation. Complied from Braathen and Davidson (2000); Papunen (1998); Lehtonen et al. (1998). Age determinations from Pihiaja and Manninen (1988), Hanski et al. (1997).

Within the Kittilä and correlative Karasjok Greenstone Belts, extrusive ultramafic

units are characterized by volcaniclastic textures (agglomerates to tuffs) associated

with massive and pillowed flows (Saverikko, 1985; Barnes and Often, 1990;

Gangopadhyay et al., 2006). Nickel mineralization, in the form of low-grade

disseminated sulfides, is identified at a number of prospects in the Central Lapland

Greenstone Belt, with the two most significant including: Hotinvaara (1.3 Mt at 0.4

wt% Ni) and Iso-Siettelöjoki (0.5 Mt at 0.29 wt% Ni), both within the Pulju

Greenstone Belt (Saltikoff et al., 2006; Makkonen et al., 2009).

7.3. Sampling and Physical Volcanology

Minimal geological relationships are visible in the field, due to the limited outcrop

exposure, and poor lateral and stratigraphic continuity at all sites. Consequently,

mapping and geochemical sampling were conducted on a coarse scale. More

intensive mapping and interpretation are available in the literature for several

locations within the Karasjok, Pulju, and Kittilä Greenstone Belts (Papunen, 1998;

Lehtonen et al., 1998; Barnes and Often, 1990; Braathen and Davidsen, 2000;

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Hanski et al., 2001; Gangopadhyay et al., 2006), and within the Enontekiö area

(Papunen et al., 1977).

Sampling was undertaken by the author at three locations with outcropping

ultramafic units in the central Karelian Craton (Fig. 7.1). Archean komatiite units

were sampled within the Enontekiö area (Fig. 7.1: Papunen et al., 1977).

Paleoproterozoic komatiitic units were sampled within the Pulju Greenstone Belt

and Karasjok Greenstone Belt (Figs. 7.1, 7.2). Samples comprise both A-zone

(spinifex/flow top breccia) and B-zones (cumulate), as defined by Pyke et al. (1973).

The samples contain no visible sulfides or primary igneous mineralogy, as alteration

of primary olivine/pyroxene mineralogy was pervasive, comprising secondary

greenschist facies (serpentine, chlorite, anthophyllite, tremolite, actinolite and talc).

a. Archean komatiites (Enontekiö area)

Archean komatiitic lithologies from the Enontekiö area (Sarvisoaivi) are associated

with amphibolites, felsic to intermediate volcanic rocks, banded iron formation and

sulfidic sediments (Papunen et al., 1977; Saltikoff et al., 2006). The komatiite

samples are characterized as thin, differentiated flows and massive cumulate units

(samples #59 to 73; Table A7.1). Within the Enontekiö area, two zones of Fe-Ni-Cu

sulfide mineralization are recognized: Ruossakero (5.5 Mt at 0.53% Ni) and

Sarvisoaivi (0.7 Mt at 0.40% Ni: Papunen et al., 1977; Saltikoff et al., 2006;

Slabunov et al., 2006; Makkonen et al., 2009). All Enontekiö samples are from the

Sarvisoaivi area.

b. Paleoproterozoic komatiites (Pulju and Karasjok Greenstone Belts)

Paleoproterozoic komatiitic lithologies in the Pulju Greenstone Belt (Nilivaara and

Hotinvaara areas: Figs. 7.1, 7.2) are part of the upper komatiite group, described as a

komatiite-picrite association by Hanski et al. (2001). The sampled komatiite units

are associated with metapelites and sillimanite schists of the Mertavaara Formation,

which overlies the quartzites of the Sietkuoja Formation (Fig. 7.2). Samples were

taken from different units comprising: thin differentiated flows (<3 m),

volcaniclastic textured units, flow units with visible fragmental flow top textures,

and massive cumulate units of unconstrained thickness (samples #44 to 56; Table

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A7.1). Nickel mineralization is identified within the Hotinvaara sample area (1.3 Mt

at 0.43% Ni: Papunen, 1998; Saltikoff et al., 2006; Makkonen et al., 2009).

Sampled komatiite lithologies from the Karasjok Greenstone Belt (samples #75 to

94; Table A7.1) are from within the Briittagielas Formation (Fig. 7.2). These

komatiites are characterized by thin and pillowed flows, with abundant fragmental

and volcaniclastic textured units (Barnes and Often, 1990). Ultramafic lithologies

are intercalated with mafic volcanics and sedimentary lithologies (slate), with cross-

cutting gabbroic units. Within the Karasjok Greenstone Belt there is no known Fe-

Ni-Cu sulfide mineralization.


Samples were split with a diamond saw and a representative slab was retained for


were cleaned and cut to remove visible weathering effects. Samples were coarse

crushed at the University of Western Australia using a jaw crusher, which was

flushed with quartz, cleaned with a wire brush, acetone and blown dry with

compressed air after each sample. Samples were packaged in clear locking plastic

bags and sent to Ultra Trace Analytical Laboratories in Perth, Western Australia for

further milling and geochemical analysis. Major and select trace elements (Al2O3,

CaO, Fe2O3, K2O, MgO, MnO, Na2O, P2O5, SiO2, TiO2, Cr2O3, SO3, Ni, Cu) were

analyzed by wavelength dispersive X-Ray fluorescence (XRF) on 0.66 gram

samples, each fused to a glass bead. Platinum group elements (Pt, Pd, Rh, Ru, Ir)

were analyzed by ICP-MS following a nickel-sulfide fire pre-concentration, Aqua

Regia dissolution of the sulfide button, and co-precipitation of the PGE with

tellurium from a 25 gram sample. Total sulfur was measured by infrared adsorption

during the combustion of the sample in an oxygen-rich environment.

The precision and accuracy of the analytical methods was evaluated through the use

of internal standards, blanks and duplicate analyses. Analytical precision was

assessed with duplicate analyses by the method of Thompson and Howarth (1976).

Major elements exhibit median errors of <1% for measured concentrations.

Chalcophile elements exhibit median errors of 17% Ir, 29% Ru, 16% Rh, 18% Pt,

13% Pd, 1% Ni and 21% for Cu over a normal unmineralized range of abundances.

Duplicate analysis of all samples was carried out for select major elements utilizing

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ICP-OES (inductively coupled plasma-optical emission spectrometry).

Concentration of major and minor elements between the original and duplicate

samples exhibit median variations of 2% TiO2, 1.5% Al2O3, 1.4% MgO, and 2.8%

Ni.

7.5. Whole-Rock Geochemistry Results

Whole-rock geochemical results for Archean komatiites from the Enontekiö area

and Paleoproterozoic komatiites from the Karasjok and Pulju Greenstone Belts are

shown in Table A7.1.

a. Archean komatiites (Enontekiö area) Major elements from the komatiitic units of the Enontekiö area (Sarvisoaivi) exhibit

a range of compositions reflecting olivine accumulation. Thin flow units are

characterized by median compositions of 17 wt% MgO, 12 wt% FeOtot, 0.5 wt%

TiO2, and 12 wt% Al2O3, with massive units exhibiting a maximum MgO content of

48 wt%. Negative correlations are observed between: MgO and TiO2, Al2O3, with

positive correlations between MgO and Cr, Ni (Fig. 7.3). Al2O3/TiO2 ratios are

variable between the komatiite units, with a median value of 29.

Figure 7.3. Bivariant plots of major elements for the ultramafic units from the three areas within the central Karelian Craton, as determined by XRF and ICP-MS. Komatiites from the Archean Enontekiö area (Sarvisoaivi), and Paleoproterozoic areas: Karasjok (Karasjok Greenstone Belt), and Nilivaara and Hotinvaara (Pulju Greenstone Belt).

271


Chalcophile element concentrations exhibit a range from <1 ppb to a strong

enrichment of 30 times primitive mantle (Table A7.1). Nickel exhibits a strong

positive correlation with MgO, whereas Cu does not exhibit any correlation. The

platinum group elements (PGEs: Pt, Pd, Ir, Ru, Rh) exhibit poor negative correlation

with MgO content (Fig. 7.4). Conversely, Ir and Ru exhibit positive correlations

with MgO. The PGE exhibit moderate positive inter-element correlations, with Ir

exhibiting the poorest positive relationship. Additionally, the PGE correlate well Ni,

and all the chalcophile elements moderately correlate with sulfur.

b. Paleoproterozoic komatiites (Karasjok and Pulju Greenstone Belts)

Komatiitic rocks from the Karasjok Greenstone Belt exhibit a range of MgO

contents from 7 to 30 wt%. The rocks that approximate liquid compositions (e.g.

thin flows, pillowed flows and volcaniclastic rocks), are characterized by median

values of 20 wt% MgO, 11 wt% FeOtot, 0.9 wt% TiO2, and 9.6 wt % Al2O3 (Table

A7.1). Titanium oxide and Al2O3 exhibit negative correlations with MgO, with TiO2

exhibiting more scatter (Fig. 7.3). The komatiitic rocks are characterized by a

subchondritic Al2O3/TiO2 ratio of 13.

Ultramafic rocks from the Pulju Greenstone Belt (Nilivaara and Hotinvaara areas)

exhibit a range of MgO contents from 13 to 43 wt%. Thin flows, pillowed flows and

volcaniclastic rocks are characterized by median values of 23 wt% MgO, 11 wt%

FeOtot, 0.7 wt% TiO2, and 7.6 wt% Al2O3 (Table A7.1). Similarly, TiO2 and Al2O3

exhibit negative correlations with MgO (Fig. 7.3); however, the sampled lithological

units have a chondritic Al2O3/TiO2 ratio of 23.

Within the Karasjok and Pulju Greenstone Belts, major element distributions display

a strong olivine control, with positive correlations observed between Ni and MgO,

and negative correlations between MgO and TiO2 and Al2O3. Chromium abundances

plot along the olivine-chromium equilibrium line for units approximating liquid

compositions (thin and pillowed flows, fragmental textured units and volcaniclastic

units); whereas the massive units plot as olivine-chromite cumulates, as described by

Barnes (1998). Komatiitic rocks in both greenstone belts exhibit elevated TiO2

contents (Karasjok komatiites: 0.9 wt%, and Pulju komatiites: 0.7 wt%) at a given

MgO content, relative to Munro- and Barberton-type komatiite compositions

272


(estimated 0.45 and 0.25 wt% TiO2, respectively for Karasjok and Pulju), as

described by Barnes and Often (1990) and Hanski et al. (2001).

Figure 7.4. Bivariant plots of chalcophile and major elements for the ultramafic units from the three areas within the central Karelian Craton, as determined by fire-assay ICP-MS. Komatiites from the Archean Enontekiö area (Sarvisoaivi) and Paleoproterozoic areas: Karasjok (Karasjok Greenstone Belt), and Nilivaara and Hotinvaara (Pulju Greenstone Belt).

Chalcophile element (Ni, Cu, Ir, Ru, Rh, Pt and Pd) abundances within the sampled

units are variable, ranging from below analytical detection limits (<1 ppb) to

enrichment of 3 to 5 times primitive mantle (Table A7.1). Nickel exhibits a strong

positive correlation with MgO, whereas Cu generally displays a negative

relationship with moderate scatter in the data. Iridium and Ru exhibit positive

correlations with MgO, whereas Rh, Pt and Pd do not show any apparent correlation

with MgO content (Fig. 7.4).

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The PGE (Pt, Pd, Ir, Ru, Rh) exhibit moderate inter-element correlations with

positive linear relationships. The PGE also exhibit varying correlations with Ni;

where negative correlations are observed in the Nilivaara samples, and positive

correlations in the Karasjok samples (Fig. 7.4). The two remaining areas

(Sarvisoaivi and Hotinvaara) exhibit no correlation. Sulfur does not correlate with

any of the PGE (Fig. 7.4).

7.6. Lithogeochemical Prospectivity Indicators

The best Ni prospectivity indicator is the presence of Fe-Ni-Cu sulfide

mineralization (Barnes et al., 2004). In the absence of Fe-Ni-Cu sulfide

mineralization other methodologies are required as prospectivity indicators. The

major and chalcophile element abundances obtained from whole-rock geochemistry

provide indicators of: 1) petrogenetic classification and initial chalcophile element

content of the magma; 2) volcanic facies; and 3) presence of chalcophile element

mineralization signatures. The following discussion covers the significance of these

three indicators to Ni prospectivity.

a. Petrogenetic classification and initial chalcophile content

A number of petrogenetic processes lead to the formation of Munro-, Barberton-,

and Karasjok-type komatiites (Arndt et al., 2008). In terms of Ni prospectivity, the

significance of the petrogenetic process leading to the formation of komatiites is

unconstrained (Fiorentini et al., in press b). However, the empirical observation is

that the majority of global komatiite Fe-Ni-Cu sulfide deposits are hosted within

Munro-type komatiites, with a small fraction of identified deposits occurring within

Barberton-type, and only a few known mineralization occurrences hosted within

Karasjok-type komatiites. Consequently, based on statistics of known deposits,

Munro-type are the most prospective, followed by Barberton- and Karasjok-type

komatiites.

Samples collected from the Archean Enontekiö area (Sarvisoaivi) are komatiitic,

with interpreted quenched textured units having MgO contents >18 wt%. The

Al2O3/TiO2 ratios are variable, as shown in an ultramafic discrimination diagram of

[Al2O3] versus [TiO2] ([ ] denotes mole proportions: Hanski, 1992: Fig. 7.5). On the

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basis of the discrimination diagram, the Enontekiö area komatiites range from

Munro-type compositions (Al-undepleted komatiites) to Barberton-type (Al-

depleted).

Figure 7.5. [Al2O3] versus [TiO2] high-MgO volcanic discrimination diagram of Hanski et al. (2001). Where [Al2O3] and [TiO2] are normalized mole proportions using the equations [Al2O3] = Al2O3/(2/3-MgO-FeO) and [TiO2] = TiO2/(2/3-MgO-FeO): (see Hanksi, 1992).

Samples from the Paleoproterozoic ultramafic units within the Karasjok Greenstone

Belt (Briittagielas Formation) and the Pulju Greenstone Belt (Mertavaara Formation)

exhibit a range of rock types (Fig. 7.5). Ultramafic units from the Karasjok

Greenstone Belt exhibit a range of liquid compositions from <18 to 26 wt% MgO,

with a median value of 20 wt% MgO, that characterize them as komatiites. Despite

the samples having a subchondritic Al2O3/TiO2 ratio (13), the samples mainly plot

as Al-undepleted and exhibit a range from normal Ti-abundances to Ti-enriched and

picrites (Fig. 7.5). This disparity between whole-rock subchondritic Al2O3/TiO2

ratios and Al-undepleted signatures was also observed in the Kittilä Greenstone Belt,

and is attributed to excess TiO2 (Hanksi et al., 2001). Resultantly, ultramafic

samples from the Briittagielas Formation are interpreted as Al-undepleted Karasjok-

type komatiites and picrites. This result is similar to that reported previously for the

formation by Barnes and Often (1990), and is similar to the ultramafic units within

the Sodankyla Group of the Kittilä Greenstone Belt (Fig. 7.2: Lehtonen et al., 1998;

Hanski et al., 2001).

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Samples from the Pulju Greenstone Belt (Mertavaara Formation) are characterized

by a narrow range of liquid compositions with a median value of 23 wt% MgO,

indicating they are komatiites. Although the majority of the samples have chondritic

Al2O3/TiO2 ratios, they plot within the Al-depleted field and exhibit both normal to

enriched TiO2 abundances. Resultantly, the ultramafic units sampled within the

Pulju Greenstone Belt are interpreted as Al-depleted Karasjok-type komatiites. This

petrogenetic classification is similar to previous analyses by Papunen (1998), who

identified the ultramafic rocks as Al-depleted.

Despite being correlated within the Central Lapland Greenstone Belt (Fig. 7.2:

Braathen and Davidsen, 2000; Papunen, 1998; Lehtonen et al., 1998; Hanski et al.,

2001), the Briittagielas and Mertavaara Formations (Karasjok and Pulju Greenstone

Belts, respectively) exhibit differing geochemistry between ultramafic units.

Titanium-enrichment is observed within the komatiitic units of both belts, and is

characteristic of Karasjok-type komatiites (Barnes and Often, 1990; Barley et al.,

2000; Hanksi et al., 2001). However, the Al-content and range of liquid MgO-

content differs, with the Karasjok ultramafic units exhibiting a range of liquidus

compositions and have Al-undepleted compositions; whereas the komatiite rocks of

Pulju exhibit a narrow range of liquidus compositions and are Al-depleted.

The cause of this geochemical difference between the Karasjok and Pulju

Greenstone Belts is unconstrained and beyond the scope of this work, but may be

related to the petrogenetic history and residual phases in the melt source area, as

observed with Barberton-type komatiites and residual garnet (Arndt et al., 2008).

Although, there is a potential petrogenetic difference between the komatiitic units

within the Pulju and Karasjok Greenstone Belts, it has not substantially affected the

chalcophile element budget. In both greenstone belts there is no evidence for

complete chalcophile element depletion, which would be indicative of residual

sulfide in the source area (Fiorentini et al., in press b). Similarly, the Munro-type

komatiites from the Enontekiö area exhibit a range of chalcophile element

abundances, and consequently also left the source area sulfur undersaturated.

Despite contrasting petrogenetic histories between komatiites from the three areas

(Enontekiö, Pulju, Karasjok) the chalcophile element contents indicate that all are

equally prospective hosts for Fe-Ni-Cu sulfide mineralization.

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b. Volcanic facies

Volcanological studies on ultramafic units associated with Fe-Ni-Cu sulfide

mineralization have identified sustained magma flow-through within lava channels

or conduits as a critical component for mineralization (Lesher et al., 1984; Lesher

and Keays, 2002; Barnes, 2006a, b; Barnes et al., 2004; 2007; Arndt et al., 2008).

Favorable volcanic environments for mineralization are recognized by the presence

of thickened (>30 m) linear olivine mesocumulate to adcumulate bodies, interpreted

to represent long-lived magma conduits within the larger developing flow field

(Lesher et al., 1984; Hill et al., 1995; Hill, 2001). The presence of a high proportion

of olivine meso- and adcumulates within komatiite sequences has also been

advanced as the critical feature of the highly mineralized eastern Yilgarn Craton in

Western Australia, as opposed to otherwise similar but much less mineralized

terranes, such as the Abitibi Greenstone terrane in Ontario, Canada (Barnes et al.,

2004; 2007).

Surficial volcanology within the Enontekiö area (Sarvisoaivi) comprises both thin

and thickened flows, with observed orthocumulate and mesocumulate bodies of at

least 5 m thickness. Previous diamond drilling in the area indicates the presence of

thickened olivine cumulate bodies (Papunen et al., 1977). These observations are

supported by the whole-rock geochemistry, as apparent with MgO contents > 40

wt% (Table A7.1). Samples plotted within volcanic facies differentiation fields

defined by Barnes (2006a: Fig. 7.6) verify the thin flows identified in the Enontekiö

field area. However, Figure 7.6 indicates the majority of thickened olivine cumulate

bodies sampled are channelized sheet flows to layered sills and lava lakes, rather

than the more prospective dunite bodies. Only one sample is characterized as dunite

(Fig. 7.6).

The volcanology of the Nilivaara and Hotinvaara areas within the Pulju Greenstone

Belt comprises thin flows and thickened (>5 m) olivine cumulate units. Exploration

diamond drilling in the Hotinvaara area identified dunitic units in excess of 100 m

thickness (Papunen, 1998). The abundance of major elements also reflects the

presence of olivine cumulates (MgO > 40 wt%). The volcanic facies plots in Figure

7.6 verify the thin flow classification and estimate olivine (Fo92-93) to be in

equilibrium with the initial magma. The remaining data plots below the defined

277


fields, possibly due to FeO loss. If FeO loss is the cause for plotting low on the

graph, the volcanological setting may have ranged from channelized sheet flow to

layered sills, and lava lakes to dunitic units.

Figure 7.6. FeO wt% versus MgO wt% recalculated to volatile free for ultramafic samples from Central Karelian Craton. Olivine compositions in equilibrium with liquid shown as solid lines (Fo91-

94) and olivine compositions in adcumulates (pure olivine) shown as diamonds (Fo95-85), with volcanic facies discrimination fields as determined by Barnes (2006a).

The Karasjok Greenstone Belt is characterized by pillowed and thin flows with

variable abundance of volcaniclastic rocks, and exhibits limited MgO enrichment

(maximum 30 wt%). These samples are predominantly classified as thin,

differentiated flow lobes in equilibrium with a maximum olivine composition of

Fo94, with a range extending to less than Fo90 (Fig. 7.6).

Sparse outcrop exposure in all sample areas limited the extent of volcanological

interpretation. However, through the use of volcanic facies differentiation based on

major element abundances, it is possible to estimate the volcanological setting.

These volcanic facies interpretations were reconcilable with more extensive

diamond drilling carried out in the two areas (Hotinvaara and Sarvisoaivi). In

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conclusion, the volcanology of the Archean Enontekiö area and the Pulju

Greenstone Belt (Nilivaara and Hotinvaara areas) is prospective to host

mineralization, as the units are characterized as higher volume flow facies.

However, the volcanological setting identified within the Karasjok Greenstone Belt

is of lower-prospectivity, due to low volume flow facies.

c. Chalcophile element mineralization signatures

Chalcophile element mineralization signatures are the result of sulfur saturation

within the magmatic system and the segregation of an immiscible sulfide melt

(Lesher et al., 1984; 2001). Since the chalcophile elements have high partition

coefficients for the sulfide liquid, they are sensitive to local sulfur saturation events.

As a result, two mineralization indicators (enrichment and associated depletion) are

generated; where both are predicable, recognizable and quantifiable. All chalcophile

elements partition into the sulfide phase; however, Pt and Pd are identified as the

preferred elements to characterize mineralization signatures (Chapters 4 and 6).

Platinum and Pd are relatively more sensitive to depletion and enrichment than Ni

and Cu. Additionally, Pt and Pd exhibit strong incompatibility, and proportionally

occur at the highest abundances relative to the other PGE (McDonough and Sun,

1995).

Mineralization signatures are apparent in the whole-rock geochemistry, with the

elimination of silicate fractionation effects through the normalization of the strongly

chalcophile elements (Pt and Pd) to incompatible elements such as Ti, Al, Zr, or Y

(Maier and Barnes, 2005; Barnes et al., 2007; Fiorentini et al., 2010). Titanium is

commonly utilized as the normalizing element (Barnes et al., 2007; Fiorentini, et al.,

2010; Chapters 4 and 6), due to strong incompatibility in ultramafic systems,

moderate abundance and good analytical precision. However, the variable TiO2

abundance of Karasjok-type komatiites results in spurious ratios and false

mineralization signatures. Consequently, Al2O3 which exhibits negative correlations

with MgO (Fig. 7.3) was used instead of TiO2. Utilizing this methodology

(PGE/Alpmn: where pmn is primitive mantle normalized), normal background

concentrations plot as a cluster of data points (Fig. 7.7). Conversely, sulfide-related

enrichment plots up-slope and depletion down-slope from the normal values (Fig.

7.7). Intrinsic to quantifying chalcophile mineralization signatures, is the

279


identification of the normal background value (PGE/incompatible) representing

sulfur free crystallization conditions.

The establishment of baseline concentrations for a specific area is data intensive.

Based on the limited number of samples from the areas examined, it is not possible

to establish background abundances. Previous work has established backgrounds for

2.7 Ga Munro-type and 2.9 Ga Barberton-type komatiites, based on Ni deposits in

Western Australia (Chapters 4 and 6). Even though Karasjok-type komatiites

display different petrogenetic histories from both Munro- and Barberton-type

komatiites, processes leading to mineralization are the same in all komatiite-hosted

Ni deposits, and result in the generation of similar mineralization signatures. The

identification of known mineralization within one of the Karasjok-type areas

(Hotinvaara) provides a relative measure of the applicability of Munro- and

Barberton-type background abundances to Karasjok-type settings.

Figure 7.7. Pt/Alpmn versus Pd/Alpmn diagram for classifying chalcophile element mineralization signatures within komatiitic systems. Fields derived from mineralized Munro- (Long-Victor deposit, Kambalda Dome) and Barberton-type (Maggie Hays deposit, Lake Johnston Greenstone Belt) komatiite systems in Western Australia (Chapters 4 and 6).

280


Plots of Pt/incompatible versus Pd/incompatible (Fig. 7.7), identify mineralization

signatures, with three prospectivity scenarios. The first and least prospective

scenario is characterized by all samples plotting as a cluster in the lower left hand

corner of Figure 7.7, indicating that the melt was derived from a sulfur saturated

source, and consequently this melt was metal depleted. The second scenario with

low prospectivity is characterized by samples plotting as a cluster of data in the

centre of Figure 7.7. The second scenario is dependent upon the initial chalcophile

budget and incompatible element abundance, indicating the melt was sulfur

undersaturated prior to leaving the source area (maximum chalcophile element

budget). However, in the second scenario the melt did not attain sulfur saturation

before crystallization, and all samples represent silicate-controlled chalcophile

element abundances. The final and most prospective scenario for Fe-Ni-Cu sulfide

mineralization, is characterized by a wide scatter of data in Figure 7.7. The wide

distribution of data indicates that the initial melt was sulfur undersaturated and: 1)

portions of the magmatic system preserve the primary chalcophile budget, and 2)

deviations from the primary chalcophile budget indicate portions of the magmatic

system attained sulfur saturation (PGE enrichment and depletion) and possibly

contain accumulated Fe-Ni-Cu sulfides.

The komatiites of the Enontekiö area are Munro-type and plot along the composite

Barberton- and Munro-type trend (Fig. 7.7). The sample data exhibit variation from

low values of <0.1 to highs of 40 in Figure 7.7, and plot as PGE-Enriched, PGE-

Depleted, and Normal-PGE, as defined by other mineralized komatiite systems

(Chapters 4 and 6). Consequently, the samples collected from surface outcrops

would indicate high prospectivity for the area (e.g. system) to host mineralization.

This interpretation is verified by exploration diamond drilling and the delineation of

mineralization within the Sarvisoaivi area (0.7 Mt at 0.4% Ni: Papunen et al., 1977;

Saltikoff et al., 2006).

Paleoproterozoic komatiitic units within the Karasjok Greenstone Belt are Karasjok-

type, however appear Al-undepleted and exhibit a range of Pt/Alpmn and Pd/Alpmn

values from 0.2 to >4 (Fig. 7.7). The majority of samples plot within and adjacent to

the Normal-PGE field, differing slightly from that of Munro and Barberton-types.

Additionally, one sample exhibits minor enrichment and three samples plot as

281


strongly depleted, reflecting the influence of sulfur saturation. Mineralization is not

documented in the sample area or within the Karasjok Greenstone Belt. However,

the PGE/Alpmn mineralization signatures, in addition to Normal-PGE values, support

the presence of Fe-Ni-Cu sulfides within the system.

Komatiite units in the Pulju Greenstone Belt (Nilivaara and Hotinvaara areas) are

Al-depleted Karasjok-type (high Fe-Ti) and exhibit a Pt/Alpmn and Pd/Alpmn range

from 0.3 to 10 (Fig. 7.7). The samples plot as both Normal-PGE concentrations and

as an enrichment trend, with both data series located below the composite Munro-

and Barberton-type trend (Fig. 7.7). Mineralization is identified within the

Hotinvaara area (1.3 Mt at 0.43% Ni: Papunen, 1998; Saltikoff et al., 2006), thus

supporting the observed PGE-enrichment trend. The deviation from the composite

Barberton-Munro trend is interpreted as a function of a differing initial chalcophile

budget, relative to the Barberton- and Munro-types. Insufficient samples with liquid

compositions (flow top breccias, chilled margins) lacking mineralization signatures

prevented the estimate of an initial chalcophile composition. Regardless, observed

mineralization signatures should parallel the documented Barberton-Munro Pt/Alpmn

and Pd/Alpmn trend. Depletion signatures are not observed in the data set from the

Pulju Greenstone Belt. However, depletion signatures typically constitute 10% of

the samples in komatiite Ni systems (Fiorentini et al., 2010; Chapters 4 and 6).

7.7. Conclusions

The Karelian Craton of northwestern Russia, northern Finland and Norway is typical

of other Archean cratons. It comprises numerous terranes, complex tectonic and

deformational histories, varying extent of alteration and metamorphic grade, and

variable outcrop exposure. Consequently, targeting Ni deposits hosted within these

settings is challenging. To aid in exploration targeting, the application of

lithogeochemistry prospectivity indicators (both major and chalcophile elements) is

evaluated to assess their practical application as a tool in Ni exploration.

Komatiitic lithologies in the central Karelian Craton are diverse in both age

(Archean and Paleoproterozoic) and geochemical type, comprising both Munro- and

Karasjok-type (Al-depleted and Al-undepleted, respectively). Nickel mineralization

282


is delineated by exploration diamond drilling within select areas, providing

necessary supporting evidence for assessing the prospectivity interpretations.

Outcrop mapping, field observations and sampling are important first step in

assessing Ni prospectivity. Within the sample areas, mixed thin flows and

unconstrained cumulate bodies were identified in the Enontekiö area and Pulju

Greenstone Belt, whereas thin and pillowed flows dominated the Karasjok

Greenstone Belt. Resultantly, the Enontekiö area and Pulju Greenstone Belt are

more prospective, as the volcanic facies reflect higher volumes of magma flow-

through.

Major element whole-rock geochemistry provides petrogenetic information and a

means to differentiate komatiites into Barberton-, Munro- and Karasjok-types (Fig.

7.5). The importance of constraining komatiite type as a prospectivity indicator is

debatable, although the majority of Ni deposits are hosted within Munro-type

komatiites. None the less, Barberton- and Karasjok-type komatiites are not barren.

Major elements analysed by whole-rock geochemistry can be utilized to differentiate

flow facies on the basis of FeO and MgO contents (Fig. 7.6: Barnes, 2006a). This

forms a critical assessment of the prospectivity of systems, since identification of

dunitic and channelized sheet flow facies warrant a more in-depth examination.

Volcanic sequences characterized by thin flows require additional positive

prospectivity indicators prior to further work being undertaken. Cumulate lithologies

identified and interpreted within the field areas have been further refined, as

observed in the Enontekiö area, where the majority of identified cumulate bodies are

not classified as dunitic bodies, but as less prospective layered sill and lava lakes

(Fig. 7.6).

The chalcophile element concentrations, specifically the PGE, are excellent

indicators of the mineralization potential of an area. However, the whole-rock data

obtained from the central Karelian Craton exhibits extensive scatter and no apparent

systematic mineralization indicators. The use of PGE/incompatible element ratios

allows for the rapid identification of strong chalcophile element enrichment and

depletion, by reducing the effects of olivine (Fiorentini et al., 2010; Chapters 4 and

6). Further classification of observed mineralization signatures, based on signature

283


fields defined by mineralized Barberton- and Munro-type komatiites of Western

Australia, allows for regional comparisons with limited pre-existing data. Within the

Enontekiö area, Munro-type komatiites host Fe-Ni-Cu sulfide mineralization, and

validate the classified chalcophile element enrichment and depletion signatures (Fig.

7.7).

Ultramafic rocks from the Karasjok and Pulju Greenstone Belts are classified as

Karasjok-type with associated picrites. Petrogenetic differences between the two

belts are identified in contrasting Al-contents within the ultramafic rocks.

Consequently, the samples show slight deviations from the defined Barberton- and

Munro-type trends (Fig. 7.7). Despite the differing baselines, it is argued that

mineralization signatures will plot parallel to the characterized Barberton and Munro

trends, at slightly higher or lower PGE/incompatible element values. This hypothesis

is supported by Fe-Ni-Cu sulfide mineralization data from the Pulju Greenstone

Belt, which exhibits an enrichment trend below the defined fields. Although, Fe-Ni-

Cu sulfide mineralization is not yet identified within the Karasjok Greenstone Belt,

the chalcophile element mineralization indicators support the presence of

mineralization. Identification of high-volume flow conduits within the system,

rather than the sampled thin flows, is the critical next step in targeting potential Fe-

Ni-Cu sulfide mineralization within the Karasjok Greenstone Belt.

This research positively identifies the Enontekiö area and the Pulju Greenstone Belt

(Nilivaara and Hotinvaara areas) as having high Ni prospectivity, based on the

application of whole-rock major and chalcophile element geochemistry. Although

whole-rock prospectivity indicators were derived from well-documented nickel

sulfide deposits, prospectivity indicators are applicable for greenfields exploration in

terranes with sparse outcrop exposure, minor to no previous work, and limited

possible volcanological interpretations.

Acknowledgements

This research was supported by a 2007 Society of Economic Geologists Foundation Inc. student research grant (Hickok-Radford Memorial Fund) to G. Heggie. The research forms part of a larger project with AMIRA, BHP-Billiton, Independence Group and Noril’sk Nickel; we are grateful for their support. The manuscript greatly benefited from the insightful and thorough reviews of Tapio Halkoaha, Martin Prendergast and Mei-Fu Zhou.

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Barley, M.E., Kerrich, R., Reudavy, I., and Xie, Q., 2000, Late Archaean Ti-rich, Al-depleted komatiites and komatiitic volcaniclastic rocks from the Murchison Terrane in Western Australia: Australian Journal of Earth Sciences, v. 47, p. 873-883.

Barnes, S-J., Naldrett, A.J., and Gordon, M.P., 1985, The origin of the fractionation of platinum-group elements in terrestrial magmas: Chemical Geology, v. 53, p. 303-323.

Barnes, S-J., and Often, M., 1990, Ti-rich komatiites from northern Norway: Contributions to Mineralogy and Petrology, v. 105, p. 42-54.

Barnes, S.J., 1998, Chromite in komatiites, I. Magmatic controls on crystallization and composition: Journal of Petrology, v. 39, p. 1689-1720.

Barnes, S.J., 2006a, Komatiites: Petrology, Volcanology, Metamorphism, and Geochemistry, in Barnes, S. J., ed., Nickel Deposits of the Yilgarn Craton: Geology, Geochemistry, and Geophysics Applied to Exploration. Special Publication No. 13, Society of Economic Geologists, Inc., p. 13-50.

Barnes, S. J., 2006b, Komatiite-hosted nickel sulfide deposits: Geology, Geochemistry, and Genesis, In, Barnes, S. J., (ed.), Nickel Deposits of the Yilgarn Craton: Geology, Geochemistry, and Geophysics Applied to Exploration. Special Publication No. 13, Society of Economic Geologists, Inc., p. 51-118.

Barnes, S.J., Hill, R.E.T., Perring, C.S., and Dowling, S.E., 2004, Lithogeochemical exploration for komatiite-associated Ni-sulfide deposits: strategies and limitations: Mineralogy and Petrology, v. 82, p. 259-293.

Barnes, S.J., Lesher, C.M., and Sproule, R.A., 2007, Geochemistry of komatiites in the Eastern Goldfields Superterrane, Western Australia, and the Abitibi Greenstone Belt, Canada, and implications for the distribution of associated Ni-Cu-PGE deposits: Applied Earth Science (Transactions of Mining and Metallurgy Series B), v. 116, p. 167-187.

Braathen, A. and Davisen, B., 2000, Structure and stratigraphy of the Palaeoproterozoic Karasjok Greenstone Belt, northern Norway – regional implications: Norsk Geologist Tidsskrift, v. 80, p. 33-50.

Fiorentini, M.L., Barnes, S.J., Lesher, C.M., Heggie, G.J., Keays, R.R., and Burnham, M.O., 2010, Platinum-group element geochemistry of mineralized and non-mineralized komatiites and basalts: Economic Geology.

Fiorentini, M.L., Barnes, S.J., Maier, W.D., Burnham, M., and Heggie, G.J., in press b, Global variability in the platinum-group element contents of komatiites: Journal of Petrology.

Gaál, G., and Gorbatschev, R., 1987, An outline of the Precambrian evolution of the Baltic Shield: Precambrian Research, v. 35, p. 15-52.

Gangopadhyay, A., Walker, R.J., Hanski, E., and Solheid, P.A., 2006, Origin of Paleoproterozoic Komatiites at Jeesiö, Kittilä Greenstone Complex, Finnish Lapland: Journal of Petrology, v. 47, p. 773-789.

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Hill, R. E. T., 2001, Komatiite volcanology, volcanological setting and primary geochemical properties of komatiite-associated nickel deposits, In Geochemical exploration for gold and nickel in the Yilgarn Craton, Western Australia: Part 2.

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Lehtonen, M., Airo, M-L., Eilu, P., Hanski, E., Kortelainen, V., Lanne, E., Manninen, T., Rastas, P., Räsänen, J.Ja., and Virronsalo, P., 1998, The stratigraphy, petrology and geochemistry of the Kittilä greenstone area, northern Finland: Geological Survey of Finland, Report of Investigations, v. 140. 144p.

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Lesher, C.M., and Keays, R.R. 2002, Komatiite-associated Ni-Cu-PGE deposits: Geology, Mineralogy, Geochemistry, and Genesis, In Cabri, L.J., (ed.), The Geology, Geochemistry, Mineralogy and Mineral Beneficiation of Platinum-Group Elements. Canadian Institute of Mining, Metallurgy and Petroleum, Special Volume 54, p. 579-617.

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List of Contents

7.1. Introduction ................................................................................................... 262 a. Volcanic facies ........................................................................................ 262 b. Mineralization indicators ......................................................................... 263 c. Test area ................................................................................................... 263

7.2. Regional Setting ............................................................................................ 264 a. Central Karelian Craton ........................................................................... 264

i. Archean komatiites (2.9-2.7 Ga) ............................................................. 266 ii. Paleoproterozoic komatiites (2.0-1.9 Ga) ............................................... 266

7.3. Sampling and Physical Volcanology ............................................................. 268 a. Archean komatiites (Enontekiö area) ...................................................... 269 b. Paleoproterozoic komatiites (Pulju and Karasjok Greenstone Belts) ...... 269

7.4. Materials and Methods .................................................................................. 270 7.5. Whole-Rock Geochemistry Results .............................................................. 271

a. Archean komatiites (Enontekiö area) ...................................................... 271 b. Paleoproterozoic komatiites (Karasjok and Pulju Greenstone Belts) ...... 272

7.6. Lithogeochemical Prospectivity Indicators ................................................... 274 a. Petrogenetic classification and initial chalcophile content ...................... 274 b. Volcanic facies ........................................................................................ 277 c. Chalcophile element mineralization signatures ....................................... 279

7.7. Conclusions ................................................................................................... 282 7.8. References ..................................................................................................... 285

Figure 7.1. Map of northern Sweden, Norway, Finland and northwestern Russia showing the distribution of the Paleoproterozoic Central Lapland Greenstone Belt (green), and associated komatiite and picritic rocks (black). Sampling areas are delineated by boxes comprising the: Archean Enontekiö Area, and Paleoproterozoic Pulju and Karasjok Greenstone Belts. Inset map of Norway, Sweden and Finland showing major tectonic divisions of the Baltic Shield. Modified from Hanski et al. (2001). ................................................................ 265

Figure 7.2. Paleoproterozoic stratigraphic sequences and correlations within the Central Lapland Greenstone Belt, comprising the Karasjok, Pulju and Kittilä Greenstone Belts; with arrows indicating formations sampled within the Karasjok and Pulju belts. Formations and Groups are identified with characteristic lithologies summarized: mf. vol. = mafic volcanic, amp. = amphibolite, vol. clast. = volcaniclastic, kom. = komatiite, psam. = psammite, thole. vol. = tholeiitic volcanic, cong. = conglomerate, fels. vol. = felsic volcanic, suf. sed. = sulfidic sediment, qutz. = quartzite, BIF = banded iron formation. Complied from Braathen and Davidson (2000); Papunen (1998); Lehtonen et al. (1998). Age determinations from Pihiaja and Manninen (1988), Hanski et al. (1997). ........................................................................................ 268

Figure 7.3. Bivariant plots of major elements for the ultramafic units from the three areas within the central Karelian Craton, as determined by XRF and ICP-MS. Komatiites from the Archean Enontekiö area (Sarvisoaivi), and Paleoproterozoic areas: Karasjok (Karasjok Greenstone Belt), and Nilivaara and Hotinvaara (Pulju Greenstone Belt). ......................................................... 271

Figure 7.4. Bivariant plots of chalcophile and major elements for the ultramafic units from the three areas within the central Karelian Craton, as determined by

289


290

fire-assay ICP-MS. Komatiites from the Archean Enontekiö area (Sarvisoaivi) and Paleoproterozoic areas: Karasjok (Karasjok Greenstone Belt), and Nilivaara and Hotinvaara (Pulju Greenstone Belt). ........................................ 273

Figure 7.5. [Al2O3] versus [TiO2] high-MgO volcanic discrimination diagram of Hanski et al. (2001). Where [Al2O3] and [TiO2] are normalized mole proportions using the equations [Al2O3] = Al2O3/(2/3-MgO-FeO) and [TiO2] = TiO2/(2/3-MgO-FeO): (see Hanksi, 1992). ..................................................... 275

Figure 7.6. FeO wt% versus MgO wt% recalculated to volatile free for ultramafic samples from Central Karelian Craton. Olivine compositions in equilibrium with liquid shown as solid lines (Fo91-94) and olivine compositions in adcumulates (pure olivine) shown as diamonds (Fo95-85), with volcanic facies discrimination fields as determined by Barnes (2006a). ................................. 278

Figure 7.7. Pt/Alpmn versus Pd/Alpmn diagram for classifying chalcophile element mineralization signatures within komatiitic systems. Fields derived from mineralized Munro- (Long-Victor deposit, Kambalda Dome) and Barberton-type (Maggie Hays deposit, Lake Johnston Greenstone Belt) komatiite systems in Western Australia (Chapters 4 and 6)....................................................... 280

List of Tables

zed by

289289

Appendix Table 7.1A. Whole-rock geochemistry of ultramafic rocks from Karasjok and Pulju Greenstone Belts and Enontekiö area. Major elements analyXRF and given in wt% oxide and chalcophile elements by ICP-MS from NiS fire assay pre-concentration with PGE concentrations in ppb and Ni, Cu in ppm. Morphology as determined from outcrop mapping: TF = Thin flow, MF = Massive flow, PF = Pillowed flow, FR = Fragmental textured, Flt = Flow top. Sample location given as decimal degrees latitude (Lat) and longitude (Long) with WGS84 datum. LOI = loss on ignition, n.d. = not determined.

Sample WP-44 WP-45 WP-46 WP-47 WP-48 WP-49 WP-50 WP-51 WP-52 WP-53 WP-54 WP-55Location Nilivaara Nilivaara Nilivaara Nilivaara Nilivaara Nilivaara Nilivaara Nilivaara Nilivaara Hotinvaara Hotinvaara Hotinvaara

Morphology* TF MF TF FR MF MF MF MF Flt MF MF MF

Lat 68.11815 68.11914 68.12009 68.11795 68.11801 68.11759 68.11764 68.11599 68.11578 68.08929 68.08776 68.08955Long 24.50947 24.50507 24.50447 24.50533 24.50417 24.50401 24.5041 24.49681 24.49693 24.42158 24.41607 24.4118

WP 44 WP 45 WP 46 WP 47 WP 48 WP 49 WP 50 WP 51 WP 52 WP 53 WP 54 WP 55SiO2 40 46.5 44.6 49.2 47 41.1 41.7 45 48.3 40.4 47.6 43.3TiO2 0.73 0.07 0.65 0.58 0.05 0.03 0.03 0.43 0.61 0.19 0.72 0.08Al2O3 8.88 2.26 6.53 5.65 0.91 0.67 1.42 4.68 5.49 3.69 6.01 4.25

FeO tot 12.06 5.43 10.98 9.00 5.08 6.22 6.58 5.26 9.90 9.00 10.71 5.88MgO 23.9 31.6 22.8 20.3 33 37.5 34.6 31.3 21.5 32.2 21.9 31.6CaO 5.98 5.3 8.05 10.3 3.42 0.3 2.58 5.7 8.96 3.21 9.41 4.29Na2O 0.18 0.07 0.32 0.37 0.06 0.04 0.05 0.08 0.21 0.13 0.28 0.21K2O 0.02 n.d. 0.03 0.03 n.d. n.d. n.d. n.d. 0.01 n.d. 0.07 0.02P2O5 0.047 0.007 0.027 0.04 0.004 0.003 0.005 0.068 0.036 0.022 0.044 0.014Cr2O3 0.404 0.374 0.31 0.228 0.348 0.327 0.398 0.261 0.249 0.572 0.275 0.662S % 0.03 0.33 0.01 0.17 0.64 0.44 0.85 0.22 0.23 0.11 0.13 0.14LOI 6.17 7.22 4.27 3.32 8.27 12.3 10.8 6.39 4.05 9.47 1.86 8.63

Al2O3/TiO2 12 32 10 10 18 22 47 11 9 19 8 53

Ni 1420 1160 1480 390 2550 2040 2370 1220 580 2030 1310 2040Cu 260 40 50 20 30 20 60 20 40 n.d. 60 50Ir 2.4 2.5 2.8 1.6 4.1 3.8 3.8 2.1 2.1 1.6 2.9 1.2

Ru 5.3 7.7 5.1 3.5 8.7 7.2 8.7 4.2 2.9 9.3 4.3 3.9Rh 1.3 1.6 1.6 1.1 1.2 1.1 1.2 1.3 0.9 1.2 1.1 2.9Pt 15 27.5 44 34 9.5 13.5 4 8 11.5 8.5 9.5 5.5Pd 6 3 11 8 2 3 2.5 7.5 9 7 7 2.5

290290

Ir 2 1 0

Sample WP-56 WP-59 WP-60 WP-61 WP-62 WP-63 WP-64 WP-65 WP-66 WP-67 WP-68 WP-69Location Hotinvaara Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi

Morphology* MF MF MF MF MF MF MF MF MF TF TF TF

Lat 68.09171 68.63982 68.6398 68.63962 68.63989 68.63686 68.63373 68.63372 68.63335 68.63202 68.632 68.63561Long 24.41275 21.90222 21.90015 21.90009 21.89256 21.89952 21.9078 21.90821 21.90921 21.91375 21.91411 21.91635

WP 56 WP 59 WP 60 WP 61 WP 62 WP 63 WP 64 WP 65 WP 66 WP 67 WP 68 WP 69SiO2 55.6 36.2 38.7 37.9 37.1 40.8 46.5 45.3 49.8 43.7 46.3 39.3TiO2 0.03 0.07 0.04 0.03 0.01 0.14 0.26 0.29 0.23 0.84 0.24 0.39Al2O3 1.72 1.97 1.25 0.96 0.51 2.79 7.55 8.21 6.34 12.7 7.48 13.2

FeO tot 4.26 10.08 10.26 8.67 7.18 9.72 9.27 11.16 9.90 16.11 10.98 13.86MgO 22.1 32.8 35.2 37.7 43.6 31.8 22.9 21.7 20.7 6.49 21.8 19CaO 13.2 0.4 0.05 0.02 0.5 1.22 7.38 7.55 8.63 16.1 6.84 6.3Na2O 0.24 0.04 0.04 0.03 0.05 0.02 0.24 0.37 0.43 0.47 0.17 0.37K2O 0.02 n.d. 0.01 n.d. n.d. n.d. 0.03 0.05 0.06 0.19 0.02 0.19P2O5 0.003 0.037 0.007 0.008 0.008 0.014 0.01 0.023 0.012 0.076 0.018 0.022Cr2O3 0.285 0.365 0.474 0.942 1.448 0.325 0.478 0.498 0.424 0.062 0.287 0.117S % 0.08 0.1 0.02 1.91 0.3 0.08 0.1 0.45 n.d. 0.06 0.08 0.01LOI 2.17 15.3 12.1 12.4 8.36 10.6 4.31 3.34 1.89 1.02 4.51 5.52

Al2O3/TiO2 57 28 31 32 51 20 29 28 28 15 31 34

Ni 1430 8410 2530 3260 3250 1800 920 1000 790 200 790 360Cu 20 240 90 n.d. n.d. 20 20 40 40 30 n.d. 20Ir 2 82.8 3 53.5 2 4.4 1 51.5 4 24.2 0 60.6 0 60.6 1 11. 0 6.6 0 2 0 5 0 50.2 0.5 0.5

Ru 5.7 33.8 14.6 5.4 25.2 5.6 5 4.5 4.4 0.5 3.1 2.1Rh 0.7 12.3 4.1 1.7 5.6 1.4 1.3 1.4 1.2 0.2 1.2 0.7Pt 2 42 13.5 6.5 11.5 8.5 8 9 8 2 7 4.5Pd 1 122 37.5 17.5 12.5 8 3.5 13 9 1 5.5 4.5

291291

Ir 7 2 1 3

Sample WP-70 WP-71 WP-72 WP-73 WP-75 WP-76 WP-77 WP-78 WP-79 WP-80 WP-81 WP-82Location Sarvisoaivi Sarvisoaivi Sarvisoaivi Sarvisoaivi Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok

Morphology* TF TF MF MF TF TF PF PF PF PF PF PF

Lat 68.63574 68.63577 68.63768 68.63777 70.04265 70.04268 70.04252 70.04077 70.04029 70.03971 70.03971 70.03906Long 21.91681 21.91715 21.91367 21.91367 25.10507 25.105 25.10551 25.11119 25.11142 25.11139 25.11138 25.1082

WP 70 WP 71 WP 72 WP 73 WP 75 WP 76 WP 77 WP 78 WP 79 WP 80 WP 81 WP 82SiO2 44 48.5 36.8 35.8 64.2 48.4 43.1 42.2 44.8 42.3 44.1 42.8TiO2 0.32 0.52 0.02 0.64 0.33 0.47 0.68 0.62 0.5 0.63 0.53 0.27Al2O3 8.96 15.2 1.21 14.6 16.2 8.28 9.18 8.15 7.91 9.75 9.45 5.18

FeO tot 8.87 9.90 8.40 8.01 3.13 9.09 10.53 10.17 9.90 10.89 10.26 8.49MgO 23.2 9.79 36.1 27.2 3.35 16.9 19.5 20.5 21.1 21.4 21 26.9CaO 7.26 11.4 0.39 3.14 2.67 10.2 9.08 10 9.13 7.8 7.89 3.91Na2O 0.2 1.71 0.05 0.05 8.08 1.54 0.95 0.81 0.85 0.9 0.66 0.04K2O 0.02 0.33 n.d. n.d. 0.29 0.23 0.09 0.1 0.12 0.12 0.04 n.d.P2O5 0.02 0.036 0.007 0.065 0.118 0.02 0.038 0.052 0.042 0.051 0.042 0.002Cr2O3 0.335 0.067 1.832 0.05 0.046 0.238 0.292 0.291 0.283 0.297 0.255 0.519S % n.d. n.d. 0.14 0.12 n.d. 0.13 n.d. n.d. n.d. n.d. n.d. n.d.LOI 5.89 1.29 13.8 9.22 0.99 3.28 5.07 5.7 4.29 4.65 4.54 10.7

Al2O3/TiO2 28 29 61 23 49 18 14 13 16 15 18 19

Ni 470 200 1860 720 110 1060 1130 1090 1090 1140 990 1520Cu 20 150 20 n.d. 50 180 80 40 20 30 20 60Ir 0 70.7 0 40.4 7 2.2 0 20.2 0 20.2 0 80.8 1 91.9 1 21. 1 1.1 1 3 0 7 31.3 0.7

Ru 4.4 2.1 16.1 0.7 0.5 2.3 7.6 3.1 2.8 4.2 3.6 6.3Rh 0.7 0.9 1.6 0.1 n.d. 0.7 2.7 0.9 0.7 1.1 1.2 1.4Pt 4.5 7.5 8 1 1 6.5 16 9 7.5 11 9.5 17Pd 11.5 7.5 6 0.5 1.5 2.5 6 8.5 2.5 6 3.5 21.5

292292

Ir 0

Sample WP-83 WP-84 WP-86 WP-87 WP-88 WP-91 WP-92 WP-93 WP-94Location Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok Karasjok

Morphology* PF PF PF FR FR TF TF TF TF

Lat 70.03894 70.03227 70.03298 70.03309 70.03311 70.03085 70.03083 70.03083 70.03039Long 25.10805 25.12266 25.12109 25.12059 25.12051 25.07221 25.07203 25.07208 25.07303

WP 83 WP 84 WP 86 WP 87 WP 88 WP 91 WP 92 WP 93 WP 94SiO2 45.9 48.5 44 47.1 40.5 43.4 44.4 36.9 43.7TiO2 0.62 0.45 1.19 0.52 0.72 1.26 1.07 3.42 1.36Al2O3 10.5 7 13.4 6.59 11.9 8.4 7.43 10.3 7.01

FeO tot 10.62 8.49 11.61 8.39 11.34 11.43 11.61 14.76 10.89MgO 16.1 20.4 13.5 22.1 20.8 20.1 20.8 17 20.1CaO 11.3 9.85 11.5 8.61 6.75 8.29 8.35 8.96 8.95Na2O 1.74 0.74 2 0.29 0.55 0.38 0.4 0.64 0.38K2O 0.08 0.05 0.25 0.02 0.04 0.04 0.05 0.1 0.04P2O5 0.013 0.01 0.055 0.034 0.05 0.081 0.077 0.426 0.108Cr2O3 0.253 0.292 0.163 0.365 0.288 0.248 0.214 0.005 0.261S % n.d. n.d. 0.04 n.d. n.d. n.d. n.d. n.d. n.d.LOI 1.37 3.08 0.79 4.84 5.53 4.8 3.98 5.26 5.76

Al2O3/TiO2 17 16 11 13 17 7 7 3 5

Ni 800 950 350 1050 1060 840 1050 130 900Cu 20 20 110 60 20 300 70 20 50Ir 0 80.8 1 11.1 0 7.7 1 41.4 0 90.9 1 81.8 1 41.4 0 1 1 70.1 1.7

Ru 3.6 3.7 2.5 2.8 3.4 3.8 3 0.5 4.1Rh 1.2 0.7 0.6 0.6 0.7 0.7 0.5 n.d. 1.1Pt 9.5 6.5 6 6.5 8 4.5 4 1 9.5Pd 2.5 2 3.5 4 1.5 4 3.5 n.d. 8

Chapter 8. Conclusions

Chapter 8. Conclusions: Application of Platinum Group Elements

in Komatiite-Hosted Nickel Exploration

8.1. Conclusions

This research thesis presents the first quantification of the spatial and genetic

correlation between nickel (Ni) sulfide mineralization and platinum group element

(PGE) ore forming signatures in komatiite systems. The understanding of this spatial

correlation provides the means to translate a mineralization indicator into a

mineralization vector, which can be used in Ni exploration. Two komatiite-hosted Ni

sulfide deposits were used as case studies to quantify PGE mineralization signatures.

The Long-Victor and Maggie Hays deposits of the Yilgarn Craton in Western

Australia differ in mineralization style (extrusive vs. intrusive), age (2.7 vs. 2.9 Ga),

and geochemistry (Munro- versus Barberton-type), and exhibit both similar and

differing spatial correlations between ore forming signatures and physical

mineralization. Within each deposit, the mineralization signatures are characterized

as proximal enrichment and distal depletion, both functioning to expand the volume

of the system. The presence of proximal and distal signatures indicates the potential

for the system to host Ni mineralization beyond the limits of the physical

mineralization.

Platinum group element enrichment signatures in both the Long-Victor and Maggie

Hays systems exhibit increasing enrichment signatures with proximity to

mineralization. This proximal enrichment is here argued to represent a primary

mineralization halo, rather than secondary diffusion away from mineralization. A

geochemical halo enrichment signature is identified within approximately 30 m of

mineralization at the Long-Victor deposit and approximately 30-50 m within the

Maggie Hays system. Despite the similarity in enrichment signatures between the

two systems, the localization of depleted environments is significantly different.

The Long-Victor and Maggie Hays systems exhibit differing physical locations for

depletion signature preservation and gradients, which are likely a product of the

differing magmatic settings. However, depletion signatures in both settings exhibit

the influence of recharging magma. Within the Long-Victor system, depletion

signatures are preserved in the flanking environments at a distance of 340 m from

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the mineralized channel, and exhibit diminishing signature magnitude both

downward through the basal komatiite flow, and with proximity to the mineralized

channel. Both observed chemical gradients are a result of magma flushing within the

basal flow. Within the Maggie Hays system, depletion signatures occur at a distance

of 320 m from mineralization, coinciding with the sulfur contaminant source. From

the initial location of sulfur saturation, depletion signatures exhibit an increasing

trend, followed by a progressive decrease with proximity to mineralization. These

observed distribution and chemical trends are here argued to be the product of

magma flow through a conduit system.

The identification and quantification of mineralization signatures requires the

establishment of initial background abundances of the chalcophile elements.

Background abundances were derived from the deposit data sets (Long-Victor and

Maggie Hays) through the use of Ti-normalization and data filtering. The use of Ti

as a normalization factor removes low-pressure olivine fractionation and

accumulation effects. As the PGE (specifically Pt, Pd, Rh) are strongly incompatible

with olivine, the removal of olivine fractionation effects results in residual PGE

abundances, which are controlled by crystallizing phases other than olivine (Barnes

and Maier, 1999). These other identified phases comprise: chromite (Fiorentini et

al., 2008), PGM alloys (Barnes and Fiorentini, 2008), and sulfide (Chapters 4 and

6; Fiorentini et al., 2010b). Sulfide is the preferred PGE collector in mineralized

systems. Samples containing accumulated sulfide typically exhibit enrichment,

whereas samples that interacted with a sulfide liquid are PGE-depleted.

Iterative filtering of the deposit geochemical data sets on the basis of Pt/Tipmn and

Pd/Tipmn, progressively removes samples which are depleted or enriched from a

median value. The median value is interpreted to represent an initial chalcophile

element: Ti ratio. The resultant filtered data sets cover a range of MgO contents with

background PGE abundances. The linear regression of the filtered data (PGE versus

MgO) was used to calculate background abundances for all samples as a function of

MgO content, with enrichment and depletion representing the positive and negative

residual anomalies from the background abundance.

Previous work has characterized background chalcophile element abundances from a

wide spectrum of non-mineralized ultramafic-mafic magmatic systems, identifying a

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narrow range of initial chalcophile abundances (Maier et al., 2009; Fiorentini et al.,

2010a). The similarity in initial chalcophile element abundances allows for the

uniform application of PGE-based mineralization signatures, between komatiites of

different petrogenesis, age (≤ 2.9 Ga), and tectonic setting.

Even though background abundances can be calculated for all chalcophile elements

(Ni, Cu, PGE: Ir, Ru, Rh, Pt, Pd), the identification of chalcophile element

mineralization signatures is not practical with all chalcophile elements, due to

varying partition coefficients (Table 8.1), relative abundance, analytical methods,

and PGE collector phases other than sulfide, as summarized in Table 8.2:

Table 8.1. Partition coefficients for the chalcophile elements between silicate liquid and sulfide liquid. 1. Francis (1990); 2. Sattari et al. (2002); 3. Gaetani and Grove (1997); 4. Peach et al. (1990); 5. Jana and Walker (1997); 6. Rajamani and Naldrett (1978); 7. Stone et al. (1990); 8. Bezmen et al. (1994); 9. Fleet et al. (1999); 10. Peach et al. (1994); 11. Helz and Rait (1988).

Ni Cu Ir Ru Rh Pt Pd Reference 315-424 913-1006 1

810-1300 >50000 >12000 >140000 >18000 >92000 2 410-580 250-313 3 575-836 1383 14000 23000 4

350-1070 5 257-274 6

130000 9100 88000 7 310000 2500 27000 55000 8

1800-51000

2400-35000

1400-20000 2900-25000 9

35000 43000 10 36000 25000 11

Table 8.2.Mineralization signature characteristics of the chalcophile elements

Nickel has the lowest Dsul-sil (partition coefficient for sulfide from silicate: Table

8.1), which limits the sensitivity of nickel as a mineralization indicator. This

thesis indicates that Ni is insensitive to sulfur saturation, with respect to the

development of depletion signatures. Poor to no visible correlations are

observed between depletion in Pt and Pd with depletion in Ni.

Copper, being strongly incompatible, moderately abundant at liquid compositions

(~ 50 ppm at 25 wt% MgO), and having moderate Dsul-sil (~ 1000: Table

8.1), makes it an ideal element for mineralization signatures, both

individually and as ratios. However, Cu is highly mobile under any hydrated

prograde, retrograde or contact metamorphic processes and during all forms

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of weathering, making it of limited to no use in the current application of

mineralization signatures. Within the two case study areas (Long-Victor:

green schist facies, and Maggie Hays: amphibolite facies), Cu no longer

displays incompatible behavior with fractionation due to extensive

mobilization. Consequently, both Cu enrichment and depletion exhibit poor

to no correlation with mineralization signature trends observed with the PGE.

Iridium has a high Dsul-sil (>10000: Table 8.1) and is potentially sensitive to sulfur

saturation within ore forming systems. However, both low initial abundance

and temperature-dependent Ir-alloy saturation (Barnes and Fiorentini, 2008)

limits the current use of Ir as a lithogeochemical vector. In comparison with

other incompatible PGE (Pt and Pd), Ir correlates well with strong depletion,

but exhibits mixed signals with weak depletion signatures. Iridium correlates

well with enrichment for all other chalcophile elements.

Ruthenium has a high Dsul-sil (> 10000: Table 8.1), occurs at relatively moderate

abundances (3-5 ppb at 25 wt% MgO) and exhibits stronger incompatibility

with olivine (negative correlation with MgO) than Ir and Ni. However, Ru

exhibits a positive correlation with Cr (chromite). This compatibility is

argued to result from solid solution within komatiite systems (Fiorentini et

al., 2008; Locmelis et al., 2009). The solid solution of Ru in chromite forms

the basis of an individual mineral mineralization indicator. The Ru-chromite

indicator is based on the pretense that chromite crystallizing in sulfur

saturated conditions will be Ru-depleted, whereas chromite from un-

saturated systems will be characterized by elevated Ru abundances

(Fiorentini et al., 2008).

Ruthenium exhibits potential as a mineralization indicator, using both

chromite separates and whole-rock geochemistry. Ruthenium as a

mineralization indicator by whole-rock analyses, identifies similar anomalies

to those observed with Pt and Pd, with good correlations for both enrichment

and depletion. The modest abundance of Ru (3-5 ppb) is a limitation. With

natural sample variability, in addition to current analytical limitations,

uncertainty of ± 1 ppb is not unreasonable for Ru. However, this uncertainty

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masks most depletion signatures occurring in lithologies that are not more

fractionated than the initial liquid.

Rhodium occurs at a low relative abundance (1-2 ppb at 25 wt% MgO); however,

has a high Dsul-sil (> 10000: Table 8.1), and exhibits incompatibility with

olivine, and has an almost constant Rh/Tipmn value. These characteristics,

(excluding low abundance) contribute to make Rh a useful mineralization

indicator. This low abundance hinders intensive interpretation of depletion

signatures, as mentioned previously with Ir and Ru. Regardless of the initial

abundance, Rh depletion and enrichment signatures correlate well with Pt

and Pd.

Platinum has a high Dsul-sil (> 10000: Table 8.1), and occurs at relatively high

abundances (8-10 ppb at 25 wt% MgO). Platinum exhibits strong negative

correlation with MgO, from initial liquid compositions to <10 wt% MgO,

due to incompatibility with olivine. Resultantly, komatiite systems exhibit a

constant Pt/Tipmn value, allowing for easy identification of enrichment and

depletion signatures and calculation of background values. As a consequence

of all these characteristics, Pt is one of two PGE that represent ideal

chalcophile elements for the identification of Ni mineralization signatures.

Platinum is very sensitive to sulfur saturation events, and occurs at a high

enough abundance that even with significant olivine accumulation and

limited trapped liquid (e.g. B-zone cumulates), quantifiable depletion

deviations are apparent.

Palladium is the second PGE that represents an ideal chalcophile element

mineralization indicator. Palladium has a high Dsul-sil (> 10000: Table 8.1),

and occurs at relatively high abundances (6-12 ppb at 25 wt% MgO).

Palladium, similar to Pt, exhibits a strong negative correlation with MgO

from liquidus compositions to < 10 wt% MgO, due to incompatibility with

olivine, and exhibits a constant Pd/Tipmn value. These similar characteristics

between Pt and Pd allow for simplified application and interpretation of Ni

mineralization signatures.

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Palladium is limited in use by suspected higher mobility under hydrated

conditions, which can lead to fractionation from the other less mobile PGE.

Within the Long-Victor and Maggie Hays data sets, several deviations were

observed in the typically perfect linear relationship between Pt and Pd. These

deviations are attributed to PGE mobility. However, these deviations are

observed with both Pt and Pd, and are not limited to Pd enrichment or

depletion. Samples which deviate from a linear relationship between Pt and

Pd comprise a minor portion of the data sets, and to-date are not constrained

by mobility distance, alteration assemblages, or fluid types.

Overall, the chalcophile elements are indicators of sulfur saturation and ore forming

processes; yet some of the chalcophile elements (Pt, Pd, Rh) present clearer

depletion signatures than others (Ni, Ir), although, all of the chalcophile elements

provide good indicators of enrichment. Approximately 40% of whole rock

geochemical samples, within the Long-Victor and Maggie Hays deposit data set

exhibit chalcophile element enrichment. This enrichment appears robust, and is not

modified by metamorphism or alteration within the two study areas. Chalcophile

element depleted samples form the smallest portion of the data set, at ~ 10%.

Despite the minor occurrence of depletion signatures within the data set, these are

critical; as no other process aside from sulfur saturation would deplete the magma to

generate false anomalies, as is possible with enrichment signatures.

Research on PGE-based mineralization signatures is ongoing and evolving (Barnes

et al., 1985; Barnes and Naldrett, 1986; Barnes et al, 1988; Barnes and Maier, 1999;

Lesher et al., 2001; Keays and Lightfoot, 2007; 2010; Fiorentini et al., 2010a; b).

Additionally, this thesis incorporates and builds upon 40 years of research covering

all aspects of komatiite systems: plume melting, komatiite geochemistry,

contamination, komatiite volcanology, and ore forming processes. Arguably, the

research will continue, and the question arises if there are further gains which can be

made towards the understanding of whole-rock chalcophile element mineralization

signatures.

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8.2. Recommendations and Further Research

It is here proposed that a different approach, other than spatial correlations of whole-

rock chalcophile element mineralization indicators, is necessary for the next step of

enlarging the recognizable Ni deposit footprint. This thesis represents a balance

between system understanding and sample density. Further sampling at a higher

resolution is possible, but will only add marginal information. Analytical techniques

will improve, and increased accuracy and precision will be attained, but will not add

more than is already apparent with the Pt and Pd mineralization signatures.

Arguably, questions still exist for both the Long-Victor and Maggie Hays systems

that can only be answered from additional diamond drilling e.g. What happens

beyond 450 m in the Victor flank? Is there another mineralized sub-parallel channel

up-dip to the Victor channel? Is there mineralization hosted stratigraphically within

the Western Ultramafic Unit of the Lake Johnston Greenstone Belt? Consequently,

additional exploration work must be carried out on these systems to add any value to

these questions and further resolve the current chalcophile element mineralization

indicators.

A number of tangential questions and potential research topics arose during the

thesis process. These questions range from Archean tectonics to deposit-specific.

Questions are briefly summarized below:

PGE lithogeochemistry

• The numerical models describing chalcophile element depletion once sulfur

saturation occurs are generally rapid and extreme. However, in reality, extreme

depletion in ore forming systems is rarely documented. Are we still looking in

the wrong parts of the flow system to have prevalent depletion signatures? Are

we overlooking a critical physical or geochemical process that limits the volume

and preservation of depletion signatures? Or are our numerical models not yet

correct?

• PGE-based mineralization vectors are applicable for Ni exploration in komatiite

systems. However, are these vectors applicable in more fractionated systems,

299


with the complexity of additional crystallizing phases potentially partitioning the

PGE?

Tectonics

• Stratigraphic work on the Maggie Hays deposit identified “arc-like” felsic

volcanic units underlying komatiite flows. Similarly, a global association

between komatiites and felsic volcanics is documented in the arc-volcanics units

of the Lake Johnston, Ravensthorpe, Forrestania, Uchi, and North Caribou

Greenstone Belts: However, the occurrence of komatiites in back-arc settings has

not been substantiated. Is it possible to have mantle magmatism without a plume

in back-arc settings? Additionally how does slab roll back, mantle inwelling and

bonninites fit into an alternative setting for komatiite generation?

• Modern day rivers systems exhibit meanders as an equilibrium between flow

velocity, sediment load, and gradient. Channelized volcanic flows on both the

Moon and Mars show strong similarities these fluvial systems (Bleacher et al.,

2010). The channel and trough structures observed in Ni mineralized systems

exhibit interesting arc-like shapes (e.g. Long-Victor, Widgiemooltha Dome,

Katiniq). Could these arc-like shapes provide evidence for meandering komatiite

flows? Is there more than structure associated with these systems? Can we

estimate the spatial distribution of curvature and meanders to better target

exploration drill holes?

Deposit Specific

• Further research should be undertaken on the “sulfur from above” hypothesis, as

presented in Chapter 6. This model works for the Maggie Hays system. Is it

also applicable for the Kambalda Dome Ni deposit systems? Arguments exist for

the presence of transient quench crusts in channelized environments, and

significant sediment accumulation is observed in the flanking environment. Is it

possible to incorporate sulfidic sedimentary material from the flow top, in

addition to the base? This question could be addressed through the use of non

mass-dependent S variation within flank sediments and massive ore profiles.

Metal profiles in the flank sediments were identified by Bavinton (1979). Are

300


these also present in the Fe-Ni-Cu sulfide ore profiles or has monosulfide solid

solution homogenized any potential geochemical indicator?

• Within the Maggie Hays Ni deposit, stratabound mineralization in Western

ultramafic unit (WUU: Chapter 5) represents a viable exploration target and

research topic. Most mineralized extrusive komatiite systems host mineralization

in the basal flow. However, the basal flow within the WUU appears barren.

Initial magmas were sulfur undersaturated, but the deposit model indicates

successive flows may have been sulfur saturated and Ni sulphides may be

present higher in the stratigraphy.

• The origin of flow top breccias in the WUU should be further investigated in the

Maggie Hays deposit. Are these breccias related to paleo-topography, or

something more controversial, such as water in the magma? Degassing features

in the form of amygdules are identified within the flow tops and spinifex.

• The thermal and geochemical influence of a mantle plume on the Lake Johnston

Greenstone Belt stratigraphy requires more investigation (Chapter 5). How far

back through the stratigraphy can plume influences be observed: komatiites,

massive sulfide, BIF, TZU, felsics? Can plume influences be identified using

isotopes (S, Fe, other).

• Research should also focus on the influence primary structural controls on the

spatial distribution of channels within the Kambalda system. Is there a sub-

parallel channel every 150-200 m?

This research thesis has concluded that it is possible to constrain the scale of Ni

mineralized systems. By doing so, it establishes the framework for the use of

applicable PGE (Pt, Pd, Rh) -based lithogeochemical vectors for Ni mineralization.

These PGE mineralization signatures, can be quantified as positive and negative

residual anomalies from a calculated normal background value. Positive and

negative anomalies, in the form of PGE enrichment and depletion, occur within

mineralization systems in varying proportions, and exhibit a spatial dependence

between the magnitude of the residual anomaly and proximity to known Ni

mineralization. These characteristics make it viable to use applicable PGE (Pt, Pd,

301


Rh) mineralization signatures as both lithogeochemical vectors and prospectivity

indicators for Ni sulfide mineralization.

If all chalcophile elements (Ni, Cu, Pt, Pd, Rh, Ru, Ir) are analyzed during routine Ni

exploration, an immense amount of useful information is gains. As Ni is largely

insensitive to sulfur saturation processes, Ni mineralization signatures are present

even when not visually apparent in the rock samples. This research thesis outlines

the volcanic facies that host the applicable PGE (Pt, Pd, Rh) mineralization

signatures, and the spatial correlations between mineralization signatures and Ni

mineralization. As such, a limited number of well-selected geochemical samples

may go a long way towards further targeting Ni within a komatiite system.

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Maier, W.D., Barnes, S.J., Campbell, IH., Fiorentini, M.L., Peltonen, P., Barnes, S-J., Smithies, R.H., 2009. Progressive mixing of meteoritic veneer into the early Earth's deep mantle: Nature, v. 460, p. 620-623.

Peach, C.L., Mathez, E.A., Keays, R.R., 1990. Sulfide melt-silicate melt distribution coefficients for noble metals and other chalcophile elements as deduced from MORB: implications for partial melting: Geochimica et Cosmochimica Acta, v. 54, p. 3379-3389.

Peach, C.L., Mathez, E.A., Keays, R.R., Reeves, S.J., 1994. Experimentally determined sulfide melt-silicate melt partition coefficients for iridium and palladium: Chemical Geology, v. 117, p. 361-377.

Rajamani, V., Naldrett, A.J., 1978. Partitioning of Fe, Co, Ni and Cu between sulfide liquid and basaltic melts and the composition of Ni-Cu sulfide deposits: Economic Geology, v. 73, p. 82-93.

Sattari, P., Brenan, J.M., Horn, I., 2002. Experimental constraints on the sulfide- and chromite-silicate melt partitioning behavior of rhenium and platinum group elements: Economic Geology, v. 97, p. 385-398.

Stone, W.E., Crocket, J.H., Fleet, M.E., 1990. Partitioning of palladium, iridium, platinum and gold between sulfide liquid and basalt melt at 1200C: Geochimica et Cosmochimica Acta, v. 54, p. 2341-2344.

304


305

Table of Contents

8.1. Conclusions ................................................................................................... 293

......................................................................................................................... 295 Nickel .............................................................................................................. 295

......................................................................................................................... 295 Copper ............................................................................................................. 295

......................................................................................................................... 296 Iridium ............................................................................................................. 296

......................................................................................................................... 296 Ruthenium ....................................................................................................... 296

......................................................................................................................... 297 Rhodium .......................................................................................................... 297

......................................................................................................................... 297 Platinum ........................................................................................................... 297

......................................................................................................................... 297 Palladium ......................................................................................................... 297

......................................................................................................................... 298 8.2. Recommendations and Further Research ...................................................... 299 References ................................................................................................................. 303

Table 8.1. Partition coefficients for the chalcophile elements between silicate liquid and sulfide liquid. 1. Francis (1990); 2. Sattari et al. (2002); 3. Gaetani and Grove (1997); 4. Peach et al. (1990); 5. Jana and Walker (1997); 6. Rajamani and Naldrett (1978); 7. Stone et al. (1990); 8. Bezmen et al. (1994); 9. Fleet et al. (1999); 10. Peach et al. (1994); 11. Helz and Rait (1988). ......................... 295

Table 8.2. Mineralization signature characteristics of the chalcophile elements ... 295

Table A.1. Location and description of samples from Long-Victor, Kambalda Dome, Western Australia

Long-Victor, Kambalda Dome, Western AustraliaLab X Y Z Dip Az

KD5051 320 Fl1 Channel Spfx Geolabs 374730.8 6549092.8 -61.8 14.9 46.2 55.5 57.6 0.6KD5051 345 Fl1 Channel BZ Geolabs 374725.0 6549093.4 -86.1 -75.2 175.2 33.9 53.1 0.5KD5073 329 Fl1 Channel Spfx Geolabs 374627.6 6549511.1 -73.6 13.1 54.6 92.7 93.2 2.5KD5073 473.5 Fl1 Channel BZ M Ultratrace 374620.1 6549508.7 -217.9 86.4 260.2 14.5 31.4 0.5KD5073 488 Fl1 Channel BZ Geolabs 374619.1 6549508.6 -232.3 86.2 264.2 29.0 31.4 0.5KD5073 558 Fl1 Channel BZ Geolabs 374614.3 6549508.1 -302.2 -86.0 81.5 1.0 1.6 2.5KD5073 562.8 Fl1 Channel BZ M Ultratrace 374614.0 6549508.0 -306.9 -86.1 79.3 0.1 1.4 6.9KD5081A 583.4 Fl1 Channel BZ M Ultratrace 374494.5 6549719.4 -325.6 83.5 265.8 0.4 1.5 2.6KD5081A 586.5 Fl1 Channel BZ M Ultratrace 374494.2 6549719.4 -328.7 -83.5 86.7 0.4 1.5 4.5KD5082 339 Fl1 Channel Spfx Geolabs 374165.3 6550412.7 -77.5 3.5 50.8 87.4 87.5 0.6KD5082 380 Fl1 Channel BZ Geolabs 374160.5 6550409.8 -118.1 -81.6 59.6 52.9 54.9 0.5KD5082 422 Fl1 Channel BZ Geolabs 374155.5 6550406.5 -159.7 -81.0 68.8 10.9 12.9 0.5KD5082 433.4 Fl1 Channel BZ M Ultratrace 374153.8 6550405.9 -171.0 81.0 251.6 0.5 1.8 0.5KD5085 434 Fl1 Channel BZ Geolabs 374362.5 6549993.8 -172.2 -0.4 168.0 34.9 37.4 1.3KD5085 498 Fl1 Channel BZ Geolabs 374345.8 6549987.9 -233.7 -32.5 54.4 25.3 26.8 1.6

Facies Texture NotesDistance

(m)Distance

(m)Ni

(wt%)

Sample identification given based on collar name and depth. Volcanology of sample described as Fl# = Flow Number (Fl1 being basal flow), BZ = B-zone cumulate, Spfx = spinifex textured, M = mineralized. Sample location given as X, Y, Z UTM coordinates as calculated. Z (depth) is relative to a local mine datum. Closest occurrence of Ni are calculated distances, directions and grades based on the 3 closest occurrences of Ni >0.4%. Az = azimuth. Analytical lab used is indicated for each sample (Lab: Geolabs or Ultratrace) refer to Appendix C for additional information on quality control and quality assurances.

UTM MGA94 Z51Closest occurrence of Ni >

0.4%Ave of 3 closest

occurrences

CollarDepth

(m) Flow #

KD5085 498 Fl1 Channel BZ Geolabs 374345.8 6549987.9 233.7 32.5 54.4 25.3 26.8 1.6KD5085 532.6 Fl1 Channel BZ M Ultratrace 374336.7 6549983.0 -266.8 -72.1 60.0 4.3 7.7 1.9KD5105 152 Fl1 Flank BZ Geolabs 374579.5 6549081.6 111.5 -48.5 345.3 102.9 126.3 0.5KD5106 190 Fl1 Channel Spfx Geolabs 374673.6 6549084.8 64.9 -41.1 54.0 46.2 54.0 0.5KD5106 244 Fl1 Channel BZ Geolabs 374669.8 6549086.2 11.1 -85.2 172.4 2.9 4.9 0.5KD5109 506 Fl1 Flank BZ Geolabs 374851.1 6549091.7 -249.9 -85.3 164.0 20.9 106.5 2.8KD5115W1 705 Fl1 Channel Spfx Geolabs 375068.4 6549095.3 -450.3 -45.2 176.1 115.0 118.3 0.6KD5115W1 710 Fl1 Channel BZ Geolabs 375068.2 6549095.5 -455.3 -43.4 176.2 111.4 115.0 0.6

A1

Long-Victor, Kambalda Dome, Western AustraliaLab X Y Z Dip AzFacies Texture Notes

Distance (m)

Distance (m)

Ni (wt%)Collar

Depth (m) Flow #

KD6025 194.3 Fl1 Flank BZ M Ultratrace 374739.0 6547267.6 59.9 36.7 169.7 332.0 510.5 0.5KD6025 200 Fl1 Flank BZ Geolabs 374738.9 6547268.3 54.3 37.5 169.9 335.2 509.6 0.5KD6025 204 Fl1 Flank BZ Geolabs 374738.8 6547268.8 50.3 38.1 170.0 337.5 509.0 0.5KD6027 316 Fl1 Flank Spfx Geolabs 374946.0 6547639.3 -58.1 -49.8 211.2 225.3 228.6 0.5KD6027 319 Fl1 Flank BZ Geolabs 374945.4 6547639.5 -61.0 -49.3 211.5 223.1 226.4 0.5KD6027 342.5 Fl1 Flank BZ M Ultratrace 374941.2 6547642.0 -84.0 -45.0 213.4 206.4 209.7 0.5KD6037 282 Fl1 Channel Spfx Geolabs 374877.2 6548165.7 -21.9 -79.6 156.9 104.9 108.3 0.5KD6037 386 Fl1 Channel BZ Geolabs 374859.9 6548173.0 -124.2 -80.0 153.4 1.0 4.3 0.5KD6037 443 Fl1 Channel BZ Geolabs 374850.7 6548177.6 -180.2 -37.6 76.7 21.9 23.9 6.5KD6037 462.8 Fl1 Channel BZ M Ultratrace 374847.3 6548179.2 -199.7 -79.0 155.0 4.2 6.2 5.6KD6037 462.8 Fl1 Channel BZ M Ultratrace 374847.3 6548179.2 -199.7 -79.0 155.0 4.2 6.2 5.6KD6039 173 Fl_n Channel Spfx Geolabs 374895.7 6548361.4 81.5 -67.8 57.4 277.4 280.3 2.0KD6039 466 Fl1 Channel BZ Geolabs 374900.8 6548379.4 -210.8 -84.8 208.0 2.9 17.3 0.5KD6039 502 Fl1 Channel BZ Geolabs 374902.4 6548382.5 -246.6 -10.0 158.2 19.7 23.2 0.6KD6041 376 Fl1 Channel BZ Geolabs 374920.5 6547961.2 -120.6 84.0 344.7 11.0 44.0 0.6KD6041W1 465 Fl1 Channel BZ Geolabs 374911.3 6547962.9 -209.1 83.4 345.0 0.6 1.5 0.4KD6041W1 503 Fl1 Channel BZ Geolabs 374906.8 6547964.0 -246.8 82.5 347.4 0.5 1.5 2.2KD6042AW1 594 Fl1 Channel BZ Geolabs 375022.2 6547939.1 -338.6 -85.1 173.6 9.9 11.3 0.8KD6042AW1 612 Fl1 Channel BZ Geolabs 375020.6 6547939.3 -356.5 85.1 345.8 2.0 2.9 1.0KD6042AW1 625.5 Fl1 Channel BZ M Ultratrace 375019.5 6547939.7 -370.0 -85.0 157.1 0.5 1.5 6.5KD6042AW1 625.5 Fl1 Channel BZ M Ultratrace 375019.5 6547939.7 -370.0 -85.0 157.1 0.5 1.5 6.5KD6048 699 Fl1 Channel BZ Geolabs 375177.1 6548738.2 -444.8 -85.3 231.9 77.9 91.9 0.5KD6048 784 Fl1 Channel BZ Geolabs 375182.7 6548742.6 -529.5 83.0 50.4 7.0 13.6 0.5KD6048 795 3 Fl1 Ch l BZ M Ult t 375183 8 6548743 5 540 7 83 0 234 3 3 6 8 3 0 7KD6048 795.3 Fl1 Channel BZ M Ultratrace 375183.8 6548743.5 -540.7 -83.0 234.3 3.6 8.3 0.7KD6051 694 Fl1 Channel Spfx Geolabs 375292.2 6547285.0 -440.0 -42.2 84.3 101.6 101.8 3.7KD6051 703 Fl1 Channel BZ Geolabs 375292.0 6547285.5 -449.0 -39.4 83.7 95.7 96.2 3.7KD6051 766 Fl1 Channel BZ Geolabs 375291.1 6547290.7 -511.7 -17.3 76.7 60.2 61.0 0.8KD6051 810 Fl1 Channel BZ Geolabs 375290.9 6547295.4 -555.5 -22.6 168.1 37.5 38.2 1.0KD6053A 634 Fl1 Channel Spfx Geolabs 375155.0 6547636.8 -379.9 -50.9 41.9 36.7 37.1 0.9KD6053A 663 Fl1 Channel BZ Geolabs 375154.2 6547639.0 -408.8 -40.7 60.0 21.5 23.1 1.3

A2


Distance (m)

Distance (m)

Ni (wt%)Collar

Depth (m) Flow #

KD6053A 693.1 Fl1 Channel BZ M Ultratrace 375154.0 6547641.6 -438.8 85.0 275.0 0.1 1.4 1.0KD6056 299 Fl1 Channel Spfx Geolabs 374817.3 6548456.5 -43.8 -58.1 284.8 177.6 185.9 2.5KD6056 365 Fl1 Channel BZ Geolabs 374809.3 6548457.2 -109.3 -60.4 308.6 125.8 133.6 2.5KD6056 388 Fl1 Channel BZ Geolabs 374806.4 6548457.7 -132.1 -53.4 310.3 107.9 118.7 2.5KD6061 243 Fl1 Channel Spfx Geolabs 374662.6 6548179.3 18.1 -68.7 289.4 57.3 151.9 0.5KD6061 271 Fl1 Channel BZ Geolabs 374660.9 6548180.6 -9.8 -48.3 292.2 34.0 139.7 0.5KD6061 294 Fl1 Channel BZ Geolabs 374659.8 6548182.1 -32.7 -5.9 293.5 24.6 135.0 0.5KD6067BW7 755 Fl1 Channel Spfx Geolabs 375237.2 6548335.5 -498.0 -80.0 211.2 100.1 102.1 0.7KD6067BW7 765 Fl1 Channel BZ Geolabs 375235.8 6548335.3 -507.9 -78.3 214.4 90.5 92.5 0.7KD6067BW7 806 Fl1 Channel BZ Geolabs 375230.1 6548334.5 -548.5 -64.8 225.2 53.1 54.3 0.5KD6067BW7 857 Fl1 Channel BZ M Ultratrace 375222.8 6548333.9 -599.0 -81.7 87.8 3.7 8.3 1.6KD6067BW7 857 Fl1 Channel BZ M Ultratrace 375222.8 6548333.9 -599.0 -81.7 87.8 3.7 8.3 1.6KD6071A 673 Fl1 Channel Spfx Geolabs 375231.0 6547465.1 -417.4 -52.3 163.0 51.2 56.2 0.7KD6071A 682 Fl1 Channel BZ Geolabs 375230.5 6547465.9 -426.3 -45.9 164.0 44.0 49.7 0.7KD6071A 726 Fl1 Channel BZ Geolabs 375227.8 6547470.0 -470.0 -83.7 120.9 16.9 19.5 0.7KD6071A 750.2 Fl1 Channel BZ M Ultratrace 375226.5 6547472.3 -494.1 -83.8 116.4 0.8 4.6 0.5KD6074 349 Fl1 Channel BZ Geolabs 374854.5 6547961.8 -94.9 -86.4 134.6 55.9 63.9 0.6KD6074 400 Fl1 Channel BZ Geolabs 374852.3 6547964.0 -145.8 -86.5 140.6 5.0 20.5 0.5KD6074 408.3 Fl1 Channel BZ M Ultratrace 374851.9 6547964.4 -154.0 86.5 321.7 3.3 18.4 0.5KD6083A 404 Fl1 Flank Spfx Geolabs 375069.1 6547261.6 -148.2 -65.0 222.4 353.8 354.3 0.9KD6083A 409 Fl1 Flank BZ Geolabs 375068.5 6547261.8 -153.2 -64.6 222.6 349.4 350.0 0.9KD6083A 422.3 Fl1 Flank BZ M Ultratrace 375066.7 6547262.2 -166.3 -63.5 223.3 338.0 338.5 0.9LNSD011 107 Geolabs 374416.7 6549332.7 151.9 -89.6 114.8 8.2 9.1 0.8VS15 019 24 Fl1 Fl k BZ G l b 375232 0 6547568 5 449 7 17 6 89 2 13 0 13 2 1 3VS15-019 24 Fl1 Flank BZ Geolabs 375232.0 6547568.5 -449.7 -17.6 89.2 13.0 13.2 1.3KD6012 295.4 Fl2 Channel BZ Geolabs 374977.3 6547269.0 -40.8 -49.7 210.2 452.5 453.4 0.6KD6012 307.2 Fl1 Channel Spfx Geolabs 374976.2 6547269.9 -52.5 -48.7 210.5 443.5 444.4 0.6KD6012 308.2 Fl1 Channel Spfx Geolabs 374976.1 6547270.0 -53.5 -48.6 210.5 442.7 443.6 0.6KD6012 312.7 Fl1 Channel BZ Geolabs 374975.7 6547270.3 -58.0 -48.2 210.6 439.3 440.2 0.6KD6012 318.7 Fl1 Channel BZ Geolabs 374975.2 6547270.7 -64.0 -47.7 210.8 434.8 435.7 0.6KD6012 322.9 Fl1 Channel BZ Geolabs 374974.9 6547271.0 -68.1 -47.4 210.9 431.7 432.6 0.6

A3


Distance (m)

Distance (m)

Ni (wt%)Collar

Depth (m) Flow #

KD6012 331 Fl1 Channel BZ Geolabs 374974.2 6547271.6 -76.2 -46.6 211.0 425.7 426.5 0.6KD6020 74.98 Fl1 Flank Spfx Geolabs 374500.3 6547985.8 180.3 -51.6 357.1 278.6 287.5 0.5KD6020 76.8 Fl1 Flank Spfx Geolabs 374500.2 6547985.9 178.5 -51.3 357.1 277.2 286.1 0.5KD6020 80.46 Fl1 Flank BZ Geolabs 374500.0 6547986.0 174.8 -50.8 357.0 274.5 283.3 0.5KD6024 170.5 Fl2 Flank BZ Geolabs 374742.3 6547626.7 83.7 -42.8 190.4 264.1 338.9 0.5KD6024 178.6 Fl2 Flank BZ Geolabs 374741.5 6547627.3 75.7 -41.5 190.7 258.3 334.5 0.5KD6024 196.3 Fl1 Flank Spfx Geolabs 374739.3 6547628.5 58.2 -38.6 191.4 246.5 325.6 0.5KD6024 198.4 Fl1 Flank BZ Geolabs 374739.0 6547628.6 56.1 -38.2 191.4 245.2 324.6 0.5KD6024 212.1 Fl1 Flank BZ Geolabs 374737.1 6547629.3 42.6 -35.7 192.1 236.8 318.6 0.5KD6026 78.2 Fl1 Flank Spfx Geolabs 374503.8 6547624.4 180.0 14.2 4.9 320.2 400.5 0.5KD6026 82.7 Fl1 Flank BZ Geolabs 374503.7 6547624.5 175.5 15.0 4.9 321.5 399.2 0.5KD6026 87.2 Fl1 Flank BZ M Ultratrace 374503.7 6547624.7 171.0 15.7 4.9 322.9 398.0 0.5KD6036 283.9 Fl1 Channel BZ Geolabs 374732.2 6548004.2 -28.6 -81.7 144.8 25.0 28.4 0.5KD6036 328 Fl1 Channel BZ Geolabs 374727.3 6548008.0 -72.3 82.0 316.2 11.0 13.0 0.6KD6037 382.2 Fl1 Channel BZ Geolabs 374860.5 6548172.7 -120.4 -80.0 153.8 4.7 8.1 0.5KD6037 401.4 Fl1 Channel BZ Geolabs 374857.6 6548174.3 -139.3 80.0 330.5 6.4 8.4 0.5KD6037 407.1 Fl1 Channel BZ Geolabs 374856.7 6548174.8 -144.9 80.0 330.8 12.1 14.1 0.5KD6037 414.8 Fl1 Channel BZ Geolabs 374855.5 6548175.4 -152.5 79.9 331.6 19.8 21.8 0.5KD6043 180 Fl1 Flank Spfx Geolabs 374682.2 6547969.1 74.3 -83.9 138.5 112.9 124.9 0.5KD6043 277 Fl1 Channel BZ Geolabs 374674.4 6547975.4 -22.1 -82.8 127.3 15.9 27.9 0.5KD6045 303.5 Fl1 Channel Spfx Geolabs 374764.9 6547969.5 -48.0 -82.0 163.3 19.4 22.1 0.5KD6045 313.1 Fl1 Channel BZ Geolabs 374763.6 6547969.9 -57.5 -82.0 164.4 9.8 12.5 0.5KD6049 480.1 Fl1 Channel Spfx Geolabs 375083.5 6547633.8 -226.0 -82.4 225.6 75.2 111.4 0.5KD6049 511 7 Fl1 Ch l BZ G l b 375082 8 6547635 9 257 5 78 0 238 2 44 0 98 6 0 4KD6049 511.7 Fl1 Channel BZ Geolabs 375082.8 6547635.9 -257.5 -78.0 238.2 44.0 98.6 0.4KD6049 544.7 Fl1 Channel BZ Geolabs 375082.1 6547638.6 -290.4 -49.3 255.3 13.4 67.0 0.4KD6054W1 857.2 Fl1 Flank BZ Geolabs 375487.5 6547293.4 -603.1 21.1 145.2 201.9 203.7 2.0KD6062A 783.1 Fl2 Channel BZ Geolabs 375275.4 6548782.6 -526.2 -81.0 217.3 23.8 55.0 0.5KD6062A 787.5 Fl1 Channel Spfx Geolabs 375275.8 6548783.2 -530.6 -81.0 217.6 19.4 51.6 0.5KD6062A 804.5 Fl1 Channel BZ Geolabs 375277.4 6548785.3 -547.4 -81.0 218.9 2.4 39.2 0.5KD6066 619.1 Fl1 Flank Spfx Geolabs 375101.3 6548711.8 -364.2 -61.4 249.4 180.2 191.9 0.5

A4


Distance (m)

Distance (m)

Ni (wt%)Collar

Depth (m) Flow #

KD6066 622.4 Fl1 Flank Spfx Geolabs 375101.1 6548711.8 -367.5 -60.9 249.5 177.4 189.0 0.5KD6066 624.4 Fl1 Flank BZ Geolabs 375101.0 6548711.7 -369.5 -60.5 249.5 175.7 187.3 0.5KD6066 636.6 Fl1 Flank BZ Geolabs 375100.4 6548711.6 -381.7 -58.2 249.5 165.5 176.8 0.5KD6068 710.9 Fl_n breccia Geolabs 375741.7 6547310.8 -454.7 -83.8 211.7 4.0 54.0 0.9KD6068 847.8 Fl_n BZ Geolabs 375749.9 6547325.0 -590.6 83.0 30.0 48.8 50.8 1.3KD6068 871.8 Fl_n BZ Geolabs 375751.3 6547327.5 -614.4 83.0 30.0 72.8 74.8 1.3KD6068AW2 826 Fl_n Spfx Geolabs 375712.7 6547314.5 -571.2 40.1 261.1 45.2 46.4 1.3KD6069AW1 616.7 Fl1 Channel BZ Geolabs 375123.6 6547997.2 -361.1 -11.1 48.8 61.2 62.9 1.3KD6069AW1 635.6 Fl1 Channel BZ Geolabs 375125.6 6547998.4 -379.8 6.4 49.2 62.7 64.7 1.3KD6070 280.3 Fl1 Flank Spfx Geolabs 374852.3 6547817.7 -25.5 -43.0 89.7 102.7 154.9 0.6KD6070 286.3 Fl1 Flank BZ Geolabs 374851.6 6547817.9 -31.4 -40.8 89.6 98.2 151.9 0.6KD6070 295.1 Fl1 Flank BZ Geolabs 374850.4 6547818.1 -40.1 -37.2 89.4 91.8 147.9 0.6KD6074 394.7 Fl1 Channel BZ Geolabs 374852.6 6547963.8 -140.5 -86.5 139.9 10.2 24.0 0.5KD6074 404.8 Fl1 Channel BZ Geolabs 374852.1 6547964.2 -150.5 -86.5 141.2 0.2 17.8 0.5KD6082 290 Fl2 Flank BZ Geolabs 374902.7 6547618.7 -35.8 -14.3 122.3 242.4 261.2 0.5KD6082 320.7 Fl1 Flank BZ Geolabs 374900.6 6547619.3 -66.4 -7.1 121.9 235.2 245.1 0.5KD6083A 391.3 Fl2 Flank BZ Geolabs 375070.9 6547261.1 -135.6 -66.0 221.8 364.9 365.3 0.8KD6083A 403.4 Fl1 Flank Spfx Geolabs 375069.2 6547261.6 -147.6 -65.1 222.4 354.3 354.9 0.9KD6083A 404.8 Fl1 Flank Spfx Geolabs 375069.0 6547261.6 -149.0 -65.0 222.4 353.1 353.6 0.9KD6083A 405.5 Fl1 Flank BZ Geolabs 375068.9 6547261.6 -149.7 -64.9 222.5 352.5 353.0 0.9KD6083A 407.8 Fl1 Flank BZ Geolabs 375068.6 6547261.7 -152.0 -64.7 222.6 350.5 351.0 0.9KD6083A 413.7 Fl1 Flank BZ Geolabs 375067.9 6547261.9 -157.8 -64.3 222.9 345.4 345.9 0.9KD6083A 416.9 Fl1 Flank BZ Geolabs 375067.4 6547262.0 -161.0 -64.0 223.0 342.6 343.2 0.9KD6084 691 7 Fl1 Fl k S f G l b 375282 0 6547998 3 437 5 42 2 182 8 214 1 215 6 0 7KD6084 691.7 Fl1 Flank Spfx Geolabs 375282.0 6547998.3 -437.5 -42.2 182.8 214.1 215.6 0.7KD6084 741.4 Fl1 Flank BZ Geolabs 375281.2 6547999.5 -487.2 -30.9 183.1 183.4 185.0 0.7KD6084 804.1 Fl1 Flank BZ Geolabs 375281.6 6548001.8 -549.8 -11.5 183.0 158.2 159.8 0.7KD6093 236.8 Fl1 Flank Spfx Geolabs 374791.3 6547620.4 17.0 -29.7 94.1 227.3 316.7 0.5KD6093 241.2 Fl1 Flank BZ Geolabs 374791.3 6547620.4 12.6 -28.7 94.1 225.2 315.3 0.5KD6093 247.2 Fl1 Flank BZ Geolabs 374791.2 6547620.4 6.6 -27.4 94.0 222.3 313.4 0.5KD6093 252.1 Fl1 Flank BZ Geolabs 374791.1 6547620.4 1.7 -26.3 94.0 220.1 311.5 0.5

A5


Distance (m)

Distance (m)

Ni (wt%)Collar

Depth (m) Flow #

KD6168 128.6 Fl2 Flank BZ Geolabs 374609.7 6547625.3 130.2 -41.5 221.1 340.5 365.5 0.5KD6168 129.7 Fl1 Flank Spfx Geolabs 374609.4 6547625.3 129.1 -41.4 221.2 339.9 365.2 0.5KD6168 133 Fl1 Flank BZ Geolabs 374608.5 6547625.5 126.0 -40.9 221.3 338.2 364.5 0.5KD6168 137.3 Fl1 Flank BZ Geolabs 374607.3 6547625.6 121.8 -40.3 221.6 336.0 363.7 0.5KD6169 370.2 Fl1 Channel BZ Geolabs 374817.4 6547989.8 -113.8 -40.6 323.7 56.8 64.1 0.7LG16-76 426.2 Fl1 Flank BZ Geolabs 375257.5 6549077.9 -737.6 31.9 174.0 225.3 227.4 0.7LG16-76 439.5 Fl1 Flank BZ Geolabs 375269.2 6549077.5 -743.9 31.7 174.3 238.6 240.6 0.7LG7-149 137.6 Fl1 Flank BZ Geolabs 374710.9 6549113.0 6.3 2.1 57.1 49.3 49.4 0.5LG7-150 129.8 Fl1 Flank BZ Geolabs 374699.7 6549103.3 -62.7 -10.5 64.2 38.7 46.9 0.6LNSD-017 1004.8 Fl2 Flank BZ Geolabs 375405.0 6548016.1 -736.7 80.8 345.3 53.8 55.8 0.5LSU-001 790.5 Fl1 Flank BZ Geolabs 375432.4 6548301.1 -859.6 51.0 41.9 127.8 129.5 0.7LSU-001W2 599.8 Fl1 Flank BZ Geolabs 375321.0 6548226.8 -720.2 -34.0 255.8 59.6 62.3 0.8LSU-001W2 606.8 Fl1 Flank Spfx Geolabs 375326.6 6548228.3 -724.2 -34.0 255.9 52.6 55.3 0.8LSU-001W2 663.5 Fl2 Flank BZ Geolabs 375372.2 6548239.7 -755.9 -33.9 256.8 2.0 2.6 1.0LSU-001W2 680.4 Fl3 Flank BZ Geolabs 375385.9 6548243.0 -765.2 33.6 76.4 8.9 10.9 0.7LSU-012 221 Fl1 Flank BZ Geolabs 375278.9 6548959.3 -653.1 29.4 80.0 157.7 163.8 0.5LSU-012 223.3 Fl1 Flank Spfx Geolabs 375280.7 6548960.5 -654.0 29.3 79.7 159.9 166.0 0.5LSU-143 542.3 Fl1 Channel BZ Geolabs 375441.3 6547659.1 -674.1 36.1 75.4 230.6 231.7 0.6LSU-143 573.3 Fl1 Channel BZ Geolabs 375452.4 6547630.2 -675.7 35.6 84.6 236.4 238.3 0.6VS15-150 157 Fl1 Flank BZ Geolabs 375332.8 6547666.3 -485.1 20.3 84.0 74.4 93.1 0.6VS15-150 168.4 Fl1 Flank Spfx Geolabs 375344.0 6547666.0 -487.1 18.9 85.0 85.6 102.8 0.6

A6

Table A.2. Location and description of samples from Maggie Hays, Lake Johnston GSB, Western Australia

Maggie Hays, Lake Johnston GSB, Western AustraliaDescription Classification Lab X Y Z Ni (wt%)

FGD91-7 318 CUU Massive Ultramafic Cumulate Ultratrace 265024.2 6430244.1 1115.8 10.2 0.8FGD92-8 177 CUU Massive Ultramafic Cumulate Ultratrace 265619.2 6429212.1 1204.2 202.8 0.8LJD0003A 231 CUU Massive Ultramafic Felsic-UM Ultratrace 264702.5 6430258.6 1161.4 102.0 0.9LJD0003A 300.5 CUU Massive Ultramafic Felsic-UM Ultratrace 264739.9 6430271.1 1104.3 93.2 0.9LJD0003A 501 CUU Massive Ultramafic Cumulate Ultratrace 264850.6 6430307.3 941.1 56.2 1.4LJD0004 332 CUU Massive Ultramafic Cumulate Ultratrace 265048.1 6429870.0 1085.8 191.7 0.5LJD0004 361.7 CUU Massive Ultramafic Cumulate Ultratrace 265063.4 6429875.9 1061.1 162.5 0.5LJD0004 413 CUU Massive Ultramafic Cumulate Ultratrace 265089.8 6429885.7 1018.2 113.1 0.5LJD0004 432 CUU Massive Ultramafic Cumulate Ultratrace 265099.4 6429889.3 1002.2 95.3 0.5LJD0004 520.5 CUU Massive Ultramafic Cumulate Ultratrace 265143.8 6429906.1 927.5 37.4 0.4LJD0005 260.5 CUU Massive Ultramafic Cumulate Ultratrace 265120.6 6429673.1 1148.2 320.6 0.5LJD0005 332.7 CUU Massive Ultramafic Cumulate Ultratrace 265162.7 6429689.9 1092.0 267.1 0.5LJD0005 356.2 CUU Massive Ultramafic Cumulate Ultratrace 265176.2 6429695.3 1073.5 251.6 0.5LJD0005 384.5 CUU Massive Ultramafic Cumulate Ultratrace 265192.3 6429701.7 1051.2 234.7 0.5LJD0005 414 CUU Massive Ultramafic Cumulate Ultratrace 265209.1 6429708.4 1027.9 219.5 0.5LJD0009 771 Felsic Vol Felsic Felsic Ultratrace 265034 3 6429961 9 702 7 193 9 0 4

Distance (m)

Sample identification given based on collar name and depth. Stratigraphic units are CUU = Central ultramafic unit, WUU = Western ultramafic unit, BIF-sill = BIF hosted sill. Sample location given as X, Y, Z UTM coordinates as calculated. Z (depth) is relative to a local mine datum. Closest occurrence of Ni are calculated distances, directions and grades based on the 3 closest occurrences of Ni >0.4%. Analytical lab used is indicated for each sample (Lab: Geolabs or Ultratrace) refer to Appendix C for additional information on quality control and quality assurances.

UTM MGA94 Z51Ave of 3 closest

occurrences

Collar Depth (m) Strat. Unit

LJD0009 771 Felsic Vol Felsic Felsic Ultratrace 265034.3 6429961.9 702.7 193.9 0.4LJD0010A 371.5 CUU Massive Ultramafic Cumulate Ultratrace 265246.1 6429830.8 1043.9 149.9 0.5LJD0010A 425.6 CUU Massive Ultramafic Cumulate Ultratrace 265221.5 6429821.0 996.7 120.3 0.5LJD0010A 538.4 CUU Massive Ultramafic Cumulate Ultratrace 265172.3 6429801.5 897.1 91.1 0.6LJD0011 464.6 CUU Massive Ultramafic Cumulate Ultratrace 264908.6 6430363.1 974.7 5.6 0.5LJD0011 479 CUU Massive Ultramafic Cumulate Ultratrace 264900.4 6430360.3 963.2 11.6 0.6LJD0011 713 BIF Sill Chill zone Chill Ultratrace 264768.8 6430315.8 774.9 134.2 0.7

A7


Distance (m)Collar Depth (m) Strat. Unit

LJD0011 716 BIF Sill Massive Ultramafic Cumulate Ultratrace 264767.0 6430315.2 772.6 137.0 0.7LJD0011 748 BIF Sill Massive Ultramafic Cumulate Ultratrace 264748.5 6430308.9 747.2 167.8 0.7LJD0011 752.1 BIF Sill Chill zone Chill Ultratrace 264746.2 6430308.1 744.0 171.7 0.7LJD0011 793.5 WUU A2-Zone top chill Chill Ultratrace 264722.3 6430300.0 711.1 211.6 0.9LJD0011 801 WUU A2-Zone top chill Chill Ultratrace 264717.9 6430298.6 705.2 218.9 0.9LJD0015 344 CUU Massive Ultramafic Cumulate Ultratrace 265093.5 6429772.5 1073.2 208.2 0.5LJD0015 398 CUU Massive Ultramafic Cumulate Ultratrace 265120.6 6429783.3 1027.7 165.8 0.5LJD0015 479 CUU Massive Ultramafic Cumulate Ultratrace 265159.3 6429798.7 958.3 116.1 0.5LJD0017 183 CUU Massive Ultramafic Cumulate Ultratrace 265766.8 6428961.8 1203.0 460.9 0.6LJD0017 210.6 CUU Border phase Chill zone Ultratrace 265753.9 6428956.7 1179.1 459.1 0.7LJD0017 222 Felsic Vol Felsic Felsic Ultratrace 265748.6 6428954.6 1169.3 458.7 0.7LJD0018 174.8 CUU Massive Ultramafic Cumulate Ultratrace 265402.9 6429347.0 1219.2 140.8 0.9LJD0018 225.5 CUU Massive Ultramafic Cumulate Ultratrace 265432.2 6429358.6 1179.5 109.3 0.9LJD0018 265.6 Felsic Vol Felsic Felsic Ultratrace 265454.8 6429367.6 1147.6 97.4 0.9LJD0048 162.3 WUU B-zone cumulate Cumulate Ultratrace 264228.0 6430726.9 1202.1 330.4 0.6LJD0048 164.05 WUU A-zone spinifex Chill Ultratrace 264227.2 6430726.7 1200.6 331.1 0.6LJD0048 164.3 WUU A-zone spinifex top Chill Ultratrace 264227.1 6430726.6 1200.4 331.2 0.6LJD0048 189.1 WUU A-zone spinifex top Chill Ultratrace 264215.3 6430723.3 1178.8 341.3 0.6LJD0051 84.6 CUU Massive Ultramafic Felsic-UM Ultratrace 264938.3 6429924.7 1281.4 281.3 0.5LJD0051 135 CUU Massive Ultramafic Felsic-UM Ultratrace 264961.3 6429934.0 1237.6 238.8 0.5LJD0051 224 CUU Massive Ultramafic Felsic-UM Ultratrace 265001.5 6429949.4 1159.6 175.1 0.5LJD0051 272 CUU Border phase Border Ultratrace 265023.2 6429957.1 1117.5 152.1 0.5LJD0051 289 CUU Massive Ultramafic Cumulate Ultratrace 265030.9 6429959.7 1102.6 146.2 0.5LJD0051 344 CUU Massive Ultramafic Cumulate Ultratrace 265055.9 6429968.1 1054.4 114.9 0.5LJD0051 441.8 CUU Massive Ultramafic Cumulate Ultratrace 265099.7 6429981.4 967.9 41.1 0.6LJD0052 100.5 CUU Massive Ultramafic Cumulate Ultratrace 264944.3 6430044.8 1271.4 223.1 0.5LJD0052 167.5 CUU Massive Ultramafic Cumulate Ultratrace 264973.9 6430056.6 1212.4 161.6 0.5LJD0052 228.8 CUU Massive Ultramafic Cumulate Ultratrace 265000.9 6430067.3 1158.5 111.2 0.5LJD0052 320 CUU Massive Ultramafic Cumulate Ultratrace 265040.6 6430083.1 1077.9 56.3 0.6LJD0052 381.2 CUU Massive Ultramafic Cumulate Ultratrace 265066.8 6430093.5 1023.6 17.8 0.5

A8



LJD0052 419 CUU Massive Ultramafic Cumulate Ultratrace 265082.9 6430099.9 990.0 19.3 0.5LJD0052 425 CUU Border phase Border Ultratrace 265085.4 6430100.9 984.6 18.7 0.5LJD0054A 248 CUU Massive Ultramafic Cumulate Ultratrace 264892.7 6430031.9 1141.5 202.1 0.7LJD0054A 333 CUU Massive Ultramafic Cumulate Ultratrace 264932.0 6430047.5 1067.8 150.6 0.5LJD0054A 372.5 CUU Massive Ultramafic Cumulate Ultratrace 264949.9 6430054.7 1033.3 124.8 0.5LJD0057 150.7 CUU Massive Ultramafic Cumulate Ultratrace 265092.4 6429889.2 1228.6 224.5 0.5LJD0057 223.5 CUU Massive Ultramafic Cumulate Ultratrace 265128.8 6429903.6 1167.2 189.7 0.5LJD0057 273.5 CUU Massive Ultramafic Cumulate Ultratrace 265153.3 6429913.4 1124.8 177.6 0.5LJD0057 345.8 CUU Massive Ultramafic Cumulate Ultratrace 265187.7 6429927.1 1062.7 115.8 0.5LJD0061 93 CUU Massive Ultramafic Cumulate Ultratrace 265203.9 6429710.6 1271.0 261.9 0.7LJD0061 181.5 CUU Massive Ultramafic Cumulate Ultratrace 265235.6 6429725.8 1189.8 212.9 0.7LJD0061 263 CUU Massive Ultramafic Cumulate Ultratrace 265264.5 6429739.7 1114.8 194.7 0.7LJD0061 347.5 CUU Massive Ultramafic Cumulate Ultratrace 265294.1 6429754.0 1037.0 175.8 0.6LJD0066 130 CUU Massive Ultramafic Cumulate Ultratrace 265075.7 6429983.0 1247.1 164.1 0.6LJD0066 304.2 CUU Massive Ultramafic Cumulate Ultratrace 265157.8 6430019.0 1097.7 87.9 0.6LJD0068 125.7 CUU Massive Ultramafic Cumulate Ultratrace 266829.7 6428307.5 1248.2 1437.0 0.6LJD0068 218.6 CUU Massive Ultramafic Cumulate Ultratrace 266870.6 6428325.4 1166.8 1457.6 0.6LJD0068 291 CUU Massive Ultramafic Cumulate Ultratrace 266901.5 6428339.0 1102.7 1477.1 0.6LJD0068 324.4 CUU Border phase Chill zone/fels Ultratrace 266915.2 6428345.0 1072.9 1487.0 0.6LJD0069 100.3 CUU Massive Ultramafic Cumulate Ultratrace 265854.3 6428789.1 1284.8 542.0 0.8LJD0069 173 CUU Massive Ultramafic Cumulate Ultratrace 265884.2 6428803.8 1220.1 530.8 0.8LJD0070 84.2 CUU Massive Ultramafic Cumulate Ultratrace 266007.3 6428626.9 1302.8 663.3 0.8LJD0070 156.2 CUU Massive Ultramafic Cumulate Ultratrace 266037.5 6428638.9 1238.5 652.4 0.6LJD0070 224 CUU Border phase Chill zone Ultratrace 266065.5 6428650.1 1177.8 653.4 0.6LJD0071 98 CUU Massive Ultramafic Cumulate Ultratrace 265384.4 6429567.7 1275.6 174.9 0.6LJD0071 193 CUU Massive Ultramafic Cumulate Ultratrace 265428.2 6429585.1 1193.1 99.3 0.9LJD0071 242.5 CUU Massive Ultramafic Cumulate Ultratrace 265451.0 6429594.2 1150.1 55.8 0.9LJD0071 299.5 CUU Massive Ultramafic Cumulate Ultratrace 265476.2 6429604.2 1100.0 38.4 0.9LJD0077 190 CUU Massive Ultramafic Cumulate Ultratrace 265398.2 6429673.3 1190.6 56.2 0.7LJD0077 345.1 CUU Border phase Chill zone Ultratrace 265466.1 6429700.3 1053.9 10.1 0.7

A9



LJD0079 147 CUU Massive Ultramafic Cumulate Ultratrace 265334.8 6429754.2 1231.3 138.7 0.7LJD0079 210 CUU Massive Ultramafic Cumulate Ultratrace 265361.7 6429764.9 1175.3 108.5 0.7LJD0079 245 CUU Massive Ultramafic Cumulate Ultratrace 265376.4 6429770.8 1144.1 105.1 0.8LJD0079 259 CUU Border phase Border Ultratrace 265382.3 6429773.1 1131.6 106.5 0.8LJD0080 142 CUU Massive Ultramafic Cumulate Ultratrace 265198.3 6429812.6 1233.0 280.2 0.7LJD0080 199.5 CUU Massive Ultramafic Cumulate Ultratrace 265223.7 6429822.7 1182.4 255.7 0.7LJD0080 261 CUU Massive Ultramafic Cumulate Ultratrace 265250.3 6429833.3 1128.0 214.1 0.5LJD0081 79.5 CUU Massive Ultramafic Cumulate Ultratrace 265132.7 6429909.8 1292.1 245.2 0.6LJD0081 161.5 CUU Massive Ultramafic Cumulate Ultratrace 265172.1 6429925.5 1221.9 199.4 0.6LJD0081 205 CUU Massive Ultramafic Cumulate Ultratrace 265192.6 6429933.6 1184.4 185.5 0.6LJD0086 106.7 CUU Massive Ultramafic Cumulate Ultratrace 265395.2 6429344.6 1262.3 160.6 0.9LJD0086 162.3 CUU Border phase Chill zone Ultratrace 265369.3 6429337.8 1213.5 174.5 0.9LJD0088 156.7 CUU Massive Ultramafic Cumulate Ultratrace 265510.8 6429397.8 1222.9 33.6 0.9LJD0088 198.5 CUU Massive Ultramafic Cumulate Ultratrace 265491.6 6429390.2 1186.5 43.5 0.9LJD0088 237 CUU Massive Ultramafic Cumulate Ultratrace 265474.3 6429383.4 1152.8 75.3 0.9LJD0104W1 126.7 CUU Massive Ultramafic Cumulate Ultratrace 265010.2 6430139.3 1252.3 131.3 0.6LJD0104W1 212 CUU Massive Ultramafic Cumulate Ultratrace 265040.6 6430152.6 1173.9 55.0 0.5LJD0107 392 CUU Massive Ultramafic Cumulate Ultratrace 264992.5 6430138.5 1025.4 38.1 0.5LJD0107 445 CUU Massive Ultramafic Cumulate Ultratrace 265013.4 6430147.6 977.6 24.4 0.6LJD0120 229.6 CUU Massive Ultramafic Cumulate Ultratrace 264862.6 6430213.4 1149.0 101.5 1.1LJD0120 229.6 CUU Massive Ultramafic Cumulate Ultratrace 264862.5 6430213.4 1148.9 101.5 1.1LJD0120 257 CUU Massive Ultramafic Cumulate Ultratrace 264851.0 6430208.8 1124.6 91.3 0.9LJD0120 284 CUU Massive Ultramafic Cumulate Ultratrace 264839.8 6430204.3 1100.4 84.5 0.9LJD0120 311.5 CUU Massive Ultramafic Cumulate Ultratrace 264828.6 6430199.9 1075.7 89.5 0.9LJD0124 130 CUU Massive Ultramafic Cumulate Ultratrace 266419.2 6428469.3 1259.2 1022.1 0.6LJD0124 164.21 CUU Border phase Chill zone Ultratrace 266434.5 6428475.3 1229.2 1027.0 0.6LJD0126 81.5 WUU A-zone spinifex Chill Ultratrace 266929.1 6428024.1 1275.4 1685.6 0.6LJD0126 119.2 WUU Breccia/spinifex Chill Ultratrace 266912.7 6428017.6 1242.1 1675.2 0.6LJD0126 157.5 WUU Breccia Chill Ultratrace 266896.1 6428010.9 1208.2 1665.4 0.6LJD0126 189.5 WUU Fragmental Chill Ultratrace 266882.3 6428005.4 1179.9 1657.9 0.6

A10



LJD0126 313.1 WUU Fragmental Chill Ultratrace 266828.6 6427984.1 1070.6 1634.9 0.6LJPD0094 101.7 CUU Massive Ultramafic Cumulate Ultratrace 265941.7 6428664.9 1259.8 601.8 0.6LJPD0094 265.4 CUU Border phase Chill Ultratrace 265934.5 6428589.4 1114.7 611.7 0.6MHD94-3 122 CUU Massive Ultramafic Cumulate Ultratrace 264807.9 6430194.1 1259.6 192.7 0.9MHD94-3 196.5 CUU Massive Ultramafic Cumulate Ultratrace 264841.6 6430206.2 1194.2 135.6 1.1MHD94-3 253 CUU Massive Ultramafic Cumulate Ultratrace 264867.0 6430215.3 1144.6 97.6 1.1MHD94-3 287.5 CUU Massive Ultramafic Cumulate Ultratrace 264882.4 6430220.8 1114.2 84.6 0.8MHD94-3 375 CUU Massive Ultramafic Cumulate Ultratrace 264920.4 6430233.9 1036.5 37.7 0.5MHD94-3 440 CUU Massive Ultramafic Cumulate Ultratrace 264947.3 6430243.7 978.1 3.0 0.5MHD94-5 243 CUU Massive Ultramafic Cumulate Ultratrace 265037.1 6429965.2 1155.9 142.3 0.5MHD94-5 428 CUU Massive Ultramafic Cumulate Ultratrace 265130.8 6430002.5 1000.8 72.3 0.6LJD0077 295.6 CUU Massive Ultramafic Cumulate Ultratrace 265444.87 6429691.9 1097.7 24.7 0.6LJD0120 344.2 CUU Massive Ultramafic Cumulate Ultratrace 264815.29 6430194.6 1046.3 99.9 0.9LJD0048 159.4 WUU A-zone spinifex Ultratrace 264229.38 6430727.3 1204.6 329.4 0.6LJD0010A 425.6 CUU Massive Ultramafic Cumulate Ultratrace 265221.5 6429821.0 996.6 120.2 0.5LJD0048 159.7 WUU A-zone spinifex Ultratrace 264229.24 6430727.3 1204.4 329.5 0.6LJD0048 171.3 WUU A-zone spinifex Ultratrace 264223.76 6430725.7 1194.3 333.9 0.6LJD0048 156 WUU B-zone cumulate Ultratrace 264230.99 6430727.7 1207.6 328.2 0.6LJD0011 785.6 WUU A-zone spinifex Ultratrace 264726.83 6430301.6 717.4 204.0 0.9LJD0069 238 CUU Border phase Pyroxenite Ultratrace 265910.56 6428816.5 1162.1 528.3 0.6LJD0011 780.2 WUU B-zone cumulate Ultratrace 264729.89 6430302.6 721.6 198.9 0.9

A11

TF = thin flow, MF = massive flow, FRG = fragmental textured, PF = pillowed flow, A1 = A1 spinifexKarasjok Greenstone Belt

Sample Area Morphology Sample type Lab Lat LongWP-44 Nilivaara TF surface grab Ultratrace 68.11815 24.50947WP-45 Nilivaara MF surface grab Ultratrace 68.11914 24.50507WP-46 Nilivaara TF surface grab Ultratrace 68.12009 24.50447WP-47 Nilivaara FRG surface grab Ultratrace 68.11795 24.50533WP-48 Nilivaara MF surface grab Ultratrace 68.11801 24.50417WP-49 Nilivaara MF surface grab Ultratrace 68.11759 24.50401WP-50 Nilivaara MF surface grab Ultratrace 68.11764 24.5041WP-51 Nilivaara MF surface grab Ultratrace 68.11599 24.49681WP-52 Nilivaara A1 surface grab Ultratrace 68.11578 24.49693WP-53 Hotinvaara MF surface grab Ultratrace 68.08929 24.42158WP-54 Hotinvaara MF surface grab Ultratrace 68.08776 24.41607WP-55 Hotinvaara MF surface grab Ultratrace 68.08955 24.4118WP-56 Hotinvaara MF surface grab Ultratrace 68.09171 24.41275WP-57 Hotinvaara gabbro surface grab Ultratrace 68.0917 24.41277WP-58 Hotinvaara gabbro surface grab Ultratrace 68.09175 24.41155WP-59 Sarvisoaivi MF surface grab Ultratrace 68.63982 21.90222WP-60 Sarvisoaivi MF surface grab Ultratrace 68.6398 21.90015WP-61 Sarvisoaivi MF surface grab Ultratrace 68.63962 21.90009WP-62 Sarvisoaivi MF surface grab Ultratrace 68.63989 21.89256WP-63 Sarvisoaivi MF surface grab Ultratrace 68.63686 21.89952WP 64 S i i i MF f b Ult t 68 63373 21 9078

Table A.3. Decription and location of samples from the Karasjok Greenstone Belt, Finland and Norway

WP-64 Sarvisoaivi MF surface grab Ultratrace 68.63373 21.9078WP-65 Sarvisoaivi MF surface grab Ultratrace 68.63372 21.90821WP-66 Sarvisoaivi MF surface grab Ultratrace 68.63335 21.90921WP-67 Sarvisoaivi TF surface grab Ultratrace 68.63202 21.91375WP-68 Sarvisoaivi TF surface grab Ultratrace 68.632 21.91411WP-69 Sarvisoaivi TF surface grab Ultratrace 68.63561 21.91635WP-70 Sarvisoaivi TF surface grab Ultratrace 68.63574 21.91681

A12

Karasjok Greenstone BeltSample Area Morphology Sample type Lab Lat LongWP-71 Sarvisoaivi TF surface grab Ultratrace 68.63577 21.91715WP-72 Sarvisoaivi MF surface grab Ultratrace 68.63768 21.91367WP-73 Sarvisoaivi MF surface grab Ultratrace 68.63777 21.91367WP-74 Sarvisoaivi MF surface grab Ultratrace 68.63774 21.91368WP-75 Karasjok TF surface grab Ultratrace 70.04265 25.10507WP-76 Karasjok TF surface grab Ultratrace 70.04268 25.105WP-77 Karasjok PF surface grab Ultratrace 70.04252 25.10551WP-78 Karasjok PF surface grab Ultratrace 70.04077 25.11119WP-79 Karasjok PF surface grab Ultratrace 70.04029 25.11142WP-80 Karasjok PF surface grab Ultratrace 70.03971 25.11139WP-81 Karasjok PF surface grab Ultratrace 70.03971 25.11138WP-82 Karasjok PF surface grab Ultratrace 70.03906 25.1082WP-83 Karasjok PF surface grab Ultratrace 70.03894 25.10805WP-84 Karasjok PF surface grab Ultratrace 70.03227 25.12266WP-85 Karasjok PF surface grab Ultratrace 70.03217 25.12237WP-86 Karasjok PF surface grab Ultratrace 70.03298 25.12109WP-87 Karasjok FRG surface grab Ultratrace 70.03309 25.12059WP-88 Karasjok FRG surface grab Ultratrace 70.03311 25.12051WP-89 Karasjok gabbro surface grab Ultratrace 70.03352 25.11921WP-90 Karasjok gabbro surface grab Ultratrace 70.03323 25.11882WP-91 Karasjok TF surface grab Ultratrace 70.03085 25.07221WP-92 Karasjok TF surface grab Ultratrace 70.03083 25.07203WP-93 Karasjok TF surface grab Ultratrace 70.03083 25.07208WP-94 Karasjok TF surface grab Ultratrace 70 03039 25 07303WP-94 Karasjok TF surface grab Ultratrace 70.03039 25.07303

A13

Long-Victor

Long-Victor

Notes: XRF = X-ray florescence, ICP-MS = Inductively coupled plasma mass spectrometry, FA-ICP-MS = Fire assay inductively coupled plasma mass spectrometry, D.L. = analytical reported detection limit, N.D. = not determined, wt% = weight percent, ppm = parts per million, ppb = parts per billion.

B1

Long-Victor

SampleKD6012-

295.4KD6012-

307.2KD6012-

308.2KD6012-

312.7KD6012-

318.7KD6012-

322.9KD6012-

331Lab Geolabs Geolabs Geolabs Geolabs Geolabs Geolabs Geolabs

Units DL Batch 07-0521 07-0521 07-0521 07-0521 07-0521 07-0521 07-0521wt% 0.01 Al2O3 2.76 9.96 9.19 4.5 2.1 2.11 3.03wt% 0.01 CaO 0.75 9.52 10.53 3.91 0.45 0.33 5.58wt% 0.01 Fe2O3 9.17 10.86 13.12 10.91 8.04 7.33 8.66wt% 0.01 K2O <0.01 2.06 2.53 0.02 0.01 <0.01 0.01wt% 0.05 LOI 14.28 1.73 2.08 13.8 14.23 20.52 16.91wt% 0.01 MgO 34.07 13.31 14.76 29.18 35.75 34.83 29.56wt% 0.01 MnO 0.1 0.18 0.2 0.15 0.07 0.12 0.19wt% 0.01 Na2O 0.03 1.97 0.87 0.04 0.03 0.02 0.03wt% 0.01 P2O5 0.01 0.03 0.04 0.02 0.01 0.01 0.01wt% 0.01 SiO2 38.92 50.05 46.93 37.6 39.97 34.82 35.93wt% 0.01 TiO2 0.15 0.46 0.45 0.21 0.12 0.11 0.15

Total 100.23 100.14 100.69 100.35 100.77 100.19 100.06ppm 0.9 Ba <0.9 602.3 732.4 4.2 2.9 <0.9 0.9ppm 0.06 Be 0.21 0.5 0.31 0.06 0.18 0.1 0.2ppm 0.009 Bi 0.483 0.165 0.202 0.112 1.077 0.888 1.091ppm 0.01 Cd 0.03 0.08 0.1 0.17 0.02 0.02 0.04ppm 0.2 Ce 1.1 3.3 3.2 1.1 0.5 0.5 0.9ppm 0.1 Co 74.4 51.2 80 106.8 94.8 91.5 104.4ppm 24 Cr >600 >600 >600 >600 >600 >600 >600ppm 0.006 Cs 0.861 26.489 35.834 0.925 1.015 0.201 0.47ppm 2 Cu 51 6 31 54 9 35 63ppm 0.02 Dy 0.5 1.9 1.7 0.7 0.3 0.3 0.5ppm 0.02 Er 0.27 1.15 1.07 0.43 0.2 0.16 0.31ppm 0.005 Eu 0.106 0.532 0.478 0.151 0.073 0.05 0.161ppm 0.05 Ga 2.82 11.33 9.62 4.4 2.16 2.04 3.13ppm 0.02 Gd 0.38 1.4 1.32 0.56 0.25 0.22 0.4ppm 0.09 Hf 0.25 0.79 0.71 0.31 0.16 0.14 0.23ppm 0.003 Ho 0.094 0.401 0.381 0.157 0.069 0.057 0.116ppm 0.09 La 0.38 1.25 1.13 0.38 0.15 0.16 0.4ppm 0.2 Li 0.7 53.8 81.3 1.9 1 0.2 0.6ppm 0.002 Lu 0.041 0.181 0.166 0.069 0.034 0.025 0.053ppm 0.03 Mo 0.51 0.79 7.7 0.18 0.12 0.07 0.13ppm 0.04 Nb 0.27 0.76 0.62 0.3 0.17 0.15 0.25ppm 0.08 Nd 0.9 2.85 2.9 1.04 0.48 0.47 0.77ppm 3 Ni 1553 114 252 537 >2000 >2000 >2000ppm 0 Pb 2 4 3 2 2 2 2ppm 0.02 Pr 0.16 0.53 0.51 0.17 0.08 0.08 0.14ppm 0.2 Rb 0.3 78 97.4 0.8 0.3 <0.2 0.2ppm 0.04 Sb 0.41 0.37 0.48 0.32 0.47 0.26 0.3ppm 0 Sc 11.2 35 36.1 18.4 10.7 8.8 13.1ppm 0.02 Sm 0.29 1.1 1.03 0.41 0.2 0.16 0.26ppm 0.08 Sn 0.26 1.51 0.75 0.18 0.12 0.13 0.12

XRF

ICP

-MS

ppm 0.08 Sn 0.26 1.51 0.75 0.18 0.12 0.13 0.12ppm 2 Sr 15 213 101 51 21 11 205ppm 0.2 Ta <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2ppm 0.003 Tb 0.065 0.273 0.253 0.099 0.048 0.038 0.074ppm 0.09 Th <0.09 0.16 0.14 <0.09 <0.09 <0.09 <0.09ppm 26 Ti 696 2202 2190 1069 539 513 737ppm 0.005 Tl 0.023 1.714 1.965 0.017 <0.005 0.045 0.021ppm 0.002 Tm 0.042 0.178 0.166 0.069 0.033 0.024 0.051ppm 0.02 U 0.02 0.05 0.06 <0.02 <0.02 <0.02 0.03ppm 10 V 37 181 181 87 33 28 32ppm 0.5 W 0.9 <0.5 <0.5 0.8 1.1 <0.5 0.7ppm 0.08 Y 2.8 11.34 10.84 4.44 1.84 1.48 3.25ppm 0.009 Yb 0.284 1.171 1.057 0.439 0.205 0.171 0.334ppm 8 Zn 80 232 219 122 66 53 51ppm 3 Zr 9 26 24 12 5 4 8ppm 1 As 1 <1 2 <1 2 2 3ppm 20 Ba <20 674.5 772.2 <20 <20 <20 <20ppm 4 Cr 1524 860 843 3277 1552 1669 1894ppm 1 Cu 46 2 32 54 11 40 71ppm 1 Ni 1456 106 219 496 2282 2165 1886ppm 1 Rb <1 77 93 1 <1 <1 1ppm 6 Sc 11 38 40 19 9 10 17ppm 2 Sr 15 221 102 51 22 12 209ppm 4 V 67 198 191 97 56 54 74ppm 1 Y 4 12 11 5 3 3 4ppm 3 Zr 10 27 25 12 7 7 9ppb 0.22 Au 2.79 0.58 1.49 2.49 7.64 3.13 7.69ppb 0.01 Ir 1.46 0.02 0.01 1.18 2.31 3.51 2.88ppb 0.12 Pd 3.88 0.14 0.15 3.98 0.98 15.55 9.55ppb 0.17 Pt 3.24 <0.17 <0.17 6.63 1.11 8.75 5.72ppb 0.02 Rh 0.6 <0.02 <0.02 0.87 0.16 2.18 1.39ppb 0.08 Ru 3.03 <0.08 <0.08 3.76 1.67 7.17 4.91Wt% 0.03 CO2 7.89 0.64 1 6.63 6.52 15.8 12.4wt% 0.01 S 0.3 0.04 0.12 0.05 0.16 0.41 0.32

XRF

FA-IC

P-M

S

B2

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI

MgOMnONa2OP2O5SiO2TiO2TotalBaBeBiCdCeCoCrCsCuDyErEuGaGdHfHoLaLiLuMoNbNdNiPbPrRbSbScSmSn

KD6020-74.98

KD6020-76.8

KD6020-80.46

KD6024-170.5

KD6024-178.6

KD6024-196.3

KD6024-198.4

KD6024-212.1

KD6026-78.2

Geolabs Geolabs Geolabs Geolabs Geolabs Geolabs Geolabs Geolabs Geolabs07-0521 07-0521 07-0521 07-0521 07-0521 07-0521 07-0521 07-0521 07-0521

2.42 9.51 3.43 2.25 2.76 8.73 3.04 2.98 8.9411.33 8.23 4.96 3.14 0.57 7.73 3.02 2.91 7.116.22 12.67 8.62 7.7 10.12 13.75 8.93 8.47 12.790.09 3.95 0.01 0.02 0.02 0.22 0.01 0.01 0.7

20.06 3.44 17.9 18.73 11.39 5.66 17.53 16.23 5.4625.51 17.19 29.15 31.09 36.48 20.07 31.24 32.88 21.380.17 0.22 0.15 0.12 0.14 0.23 0.16 0.14 0.240.04 0.5 0.05 0.03 0.05 0.26 0.03 0.02 0.250.01 0.04 0.02 0.01 0.01 0.02 0.01 0.01 0.02

33.99 44.13 35.61 36.45 39.01 43.97 35.71 37.19 43.380.13 0.51 0.19 0.13 0.16 0.38 0.15 0.15 0.37

99.96 100.39 100.09 99.65 100.72 101.03 99.82 100.99 100.635.8 588.7 73.4 113.5 1.8 20.3 1.2 <0.9 29.5

0.68 1.9 0.35 0.32 0.11 0.25 0.07 0.1 0.340.432 0.25 1.033 0.943 0.685 0.329 0.821 0.577 0.4120.04 0.1 0.05 0.04 0.02 0.08 0.03 0.02 0.041.4 2.9 1.5 1 1 2.2 0.8 0.7 2.3

76.3 66.9 89.9 134.4 91.4 112.5 100.9 98.9 74.5>600 >600 >600 >600 >600 >600 >600 >600 >6001.349 >120 0.286 0.392 1.476 7.769 0.328 0.674 41.52

9 7 92 63 68 82 40 18 810.5 1.7 0.7 0.4 0.5 1.4 0.5 0.5 1.50.3 1.08 0.39 0.27 0.3 0.88 0.31 0.3 0.92

0.144 0.367 0.148 0.108 0.086 0.306 0.104 0.096 0.2982.61 9.41 3.24 2.52 2.93 9.35 3.03 2.79 11.960.39 1.23 0.48 0.32 0.37 1.06 0.4 0.37 1.090.22 0.83 0.27 0.18 0.27 0.61 0.25 0.21 0.56

0.104 0.385 0.143 0.09 0.102 0.307 0.114 0.108 0.3240.77 1.3 0.58 0.38 0.37 0.86 0.3 0.22 0.924.4 72.7 0.5 0.7 2.2 26.7 0.6 <0.2 51.1

0.05 0.179 0.059 0.039 0.053 0.141 0.047 0.044 0.1481.08 50.29 0.75 0.2 0.31 0.5 0.19 0.07 0.460.28 1.07 0.33 0.28 0.29 0.58 0.23 0.2 0.510.93 2.28 1.22 0.8 0.84 2.27 0.79 0.71 2.061727 283 >2000 >2000 1642 709 1772 1926 435

2 11 2 2 1 2 2 1 10.2 0.44 0.24 0.15 0.16 0.39 0.14 0.11 0.376.2 234.5 0.4 0.6 0.8 12.8 0.2 0.2 54.9

0.13 0.45 0.27 0.25 0.24 0.25 0.29 0.28 0.199.7 34.7 11.7 8.1 13.4 34.7 12.8 11.6 39.7

0.33 0.89 0.38 0.26 0.26 0.83 0.3 0.29 0.810.15 1.19 0.15 0.15 0.15 0.41 0.43 0.09 1.3Sn

SrTaTbThTiTl

TmUVWY

YbZnZrAsBaCrCuNiRbScSrVYZrAuIr

PdPtRhRu

CO2S

0.15 1.19 0.15 0.15 0.15 0.41 0.43 0.09 1.3177 34 93 72 <2 7 49 13 4<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2

0.076 0.246 0.092 0.06 0.065 0.208 0.074 0.074 0.203<0.09 0.17 <0.09 <0.09 <0.09 0.1 <0.09 <0.09 <0.09608 2471 886 584 769 1936 769 730 1954

0.077 7.701 0.011 0.008 0.193 0.221 0.067 0.016 1.1010.047 0.175 0.061 0.042 0.045 0.138 0.049 0.045 0.140.02 0.11 0.02 0.02 0.02 0.03 <0.02 <0.02 0.0231 189 55 38 49 161 74 56 1890.6 2.9 0.8 0.8 1 <0.5 <0.5 0.8 1.1

2.92 10.62 3.84 2.63 2.94 8.62 3.04 3.01 9.040.321 1.131 0.382 0.245 0.332 0.88 0.316 0.295 0.94211 219 437 258 83 208 71 61 268

7 29 10 7 10 19 9 7 172 <1 <1 1 1 1 2 2 5

<20 655.9 131.1 152.9 <20 33 <20 <20 51.31381 1062 1897 1566 1569 1520 2293 1674 1636

13 7 81 52 62 76 36 17 892073 264 2353 2194 1539 617 1572 1935 374

7 227 1 1 1 12 <1 <1 5219 37 16 12 11 34 14 15 39

187 37 102 77 <2 7 50 14 561 221 79 64 66 175 76 72 1974 11 5 4 4 9 4 4 109 29 12 8 10 21 9 9 19

2.14 1 7.91 21.43 1.08 2.81 2.56 4.29 5.612.54 0.21 0.54 2.6 2.81 0.93 1.93 3.13 0.982.06 2.98 3.65 5.24 3.03 4.08 4.93 3.49 14.922.86 2.75 2.84 6.4 2.11 7.04 4.61 3.68 14.880.53 0.28 0.47 0.93 0.48 0.7 0.75 0.63 1.432.73 0.55 2 3.94 2.25 2.32 3.07 3.72 3.116 1.44 12.9 15.1 0.12 0.24 13.1 9.89 0.08

0.29 0.06 0.17 0.51 0.15 0.25 0.38 0.13 0.48

B3

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6026-82.7

KD6026-87.2

KD6036-283.9

KD6036-328

KD6037-382.2

KD6037-401.4

KD6037-407.1

KD6043-180

KD6043-277


3.95 2.53 2.66 2.42 3.16 3.76 2.08 9.16 1.986.83 1.07 0.49 0.83 3.74 3.22 3.23 7.83 0.946.85 9.68 7.53 8.53 8.78 8.33 7.12 13.22 6.930.02 0.01 0.01 0.01 0.01 0.09 0.01 2.3 0.047.85 12.74 12.45 8.69 13.64 9.7 14.76 3.95 12.33

29.59 36.44 38 39.6 32.84 30.97 36.19 18.46 38.80.16 0.11 0.13 0.11 0.15 0.12 0.13 0.22 0.110.12 0.04 0.04 0.07 0.03 0.06 0.03 0.32 0.040.01 0.01 0.01 0.01 0.01 0.01 0.01 0.03 0.01

45.83 36.96 38.29 40.2 38.07 44.28 37.31 45.01 39.120.14 0.13 0.14 0.13 0.16 0.18 0.12 0.46 0.12

101.35 99.71 99.74 100.59 100.61 100.71 100.99 100.95 100.41<0.9 <0.9 <0.9 2.5 <0.9 2 <0.9 224.9 1.30.32 0.06 0.11 <0.06 0.23 0.29 0.13 1.12 0.09

0.986 0.434 1.268 0.815 1.922 0.772 0.634 0.157 1.090.05 0.02 0.02 0.03 0.03 0.06 0.03 0.05 0.021.1 0.6 0.6 0.7 0.9 1.3 0.6 2.2 0.5

78.2 117.9 101.4 109.6 106.8 93.6 92.1 83.9 93.2>600 >600 >600 >600 >600 >600 >600 >600 >6002.038 0.3 0.168 1.352 1.444 4.185 0.776 42.649 0.252

17 <2 21 13 46 18 20 83 90.6 0.4 0.4 0.5 0.5 0.6 0.3 1.7 0.4

0.37 0.27 0.28 0.28 0.34 0.4 0.22 1.07 0.230.177 0.075 0.091 0.099 0.114 0.107 0.078 0.388 0.0696.84 2.48 2.76 2.39 3.36 4.31 2.02 8.84 1.920.47 0.33 0.33 0.35 0.45 0.47 0.27 1.24 0.290.2 0.2 0.21 0.19 0.24 0.27 0.19 0.66 0.15

0.133 0.096 0.095 0.099 0.12 0.135 0.075 0.378 0.0730.45 0.23 0.2 0.25 0.3 0.54 0.26 0.9 0.172.6 0.7 0.3 5.7 1.4 4.4 0.3 65.5 0.5

0.06 0.047 0.048 0.046 0.061 0.068 0.036 0.168 0.0360.28 0.35 0.34 0.34 0.51 9.11 0.21 0.58 0.140.21 0.17 0.18 0.18 0.26 0.43 0.17 0.67 0.141.11 0.57 0.65 0.73 0.89 0.98 0.53 2.23 0.51273 >2000 >2000 >2000 >2000 1930 2233 528 >2000

1 1 1 1 3 20 1 2 10.2 0.11 0.11 0.12 0.15 0.19 0.1 0.38 0.080.7 0.2 <0.2 0.8 1.1 6.5 0.2 153.1 0.7

0.46 0.42 0.45 0.26 0.72 0.46 0.19 0.17 0.8210 12.7 12.5 12.5 12.7 15 9.4 34.6 9.9

0.36 0.25 0.26 0.27 0.33 0.37 0.19 0.88 0.210.24 0.08 0.08 0.08 0.24 0.33 0.11 0.36 <0.08Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.24 0.08 0.08 0.08 0.24 0.33 0.11 0.36 0.089 <2 3 <2 30 15 32 16 2

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.088 0.061 0.065 0.066 0.082 0.085 0.051 0.247 0.051<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09676 637 661 617 730 828 508 2251 492

0.099 0.058 0.072 0.013 0.056 0.242 <0.005 2.199 0.0920.059 0.042 0.044 0.045 0.056 0.062 0.035 0.16 0.034<0.02 <0.02 <0.02 <0.02 <0.02 0.05 0.02 0.03 <0.02

59 58 57 50 63 84 39 180 331.6 8.1 3.6 2.8 <0.5 <0.5 <0.5 0.6 4.3

3.88 2.63 2.7 2.75 3.43 3.74 2.13 10.73 2.120.378 0.289 0.294 0.293 0.381 0.411 0.223 1.065 0.22159 82 64 61 71 104 150 224 121

7 7 7 7 9 9 7 23 58 11 14 9 3 4 5 4 22

<20 <20 <20 <20 <20 <20 <20 269.7 <201028 1650 1780 1820 2135 1906 1406 2040 1594

22 5 30 20 60 20 21 72 121126 2478 2211 2291 2629 1933 2422 494 2180

1 <1 <1 1 1 6 <1 145 112 12 11 11 16 16 10 36 99 <2 3 <2 30 16 33 17 2

62 60 62 60 87 101 56 207 505 4 4 4 4 5 3 11 39 8 9 8 9 10 7 24 7

1.5 2.02 2.63 1.71 27.26 11.01 19.15 29.56 5.682.36 3.45 5.87 6.17 12.2 1.32 4.04 0.74 8.61

15.35 3.93 3.4 3.5 61.05 5.9 5.32 11.11 2.5110.55 3.42 3.12 1.25 41.87 5.43 2.74 11.62 2.261.97 0.6 0.62 0.36 8.21 0.91 0.53 1.37 0.535.58 4.06 4.2 2.96 27.82 4.18 2.79 3.17 3.360.58 1.01 0.31 0.13 5.44 2.45 4.97 0.75 0.740.09 0.17 0.27 0.26 0.17 0.06 0.07 0.07 0.26

B4

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6045-303.5

KD6045-313.1

KD6049-480.1

KD6049-511.7

KD-6049-544.7

6054W1-857.2

6054W1-789.3

KD6062A-787.5

KD6062A-804.5


3.95 2.36 5.82 1.7 2.15 4.68 9.54 8.04 7.036.06 3.76 5.83 0.7 0.8 4.34 5.57 5.84 7.11

11.19 7.42 10.6 6.58 7.47 9.59 12.28 12.08 11.660.02 0.03 0.03 <0.01 0.02 2.93 4.95 0.37 0.03

10.05 9.37 7.23 16.86 15.8 8.38 2.16 6.16 5.4531.2 37.71 28.31 36.98 37.09 23.91 19.34 25.82 24.020.17 0.12 0.15 0.12 0.16 0.17 0.16 0.18 0.20.08 0.07 0.09 0.02 0.03 0.05 0.3 0.13 0.110.02 0.03 0.02 0.01 <0.01 0.02 0.04 0.03 0.02

36.44 39.12 42.24 37.58 37.1 44.21 45.45 41.38 44.440.21 0.14 0.3 0.1 0.12 0.24 0.45 0.38 0.35

99.39 100.13 100.63 100.66 100.74 98.53 100.23 100.41 100.420.9 1.2 <0.9 <0.9 2.2 328.8 528.4 10.6 0.9

0.06 0.06 0.1 0.23 0.29 0.63 0.69 0.19 0.330.848 0.786 0.612 1.703 0.703 0.744 0.852 0.322 0.5950.03 0.03 0.03 0.02 0.02 0.03 0.09 0.06 0.071.3 0.7 1.1 0.5 0.9 1.2 11.5 3.1 1.4

111.8 99.3 94 91.6 100.1 94.9 77.8 81.3 91.8>600 >600 >600 >600 >600 >600 >600 >600 >6000.533 0.314 1.846 0.973 1.388 20.876 40.039 7.5 1.638

40 30 24 17 27 83 117 71 540.8 0.5 0.9 0.3 0.4 0.9 1.8 1.2 1.1

0.48 0.3 0.59 0.15 0.25 0.57 1.09 0.79 0.720.126 0.115 0.161 0.052 0.069 0.175 0.472 0.254 0.2443.88 2.17 5.3 1.85 2.48 4.93 11.91 8.16 6.680.62 0.36 0.67 0.2 0.29 0.64 1.61 0.89 0.890.28 0.18 0.43 0.14 0.17 0.34 1.02 0.57 0.51

0.167 0.102 0.193 0.055 0.084 0.2 0.378 0.266 0.2450.46 0.26 0.41 0.18 0.34 0.57 5.65 1.32 0.422.1 3.6 0.3 1.4 1.7 18.6 44.9 1.1 0.3

0.078 0.046 0.094 0.028 0.041 0.087 0.175 0.127 0.1141.08 0.14 0.1 4.22 0.12 1 1.37 2.06 4.950.28 0.17 0.41 0.17 0.36 0.45 1.17 0.76 0.551.33 0.74 1.27 0.35 0.68 1.22 6.35 2.05 1.521760 >2000 1439 >2000 >2000 1660 468 922 1146

1 1 1 1 2 5 5 2 20.24 0.14 0.22 0.08 0.13 0.21 1.49 0.45 0.260.5 0.2 0.8 0.3 0.4 84.2 189.9 20.8 1.8

0.22 0.17 0.53 0.23 0.42 0.12 0.44 0.34 0.214.5 10.9 19 8.7 10.5 18.4 28.7 24.4 22.40.47 0.28 0.49 0.14 0.22 0.46 1.5 0.73 0.630.09 0.08 0.13 0.09 0.15 0.27 1.39 0.78 0.39Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.09 0.08 0.13 0.09 0.15 0.27 1.39 0.78 0.3923 22 4 9 38 117 9 9 28

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.109 0.072 0.131 0.036 0.055 0.125 0.28 0.173 0.167<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 0.54 <0.09 <0.09898 572 1389 430 526 1134 2234 1862 16950.01 <0.005 0.025 <0.005 0.097 1.082 3.314 0.309 0.048

0.074 0.043 0.086 0.027 0.039 0.089 0.171 0.122 0.109<0.02 <0.02 <0.02 0.04 0.06 0.06 0.21 0.07 0.02

96 50 99 37 42 106 168 136 1302.6 2.9 1.6 1.4 1.7 <0.5 <0.5 1 0.5

4.92 2.96 5.5 1.57 2.42 5.67 10.35 7.54 70.505 0.299 0.592 0.166 0.25 0.575 1.11 0.799 0.725119 57 57 60 72 110 240 172 120

9 7 14 5 6 12 36 19 166 7 7 5 4 5 <1 <1 3

<20 <20 <20 <20 <20 414 640.1 27.8 <204543 1645 2457 1573 1985 2077 1726 2641 2855

50 33 27 18 28 66 99 60 521638 2021 1310 2231 2217 1495 345 758 1025

<1 1 1 <1 <1 82 185 20 218 14 25 8 11 19 29 29 2627 28 4 9 38 118 11 9 29

107 60 122 50 58 107 184 158 1496 4 7 3 4 6 11 9 8

11 8 16 6 7 14 36 19 182.89 6.41 5.33 3.04 7.64 10.07 4.86 3.13 8.712.62 6.38 1.27 6.54 6 2.39 1.12 0.83 1.24.43 0.48 7.33 2.44 3.85 5.67 8.37 9.71 8.824.52 0.46 7.33 2.43 2.96 5.62 9.08 9.2 8.80.9 0.12 1.07 0.5 0.62 0.9 1.2 1.28 1.25

4.98 1.7 3.68 3.33 3.69 4.14 3.17 3.84 4.092.33 0.87 0.27 9.25 6.72 6.01 0.07 0.04 0.320.28 0.46 0.2 0.26 0.13 1.24 0.79 0.25 0.56

B5

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6066-619.1

KD6066-622.4

KD6066-624.4

KD6066-636.6

KD6068-710.9

KD6068-847.8

KD6068-871.8

6068AW2-826

6069AW1-616.7


3.15 1.68 7.75 3.17 4.05 1.98 1.85 6.38 2.372.33 0.83 7.23 3.08 0.47 1.56 2.09 5.56 1.557.98 14.46 12.6 8.25 7.21 6.98 5.91 11.37 7.56

<0.01 0.01 0.11 <0.01 0.02 0.01 0.02 0.29 0.0117.03 14.56 5.33 16.63 11.7 14.98 13.64 7.15 14.4431.9 33.06 22.65 31.16 35.52 35.76 34.68 26.02 36.930.15 0.1 0.21 0.16 0.09 0.11 0.09 0.16 0.140.02 0.03 0.15 0.02 0.06 0.04 0.04 0.39 0.030.01 <0.01 0.02 0.01 0.01 <0.01 <0.01 0.02 0.01

37.02 34.13 44.09 36.48 40.59 38.77 42.06 42.88 37.370.16 0.09 0.38 0.16 0.15 0.09 0.08 0.34 0.13

99.75 98.96 100.53 99.13 99.87 100.28 100.47 100.56 100.53<0.9 <0.9 4.4 <0.9 <0.9 <0.9 <0.9 <0.9 <0.90.14 0.19 0.54 0.11 0.12 0.15 0.13 0.24 0.14

0.692 1.909 0.572 0.92 0.098 0.109 0.115 0.052 1.2240.03 0.17 0.12 0.04 0.02 0.02 0.03 0.05 0.020.8 0.5 1.5 0.6 1 0.6 0.7 1 0.6

96.9 203.8 101.8 94.5 103.9 93.4 83.9 86.5 89.7>600 >600 >600 >600 >600 >600 >600 >600 >6000.24 1.142 3.918 0.219 3.797 2.251 2.131 11.785 1.39422 308 51 28 58 37 36 62 190.5 0.3 1.2 0.5 0.7 0.3 0.4 1.2 0.40.3 0.2 0.78 0.35 0.47 0.24 0.23 0.72 0.27

0.108 0.071 0.338 0.089 0.13 0.131 0.163 0.178 0.0692.91 1.96 7.17 2.85 4.45 2.19 2.1 6.07 2.260.41 0.24 0.92 0.41 0.52 0.25 0.26 0.9 0.330.22 0.14 0.54 0.24 0.24 0.11 0.12 0.54 0.18

0.111 0.069 0.27 0.123 0.158 0.081 0.079 0.261 0.0870.29 0.19 0.48 0.23 0.34 0.23 0.26 0.35 0.190.2 0.6 1.1 <0.2 5.1 2.9 2.8 7.6 0.5

0.046 0.031 0.121 0.047 0.078 0.041 0.041 0.117 0.0436.51 0.46 0.17 0.18 0.1 0.08 0.07 0.07 0.070.22 0.13 0.5 0.22 0.37 0.19 0.16 0.57 0.170.84 0.44 1.71 0.69 0.86 0.5 0.6 1.27 0.571866 >2000 1113 1790 2221 >2000 >2000 1262 >2000

3 6 3 3 1 2 2 1 10.15 0.08 0.29 0.12 0.15 0.1 0.11 0.2 0.10.2 0.6 6.8 <0.2 2 1 1 12.3 0.4

0.24 0.62 0.66 0.25 1.73 0.97 0.91 0.32 1.2612.6 9 26.6 11.7 14.1 8.4 7.9 20.2 9.70.29 0.17 0.66 0.29 0.4 0.2 0.21 0.59 0.220.19 0.62 0.95 0.23 0.24 0.16 0.15 0.25 0.12Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.19 0.62 0.95 0.23 0.24 0.16 0.15 0.25 0.1231 7 13 45 6 23 36 18 6

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.077 0.045 0.176 0.078 0.099 0.051 0.056 0.172 0.06<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09764 430 1885 766 664 396 320 1690 548

0.072 0.051 0.211 0.115 0.157 0.083 0.085 0.085 0.0070.047 0.032 0.118 0.052 0.076 0.04 0.043 0.112 0.043<0.02 <0.02 <0.02 <0.02 0.02 <0.02 <0.02 0.02 <0.02

71 65 142 75 94 57 50 121 500.7 5.7 0.9 0.6 2.5 1.6 1.4 0.6 2.6

3.15 1.89 7.52 3.44 4.54 2.34 2.45 6.86 2.40.307 0.207 0.819 0.331 0.497 0.282 0.252 0.785 0.259

60 127 127 61 52 40 37 68 658 5 18 8 6 4 4 18 62 3 4 5 18 16 15 <1 16

<20 <20 <20 <20 <20 <20 <20 <20 <201897 2802 3239 1792 1582 1480 1258 2578 1679

21 328 44 25 60 38 34 52 221775 3027 887 1649 2265 2266 2083 1087 2373

<1 1 7 <1 2 1 1 13 <114 8 31 15 14 8 7 25 1033 7 13 50 6 24 35 19 770 51 168 73 98 58 52 139 605 3 9 4 5 3 4 9 49 6 20 9 7 6 5 20 8

10.28 19.79 6.85 6.55 1.8 3.09 3.5 1.33 26.381.54 7.08 1.27 2.45 1.92 4.1 2.96 1.21 5.824.53 25.85 10.01 4.19 3.37 2.16 1.83 8.21 3.314.01 13.6 9.52 3.89 3.24 2.31 1.77 8.12 2.970.74 3.27 1.38 0.73 0.69 0.5 0.42 1.15 0.573.43 10.22 4.71 3.65 3.92 3.31 2.95 4.12 3.6312.8 7.18 0.06 13.1 0.86 5.89 5.05 0.12 4.570.42 2.21 0.76 0.92 0.16 0.26 0.27 0.06 0.14

B6

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


6069AW1-635.6

KD6070-280.3

KD6070-286.3

KD6070-295.1

KD6074-394.7

KD6074-404.8

KD6082-290

KD6082-320.7

KD6083A-403.4


2.41 8.31 3.52 5.76 1.65 2.18 4.11 4.52 10.272.01 6.95 5.7 3.67 0.19 0.87 6.38 4.97 6.087.24 12.8 9.7 11.11 6.15 7.61 9.1 9.88 14.77

<0.01 0.03 0.01 0.01 0.01 0.01 0.01 0.01 4.1511.69 6.01 12.22 8.14 13.2 10.26 13.62 12.95 338.64 22.76 27.57 28.29 39.26 39.18 26.72 28.31 17.680.12 0.21 0.16 0.16 0.1 0.15 0.18 0.16 0.250.04 0.25 0.05 0.07 0.04 0.06 0.04 0.04 0.250.01 0.02 0.01 0.02 <0.01 0.01 0.02 0.01 0.0438.3 42.41 41.27 42.27 39.34 38.86 38.84 38.91 43.410.12 0.41 0.18 0.28 0.1 0.13 0.24 0.23 0.47

100.57 100.16 100.39 99.8 100.02 99.31 99.26 99.99 100.361.1 1.4 <0.9 <0.9 14.6 1 <0.9 <0.9 365.2

0.06 0.33 <0.06 0.32 0.15 0.23 0.15 0.24 0.440.292 0.265 0.408 1.941 1.024 0.372 0.874 1.631 0.1130.02 0.06 0.03 0.04 0.02 0.02 0.11 0.06 0.060.6 2.2 1.4 1.1 0.5 0.6 1.4 1.4 8.8

91.8 95.3 96.8 99.9 89.3 96.6 77.8 86.4 99.3>600 >600 >600 >600 >600 >600 >600 >600 >6000.195 1.599 0.336 1.686 0.551 2.038 0.681 1.226 49.748

3 85 47 59 15 12 63 88 460.4 1.5 0.7 0.9 0.3 0.4 0.7 0.8 2

0.26 0.91 0.42 0.55 0.2 0.24 0.47 0.45 1.220.105 0.332 0.162 0.119 0.067 0.098 0.192 0.152 1.172.17 8.43 3.46 5.56 1.77 2.17 4.07 4.27 12.730.31 1.09 0.59 0.66 0.24 0.27 0.6 0.58 1.70.19 0.61 0.26 0.47 0.15 0.15 0.36 0.32 0.86

0.096 0.318 0.154 0.195 0.069 0.087 0.173 0.157 0.4370.22 0.84 0.5 0.38 0.16 0.2 0.51 0.5 4.790.8 3.9 <0.2 0.7 0.6 3.3 0.4 0.2 57.5

0.042 0.142 0.067 0.092 0.031 0.039 0.075 0.07 0.1950.17 0.28 2.79 0.48 0.62 0.24 1.49 0.83 2.880.16 0.68 0.24 0.4 0.11 0.16 0.41 0.33 1.170.66 2.21 1.33 1.25 0.5 0.57 1.29 1.25 5.13

>2000 868 1486 1520 >2000 >2000 1418 1602 6311 2 2 4 1 1 2 2 1

0.12 0.39 0.23 0.21 0.08 0.1 0.24 0.22 1.1<0.2 1 0.2 0.8 0.5 1.2 0.3 0.4 146.90.5 0.17 0.35 0.35 0.43 0.3 0.19 0.25 0.4810 27.8 12.6 21.3 8 10.3 14.8 14.7 33.8

0.25 0.84 0.49 0.49 0.17 0.22 0.47 0.46 1.32<0.08 0.42 <0.08 0.19 <0.08 0.13 0.23 0.34 1.75Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.08 0.42 0.08 0.19 0.08 0.13 0.23 0.34 1.7510 21 83 30 14 2 145 26 16

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.061 0.213 0.113 0.13 0.044 0.055 0.113 0.114 0.307<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 0.23531 1985 805 1389 361 482 1082 1028 2308

<0.005 0.031 0.032 0.044 0.047 0.019 0.062 0.112 3.7360.043 0.145 0.067 0.091 0.032 0.04 0.077 0.07 0.192<0.02 0.03 <0.02 0.02 <0.02 <0.02 0.02 <0.02 0.08

49 170 76 126 36 36 81 80 1901.2 <0.5 0.6 0.8 2.2 1.8 0.5 0.7 0.7

2.57 8.79 4.6 5.52 1.89 2.34 4.74 4.43 12.940.272 0.918 0.451 0.568 0.197 0.265 0.488 0.446 1.246

47 167 84 75 48 75 83 87 5716 20 8 16 5 5 12 10 30

13 4 3 3 8 8 1 3 <1<20 <20 <20 <20 20.8 <20 <20 <20 488.5

1524 2545 2286 2198 1368 1658 1720 1893 26016 75 43 55 17 14 40 80 40

2149 743 1371 1368 2388 2385 1128 1497 565<1 1 1 1 1 1 1 <1 14010 32 18 22 7 11 21 21 3510 22 90 32 15 2 153 28 1755 184 87 120 42 54 92 96 2054 10 6 7 3 4 6 6 137 21 10 15 6 7 14 13 31

6.63 2.9 4.21 3.85 2.21 3.95 3.42 16 0.895.3 0.75 1.7 1.71 4.95 5.34 1.32 1.89 0.323 10.65 4.12 6.86 5.02 2.67 4.6 7.27 0.75

2.68 10.29 4.08 6.92 3.69 2.62 3.79 6.09 1.650.55 1.36 0.8 1.02 0.51 0.57 0.58 1.14 0.133.29 3.91 3.64 3.81 2.53 3.8 2.49 2.93 0.380.69 0.03 7.51 2.29 0.22 0.48 9.67 8.04 0.070.27 0.61 0.34 0.27 0.2 0.2 0.64 0.25 0.23

B7

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6083A-405.5

KD6083A-407.8

KD6083A-413.7

KD6084-691.7

KD6084-741.4

KD6084-804.1

KD6093-236.8

KD6093-241.2

KD6093-247.2


3.48 2.89 2.21 6.18 3.15 2.21 10.33 2.7 2.583.43 2.53 0.93 7.66 2.17 0.43 8.49 1.01 0.839.77 9.37 7.91 10.13 8.52 6.86 13.67 7.81 8.690.04 0.01 0.01 0.01 <0.01 <0.01 0.09 <0.01 0.02

12.63 13.22 14.17 5.89 16.29 20.27 4.95 12.8 13.0132.85 34.67 37.45 24.76 31.13 34.37 19.15 36.54 36.660.18 0.16 0.14 0.17 0.14 0.14 0.21 0.14 0.120.04 0.03 0.03 0.08 0.03 0.01 0.36 0.05 0.070.01 0.01 0.01 0.02 0.01 0.01 0.02 0.01 0.01

37.98 37.74 37.52 44.76 37.77 35.42 43.07 39.15 38.060.16 0.15 0.12 0.29 0.16 0.12 0.44 0.13 0.14

100.56 100.78 100.5 99.96 99.37 99.84 100.79 100.34 100.22.9 <0.9 2.9 <0.9 <0.9 <0.9 11.1 1.4 <0.90.1 0.1 <0.06 0.2 0.09 0.08 0.23 <0.06 0.09

0.281 0.553 0.24 0.943 1.768 1.029 0.048 0.624 0.5420.06 0.05 0.03 0.05 0.05 0.02 0.08 0.03 0.031.1 1 0.9 1.1 0.5 0.6 3.6 0.8 0.7

87.1 95.1 105.7 88.3 100.6 86.5 93.3 109.5 99.3>600 >600 >600 >600 >600 >600 >600 >600 >6001.637 2.012 1.485 0.877 0.295 0.311 1.646 2.731 2.18

46 50 23 49 76 9 2 26 140.5 0.5 0.4 1.1 0.4 0.2 1.7 0.6 0.4

0.33 0.31 0.26 0.66 0.25 0.14 1.03 0.37 0.290.099 0.106 0.08 0.171 0.065 0.054 0.62 0.088 0.0863.33 2.68 2.21 6.31 3.2 2.17 9.78 2.71 2.640.39 0.38 0.32 0.81 0.32 0.21 1.36 0.39 0.340.21 0.23 0.17 0.45 0.22 0.16 0.7 0.2 0.2

0.116 0.11 0.091 0.239 0.088 0.049 0.367 0.125 0.0970.46 0.4 0.36 0.32 0.17 0.25 1.45 0.32 0.25

1 1.1 1.3 <0.2 0.3 0.7 30.2 0.6 0.40.062 0.05 0.042 0.104 0.039 0.024 0.161 0.059 0.0460.14 0.07 0.06 5.52 0.09 0.09 0.11 0.09 0.090.23 0.21 0.17 0.45 0.22 0.16 0.68 0.19 0.180.9 0.83 0.77 1.31 0.5 0.56 2.92 0.77 0.65

1059 1842 >2000 1138 >2000 >2000 435 >2000 >20002 3 1 5 3 1 1 2 1

0.18 0.14 0.14 0.22 0.09 0.11 0.57 0.14 0.111.4 0.6 0.5 0.4 0.2 0.3 1.2 0.4 0.4

0.52 0.87 1.15 0.41 0.3 0.22 0.51 0.98 1.2413.5 11.8 9.4 20.2 12 8.7 30.3 12.3 11.80.31 0.28 0.24 0.56 0.23 0.19 1.05 0.3 0.250.21 0.14 0.08 0.27 0.17 0.14 1.07 0.36 0.14Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.21 0.14 0.08 0.27 0.17 0.14 1.07 0.36 0.149 12 11 5 31 7 13 7 4

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.076 0.073 0.065 0.152 0.06 0.038 0.248 0.083 0.068<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 0.1 <0.09 <0.09666 653 491 1466 753 529 2166 585 5910.07 0.031 <0.005 0.061 0.033 0.005 0.034 0.081 0.041

0.056 0.052 0.042 0.106 0.04 0.023 0.163 0.058 0.0460.02 <0.02 0.02 0.02 <0.02 <0.02 0.03 <0.02 <0.0272 56 41 113 103 58 191 72 641.3 1.7 1.9 0.7 0.5 0.6 0.6 1.6 2.1

3.29 2.99 2.64 6.8 2.55 1.33 10.24 3.46 2.830.389 0.321 0.27 0.667 0.255 0.153 1.024 0.376 0.294203 180 83 87 120 59 133 92 84

7 8 6 15 8 6 22 7 710 11 9 <1 2 1 4 12 23

<20 <20 <20 <20 <20 <20 22.8 <20 <202043 1917 1686 2328 3851 1670 1266 1508 1489

48 51 26 45 65 8 1 27 151117 1972 2333 822 2205 2171 416 2533 2262

2 1 <1 1 <1 <1 1 <1 <117 14 9 23 14 9 33 12 1210 12 11 5 33 8 12 7 489 67 59 135 74 56 195 72 635 4 4 8 4 4 10 4 49 9 8 16 10 8 24 8 8

7.9 7.34 7.24 8.21 9.58 3.14 0.56 20.44 17.161.05 2.24 3.78 0.67 1.53 3.04 0.29 2.3 2.482.27 3.62 13.23 4.85 32.08 3.99 1.18 15.85 4.493.22 1.75 12.06 4.94 29.92 3.53 2.73 9.8 3.640.58 0.46 1.9 0.7 4.25 0.6 0.22 2.07 0.921.67 1.42 3.28 2.92 7.61 1.28 0.31 3.03 1.423.71 3.9 4.1 0.05 11.8 16 0.15 2.21 2.450.16 0.22 0.15 1.92 0.34 0.34 0.02 0.13 0.1

B8

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6093-252.1

KD6168-128.6

KD6168-129.7

KD6168-133

KD6168-137.3

KD6169-370.2

LG16-76-426.2

LG16-76-439.5

LG7-149-137.6


5.4 2.99 6.15 2.34 2.24 2.59 3.55 2.56 3.624.25 1.07 5.88 0.25 0.36 2.71 4.85 7.66 2.83

10.28 10.28 11.26 8.23 8.27 9.14 8.57 7.15 8.730.01 0.01 0.02 0.01 0.01 0.01 <0.01 0.01 0.018.74 11.23 6.34 12.82 12.93 13.34 14.53 15.71 10.82

27.67 35.04 26.08 37.33 37.2 34.14 29.2 27.15 33.440.2 0.08 0.27 0.13 0.13 0.16 0.17 0.18 0.15

0.07 0.05 0.09 0.05 0.04 0.03 0.04 0.03 0.060.02 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01

42.99 38.27 44.22 38.03 38.35 37.58 38.51 39.16 40.350.27 0.16 0.29 0.13 0.13 0.14 0.19 0.16 0.19

99.89 99.19 100.62 99.33 99.66 99.86 99.63 99.78 100.19<0.9 <0.9 1.9 <0.9 <0.9 <0.9 <0.9 <0.9 <0.90.11 0.12 0.23 0.15 0.18 0.22 0.07 <0.06 0.22

0.371 0.827 0.407 0.995 1.256 2.062 1.933 1.701 1.0250.05 0.03 0.1 0.02 0.02 0.02 0.03 0.04 0.031.6 1 1.5 0.7 0.6 0.8 0.9 1.2 0.9

101.8 117.1 105.4 104 98.6 101.7 92.2 79.6 95.5>600 >600 >600 >600 >600 >600 >600 >600 >6001.023 2.797 2.24 1.63 1.949 1.535 0.419 0.443 1.899121 36 33 7 35 39 20 32 571.2 0.5 1.1 0.4 0.4 0.6 0.6 0.5 0.6

0.67 0.35 0.7 0.26 0.28 0.35 0.39 0.34 0.360.091 0.162 0.161 0.116 0.088 0.096 0.117 0.165 0.1285.16 2.97 7.1 2.28 2.25 2.71 3.59 2.78 3.620.89 0.42 0.85 0.34 0.32 0.43 0.46 0.46 0.460.43 0.24 0.41 0.17 0.18 0.19 0.3 0.25 0.26

0.245 0.123 0.247 0.093 0.095 0.127 0.141 0.119 0.1260.66 0.43 0.5 0.26 0.2 0.28 0.38 0.49 0.351.1 3.5 1.5 3.6 3.5 0.6 <0.2 0.4 0.8

0.099 0.055 0.104 0.045 0.046 0.058 0.062 0.052 0.0573.28 0.2 20.6 0.38 1.71 4.5 0.39 3.65 0.410.42 0.21 0.42 0.16 0.15 0.2 0.33 0.26 0.271.59 0.94 1.62 0.64 0.58 0.85 0.84 1.1 0.91572 1999 1227 >2000 2571 >2000 1824 1835 1998

0 1 <0.4 1 1 1 2 4 20.27 0.17 0.28 0.11 0.11 0.14 0.16 0.2 0.160.4 0.5 1.4 0.4 0.4 0.7 0.2 0.2 0.5

0.25 0.87 0.11 0.57 0.7 0.4 0.2 0.16 0.6418.3 14.8 25 11.1 10.6 12.8 13 10 13.90.68 0.32 0.64 0.25 0.26 0.32 0.33 0.38 0.320.23 0.26 0.29 0.09 0.08 0.15 0.11 <0.08 0.17Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.23 0.26 0.29 0.09 0.08 0.15 0.11 0.08 0.1714 4 11 2 3 14 35 178 27

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.175 0.081 0.155 0.063 0.061 0.081 0.088 0.085 0.083<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.091289 716 1415 544 526 590 939 707 8440.148 0.289 0.085 0.516 0.397 0.122 0.117 0.034 0.1310.107 0.054 0.112 0.043 0.044 0.061 0.064 0.054 0.0590.02 <0.02 0.07 <0.02 <0.02 <0.02 <0.02 0.05 <0.02107 74 151 51 45 58 70 65 860.7 1.7 <0.5 4.5 1.8 0.8 0.6 <0.5 1.2

6.42 3.3 6.8 2.62 2.59 3.45 3.87 3.24 3.490.646 0.345 0.69 0.258 0.273 0.361 0.385 0.334 0.375158 100 111 33 38 51 56 66 6014 8 14 6 6 6 10 9 10<1 6 3 6 8 10 2 1 8

<20 <20 <20 <20 <20 <20 <20 <20 <202160 2692 2957 1547 1610 1977 1775 1586 1856111 51 42 10 51 55 23 37 70

1336 1928 1116 2222 2546 2182 1487 1612 1936<1 1 2 1 1 1 <1 1 118 14 26 11 12 14 18 16 1615 4 11 2 3 14 36 192 28

112 68 127 58 56 67 84 69 908 4 8 4 4 5 5 4 5

16 9 15 8 8 8 12 10 1110.33 1.21 1.26 1.69 1.14 18.21 5.92 3.58 13.221.61 1.6 0.67 1.4 1.85 3.55 3.33 3.7 2.1

10.48 2.24 0.63 1.32 3.14 1.05 4.3 3.29 8.597.2 2.67 2.51 0.65 1.96 1.3 4.26 3.04 7.05

1.25 0.64 0.12 0.19 0.43 0.48 0.72 0.56 1.772.02 1.41 0.4 0.66 1.38 2.02 1.38 1.34 3.093.22 0.25 0.07 0.26 0.38 4.21 10.2 11.4 20.33 0.27 0.04 0.15 0.18 0.16 0.48 0.32 0.22

B9

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LG7-150-129.8

LNSD-17-1027.8

LSU-1-790.5

LSU-1W2-599.8

LSU-1W2-606.8

LSU-1W2-663.5

LSU-1W2-680.4

LSU-12-221

LSU-12-223.3


2.3 5.26 3.88 2.4 6 4.98 2.76 3.35 3.021.65 2.87 1.07 3.19 8 9.12 2 9.04 12.47.79 10.27 9.73 7.25 12.17 9.11 7.62 7.33 8.780.01 <0.01 0.01 0.01 3.58 0.02 <0.01 2.26 1.25

10.21 12.39 6.01 19.55 2.37 4.66 13.54 2.83 10.9939.3 29.21 27.93 31.09 19.07 22.97 31.46 22.11 22.490.13 0.14 0.12 0.13 0.2 0.23 0.15 0.17 0.270.04 0.04 0.06 0.02 0.17 0.11 0.04 0.13 0.060.01 0.02 0.02 0.01 0.02 0.01 0.01 <0.01 <0.01

39.05 38.93 50.29 36 48.63 47.68 41.54 53.3 37.520.13 0.27 0.24 0.14 0.3 0.24 0.12 0.15 0.19

100.63 99.41 99.37 99.79 100.51 99.14 99.23 100.66 96.981 <0.9 <0.9 1 374.8 <0.9 <0.9 100.9 146.5

0.13 <0.06 0.09 0.13 0.41 0.28 0.06 0.6 0.910.437 0.502 1.411 0.629 2.992 0.59 0.177 23.163 3.4360.02 0.07 0.03 0.04 0.07 0.14 0.05 0.12 0.140.6 1.3 1.6 0.8 1.4 1 7.5 1.1 4.2

99.3 95.4 114.4 86.3 106 92.9 89.8 84.9 84.1>600 >600 >600 >600 >600 >600 >600 >600 >6000.149 0.349 0.533 0.265 39.901 0.693 0.193 26.183 14.613

14 33 106 29 145 62 30 96 590.4 0.9 0.7 0.5 1.2 1.4 0.9 0.6 1.2

0.24 0.51 0.41 0.28 0.68 0.8 0.57 0.36 0.640.081 0.102 0.048 0.111 0.217 0.199 0.08 0.189 0.3552.13 5.15 4.27 2.37 7.79 7 3.59 5.41 5.620.32 0.67 0.52 0.34 0.83 1.02 0.89 0.47 0.950.17 0.44 0.4 0.2 0.44 0.35 0.91 0.34 0.41

0.087 0.184 0.151 0.104 0.248 0.291 0.204 0.126 0.2390.21 0.45 0.67 0.29 0.48 0.31 3.46 0.37 1.820.6 0.6 1.9 1 44.5 0.4 1.1 5.3 1.8

0.041 0.079 0.064 0.044 0.111 0.109 0.096 0.058 0.0942.25 0.09 0.65 0.05 0.87 0.11 0.07 3.71 62.360.16 0.46 0.48 0.16 0.54 0.8 1.88 0.37 0.410.58 1.28 1.32 0.78 1.48 1.53 3.83 1.09 2.9

>2000 1426 2261 1869 831 1566 1974 1503 14781 1 0 1 4 6 2 36 8

0.1 0.23 0.25 0.14 0.26 0.25 0.91 0.2 0.60.4 <0.2 0.2 0.2 138.6 0.6 <0.2 127.1 77.20.8 0.24 0.16 0.15 0.15 0.28 0.2 0.18 0.15

10.7 18.8 16.9 10.8 22.9 16.3 8 8.7 15.10.23 0.5 0.42 0.28 0.61 0.7 0.89 0.37 0.88

<0.08 0.34 0.08 <0.08 0.51 1.04 0.2 0.31 0.56SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.08 0.34 0.08 0.08 0.51 1.04 0.2 0.31 0.564 22 2 87 8 8 10 27 341

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 0.2 <0.2 <0.20.055 0.124 0.101 0.067 0.157 0.191 0.161 0.086 0.173<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 1.34 0.14 0.1517 1312 1220 582 1437 1120 518 698 1019

0.005 0.028 0.211 0.029 2.385 0.075 0.018 1.38 0.8850.04 0.081 0.069 0.045 0.109 0.123 0.098 0.058 0.099

<0.02 0.02 0.03 <0.02 0.07 0.07 0.58 0.11 0.0546 108 98 66 146 88 56 66 2014.1 0.8 <0.5 0.6 <0.5 <0.5 0.5 <0.5 12.7

2.47 4.93 3.94 2.82 6.95 8.32 5.61 3.86 7.240.274 0.493 0.448 0.299 0.7 0.786 0.617 0.365 0.614

39 131 135 46 198 416 129 160 1606 15 14 7 15 12 29 11 14

13 2 10 <1 <1 2 <1 3 8<20 <20 <20 <20 508.2 <20 <20 188.5 214.3

1592 2843 1668 1633 4009 2462 1482 1474 161921 34 108 27 136 67 29 101 61

2360 1174 1494 1625 745 1159 1456 1361 1215<1 <1 <1 1 134 1 <1 122 6811 21 14 12 25 23 10 14 244 23 2 89 9 8 10 29 313

57 110 80 65 138 92 46 73 2014 6 5 4 8 9 7 5 78 16 16 8 17 14 33 13 13

2.18 2.57 25.8 1.28 47.16 6.14 8.14 30.98 7.244.98 0.97 2.26 2.75 0.79 1.85 3.84 0.82 1.373.18 3.62 14.03 8.6 8.74 5.84 6.79 3.99 8.642.56 4.23 9.96 5.26 7.98 5.91 11.91 3.77 7.310.56 0.63 1.9 1.19 1.07 0.86 0.99 0.71 11.83 1.71 3.24 2.21 4.42 3.72 3.2 3.64 3.690.69 6.99 0.1 15.3 0.08 0.03 9.78 0.07 9.720.16 0.42 1.08 0.29 0.88 1.08 0.71 0.92 1.97

B10

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LSU-143-542.3

LSU-143-573.3

VS15-150-157

VS15-150-168.4

KD6037-414.8

KD6083A-391.3

KD6083A-404.8

KD6083A-416.9

LJD0010A-425.6


1.8 1.89 4.75 2.49 1.96 2.36 4.49 2.46 2.440.66 0.31 5.2 2.02 0.43 0.21 7.68 1.05 2.067.63 7.14 9.79 9.21 7.02 8.36 13.27 8.25 11.120.01 0.01 <0.01 0.01 0.01 <0.01 0.01 <0.01 0.01

12.35 16.54 12.14 12.46 17.12 13.18 4.95 12.11 9.6936.95 37.92 27.9 34.74 37.89 37.27 24.46 36.82 34.16

0.1 0.15 0.23 0.14 0.12 0.13 0.18 0.12 0.130.05 0.02 0.04 0.03 0.02 0.03 0.08 0.04 0.05

<0.01 <0.01 0.02 0.01 0.01 0.01 0.02 0.01 0.0239.87 36.04 38.3 38.83 35.84 39.08 44.94 39.28 39.41

0.1 0.11 0.24 0.13 0.11 0.1 0.21 0.13 0.2399.51 100.13 98.61 100.06 100.51 100.73 100.28 100.28 99.33

1.8 0.9 <0.9 <0.9 1 <0.9 <0.9 <0.9 <0.90.21 0.24 0.07 0.2 0.22 0.11 0.06 0.08 0.58

0.225 0.08 0.552 0.176 0.255 0.395 0.441 0.425 0.4330.02 0.03 0.07 0.04 0.02 0.02 0.12 0.02 0.030.5 0.9 1.2 0.8 0.4 0.6 0.7 0.7 1.7

112.3 105.4 86.7 105.4 93.6 105.3 168.6 103.1 111.4>600 >600 >600 >600 >600 >600 >600 >600 >6003.357 2.005 0.257 0.925 1.002 1.078 0.432 1.188 0.537

42 11 88 71 10 24 223 33 30.3 0.3 0.8 0.4 0.3 0.4 0.7 0.4 0.60.2 0.21 0.45 0.28 0.19 0.26 0.42 0.26 0.35

0.065 0.064 0.09 0.085 0.057 0.066 0.062 0.087 0.1252.06 1.98 4.48 2.53 1.82 2.21 4.72 2.31 4.070.21 0.23 0.53 0.33 0.2 0.3 0.48 0.31 0.520.12 0.16 0.34 0.2 0.16 0.15 0.31 0.2 0.470.07 0.069 0.168 0.1 0.066 0.09 0.153 0.093 0.1290.16 0.26 0.45 0.3 0.13 0.19 0.27 0.25 0.663.3 1.6 0.2 0.6 0.7 0.7 <0.2 1.8 4.8

0.035 0.035 0.076 0.047 0.034 0.044 0.059 0.042 0.0510.34 0.15 0.95 0.1 0.11 0.09 5.28 0.16 0.150.15 0.19 0.34 0.19 0.13 0.14 0.26 0.18 0.420.44 0.48 1.13 0.68 0.37 0.55 0.88 0.63 1.36

>2000 >2000 1410 1638 >2000 >2000 >2000 >2000 19366 2 4 4 1 3 3 2 1

0.08 0.09 0.2 0.12 0.06 0.09 0.15 0.11 0.251.4 0.7 <0.2 0.3 0.4 0.4 0.2 0.3 0.3

2.09 1.51 0.15 0.52 0.3 1.12 0.26 1.14 0.049 7.8 16 11.8 8.6 9.8 13.8 10.1 10.4

0.17 0.18 0.42 0.25 0.14 0.22 0.38 0.23 0.450.08 0.09 0.14 0.14 <0.08 0.09 0.15 0.08 <0.08Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.08 0.09 0.14 0.14 0.08 0.09 0.15 0.08 0.0825 12 84 48 4 <2 4 9 2

<0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.2 <0.20.044 0.044 0.107 0.067 0.04 0.058 0.097 0.057 0.094<0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 <0.09 0.12405 469 1122 598 448 455 929 567 1131

0.009 <0.005 0.022 0.071 <0.005 0.02 0.138 <0.005 0.0230.034 0.035 0.074 0.046 0.031 0.042 0.067 0.042 0.053<0.02 0.02 <0.02 <0.02 <0.02 0.03 <0.02 <0.02 0.07

50 41 92 55 40 41 89 53 1004.4 3.2 0.6 1.5 1.1 2.2 <0.5 2 16.1

1.94 1.92 4.59 2.88 1.73 2.46 4.11 2.45 3.40.234 0.222 0.467 0.297 0.197 0.274 0.394 0.269 0.331

50 58 78 73 155 84 236 78 424 5 11 7 6 5 10 7 18

22 17 4 8 <1 9 17 18 5<20 <20 <20 <20 <20 <20 <20 <20 <20

1597 1706 1827 1676 1621 1522 2217 1725 238457 15 100 69 11 23 214 36 5

2952 2507 1204 1621 2456 2324 1651 2729 18681 1 1 <1 <1 1 <1 <1 <18 9 22 11 9 10 17 11 11

25 11 93 49 4 <2 4 8 251 51 102 66 52 61 100 62 913 3 6 4 3 4 6 4 56 7 13 8 7 7 12 8 18

8.8 1.43 2.99 9.28 4.32 44.16 20.67 15.15 <0.222.02 2.48 1.34 3.23 4.7 3 0.5 7.33 2.182.16 0.55 5.72 1.76 3.08 1.41 2.3 46.31 2.941.18 0.58 6.29 4.04 2.59 1.97 2.11 33.24 3.090.15 0.12 0.85 0.55 0.52 0.31 0.27 5.38 0.71.16 1.3 3.35 3.6 2.75 2.03 1.31 18.46 5.021.23 6.69 8.39 3.65 9.55 2.96 0.17 1.69 0.110.17 0.07 0.89 0.48 0.1 0.61 1.04 0.19 0.11

B11

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0011-780.2

LJD0068-125.7

KD5051-320.2

KD5051-345.4

KD5073-328.9

KD5073-487.5

KD5073-558.4

KD5082-339.1

KD5082-379.9


3.62 0.95 6.96 2.06 9.96 3.36 5.41 8.7 2.6510.14 0.91 6.36 0.75 3.55 2.21 4.17 7.48 0.5211.2 8.07 11.31 7.51 11.21 8.56 11.75 13.03 6.730.02 <0.01 0.95 N.D. 6.31 0.01 N.D. 0.58 0.293.59 11.33 5.37 9.62 1.81 16.16 11.06 4.69 15.75

22.64 42 25.47 41.68 20.42 31.83 28.41 22.24 32.920.19 0.24 0.18 0.14 0.11 0.11 0.13 0.2 0.080.2 0.11 0.14 0.04 0.16 0.02 0.03 0.38 0.03

0.03 0.01 0.02 N.D. N.D. 0.01 0.01 0.03 N.D. 48.55 36.71 44.69 39.7 46.88 38.47 38.96 43.64 42.570.33 0.09 0.34 0.12 0.42 0.17 0.25 0.43 0.13

100.51 100.42 101.8 101.63 100.85 100.92 100.2 101.38 101.69<0.9 <0.9 30.42 1.76 280.02 1.69 N.D. 23.74 6.581.71 0.06 N.D. N.D. 4.28 N.D. N.D. 1.54 0.5

0.602 0.0760.04 0.03 0.032 N.D. 0.078 N.D. 0.227 0.045 N.D. 3.6 1.6 1.53 0.54 3.58 0.57 1.43 2.13 0.38

96.4 131.8 87.92 98.7 64.08 97.4 114.27 91.07 84.46>600 >600 2350 1508 1388 1910 1989 2496 15570.146 0.05 19.449 0.424 93.158 0.497 0.679 9.254 5.586

60 4 34 6 11 21 193 21 141.3 0.4 1.189 0.372 1.301 0.353 0.58 1.637 0.217

0.76 0.21 0.781 0.236 0.829 0.214 0.353 1.061 0.1340.249 0.096 0.241 0.071 0.334 0.08 0.149 0.354 0.044.97 1.25 6.32 1.92 10.85 3.15 4.71 8.2 2.811.14 0.33 0.934 0.29 0.994 0.307 0.525 1.292 0.1810.54 0.17 0.5 0.2 0.7 0.3 0.4 0.7 0.3

0.287 0.079 0.272 0.079 0.283 0.073 0.125 0.359 0.0461.34 0.73 0.54 0.2 1.59 0.2 0.53 0.72 0.140.6 6.1 3.13 0.82 40.96 N.D. N.D. 14.39 12.79

0.109 0.028 0.125 0.038 0.125 0.032 0.056 0.165 0.0250.41 0.17 0.84 1.02 8.14 0.15 0.2 2.28 0.450.42 0.12 0.6 0.3 1.3 0.4 0.5 0.7 0.72.91 1.01 1.56 0.54 2.39 0.59 1.23 2.19 0.351295 >2000 1063 2407 471 1988 3602 808 2187

1 <0.4 1 1 2 3 4 2 10.58 0.23 0.283 0.094 0.536 0.107 0.238 0.408 0.0650.2 <0.2 56.54 0.7 >150.00 0.76 0.74 40.63 22.39

<0.04 0.08 0.62 0.86 0.09 0.3 0.3 0.37 0.2114.5 5.6 22.51 8.22 17.11 8.45 16.29 29.17 4.620.98 0.28 0.64 0.2 0.73 0.22 0.42 0.86 0.140.23 <0.08 0.16 N.D. 2.21 0.12 1.5 0.3 0.19Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.23 0.08 0.16 N.D. 2.21 0.12 1.5 0.3 0.1933 13 23.1 2.8 7.5 28.4 16.7 19.5 9.3

<0.2 <0.2 N.D. N.D. N.D. N.D. N.D. N.D. N.D. 0.201 0.058 0.182 0.054 0.193 0.053 0.093 0.243 0.0330.12 <0.09 N.D. N.D. 0.14 N.D. N.D. N.D. 0.081602 336 1520 456 1869 763 1086 1926 5340.016 <0.005 0.82 0.01 5.29 0.02 0.06 0.45 0.240.116 0.029 0.119 0.035 0.124 0.031 0.055 0.162 0.0210.04 <0.02 0.026 0.008 0.221 0.02 0.019 0.025 0.058101 26 160.01 47.91 161.89 88.18 108.05 202.03 62.6517.4 62.6 1.38 2.26 0.26 0.55 0.56 7.65 0.847.8 2.08 6.35 1.9 5.68 1.6 2.9 8.54 1.05

0.719 0.178 0.79 0.24 0.79 0.2 0.35 1.03 0.1562 49 102.55 56.37 332.38 52.21 82.22 169.91 198.118 6 17 8.1 24.6 11.2 13.8 24.4 9.5<1 2 2 13 N.D. 1 1 N.D. 1

<20 <20 60 N.D. 327 N.D. N.D. 45 N.D. 1869 1362 2662 1614 1599 1835 2128 3078 1626

62 6 37 9 10 21 169 20 131119 2943 1133 2170 485 1833 2821 861 2072

<1 1 57 N.D. 337 N.D. N.D. 42 2323 6 29 10 20 15 23 35 936 14 23 2 11 29 17 19 999 34 139 45 168 71 98 185 569 3 8 2 10 2 5 11 2

21 6 20 10 28 12 15 26 10<0.22 0.42 2.96 9.38 1.42 9.82 7.76 4.07 1.71.33 4.05 1.48 6.62 0.61 4.34 8.86 0.95 3.345.3 1.53 7.76 2.56 8.38 4.56 44.54 10.31 1.42

5.96 2.59 8.12 2.49 9.03 4.25 20.3 9.75 1.740.74 0.7 1.14 0.53 1.14 0.75 6.94 1.34 0.323.33 5.56 4.4 4.35 2.38 4.85 22.31 4.47 2.710.02 2.960.66 0.18 0.03 0.12 N.D. 0.14 0.81 N.D. 0.04

B12

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD5082-421.6

KD5085-434.1

KD5085-497.5

KD5105-152.2

KD5106-189.8

KD5106-244.5

KD5109-506.3

KD5115-704.8

KD5115-709.6


1.57 2.44 1.68 1.59 8.73 2.52 3.01 8.05 4.990.21 2.2 0.08 0.08 5.71 2.18 2.33 7.46 6.166.74 7.1 8.51 7.02 12.98 7.53 8.95 11.68 8.650.01 N.D. N.D. N.D. 1.66 N.D. 0.03 1.17 0.01

22.02 20.57 21.24 18.82 4.72 14.25 12.26 4.43 14.0139 33.58 39.19 39.41 24.57 36.57 34.27 21.86 27.53

0.16 0.12 0.26 0.15 0.18 0.14 0.15 0.2 0.17 N.D. 0.01 N.D. 0.01 0.22 0.03 0.03 0.24 0.04 N.D. N.D. N.D. N.D. 0.02 N.D. N.D. 0.02 N.D.

32.16 35.58 31.08 35.07 42.11 38.82 40.29 46.34 39.410.08 0.12 0.09 0.09 0.35 0.14 0.16 0.39 0.23

101.95 101.74 102.14 102.24 101.26 102.19 101.5 101.84 101.2 N.D. 0.98 1.35 N.D. 73.78 N.D. N.D. 48.93 N.D.

0.5 N.D. N.D. 0.39 0.55 N.D. N.D. N.D. N.D.

N.D. N.D. N.D. N.D. N.D. N.D. 0.033 0.044 0.0350.35 1.1 0.43 0.44 1.88 0.62 0.88 2.3 1.34

99.82 85.59 117.64 98.85 96.58 94.28 86.99 85.33 73.211359 1429 1402 1424 2735 1550 1844 2112 19230.964 0.37 0.833 0.683 20.745 1.058 1.15 18.154 1.001

26 24 56 8 28 33 42 97 590.228 0.464 0.164 0.207 1.242 0.412 0.578 1.42 0.7510.155 0.291 0.122 0.146 0.801 0.261 0.379 0.932 0.4810.037 0.108 0.029 0.034 0.208 0.092 0.112 0.305 0.1622.02 2.22 1.9 1.71 7.79 2.45 3.08 7.3 4.71

0.165 0.373 0.126 0.158 0.986 0.339 0.459 1.162 0.6170.2 0.3 0.2 0.2 0.6 0.3 0.3 0.6 0.4

0.05 0.1 0.037 0.046 0.281 0.09 0.127 0.317 0.1660.14 0.43 0.18 0.19 0.71 0.22 0.3 0.87 0.522.14 1.2 N.D. 0.7 9.37 N.D. N.D. 24.71 0.88

0.029 0.04 0.024 0.026 0.133 0.042 0.065 0.144 0.0770.14 0.09 0.28 0.17 0.21 0.12 0.18 0.26 0.090.5 0.4 0.4 0.3 0.6 0.3 0.4 0.6 0.4

0.29 0.9 0.3 0.28 1.69 0.61 0.84 2 1.212252 1930 2751 2438 1079 2293 1733 853 1764

2 1 3 1 1 1 4 5 30.057 0.18 0.063 0.065 0.325 0.115 0.157 0.408 0.2351.23 0.3 0.3 0.26 92.22 0.6 0.4 47.91 0.70.25 0.25 0.45 0.61 0.93 0.57 0.64 0.36 0.268.05 7.96 7.97 7.82 23.72 8.79 7.42 25.82 14.540.11 0.29 0.1 0.11 0.65 0.23 0.31 0.78 0.430.24 0.1 0.08 0.08 0.18 0.07 0.12 0.25 0.16Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.24 0.1 0.08 0.08 0.18 0.07 0.12 0.25 0.161.2 60.2 2 0.9 16.1 7.7 25.7 19.5 196.2

N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 0.031 0.069 0.023 0.03 0.188 0.061 0.087 0.214 0.114

N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 332 492 410 326 1591 566 661 1688 10640.01 N.D. N.D. N.D. 1.47 N.D. 0.16 0.95 0.08

0.025 0.044 0.021 0.023 0.124 0.041 0.06 0.143 0.0740.072 0.027 0.041 0.05 0.015 0.02 0.016 0.027 0.03141.89 62.48 47.73 42.35 213.5 59.08 91.49 159.33 114.880.95 0.7 1.24 1.35 0.66 1.2 1.82 0.61 0.761.22 2.36 0.89 1.07 6.28 2.08 2.84 7.75 3.850.17 0.27 0.14 0.16 0.8 0.26 0.4 0.91 0.48

46.46 50.96 54.85 71.05 112.6 53.39 80.41 74.65 63.439.2 8.8 8.1 7.3 18.9 9.1 12.1 21 14.42 N.D. 3 2 2 4 3 N.D. N.D.

N.D. N.D. N.D. N.D. 102 N.D. N.D. 85 N.D. 1437 1425 1496 1508 2939 1502 1785 2487 1915

24 12 53 9 27 32 40 95 501991 1723 2497 2219 1198 2113 1567 915 1437

N.D. N.D. N.D. N.D. 105 N.D. N.D. 48 N.D. 8 13 8 8 32 12 14 30 23

N.D. 62 N.D. N.D. 17 7 25 19 19942 51 49 40 169 55 76 157 932 2 1 2 8 2 4 9 48 9 9 8 21 10 11 24 11

12.54 2.71 23.14 7.94 2.21 13.85 11.08 19.1 5.735.29 3.93 3.45 7.12 0.97 5.6 2.86 1.16 2.158.29 2.76 10.25 2.41 13.5 5.88 2.66 8.98 7.994.95 2.84 5.89 2.02 8.69 3.87 3.16 9.09 3.291.31 0.56 1.31 0.48 1.18 0.6 0.5 1.21 0.976.12 3.84 5.41 3.83 4.7 3.64 3.04 3.83 4.51

0.15 0.16 0.3 0.1 N.D. 0.13 0.26 0.08 0.39

B13

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6025-199.8

KD6025-204.4

KD6027-315.5

KD6027-318.9

KD6037-282.1

KD6037-385.5

KD6037-443.4

KD6039-172.8

KD6039-465.8


2.77 5.38 10.72 3.03 8.18 4.43 1.67 7.94 3.60.6 7.97 15.73 0.91 6.77 4.01 0.27 6.08 3.537.5 9.75 11.83 9.56 12.46 10.47 7.34 12.63 10.14

N.D. 0.01 0.12 N.D. 0.04 0.02 0.02 1.38 N.D. 11.33 8 3.49 16.46 7.13 12.06 17.36 7.07 14.8437.9 26.72 8.91 33.75 23.22 31.12 38.16 22.9 30.540.14 0.22 0.21 0.13 0.19 0.16 0.11 0.17 0.180.03 0.09 1.96 0.03 0.17 0.04 0.01 0.2 0.03

N.D. 0.01 0.04 N.D. 0.02 0.01 N.D. 0.02 0.0140.96 42.59 47.58 36.56 42.33 38.55 36.26 42.31 35.580.15 0.26 0.53 0.15 0.4 0.22 0.1 0.39 0.18

101.4 101.01 101.12 100.6 100.92 101.1 101.3 101.09 98.64 N.D. N.D. 28.94 N.D. 1.78 N.D. 0.78 102.7 N.D. N.D. N.D. 0.86 N.D. 0.46 N.D. N.D. 1.03 N.D.

N.D. 0.03 0.054 N.D. 0.033 N.D. N.D. 0.035 0.0480.77 1.4 3.65 0.84 2.16 0.8 0.37 1.56 1.09

100.33 92.72 55.79 103.36 82.47 94.5 77.75 96.65 182.231803 2448 332 4483 2556 2707 1424 3083 16970.145 1.068 0.23 0.433 2.186 1.815 0.867 26.263 0.439

1 34 106 35 2 13 17 44 3740.472 0.917 2.073 0.427 1.421 0.66 0.251 1.311 0.6060.306 0.6 1.331 0.257 0.935 0.417 0.167 0.84 0.3990.089 0.173 0.449 0.072 0.193 0.126 0.038 0.311 0.1172.61 5.05 10.69 2.89 7.7 3.95 1.55 7.29 3.33

0.376 0.762 1.644 0.383 1.118 0.543 0.205 1.057 0.4970.3 0.4 0.8 0.2 0.6 0.3 0.2 0.6 0.3

0.104 0.207 0.444 0.098 0.325 0.147 0.058 0.299 0.1380.3 0.5 1.41 0.32 0.74 0.26 0.14 0.48 0.42

N.D. N.D. 6.61 0.84 8.98 2.01 N.D. 28.77 N.D. 0.044 0.093 0.202 0.042 0.148 0.066 0.028 0.133 0.0630.17 0.16 1.52 0.09 0.28 0.12 0.11 0.76 0.180.3 0.4 1 0.2 0.6 0.3 N.D. 0.6 0.3

0.73 1.31 3.22 0.8 2.02 0.86 0.34 1.79 0.922336 1385 53 1980 625 1552 3111 873 5045

1 1 8 2 1 1 2 7 30.134 0.252 0.625 0.145 0.389 0.151 0.066 0.312 0.1850.12 0.4 0.64 0.22 2.47 1.78 0.3 78.82 0.461.14 0.35 0.79 0.31 0.36 0.36 0.54 0.29 0.54.44 18.7 35.96 11.78 27.54 16.68 7.69 28.56 12.190.26 0.51 1.17 0.27 0.77 0.35 0.13 0.71 0.340.09 0.25 0.86 N.D. 0.25 0.11 N.D. 3.47 0.35Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.09 0.25 0.86 N.D. 0.25 0.11 N.D. 3.47 0.351.2 16.2 206.7 8 16.2 21 13.7 57.6 29.2


N.D. N.D. 0.18 N.D. N.D. N.D. N.D. 0.07 N.D. 619 1152 2406 659 1941 940 359 1782 7870.04 0.18 0.01 0.06 0.04 0.02 N.D. 1.25 0.45

0.046 0.09 0.202 0.042 0.143 0.064 0.027 0.132 0.0610.01 0.013 0.209 0.012 0.026 0.017 0.011 0.06 0.019

65.03 148.94 211.44 18.25 229.34 148.94 47.77 216.55 90.80.84 0.9 30.76 0.64 0.48 0.51 0.99 1.84 0.452.2 5.05 11.18 2.21 7.77 3.36 1.38 7.12 3.21

0.29 0.57 1.26 0.26 0.91 0.4 0.18 0.83 0.3986.03 148.14 161.2 97.25 138.12 76.57 61.23 123.9 78.47

9.1 13.6 28.7 8.6 21.3 11.9 7.9 20.2 11.715 4 3 N.D. N.D. 1 4 N.D. 1

N.D. N.D. 37 N.D. N.D. N.D. N.D. 142 N.D. 1627 2486 392 3896 3046 2716 1511 3604 1866

3 39 108 33 3 16 17 14 3372055 1381 53 1730 653 1644 2650 933 3686

N.D. N.D. N.D. N.D. N.D. 2 N.D. 83 N.D. 13 25 44 14 34 20 8 32 17

N.D. 16 205 8 15 20 14 59 3051 109 224 67 177 90 37 177 753 6 12 2 9 4 2 9 4

11 16 28 11 25 14 8 21 132.18 6.66 14.5 8.38 0.51 6.22 10.71 6.99 22.653.53 1.94 0.01 1.42 0.89 1.56 3.61 1.78 8.524.37 7.37 N.D. 3.55 9.53 3.78 7.56 10.6 57.124.18 6.72 N.D. 3.45 9.57 2.47 3.98 10.1 19.580.44 1.07 N.D. 0.59 1.25 0.91 0.89 1.38 6.823.32 4.22 N.D. 4.36 3.55 5.91 3.94 5.01 24.5

0.14 0.16 0.49 0.35 N.D. 0.04 0.17 0.01 1.39

B14

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6039-501.5

KD6041-376.2

KD6041-464.6

KD6041-502.8

KD6042A-594.1

KD6042A-611.5

KD6048-698.6

KD6048-783.8

KD6051-694.5


2.34 3.44 2.79 3.66 1.43 2.74 3.35 2.27 6.541.07 5.93 1.24 1.81 0.22 1.81 0.65 2.1 3.067.73 8.54 9.42 9.12 7.1 7.29 7.26 8.51 9.9

N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 12.81 11.29 6.67 8.29 15.39 18.11 15.54 22.14 9.5538.7 27.52 41.36 37.78 38.99 33.33 36.84 33.24 27.670.12 0.18 0.19 0.17 0.15 0.19 0.08 0.16 0.140.03 0.04 0.06 0.05 0.02 0.03 0.03 N.D. 0.04

N.D. 0.01 N.D. 0.01 N.D. N.D. 0.01 N.D. 0.0238.45 41.87 39.37 40.44 37.67 36.3 37.24 32.34 42.150.13 0.17 0.15 0.16 0.09 0.14 0.17 0.11 0.34

101.39 99.01 101.25 101.48 101.06 99.95 101.18 100.9 99.42 N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 1.52 N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D.

N.D. 0.036 N.D. 0.066 N.D. 0.061 N.D. 0.03 0.0520.72 1.31 0.54 0.8 0.72 0.62 0.64 0.98 0.8493.7 84.38 103.45 91 98.32 109 89.16 74.06 87.981462 1633 1784 1744 1331 1096 1886 1318 25280.088 0.431 0.761 0.381 1.842 0.623 0.967 0.606 0.988

25 58 27 63 15 130 18 69 750.408 0.648 0.417 0.559 0.262 0.389 0.494 0.452 0.6440.266 0.386 0.289 0.351 0.167 0.242 0.324 0.288 0.4410.087 0.235 0.085 0.118 0.07 0.088 0.08 0.112 0.0952.25 3.25 2.81 3.25 1.61 2.4 2.76 2.14 7.17

0.337 0.579 0.315 0.474 0.223 0.321 0.388 0.387 0.510.2 0.3 0.2 0.3 0.2 0.3 0.3 0.2 0.5

0.096 0.142 0.1 0.125 0.059 0.088 0.108 0.102 0.1520.26 0.5 0.2 0.29 0.31 0.23 0.22 0.37 0.3

N.D. 1.6 2.88 3.21 0.8 N.D. N.D. 1.09 3.60.045 0.062 0.049 0.056 0.029 0.034 0.051 0.045 0.080.12 0.13 0.22 0.14 0.14 0.1 0.16 0.07 6.470.2 0.3 0.2 0.3 0.3 N.D. 0.2 N.D. 0.5

0.66 1.12 0.52 0.81 0.5 0.57 0.66 0.9 0.852750 1622 2830 2980 2731 4554 2196 1850 1178

4 4 1 1 1 1 3 1 50.125 0.224 0.092 0.15 0.111 0.107 0.114 0.164 0.1550.24 0.23 0.44 0.2 0.56 0.21 0.46 0.62 0.68

>2.00 0.38 0.23 0.37 0.67 0.25 0.39 0.3 0.229.27 12.35 4.06 3.02 7.18 8.68 9.79 10.22 22.310.23 0.42 0.21 0.31 0.17 0.2 0.25 0.28 0.32

N.D. 0.43 N.D. 0.29 N.D. 0.16 0.14 N.D. 0.16SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

N.D. 0.43 N.D. 0.29 N.D. 0.16 0.14 N.D. 0.169.2 145.7 2.6 2.7 8.1 20.3 10.2 45.1 72.7


N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 538 736 581 736 324 564 727 516 14900.3 0.02 0.04 0.02 N.D. 0.18 0.18 0.06 0.03

0.042 0.06 0.045 0.056 0.026 0.036 0.047 0.045 0.070.016 0.03 0.01 0.013 0.07 0.011 0.011 0.016 0.0263.47 90.39 99.05 88.02 45 54.55 75.26 73.71 185.222.55 0.64 3.96 1.87 1.7 0.39 1.46 0.87 0.882.23 3.36 2.01 2.58 1.45 2.04 2.59 2.47 3.640.27 0.38 0.29 0.34 0.17 0.23 0.32 0.29 0.46

70.25 129.12 72.42 71.22 44.58 59.6 43.5 58.36 125.968.4 10.5 7.9 10.9 5.7 9.7 10.4 7.8 17.744 2 4 16 9 1 3 N.D. N.D.

N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 1450 1811 1787 1749 1440 973 1883 1321 2974

29 55 33 73 17 102 17 58 632418 1264 2633 2541 2448 2980 1836 1414 960

N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 11 18 15 16 8 13 13 13 259 146 2 3 8 21 9 45 71

48 73 71 75 37 54 55 61 1473 4 3 4 2 3 3 2 5

10 9 11 13 9 10 13 8 191.52 5.65 2.42 9.37 17.2 16.63 9.84 0.73 4.311.81 0.75 4.77 5.01 4.94 12.87 3.85 3.32 1.40.2 4.1 3.53 33.74 3.36 98.93 2.06 3.02 8.08

N.D. 4.07 4.6 16.99 2.28 45.84 2.17 2.39 8.060.03 0.69 0.6 2.32 0.52 13.76 0.5 0.45 1.130.23 2.55 2.86 6.87 1.98 30 2.34 2.4 3.26

0.26 1.46 0.16 0.33 0.12 0.65 0.65 0.24 1.1

B15

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6051-702.7

KD6051-766.3

KD6051-809.5

KD6053A-634.5

KD6053A-663.4

KD6056-299.3

KD6056-364.8

KD6056-387.8

KD6061-243


5 1.67 2.03 9.28 1.56 6.84 2.86 1.61 7.182.84 0.34 0.6 13.8 0.63 3.09 6.52 2.3 6.9

10.23 6.39 7.21 12.56 7.56 10.8 6.44 6.49 11.37 N.D. N.D. N.D. 0.27 N.D. 3.98 0.02 N.D. 2.28

11.83 18.01 16.13 6.39 10.19 8.19 15.19 17.4 6.229.09 37.66 37.11 16.17 40.76 26.41 28.44 37.58 22.660.12 0.13 0.1 0.23 0.17 0.11 0.16 0.13 0.180.04 0.01 0.02 0.9 0.03 0.07 0.03 0.02 0.130.01 N.D. N.D. 0.03 N.D. 0.02 N.D. N.D. 0.02

41.62 36.96 37.93 41.17 39.89 41.58 41.57 36.15 43.690.26 0.09 0.12 0.5 0.1 0.33 0.15 0.09 0.33

101.03 101.28 101.26 101.29 100.91 101.43 101.39 101.77 100.941.05 N.D. N.D. 8.6 N.D. 121.37 N.D. 1.75 121.17

N.D. N.D. N.D. 0.31 N.D. 2.63 N.D. 0.4 0.81

N.D. N.D. N.D. N.D. N.D. N.D. 0.034 N.D. N.D. 1.08 0.35 0.51 2 0.37 1.8 1.4 0.62 1.46

91.21 88.1 90.46 83.68 103.17 82.78 78.73 91.81 84.182202 1296 1546 916 1698 2234 1828 1385 25940.47 2.009 3.554 0.757 0.151 47.841 0.777 1.338 40.81218 11 5 2 19 46 22 16 159

0.671 0.274 0.346 1.936 0.245 1.155 0.565 0.268 1.3190.446 0.186 0.221 1.229 0.178 0.788 0.355 0.189 0.8450.118 0.053 0.084 0.595 0.069 0.166 0.148 0.061 0.1864.79 1.72 1.95 7.72 1.54 6.06 3.04 1.55 6.72

0.546 0.201 0.282 1.6 0.199 0.864 0.498 0.216 1.090.4 0.1 0.1 0.8 0.2 0.6 0.3 0.2 0.5

0.156 0.064 0.08 0.437 0.058 0.272 0.13 0.063 0.3070.4 0.13 0.19 0.67 0.13 0.69 0.57 0.28 0.47

1.23 1.85 1.02 26.73 0.92 35.99 2.36 0.52 20.820.075 0.033 0.037 0.189 0.031 0.115 0.057 0.032 0.1320.32 0.1 0.08 0.18 0.17 3.91 2.38 0.74 4.110.3 N.D. N.D. 0.8 N.D. 0.5 0.2 N.D. 0.5

1.04 0.34 0.51 2.45 0.37 1.62 1.14 0.43 1.861556 2318 2402 349 3250 1029 1761 2565 906

1 2 1 1 1 1 2 4 10.197 0.063 0.092 0.41 0.064 0.313 0.224 0.091 0.2980.34 0.62 0.76 1.12 0.15 >150.00 1.2 0.6 138.970.35 1.03 1.37 0.25 1.02 0.58 0.42 1.36 0.11

19.28 8.02 10 34.99 8.44 22.19 10.04 8.94 25.230.36 0.13 0.18 1 0.13 0.52 0.36 0.14 0.70.11 N.D. 0.09 0.32 N.D. 0.29 0.1 0.1 0.18Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

0.11 N.D. 0.09 0.32 N.D. 0.29 0.1 0.1 0.1851.4 3.4 4.7 23.4 4.1 46.1 104.5 132.3 35.8


N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 1196 356 463 2242 360 1464 611 341 15020.02 N.D. N.D. 0.02 N.D. 3.7 0.04 0.02 1.870.07 0.03 0.035 0.189 0.029 0.119 0.055 0.03 0.133

0.011 0.008 0.007 0.018 N.D. 0.037 0.065 0.015 0.016160.96 52.06 54.05 188.44 49.09 160.44 74.65 52.36 199.38

0.77 2.05 2.71 0.65 4.28 0.7 0.12 1.64 0.373.78 1.58 1.93 10.42 1.43 6.74 2.96 1.56 7.190.45 0.2 0.22 1.17 0.19 0.72 0.34 0.2 0.82

72.51 55.01 37.31 93.29 47.7 103.83 65.33 55.86 106.0814.5 5.4 4.9 26.2 6.3 20.2 9 8.5 18.7

N.D. 7 9 N.D. 24 2 N.D. 25 1 N.D. N.D. N.D. N.D. N.D. 135 N.D. N.D. 176

2249 1374 1628 1053 1870 2250 1796 1463 278616 11 5 4 23 41 18 17 151

1336 2128 2131 352 2605 1110 1652 2222 967 N.D. N.D. N.D. N.D. N.D. 295 N.D. N.D. 153

21 9 10 42 9 28 15 9 2550 3 5 23 4 49 108 131 37

107 39 45 193 36 118 58 34 1545 2 3 11 1 9 3 2 9

15 8 9 29 8 19 10 6 210.72 11.92 74.28 0.44 13.96 2.57 6.89 20.61 33.721.38 6.71 3.55 0.29 4.51 0.88 3.39 3.59 0.896.59 2.38 4.61 11.07 4.45 7.84 5.31 4.28 8.596.19 2.16 3.84 11.4 2.98 7.79 8.14 2.53 8.30.93 0.45 0.62 1.27 1.66 1.05 3.27 0.66 1.133.05 2.75 2.86 1.29 5.31 2.57 8.18 2.05 3.11

0.25 0.16 0.18 N.D. 0.4 0.01 0.04 0.18 0.2

B16

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6061-271.2

KD6061-294.3

KD6067-754.9

KD6067-765

KD6067-805.5

KD6071A-672.6

KD6071A-681.8

KD6071A-726.4

KD6074-348.6


2.3 1.44 8.34 3.13 1.92 10.33 3.23 2.63 2.111.24 2.43 7.47 2.06 0.31 9.6 0.43 1.12 2.587.18 8 12.22 7.82 7.02 13.45 10.23 7.85 7.02

N.D. N.D. 0.93 N.D. N.D. 0.14 N.D. 0.01 N.D. 14.06 13.42 4.43 16.84 20.42 3.67 10.65 20.87 13.9638.21 36.64 21.71 32.56 34.83 17.67 34.72 33.96 37.390.15 0.1 0.2 0.12 0.13 0.21 0.11 0.15 0.150.02 0.03 0.19 0.02 0.01 0.7 0.05 0.03 0.03

N.D. N.D. 0.02 N.D. N.D. 0.03 0.01 N.D. N.D. 38.35 38.77 44.63 38.48 35.98 45 41.01 33.36 38.080.13 0.09 0.4 0.14 0.11 0.49 0.19 0.14 0.12

101.66 100.92 100.54 101.19 100.75 101.28 100.64 100.12 101.470.68 N.D. 66.6 0.69 N.D. 4.4 N.D. N.D. N.D. 0.48 N.D. 0.55 N.D. N.D. N.D. N.D. N.D. N.D.

N.D. N.D. 0.031 N.D. N.D. N.D. N.D. N.D. N.D. 0.64 0.61 2.02 0.72 0.46 2.4 0.57 0.41 0.44

88.09 70.85 77.4 90.19 94.24 89.38 126.47 92.21 88.611422 1346 2553 1662 1488 1443 6786 1556 13850.179 0.908 8.917 0.175 0.138 0.991 1.614 0.33 0.185

17 49 26 27 9 2 7 42 10.403 0.324 1.451 0.397 0.215 1.844 0.477 0.481 0.310.252 0.205 0.969 0.266 0.142 1.191 0.34 0.315 0.2110.093 0.054 0.326 0.075 0.038 0.232 0.084 0.087 0.0562.14 1.6 7.59 2.92 1.87 9.19 3.48 2.77 1.92

0.322 0.265 1.166 0.345 0.179 1.518 0.351 0.36 0.2530.3 0.2 0.6 0.3 0.3 0.7 0.3 0.2 0.2

0.086 0.07 0.334 0.092 0.048 0.427 0.115 0.108 0.0720.23 0.23 0.68 0.27 0.17 0.8 0.2 0.14 0.16

N.D. 0.5 7.44 N.D. 0.53 63.37 2.59 1.12 N.D. 0.041 0.031 0.156 0.041 0.028 0.178 0.059 0.045 0.0350.38 1.94 0.11 0.09 0.09 0.13 0.12 0.08 0.16

N.D. N.D. 0.7 0.2 N.D. 0.8 0.3 N.D. N.D. 0.61 0.5 2.1 0.68 0.38 2.52 0.55 0.46 0.492137 2711 737 1988 2235 519 2398 1666 2607

1 2 1 2 0 1 1 1 10.114 0.104 0.373 0.126 0.074 0.454 0.098 0.076 0.080.46 0.22 33.9 0.26 0.18 0.94 0.56 0.2 0.161.54 0.63 0.36 0.24 0.24 0.67 1.21 0.17 1.54

10.56 8.48 26.76 11.54 9.11 36.98 5.22 9.83 7.880.22 0.18 0.74 0.23 0.13 0.98 0.22 0.2 0.19

N.D. N.D. 0.26 0.07 N.D. 0.16 N.D. N.D. N.D. SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

N.D. N.D. 0.26 0.07 N.D. 0.16 N.D. N.D. N.D. 21.1 8.8 17.5 31.1 4.5 7 1.4 11.9 20


N.D. N.D. N.D. N.D. N.D. 0.07 N.D. N.D. N.D. 534 296 1768 665 451 2352 862 614 448

N.D. 0.47 0.55 0.1 N.D. 0.01 N.D. 0.03 0.040.039 0.03 0.151 0.042 0.024 0.181 0.057 0.047 0.0320.015 0.007 0.038 0.013 0.013 0.021 0.011 0.017 N.D. 65.81 48.53 212.26 83.19 56.67 213.82 N.D. 121.28 59.72.22 1.7 0.54 0.69 0.88 0.4 2.53 0.35 4.12.14 1.74 7.96 2.12 1.14 10.1 2.41 2.67 1.70.26 0.2 0.95 0.26 0.17 1.14 0.36 0.28 0.2

63.52 32.33 72.68 43.65 37.55 70.44 67.89 67.44 48.248.9 6.4 21.6 11.3 11.4 21.3 10.4 8.9 6.89 3 N.D. 1 N.D. 1 9 N.D. 25

N.D. N.D. 151 N.D. N.D. N.D. N.D. N.D. N.D. 1530 1343 3222 1659 1567 1625 7002 1704 1553

19 48 25 26 7 5 9 33 31900 2353 793 1658 1786 519 2159 1364 2296

N.D. N.D. 34 N.D. N.D. N.D. N.D. N.D. N.D. 11 9 36 15 10 39 16 12 1219 7 17 32 4 6 N.D. 13 2153 37 181 62 41 200 88 44 452 2 9 3 2 10 3 3 2

10 8 24 10 9 30 13 10 931.44 11.89 2.32 9.08 6.27 0.49 1.27 1.06 22.875.65 9.17 0.61 1.55 3.46 0.34 2.3 3.91 6.755.59 61.03 10.45 3.64 1.31 12.47 4.65 5.83 6.260.82 34.25 10.12 3.67 1.67 12.03 4.46 8.6 1.920.07 9.61 1.3 0.64 0.46 1.43 1.04 1.96 0.381.29 20.97 2.47 2.44 2.37 2.04 5.52 5.27 2.49

0.16 0.29 N.D. 0.21 0.21 N.D. 0.04 0.42 0.17

B17

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6074-400.5

KD6083A-404

KD6083A-409

LNSD011-106.6

VS15-019-24.5

KD5073 473.5

KD5085 532.6

KD6025 194.3

KD6027 342.5

Geolabs Geolabs Geolabs Geolabs Geolabs Ultratrace Ultratrace Ultratrace Ultratrace06-0363 06-0363 06-0363 06-0363 06-0363 u92996 u92996 u92996 u92996

2.3 10.77 2.78 2.1 1.59 2.19 2.24 3.36 4.630.17 5.85 1.14 0.46 1.48 0.72 0.99 4.13 3.447.6 14.77 9.32 7.48 7.01 7.09 7.55 10.2 9.18

0.05 0.03 N.D. N.D. N.D. -0.01 -0.01 0.01 -0.019.94 6.23 12.15 11.83 10.81

39.71 22.06 35.98 38.83 35.47 34.1 32.6 27.3 28.90.11 0.27 0.15 0.06 0.11 0.09 0.05 0.12 0.130.05 0.15 0.03 0.05 0.04 0.08 0.07 0.1 0.09

N.D. 0.03 N.D. N.D. N.D. 0.009 0.005 0.013 0.01840.56 40.27 39.61 40.08 44.67 34.4 37.9 48.1 39.50.14 0.48 0.15 0.1 0.07 0.1 0.11 0.16 0.22

100.63 100.92 101.32 101.01 101.262.33 1.06 N.D. N.D. N.D. -1 -1 -1 -1

N.D. N.D. N.D. 0.32 N.D. 1.1 1 0.6 0.7

N.D. 0.059 0.034 N.D. 0.0710.54 3.94 0.79 0.52 0.48

97.24 80.77 84.33 90.33 79.481498 2647 1843 1390 13503.647 0.695 1.491 0.159 1.81

7 2 29 11 190.386 1.626 0.425 0.354 0.3430.26 1.079 0.288 0.238 0.229

0.074 0.351 0.085 0.061 0.0722.02 9.8 2.74 1.93 2.090.31 1.305 0.347 0.286 0.2820.2 0.8 0.2 0.2 0.2

0.089 0.376 0.098 0.083 0.0770.19 1.75 0.35 0.18 0.173.43 7.59 1.89 N.D. 1.05

0.042 0.176 0.048 0.038 0.0350.25 2.06 0.11 0.25 0.46

N.D. 0.7 N.D. N.D. N.D. 0.56 3.01 0.68 0.52 0.512527 309 1353 2464 2266

1 1 1 N.D. 2 2 1 -1 -10.097 0.573 0.126 0.09 0.0894.37 0.68 0.46 0.19 0.38 0.2 1.4 0.4 0.40.37 0.3 0.99 0.95 0.346.2 33.99 8.58 9.69 8.18

0.21 0.98 0.25 0.2 0.2 N.D. 0.71 0.08 N.D. N.D. Sn

SrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

N.D. 0.71 0.08 N.D. N.D. 0.8 3.1 3.9 2.1 9.4 9 30.5 6 60.5

N.D. N.D. N.D. N.D. N.D. 0.058 0.247 0.065 0.053 0.053

N.D. 0.15 N.D. N.D. N.D. 0.1 0.1 -0.1 0.2514 2157 612 443 3390.09 0.01 N.D. 0.04 N.D. 0.04 0.167 0.045 0.037 0.035

N.D. 0.037 0.024 0.008 0.007 -0.1 -0.1 -0.1 -0.162.05 276.3 73.89 66.99 50.36 50 60 90 952.09 0.42 2 2.78 1.312.07 9.12 2.43 2.04 1.930.26 1.07 0.29 0.24 0.22

53.49 >150.00 128.19 25.42 39.91 50 40 95 1456.6 27.7 8.9 8.4 6.2 11 4 3 129 N.D. 6 22 5

N.D. N.D. N.D. N.D. N.D. 1524 2744 1722 1409 1428 1320 940 1790 1080

8 2 31 13 20 50 40 375 252301 328 1371 2131 2053 2060 2330 1880 1500

4 N.D. N.D. N.D. N.D. 11 39 13 10 10

N.D. 3 4 N.D. 1047 229 59 49 382 10 3 2 2

10 31 11 9 83.35 0.35 8.32 1.09 6.969.44 0.04 2.19 5.62 5 2 3 -1 23.17 N.D. 0.17 4.38 4.12 -1 2 6 51.88 0.28 0.32 2.91 4.55 -1 -1 2 40.82 N.D. 0.02 0.47 0.42 -1 -1 -1 -16.47 N.D. 0.52 3.6 4.18 1 3 3 4

0.13 N.D. 0.04 0.22 0.24 -0.40 -0.40 0.80 1.60

B18

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6037 462.8

KD6037 462.8 Rpt

KD5082 433.4

KD6048 795.3

KD6053A 693.1

KD6071A 750.2

KD6074 408.3

KD6083A 422.3

KD6067 BW7 857

Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu92996 u92996 u92996 u92996 u92996 u92996 u92996 u92996 u92996

2.27 2.27 1.78 2.39 1.64 2.94 2.44 4.01 2.630.65 0.64 1.44 0.15 2.02 8.42 2.22 2.22 1.557.82 7.82 8.92 6.01 8.47 10.7 7.81 10.1 8.03-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 0.01 0.02 -0.01

37.1 37 32.4 34.6 34.4 25.2 38.3 30.6 32.90.1 0.1 0.09 0.07 0.1 0.13 0.09 0.1 0.1

0.09 0.06 0.08 0.08 0.11 0.07 0.11 0.12 0.070.01 0.011 0.01 0.012 0.01 0.015 0.012 0.021 0.0137.4 37.5 32.5 36.4 32.9 33 39.5 41.5 33.40.11 0.11 0.08 0.12 0.08 0.15 0.12 0.23 0.13

7 6 -1 -1 -1 -1 -1 -1 -1

0.7 0.7 2.3 0.5 4.5 1.8 2.2 2.7 1.5

2 3 3 2 1 -1 3 3 -1

0.4 0.6 -0.2 0.6 0.2 -0.2 0.2 0.8 -0.2

SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

156 154 11.5 3 7.5 52 2.5 50.5 21

0.4 0.4 0.1 0.1 0.2 -0.1 0.2 0.1 -0.1

-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.155 55 45 45 40 90 60 100 65

55 55 60 40 40 25 30 25 2514 14 2 2 5 -1 8 3 1

1060 1090 810 1030 990 870 1240 800 90030 30 340 45 990 365 45 70 45

2520 2440 5780 1990 12000 7430 2980 2270 2850

5 6 11 3 29 17 3 3 42 3 70 2 218 112 11 15 132 1 33 1 102 63 5 7 5-1 -1 8 -1 30 16 2 2 24 3 23 3 92 54 7 7 6

0.80 0.40 13.23 0.40 40.88 25.25 2.00 2.81 2.00

B19

Long-Victor

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


KD6067 BW7 857 Rpt

KD5073 562.8

KD5081A 583.4

KD5081A 586.5

KD6042A W1 625.5

KD6042A W1 625.5 Rpt

Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu92996 u92996 u92996 u92996 u92996 u92996

2.62 -0.01 1.35 0.22 0.18 0.171.54 0.16 1.13 1.42 1.04 1.058.04 68.8 26.2 59.2 61.4 61.5-0.01 -0.01 -0.01 -0.01 -0.01 0.01

32.9 -0.01 22.6 0.1 -0.01 -0.010.1 0.02 0.12 0.05 0.05 0.05

0.07 0.28 0.13 0.23 0.25 0.250.01 0.048 0.021 0.049 0.048 0.0533.3 2.66 25.1 5.91 4.07 4.070.13 0.03 0.08 0.04 0.04 0.04

-1 -1 -1 -1 -1 -1

1.6 11.3 9.6 14.5 13 13

-1 31 9 10 8 10

0.2 0.2 -0.2 0.2 -0.2 -0.2

SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

22.5 1 11.5 4 3.5 3.5

-0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.1 -0.1 -0.1 -0.1 -0.1 -0.165 35 50 50 150 150

30 275 75 80 75 70-1 -1 -1 2 -1 -1

930 80 980 580 2080 201040 2020 1930 525 6630 6730

2860 149000 40200 171000 167000 165000

4 141 66 210 506 52913 889 298 349 951 9156 23 420 240 124 1242 183 76 226 516 5227 278 151 672 1460 1560

2.40 9.22 168.34 96.19 49.70 49.70

B20

Maggie Hays

Maggie Hays

Notes: XRF = X-ray florescence, ICP-MS = Inductively coupled plasma mass spectrometry, FA-ICP-MS = Fire assay inductively coupled plasma mass spectrometry, D.L. = analytical reported detection limit, N.D. = not determined, wt% = weight percent, ppm = parts per million, ppb = parts per billion.

B21

Maggie Hays

SampleFGD91-7-

318FGD93-9-

410 LJD3A-231LJD3A-231

Rpt LJD3A-524 LJD4-432 LJD5-384.5Lab Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace

Units D.L. Batch u118354 u118354 u118354 u118354 u118354 u118354 u118354wt% 0.01 Al2O3 4.76 1.53 0.93 0.93 1.21 1.08 0.63wt% 0.01 CaO 6.86 0.17 0.09 0.09 0.35 0.06 0.05wt% 0.01 Fe2O3 8.89 11.4 7.47 7.46 9.08 7.75 8.14wt% 0.01 K2O -0.01 -0.01 -0.01 -0.01 0.16 -0.01 -0.01wt% LOI 5.75 4.16 9.25 9.2 11.4 9.63 6.88wt% 0.01 MgO 27.6 40.6 43.5 43.5 38.1 41.3 44.6wt% 0.01 MnOwt% 0.01 Na2O 0.08 0.05 0.09 0.09 0.06 0.04 0.04wt% 0.001 P2O5 0.01 0.013 0.009 0.009 0.013 0.007 0.007wt% 0.01 SiO2 45.02 41.99 38.68 38.7 39.46 39.77 39.68wt% 0.01 TiO2 0.433 0.119 0.05 0.052 0.086 0.076 0.052

Total 99.3 100 100 100 99.9 99.7 100ppm Bappm Beppm Bippm Cdppm Ceppm Coppm Crppm Csppm Cuppm 0.05 Dy 2.45 0.3 0.2 0.2 0.35 0.15 0.2ppm 0.05 Er 1.3 0.2 0.15 0.1 0.2 0.1 0.1ppm 0.05 Eu 0.25 0.05 -0.05 -0.05 0.1 -0.05 -0.05ppm 0.2 Gappm 0.2 Gd 2.55 0.3 0.2 0.2 0.35 0.1 0.2ppm 0.1 Hf 0.6 0.1 -0.1 -0.1 0.1 -0.1 -0.1ppm 0.02 Ho 0.5 0.05 -0.05 -0.05 0.05 -0.05 -0.05ppm 0.05 La 3.55 0.6 0.15 0.15 2.7 0.1 0.6ppm 0.5 Lippm 0.02 Lu 0.15 -0.05 -0.05 -0.05 -0.05 -0.05 -0.05ppm 0.2 Moppm 0.5 Nb 1.2 0.3 0.2 0.1 0.2 0.1 0.1ppm 0.5 Nd 8.6 0.8 0.35 0.3 2 0.2 0.8ppm Nippm Pbppm 0.02 Pr 1.8 0.15 0.05 -0.05 0.55 -0.05 0.15ppm 0.02 Rb 70 50 40 40 50 -10 20ppm Sbppm Scppm 0.05 Sm 2.3 0.2 0.15 0.15 0.4 0.05 0.2ppm Sn

XRF

ICP

-MS

ppm Snppm 0.1 Sr 80 40 30 40 30 30 30ppm 0.05 Ta -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1ppm 0.02 Tb 0.4 -0.05 -0.05 -0.05 0.05 -0.05 -0.05ppm 0.05 Th 0.65 0.1 0.05 0.05 0.2 -0.05 -0.05ppm Tippm Tlppm Tmppm 0.05 Uppm Vppm Wppm 0.1 Y 12.4 1.95 1.2 1.2 1.9 0.9 1ppm 0.05 Yb 0.95 0.25 0.15 0.15 0.2 0.15 0.1ppm Znppm 1 Zrppm Asppm 20 Ba 80 40 40 40 60 20 40ppm 7 Cr 3430 4830 2000 2010 2270 2430 1850ppm 8 Cu 90 40 20 20 50 -10 -10ppm 8 Ni 1310 2850 2990 3000 4520 2250 3310ppm 10 Rb 70 50 40 40 50 -10 20ppm Scppm 10 Sr 80 40 30 40 30 30 30ppm 40 V 250 80 50 60 60 20 40ppm Yppm 25 Zr 75 35 25 30 30 25 25ppb Auppb 0.2 Ir 1.9 4 1.2 1.4 3 2.7 4.5ppb 0.3 Pd 2 13.5 1 0.5 8 7.5 1ppb 0.3 Pt 3 12.5 1 1 7 6.5 3ppb 0.1 Rh 0.8 2.2 0.4 0.5 1.2 1.2 1.8ppb 0.2 Ru 4.8 11.7 4.1 4.7 4.7 5.6 8.3Wt% CO2wt% S 0.22 0.37 0.2 0.2 0.75 -0.01 0.04

XRF

FA-IC

P-M

S

B22

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD10A-487.9

LJD10A-559.7 LJD11-479 LJD15-344 LJD15-398 LJD51-170 LJD51-344

BSD086-311

Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu118354 u118354 u118354 u118354 u118354 u118354 u118354 u118354

1.22 1.62 0.57 1.44 0.64 14.4 0.51 1.825.25 1.37 0.09 0.72 0.16 6.49 0.02 0.127.46 11.4 7.66 9.18 6.79 9.17 7.09 6.77-0.01 -0.01 -0.01 -0.01 -0.01 0.71 -0.01 -0.017.95 3.78 2.13 10.2 9 0.87 9.21 20.731.6 41.8 47.6 38.6 42.5 3.39 44.7 34.4

0.06 0.05 0.05 0.1 0.21 3.62 0.06 0.030.01 0.016 0.011 0.013 0.008 0.108 0.005 0.002

46.09 39.77 41.99 39.33 40.72 60.05 38.43 36.270.089 0.098 0.042 0.115 0.04 0.838 0.032 0.09199.7 99.8 100 99.6 100 99.6 100 100

1.15 0.5 0.15 0.4 0.15 3.25 0.05 0.30.7 0.25 0.1 0.25 0.1 1.75 0.05 0.20.2 0.05 -0.05 0.1 -0.05 0.85 -0.05 0.05

1 0.4 0.1 0.4 0.15 3.05 -0.05 0.25-0.1 -0.1 -0.1 -0.1 -0.1 3.6 -0.1 0.10.2 0.1 -0.05 0.1 -0.05 0.6 -0.05 0.051.1 1.7 0.3 0.35 0.15 11.2 0.1 0.25

0.1 -0.05 -0.05 -0.05 -0.05 0.2 -0.05 -0.05

0.4 0.4 0.1 0.3 -0.1 3.7 -0.1 -0.12.8 1.7 0.3 0.85 0.3 12 0.05 0.5

0.6 0.35 0.05 0.15 -0.05 2.85 -0.05 0.130 70 40 40 20 100 30 -10

0.9 0.4 0.1 0.3 0.1 2.95 -0.05 0.2SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

50 70 70 40 50 250 30 -10-0.1 -0.1 -0.1 -0.1 -0.1 0.2 -0.1 -0.10.15 0.05 -0.05 0.05 -0.05 0.5 -0.05 -0.050.2 0.2 -0.05 0.1 -0.05 3.9 0.05 -0.05

6.4 2.3 0.85 2.45 1 17.6 0.4 1.80.6 0.2 0.1 0.2 0.1 1.6 0.05 0.2

20 80 80 20 40 540 60 -201930 3030 1730 2210 1720 280 1750 2060

30 40 20 -10 -10 80 -10 -102600 2850 3200 2440 3030 120 2890 2350

30 70 40 40 20 100 30 -10

50 70 70 40 50 250 30 -1070 80 40 80 20 390 20 60

30 50 40 20 25 190 25 -5

2 2.3 3.4 2.4 1.5 -0.1 1.3 4.14.5 12 1 2 2 0.5 0.5 3.52 4 2.5 2.5 3.5 -0.5 1 4

0.5 0.6 0.5 0.5 1.1 -0.1 0.2 0.83.1 4.2 4.2 4.3 3.9 0.1 1.5 5.1

0.3 0.49 0.15 0.11 0.05 0.5 0.02 -0.01

B23

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD51-441.8

LJD52-167.5

LJD52-167.5 Rpt

LJD52-228.8 LJD52-320

LJD54A-248

LJD54A-546

LJD57-150.7


0.67 2.3 2.3 1.19 0.59 5.41 1.32 3.170.03 1.62 1.62 0.04 0.08 10.3 0.86 5.869.26 11.8 11.9 7.74 6.19 14.8 17.4 9.77-0.01 -0.01 -0.01 -0.01 -0.01 0.78 -0.01 -0.012.15 9.36 9.35 8.97 13.9 1.23 5.02 3.1245.7 32.5 32.5 41.6 40 15.8 38 24.8

0.08 0.21 0.21 0.18 0.23 0.56 0.05 0.130.011 0.02 0.021 0.012 0.004 0.048 0.012 0.02641.94 41.59 41.57 39.87 39.28 50.19 37.23 52.780.05 0.152 0.153 0.084 0.034 0.502 0.129 0.22199.8 99.5 99.6 99.6 100 99.6 100 99.8

0.15 0.6 0.6 0.25 -0.05 1.7 0.5 10.1 0.35 0.35 0.15 -0.05 1.05 0.25 0.65

-0.05 0.15 0.15 0.05 -0.05 0.7 0.15 0.15

0.15 0.6 0.6 0.25 -0.05 1.5 0.65 0.85-0.1 0.3 0.3 0.2 -0.1 0.7 0.2 0.4

-0.05 0.1 0.15 0.05 -0.05 0.35 0.1 0.20.15 1.95 2.05 0.35 -0.05 2.7 3.85 1.8

-0.05 0.05 0.05 -0.05 -0.05 0.15 -0.05 0.1

-0.1 0.2 0.2 0.2 -0.1 1.1 0.2 0.50.35 1.6 1.7 0.6 0.05 3.75 3.85 2.55

0.05 0.35 0.4 0.1 -0.05 0.75 0.95 0.5530 30 50 40 -10 140 80 50

0.1 0.45 0.5 0.2 -0.05 1.15 0.75 0.75SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

50 30 30 30 -10 90 80 50-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 0.1 0.1 -0.05 -0.05 0.25 0.1 0.150.05 0.35 0.4 0.1 -0.05 0.3 0.2 0.35

0.85 3.3 3.45 1.45 0.3 9.9 2.45 5.650.1 0.3 0.35 0.15 0.05 0.95 0.2 0.65

40 -20 -20 60 20 320 60 -202000 10700 10900 2060 1720 2040 11700 3650

20 -10 -10 -10 -10 20 130 203610 1720 1740 2990 2400 260 2400 565

30 30 50 40 -10 140 80 50

50 30 30 30 -10 90 80 5060 130 150 60 10 330 170 180

35 25 35 25 -5 90 50 60

1.1 1.2 1.3 1.6 2.1 1.4 3.8 0.51.5 3.5 4 1.5 -0.5 5.5 10 -0.51.5 4 5.5 3.5 3 10 10 10.3 0.7 0.8 0.6 1.5 1.1 1.9 0.31.9 6.3 6.8 5.7 5.6 3 8.4 3

0.02 0.04 0.04 0.2 0.02 0.07 1.95 -0.01

B24

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD57-223.5

LJD57-273.5

LJD57-345.8

LJD57-384.7 LJD61-263

LJD61-347.5 LJD66-130

LJD66-304.2


0.27 0.38 2.06 4.98 0.39 1.46 0.24 4.75-0.01 0.03 0.08 5.16 0.06 2.43 4.74 5.226.24 7.74 8.84 11.9 8.33 7.96 5.35 120.08 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.0114.8 14.9 12.2 10.3 11.7 11 17.7 4.5142.9 43.2 38.2 26.7 43.3 37.5 38.6 26.2

0.24 0.08 0.06 0.06 0.05 0.24 0.06 0.070.005 0.013 0.006 0.045 0.005 0.011 0.005 0.04835.35 36.07 40.3 45.52 36.16 39.09 33.17 46.980.021 0.036 0.13 0.425 0.029 0.105 0.013 0.40999.8 102 101 105 100 99.7 99.8 100

-0.05 -0.05 0.3 1.6 0.1 0.5 0.1 1.65-0.05 -0.05 0.2 0.95 0.05 0.3 -0.05 0.95-0.05 -0.05 -0.05 0.2 -0.05 0.1 -0.05 0.2

-0.05 -0.05 0.2 1.4 0.1 0.4 0.05 1.4-0.1 -0.1 -0.1 0.4 -0.1 -0.1 -0.1 0.4

-0.05 -0.05 0.05 0.3 -0.05 0.1 -0.05 0.3-0.05 0.4 0.15 2.25 0.15 0.5 0.05 2.5

-0.05 -0.05 0.05 0.1 -0.05 -0.05 -0.05 0.1

-0.1 0.2 0.3 0.8 -0.1 0.2 -0.1 0.8-0.05 0.05 0.4 3.45 0.2 1.1 0.05 3.45

-0.05 -0.05 0.05 0.7 -0.05 0.2 -0.05 0.720 -10 -10 60 -10 40 30 60

-0.05 -0.05 0.15 1.15 0.05 0.35 -0.05 1.1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 -10 30 40 30 20 100 40-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 -0.05 0.25 -0.05 0.1 -0.05 0.250.05 -0.05 0.1 0.25 -0.05 0.05 -0.05 0.25

0.15 0.15 1.8 8.45 0.6 2.65 0.4 8.30.05 -0.05 0.2 0.85 0.05 0.25 -0.05 0.8

60 40 40 60 40 20 40 401490 1670 2050 3320 1830 1680 1130 3270-10 -10 -10 40 -10 -10 -10 30

2940 2660 2440 1300 2990 2590 2630 134020 -10 -10 60 -10 40 30 60

-10 -10 30 40 30 20 100 4020 30 50 240 30 70 -10 250

10 10 30 75 25 20 10 60

1.5 2 1.9 1.9 1.9 2.2 2.6 1.81.5 1.5 4 6.5 1 2 2.5 62 4.5 3.5 8 2.5 2.5 2 8

0.7 0.7 0.5 1.1 0.7 0.5 0.8 1.13.8 2.9 4.8 5.4 4.7 4 5.7 5

0.06 0.04 0.06 0.43 0.02 0.04 0.08 -0.01

B25

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD71-242.5

LJD71-299.5

BSD091-113.3

BSD091-113.3 Rpt LJD79-147 LJD79-210 LJD79-245 LJD80-142


0.28 0.43 1.08 1.08 0.97 1.46 2.02 0.30.02 0.02 0.32 0.32 2.85 2.4 2.56 0.417.07 7.75 6.01 5.99 7.17 11.5 11.5 7.02-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 0.05 -0.0113.3 7.62 30.6 30.5 13.3 8.9 4.77 17.343.1 45.1 33.7 33.7 37 34.5 36.6 40.6

0.17 0.25 0.02 0.02 0.26 0.15 0.08 0.260.009 0.013 0.006 0.006 0.014 0.015 0.022 0.00536.03 38.68 28.08 27.97 38.07 40.66 42.14 33.70.021 0.027 0.037 0.036 0.071 0.091 0.173 0.02599.9 99.8 99.8 99.6 99.6 99.6 99.9 99.6

0.05 -0.05 0.15 0.15 0.3 0.65 0.85 0.05-0.05 -0.05 0.1 0.1 0.15 0.35 0.5 -0.05-0.05 -0.05 -0.05 -0.05 -0.05 0.15 0.25 -0.05

0.05 -0.05 0.1 0.1 0.25 0.6 0.7 0.05-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 -0.05 -0.05 0.05 0.15 0.15 -0.050.1 0.05 0.1 0.15 0.7 0.75 1.2 0.1

-0.05 -0.05 -0.05 -0.05 -0.05 -0.05 0.05 -0.05

-0.1 -0.1 -0.1 0.1 0.1 0.2 0.4 -0.10.1 0.05 0.25 0.25 0.65 1.6 2.3 0.1

-0.05 -0.05 -0.05 -0.05 0.1 0.3 0.5 -0.05-10 20 -10 -10 50 30 80 -10

-0.05 -0.05 0.05 0.1 0.15 0.5 0.6 -0.05SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 40 -10 -10 50 30 60 20-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 -0.05 -0.05 0.05 0.1 0.1 -0.05-0.05 0.05 -0.05 -0.05 -0.05 0.15 0.1 -0.05

0.3 0.3 0.95 1 1.6 3.25 4.1 0.550.05 0.05 0.1 0.1 0.15 0.3 0.4 0.1

40 80 -20 -20 40 -20 60 401530 1810 1600 1580 4800 3680 4490 1390-10 -10 -10 -10 -10 -10 20 -10

2830 3100 2170 2160 2410 2690 2150 2930-10 20 -10 -10 50 30 80 -10

-10 40 -10 -10 50 30 60 2020 30 10 -10 50 90 120 10

10 15 -5 -5 20 15 50 10

2.2 2.3 8.5 9.8 2.9 3 2.9 1.71 1.5 2 2.5 2 2 3.5 3.5

1.5 2.5 3 3 2 3 7 5.50.5 0.5 0.6 0.5 0.8 0.7 1.1 0.64.2 3.7 5.8 5.5 4.9 6.8 8.4 4.1

0.08 0.05 -0.01 -0.01 0.13 0.09 0.02 0.29

B26

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD80-199.5 LJD81-79.5

LJD81-161.5

LJD104W1-126.7

LJD104W1-212

LJD107-445

LJD120-229.6

LJD120-257


0.48 0.33 0.47 0.2 0.54 0.72 4.23 0.69-0.01 0.03 0.05 0.02 0.04 0.07 9.08 0.497.33 6 7.62 6.16 6.89 8.52 8.77 7.12-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.0112.5 17.3 13.7 14.5 14.5 4.35 4.7 12.342.7 41.4 42.6 42.9 43.7 46.2 25 42.3

0.13 0.03 0.32 0.17 0.03 0.1 0.1 0.050.006 0.005 0.007 0.004 0.005 0.013 0.026 0.00736.46 34.61 34.84 36.03 34.38 40.09 47.4 36.950.026 0.019 0.04 0.01 0.042 0.063 0.335 0.0599.6 99.7 99.6 99.9 100 100 99.6 99.9

0.05 0.1 -0.05 -0.05 -0.05 0.4 1.85 0.150.05 0.05 0.05 -0.05 -0.05 0.2 1 0.1-0.05 -0.05 -0.05 -0.05 -0.05 0.05 0.2 -0.05

0.05 0.1 -0.05 -0.05 -0.05 0.5 1.5 0.1-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 0.4 -0.1

-0.05 -0.05 -0.05 -0.05 -0.05 0.05 0.3 -0.050.15 0.2 -0.05 0.1 -0.05 0.95 1.25 0.1

-0.05 -0.05 -0.05 -0.05 -0.05 -0.05 0.1 -0.05

-0.1 -0.1 0.1 -0.1 -0.1 0.2 0.6 0.10.25 0.3 -0.05 -0.05 -0.05 1.85 2.95 0.3

-0.05 0.05 -0.05 -0.05 -0.05 0.35 0.6 -0.05-10 -10 -10 -10 -10 40 70 30

0.1 0.1 -0.05 -0.05 -0.05 0.4 1.05 0.1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 -10 20 -10 20 50 60 30-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 -0.05 -0.05 -0.05 0.05 0.25 -0.050.05 -0.05 0.1 -0.05 -0.05 0.1 0.4 -0.05

0.55 0.65 0.35 0.15 0.1 1.95 8.8 0.850.05 -0.05 0.1 -0.05 -0.05 0.15 0.85 0.1

40 -20 60 40 40 100 60 401690 1360 1890 1280 1370 1920 2940 1550-10 -10 -10 -10 -10 20 20 -10

2670 2480 3260 2420 2730 3020 565 2470-10 -10 -10 -10 -10 40 70 30

-10 -10 20 -10 20 50 60 30-10 -10 20 -10 20 40 220 20

5 -5 10 10 10 35 60 15

1.2 1.6 2.1 1.8 2.8 3.1 1.2 0.91.5 0.5 2.5 1 5.5 5.5 4 22 1 3.5 1 4.5 6.5 8 2

0.8 0.3 0.5 0.1 1.2 0.7 0.4 0.32.8 3.5 3 0.9 4.6 4.6 1.5 1.3

0.05 0.09 0.16 0.23 0.03 0.11 0.02 0.18

B27

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


MHD94-3-253

MHD94-3-440

MHD94-3-440 Rpt

FGD91-368.5

FGD93-9-364

LJD0003A-300.5

LJD0003A-454.5

LJD0003A-501


1.49 0.51 0.51 0.5 2.1 0.81 8.01 1.380.65 -0.01 -0.01 0.02 1.92 0.1 20 0.269.45 9.11 9.13 13 11.3 8.83 12.3 8.56-0.01 -0.01 -0.01 -0.01 0.34 -0.01 1.27 -0.017.2 0.88 0.82 7.94 9.23 10.7 1.23 11.6

42.7 48 47.9 41 33.8 39.6 8 37.7

0.12 0.05 0.05 0.08 0.07 0.06 0.4 0.080.015 0.017 0.018 0.008 0.007 0.018 0.025 0.0138.3 41.7 41.61 35.88 39.61 38.46 47.31 39.34

0.107 0.05 0.053 0.041 0.166 0.063 0.263 0.114100 100 100 98.4 98.5 98.6 98.8 99

0.4 -0.05 -0.05 -0.05 0.65 0.2 2.45 0.350.3 -0.05 -0.05 -0.05 0.3 0.15 1.2 0.20.1 -0.05 -0.05 -0.05 0.15 0.05 0.5 -0.05

0.3 -0.05 -0.05 -0.05 0.65 0.25 2.6 0.30.2 -0.1 -0.1 -0.1 -0.1 0.1 1.3 0.10.1 -0.05 -0.05 -0.05 0.15 -0.05 0.5 0.1

0.75 -0.05 -0.05 0.1 0.7 2.55 6.05 0.55

-0.05 -0.05 -0.05 -0.05 -0.05 -0.05 0.15 -0.05

0.3 0.1 0.1 0.1 0.3 0.2 2.2 0.30.9 -0.05 -0.05 0.1 1.95 1.55 9.8 0.75

0.2 -0.05 -0.05 -0.05 0.35 0.5 2.15 0.1540 50 60 50 50 -10 160 40

0.25 -0.05 -0.05 -0.05 0.5 0.25 2.35 0.2SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

30 40 70 30 20 20 320 30-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 0.3 -0.10.05 -0.05 -0.05 -0.05 0.1 -0.05 0.4 0.050.15 0.05 -0.05 0.05 0.1 0.25 16.6 0.5

2.65 0.2 0.15 0.2 3.4 1.2 12.7 2.150.25 0.05 0.05 -0.05 0.3 0.1 1 0.2

40 60 60 40 100 300 540 4601800 2450 2430 1110 4470 1650 20 2100-10 -10 20 300 60 60 210 20

2900 3950 3960 4490 3100 2480 3250 248040 50 60 50 50 -10 160 40

30 40 70 30 20 20 320 3030 60 60 20 100 50 150 40

30 30 40 5 15 5 165 10

4.2 3.4 3.1 3.5 4.1 4 -0.1 5.72 1 1 33 14.5 1.5 -0.5 7.5

3.5 4 3.5 21 11.5 2.5 0.5 120.5 0.4 0.4 2 2.2 0.8 -0.1 1.92.8 4.2 4.2 6.5 13.2 5.7 0.2 9.9

0.13 0.04 0.04 1.51 0.73 0.26 1.49 0.27

B28

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0003A-555

LJD0004-361.7

LJD0005-300.3

LJD0005-332.7

LJD0005-414

LJD0011-464.6

LJD0015-479

BSD086-167.3


2.24 0.75 2.53 1.28 0.49 0.65 0.83 5.185.98 0.43 1.41 1.1 0.06 -0.01 0.12 5.629.81 7.93 12.1 9.73 9.49 8.12 9.1 16.30.02 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.016.98 9.29 8.48 9.3 4.21 1.41 8.52 20.629.7 41.5 34 39.9 45 46.9 42 12.7

0.11 0.11 0.04 0.09 0.05 0.05 0.04 0.020.007 0.008 0.022 0.014 0.006 0.009 0.009 0.01643.5 38.63 40.12 37.5 39.73 41.96 38.1 37.32

0.186 0.058 0.179 0.097 0.046 0.052 0.065 0.25298.5 98.6 98.8 99 99 99.1 98.7 97.9

0.9 0.25 0.6 0.4 0.05 0.05 0.2 1.20.4 0.1 0.3 0.25 -0.05 0.05 0.1 0.8

0.15 0.05 0.15 0.1 -0.05 -0.05 -0.05 0.3

0.85 0.2 0.5 0.35 0.05 0.05 0.2 10.2 -0.1 0.2 -0.1 -0.1 -0.1 -0.1 0.5

0.15 -0.05 0.1 0.1 -0.05 -0.05 -0.05 0.251.3 0.35 1.05 0.4 0.1 0.05 0.5 1.15

0.05 -0.05 0.05 0.05 -0.05 -0.05 -0.05 0.1

0.6 0.1 0.4 0.2 -0.1 0.1 0.1 0.33.5 0.65 1.25 0.9 0.15 0.1 0.7 2.2

0.75 0.15 0.25 0.15 -0.05 -0.05 0.15 0.450 -10 30 40 50 -10 -10 20

0.8 0.15 0.4 0.25 -0.05 -0.05 0.2 0.65SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

50 30 20 50 40 30 -10 80-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.10.15 -0.05 0.1 0.05 -0.05 -0.05 -0.05 0.150.35 0.05 0.2 0.05 -0.05 -0.05 -0.05 0.1

4.3 1.2 3.25 2.5 0.3 0.4 1.25 8.20.4 0.1 0.3 0.25 0.05 0.05 0.1 0.7

40 60 -20 60 80 80 -20 -203190 2090 5400 2050 2150 1560 2160 4170

50 -10 20 -10 20 30 -10 502250 2560 1350 945 3430 2920 3280 1770

50 -10 30 40 50 -10 -10 20

50 30 20 50 40 30 -10 80110 30 90 50 40 40 30 180

30 5 15 15 20 25 -5 20

4.4 2.6 0.4 0.5 3 5.4 2.9 3.411.5 2.5 1 1 2.5 12 6 7.510.5 3 1 1 3 13 5.5 92.7 0.4 0.3 0.2 1.1 2 0.9 1.68.1 3.6 3 2.2 6.7 10.6 5.6 7.5

0.9 0.12 0.33 0.12 0.02 0.21 0.04 0.09

B29

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


BSD086-167.3 Rpt

LJD0051-84.6

LJD0051-135

LJD0051-224

LJD0051-272

LJD0051-289

LJD0051-289 Rpt

LJD0052-100.5


5.19 8.61 10.6 12.9 6.63 1.48 1.46 14.45.62 5.54 9.55 9.92 12.1 0.28 0.28 7.5816.3 12 12.4 11.5 11.5 9.11 9.13 10.2-0.01 0.64 0.12 0.66 0.7 -0.01 -0.01 0.3520.6 3.13 0.78 2.02 1.16 11.2 11.2 1.4112.7 13 8.88 2.38 15.4 37.3 37.2 5.81

0.03 0.76 1.92 0.87 0.26 0.14 0.14 2.560.015 0.061 0.078 0.136 0.045 0.013 0.013 0.07737.36 54.91 54.45 58.14 51.13 39.25 39.18 56.030.251 0.603 0.736 1.149 0.495 0.122 0.119 0.66

98 99.2 99.5 99.6 99.4 98.8 98.7 99

1.15 2.35 2.95 4.35 1.95 0.5 0.5 2.550.75 1.25 1.65 2.45 1.05 0.25 0.25 1.450.3 0.5 0.75 1.15 0.55 0.05 0.05 0.75

1.05 2.1 2.7 4.35 1.7 0.35 0.4 2.350.5 1.6 2.1 3.3 1 0.1 0.1 2.2

0.25 0.5 0.65 0.9 0.35 0.1 0.1 0.551.1 2.15 6.65 12.7 2.45 0.6 0.55 2.85

0.1 0.15 0.2 0.3 0.15 -0.05 -0.05 0.2

0.3 1.4 2 3.5 1.2 0.3 0.3 1.92.1 5.1 7.75 14.3 4.3 0.9 0.9 6

0.35 1 1.75 3.4 0.9 0.2 0.2 1.2520 80 70 90 120 30 20 60

0.7 1.5 2.1 3.5 1.3 0.3 0.3 1.8SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

80 60 150 190 110 20 20 210-0.1 -0.1 0.1 0.2 -0.1 -0.1 -0.1 0.10.2 0.35 0.5 0.7 0.25 0.05 0.05 0.350.1 1.05 1.55 3.55 0.9 0.05 0.05 1.95

8.25 12.1 16 23.1 9.9 2.6 2.55 13.40.7 1.2 1.55 2.25 0.95 0.2 0.25 1.25

-20 140 60 180 120 -20 -20 1404210 1540 735 -5 3070 2060 2080 220

40 30 70 30 130 -10 -10 201760 260 110 -5 370 2460 2430 100

20 80 70 90 120 30 20 60

80 60 150 190 110 20 20 210180 330 410 530 340 70 80 340

15 90 125 200 80 10 10 115

3 1.1 -0.1 -0.1 -0.1 3.3 2.4 -0.17 6 -0.5 -0.5 1.5 2.5 3 -0.5

8.5 8 1 0.5 2.5 4 4.5 0.51.6 0.9 -0.1 -0.1 0.2 1 1 -0.17.7 3.2 0.3 0.2 0.4 4.1 4.5 0.3

0.08 0.01 0.11 0.14 0.01 0.11 0.12 0.01

B30

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0052-381.2

LJD0052-419

LJD0052-425

LJD0054A-208.3

LJD0054A-287

LJD0054A-333

LJD0054A-372.5

LJD0061-93


1.21 4.5 4.38 2.96 2.42 2.32 1.02 0.370.32 4.54 6.98 4.5 9.09 1.68 0.04 -0.0112.1 12.1 11 13.5 12.7 7.54 7.91 6.88-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.016.11 4.87 4.28 3.62 4.81 11 11.7 14.739.6 27.9 25.3 31.8 27.1 36.7 40 41.2

0.23 0.05 0.07 0.09 0.15 0.07 0.33 0.110.01 0.01 0.039 0.027 0.025 0.016 0.007 0.007

38.95 44.74 46.78 42.13 42.08 39.64 37.78 35.460.095 0.379 0.373 0.252 0.208 0.125 0.068 0.02998.6 99 99.1 98.8 98.5 99 98.8 98.7

0.25 1.15 1.4 0.85 0.85 0.25 0.15 0.050.15 0.75 0.75 0.55 0.4 0.15 0.1 -0.050.05 0.25 0.3 0.15 0.15 0.1 -0.05 -0.05

0.25 1.2 1.3 0.85 0.8 0.2 0.15 0.10.1 0.2 0.5 0.3 0.3 0.1 -0.1 -0.1

0.05 0.25 0.25 0.15 0.15 -0.05 -0.05 -0.050.65 2.9 2 1.1 1.05 0.3 0.15 -0.05

-0.05 0.1 0.1 0.05 0.05 -0.05 -0.05 -0.05

0.2 0.7 0.6 0.6 0.4 0.8 0.1 -0.10.75 4.2 3.85 2 1.7 0.5 0.3 0.15

0.15 1 1 0.4 0.35 0.1 0.05 -0.0540 40 50 50 50 -10 -10 -10

0.2 1 0.95 0.6 0.6 0.15 0.1 -0.05SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

30 20 30 70 100 -10 -10 -10-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 0.2 0.2 0.15 0.1 -0.05 -0.05 -0.050.05 0.2 0.2 0.1 0.1 0.1 -0.05 -0.05

1.55 6.5 7 4.55 4.2 1.4 0.95 0.350.15 0.7 0.7 0.5 0.4 0.15 0.1 -0.05

40 -20 -20 40 -20 -20 -20 -204940 3410 3170 4870 5500 1670 2250 1480

20 40 20 50 40 -10 -10 -102360 1450 1230 2000 1220 2010 2360 2540

40 40 50 50 50 -10 -10 -10

30 20 30 70 100 -10 -10 -1080 190 240 170 140 20 -10 -10

15 35 35 45 45 5 -5 -5

2.8 1.9 1.7 1.3 1.3 1.8 0.8 2.31 3 7.5 8.5 4.5 1.5 -0.5 5

1.5 6.5 10 7 6 2.5 1 6.50.7 1.1 1.3 1.2 1 0.4 0.2 1.18.6 5.8 5.4 6.9 6.8 3.5 1.1 6.6

0.2 0.29 0.28 0.73 0.75 0.27 0.12 0.18

B31

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0061-181.5

LJD0066-228

LJD0071-98

LJD0071-193

LJD0079-259

LJD0080-261

BSD088-372.8

BSD088-372.8 Rpt


0.32 0.8 0.3 0.33 4.76 1.25 9.45 9.40.02 3.9 0.1 -0.01 5.65 0.42 7.9 7.98.04 7.27 8.94 7.24 11.6 8.55 10.2 10.3-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 0.02 0.0212.1 11 9.71 16.2 4.75 9.62 15.4 15.441.8 34.9 43.1 39.8 26.2 38.2 11.7 11.8

0.11 0.12 0.04 0.16 0.06 0.17 1.31 1.290.006 0.012 0.006 0.006 0.042 0.012 0.037 0.03736.46 40.66 37.13 34.48 45.48 40.67 42.79 42.850.03 0.038 0.021 0.021 0.386 0.09 0.472 0.47498.8 98.6 99.3 98.2 98.9 98.9 99.2 99.4

-0.05 0.15 0.1 0.05 1.55 0.35 0.75 0.75-0.05 0.1 0.05 -0.05 0.85 0.2 0.35 0.4-0.05 0.05 -0.05 -0.05 0.2 0.05 0.3 0.3

-0.05 0.15 0.1 -0.05 1.5 0.35 0.9 0.9-0.1 -0.1 -0.1 -0.1 0.7 -0.1 0.7 0.8

-0.05 -0.05 -0.05 -0.05 0.3 0.1 0.15 0.15-0.05 0.85 0.15 0.1 2.45 0.25 1.85 1.75

-0.05 -0.05 -0.05 -0.05 0.1 -0.05 0.1 0.1

-0.1 0.1 -0.1 -0.1 0.8 0.2 0.7 0.70.1 0.7 0.2 0.15 3.2 0.75 3.6 3.4

-0.05 0.15 -0.05 -0.05 0.7 0.1 0.75 0.7-10 -10 -10 -10 50 -10 -10 20

-0.05 0.1 0.05 -0.05 1 0.25 0.95 0.95SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 20 -10 -10 30 -10 80 100-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 -0.05 -0.05 0.25 0.05 0.1 0.1-0.05 -0.05 -0.05 -0.05 0.15 -0.05 0.15 0.15

0.3 0.8 0.55 0.35 8.85 2.05 3.6 3.60.05 0.05 0.05 -0.05 0.8 0.2 0.5 0.5

-20 40 -20 -20 -20 -20 -20 -201580 1280 1490 1480 3170 2080 870 885-10 180 -10 -10 50 -10 40 40

3240 2970 3060 2590 1290 2700 195 205-10 -10 -10 -10 50 -10 -10 20

-10 20 -10 -10 30 -10 80 100-10 30 -10 -10 230 40 290 300

-5 -5 -5 -5 45 -5 25 30

1.3 5.1 4 4.5 1.7 2.5 0.2 0.10.5 5 1 1 6 2 10.5 110.5 7.5 1.5 2 7.5 2.5 10 100.5 4.8 0.5 0.5 1 0.6 1.2 1.13.9 5.6 5.1 8.1 5.7 5 2 1.9

0.05 0.39 0.05 0.14 0.15 0.16 0.01 0.01

B32

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0081-205

LJD0104W1-285.5

LJD0104W1-329.5

LJD0107-392

LJD0107-491

LJD0107-520.5

LJD0120-311.5

MHD94-3-122


1.01 0.87 4.73 0.56 3.41 4.27 0.67 2.420.31 0.27 5.65 0.07 4.42 6.68 0.11 6.8711.9 14.6 11.2 7.93 11.8 11.3 8.22 9.95-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.0110.5 7.07 5.02 6.76 4.86 4.55 11.7 6.4436 38.8 26.8 44.6 31.3 25.2 41.3 28.9

0.16 0.05 0.08 0.06 0.06 0.07 0.04 0.150.02 0.01 0.013 0.012 0.024 0.035 0.016 0.019

38.83 37.21 45.38 39.19 42.98 46.74 37.2 44.170.08 0.074 0.384 0.043 0.274 0.359 0.065 0.21398.8 98.9 99.2 99.2 99.1 99.1 99.3 99.1

0.2 0.2 1.45 0.05 1 1.8 0.15 0.80.15 0.15 0.8 -0.05 0.55 0.9 0.15 0.5-0.05 -0.05 0.25 -0.05 0.25 0.25 -0.05 0.15

0.15 0.2 1.3 -0.05 0.95 1.7 0.2 0.7-0.1 -0.1 0.3 -0.1 -0.1 0.5 -0.1 0.1

-0.05 -0.05 0.25 -0.05 0.2 0.35 -0.05 0.150.2 0.3 2.5 0.05 1.55 3.4 1.25 0.95

-0.05 -0.05 0.1 -0.05 0.05 0.15 -0.05 0.05

0.1 0.2 0.7 -0.1 0.6 0.6 0.3 0.30.35 0.5 3.45 0.1 2.85 4.3 1.05 1.65

0.1 0.1 0.85 -0.05 0.7 1 0.25 0.330 50 50 -10 50 50 -10 40

0.15 0.15 0.95 -0.05 0.75 1.2 0.2 0.5SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 30 40 -10 20 40 -10 70-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 0.2 -0.05 0.15 0.25 -0.05 0.1-0.05 -0.05 0.2 -0.05 0.05 0.2 0.05 -0.05

1.25 1.2 7 0.4 5.3 8.95 1.05 4.150.2 0.1 0.7 0.05 0.5 0.85 0.1 0.4

-20 -20 -20 -20 -20 -20 -20 404910 915 3180 1830 3990 3170 1460 2610-10 290 40 -10 30 60 -10 60

2430 4500 1250 2810 1740 1270 2690 197030 50 50 -10 50 50 -10 40

-10 30 40 -10 20 40 -10 7080 40 220 20 140 240 -10 110

5 10 50 -5 25 50 -5 25

2.4 2.5 1.7 2 2.3 1.8 2.5 1.81 41 5 0.5 7 5.5 0.5 3

2.5 24 8 1 7 7.5 2 50.7 2.1 1.1 0.8 1 1.1 0.4 0.87.8 5.8 5 4.2 7.9 5.5 4.9 4.5

0.1 2.33 0.42 0.03 0.42 0.36 0.22 0.14

B33

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


MHD94-3-196.5

MHD94-3-287.5

MHD94-3-375

MHD94-5-243

MHD94-5-428

MHD94-6-249.5

LJD3A-319.5

LJD3A-319.5 Rpt


1.54 0.75 0.79 2.03 2.52 2.27 0.82 0.830.56 0.07 0.03 3.03 6.43 1.72 4.44 4.478.96 8.03 5.05 12.6 10.5 7.67 11.5 11.5-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.018.45 13.3 5.98 8.98 4 6.35 4.43 4.4137.9 38.9 44.2 32.6 29.7 39.3 31.3 31.4

0.05 0.03 0.07 0.03 0.05 0.05 0.06 0.070.014 0.01 0.005 0.017 0.01 0.024 0.015 0.01541.36 37.91 43.19 39.46 45.62 41.41 46.07 46.140.123 0.048 0.046 0.157 0.233 0.198 0.127 0.12898.9 99 99.3 98.8 99 98.9 98.7 98.9

0.3 0.1 0.2 0.55 0.9 0.7 0.35 0.30.2 0.1 0.15 0.3 0.5 0.35 0.15 0.2

-0.05 -0.05 -0.05 0.1 0.2 -0.05 0.15 0.15

0.25 0.1 0.1 0.5 0.8 0.6 0.3 0.30.2 -0.1 -0.1 0.1 0.2 0.2 0.1 0.1

0.05 -0.05 0.05 0.1 0.15 0.1 0.05 0.050.35 0.15 0.4 0.65 1.15 0.35 0.95 1

-0.05 -0.05 -0.05 0.05 0.05 0.05 -0.05 -0.05

0.3 0.2 -0.1 0.3 0.3 0.5 0.3 0.30.65 0.2 0.3 1.1 2.05 0.85 1.2 1.15

0.1 0.05 0.1 0.2 0.4 0.15 0.25 0.3-10 -10 -10 50 40 -10 40 50

0.15 0.1 0.05 0.3 0.65 0.35 0.25 0.3SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 -10 -10 30 40 20 20 30-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1

-0.05 -0.05 -0.05 0.05 0.15 0.1 0.05 0.050.1 -0.05 -0.05 0.05 0.05 0.2 -0.05 -0.05

1.85 0.6 1.1 2.9 4.45 3.5 1.85 1.850.2 0.1 0.1 0.3 0.4 0.35 0.15 0.2

-20 -20 60 -20 40 40 -20 -201840 1670 1470 6110 4020 1210 4380 4430-10 -10 -10 30 20 20 30 30

2660 2810 3100 1210 1520 2830 2120 2120-10 -10 -10 50 40 -10 40 50

-10 -10 -10 30 40 20 20 3050 20 -10 110 170 80 90 90

-5 -5 -5 20 30 15 10 15

1.1 1.2 1.2 0.5 2.1 1.6 6.8 7.12 0.5 1 2.5 4 2 25.5 252 1 2 3.5 9 5 24.5 24

0.4 0.3 0.5 0.4 1.6 0.9 8.3 8.22.7 2.9 3.3 3.9 6.7 3.1 28.7 30

0.1 0.16 0.05 0.14 0.11 0.39 0.45 0.44

B34

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


FGD92-8 177.00

LJD0004 332.00

LJD0004 413.00

LJD0004 520.50

LJD0005 260.50

LJD0005 356.20

LJD0009 771.00

LJD0010A 371.50


1.15 19.5 0.77 0.99 1.62 0.57 13.5 4.480.3 0.02 0.08 0.33 0.79 1.12 0.58 5.97

10.8 12.3 7.97 8.79 8.88 7.28 0.98 10.1-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 2.3 -0.0111.8 11.6 6.6 2.19 9.28 13.9 0.7 4.7438 27.9 42.3 44.3 38.5 40.4 0.12 27.8

0.1 -0.01 0.63 0.04 0.03 0.06 4.73 0.030.017 0.005 0.007 0.007 0.014 0.007 0.008 0.03136.9 28.6 41.3 43.2 40.1 36 76.9 45.90.1 0.03 0.05 0.07 0.11 0.04 0.02 0.35

99.1 99.9 99.6 99.9 99.3 99.3 99.8 99.3

-0.5 -0.5 -0.5 -0.5 -0.5 -0.5 3 1.5-0.5 -0.5 -0.5 -0.5 -0.5 -0.5 2 1-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 0.6 -0.2

-2 -2 -2 -2 -2 -2 2 -2-0.1 6.9 -0.1 -0.1 -0.1 -0.1 3.8 0.7-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 0.6 0.4-0.5 1 -0.5 -0.5 0.5 -0.5 14 1.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 0.4 -0.2

0.2 1.4 0.1 0.2 0.2 -0.1 5.3 0.70.5 1.5 -0.5 -0.5 1 -0.5 12 3

-0.2 0.4 -0.2 -0.2 -0.2 -0.2 3 0.6-10 -10 -10 -10 -10 -10 50 -10

-0.5 -0.5 -0.5 -0.5 -0.5 -0.5 3 1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 -10 -10 20 -10 -10 30 -10-0.1 0.7 -0.1 -0.1 -0.1 -0.1 0.5 -0.1-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 0.6 0.20.15 27.2 0.4 0.35 0.35 0.25 14.3 3.65

1.85 3.25 1 1.8 2.45 1 20.3 7.95-0.5 0.5 -0.5 -0.5 -0.5 -0.5 2 0.5

-0.01 -0.01 -0.01 0.01 -0.01 -0.01 0.1 -0.011977 -7 1293 1607 1676 1204 2326-10 -10 -10 20 20 -10 -10 40

2770 340 2350 3010 3000 2840 -10 1390-10 -10 -10 -10 -10 -10 50 -10

-10 -10 -10 20 -10 -10 30 -1017 13 -10 13 17 -10 -10 96

3.2 -0.1 3 9.9 9.1 2.4 -0.1 2.72.5 -0.5 4.5 12.5 22.5 4 -0.5 9.55 -0.5 6 17.5 25 8.5 0.5 16.5

0.8 -0.1 1.5 5.6 4.5 1.9 0.2 3.35.7 -0.1 7 25.5 21.6 5.4 0.2 9.2

0.07 0.01 0.07 0.14 0.17 0.12 0.02 0.07

B35

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0010A 425.60

LJD0010A 538.40

LJD0011 713.00

LJD0011 716.00

LJD0011 748.00

LJD0011 748.00 Rpt

LJD0011 752.10

LJD0011 780.20


2.5 1.39 3.53 4.61 2.88 2.87 5.73 3.642.07 1.97 11.7 10.3 2.64 2.64 10.8 10.111.2 9.31 26.1 9.63 12.1 12 13.4 10.9-0.01 -0.01 0.05 0.02 -0.01 -0.01 0.67 0.028.43 2.6 0.54 3.49 6.96 6.96 1.25 3.4234.8 41.1 9.96 22.4 34.8 34.9 16.2 22.6

0.03 0.05 0.27 0.18 0.04 0.04 0.47 0.170.023 0.007 0.154 0.034 0.026 0.026 0.045 0.032

40 43.6 47.1 48.6 39.7 39.8 50.3 48.50.22 0.09 0.28 0.33 0.26 0.26 0.56 0.3199.2 100 99.6 99.5 99.3 99.4 99.4 99.6

0.5 -0.5 2 1.5 1 1 2 1.5-0.5 -0.5 1 1 0.5 0.5 1 1-0.2 -0.2 1.2 0.4 -0.2 -0.2 0.6 0.2

-2 -2 -2 -2 -2 -2 -2 -20.3 0.2 0.7 0.5 0.2 0.2 0.6 0.3-0.2 -0.2 0.4 0.4 -0.2 -0.2 0.4 0.2

1 -0.5 5 1.5 0.5 0.5 2 1.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.5 0.3 0.6 0.9 0.4 0.4 0.9 0.31.5 0.5 5.5 3 1.5 1.5 4 3

0.2 -0.2 1.2 0.6 0.2 0.4 0.8 0.6-10 -10 40 -10 -10 -10 70 -10

-0.5 -0.5 1.5 1 0.5 0.5 1.5 1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 20 50 30 -10 -10 70 40-0.1 -0.1 -0.1 0.4 -0.1 -0.1 -0.1 -0.1-0.2 -0.2 0.4 0.2 -0.2 -0.2 0.4 0.22.1 1.3 2.1 1.15 0.85 0.75 0.9 0.7

3.4 1.95 12.7 7.85 4.7 4.8 11.2 7.75-0.5 -0.5 1 0.5 -0.5 -0.5 1 0.5

-0.01 0.01 -0.01 -0.01 -0.01 0.01 0.03 0.012511 1587 930 1813 2032 2025 2155 2155-10 30 30 20 -10 -10 20 70

2120 2500 540 940 1970 1990 530 1360-10 -10 40 -10 -10 -10 70 -10

-10 20 50 30 -10 -10 70 4070 21 57 96 61 61 149 87

3.2 2.4 1.2 0.8 2.7 2.5 1.6 1.84 1 5.5 1.5 4 4 6 4.57 5 11.5 1.5 5.5 6 10.5 6.5

1.6 1.1 1.6 0.2 0.8 0.9 1.2 17.9 4.8 4.3 1.2 4.7 5 4.4 4.7

0.08 0.25 0.01 0.07 0.03 0.03 -0.01 0.6

B36

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0011 785.60

LJD0011 793.50

LJD0011 801.00

LJD0017 183.00

LJD0017 210.60

LJD0017 222.00

LJD0018 174.80

LJD0018 225.50


3.99 4.15 4.96 1.03 3.6 17.9 0.91 0.76.42 8.37 9.33 0.91 5.39 1.61 0.25 10.111.4 10.4 11.8 8.12 8.02 0.7 8.44 7-0.01 0.03 0.09 -0.01 -0.01 1.48 -0.01 -0.014.17 3.22 3.71 11.1 3.77 0.82 8.24 20.226.4 25.7 22.7 38.5 26.3 0.81 42.7 31.1

0.11 0.16 0.25 0.09 0.08 8.08 0.04 0.050.027 0.032 0.03 0.01 0.012 0.24 0.006 0.00346.5 47 46 39.8 52.1 67.5 39 30.10.37 0.37 0.44 0.06 0.32 0.79 0.07 0.0699.3 99.4 99.3 99.6 99.5 99.9 99.6 99.3

1 1.5 1.5 -0.5 1 1.5 -0.5 0.50.5 0.5 1 -0.5 0.5 1 -0.5 -0.50.2 0.2 0.4 -0.2 0.2 0.8 -0.2 -0.2

-2 -2 -2 -2 -2 2 -2 -20.3 0.3 0.4 0.1 0.4 5.9 0.3 0.20.2 0.2 0.4 -0.2 0.2 0.4 -0.2 -0.21 1 1.5 0.5 1.5 16.5 -0.5 0.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

1.3 0.7 0.6 0.1 0.5 1.8 0.6 0.12.5 2.5 3 0.5 4.5 12 -0.5 1.5

0.4 0.4 0.6 -0.2 0.8 3.2 -0.2 0.2-10 20 20 -10 -10 -10 -10 -10

1 1 1 -0.5 1 2.5 -0.5 -0.5SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

40 60 80 -10 -10 340 -10 20-0.1 -0.1 -0.1 -0.1 -0.1 0.2 -0.1 -0.1-0.2 0.2 0.2 -0.2 -0.2 0.2 -0.2 -0.20.65 0.6 0.65 0.4 0.65 9.85 0.95 0.7

6.35 6.75 7.8 1.45 6.25 8.2 1.5 2.80.5 0.5 1 -0.5 0.5 1 -0.5 -0.5

-0.01 -0.01 -0.01 -0.01 -0.01 0.05 -0.01 -0.012346 2189 2620 1423 2230 14 1224 1327

60 90 80 30 20 20 -10 -101540 1520 1270 2630 1260 80 2880 1510-10 20 20 -10 -10 -10 -10 -10

40 60 80 -10 -10 340 -10 2065 109 118 13 70 35 8 13

2.3 1.5 1.9 2 1.5 -0.1 2.2 2.25 5 7.5 -0.5 1.5 -0.5 0.5 0.5

7.5 7.5 10 1 6 -0.5 2 11.1 1 1.2 0.6 0.8 -0.1 0.6 0.45.6 4.4 5.7 3.4 4.3 -0.1 3.5 4.7

0.38 0.19 0.1 0.23 0.14 0.01 0.15 0.14

B37

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0018 265.60

LJD0048 156.00

LJD0048 159.40

LJD0048 159.70

LJD0048 162.30

LJD0048 164.05

LJD0048 164.30

LJD0048 171.30


14.3 2.74 5 5.21 2.19 4.86 5.64 5.111.4 7.85 4.42 6.69 6.63 7.63 7.2 8.18

1.06 9.58 14.4 13.3 8.61 12.5 13.6 12.92.87 0.02 -0.01 -0.01 -0.01 -0.01 -0.01 -0.010.38 7.38 6.98 5.12 6.94 4.78 4.97 4.320.77 27.7 28.1 25.1 28.6 24.3 24.1 23.3

4.83 0.18 0.08 0.09 0.11 0.08 0.06 0.080.091 0.049 0.021 0.042 0.044 0.038 0.044 0.03573.7 44.1 40 43.4 46.3 44.9 43.4 44.80.35 0.26 0.49 0.5 0.23 0.45 0.51 0.5199.7 99.8 99.4 99.4 99.6 99.5 99.5 99.2

1 1 1.5 2 1 2 1.5 20.5 0.5 1 1 0.5 1 1 10.4 0.2 0.4 0.4 0.4 0.4 0.4 0.4

-2 -2 -2 -2 -2 -2 -2 -22 0.3 0.3 0.3 0.1 0.3 0.4 1.6

0.2 0.2 0.4 0.4 0.2 0.4 0.4 0.47 1.5 1 1.5 1 1.5 1.5 2.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

1 0.4 0.7 0.6 0.3 0.6 0.6 3.66 2.5 3 3.5 2 3.5 3 4

1.4 0.6 0.6 0.6 0.4 0.6 0.6 0.820 -10 -10 -10 -10 -10 -10 -10

1 1 1 1.5 0.5 1 1 1.5SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

90 60 40 50 60 30 30 500.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 0.1-0.2 -0.2 0.2 0.4 -0.2 0.2 0.2 0.26.65 1.15 1.05 0.85 0.5 0.65 0.65 1

5.5 6.25 9 9.6 5.5 9.4 8.95 9.250.5 0.5 1 1 -0.5 1 1 1

0.02 0.01 0.01 -0.01 -0.01 -0.01 -0.01 -0.0121 1929 2928 2887 1669 2661 3003 2914-10 50 60 70 90 60 70 10030 1600 1360 1120 1570 1080 1190 89020 -10 -10 -10 -10 -10 -10 -10

90 60 40 50 60 30 30 5039 83 145 149 70 118 140 145

-0.1 2.9 2.5 2.3 2 1.9 2.1 20.5 4 8.5 8 3.5 6.5 9 7.51 6 11.5 11 5 10.5 11.5 14.5

-0.1 0.9 1.5 1.4 0.8 1.4 1.4 1.3-0.1 4.7 6.9 5.9 4.3 6.4 6.3 5.5

-0.01 1.44 0.48 0.46 0.23 0.18 0.21 0.79

B38

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0048 189.10

LJD0068 125.70

LJD0068 218.60

LJD0068 291.00

LJD0068 324.40

LJD0068 324.40 Rpt

LJD0069 100.30

LJD0069 173.00


5.35 0.96 1.06 1.13 3.8 3.82 0.44 0.68.09 0.91 0.18 0.19 28.9 28.9 -0.01 -0.0113.5 8.2 8.4 8.34 7.45 7.43 7.96 7.03-0.01 -0.01 -0.01 -0.01 0.28 0.29 -0.01 -0.014.35 8.9 6.7 6.48 19.1 19.1 12.5 11.623.1 43 40.7 42.7 11.5 11.4 41.7 42.3

0.09 0.18 0.03 0.06 0.37 0.37 0.04 0.130.038 0.011 0.01 0.007 0.031 0.031 0.009 0.00744.2 37.5 42.7 40.6 27.9 27.8 37.1 37.90.53 0.06 0.08 0.07 0.31 0.31 0.02 0.0399.2 99.7 99.8 99.5 99.6 99.4 99.7 99.5

2 -0.5 -0.5 -0.5 1 1 -0.5 -0.51 -0.5 -0.5 -0.5 0.5 0.5 -0.5 -0.5

0.4 -0.2 -0.2 -0.2 0.4 0.4 -0.2 -0.2

-2 -2 -2 -2 -2 -2 -2 -20.5 0.5 0.3 0.2 0.7 0.7 0.2 0.10.4 -0.2 -0.2 -0.2 0.2 0.2 -0.2 -0.21.5 1 -0.5 -0.5 3 3 -0.5 -0.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.8 1.2 0.7 0.5 0.8 0.8 0.2 0.23.5 1 -0.5 0.5 3 3 -0.5 -0.5

0.6 0.2 -0.2 -0.2 0.6 0.6 -0.2 -0.2-10 -10 -10 -10 -10 -10 -10 -10

1.5 -0.5 -0.5 -0.5 1 1 -0.5 -0.5SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

40 -10 -10 -10 130 120 -10 -100.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.10.4 -0.2 -0.2 -0.2 0.2 0.2 -0.2 -0.2

1.55 0.95 0.65 0.5 0.75 0.7 0.35 0.25

9.95 2.05 1.5 1.55 6.55 6.75 0.5 0.451 -0.5 -0.5 -0.5 0.5 0.5 -0.5 -0.5

-0.01 -0.01 -0.01 -0.01 0.02 0.01 -0.01 -0.013058 1382 1498 1573 1813 1806 1156 1169

60 -10 -10 -10 130 130 -10 -10960 3090 2810 2770 990 980 2700 2770-10 -10 -10 -10 -10 -10 -10 -10

40 -10 -10 -10 130 120 -10 -10149 13 21 13 79 79 -10 -10

2.4 4 3.8 3.5 1.3 1.3 2.5 3.19.5 1.5 1.5 1 4 4.5 1 0.5

12.5 2 2 2 5.5 5.5 0.5 11.5 0.6 0.7 0.7 0.9 0.9 0.3 0.46.6 8.1 5.3 5.3 4 4.3 5.3 4.7

0.22 0.15 0.07 0.08 0.01 0.01 0.05 0.04

B39

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0069 238.00

LJD0070 84.20

LJD0070 156.20

LJD0070 224.00

LJD0077 95.56

LJD0077 190.00

LJD0077 295.60

LJD0077 345.10


4.31 0.61 0.85 5.08 0.37 0.51 1.06 4.259.04 0.13 0.22 23.1 -0.01 0.04 0.27 12.19.9 8.94 9.15 10.8 7.36 6.91 14.8 10.5

0.05 -0.01 -0.01 0.92 -0.01 -0.01 0.52 0.054.19 8.27 7.88 19.7 16.7 15.3 4.51 5.9724 44.3 44.5 14.2 41.5 41 40.1 23.3

0.14 0.08 0.27 0.08 0.02 0.3 0.1 0.150.034 0.012 0.012 0.041 0.007 0.01 0.013 0.03947.7 37.4 36.8 24.6 33.6 35.5 38.4 42.60.36 0.03 0.06 0.41 0.02 0.03 0.08 0.3699.7 99.7 99.7 98.9 99.5 99.5 99.8 99.3

1.5 -0.5 -0.5 2 -0.5 -0.5 -0.5 1.51 -0.5 -0.5 1 -0.5 -0.5 -0.5 1

0.4 -0.2 -0.2 0.4 -0.2 -0.2 -0.2 0.4

-2 -2 -2 -2 -2 -2 -2 -20.6 0.2 -0.1 0.9 -0.1 0.5 0.1 0.50.4 -0.2 -0.2 0.4 -0.2 -0.2 -0.2 0.43.5 -0.5 0.5 4 -0.5 2 1.5 3

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.7 0.3 0.2 0.7 0.1 1 0.3 0.64.5 -0.5 0.5 4.5 -0.5 3.5 1.5 3.5

1 -0.2 -0.2 1 -0.2 0.8 0.4 0.8-10 -10 -10 50 -10 -10 30 -10

1 -0.5 -0.5 1.5 -0.5 0.5 -0.5 1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

30 -10 -10 130 -10 -10 -10 30-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.10.2 -0.2 -0.2 0.4 -0.2 -0.2 -0.2 0.2

0.55 0.3 0.25 0.55 0.25 0.65 0.35 0.5

8.15 0.75 1.15 10.5 0.3 0.85 2 8.251 -0.5 -0.5 1 -0.5 -0.5 -0.5 1

0.01 -0.01 -0.01 0.02 -0.01 -0.01 0.01 0.012209 1464 1388 2476 862 1224 896 2127

80 -10 -10 250 -10 -10 140 501680 3070 3170 1380 3110 2680 3430 1230-10 -10 -10 50 -10 -10 30 -10

30 -10 -10 130 -10 -10 -10 30101 8 13 109 -10 -10 30 96

3.6 3.4 3 1.9 2.3 4.8 3.3 1.514 1.5 1.5 5.5 1.5 1.5 17 4.5

13.5 0.5 1 8.5 5 5 12.5 72.6 0.6 0.6 1.1 0.6 1.4 1.6 0.99.7 5.5 4.8 5.6 4.6 9.8 6.7 4.3

0.45 0.11 0.2 0.16 0.3 0.04 0.84 0.18

B40

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0077 345.10 Rpt

LJD0086 106.70

LJD0086 162.30

LJD0088 156.70

LJD0088 198.50

LJD0088 237.00

LJD0120 229.60

LJD0120 284.00


4.24 0.54 4.15 0.73 1.4 1.69 5.22 0.9512.1 0.35 9.31 0.09 0.07 2.53 8.82 0.5810.6 8.53 8.53 14.1 8.33 9.72 9.35 8.390.05 -0.01 0.06 -0.01 0.04 -0.01 -0.01 -0.016.02 8.23 7.85 13.2 12.6 11.3 4.87 9.9423.2 43.5 24.8 36 38.2 34.5 24.5 40.4

0.15 0.14 0.34 0.32 0.16 0.19 0.11 0.040.039 0.007 0.033 0.021 0.005 0.017 0.039 0.01142.5 38.5 44.6 35.4 38.7 39.4 46.3 39.20.36 0.04 0.34 0.06 0.08 0.13 0.42 0.0699.2 99.8 100 99.9 99.5 99.4 99.6 99.5

1.5 -0.5 1 -0.5 -0.5 1 1.5 -0.51 -0.5 0.5 -0.5 -0.5 0.5 1 -0.5

0.4 -0.2 0.2 -0.2 -0.2 -0.2 0.2 -0.2

-2 -2 -2 -2 -2 -2 -2 -20.5 0.1 0.4 -0.1 -0.1 0.1 0.5 0.10.4 -0.2 0.2 -0.2 -0.2 -0.2 0.4 -0.23 -0.5 1 1 0.5 2 1.5 -0.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.7 0.1 0.5 0.1 0.2 0.3 0.7 0.23.5 -0.5 2.5 1 1 5.5 3 -0.5

0.8 -0.2 0.4 0.2 0.2 1.2 0.6 -0.2-10 -10 -10 -10 -10 -10 -10 -10

1 -0.5 1 -0.5 -0.5 1.5 1 -0.5SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

40 -10 20 -10 -10 -10 -10 -10-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.10.2 -0.2 0.2 -0.2 -0.2 -0.2 0.2 -0.2

0.45 0.2 0.35 0.3 0.3 0.5 0.55 0.2

8.4 1.05 6.6 1.25 2.15 5.05 7.8 1.150.5 -0.5 0.5 -0.5 -0.5 -0.5 1 -0.5

0.01 0.01 0.01 -0.01 -0.01 -0.01 -0.01 -0.012134 1443 1888 745 1108 2593 1826 1361

50 -10 30 330 -10 80 20 -101230 2880 1200 2090 2050 2270 530 2940-10 -10 -10 -10 -10 -10 -10 -10

40 -10 20 -10 -10 -10 -10 -1096 8 83 17 13 39 105 13

1.6 3.5 1.4 1.9 3.5 3 0.4 1.34.5 1 4.5 8.5 -0.5 2.5 -0.5 76.5 2 7.5 7 1 4.5 1.5 80.9 0.9 0.9 1 0.7 1.1 -0.1 1.14.3 6.6 3.9 3.7 10.1 5.8 0.7 5.2

0.18 0.09 0.03 0.19 0.14 0.1 0.01 0.16

B41

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJD0120 344.20

LJD0124 130.00

LJD0124 164.21

LJD0126 81.50

LJD0126 119.20

LJD0126 157.50

LJD0126 189.50

LJD0126 313.10


1.13 0.61 3.4 4.31 3.87 4.46 4.64 3.910.4 0.12 14.1 6.18 7.66 6.54 5.85 6.2

9.29 9.4 8.65 11 10.4 13.1 11.7 12.6-0.01 -0.01 -0.01 0.03 0.06 0.07 0.03 0.0212.4 8.39 12.6 5.98 4.65 4.54 4.92 4.0438.3 43.7 22.1 26.4 24 25.4 27.9 27.4

0.11 0.11 0.06 0.17 0.19 0.17 0.14 0.160.009 0.008 0.031 0.034 0.012 0.03 0.048 0.03537.8 37.4 38.1 45 48.3 44.4 43.8 44.40.04 0.05 0.27 0.37 0.4 0.51 0.38 0.4599.4 99.7 99.3 99.4 99.5 99.2 99.4 99.2

-0.5 -0.5 1 1.5 1 1.5 1.5 1.5-0.5 -0.5 0.5 1 0.5 1 1 1-0.2 -0.2 -0.2 0.4 0.4 0.4 0.4 0.4

-2 -2 -2 -2 -2 -2 -2 -20.1 -0.1 0.5 0.3 0.3 0.4 0.2 0.2-0.2 -0.2 0.2 0.4 0.2 0.4 0.4 0.41.5 -0.5 1.5 1 1 1 1.5 1

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.2 0.1 0.5 0.5 0.6 0.6 0.6 0.70.5 -0.5 2 3 2 3 2.5 3

0.2 -0.2 0.4 0.6 0.4 0.6 0.6 0.6-10 -10 -10 -10 -10 -10 -10 -10

-0.5 -0.5 1 1 1 1 1 1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 -10 20 30 20 30 40 30-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1-0.2 -0.2 -0.2 0.2 -0.2 0.2 0.2 0.20.25 0.2 0.3 0.25 0.25 0.3 0.25 0.25

1.15 0.95 6.4 7.8 6.3 7.55 7.9 7.8-0.5 -0.5 0.5 0.5 -0.5 0.5 0.5 0.5

-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.011457 1464 1765 2251 2052 2969 2244 2599

30 -10 160 60 70 20 40 703560 2950 1190 1390 1160 1080 1510 1590-10 -10 -10 -10 -10 -10 -10 -10

-10 -10 20 30 20 30 40 304 8 70 96 92 136 101 109

1.2 3.3 3.8 1.6 1.5 2.1 1.9 2.34 1.5 11.5 -0.5 7 13 6 85 1 13 0.5 10 14 8.5 11

0.8 0.4 1.8 0.3 1.5 2 1.1 1.43.7 4.8 9.6 2.9 5.9 6.4 5.1 6.5

0.44 0.05 0.07 0.05 0.04 0.03 0.04 0.06

B42

Maggie Hays

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


LJPD0094 101.70

LJPD0094 265.40

Ultratrace Ultratraceu106349 u106349

0.53 7.520.02 7.138.84 12.1-0.01 0.028.9 4.3

42.3 21.9

0.07 0.140.01 0.03738.8 45.90.04 0.4699.5 99.5

-0.5 1.5-0.5 1-0.2 0.2

-2 -20.1 0.4-0.2 0.4-0.5 1.5

-0.2 -0.2

0.1 0.9-0.5 3

-0.2 0.6-10 -10

-0.5 1SnSrTaTbThTiTl

TmUVWY


PdPtRhRu

CO2S

-10 -10-0.1 -0.1-0.2 0.20.15 0.3

0.25 9.9-0.5 1

-0.01 -0.011375 2654-10 20

2930 990-10 -10

-10 -104 136

3.7 1.6-0.5 111.5 140.7 1.85.4 5.6

0.08 0.01

B43

Karelian Craton

Karelian Craton

Notes: XRF = X-ray florescence, ICP-MS = Inductively coupled plasma mass spectrometry, FA-ICP-MS = Fire assay inductively coupled plasma mass spectrometry, OES = optical emission spectrometry, D.L. = analytical reported detection limit, N.D. = not determined, wt% = weight percent, ppm = parts per million, ppb = parts per billion.

B44

Karelian Craton

Sample WP 44 WP 45 WP 46 WP 47 WP 48 WP 49Lab Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace

Units D.L. Batch u116494 u116494 u116494 u116494 u116494 u116494wt% 0.01 Al2O3 8.88 2.26 6.53 5.65 0.91 0.67wt% 0.01 CaO 5.98 5.30 8.05 10.30 3.42 0.30wt% 0.01 Fe2O3 13.40 6.04 12.20 10.00 5.65 6.91wt% 0.01 K2O 0.02 -0.01 0.03 0.03 -0.01 -0.01wt% LOI 6.17 7.22 4.27 3.32 8.27 12.30wt% 0.01 MgO 23.90 31.60 22.80 20.30 33.00 37.50wt% 0.01 MnOwt% 0.01 Na2O 0.18 0.07 0.32 0.37 0.06 0.04wt% 0.001 P2O5 0.05 0.01 0.03 0.04 0.00 0.00wt% 0.01 SiO2 40.00 46.50 44.60 49.20 47.00 41.10wt% 0.01 TiO2 0.73 0.07 0.65 0.58 0.05 0.03wt% Total 99.30 99.00 99.40 99.70 98.30 98.80ppm Bappm Beppm Bippm Cdppm 0.05 Ce 2 2 1.5 3 1.5 -0.5ppm Coppm Crppm Csppm Cuppm 0.05 Dy 1.5 0.5 2 2.5 -0.5 -0.5ppm 0.05 Er 1 -0.5 1 1.5 -0.5 -0.5ppm 0.05 Eu 0.4 0.2 0.4 0.4 -0.2 -0.2ppm 0.2 Gappm 0.2 Gd -2 -2 2 2 -2 -2ppm 0.1 Hf 0.3 -0.1 0.4 0.9 -0.1 -0.1ppm 0.02 Ho 0.4 -0.2 0.4 0.6 -0.2 -0.2ppm 0.05 La 0.5 1 -0.5 1 -0.5 -0.5ppm 0.5 Lippm 0.02 Lu -0.2 -0.2 -0.2 -0.2 -0.2 -0.2ppm 0.2 Moppm 0.5 Nb 1.6 0.5 0.4 0.5 0.3 0.1ppm 0.5 Nd 2.5 1.5 2.5 4 1 -0.5ppm Nippm Pbppm 0.02 Pr 0.4 0.4 0.4 0.6 0.2 -0.2ppm 0.02 Rbppm Sbppm Scppm 0.05 Sm 1 -0.5 1.5 1.5 -0.5 -0.5ppm Snppm 0.1 Sr

XRF

ICP-

MS

ppm 0.1 Srppm 0.05 Ta -0.1 -0.1 -0.1 -0.1 -0.1 -0.1ppm 0.02 Tb 0.2 -0.2 0.4 0.4 -0.2 -0.2ppm 0.05 Th 0.2 0.05 0.05 0.15 -0.05 -0.05ppm Tippm Tlppm Tm -0.2 -0.2 -0.2 -0.2 -0.2 -0.2ppm 0.05 Uppm Vppm Wppm 0.1 Y 7.4 3.4 9.5 12.1 2.65 0.6ppm 0.05 Yb 0.5 -0.5 1 1 -0.5 -0.5ppm Znppm 1 Zrppm Asppm 20 Ba -0.01 -0.01 -0.01 -0.01 -0.01 -0.01ppm 7 Cr 2764 2558 2121 1559 2381 2237ppm 8 Cu 260 40 50 20 30 20ppm 8 Ni 1420 1160 1480 390 2550 2040ppm 10 Rb -10 -10 20 -10 -10 -10ppm Scppm 10 Sr 30 30 50 30 -10 -10ppm 40 V 175 35 140 184 21 13ppm Yppm 25 Zr - - - - - -ppb Auppb 0.2 Ir 2.40 2.50 2.80 1.60 4.10 3.80ppb 0.3 Pd 6.00 3.00 11.00 8.00 2.00 3.00ppb 0.3 Pt 15.00 27.50 44.00 34.00 9.50 13.50ppb 0.1 Rh 1.30 1.60 1.60 1.10 1.20 1.10ppb 0.2 Ru 5.30 7.70 5.10 3.50 8.70 7.20Wt% CO2wt% S 0.03 0.33 0.01 0.17 0.64 0.44

FA-IC

P-M

SXR

F

B45

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI

MgOMnONa2OP2O5SiO2TiO2TotalBaBeBiCdCeCoCrCsCuDyErEuGaGdHfHoLaLiLuMoNbNdNiPbPrRbSbScSmSnSr

WP 50 WP 51 WP 51 Rpt WP 52 WP 53 WP 53 Rpt WP 54 WP 55Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

1.42 4.68 4.70 5.49 3.69 3.68 6.01 4.252.58 5.70 5.71 8.96 3.21 3.21 9.41 4.297.31 5.85 5.84 11.00 10.00 10.10 11.90 6.53-0.01 -0.01 -0.01 0.01 -0.01 -0.01 0.07 0.0210.80 6.39 6.38 4.05 9.47 9.48 1.86 8.6334.60 31.30 31.30 21.50 32.20 32.20 21.90 31.60

0.05 0.08 0.08 0.21 0.13 0.14 0.28 0.210.01 0.07 0.07 0.04 0.02 0.02 0.04 0.0141.70 45.00 45.00 48.30 40.40 40.30 47.60 43.300.03 0.43 0.43 0.61 0.19 0.19 0.72 0.0898.40 99.40 99.40 100.00 99.30 99.30 99.70 98.90

1.5 4 4 2 6 6 2.5 4

-0.5 1 1 2.5 1 1 2 1-0.5 -0.5 -0.5 1 0.5 0.5 1 -0.5-0.2 0.2 0.2 0.2 0.2 0.2 0.6 0.2

-2 -2 -2 2 -2 -2 2 -2-0.1 0.8 0.7 0.5 0.4 0.4 0.6 0.1-0.2 -0.2 -0.2 0.4 -0.2 -0.2 0.4 -0.20.5 1.5 2 0.5 2.5 2.5 0.5 1.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.3 1.6 1.6 0.4 0.8 0.8 0.4 0.51 3 3 2.5 3 3 3.5 2.5

0.2 0.6 0.6 0.4 0.8 0.8 0.6 0.6

-0.5 1 1 1.5 1 0.5 1.5 0.5

SrTaTbThTiTl

TmUVWYYbZnZrAsBaCrCuNiRbScSrVYZrAuIr

PdPtRhRu

CO2S

-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1-0.2 -0.2 -0.2 0.4 -0.2 -0.2 0.4 -0.20.05 0.25 0.25 0.05 0.35 0.4 0.05 0.1

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

1.7 4 4 10.5 5.05 4.8 9.1 4.45-0.5 -0.5 -0.5 1 -0.5 0.5 1 0.5

-0.01 -0.01 0.01 -0.01 -0.01 -0.01 -0.01 -0.012723 1785 1778 1703 3913 3920 1881 4529

60 20 20 40 -10 -10 60 502370 1220 1200 580 2030 2010 1310 2040-10 -10 -10 20 -10 -10 20 -10

-10 -10 -10 30 -10 -10 70 -1021 105 105 136 57 61 167 57

- - - 74 - - - -

3.80 2.10 2.60 2.10 1.60 2.10 2.90 1.202.50 7.50 11.00 9.00 7.00 7.00 7.00 2.504.00 8.00 14.50 11.50 8.50 9.00 9.50 5.501.20 1.30 3.30 0.90 1.20 1.30 1.10 2.908.70 4.20 10.90 2.90 9.30 9.00 4.30 3.90

0.85 0.22 0.23 0.11 0.13 0.14 0.08 0.58

B46

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP 56 WP 57 WP 58 WP 59 WP 60 WP 61 WP 62 WP 63Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

1.72 22.10 15.50 1.97 1.25 0.96 0.51 2.7913.20 2.74 13.00 0.40 0.05 0.02 0.50 1.224.73 11.10 6.83 11.20 11.40 9.64 7.98 10.800.02 0.03 0.26 -0.01 0.01 -0.01 -0.01 -0.012.17 9.04 1.18 15.30 12.10 12.40 8.36 10.6022.10 27.60 12.90 32.80 35.20 37.70 43.60 31.80

0.24 0.11 1.99 0.04 0.04 0.03 0.05 0.020.00 0.59 0.02 0.04 0.01 0.01 0.01 0.0155.60 25.00 47.60 36.20 38.70 37.90 37.10 40.800.03 0.88 0.32 0.07 0.04 0.03 0.01 0.1499.80 99.10 99.50 98.00 98.70 98.60 98.10 98.10

0.5 2.5 2 -0.5 -0.5 -0.5 -0.5 0.5

-0.5 1.5 1.5 -0.5 -0.5 -0.5 -0.5 -0.5-0.5 1 1 -0.5 -0.5 -0.5 -0.5 -0.50.2 -0.2 0.4 -0.2 -0.2 -0.2 -0.2 -0.2

-2 -2 -2 -2 -2 -2 -2 -2-0.1 1.2 0.5 -0.1 -0.1 -0.1 -0.1 -0.1-0.2 0.4 0.4 -0.2 -0.2 -0.2 -0.2 -0.2-0.5 1.5 1 -0.5 -0.5 -0.5 -0.5 -0.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

-0.1 3.4 0.7 0.4 0.3 0.1 0.1 0.70.5 2 2 -0.5 -0.5 -0.5 -0.5 -0.5

-0.2 0.4 0.4 -0.2 -0.2 -0.2 -0.2 -0.2

-0.5 0.5 0.5 -0.5 -0.5 -0.5 -0.5 -0.5

SrTaTbThTiTl


PdPtRhRu

CO2S

-0.1 0.2 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1-0.2 0.2 0.2 -0.2 -0.2 -0.2 -0.2 -0.2

-0.05 0.1 0.1 0.1 0.1 0.05 -0.05 0.05

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

1.7 7.1 8.15 1.6 0.95 0.4 0.45 1.4-0.5 0.5 1 -0.5 -0.5 -0.5 -0.5 -0.5

0.01 -0.01 0.02 -0.01 -0.01 -0.01 -0.01 -0.011949 3161 1888 2497 3243 6445 9907 2223

20 -10 20 240 90 -10 -10 201430 570 310 8410 2530 3260 3250 1800-10 -10 30 -10 -10 -10 -10 -10

40 -10 160 -10 -10 -10 -10 -1039 74 136 17 8.7 13 13 39

- - - - - - - -

2.80 0.80 0.30 3.50 2.40 1.50 4.20 0.601.00 2.00 53.00 122.00 37.50 17.50 12.50 8.002.00 40.50 23.50 42.00 13.50 6.50 11.50 8.500.70 3.00 1.30 12.30 4.10 1.70 5.60 1.405.70 2.00 0.80 33.80 14.60 5.40 25.20 5.60

0.15 0.1 0.02 1.91 0.3 0.08 0.1 0.45

B47

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP 64 WP 65 WP 66 WP 67 WP 68 WP 69 WP 70 WP 71Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

7.55 8.21 6.34 12.70 7.48 13.20 8.96 15.207.38 7.55 8.63 16.10 6.84 6.30 7.26 11.4010.30 12.40 11.00 17.90 12.20 15.40 9.86 11.000.03 0.05 0.06 0.19 0.02 0.19 0.02 0.334.31 3.34 1.89 1.02 4.51 5.52 5.89 1.2922.90 21.70 20.70 6.49 21.80 19.00 23.20 9.79

0.24 0.37 0.43 0.47 0.17 0.37 0.20 1.710.01 0.02 0.01 0.08 0.02 0.02 0.02 0.0446.50 45.30 49.80 43.70 46.30 39.30 44.00 48.500.26 0.29 0.23 0.84 0.24 0.39 0.32 0.5299.40 99.20 99.00 99.40 99.50 99.60 99.70 99.70

2.5 3 3 7.5 1.5 9 1.5 4.5

1 1 1 3.5 1 2 1 20.5 0.5 0.5 2 0.5 1.5 0.5 1.5-0.2 0.2 0.2 0.6 -0.2 0.6 -0.2 0.4

-2 -2 -2 2 -2 -2 -2 -20.4 0.5 0.4 1.1 0.3 0.3 0.3 0.50.2 0.2 0.2 0.8 0.2 0.4 0.2 0.41 1 1 3 -0.5 5 0.5 2

-0.2 -0.2 -0.2 0.4 -0.2 -0.2 -0.2 -0.2

0.7 0.8 0.5 2.6 0.7 1 0.8 1.31.5 2 2 6.5 1 6 1.5 3.5

0.4 0.4 0.4 1.2 0.2 1.4 0.2 0.8

0.5 0.5 0.5 2 -0.5 1.5 0.5 1

SrTaTbThTiTl


PdPtRhRu

CO2S

-0.1 -0.1 -0.1 0.1 -0.1 -0.1 -0.1 -0.1-0.2 -0.2 -0.2 0.4 -0.2 0.4 -0.2 0.20.15 0.2 0.2 0.25 0.15 0.15 0.1 0.15

-0.2 -0.2 -0.2 0.4 -0.2 -0.2 -0.2 -0.2

6 6.1 5.95 18.4 5.65 11.5 5.75 11.70.5 0.5 0.5 2 0.5 1.5 0.5 1

-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 0.013270 3407 2901 424 1963 800 2292 458

20 40 40 30 -10 20 20 150920 1000 790 200 790 360 470 200-10 -10 30 40 20 40 -10 30

30 40 50 130 -10 -10 -10 130127 127 109 202 105 109 96 153

- - - - - - - 74

0.60 1.10 0.60 0.20 0.50 0.50 0.70 0.403.50 13.00 9.00 1.00 5.50 4.50 11.50 7.508.00 9.00 8.00 2.00 7.00 4.50 4.50 7.501.30 1.40 1.20 0.20 1.20 0.70 0.70 0.905.00 4.50 4.40 0.50 3.10 2.10 4.40 2.10

-0.01 0.06 0.08 0.01 -0.01 -0.01 0.14 0.12

B48

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP 72 WP 73 WP 74 WP 75 WP 76 WP 77 WP 78 WP 78 RptUltratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

1.21 14.60 0.65 16.20 8.28 9.18 8.15 8.160.39 3.14 2.16 2.67 10.20 9.08 10.00 10.009.34 8.90 10.60 3.48 10.10 11.70 11.30 11.30-0.01 -0.01 -0.01 0.29 0.23 0.09 0.10 0.0913.80 9.22 12.70 0.99 3.28 5.07 5.70 5.7836.10 27.20 34.50 3.35 16.90 19.50 20.50 20.50

0.05 0.05 0.04 8.08 1.54 0.95 0.81 0.790.01 0.07 0.01 0.12 0.02 0.04 0.05 0.0536.80 35.80 37.00 64.20 48.40 43.10 42.20 42.200.02 0.64 0.01 0.33 0.47 0.68 0.62 0.6197.70 99.60 97.60 99.70 99.40 99.30 99.40 99.40

-0.5 2 1 14 3.5 3.5 4 4.5

-0.5 0.5 -0.5 1.5 2 2 2 2-0.5 -0.5 -0.5 1 1 1 1.5 1.5-0.2 -0.2 -0.2 0.8 0.6 0.6 0.6 0.6

-2 -2 -2 2 -2 -2 -2 2-0.1 3.6 -0.1 2.3 0.6 0.5 0.6 0.5-0.2 -0.2 -0.2 0.4 0.4 0.4 0.4 0.4-0.5 1 -0.5 6.5 1.5 1 1.5 1.5

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.3 8.1 1.3 5.7 1.7 1.7 1.6 1.6-0.5 1.5 0.5 6.5 2.5 3 3.5 3.5

-0.2 0.4 -0.2 1.6 0.6 0.6 0.6 0.6

-0.5 -0.5 -0.5 2 1 1 1.5 1.5

SrTaTbThTiTl


PdPtRhRu

CO2S

0.1 0.4 -0.1 0.4 -0.1 -0.1 -0.1 -0.1-0.2 -0.2 -0.2 0.2 0.2 0.2 0.4 0.40.05 0.7 0.1 6.4 0.3 0.2 0.15 0.1

-0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

0.55 2.95 1.7 9.6 10.6 10 11.6 12.4-0.5 -0.5 -0.5 1 1 1 1 1

-0.01 -0.01 -0.01 0.04 0.01 -0.01 -0.01 -0.011253 342 1201 314 1628 1997 1991 1984

20 -10 50 50 180 80 40 401860 720 2510 110 1060 1130 1090 1090-10 -10 -10 -10 30 -10 20 -10

-10 -10 -10 380 100 60 80 7017 92 21 43 131 149 153 149

- 74 - 148 74 74 - 74

7.20 0.20 7.70 0.20 0.80 1.90 1.20 1.306.00 0.50 0.50 1.50 2.50 6.00 8.50 9.008.00 1.00 1.50 1.00 6.50 16.00 9.00 9.001.60 0.10 1.20 -0.10 0.70 2.70 0.90 0.9016.10 0.70 14.80 0.50 2.30 7.60 3.10 2.90

0.06 -0.01 0.13 -0.01 -0.01 -0.01 -0.01 -0.01

B49

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP 79 WP 80 WP 81 WP 82 WP 82 Rpt WP 83 WP 84 WP 85Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

7.91 9.75 9.45 5.18 5.20 10.50 7.00 13.809.13 7.80 7.89 3.91 3.91 11.30 9.85 10.9011.00 12.10 11.40 9.44 9.46 11.80 9.44 11.100.12 0.12 0.04 -0.01 -0.01 0.08 0.05 0.204.29 4.65 4.54 10.70 10.60 1.37 3.08 2.2821.10 21.40 21.00 26.90 26.90 16.10 20.40 7.10

0.85 0.90 0.66 0.04 0.04 1.74 0.74 4.190.04 0.05 0.04 0.00 0.00 0.01 0.01 0.0944.80 42.30 44.10 42.80 42.90 45.90 48.50 48.000.50 0.63 0.53 0.27 0.27 0.62 0.45 1.3399.70 99.70 99.60 99.20 99.20 99.40 99.50 98.90

4 3.5 3 1.5 1.5 4.5 3.5 11.5

2 2 2 0.5 0.5 2.5 1.5 4.51 1 1.5 -0.5 -0.5 1.5 1 3

0.6 0.4 0.4 -0.2 -0.2 0.6 0.4 1

-2 -2 -2 -2 -2 2 -2 40.6 0.6 0.4 -0.1 -0.1 0.7 0.4 0.70.4 0.4 0.4 -0.2 -0.2 0.6 0.4 11.5 1 1 0.5 0.5 1.5 1 3.5

-0.2 -0.2 -0.2 -0.2 -0.2 0.2 -0.2 0.4

1.3 1.4 1.3 1 1 1.7 1.3 6.53 3 2.5 1 1 4 3 8.5

0.6 0.6 0.4 0.2 0.2 0.8 0.6 1.6

1 1 1 -0.5 -0.5 1.5 1 3

SrTaTbThTiTl


PdPtRhRu

CO2S

-0.1 -0.1 -0.1 -0.1 -0.1 -0.1 -0.1 0.20.4 0.2 0.2 -0.2 -0.2 0.4 0.2 0.60.1 0.1 0.05 0.05 0.05 0.1 0.1 0.3

-0.2 -0.2 -0.2 -0.2 -0.2 0.2 -0.2 0.4

11.8 10.3 11.1 3.7 3.9 12.8 9.35 23.91 1 1 -0.5 -0.5 1.5 1 2.5

-0.01 -0.01 -0.01 -0.01 -0.01 0.01 -0.01 0.011936 2032 1744 3550 3571 1731 1997 171

20 30 20 60 60 20 20 97501090 1140 990 1520 1530 800 950 90

20 -10 20 -10 -10 -10 20 40

80 80 60 50 80 60 80 250145 153 145 57 61 184 114 272

74 74 - - - 74 - 74

1.10 1.30 0.70 3.00 3.10 0.80 1.10 0.202.50 6.00 3.50 21.50 22.00 2.50 2.00 1.007.50 11.00 9.50 17.00 16.00 9.50 6.50 1.500.70 1.10 1.20 1.40 1.60 1.20 0.70 0.202.80 4.20 3.60 6.30 6.40 3.60 3.70 1.60

-0.01 -0.01 -0.01 0.04 0.04 -0.01 -0.01 0.23

B50

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP 85 Rpt WP 86 WP 87 WP 88 WP 89 WP 90 WP 91 WP 92Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

13.70 13.40 6.59 11.90 13.60 16.90 8.40 7.4310.90 11.50 8.61 6.75 17.60 16.20 8.29 8.3511.10 12.90 9.32 12.60 12.10 12.40 12.70 12.900.21 0.25 0.02 0.04 0.08 0.07 0.04 0.052.23 0.79 4.84 5.53 3.19 2.27 4.80 3.987.13 13.50 22.10 20.80 8.77 8.44 20.10 20.80

4.21 2.00 0.29 0.55 0.81 0.73 0.38 0.400.09 0.06 0.03 0.05 0.04 0.09 0.08 0.0848.00 44.00 47.10 40.50 42.90 41.60 43.40 44.401.33 1.19 0.52 0.72 0.79 1.22 1.26 1.0798.90 99.50 99.40 99.40 99.80 99.90 99.40 99.40

12 8.5 3.5 4.5 5.5 10 16.5 12.5

4.5 3.5 1.5 2.5 2.5 4 2 23 2 1 1.5 1.5 2 1 11 1 0.2 0.6 0.8 1 0.8 0.6

4 4 -2 2 2 4 2 20.8 0.9 0.2 0.5 0.6 1 0.9 0.81 0.8 0.4 0.6 0.6 0.8 0.4 0.44 3 1 1.5 2 3.5 6.5 5

0.4 0.2 -0.2 -0.2 -0.2 0.2 -0.2 -0.2

6.5 3.8 1.4 1.4 2.2 4 5.1 4.49 7.5 3 4 4.5 7.5 11 8.5

2 1.4 0.6 0.8 0.8 1.6 2.4 1.8

3 2 1 1.5 1.5 2.5 2.5 2.5

SrTaTbThTiTl


PdPtRhRu

CO2S

0.2 0.2 -0.1 -0.1 0.1 0.2 0.3 0.30.6 0.6 0.2 0.4 0.4 0.6 0.4 0.40.3 0.3 0.1 0.15 0.15 0.3 0.5 0.45

0.4 0.4 -0.2 -0.2 0.2 0.4 -0.2 -0.2

24.9 18.1 8.05 12.2 13.7 19.7 10.1 10.32.5 2 1 1.5 1.5 2 1 1

0.01 0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01171 1115 2497 1970 75.2 218 1696 14649750 110 60 20 30 60 300 70100 350 1050 1060 140 130 840 105040 20 -10 20 30 50 20 -10

260 130 70 80 350 610 80 90281 250 118 162 219 259 180 158

74 74 - 74 - - - 74

0.40 0.70 1.40 0.90 -0.10 -0.10 1.80 1.401.00 3.50 4.00 1.50 -0.50 1.00 4.00 3.501.50 6.00 6.50 8.00 0.50 0.50 4.50 4.000.20 0.60 0.60 0.70 -0.10 -0.10 0.70 0.501.30 2.50 2.80 3.40 0.50 0.20 3.80 3.00

0.24 -0.01 -0.01 -0.01 -0.01 -0.01 0.02 0.02

B51

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP 93 WP 94 WP 95 WP 96 WP 97 WP 98 WP 99 WP 99 RptUltratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu116494 u116494 u116494 u116494 u116494 u116494 u116494 u116494

10.30 7.01 0.55 0.63 2.39 0.47 1.18 1.178.96 8.95 2.61 3.46 0.26 1.96 1.13 1.1316.40 12.10 7.09 9.39 9.05 8.34 9.55 9.510.10 0.04 -0.01 -0.01 -0.01 -0.01 -0.01 -0.015.26 5.76 13.40 14.70 12.00 13.20 12.30 12.3017.00 20.10 35.50 35.30 37.40 36.40 36.30 36.30

0.64 0.38 0.03 0.02 0.02 0.02 0.02 0.010.43 0.11 0.01 0.01 0.03 0.00 0.01 0.0136.90 43.70 40.30 35.50 37.90 38.90 38.80 38.803.42 1.36 0.08 0.06 0.18 0.05 0.09 0.0899.40 99.50 99.50 99.00 99.20 99.30 99.30 99.30

50.5 12.5 4.5 1 6.5 1.5 1.5 1.5

6 2.5 -0.5 -0.5 1 -0.5 -0.5 -0.53 1.5 -0.5 -0.5 -0.5 -0.5 -0.5 -0.5

1.4 0.6 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

8 2 -2 -2 -2 -2 -2 -21.9 1.2 -0.1 -0.1 0.4 -0.1 -0.1 -0.11.2 0.4 -0.2 -0.2 0.2 -0.2 -0.2 -0.219 5 2 0.5 2.5 0.5 0.5 0.5

0.4 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

22.8 7.4 1.7 0.5 2 0.5 0.7 0.529.5 8.5 1.5 0.5 4.5 1 1 1

7.4 1.8 0.4 -0.2 1 0.2 0.2 0.2

7 2.5 -0.5 -0.5 1.5 -0.5 -0.5 -0.5

SrTaTbThTiTl


PdPtRhRu

CO2S

0.7 0.6 0.1 -0.1 0.1 -0.1 -0.1 -0.11 0.4 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

2.3 0.8 0.05 0.05 0.25 -0.05 0.1 0.1

0.4 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2 -0.2

27.1 11.7 1.25 0.95 5.15 1.05 1.65 1.52.5 1 -0.5 -0.5 -0.5 -0.5 -0.5 -0.5

-0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.01 -0.0134.2 1785 1074 1149 1354 937 2045 202520 50 -10 20 -10 20 20 20

130 900 2100 2700 2410 3150 2830 280020 20 -10 -10 -10 -10 -10 -10

190 80 -10 -10 -10 -10 -10 -10268 175 4.3 4.3 17 -4 17 13

222 74 - - - - - -

0.10 1.70 0.90 1.30 1.40 1.40 1.90 1.80-0.50 8.00 1.50 1.00 8.00 2.00 4.50 6.001.00 9.50 2.50 1.50 7.50 2.00 4.00 6.00-0.10 1.10 0.40 0.30 0.60 0.30 0.50 0.600.50 4.10 3.40 3.60 3.70 3.10 5.50 5.40

-0.01 0.01 0.03 0.12 0.1 0.12 0.07 0.08

B52

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


Sample WP-44 WP-45 WP-46 WP-47 WP-48 WP-49Lab Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace

Units D.L Batch u157155 u157155 u157155 u157155 u157155 u157155wt% 0.002 Al2O3 8.90 2.10 5.90 5.50 0.90 0.60wt% 0.001 CaO 5.90 5.00 7.80 10.00 3.30 0.30wt% 0.001 Fe2O3 6.56 2.91 2.91 5.83 4.76 2.81wt% K2Owt% LOIwt% 0.002 MgO 24.20 31.50 22.70 20.00 33.60 38.60wt% MnOwt% Na2Owt% P2O5wt% SiO2wt% 0.001 TiO2 0.74 0.06 0.65 0.57 0.05 0.03wt% Totalppm Bappm Beppm Bippm Cdppm Ceppm Coppm 5 Cr 1580 1190 990 730 1230 1310ppm Csppm Cuppm Dyppm Erppm Euppm Gappm Gdppm Hfppm Hoppm Lappm Lippm Luppm Moppm Nbppm Ndppm 1 Ni 1400 1010 1390 288 2550 2050ppm Pbppm Prppm Rbppm Sbppm Scppm Smppm Snppm Sr

OES

OES

SrTaTbThTiTl


PdPtRhRu

CO2S

ppm Srppm Tappm Tbppm Thppm Tippm Tlppm Tmppm Uppm Vppm Wppm Yppm Ybppm Znppm Zrppm Asppm Bappm Crppm Cuppm Nippm Rbppm Scppm Srppm Vppm Yppm Zr

B53

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP-50 WP-51 WP-52 WP-53 WP-54 WP-55 WP-56 WP-57Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratrace Ultratraceu157155 u157155 u157155 u157155 u157155 u157155 u157155 u157155

1.30 4.60 5.50 - 5.80 4.20 1.70 18.002.40 5.50 8.90 - 9.10 4.20 12.00 2.703.35 3.51 2.87 - 5.55 3.23 2.29 5.04

34.90 31.60 21.50 - 21.00 32.10 21.80 26.60

0.03 0.44 0.61 - 0.71 0.08 0.02 0.87

1600 850 990 - 920 2170 960 1220

2340 1160 560 - 1230 2050 1390 464

SrTaTbThTiTl


PdPtRhRu

CO2S

B54

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI



15.00 1.80 1.20 0.90 0.40 2.70 7.60 8.2012.00 0.40 0.00 0.00 0.50 1.20 7.30 7.503.27 5.26 5.56 4.66 3.81 5.18 4.93 5.97

12.30 31.80 35.80 37.80 43.90 31.60 22.50 21.70

0.31 0.06 0.04 0.03 0.01 0.14 0.25 0.29

810 1700 2270 3820 5580 1370 1710 1860

274 8170 2510 3180 3170 1770 876 958

SrTaTbThTiTl


PdPtRhRu

CO2S

B55

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI



6.20 13.00 7.50 13.00 8.90 14.00 1.20 14.008.50 15.00 6.70 6.30 7.20 11.00 0.40 3.005.25 8.81 5.87 7.62 4.85 5.32 4.58 4.28

20.50 6.36 21.30 19.00 23.30 9.63 36.80 27.30

0.22 0.85 0.24 0.39 0.33 0.51 0.02 0.63

1560 260 1280 490 1220 220 8270 210

748 178 756 346 458 168 1860 710

SrTaTbThTiTl


PdPtRhRu

CO2S

B56

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI



0.70 15.00 8.20 9.00 8.00 7.80 9.60 9.302.00 2.70 10.00 9.10 10.00 9.00 7.80 7.805.14 1.77 4.90 5.73 5.51 5.30 5.83 5.61

34.60 3.29 16.50 19.30 20.20 20.80 21.30 20.70

0.01 0.34 0.40 0.64 0.62 0.49 0.62 0.55

9180 210 1090 1220 1310 1090 1060 900

2440 102 1040 1090 1060 1060 1110 948

SrTaTbThTiTl


PdPtRhRu

CO2S

B57

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI



5.10 10.00 - 13.00 13.00 6.70 12.00 14.003.80 11.00 - 11.00 11.00 8.60 6.60 17.004.64 5.67 - 5.47 6.44 4.56 6.20 6.01

26.60 15.30 - 6.93 13.20 22.20 21.20 9.06

0.26 0.58 - 1.35 1.20 0.50 0.64 0.80

2110 940 - 120 740 1710 1260 60

1490 758 - 84 340 1020 1040 116

SrTaTbThTiTl


PdPtRhRu

CO2S

B58

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI



17.00 8.60 7.60 10.00 7.50 0.60 0.60 2.4015.00 8.10 8.20 8.90 9.10 2.60 3.40 0.206.11 6.21 6.30 8.11 6.00 3.46 4.56 4.51

8.45 20.00 20.80 17.40 20.80 37.80 37.40 38.70

1.25 1.30 1.10 3.48 1.32 0.09 0.06 0.18

130 960 880 40 1120 710 870 970

94 790 996 122 884 2070 2680 2430

SrTaTbThTiTl


PdPtRhRu

CO2S

B59

Karelian Craton

SampleLab

BatchAl2O3CaO

Fe2O3K2OLOI


WP-98 WP-99Ultratrace Ultratraceu157155 u157155

0.40 1.101.90 1.104.11 4.73

37.40 37.90

0.05 0.09

610 1300

3150 2820

SrTaTbThTiTl


PdPtRhRu

CO2S

B60

Appendix C. Data Quality

a. Sampling Techniques The majority of samples utilized in this thesis represent new data and were collected by the author, with additional data collected from external sources (e.g. Long-Victor and Maggie Hays) and as such sampled by other individuals. External samples were cross-checked for depth, lithology, and volcanic facies if the drill hole was sampled additionally by the author. Prior to sampling, 3D computer models (Leapfrog®, Surpac®, and Fracis®) were developed to identify diamond drill holes characterizing a specific area or contained units and lithologies of interest. Sampling by the author was restricted to diamond drill core within the Long-Victor and Maggie Hays deposits, and outcrop samples in northern Finland and Norway. The drill core sampling procedure comprised: laying out the drill core, core preparation (washing, checking meterage, and identification of missing intervals), lithological contact identification, compiling a brief summary lithological log, sample selection, mark core trays/boxes, and sample core. Core samples were typically split with a diamond saw and the representative sample left in the core tray. However, due to the unavailability of splitting equipment, the complete core was sampled in minor number of circumstances.

b. Chemical Analysis

Laboratory Sample Preparation

Rock samples were cut with a diamond saw to obtain a representative slab and material for thin section preparation. Sample material for whole-rock geochemical analyses was cut additionally to remove both weathering rinds and crosscutting veins. Prepared material was coarse crushed (<5 mm-15 mm) using a jaw crusher at the University of Western Australia. The jaw crusher was flushed with quartz, cleaned with a wire brush, wiped down with acetone and dried with compressed air after each sample. Crushed samples were packed in clear locking plastic sample bags for transport to the analytical labs.

Two geochemical labs were utilized for this research: Ultratrace located in Perth, Western Australia and Geolabs located in Sudbury, Ontario, Canada. Although, preferably one lab would be used for data consistency, the cost and long turn around time on sample analyses prevented this and warranted the use of a second analytical facility. Both labs carried out the same analytical methods and samples were

C. . 1

analyzed for both major, trace and chalcophile elements utilizing the following techniques. Reported lowest levels of detection (LLD) for the analytical methods are shown in Table C.1.

Analytical Methods

Ultratrace

XRF (X-ray florescence)

Major elements (Al2O3, CaO, Fe2O3, K2O, MgO, MnO, Na2O, P2O5, SiO2, TiO2 , Cr2O3, SO3) and select trace elements (Ni, Cu, Ba, Rb, Sr, V, Zr,) were analyzed by wavelength dispersive X-Ray fluorescence (XRF) on a 0.66 gram sample fused to a glass bead.

ICP-MS (Inductively coupled plasma-mass spectrometry)

Minor elements (Y, Th, Nb, Hf, Ta, La, Ce, Pr, Nd, Sm, Eu, Gd, Dy, Tb, Ho, Er, Tm, Tb, Lu, Te, Se) were analyzed by ICP-MS following four acid (hydrofluoric, hydrochloric, perchloric, and nitric) digestion of a 0.3g sample.

ICP-OES (Inductively coupled plasma-optical emission spectrometry)

Major and minor elements (Al, Ti, Mg, Mn, Na, K, Ca, Fe, P, Cr, V, Ni, Cu, Co, S, Zr) were analyzed by ICP-OES following four acid (hydrofluoric, hydrochloric, perchloric, and nitric) digestion of a 0.3g sample.

Fire Assay ICP-MS

Platinum-group elements (Au, Pt, Pd, Rh, Ru and Ir) were analyzed by ICP-MS following a nickel-sulfide fire pre-concentration, aqua regia dissolution of the sulfide button and co-precipitation of the PGE with tellurium from a 25g sample.

Geolabs

XRF (X-ray florescence)

Major elements (Al2O3, CaO, Fe2O3, K2O, MgO, MnO, Na2O, P2O5, SiO2, TiO2) were analyzed by wavelength dispersive X-Ray fluorescence (XRF) on a 4g sample which was fused to a glass bead with a borate flux. Additional and duplicate analyses of select trace elements (As, Ba, Cr, Cu, Ni, Rb, Sc, Sr, V, Y, Zr) were analysed by XRF on a 10g sample pressed into a 40mm pellet excited by a Rh target.

C. . 2

ICP-MS (Inductively coupled plasma-mass spectrometry)

Minor elements and some major elements (Al, Sb, Ba, Be, Bi, Cd, Ca, Ce, Cs, Cr, Co, Cu, Dy, Er, Eu, Gd, Ga, Hf, Ho, Fe, La, Pb, Li, Lu, Mg, Mn, Mo,Nd, Ni, Nb, P, K, Pr, Rb, Sm, Sc, Na, Sr, S, Ta, Tb, Tl, Tm, Sn, Ti, W, U, V, Yb, Y, Zn, Zr) were analyzed by ICP-MS following a four acid (hydrofluoric, hydrochloric, perchloric, and nitric) closed beaker digestion of 0.5g sample.

Fire Assay ICP-MS

Platinum-group elements (Pt, Pd, Rh, Ru and Ir) were analyzed by ICP-MS following a nickel-sulfide fire assay pre-concentration step, aqua regia dissolution of the sulfide button and co-precipitation of the PGE with tellurium from a 15g sample.

Table C.1. Lowest Level of Detection (LLD) reported by both analytical labs for each analytical technique.

Ultra trace Geolabs

ICP-OES (ICP102)

ICP-MS (ICP302)

XRF (XRF202)

FA ICP-MS

NSF001

FA ICP-MS

(IMP-200)

ICP-MS (IMC-100)

XRF (XRF-M01)

Element ppm ppm % Oxide ppb ppb ppm % Oxide SiO2 0.01 0.01 Al2O3 10 0.01 0.01 TiO2 1 0.01 7 0.01 MgO 10 0.01 0.01

Fe2O3 10 0.01 0.01 K2O 20 0.01 0.01 CaO 10 0.01 0.01 Na2O 10 0.01 0.01 MnO 1 0.01 0.01 BaO 1 1 0.01 0.8 P2O5 10 0.001 0.01 Cr2O3 5 0.001 3 V2O5 2 0.01 0.8

Ni 1 2 0.001 1.6 Cu 1 1 0.001 1.4 Co 2 1 0.001 0.13 Au 1 0.22 Pt 1 0.17 Pd 1 0.12 Rh 1 0.02 Ru 1 0.08 Ir 1 0.01

C. . 3

Table C.1 continued. Lowest Level of Detection (LLD) reported by both analytical labs for each analytical technique.

Ultra trace Geolabs

ICP-OES (ICP102)

ICP-MS (ICP302)

XRF (XRF202) NSF001

FA ICP-MS (IMP-200)

ICP-MS (IMC-100)

XRF (XRF-M01)

Element ppm ppm % Oxide ppb ppb ppm % Oxide S 10 0.001 Zn 1

2 0.001 7 Sb 0.1 0.01 0.04 Ag 1 0.5 Pb 5 1 0.001 0.6 Sn 5

1

0.01

0.16

As 5 0.5 Be 1 0.1 0.04 Cd 0.5

0.013 Cs 0.1

0.013 Dy 0.05

0.009 Er 0.05

0.007 Eu 0.05

0.0031 Ga 0.2

0.04 Gd 0.2 0.009 Ho 0.02 0.0025 In 0.02 0.0018 La 0.05 0.04 Li 10 0.5

0.4 Lu 0.02 0.002 Mo 2 0.2

0.08 Nb 0.5 0.001 0.028 Nd 0.05 0.06 Pr 0.02 0.014 Rb 0.02 0.01 0.23 Re 0.1 Sc 1 2

1.1 Se 1 Sm 0.05 0.012 Sr 2 0.1 0.01 0.6 Ta 0.05 0.01 0.023 Tb 0.02 0.0023 Te 0.2 Th 0.05 0.001 0.018 Tl 0.1

0.005 Tm 0.02 0.0019 U 0.05

0.001

0.011 Y 5 0.1

0.05 Yb 0.05

0.009 Zr 1 1 0.01 6

C. . 4

c. Error in Data Monitoring of analyses quality was carried out by blanks, standards and duplicate samples. Duplicate sample analyses are presented in Figure C.1 as coefficient of variation (CV) as defined by:

CV=2/√2*(abs(Ni-Di))/(Ni+Di)

And as half absolute relative difference (HARD) as defined by:

HARD= (abs(Ni-Di))/(Ni+Di)

Figure C.1. Duplicate analyses plots for the platinum group elements. Coefficient of Variation (CV) vs. Duplicate Mean, with Relative Error shown (RE). Half absolute relative difference (HARD) vs. Ranked Percentile with vertical line demarking 95th percentile and horizontal lines 2 standard deviations (2s).

C. . 5

d. Quality Assurance and Control

Quality assurance and control at Ultratrace and Geolabs included the analysis of two

internal quality control samples for each analysis batch. In addition to these control

samples, in-house reference samples were analysed periodically to provide

information on the quality of the measurement data. In addition, laboratory duplicate

samples were analyzed through out the duration of the larger AMRIA P710A

project. These laboratory duplicate samples include uncertainty due to sample

heterogeneity, sample preparation and analytical measurements, and thus provide the

most realistic estimate of the quality of the data. The laboratory duplicate samples

were used to quantify the precision of the concentration data reported by the

laboratory.

Precision

Laboratory duplicate samples were utilized to evaluate precision, taking into

consideration all contributing factors listed previously. Precision evaluation was

conducted using the method of Thompson and Howarth (1976). This method

recognizes the precision of an analytical method varies as a function of

concentration. Two parameters are derived from this methodology: 1) a limiting

value for method precision at high concentrations; and 2) the method precision at

zero concentration (detection limit of the analytical method) as outlined below. The

limitations recognized by this methodology are a function of the number of duplicate

samples (Stanley and Lawie, 2007). However, the Thompson and Howarth (1976)

method provides a means to assess a measurement of error at any concentration.

Thompson and Howarth (1976) Precision Method

1. Calculate the mean and absolute difference of each pair of duplicate analyses:

(Ni+Di)/2 and

Absolute(Ni-Di) Where Ni represents the normal sample i, and Di the duplicate analysis of sample i.

2. Sort these results into order of increasing means.

3. Split the means into groups of 11.

4. Calculate the median value of the 11 means in each group.

C. . 6

5. Calculate the mean value of the 11 means in each group.

6. Regress the median values against the mean values using a least-squares regression.

7. Multiply the regression coefficients by 1.047 to give estimates of the slope (k) and the intercept (s) of the precision model.

The intercept (s) gives an estimate of the standard deviation at zero concentration.

The slope coefficient (k), gives an estimate of the precision to which the method

approaches at high concentrations. The application of this methodology is

graphically shown for Pt analyses by nickel sulfide fire extraction ICP-MS carried

out by Geolabs (Fig. C.2 and C.3) and summarized for all the chalcophile elements

in Table C.2. Evaluation of precision by this methodology was not possible for all

elements at both analytical facilities. Platinum group element analyses were

evaluated for both laboratories, but majors only for Ultratrace, as insufficient

duplicate analyses prevented similar evaluation for Geolabs.

Figure C.2. A graphical representation of the estimation of precision by the regression of duplicate analyses using analytical data determined by FAICP-MS by G

-eolabs.

Precision as a function of concentration was determined utilizing the following equation:

Pc = 2*S0/C + 2*K Where S0 = Y intercept, K = slope, and C = concentration.

Precision as a function of concentration was determined for the observed range of

sample compositions (Fig. C.3) and is summarized in Table C.2 as median values

over the observed range.

C. . 7

Figure C.3. Calculated precision as a function of concentration for platinum determined by FA-ICP-MS by Geolabs.

Table C.2. Summary of Precision as determined for major and chalcophile elements through duplicate analyses. S0 and K are Y-intercept and slope, respectively, from linear regressed duplicate analyses described previously. MDL = method detection limit. Precision (%) is a median value over the compositional range given. Range in wt% for oxides, ppm for Cr, Ni, and Cu, and ppb for the PGE.

Geolabs Ultratrace

Element Range S0 K MDL (3*S0)

Precision S0 K

MDL (3*S0)

Precision

SiO2 20 – 62 - - - - 0.032 0.001 0.10 0.004 TiO2 0.03 - 1.5 - - - - 0.002 0.002 0.01 0.009

Al2O3 0.2 - 18 - - - - 0.010 0.005 0.03 0.012 MgO 10 - 50 - - - - 0.010 0.004 0.03 0.009

Cr 50 - 4000 - - - - 14.400 0.004 43.20 0.019 Ni 50 - 5000 - - - - 7.170 0.001 21.51 0.008

Cu 5 - 300 - - - - 1.740 0.100 5.22 0.214 Ir 0.2 - 5 0.002 0.041 0.01 0.08 0.112 0.049 0.34 0.17

Ru 0.2 - 5 0.198 0.023 0.59 0.19 0.302 0.040 0.91 0.29 Rh 0.2 - 5 0.019 0.064 0.06 0.13 0.171 0.018 0.51 0.16 Pt 0.2 - 12 0.114 0.038 0.34 0.11 0.341 0.039 1.02 0.18 Pd 0.2 - 12 0.170 0.009 0.51 0.07 0.362 0.011 1.09 0.13

The method detection limits (MDL) as determined by 3 x the Y-intercept (3*S0),

specifically for the PGE are comparable to the lowest level of detection (LLD)

reported by the analytical labs (Table C.1). The major elements exhibit low values

for both intercept and slope coefficients and are not statistically different from zero

at a specified confidence level. The low values indicate that the error in the

analytical method is not changing significantly over the range of concentrations

studied and MML calculated by (3*S0) is not a accurate estimate. Consequently,

analytical precision is estimated by a different approach utilizing an average

precision over the entire range (grand median difference/grand mean). This approach

results with an average precision of 0.0012% for MgO and 0.008% for TiO2.

C. . 8

The precision estimates for major elements and precision as a function of

concentration for the PGE were used to calculate total maximum uncertainties in the

derived equations for chalcophile elements as a function of MgO content (Appendix

D.). Total maximum uncertainties for the PGE are summarized in Table C.3 and

shown graphically in Figure C.4 for Pt. Total maximum uncertainties are a

combination of major element precision estimates from Ultratrace and PGE

precision estimates from Geolabs and Ultratrace. Insufficient major element data

was available from Geolabs to calculate both labs independent of each other.

However, it is assumed the major element precision from Geolabs is at a minimum

equivalent to Ultratrace.

Figure C.4. Plots of calculated Pt (see Appendix D) versus MgO and calculated Pt/Tipmn versus MgO, with total maximum sample uncertainty shown by dashed lines. Total uncertainty includes precision estimates of MgO and TiO2 as derived from Ultratrace duplicate analyses and Pt as derived from Geolabs.

Table C.3. Calculated total maximum uncertainty for the chalcophile elements. Values are median values covering the range of compositions observed in the Long-Victor system (9-48 wt% MgO).

Geolabs Ultratrace ppb PGE/Ti ppb PGE/Ti

Pt 1.23 0.17 2.15 0.31 Pd 1 0.26 1.8 0.47 Rh 0.29 0.33 0.763 0.87 Ru 1.27 0.26 1.96 0.4

Ir 0.8 0.25 1.37 0.44 Ni (ppm) - - 959 0.5

As a consequence of the difference in data quality from the two labs an uncertainly

of 2 ppb is used for Pt and Pd, even though the uncertainly obtained for Geolabs is

below this. A 0.5 ppb uncertainty value used for Rh, which is above the value for

Geolabs, but below that calculated for Ultratrace. An uncertainly of >900 ppm is

identified for Ni and may partially explain the lack of correlation in mineralization

C. . 9

signatures between Pt, Pd and Ni as the uncertainty exceeds the magnitude of the

signature.

Accuracy

The accuracy of the analytical results can be roughly inferred from the comparison

of elemental analyses by two different analytical methods. This is possible for five

elements (MgO, Al2O3, TiO2, CaO, and Ni) for a subset of data analyzed by

Ultratrace with both ICP-MS and ICP-OES and for three elements (Cr, Cu, Ni) for

samples analyzed by Geolabs by ICP-MS and XRF. All methods produce total

element concentrations. However, if minerals are not completely dissolved the ICP-

MS and OES concentrations will be less than the XRF concentrations. Linear

regressions of the duplicate samples by differing analytical techniques show that the

proportional bias between the three techniques was generally <1% for major

elements. The only bias identified was that for Cr. This Cr discrepancy was

attributed to low recoveries of Cr by ICP-MS in samples prepared by acid

dissolution caused by insoluble residues of chromite (Fig. C.5).

C. . 10

Figure C.5. Comparison between ICP-MS and ICP-OES, and ICP-MS and XRF analyses for elements that were analyzed by both analytical methods. Linear regressions and r2 correlation coefficients shown for each data set from Ultratrace and Geolabs.

C. . 11

e. References Stanely, C.R., Lawie, D., 2007. Average Relative Error in geochemical determinations: Clarification, calculation, and a Plea for consistency. Exploration and Mining Geology, v. 16, p. 267-275.

Thompson, M., Howarth, R.J., 1976. Duplicate analysis in geochemical practice, Part 1. Theoretical approach and estimation of analytical reproducibility. Analyst, v. 101, p. 690-698.

C. . 12

C. . 13

Index

Appendix C. Data Quality ........................................................................................... 1 a. Sampling Techniques ...................................................................................... 1 b. Chemical Analysis ........................................................................................... 1

Laboratory Sample Preparation .......................................................................... 1 Analytical Methods .............................................................................................. 2

c. Error in Data .................................................................................................... 5 d. Quality Assurance and Control ....................................................................... 6 e. References ..................................................................................................... 12

List of Figures

Figure C.1. Duplicate analyses plots for the platinum group elements. Coefficient of Variation (CV) vs. Duplicate Mean, with Relative Error shown (RE). Half absolute relative difference (HARD) vs. Ranked Percentile with vertical line demarking 95th percentile and horizontal lines 2 standard deviations (2s). ........ 5

Figure C.2. A graphical representation of the estimation of precision by the regression of duplicate analyses using analytical data determined by FA-ICP-MS by Geolabs. ................................................................................................... 7

Figure C.3. Calculated precision as a function of concentration for platinum determined by FA-ICP-MS by Geolabs. ............................................................. 8

Figure C.4. Plots of calculated Pt (see Appendix D) versus MgO and calculated Pt/Tipmn versus MgO, with total maximum sample uncertainty shown by dashed lines. Total uncertainty includes precision estimates of MgO and TiO2 as derived from Ultratrace duplicate analyses and Pt as derived from Geolabs. ..... 9

Figure C.5. Comparison between ICP-MS and ICP-OES, and ICP-MS and XRF analyses for elements that were analyzed by both analytical methods. Linear regressions and r2 correlation coefficients shown for each data set from Ultratrace and Geolabs. ..................................................................................... 11

List of Tables

Table C.1. Lowest Level of Detection (LLD) reported by both analytical labs for each analytical technique. .................................................................................... 3

Table C.2. Summary of Precision as determined for major and chalcophile elements through duplicate analyses. S0 and K are Y-intercept and slope, respectively, from linear regressed duplicate analyses described previously. MDL = method detection limit. Precision (%) is a median value over the compositional range given. Range in wt% for oxides, ppm for Cr, Ni, and Cu, and ppb for the PGE. 8

Table C.3. Calculated total maximum uncertainty for the chalcophile elements. Values are median values covering the range of compositions observed in the Long-Victor system (9-48 wt% MgO). ............................................................... 9

Appendix D. Methodology of PGE as a Fn(MgO)

a. Purpose

A baseline that represents the chalcophile element abundance of a sample that

crystallized under sulfur-undersaturated conditions (e.g. no sulfide influence) is

required to quantify residual positive and negative anomalies. Crystallization is

inclusive of both fractionation and accumulation, with the latter being dominant in

komatiite systems. Natural data sets were utilized instead of numerical fractionation-

accumulation models to circumvent unresolved partition coefficients. The

methodology presented builds upon TiO2 normalization presented by Barnes et al.

(2004; 2007) and Fiorentini et al. (2010).

b. Assumptions

Platinum, palladium and titanium act as incompatible elements during

komatiite crystallization (olivine and chromite dominant phases).

Limited mobility of TiO2 and MgO within the system

Figure D.1. A. Plot of Pt (ppb) versus MgO (wt%) for all Kambalda data with sulfur < 0.25 wt%, showing general negative correlation with MgO with potential Pt depletion (D) and enrichment (E) overprinting trend as shown by arrows. B. TiO2 versus MgO (wt%) showing strong negative correlation between the two elements.

D. . 1

c. Procedure

Data filtered for sulfur content less than 0.25 wt%

o Interpretation: Pt and Pd are strongly chalcophile elements any sulfur

present at time of crystallization will strongly partition Pt and Pd,

elevating the PGE content of the sample.

Pt and Pd values are normalized by TiO2 content of the sample and mantle

values (mantle normalizing values from McDonough and Sun, 1995: Pt 7.1

ppb, Pd 3.9 ppb, Ti 1205 ppm) see Figure D.1

o Interpretation: if Pt, Pd and Ti are all incompatible with olivine the

relative ratio between the two (Pt or Pd/Tipmn) should remain a

constant.

Pt/Ti pmn is plotted against Pd/Ti pmn. See Figure D.2

Figure D.2. Plots Pt/Tipmn versus MgO wt% and Pd/Tipmn versus MgO wt% of all Kambalda data with sulfur < 0.3 wt%, showing constant value with varying MgO content. Deviation from a constant value shown as D (depletion) and E (enrichment).

o Interpretation: samples which preserve initial Pt/Tipmn and Pd/Tipmn

ratios will plot as a cluster of data (Fig. D.3). Deviations from this are

attributed to the following

1. Pt and Pd enrichment do to low S mineralization

2. Pt and Pd depletion do to previous S-saturation and

chalcophile element removal.

3. Pt and or Pd mobility within the system (metamorphic)

D. . 2

Figure D.3. Plot of Pd/Ti pmn versus Pt/Ti pmn for all Kambalda samples with S<0.3wt%. Trend lines shown for low sulfur Pt and Pd enrichment/mineralization (Pt+Pd En), Pt and Pd depletion (Pt+Pd De) and enrichment or depletion of either Pt or Pd from a constant value.

Data set is filtered to remove samples which exhibit Pt and/or Pd enrichment

and depletion. Filtering process involved plotting histograms to examine the

distribution of values and sequentially removing the outlying data points, as

summarized in Table D.1.

Table D.1. Step results of iteratively filtered Kambalda Dome data set.

Step 1 Step 2 Step 3 Final

Pt/Ti Pd/Ti Pt/Ti Pd/Ti Pt/Ti Pd/Ti

Median 0.63 1.24 0.66 1.24 0.67 1.24

Number 203 204 123 133 113 111

Final median values were obtained for Pt/Tipmn (0.67) based on 113 samples

and Pd/Tipmn (1.24) on a 111 samples.

o Interpretation: These ratios represent Pt and Pd content of the magma

at any point along its fractionation and cumulate history (10 to 50

wt% MgO).

Given an estimated precision of ±2 ppb on Pt and Pd with current analytical

techniques, and natural variability ranges of Pt/Tipmn are calculated (≥ 0.46 to

≤ 0.88) and Pd/Tipmn (≥ 0.89 to ≤ 1.65), as shown in Figure D.4.

D. . 3

Figure D.4. Final data set (n=75) from Kambalda which falls within ± 2 ppb of calculated Pt/Tipmn and Pd/Tipmn ratios.

Figure D.5. Primitive mantle normalized noble metal plot of select samples.

These samples were then utilized to plot chalcophile element (Pt, Pd, Ir, Rh, Ru,

Ni, Cu) versus MgO and determine best fit lines and line equations, providing

chalcophile element abundance as a function of MgO content.

Figure D.6. Ni (ppm) versus MgO and Ir/Tipmn versus MgO for Kambalda samples with linear regressions and R2 values.

D. . 4

d. Results

The derived equations for chalcophile element as Fn(MgO) for the Kambalda Dome

are listed in Table D.2. Calculated values based on the equations (at 24 wt% MgO)

are shown in Table D.3 with data from Kambalda Dome spinifex textured samples

for comparison.

The same methodology was utilized on the Maggie Hays System, although with a

smaller data set. Derived equations for the Maggie Hays system are listed in Table

D.4. Calculated values based on the equations are shown in Table D.5 with data

from Western Ultramafic Unit spinifex textured samples for comparison.

Table D.2. Chalcophile elements as a function of MgO as derived for the Kambalda Dome system (2.7 Ga Munro-type) with calculated R2 values

Ni Fn(MgO) = 90.04(MgO)-1175 r2 = 0.92 Pt Fn(MgO) = -0.369(MgO)+17.99 r2 = 0.77

Pd Fn(MgO) = -0.36(MgO)+18.0 r2 = 0.75

Ir Fn(MgO)a = 0.2125(MgO)-3.8694 r2 = 0.66

or Ir/Ti pmn Fn(MgO)b = 0.005e0.1473(MgO) r2 = 0.83

Rh Fn(MgO)a = -0.0366(MgO)+2.1166 r2 = 0.57

or Rh/Ti pmn Fn(MgO)b = 0.437e0.0209(MgO) r2 = 0.46

Ru Fn(MgO)a = 0.0234(MgO)+3.0641 r2 = 0.03

or Ru/Ti pmn Fn(MgO)b = 0.0794e0.0653(MgO) r2 = 0.75

Table D.3. Calculated chalcophile content of a theoretical Kambalda primitive magma (24 wt% MgO) compared with median spinifex textured samples (n=15: filtered to remove mineralizing signatures) from Kambalda Dome.

MgO Pt Pd Ir Ru Rh Ni Cu

Calculated 24 9.7 9.3 1.09 b 3.74 b 1.27 a 1001 48

Median Spfx 24.1 9.0 9.3 0.99 4.0 1.23 885 56

D. . 5

Table D.4. Chalcophile elements as a function of MgO as derived for the Maggie Hays System (2.9 Ga Barberton-type) with calculated R2 values

Ni Fn(MgO) = 83.516(MgO)-823.39 r2 = 0.85 Pt Fn(MgO) = -0.3379(MgO)+17.752 r2 = 0.70

Pd Fn(MgO) = -0.2304(MgO)+12.168 r2 = 0.67

Ir Fn(MgO)a = 0.031(MgO)+1.1314 r2 = 0.15

or Ir/Ti pmn Fn(MgO)b = 0.092(MgO)-2.1132 r2 = 0.61

Rh Fn(MgO)a = -0.0269(MgO)+1.8575 r2 = 0.49

or Rh/Ti pmn Fn(MgO)b = 0.058(MgO)-0.8863 r2 = 0.49

Ru Fn(MgO)a = -0.0327(MgO)+6.3018 r2 = 0.04

or Ru/Ti pmn Fn(MgO)b = 0.0995(MgO)-2.0514 r2 = 0.56

Table D.5. Calculated chalcophile content of a theoretical Maggie Hays primitive magma (26.8 wt% MgO) compared with median spinifex textured samples (n=7: filtered to remove mineralizing signatures) from Western Ultramafic Unit.

MgO Pt Pd Ir Ru Rh Ni Cu

Calculated 26.8 8.7 6.0 2.0 b 5.4 b 1.1 a 1406 32

Median Spfx 26.8 11.6 8.0 2.2 6.3 1.5 1275 74

Checks

Pt and Pd plot as strongly incompatible elements

Check TiO2 and MgO mobility against other incompatible major element

(Al2O3) and REE (Y, Yb, Dy, Gd)

Check Pt/Ti pmn versus Pd/Ti pmn against Pt/Dy pmn versus Pd/Gd pmn.

Check Pt/Tipmn and Pd/Tipmn against Pt/Alpmn and Pd/Alpmn.

D. . 6

e. References Barnes, S.J., Hill, R.E.T., Perring, C.S., Dowling, S.E., 2004A. Lithogeochemical exploration for

komatiite-associated Ni-sulfide deposits: strategies and limitations: Mineralogy and Petrology, v. 82, p. 259-293.

Barnes, S.J., Lesher, C.M., Sproule, R.A. 2007. Geochemistry of komatiites in the Eastern Goldfields Superterrane, Western Australia and the Abitibi Greenstone Belt, Canada, and implications for the distribution of associated Ni-Cu-PGE deposits: Applied Earth Science 116, p. 167-187.

Fiorentini, M.L., Barnes, S.J., Lesher, C.M., Heggie, G.J., Keays, R.R., Burnham, O.M., 2010. Platinum-group element geochemistry of mineralized and non-mineralized komatiites and basalts: Economic Geology,

McDonough, W.F., Sun, S-S., 1995. The Composition of the Earth: Chemical Geology, v. 120, p. 223-253.

D. . 7

D. . 8

D. . 9

Index

Appendix D. Methodology of PGE as Fn(MgO) .......................................................... 1 a. Purpose .............................................................................................................. 1 b. Assumptions ...................................................................................................... 1 c. Procedure ........................................................................................................... 2 d. Results ............................................................................................................... 5 e. References ......................................................................................................... 7

List of Figures

Figure D.1. A. Plot of Pt (ppb) versus MgO (wt%) for all Kambalda data with sulfur < 0.25 wt%, showing general negative correlation with MgO with potential Pt depletion (D) and enrichment (E) overprinting trend as shown by arrows. B. TiO2 versus MgO (wt%) showing strong negative correlation between the two elements. ................................................................................................................ 1

Figure D.2. Plots Pt/Tipmn versus MgO wt% and Pd/Tipmn versus MgO wt% of all Kambalda data with sulfur < 0.3 wt%, showing constant value with varying MgO content. Deviation from a constant value shown as D (depletion) and E (enrichment). ......................................................................................................... 2

Figure D.3. Plot of Pd/Ti pmn versus Pt/Ti pmn for all Kambalda samples with S<0.3wt%. Trend lines shown for low sulfur Pt and Pd enrichment/mineralization (Pt+Pd En), Pt and Pd depletion (Pt+Pd De) and enrichment or depletion of either Pt or Pd from a constant value. ........................ 3

Figure D.4. Final data set (n=75) from Kambalda which falls within ± 2 ppb of calculated Pt/Tipmn and Pd/Tipmn ratios. ................................................................. 4

Figure D.5. Primitive mantle normalized noble metal plot of select samples. ............. 4 Figure D.6. Ni (ppm) versus MgO and Ir/Ti pmn versus MgO for Kambalda samples

with linear regressions and R2 values. ................................................................... 4

List of Tables

Table D.1. Step results of iteratively filtered Kambalda Dome data set. ..................... 3 Table D.2. Chalcophile elements as a function of MgO as derived for the

Kambalda Dome system (2.7 Ga Munro-type) with calculated R2 values ........ 5 Table D.3. Calculated chalcophile content of a theoretical Kambalda primitive

magma (24 wt% MgO) compared with median spinifex textured samples (n=15: filtered to remove mineralizing signatures) from Kambalda Dome. ......... 5

Table D.4. Chalcophile elements as a function of MgO as derived for the Maggie Hays System (2.9 Ga Barberton-type) with calculated R2 values ..................... 6

Table D.5. Calculated chalcophile content of a theoretical Maggie Hays primitive magma (26.8 wt% MgO) compared with median spinifex textured samples (n=7: filtered to remove mineralizing signatures) from Western Ultramafic Unit. ....................................................................................................................... 6

Documents

The Application of Platinum Group Elements€¦ · prospective to host mineralization. The potential use of the chalcophile elements, specifically the platinum group elements (PGE: