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GEOMETALLURGICAL AND GEOLOGICAL EVALUATION OF THE HIGH-GRADE
POLYMETALLIC UNCONFORMITY-RELATED CIGAR LAKE URANIUM DEPOSIT
By
Andrew Joseph Kaczowka
A thesis submitted to the Graduate Program in the
Department of Geological Sciences & Geological Engineering
in conformity with the requirements for the
Degree of Master of Science
Queen’s University
Kingston, Ontario, Canada
December, 2017
Copyright © Andrew Joseph Kaczowka, 2017
ii
Abstract
The high-grade, polymetallic, unconformity-related, Cigar Lake uranium deposit, located in
northern Saskatchewan, Canada, is the focus of a detailed mineralogical, geochemical and geospatial
study aimed at geometallurgical characterization and reconstruction of its underlying geological history.
The main mineralization event occurred before 1468±93 Ma and was likely syngenetic with the basin
wide (ca. 1590 Ma) U mineralization event. Uraninite co-precipitated with common Pb-bearing arsenides,
sulpharsenides and sulphides. Molybdenum was mobile during primary mineralization, crystallizing
within Stage 1 uraninite. Relatively high and consistent δ34S values up to 14.6‰ indicate a significant
basinal and marine SO42- contribution and an open-system with a well-mixed source. Low 207Pb/206Pb
values (0.56–0.86) is imprinted on co-precipitated chalcopyrite, spatially and texturally associated with
Stage 1 uraninite, suggesting that the U and Cu were sourced from Archean minerals. Syn-ore calcite
equilibrated with a fluid having high δ18O values between +1.8 to +7.2‰, typical of diagenetic basinal
brine in the Athabasca Basin, and low δ13C values ranging from -22.4 to -21.8‰, consistent with a source
from underlying graphite.
Stage 2 fluid incursion, constrained by U/Pb dates at 1270-1163 Ma resulted in crystallization,
recrystallization and Pb absorption forming radiogenic (206Pb/204Pb ~1000) and selenite-bearing sulphides
and sulpharsenides concurrent with enhanced clay alteration. Stage 3 alteration, with U/Pb and Pb/Pb
dates between 947-755 Ma was responsible for extensive hematization and crystallization of anomalously
radiogenic (206Pb/204Pb ~3000–20,000) bornite, chalcocite and galena. The most recent deposit alteration
event, stage 4 (242-0 Ma), manifests as extensive coffinitization and re-mobilization of U as perched
mineralization. The paragenetic stages coincide with major, far-field tectonic and regional geological
events: (1) Stage 2 with the Grenville Orogeny, the Mackenzie dyke swarm and the Moore Lake olivine
diabase, (2) Stage 3 with the breakup of the supercontinent Rodinia, and (3) Stage 4 with recent meteoric
and glacial meltwaters.
iii
The paragenesis, geochronology and stable isotopes reveal a protracted history, with ongoing
episodic fluid incursion, driven by far-field tectonics that resulted in alteration and re-mobilization of
selected elements. Using the geometallurgical paradigm presented here, the geology at Cigar Lake can be
used to optimize and reduce risk during long-term mine and mill planning.
iv
Co-Authorship
This thesis and the manuscripts contained herein are the works of Andrew Joseph Kaczowka.
Chapter 2 and 3 are co-authored by Kurt Kyser, Tom Kotzer (thesis supervisors) and Cliff Revering. The
latter is the Chief Geologist at Cigar Lake and an industry sponsor who has provided logistical and
technical support to the project.
v
Acknowledgements
I extend my sincere gratitude to my supervisors Kurt Kyser and Tom Kotzer for providing me
with the opportunity to work on this project and for their endless support and encouragement. Pursuing
this work concurrently while working full time at Cigar Lake added logistical challenges for the project,
so I thank them for their constant guidance and flexibility.
I would like to thank Matthew Leybourne and Dan Layton-Matthews who stepped in as
supervisors for the final revisions and for preparing me for my thesis defense. This work was greatly
improved from their feedback and without their encouragement and support I would have never
completed the project.
I would like to thank Cliff Revering for making this project possible and for his continued
guidance throughout the project. Cliff provided technical, financial and logistical support and helped to
steer the project expanding the works applicability by targeting ongoing geometallurgical concerns.
I am indebted to the entire Cigar Lake geology department, specifically Levi Kalinsky, Mikkel
Tetland, Yi Wang, Andrew Masurat, Stephen Zubowski, Innis Hook, Greg Curry, Elaine Ruff, Adam
Gobeil and Sandy Ratt for collecting samples, collecting drillcore data and for insightful discussions on
the geology at Cigar Lake. This work was significantly improved by ongoing field discussions with the
team. This study utilized structural interpretations made by Shawn Harvey, and I would like to thank him
for his work and for ongoing structural discussions.
I would like to sincerely thank Queen’s Facility for Isotope Research (QFIR) lab staff,
specifically Donald Chipley, April Vuletich and Christabel Jean who provided guidance and support with
all the isotopic analyses. Agatha Dobosz and Brian Joy for assistance with the XRD, SEM and EMPA.
Thank you to Agatha for her constant remote VPN support over the last three years allowing me to
connect securely to the Queen’s Server from Saskatoon.
vi
Thanks to the Cameco Exploration department for allowing me access to their microscopes, quiet
office space, and meeting rooms. Discussions with many of the staff over the last three years has helped
immensely formulating the ideas and interpretations presented.
I would like to thank the Saskatchewan Research Council (SRC) for their work on the whole-rock
geochemistry. The SRC went above and beyond with two major analysis campaigns for additional arsenic
and sulphur data that went into the development of this thesis. They also provided me with lab space and
logistical support for shipping samples.
I would like to thank my parents Mike and Gail Kaczowka for their support and guidance not
only over the last three years but throughout my lifetime. They have raised me into the person I am today
and I would not have been in position to pursue my Master’s without their endless unconditional love and
support.
I would like to thank my wife Kyla Kaczowka who has endured this thesis right by my side. She
has stood by me through both the highs and lows but has never stopped encouraging me. I could not have
done this without her and this work is as much her accomplishment, as it is mine.
I would like to dedicate this work to Adam Gobeil. Adam was a geological technician in the
Cigar Lake geology department and he collected some of the samples and geological data used for this
research. Adam was tragically taken from us far too soon, but his attention to details, creativity and
insightfulness live on within the department.
I would also like to dedicate this work to my supervisor Kurt Kyser who passed away shortly
before my thesis defense. I am forever grateful to have had the opportunity to learn from such a great
geochemist.
vii
Table of Contents
Abstract……………………………………………………………………………………………………..ii
Co-Authorship……………………………………………………………………………………………...iv
Acknowledgements…………………………………………………………………………………………v
Table of Contents…………………………………………………………………………………….……vii
List of Figures……………………………………………………………………………………………....x
List of Tables……………………………………………………………………………………………..xiii
Chapter 1 INTRODUCTION…………………………………………………………………..…………...1
1.1 Overview………………………………………………………………………….…….……...1
1.2 Geology…………………………………………………………………………...…….……...3
1.2.1 Geology of the Archean and Proterozoic Basement………………………….……..3
1.2.2 Geology of the Athabasca Basin Group………………………………………..…....7
1.3 Unconformity-Related U Deposit Models………………………………………………….….9
1.4 Cigar Lake Deposit…………………………………………………………………………...13
1.5 Geometallurgy………………………………………………………………………………...16
1.6 Thesis Objectives and Rationale……………………………………………………………...17
1.7 Structure of Thesis…………………………………………………………………………....19
Chapter 2 CIGAR LAKE: GEOMETALLURGICAL ORE CHARACTERIZATION IN SUPPORT OF
MINING AND MILLING……………………………………………………………………………...…21
2.1 Abstract…………………………………………………………………………………….…21
2.2 Introduction………………………………………………………………………………...…23
2.2.1 Geological Setting……………………………………………………………….…25
2.2.2 The Cigar Lake Deposit……………………………………………………………26
2.2.3 Geometallurgical Considerations at Cigar Lake………………………………...…29
2.3 Methods…………………………………………………………………………………...…..30
2.3.3 Normative Error, Probability and Limitations………………………………..……34
2.4 Results…………………………………………………………………………………...……38
2.4.1 Mineralogy…………………………………………………………………………38
2.4.1.1 Uranium Ore Mineralogy………………………………………..………38
2.4.1.1.1 Uranium Mineral Chemistry…………………………….……42
2.4.1.2 Arsenide and Sulpharsenides……………………………………………46
2.4.1.2.1 Arsenide and Sulpharsenide Mineral Chemistry………..……49
viii
2.4.1.3 Sulphides………………………………………………………….……52
2.4.1.4 Zirconium…………………………………………………………..…..56
2.4.1.5 Clay Mineralogy……………………………………………………..…56
2.4.2 Three Dimensional Modelling of Element and Ore Distribution…………………57
2.4.3 Three Dimensional Modelling of Mineral Distribution……………………...……60
2.5 Discussion………………………………………………………………………………....…64
2.6 Conclusions………………………………………………………………………………..…66
Chapter 3 EVOLUTION OF THE HIGH-GRADE POLYMETALLIC UNCONFORMITY-RELATED
URANIUM CIGAR LAKE ORE BODY…………………………………………………………………68
3.1 Abstract………………………………………………………………………………...……..68
3.2 Introduction……………………………………………………………………………..…….70
3.3 Geological Setting……………………………………………………………………….……72
3.3.1 Regional Geology…………………………………………………………….……72
3.3.2 The Cigar Lake Deposit……………………………………………………………74
3.4 Methods………………………………………………………………………………………78
3.4.1 U/Pb and Pb Isotope Systematics……………………………………….…………82
3.5 Results………………………………………………………………………………...………83
3.5.1 Mineralogy and Textural Paragenesis………………………………………...……83
3.5.1.1 Uranium Ore Mineralogy…………………………………………..……83
3.5.1.2 Arsenide and Sulpharsenides……………………………………………88
3.5.1.3 Sulphides…………………………………………………………………….…...91
3.5.2 Mineral Geochemistry………………………………………………………..……93
3.5.2.1 Uranium Mineral Chemistry……………………………………….……93
3.5.2.2 Arsenides and Sulpharsenide Mineral Chemistry………………….……97
3.5.3 Geochemical and Mineral Modelling……………………………………….…..…99
3.5.4 Geochronology of U-Bearing Mineral Phase………………………………….…103
3.5.4.1 207Pb/206Pb Systematics of U-Bearing Mineral Phase…………….……103
3.5.4.2 U–Pb Systematics of U-Bearing Minerals…………………………..…104
3.5.5 Pb Isotopes of Sulpharsenides, Sulphides and Non-Metallic Gangue Minerals...106
3.5.6 Stable Isotopes……………………………………………………………………110
3.5.6.1 Sulphur Isotopes Systematics of Sulphides and Sulpharsenides………110
3.5.6.2 Carbon and Oxygen Isotopes in Carbonates………………………...…110
3.6 Discussion………………………………………………………………………………...…114
3.6.1 Geochronology of Far-Field Tectonics………………………………………...…114
ix
3.6.2 Pb Isotopes of Sulpharsenides, Sulphides and Non-Metallic Gangue Minerals…116
3.6.3 Characterization of Fluids……………………………………………………..…117
3.6.3.1 Oxygen Isotopes…………………………………………………….…117
3.6.3.2 S Isotopes on Sulphides and Sulpharsenides…………...……………...118
3.6.3.3 C and O Isotopes in Carbonates and Hydrocarbons………………...…119
3.6.4 Genetic Model…………………………………………………………….………121
3.7 Exploration Implications…………………………………………………………….………126
3.8 Conclusions………………………………………………………………………….………126
Chapter 4 GENERAL DISCUSSION, SUMMARY OF CONTRIBUTIONS AND
RECOMMENDATIONS FOR FUTURE WORK………………………………………………………129
4.1 General Discussion…………………………………………………………………….……129
4.2 Significant Contributions……………………………………………………………………131
4.2.1 Geometallurgical Contributions for Mining, Milling and Tailings Management...131
4.2.2 Cigar Lake Deposit Evolutionary Model Contributions……………………….…132
4.2.3 Contributions to Mineral Exploration………………………………………….…135
4.3 Recommendations for Future Work…………………………………………………………135
References………………………………………………………………………………………….…….137
Appendix A Summary of whole-rock geochemistry…………………………………………………….151
Appendix B Whole-rock geochemistry of XRD samples………………………………………………..152
Appendix C EMPA results from uraninite, coffinite, arsenides and sulpharsenides……….…..………..156
Appendix D XRD RIR mineralogy results………………………………………………………………166
Appendix E SEM-MLA mineralogy results………………………..……………………………………170
Appendix F LA-ICP-MS major element analysis on sphalerite…………………………………………172
Appendix G Normative algorithm design.........……………………………………………………….…173
Appendix H Normative mineral proportions ...……………………………………………….…………179
Appendix I SWIR clay mineralogy…...……………..…………………………………………………...181
Appendix J U-Pb isotopes from uraninite and coffinite………..………………...……………………...186
Appendix K Pb-Pb isotopes from sulphides and sulpharsenides………..………………….……………189
Appendix L LA-ICP-MS trace element analysis on uraninite and coffinite…...….……………………..196
x
List of Figures
Figure 1.1. Location of the Athabasca Basin and Cigar Lake deposit……………………………………..1
Figure 1.2. Air photograph of the Cigar Lake mine site with outlined study area………………..…..……2
Figure 1.3. Generalized Geological map of Northern Saskatchewan………………………………...…….3
Figure 1.4. Sub-Basins within the Athabasca Basin……………………………………………….….……7
Figure 1.5. Cross section of the Athabasca Basin…………………………………………………..………8
Figure 1.6. Genetic models for unconformity-related uranium deposits………………………………….11
Figure 1.7. Schematic illustration of the Cigar Lake deposit and surrounding alteration……………...…14
Figure 2.1. Location map of the Athabasca Basin and the Cigar Lake deposit with regional geological
provinces………………………………………………………………………………………………..…24
Figure 2.2. Cigar Lake mine site with outlined study area and mineral resource estimations……………25
Figure 2.3. Schematic illustration of the Cigar Lake deposit and surrounding alteration……………...…27
Figure 2.4. East Pod cross section along 10898 mine grid showing orebody facies and structural
interpretation………………………………………………………………………………………………28
Figure 2.5. Work-flow diagram outlining mineralogical and geochemical characterization methods……33
Figure 2.6. Sample location map…………………………………………………………………….……33
Figure 2.7. Flow chart demonstrating the steps taken to calculate normative mineralogy from whole-rock
geochemistry………………………………………………………………………………………………35
Figure 2.8. Relative error in normative algorithm versus mineral MLA percent…………………………36
Figure 2.9. Correlation between normative mineral proportions (wt. %) and MLA (wt. %) results……...37
Figure 2.10. Mineral paragenesis for the Cigar Lake ore body………………………………………...…39
Figure 2.11. BSE images of uraninite and coffinite………………………………………………………40
Figure 2.12. Bivariate diagrams showing the relationship between UO2 and the alteration elements (SiO2,
CaO, MnO, FeO), PbO and the EOCs (MoO3, SeO2, ZrO2) within uraninite and coffinite………………44
Figure 2.13. Bivariate diagrams showing the relationship between PbO and the chemical ages, the
alteration elements (SiO2, CaO, MnO, FeO), and the EOCs (MoO3, SeO2, ZrO2) within uraninite and
coffinite……………………………………………………………………………………………………45
Figure 2.14. BSE images of arsenides and sulpharsenides……………………………………………..…47
Figure 2.15. Ni+Co molar % versus As molar % plotted from whole-rock geochemical samples within the
Phase 1 orebody…………………………………………………………………………………...………48
Figure 2.16. Compositional plots for the arsenides and sulpharsenides…………………………..………51
Figure 2.17. Heat maps showing Bi, Se, and Co content within Stage 2 gersdorffite……………….……52
xi
Figure 2.18 Petrographic and BSE images of common sulphide minerals and textures……………….…54
Figure 2.19 Molybdenite shown in BSE and with MLA interpretation………………………………..…55
Figure 2.20 BSE images of pristine and altered zircon crystals…………………………………………..56
Figure 2.21 Three dimensional implicitly modelled geochemical grade shells for the Phase 1 pods….…59
Figure 2.22 Cross-section showing MLA/XTD mineral proportions with implicitly modelled grade shells
along 10930 mine grid……………………………………………………………………………….……62
Figure 2.23 Three dimensional implicitly modelled grade shells of normatively derived As-bearing
mineral: gersdorffite, cobaltite, niccolite and rammelsbergite……………………………………………63
Figure 2.24 Three dimensional implicitly modelled grade shells of normatively derived sulphides:
chalcopyrite, bornite, chalcocite and pyrite/pyrrhotite……………………………………………………64
Figure 3.1. Location of the Athabasca Basin and Cigar Lake……………………………………….……71
Figure 3.2. Air photograph of the Cigar Lake mine site with outlined study area…………………..……72
Figure 3.3. Generalized geological map of northern Saskatchewan………………………………………73
Figure 3.4. Schematic illustration of the Cigar Lake deposit and surrounding alteration……………...…75
Figure 3.5. East Pod cross section showing orebody facies with structural interpretation…………..……76
Figure 3.6. Mineral paragenesis for the Cigar Lake deposit and host rocks………………………………85
Figure 3.7. BSE images of uraninite and coffinite…………………………………………………..…….87
Figure 3.8 BSE images of arsenides and sulpharsenides………………………………………………….89
Figure 3.9 Petrographic and BSE images of sulphides……………………………………………………92
Figure 3.10 Bivariate plot showing the linear relationship between U-Pb chemical dates and alteration
elements. Chondrite normalized REE plots for uraninite and coffinite……………………………...……95
Figure 3.11 Compositional plots for the arsenides and sulpharsenides displaying Ni, Co, S, As, Se and Bi
contents……………………………………………………………………………………………………99
Figure 3.12 Three dimensional implicitly modelled geochemical grade shells for the Phase 1 pods…...101
Figure 3.13 Three dimensional implicitly modelled normative mineral proportions of Cu-bearing mineral
phase………………………………………………………………………………………………...……103
Figure 3.14 Histograms showing dates obtained using 207Pb/206Pb values from LA-ICP-MS data and
chemical ages obtained by EMPA…………………………………………………………………….…104
Figure 3.15 Dates obtained from U-Pb ICP-MS data……………………………………………………106
Figure 3.16 207Pb/206Pb versus 206Pb/204Pb and 207Pb/204Pb versus 206Pb/204Pb plots for common sulphides
and sulpharsenides……………………………………………………………………………………….108
Figure 3.17 Pb-Pb model ages for hematized illitic clay and bornite……………………………………110
Figure 3.18 Hand sample and petrographic images of carbonates and hydrocarbon buttons………...….113
Figure 3.19 δ34S values from Cigar Lake samples…………………………………………………….…119
xii
Figure 3.20 Genetic and evolutionary model for the Cigar Lake deposit.…………………………….…122
xiii
List of Tables
Table 2.1. Averaged EMPA results for uraninite and coffinite…………………………………...………43
Table 2.2. Summarized EMPA results for arsenides and sulpharsenides………………………………....50
Table 3.1. Averaged EMPA results for uraninite and coffinite…………………………………...………94
Table 3.2. Averaged EMPA results for arsenides and sulpharsenides……………………………………98
Table 3.3. Ranges in Pb isotopic concentrations..…………………………………………………….…107
Table 3.4. Stable isotopes…………………………………………………………………………..……112
1
Chapter 1
INTRODUCTION
1.1 Overview
Unconformity-related uranium deposits in the Athabasca Basin of Northern Saskatchewan are
unrivalled as the highest grade uranium deposits in the world (IAEA, 2009). Uranium mines in
Saskatchewan produce approximately 16% of total current global production of U3O8 (World-nuclearorg,
2016). Cigar Lake, with the highest mined uranium grades in the world, is the newest operation in the
mining district with first ore production beginning in 2014 and achieving commercial production status in
2015 (Cameco, 2015). Cigar Lake is located approximately 650 km north of the city of Saskatoon (Figure
1.1) and approximately 40 km west of the eastern margin of the Athabasca Basin (Bishop et al., 2016).
Figure 1.1: Location of the Athabasca Basin and Cigar Lake (black star). Also shown are the
locations of several other high-grade unconformity-type uranium deposits (black circles) and
northern communities (white squares) (modified from Bishop et al., 2016).
2
The ore at Cigar Lake is highly variable with a complex polymetallic geochemistry containing
elevated concentrations of As, Co, Cu, Mo, Ni, Se, and Zr (Bruneton, 1987; Reyx and Ruhlmann, 1993).
Elements such as As, Ni, Co, Mo and Se can be problematic during mining, milling and tailings
management and have been identified as elements of concern (EOC; Bishop et al., 2016). Accurate
reporting of mineral and mineraloid phases, geochemistry, and the empirical spatial distribution of these
attributes is crucial for mining and milling optimization, and operational reliability. The objective of this
research is to conduct geometallurgical characterization in support of the current geologic modelling that
provides predictive ore characteristics for mining and milling of the Phase 1 Cigar Lake deposit (Figure
1.2).
Figure 1.2: Air photograph of the Cigar Lake mine site with outlined study area, the Phase 1
Cigar Lake ore body pods. The Phase 1 deposit is divided into the East Pod and the West Pod.
3
1.2 Geology
1.2.1 Geology of the Archean and Proterozoic Basement
The crystalline basement rocks underlying the Athabasca Basin can be divided into three
lithotectonic zones (Figure 1.3): (1) Taltson magmatic zone that underlies the westernmost side of the
basin, (2) the Rae province that underlies the central basin, and (3) the Hearne Province on the
easternmost side (Card et al, 2007).
Figure 1.3: Geological map of Northern Saskatchewan with the stratigraphic divisions of the
Athabasca Group and basement geology. Major unconformity-related U deposits (Square),
including the Cigar Lake deposit (star), are indicated (Modified from Card et al., 2007, Ramaekers
et al., 2007).
4
The Taltson magmatic zone (Figure 1.3) is considered to be the southern extension of the Thelon
tectonic zone (Hoffman, 1988; Card et al., 2007). The north-south trending Taltson magmatic zone was
emplaced within Archean to Paleoproterozoic basement rocks comprising granitic gneisses, amphibolites
and pelitic gneisses ranging in age from 3.2–2.14 Ga (McNicoll et al., 2000; Card et al., 2007). During the
2.02–1.91 Ga Thelon Orogeny, continental magmatic arcs emplaced two intrusive suites, a 1.99–1.96 Ga
I-type quartz-diorite and granodiorite and a 1.95–1.92 Ga S-Type granite (McNicoll et al., 2000; Card et
al., 2007).
The Rae province is subdivided into five lithotectonic domains with possible extension beneath
the Athabasca Basin (Figure 1.3): (1) Zemlak, (2) Beaverlodge, (3) Tantato, (4) Lloyd, and (5) Clearwater
(Card et al., 2007). The Zemlak Domain is dominated by 2.71–2.33 Ga highly deformed mylonitic and
migmatitic upper amphibolite-facies orthogneiss with fragmented, dispersed psammitic remnants of the
older Murmac Bay Group (Ashton et al., 2007; Card et al., 2007). The central Zemlak contains the
younger Thluicho Group, a greenschist facies succession of conglomerate, arkose and argillite (Card et
al., 2007). The Zemlak Domain was intruded by 1.97–1.93 Ga leucogranites, regarded as a distal
emplacement of Taltson magmatics (Hartlaub et al., 2005; Ashton et al., 2009). The brittle to ductile
Black Bay Fault separates the Zemlak Domain from the Beaverlodge Domain (Ashton et al., 2009).
The western Beaverlodge Domain is composed of 3.0 Ga and 2.33–2.29 Ga granites
unconformably overlain by a Paleoproterozoic succession of metabasalt-pelite-quartzite cataloged within
the Murmac Bay Group (O’Hanley et al., 1994; Ashton and Hunter, 2003; Ashton et al., 2009). All of the
rocks are affected by upper to lower amphibolite metamorphism (Ashton et al., 2009). The East and
central parts of the domain are affected by higher grade granulite-facies metamorphism containing
granites, orthogneisses and migmatites (Ashton et al., 2007).
The western Tantato Domain comprises of 2.63–2.58 Ga granitoids, garnet-bearing migmatites,
and metabasites (Hanmer, 1997). The eastern side is dominated by the 3.4–3.1 Ga Chipman tonalite
batholith (Martel et al., 2008). The Tantato Domain was subjected to two granulite facies metamorphic
5
events at 2.55–2.52 Ga (Mahan et al., 2003; Ashton et al., 2009) and 1.91–1.90 Ga (Ashton et al., 2009),
the latter attributed to the Thelon Orogeny (Card et al., 2007).
On the Southern flank of the Athabasca Basin the poorly exposed Clearwater Domain is largely
defined by a prominent north to north-northeast trending electromagnetic signature and is interpreted as
intrusive, 1.84 Ga (Stern et al., 2003) granite and an older 2.5 Ga (Stern et al., 2003) granitoid gneiss
(Card et al., 2007). The Lloyd Domain is dominated by intrusive granitic to gabbroic orthogneiss yielding
ages between 1.985–1.975 Ga (Stern et al., 2003) with minor high-grade, 2.13–2.09 Ga (Bostock and Van
Breemen, 1994) supracrustal Careen Lake Group psammites to pelites (Card et al., 2007).
The Rae and Hearne Provinces are separated by a continental-scale lineament, the Snowbird
Tectonic Zone (Card et al., 2007). Debate still looms over whether the zone is a suture between the two
cratons (Hoffman, 1988), or a Paleoproterozoic to Archean intracratonic fault zone (Lewry and Sibbald,
1980; Ashton et al., 2009). The Hearne province contains three northeast trending lithotectonic domains
with possible extension beneath the Athabasca Basin (Figure 1.3): (1) Virgin River, (2) Mudjatik, and (3)
Wollaston Domains (Card et al., 2007).
The Virgin River and Mudjatik domains are similar lithologically, comprising Archean granitoid
gneisses overlain by discontinuous supracrustal, amphibolite facies, pelitic to psammitic gneisses, mafic
granulites and rare banded iron formations (Lewry and Sibbald, 1980; Card et al., 2007). U/Pb zircon
dates have indicated the underlying gneiss could be as old as 2.9–2.8 Ga (Orrell et al., 1999) with younger
2.64–2.58 Ga (Annesly et al., 1999) dates reported proximal to the Wollaston-Mudjatik boundary (Card et
al., 2007). The Virgin River and Mudjatik domains are intruded by late 2.68–2.6 Ga magnetite-bearing
monzogranites (Orrell et al., 1999). The boundary between the Virgin River and Mudjatik Domains is
defined by the Cable Bay shear zone, whereas the Mudjatik-Wollaston Domain boundary is transitional
with an increasing proportion of supracrustal rocks associated with the Wollaston Supergroup (Lewry and
Sibbald, 1980; Yeo and Delaney, 2007). The Virgin River and Wollaston Domains are distinguished from
6
the Mudjatik Domain by their linear structural styles in contrast to the refolded dome and basin structural
style of the Mudjatik (Lewry and Sibbald, 1980).
The basement rocks of the Wollaston Domain are similar to the Mudjatik and Virgin River
Domains, comprising slightly younger 2.59–2.56 Ga (Annesley et al., 1999) granitoid and amphibolite
gneisses (Yeo and Delaney, 2007). The Wollaston Supergroup unconformably overlies the Archean
basement as graphitic pelite, calc-silicate, calc-arkose, and arkose with anatectic granitoid segregations
having undergone amphibolite facies metamorphism during the Hudsonian Orogeny (Bruneton, 1993).
Yeo and Delaney (2007) interpret the supracrustal rocks as one complete Wilson Cycle with the opening
and closing of the Mannikewan Ocean. The lowermost Courtney Lake Group was deposited during the
initial stage of continental rifting (Yeo and Delaney, 2007). Drifting during open ocean/passive margin
stages was responsible for the deposition of the Souter Lake Group (Yeo and Delaney, 2007). The final
stages of the Wilson Cycle, ocean closure, resulted in passive margin to foreland basin environments
responsible for the deposition of the Geikie River Group (Yeo and Delaney, 2007). The age of the
Wollaston Supergroup is constrained by intrusion of the Wathaman Batholiths at 1.865–1.850 Ga (Ray
and Wanless, 1980; Van Schmus et al., 1987).
A regionally pervasive paleo-weathering profile exists at the unconformity between the
Athabasca Group and the underlying Archean and Paleoproterozoic basement (Macdonald, 1980). This
profile is characterized by an upper hematite and kaolinite-altered oxidized zone that transitions into a
green zone dominated by chlorite and illite (Hoeve and Sibbald, 1978; Macdonald, 1980).
7
1.2.2 Geology of the Athabasca Group
The Athabasca Basin covers an area of approximately 100,000 km2. This Paleo to
Mesoproterozoic intracratonic basin unconformably overlies the western Churchill Province between the
remnants of two orogenic belts, the 1.9 Ga Taltson Magmatic Zone and the 1.8 Ga Trans-Hudson
(Ramaekers, 1980). The initial accommodation and subsequent exhumation of the basal Athabasca Group
(Manitou Falls and Fair Point Formations) occurred in NE–SW trending Hudsonian basement faults
(Hoeve and Quirt, 1984, Armstrong and Ramaekers, 1985; Kyser et al., 2000). Ramaekers (1980)
identified three structurally controlled NE–SW trending elliptical sub-basins (Figure 1.4): the Jackfish,
Mirror and the Cree. These basins eventually coalesced to form the full extent of the Athabasca Basin.
Rapid uplift during the Trans-Hudson Orogeny provided siliciclastic input for the Athabasca Basin
beginning at ~1750–1700 Ma (Armstrong and Ramaekers, 1985; Kotzer et al., 1992; Kyser et al., 2000).
The end of sedimentation, based on Re–Os dating of the Douglas Formation occurred after 1,540 Ma
(Creaser & Stasiuk, 2007).
Figure 1.4: Position of sub-basins of the Athabasca Basin with major faults and dykes (From Kyser
et al., 2000; after Hoeve and Quirt, 1984)
8
Basin fill consists predominantly of unmetamorphosed quartz arenitic sandstone and
conglomerate overlain by siltstone, mudstone and dolostone (Ramaekers, 1990). Paleocurrent studies
infer sedimentation within the Athabasca Basin primarily derived from the northeast, east, and south
(Ramaekers et al., 2001). The depositional environment of the flat-lying, upward-fining, red-bed
succession was initially interpreted as fluviatile near-shore shallow-shelf environments (Ramaekers,
1990). In contrast, Ramaekers and Catuneanu (2004) interpreted the formation as lake and eolian
sediments. Due to the lack of interbedded volcanics and absence of fossils, paleoenvironmental
determination of the basin is difficult (Hiatt and Kyser, 2007). The Athabasca Group is divided into
eleven lithostratigraphic formations (Figure 1.5): Fair Point, Reilly, Read (formerly the MFa), Smart,
Manitou Falls, Lazenby Lake, Wolverine Point, Locker Lake, Otherside, Douglas, and the Carswell
(Ramaekers et al., 2007).
Figure 1.5: Cross section of the Athabasca Basin. D = Douglas Formation, C = Carswell Formation,
RY = Reilly Lake Formation, VR = Virgin River Magnetic High, STZ = Snowbird Tectonic Zone
(Modified from Jefferson et al., 2007).
9
Paleomagnetic reconstructions indicate that the Athabasca Basin occupied low to intermediate
northerly latitudes at ~1770 Ma, intermediate to high northerly latitudes at ~1650 Ma, and returned to low
latitude positions by about 1500 Ma (Pesonen et al., 2003). Fluid inclusion studies suggest the Athabasca
Basin reached a maximum burial depth of 5–7 km but has been uplifted and eroded to its current
thickness of 1–2 km (Pagel, 1975; Pagel et al., 1980; Ramaekers et al., 2007).
Subsequent to deposition, the Athabasca Group and underlying basement rocks were intruded by
mafic dykes of the Mackenzie dyke swarm at ca. 1267 Ma (LeCheminant and Heaman, 1989) and the
Moore Lake olivine diabase lopolith, in the southeastern Athabasca Basin at 1100 Ma (MacDougall and
Williams, 1993; French et al., 2002). The Athabasca Basin has been affected by at least two meteorite
impacts, the 478 Ma Carswell structure on the western side of the basin (Collier, 2007) and the Pasfield
Lake basement high on the eastern side (Bosman et al., 2011).
1.3 Unconformity-Related U Deposit Models
Unconformity-related uranium deposits are semi-massive pods and veins of uranium that occur
proximal to an unconformity between siliciclastic sedimentary basins and generally crystalline
metasedimentary basement (Kyser and Cuney, 2008; IAEA, 2009). Geochronology of uranium deposits
in the Athabasca Basin indicates that there have been three major fluid events at ca. 1590 Ma (the initial
and main mineralizing event), ca. 950 Ma and ca. 300 Ma with the latter two being dominantly
mobilization of the primary mineralization (Bruneton, 1987; Reyx and Ruhlmann, 1993; Philippe et al.,
1993; Fayek and Kyser, 1993; Fayek et al., 1997; Alexandre et al., 2009). These deposits can be divided
into two end members based on their relative location with respect to the unconformity: sandstone-hosted
and basement-hosted (Figure 1.6A).
Sandstone-hosted deposits occur within the basal siliciclastic sediments typically overlying
activated basement fault zones. These deposits occur as flattened, elongated lenses with a high-grade core
(>1% U3O8), surrounded by a lower grade halo (<1% U3O8), and clay-bounded at the periphery (e.g.
Bruneton, 1987). The ore consists of a complex, polymetallic assemblage containing sulphide and
10
arsenide mineral phases and elevated concentrations of Ni, Co, As, Cu, Zn, Mo, and in some deposits Se,
Ag, Au, and platinum-group elements (Fayek and Kyser, 1997; Kyser and Cuney, 2008). Sandstone
hosted deposits occur in association with large alteration zones extending up to 400 m wide, thousands of
meters along strike, and several hundred meters above the deposit consisting of Mg–Fe to Mg–Al
chlorite, muscovite, illite, kaolinite, and dravite (Hoeve and Sibbald, 1978; Hoeve and Quirt, 1984;
Kotzer and Kyser, 1995; Fayek and Kyser, 1997).
In contrast, basement-hosted deposits occur in crystalline basement rocks hosted within brittle to
semi-brittle reactivated fault zones proximal to the unconformity. These deposits typically consist of a
simple, monometallic ore assemblage containing only U and Cu with only traces of other metals (Fayek
and Kyser, 1997; Jefferson et al., 2007). Host-rock alteration is typically restricted to fault zones, and
consists of an inner zone of Mg–Fe chlorite with an outer zone dominated by muscovite and illite
(Alexandre et al., 2005; Kyser and Cuney, 2008).
Early models for the formation of unconformity deposits included laterite-like pre-Athabasca
Group paleo-weathering of the basement rocks and subsequent supergene enrichment (Derry, 1973;
Knipping, 1974; Langford, 1974, 1977). Alternatively, magmatic or metamorphic-hydrothermal processes
were proposed (Little, 1974; Morton, 1977; Munday, 1978). However, geochronology of the deposits and
surrounding host-rocks and a lack of evidence for syngenetic magmatism rendered these models
implausible leading to the development and wider acceptance of the diagenetic-hydrothermal model first
purposed by Hoeve and Sibbald (1976, 1978). The diagenetic-hydrothermal model postulated that
oxidized U-bearing diagenetic basinal brines were focused by reactivated structures and reacted with
basement rocks or basement-sourced fluids at the unconformity to produce mineralization (Hoeve and
Sibbald, 1976, 1978). Geochemical, geochronological and stable isotopic characterization of the deposits
and host-rocks over the last forty years has led to a refinement of the diagenetic-hydrothermal model and
characterization of the fluids involved in their formation (e.g. Kotzer and Kyser, 1995; Fayek and Kyser,
1997; Alexandre et al., 2005; Cloutier et al., 2011; Kyser and Cuney, 2015).
11
Figure 1.6: Genetic models for unconformity-related uranium deposits for (A) sandstone-hosted
complex-type deposits and (A, B and C) basement-hosted simple type deposits.
Basement-sourced egress fluids are postulated as the reductant in sandstone-hosted, complex-type
U deposits (Fayek and Kyser, 1997). This two-fluid model suggests that oxidized basinal brines mix
along re-activated basement-faults at the unconformity with egress of basement-sourced, or chemically-
evolved reduced basinal brines that reacted with the graphitic metasedimentary basement rocks (Figure
1.6A; Hoeve and Sibbald, 1978; Kotzer and Kyser, 1995; Fayek and Kyser, 1997). In contrast, ingress
fluids have been proposed for basement-hosted deposits (Figure 1.6A; Hoeve and Quirt, 1984; Fayek and
Kyser, 1997). Here, oxidized basinal brines descended along re-activated faults into the basement and
were reduced through reactions with graphite or Fe-rich metasediments (Alexandre et al., 2005; Fayek
and Kyser, 1997).
Hetch and Cuney, 2000; Derome et al., 2005 Alexandre et al., 2005
Fayek and Kyser, 1997
12
The source of uranium for these high-grade deposits has been a contentious enigma for
developing a broadly excepted genetic model. In both basement and sandstone-hosted U deposits, the
primary source of the U was likely radiogenic S-type granites, pegmatites and metasediments surrounding
the Athabasca Basin (Thomas, 1983; Madore et al., 2000; Mercadier et al., 2013) or the McArthur Basin
(Kyser et al., 2000). However, differing views have emerged on whether the U was leached from U-
bearing detrital minerals (Figure 1.6A, B), such as apatite, monazite, zircon and clays in the basinal
sediments by oxidizing basinal brines (Kyser and Cuney, 2015), or directly from primary monazite and
zircon (Hecht and Cuney, 2000; Derome et al., 2003) or U bearing protores (Mercadier et al., 2013) in the
basement rocks by percolating basinal brines (Figure 1.6C). The most convincing evidence for the
basement-sourced model is the higher concentration of U in the basement rocks, however the immobility
of U under reducing basement environments suggests that basement U-leaching and transport are
improbable (Alexandre et al., 2005).
Possible drivers for basin fluid flow include thermal convection (Hoeve and Sibbald, 1978;
Hoeve and Quirt, 1984; Raffensperger & Garven, 1995; Boiron et al., 2010), compaction driven pore fluid
expulsion (Hiatt and Kyser, 2007), gravity (Alexandre and Kyser, 2012; Derome et al., 2005) and
tectonic-induced fluid flow (Cui et al., 2012). All of the various fluid flow models remain possible,
however numerical modelling favors low fluid overpressure because higher fluid overpressure would
have hindered circulating fluids in the basal sandstone or into the basement (Chi et al., 2013). Transport
of the U-bearing basinal diagenetic brine was stratigraphically focused along basal aquifers that onlap
basement rock units along the eastern margins of the Athabasca Basin (Holk et al., 2003; Hiatt and Kyser,
2007). Cross-formational fluid flow along stratigraphic pathways was likely modified by active fault
zones with the potential to focus diagenetic brines (Kyser et al., 2000). Brittle to semi-brittle reactivated
basement faults provided the pathway for fluid flow into and out of the basement. Numerical fluid flow
modelling suggests ingress and egress mineralization may form along the same fault but are generated at
different stages of deformation (Li et al., 2017). Under a compressional regime, early stage compression
13
with high stress and little displacement, low bulk shortening, is more conducive with egress fluid flow
whereas subsequent high bulk shortening from displacement and dilation favors ingress fluid flow (Li et
al., 2017).
The postulated trapping mechanisms responsible for reducing the U-bearing fluids is graphite
dissolution (Hoeve & Sibbald, 1978; Alexandre et al., 2005) and ferrous iron liberated from mafic
minerals in the basement (Quirt, 1989; Hetch & Cuney, 2000; Derome et al., 2003; Alexandre et al.,
2005; Acevedo and Kyser, 2015).
1.4 Cigar Lake Deposit
The Cigar Lake uranium deposit occurs 410 to 450 m below surface within the Athabasca Basin
along the unconformity between the Helikian Athabasca Group sediments and the underlying Aphebian
graphitic metasediments of the Wollaston Domain (Figure 1.7; Bruneton, 1987; Bishop et al., 2016). The
narrow, flat-lying, cigar-shaped deposit is approximately 1,950 m long, 20 to 100 m wide and has a
maximum thickness of 13.5 m with an average thickness of about 5.4 m (Bishop et al., 2016). As of
December 31, 2015 Cigar Lake has a total reserve of 100,501 tonnes (221.6 M lbs.) U3O8 and a total
resource (measured, indicated and inferred) of 48,412 tonnes (106.7 M lbs.) U3O8 (Bishop et al., 2016).
In the Cigar Lake area, the basin fill is unmetamorphosed quartz arenitic sandstone and
conglomerate of the Manitou Falls Formation (MF). Only the MFd, MFc and MFb are observed proximal
to the deposit (Bruneton, 1987). Basal conglomerates of the MFb are observed locally. The sandstone
units represent a finning upward, transgressive succession.
14
Figure 1.7: Schematic illustration of the Cigar Lake deposit and surrounding alteration. Modified
from Jefferson et al. (2007) and Cameco (2015) with drill core data and field observations.
Regionally the basement rocks to the Cigar Lake deposit are supra-crustal amphibolite facies
metasedimentary gneisses. These metasediments unconformably overly Archean granitoid gneisses and
form a folded NE–SW oriented dome and basin regional geological landscape (Bruneton, 1993).
Bruneton (1993), interpreted the Cigar Lake basin to be a syncline reflected in the axial-planner regional
NE-striking foliation. Cigar Lake occurs within the gradational Wollaston-Mudjatik Domain transition
zone (Bruneton, 1993). Directly underlying the deposit the rocks are moderately graphitic (3-10%),
locally anatectic, cordieritic protomylonitic pelites that have undergone extensive shearing and local
semi-brittle fault reactivation (Bruneton, 1987; Andrade, 2002). The regional foliation in the area strikes
northeast, however the shear zone underlying the deposit is oriented east–west (Bruneton, 1983). The
local, roughly 10 km long, reverse dextral shear zone has been interpreted to be Hudsonian (Bruneton,
1987, 1993; Andrade, 2002). Post peak-metamorphic reactivation occurred in association with retrograde
15
greenschist facies metamorphism and is dated at 1780 Ma (Philippe et al., 1993; Andrade, 2002). Local
foliation concordant discontinuous lenses of amphibole and pyroxene bearing calcic-magnesium rich
gneisses and granulites, occur adjacent to the shear-zone (Bruneton, 1993).
The Cigar Lake deposit is situated directly on top of an unconformity structure-contour high
interpreted as a pre-Athabasca paleo-topographic ridge based primarily on regional drilling that shows the
conglomeratic MFb units wedging out/on-lapping along the flanks of the east-west oriented Cigar Lake
ridge (Bruneton, 1993). Alternatively, post-Athabasca extensional tectonics has been proposed to explain
the coincidental unconformity high (Andrade, 2002; Jefferson et al., 2007).
The Cigar Lake deposit is located within an extensive hydrothermal alteration zone characterized
by interstitial illitization, in contrast to the regional dickite, forming a sub-cropping chimney around the
deposit (Wasyliuk, 2002). Alteration intensifies 100–200 m above the unconformity with intense
pervasive bleaching (Fe removal), local fine-grained sulphidication, silicification, and structurally
controlled quartz dissolution and clay alteration (Bruneton, 1987; Andrade, 2002). Intense structure in the
basal sandstone (~100 m) and sagging sedimentary marker horizons suggest extensive volume loss and
the development of collapse structures from the mineralizing system (Andrade, 2002). Proximal to the
mineralization the clay alteration becomes intense. The orebody is commonly capped by hematite-rich
massive mixtures of illite, muscovite, and kaolinite with local Fe–Mg chlorite (Bruneton, 1987; Percival
and Kodama 1989; Percival et al., 1993; Philippe et al., 1993). Local paragenetically late induration of
clay by calcite and siderite is common (Bruneton, 1987). An extensive argillitized basement alteration
halo extends more than 50 m below the deposit, masking the pre-Athabasca paleoweathering, and consists
of Mg-chlorite (sudoite and chlinochlore) and Mg- and Fe-rich illite (Bruneton, 1987; Percival and
Kodama, 1989). Graphite destruction directly below the deposit is extensive with traces of remobilized
carbonaceous material occurring proximal to the mineralization as irregular aggregates of bituminous
carbon or 1–5 mm black flakes that form hydrocarbon buttons (Bruneton, 1987; Landais et al., 1993).
16
The mineralization at Cigar Lake predominantly contains the uranium oxide and silicate minerals:
uraninite and coffinite (Bruneton, 1987; Janeczek and Ewing, 1992; Reyx and Ruhlmann, 1993; Cramer
and Smellie, 1994). Uraninite forms euhedral, radiating, botryoidal and massive aggregates and occurs in
association with Ni–Co arsenides, sulpharsenides and sulphides (Bruneton, 1987; Reyx and Ruhlmann,
1993; Cramer and Smellie, 1994). Reyx and Ruhlmann (1993) interpreted that the first major stage of
mineralization, responsible for the unconformity-hosted uranium, was a polyphased hydrothermal system
that deposited: U–Ni–Co–As–S–Bi–Cu–Zn and Pb. Two subsequent stages of U crystallization have been
identified reflecting mobilization of Stage 1, the primary mineralizing event. Stage 2 uranium oxide is
associated with secondary Ni–Co arsenides, sulpharsenides and Fe–Cu-rich sulphides (Bruneton, 1987;
Phillipe et al., 1993). The third stage occurs with extensive Fe-oxides replacement and is responsible for
coffinitization and the redistribution of U as perched mineralization (Bruneton, 1987; Reyx and
Ruhlmann, 1993).
The oldest reported age for the major mineralizing event (Stage 1) at Cigar Lake is 1468 Ma, but
this is interpreted to be a minimum age for mineralization (Fayek et al., 2000). Numerous younger ages
have been reported for the deposit (e.g. Cumming and Krstic, 1992; Janeczek and Edwing, 1992; Philippe
et al., 1993; Fayek et al., 1997; Fayek et al., 2002) and Pb-loss has consistently been reported for the
deposit and likely resulted from episodic hydrothermally-enhanced volume diffusion (e.g. Janeczek and
Edwing, 1992; Fayek et al., 1997). Clay mineral dating using K–Ar has yielded similar results for illite
(1255–1148 Ma) and sudoite (850 Ma) due to the episodic hydrothermal fluids that have accessed the
deposit along structures (Percival et al., 1993).
1.5 Geometallurgy
Geometallurgy is the application of mineralogy, geology, and material characterization for
predictive metallurgy during mineral processing (Bowell et al., 2011). Recognition of the underlying
mineralogical and geological controls on metallurgical performance and integration with empirical
17
geospatial characterization can provide predictive support, reducing capital expenditures and operational
disturbances in mining, milling and mine tailings management (e.g. Pownceby and Johnson, 2014;
Adams, 2007).
The performance and operational reliability of mining and milling operations is intrinsically
dependent on the underlying mineralogy and geology of the orebody (Adams, 2007). Maintaining mining
and milling throughput in geochemically and mineralogical variable orebodies can be challenging due to
grade variability, EOC concentration spikes and metallurgical performance (e.g. reagent consumption,
clay settling efficiency). Plant designs can be engineered to account for variable orebodies, however these
capital expenditures can significantly reduce the internal rate of return of the operation. Mitigation
strategies such as production and mill feed blending employed to reduce ore grade, EOC, or gangue
mineral variability is the preferred strategy for mine and mill process optimization without significant
capital. The integrated approach of material characterization and geological modelling in geometallurgy
can provide an important framework for processing capricious ores.
1.6 Thesis Objectives and Rationale
Ongoing delineation and operational drilling of the Phase 1 pods has improved the spatial
coverage of the orebody allowing for renewed insight into the deposit and ore forming system. Since the
initial discovery of the Cigar Lake orebody in 1981, many geological, mineralogical, geochemical and
geochronological aspects of the deposit have been studied in detail (e.g. Bruneton, 1987, 1993; Percival
and Kodama, 1989; Cumming and Krstic, 1992; Landais et al., 1993; Pacquet and Weber, 1993; Pagel et
al., 1993; Percival et al., 1993; Philippe et al., 1993, 2002; Reyx and Rulmann, 1993; Toulhoat and
Beaucaire, 1993; Cramer and Smellie, 1994, 1995; Janeczek and Ewing, 1992, 1994; Cramer, 1995;
Mosser et al., 1996; Fayek and Kyser, 1993, 1997; Fayek et al., 1997, 2000, 2002). However, research
focused directly on the uranium ore and associated metals, and the underlying paragenetic model is sparse
(e.g. Bruneton, 1987; Reyx and Rulmann, 1993), with studies typically lacking access to high-grade
18
uranium ore samples. Quality polymetallic samples are particularly challenging to obtain due to the
heterogeneity of the ore and high clay content that masks the sulphides and arsenides in drill core.
The overall objective of this research is to conduct geometallurgical characterization in support of
geologic modelling that provides predictive ore characteristics for mining and milling of the Phase 1
Cigar Lake deposit (Figure 1.2). Emphasis is placed on minerals containing EOC for mining and milling
operations. Four specific targets are established:
Determine the mineral phases controlling the geochemistry of select uranium and metal-
bearing ores across the Cigar Lake Phase 1 deposit and provide semi-quantitative
mineral proportions.
Design a normative algorithm to calculate inferred mineral proportions using mineral
stoichiometry and whole-rock geochemistry.
Access the empirical spatial distribution of EOCs and host mineral phases and determine
structural or geochemical controls.
Develop a genetic model for the Cigar Lake ore incorporating the paragenesis,
geochronology and stable isotopes.
This research provides Cameco with more predictive capabilities for mine planning and
scheduling of Jet Boring System (JBS) cavities potentially reducing grade variability, limiting EOC
concentration spikes for mine water effluent control and improving metallurgical performance. This
project supports milling operations with improved characterization of the mill feed. Identifying the
mineralogical structure of elements provides insight into slurry performance, including reagent
consumption during acid leaching and thermodynamic properties during oxidization. An important
environmental concern in the uranium mining industry is the potential for long-term mobilization of
EOCs (Ni, Co, As, Se, and Mo) from tailings deposited in in-pit tailings management facilities (TMFs)
into the regional groundwater systems. Understanding the mineralogical structure of the ore and waste
19
rock will enhance stakeholder’s ability to control leaching and assess environmental liability from waste
stockpiles and tailings.
1.7 Structure of Thesis
The results from this thesis are presented as two manuscripts, Chapters 2 and 3. Chapter 4 is a
general discussion of the conclusions derived from the thesis and the implication for mining, milling and
exploration. Areas for further research are highlighted. The general outline for the chapters is as follows:
Chapter 2: Cigar Lake: Geometallurgical ore characterization in support of mining and milling
By: A. Kaczowka, T. Kotzer, K. Kyser and C. Revering
Outline: Geochemical, mineralogical, geological and geospatial characteristics of the Cigar Lake orebody
are integrated to provide a geometallurgical, or geologically predictive overview to guide uranium mining
and milling. The mineralogy of the uranium ores is characterized using semi-quantitative techniques
including XRD, SEM, mineral liberation analysis (MLA), shortwave infrared spectroscopy (SWIR), and
optical petrography. Electron microprobe analysis (EMPA) and laser ablation inductively coupled plasma
mass spectrometry (LA-ICP-MS) are used to measure the chemical compositions and element deportment
of selected minerals. A normative algorithm is designed to calculate mineral proportions using mineral-
stoichiometry and whole-rock geochemistry. The quantitative mineralogy obtained is used assess the
empirical spatial distributing of mineral phases and their corresponding EOCs to identify the
mineralogical zonation within the ore body.
Chapter 3: Evolution of the high-grade polymetallic unconformity-related uranium Cigar Lake ore body
By: A. Kaczowka, K. Kyser, T. Kotzer and C. Revering
Outline: A study of the mineralogy, geochemistry, and geochronology of the Cigar Lake orebody with a
focus on defining the paragenesis and understanding ore forming processes. The mineralogy and
20
paragenesis of the uranium ores are characterized using XRD, SEM, MLA, SWIR, and optical
petrography. The empirical spatial distribution of elements and minerals is interpreted from a geological
perspective to develop a genetic model for the deposit. EMPA and LA-ICP-MS are used to measure the
chemical compositions, U/Pb and Pb-isotope ratios of selected U-bearing, arsenide and sulphide mineral
phases to establish element deportment through paragenesis and absolute ages. Stable isotope
concentrations have been measured on the U-bearing, sulphide and C-bearing phases to confirm
paragenetic observations and to ascertain the age, source and evolution of fluids that formed and altered
the Cigar Lake deposit.
21
Chapter 2
CIGAR LAKE: GEOMETALLURGICAL ORE CHARACTERIZATION IN
SUPPORT OF MINING AND MILLING
2.1 Abstract
Cigar Lake is a polymetallic, unconformity-related uranium deposit with complex geochemistry
and mineralogy. Variable concentrations and spatial distributions of elements of concern (EOC) such as
As, Mo, Ni, Co, Se and Zr associated with the high-grade and tonnage tetravalent uranium ores (UO2+x;
U(SiO4)1-x(OH)4x) present unique mining, metallurgical and environmental challenges. Sulphides and
arsenides have significant As, Mo, Ni, Co and Se control and have properties that affect EOC mobility,
thus requiring consideration during mineral processing, mine-effluent water treatment and long-term
tailings management. Here geochemical, mineralogical, geological and geospatial characteristics of the
Cigar Lake orebody are integrated to provide a geometallurgical, or geologically predictive, overview to
guide uranium mining and milling.
The mineralogy of the uranium ores from Cigar Lake were characterized using semi-quantitative
to quantitative techniques including X-ray diffraction (XRD), scanning electron microscopy (SEM),
mineral liberation analysis (MLA), shortwave infrared spectroscopy (SWIR), and optical petrography.
Electron microprobe analysis (EMPA) and laser ablation inductively coupled plasma mass spectrometry
(LA-ICP-MS) were used to determine the chemical compositions and element deportment of selected
minerals. Integration of mineralogy, stoichiometry and bulk-rock geochemistry was used to develop
normative algorithms for quantification of the metal-bearing minerals. Geostatistical implicit modelling
was undertaken to spatially delineate the distribution of EOCs and derived mineral proportions.
The U-bearing (uraninite, coffinite) and metallic arsenide (niccolite), sulpharsenide (gersdorffite,
cobaltite) and sulphide (chalcopyrite, pyrite, galena, bornite, chalcocite, sphalerite, pyrrhotite) minerals
provide the main EOC control. Arsenic, Ni, and Co occur primarily in a reduced state as 1:1 molar ratio,
22
Ni–Co:As arsenides and sulpharsenides such as gersdorffite, niccolite and cobaltite. Molybdenum occurs
within molybdenite and uraninite. Selenite occurs within sulphides, sulpharsenides and is co-located
within coffinite. Zirconium is found within detrital zircon crystals and within coffinite.
The spatial distribution and paragenesis of U-, As- and S-bearing minerals are a result of the
elemental composition, pH and redox conditions of early formational and later meteoric fluids that have
accessed the deposit along lithostratigraphic permeability and tectonic structures. Using the holistic
geometallurgical paradigm presented here, the geology at Cigar Lake can be used to optimize and reduce
risk during long-term mine and mill planning.
23
2.2 Introduction
Unconformity-related uranium deposits in the Athabasca Basin of northern Saskatchewan,
Canada are unrivalled as the highest-grade uranium deposits in the world (IAEA, 2009). Uranium mines
in Saskatchewan produce approximately 16% of total current global production of U3O8 (World-
nuclearorg, 2016). Cigar Lake, with the highest mined uranium grades in the world, is the newest
operation in the mining district with first ore production beginning in 2014 and achieving commercial
production status in 2015 (Cameco, 2015) (Figure 2.1).
The ore at Cigar Lake is highly variable with a complex polymetallic geochemistry containing
elevated concentrations of As, Co, Cu, Mo, Ni, Se, and Zr (Bruneton, 1987; Reyx and Ruhlmann, 1993).
Elements such as As, Ni, Co, Mo, and Se, which can be problematic during mining, milling and tailings
management, have been identified as elements of concern (EOC; Bishop et al., 2016). Overall, the
minerals and mineraloid phases represent significant elemental control and have properties that affect
mineral processing and mobility of EOCs in process waters and long-term tailings management facilities.
Maintaining mining and milling throughput in geochemically and mineralogical variable orebodies can be
challenging due to grade variability, EOC concentration spikes and metallurgical performance (e.g.
reagent consumption, clay settling efficiency). Plant designs can be engineered to account for variable
orebodies, however these capital expenditures can significantly reduce the internal rate of return for the
operation. Mitigation strategies such as production and mill feed blending, employed to reduce ore grade,
EOC, or gangue mineral variability, is the preferred strategy for optimization without significant capital.
Geometallurgy is the application of mineralogy, geology, and material characterization for
predictive metallurgy during mineral processing (Bowell et al., 2011). Recognition of the underlying
mineralogical and geological controls on metallurgical performance and integration with empirical
geospatial characterization can provide predictive support reducing capital expenditures and operational
disturbances in mining, milling and mine tailings management (e.g. Pownceby and Johnson, 2014;
24
Adams, 2007). This integrated approach of material characterization and geological modelling in
geometallurgy can provide an important framework for processing capricious ores.
At Cigar Lake, we used the integration of geochemical, mineralogical, geological and geospatial
characterization of the high-grade U and Ni–Co–As–S ores to support the mining and milling of the Phase
1 Cigar Lake pods (Figure 2.2). This holistic characterization program supports current geologic
modelling by providing predictive ore characteristics for mining and milling.
Figure 2.1: Location of the Athabasca Basin and the Cigar Lake deposit (yellow star) with the
underlying regional geologic provinces. Also shown are the locations of several other high-grade
unconformity-type uranium deposits (black squares) and northern communities (white circles).
Regional geology is modified from Card et al. (2007) and Ramaekers et al. (2007).
25
Figure 2.2: Air photograph of the Cigar Lake mine site with outlined study area, the Phase 1
Cigar Lake ore body. The Phase 1 deposit is divided into the East Pod and the West Pod. Mineral
reserves and resources are effective as of December 31, 2015 as reported by Bishop et al. (2016).
2.2.1 Geological Setting
The Cigar Lake uranium deposit occurs within the intracratonic Paleo to Mesoproterozoic
Athabasca Basin, which unconformably overlies the remnants of two orogenic belts, the Taltson
Magmatic Zone to the West and the Trans-Hudson to the East (Ramaekers, 1980). The initial
accommodation for the basal Athabasca Group (Manitou Falls and Fair Point Formations) occurred in
NE-SW trending Hudsonian (1.7 Ga) basement faults (Armstrong and Ramaekers, 1985; Kyser et al.,
2000). Rapid uplift during the Trans-Hudson Orogeny provided the siliciclastic input for the Athabasca
Basin with basin fill comprising unmetamorphosed quartz arenitic sandstone and conglomerate overlain
by siltstone, mudstone and dolostone (Ramaekers, 1990). The depositionional environment of the
upward-fining, red bed succession is interpreted as major fluviatile and near-shore shallow marine
environments (Ramaekers, 1990). Within the Cigar Lake region, the crystalline basement is comprised of
granites and granitoid gneisses unconformably overlain and folded with upper amphibolite facies
metasedimentary gneisses (Tran and Smith, 1999; Card et al., 2007).
26
2.2.2 The Cigar Lake Deposit
The Cigar Lake uranium deposit occurs 410 to 450 m below surface along the unconformity
between the underlying crystalline basement rocks and the Athabasca Group sediments (Figure 2.3;
Bruneton, 1997). The narrow, flat-lying, cigar-shaped deposit is approximately 1,950 m long, 20 to 100 m
wide and has a maximum thickness of 13.5 m with an average thickness of approximately 5.4 m (Bishop
et al., 2016). Local basement-hosted mineralization and perched-mineralization occur but are lower in
grade and spatially confined to structures resulting in limited mining potential (Bishop et al., 2016). As of
December 31, 2015, Cigar Lake has a total reserve of 100,501 tonnes (221.6 M lbs.) U3O8 and a total
resource (measured, indicated and inferred) of 48,412 tonnes (106.7 M lbs.) U3O8 (Figure 2.2; Bishop et
al., 2016).
At Cigar Lake, the basin fill is unmetamorphosed quartz arenitic sandstone and conglomerate of
the Manitou Falls Formation. Directly underlying the deposit the rocks are moderately graphitic (3–10%),
cordieritic pelites (Bruneton, 1987; Andrade, 2002). The regional foliation in the area strikes northeast,
however the shear zone underlying the deposit is oriented east-west (Bruneton, 1993). The Cigar Lake
deposit is situated directly on top of an unconformity structure-contour high interpreted as pre-Athabasca
paleo-topography (Bruneton, 1993).
The Cigar Lake deposit is located within an extensive hydrothermal alteration zone characterized
by interstitial illitization, in contrast to the regional dickite, forming a sub-cropping chimney around the
deposit (Wasyliuk, 2002). Alteration intensifies 100–200 m above the unconformity with intense
pervasive bleaching (Fe removal), local fine-grained sulphidication, slilicification, and structurally
controlled quartz dissolution and clay alteration (Figure 2.3; Bruneton, 1987; Andrade, 2002). Proximal to
the mineralization the clay alteration becomes intense around the periphery of the deposit (Figure 2.4).
The orebody is commonly capped by hematite-rich massive mixtures of illite, muscovite, and kaolinite
with local Fe–Mg chlorite (Bruneton, 1987; Percival and Kodama, 1989; Philippe et al., 1993). Local
induration of clay by calcite and siderite is common (Bruneton, 1987). An extensive argillitized basement
alteration halo consisting of Mg-chlorite (sudoite and clinochlore) and Mg- and Fe-rich illite extends
27
more than 50 m below the deposit, masking pre-Athabasca paleoweathering (Bruneton, 1987; Percival
and Kodama, 1989). Graphite destruction directly below the deposit is extensive with traces of
remobilized carbonaceous material occurring proximal to the mineralization as irregular aggregates of
bituminous carbon or 1–5 mm hydrocarbon buttons (Bruneton, 1987; Landais et al., 1993).
The uranium mineralization is characterized predominantly by uraninite and coffinite (Bruneton,
1987; Janeczek and Ewing, 1992; Reyx and Ruhlmann, 1993; Cramer and Smellie, 1994). Uranium oxide
forms euhedral, radiating, botryoidal and massive aggregates and occurs in association with Ni–Co
arsenides, sulpharsenides and sulphides (Bruneton, 1987; Reyx and Ruhlmann, 1993; Cramer and
Smellie, 1994).
Figure 2.3: Schematic illustration of the Cigar Lake deposit and surrounding alteration. Modified
from Jefferson et al. (2007) and Cameco (2015) with drill core data and field observations.
28
Figure 2.4: East Pod section along line 10898 (mine grid) showing orebody facies and structural
interpretation. Orebody outline at 1% U3O8 cutoff highlighted in red. Representative whole-rock
geochemical samples illustrating the variable chemistry of the orebody.
Whole-Rock ICP-OES Geohemistry
SAMPLEID HOLEID FROM TO U3O8 S* As Ni Co Mo Se Cu Pb Zn Al2O3 MgO K2O Fe2O3 CaO
(%) (%) (%) (%) (%) (%) (ppm) (%) (%) (%) (%) (%) (%) (%) (%)
1 CAM052883 SF898_14 433 433.1 1.77 13.4 26.3 21.3 1.06 0.31 57 0.20 0.25 0.072 6.60 3.41 0.41 5.25 0.19
2 CAM052877 SF898_14 430.9 431.3 19.8 8.06 2.05 1.63 0.32 1.80 182 5.96 3.20 0.041 12.0 4.05 1.38 11.0 0.62
3 CAM052797 SF898_12 426 426.4 1.84 0.48 0.14 0.04 0.02 <DL <DL 0.01 0.06 0.010 28.8 2.30 8.20 6.44 0.23
4 CAM083166 SF898_12 430.5 430.9 35.2 10.2 18.8 14.9 2.45 0.16 301 1.20 3.08 0.047 2.00 0.73 0.18 3.32 0.73
5 CAM052852 SF898_10 424.4 424.9 1.27 0.35 0.02 0.04 <DL <DL <DL 0.02 0.09 0.004 15.8 4.56 3.81 30.8 0.53
6 CAM052859 SF898_10 427.5 427.9 26.9 5.62 2.34 1.82 0.16 1.05 143 3.98 2.89 0.007 10.5 4.89 0.19 14.6 0.91
7 CAM052862 SF898_10 428.5 429 70.5 3.56 4.48 3.37 0.40 0.06 178 1.58 6.28 0.087 0.66 0.46 0.01 2.76 1.60
8 CAM052867 SF898_10 430.9 431 8.99 8.86 16.0 13.2 1.07 0.80 59 0.19 0.78 0.066 10.6 5.38 0.17 4.64 0.48
9 CAM052764 SF898_08 429.7 429.8 44.3 4.16 0.16 0.14 <DL 1.97 1590 3.59 6.72 0.009 9.76 0.58 1.46 4.21 0.82
10 CAM052766 SF898_08 430.2 430.6 76.4 1.84 0.28 0.17 0.07 0.15 568 1.77 8.65 0.003 2.79 0.09 0.23 2.33 1.25
11 CAM083152 SF898_04 437.1 437.6 1.16 1.52 0.04 0.01 <DL <DL <DL <DL 0.039 <DL 25.7 1.88 5.23 5.65 0.30
<DL denotes below lower detection limit
*Analyzed with leco induction furnace
29
2.2.3 Geometallurgical Considerations at Cigar Lake
In northern Saskatchewan, the polymetallic unconformity-related uranium deposits contain
elevated levels of As, Ni and Co within minerals intergrown with the uranium ores. During milling,
arsenides and sulpharsenides consume reagent and produce heat during oxidization (Pankratz et al., 1984;
Reimers and Hjelmstad, 1987; Wang, 2007; Pownceby and Johnson, 2014). Furthermore, discharged
hydrometallurgical waste solutions (neutralized raffinate) from the milling process and attendant leached
residues are ultimately treated and combined into mill tailings that contain elevated concentrations of Ni
and As, in both neo-formed precipitates and complexes as well as unreacted primary arsenide and
sulphide minerals. A concern in the uranium mining industry is the possibility of long-term mobilization
of As and Ni from tailings management facilities (TMFs) into the surrounding regional groundwater
systems. At uranium milling operations in the Athabasca Basin, the behavior of these elements within the
milling and tailings processes have been, and are, routinely monitored and studied (e.g. Donahue et al.,
2000; Cutler et al., 2003; Essilfie-Deughan et al., 2012; Bissonnette, 2015). As such, the geometallurgy of
these minerals has been studied here to understand their compositional and spatial variabilities within the
Cigar Lake orebody.
In addition to Ni, Co and As, tailing and metallurgical solutions from uranium mines in northern
Saskatchewan can contain elevated concentrations of Mo and Se. Mineralogically, Mo is not found in its
pure metallic state but occurs in various mineral forms, primarily as molybdenite (MoS2) and jordisite
(amorphous MoS2), commonly in association with U, V, As, or Cu (Heinrich et al., 2010). Selenium
occurs in reduced form substituting for sulfur in sulphides as well as within selenite and selenate
complexes (e.g. Blaise and Koning, 1985). As with As and Ni, the potential for long-term mobilization of
Mo and Se from the tailings to regional groundwater systems is an important environmental concern and
these elements are monitored relative to water quality guidelines and studied routinely within the tailings
mass (e.g. Shaw et al., 2011; Essilfie-Dughan, 2011; Bissonnette, 2015) and surround aquatic
environments (e.g. Muscatello and Janz, 2009; Wiramanaden et al., 2010).
30
Ziconium is not an environmental EOC, however it is problematic for uranium processing and is
therefore included in the study accordingly. Zirconium reduces the quality of yellow cake and is notorious
for creating solvent extraction problems during the preparation of UF6 (IAEA, 1980). As such, Zr is a
penalty element during the sale of yellow cake if concentrations exceed contract limits. Within this study,
the mode of occurrence and spatial distribution of Zr has been detailed to provide the appropriate
reporting to the uranium extraction circuits.
The Cigar Lake ores contain relatively high clay contents, consisting primarily of illite,
muscovite, kaolinite and chlorite (Bruneton, 1987; Percival and Kodama, 1989; Percival et al., 1993;
Philippe et al., 1993). Metallurgically, process waters and slurries containing clays require prolonged
solid-liquid separation time due to complex colloid-liquid interactions. During mining, the Cigar Lake
operation uses a novel, water intensive, Jet Boring System (JBS) to extract the ore with high pressured
water while process waters hydraulically transport the ore through the slurry comminution circuit
(Cameco, 2015). The relative amount and spatial distributions of the various clay minerals can result in
varied settling rates in clarifiers used to recycle the process waters. Furthermore, during milling,
discharged neutralized raffinate and leached residues form slurries that are deposited sub-aqueously
within contained tailings management facilities. Volume loss achieved through slurry dewatering is
partially dependent on the clay content and mineralogy of the ore (Ito and Azam, 2017). Within
geometallurgy, defining the mineralogy and overall extent of clay alteration associated with the ore is
critical for process efficiency and appropriate design.
2.3 Methods
This study incorporates various analytical techniques on select uranium and metal-bearing ores
across the Cigar Lake Phase 1 pods to provide mineralogical and geochemical characterization.
Integration of mineral identification and stoichiometry (X-ray diffraction, XRD; scanning electron
microscopy-mineral liberation analysis, SEM/MLA; and electron microprobe analysis, EMPA – Figures
2.5 and 2.6) with bulk-rock geochemistry (Whole Rock inductively coupled plasma mass spectrometry,
31
ICP-MS; Leco; and Titration – Figure 2.5) were used to quantify mineral proportions. A normative
algorithm was created to provide predictive quantification of the dominant sulphide, sulpharsenide and
arsenide minerals identified within the Phase 1 pods: sphalerite, gersdorffite, niccolite, rammelsbergite,
chalcopyrite, bornite, chalcocite, galena and pyrite/pyrrhotite (Appendix H).
Several geochemical stoichiometric techniques are utilized in the calculations. Molar element
ratios are used to differentiate element control by mineral phases. Minerals with element-constrained
ratios, exhibiting the sole control over an element (e.g. Zn in sphalerite) or with more than one element-
ratio (e.g. Ni:As and Ni:S in gersdorffite NiAsS) are calculated first in the linear algorithm. A subtractive
method, of calculating the element consumption by element-constrained minerals before calculating the
concentration of non-element constrained minerals is used to help differentiate between mineral phases.
The exact mathematical steps used to quantify mineral proportion is provided in Appendix G. An
overview of the normative algorithm is provided in Figure 2.7.
Prior to the study, Cameco Corporation (Cameco), 50% owner and operator of the Cigar Lake
mine, possessed an extensive multi-element whole-rock geochemical dataset containing over 10,000
samples (Whole Rock ICP-MS – Figure 2.5; Appendix A). To assist in mineral-stoichiometry
geochemical quantification of mineral proportions, pulp material from 3,527 spatially representative
samples were analyzed by the Saskatchewan Research Council (SRC) with a Leco induction furnace for
S% and C% and titration was used for FeO wt. % on 53 select samples (Figure 2.5; Appendix B).
The mineralogy of the ores was initially characterized by analyzing 53 spatially representative
samples with XRD (Figures 2.5 and 2.6; Appendix B, D). The analysis was undertaken at Queen’s
University with an Xpert Pro Philips powder diffractometer equipped with a cobalt X-ray tube and an
X’celerator area detector. The X-ray beam was in Bragg-Brentano configuration. To minimize the effects
of preferred mineral orientation, samples were loaded into jacket-style holders and spun during the
procedure. Mineral identification was performed by pattern-matching using PANalytical HighScore
32
software. Semi-quantitative percentages were determined using the reference intensity ratio method (RIR;
Hubbard and Snyder, 1988).
Petrographic analyses (reflected and transmitted light microscopy) were performed on a select
suite of samples collected from active on-going drilling to enhance the coverage of mineralogical data
(Figures 2.5 and 2.6). Petrographic sections provided in-situ mineral relationships and textures to assess
the relative sequence and timing of minerals during formation of the Cigar Lake ore deposit.
XRD (25) and petrographic samples (12) with mineral phases of interest were scanned with a
MLA-equipped SEM for further mineral confirmation, to improve semi-quantitative mineral proportions
and for mineral textures (Figure 2.5; Appendix E). Material was mounted into epoxy, polished, carbon
coated and scanned using a MLA 650 FEG ESEM at Queen’s University. Back-scattered electron (BSE)
images and energy-dispersive (EDS) spectra facilitated mineral identification and were used to establish a
customized Cigar Lake EDS mineral library for MLA.
A subset of representative samples was analyzed by EMPA to measure the chemical
compositions and element deportment (Figure 2.5, Appendix C). The EMPA work was undertaken on
select minerals including uraninite, coffinite, gersdorffite and niccolite using a JEOL JXA-8230 equipped
with five wavelength dispersive spectrometers (WDS). Uraninite was analyzed using 15 kV accelerating
potential, 100 nA beam current and a 7 μm beam diameter. Acquisitions of coffinite were acquired with a
15 kV accelerating potential, 10 nA beam current and a 3.5–7 μm beam diameter. Arsenides were
analyzed using a 20 kV accelerating potential, 30 nA beam current with a focused beam.
A Thermo Scientific ELEMENT XR LA-ICP-MS was used to analyze major elements on
sphalerite (Appendix F). LA-ICP-MS element concentrations were quantified using external glass
standards NIST610, NIST612 and an in-house galena calibrated to NIST610. A sample set started with
the NIST glasses and calibrated galena, followed by ten sample analyses. Internal standardization was
performed by normalizing measured intensities to an idealized chemical formula (Appendix F).
33
XRD SEM/MLAPetrographic
SectionsSWIR EMPA LA-ICP-MS
Whole Rock
ICP-MS
Leco (S%/C%/OC%)
Titration (FeO%)
53 Samples
for RIR semi-
quantitative
mineralogy
25 Epoxy
Grain
Mounts 12
Petrographic
Sections for
quantitative
mineralogy
and mineral
textures
50
Petrographic
Sections for
mineralogy
and textural
interpretation
53 Pulp
Samples
>200 Core
Samples for
clay
mineralogy
Select
Samples on
Uraninite
Coffinite and
Ni-arsenides
for detailed
mineral
geochemistry
Select
Samples on
sphalerite for
detailed
mineral
geochemistry
> 10,000
Sample
historic multi-
element whole
rock dataset
3527 S%
Samples
53 C%/OC%
Samples for
normative
mineral
quantification
53 Samples
for normative
mineral
quantification
Perdictive Mineralogical Normative
Leap Frog 3D Implicit Model
Mineralogical Characterization Geochemical Characterization
Cigar Lake Ore Characterization
Leapfrog Geo software was utilized to determine the empirical spatial distribution of EOCs: As,
Co, Mo, Ni, Se, Zr and normative mineral proportions. Geostatistics based variography was utilized to
identify geochemical domaining, and structural controls.
Figure 2.5: Cigar Lake mineralogical and geochemical ore characterization work-flow diagram
(Acronyms: XRD- X-ray diffraction, SEM- Scanning electron microscope, MLA- mineral liberation
analysis, SWIR- short-wave near-infrared spectral device, EMPA- electron microprobe analysis,
LA-ICP-MS- laser ablation inductively coupled plasma mass spectrometry).
Figure 2.6: Sample location map of mineralogical work. Each symbol typically represents two to
three samples through the deposit.
34
2.3.3 Normative Error, Probability and Limitations
The normative algorithm intrinsically contains errors and an attempt has been made to quantify
these to provide a realistic expectation regarding accuracy and precision. Cigar Lake samples proved
difficult to quantify with high clay content and poorly crystallized phases making most samples
unsuitable for Rietveld refinement whereas the RIR method provided only semi-quantitative
concentrations. Therefore, the normative mineral proportions have been compared here against MLA
results, assuming the MLA results represent a fairly well-constrained standard analysis. The comparison
with the normative algorithm yielded an overall R2 of 0.87 (n = 145) and suggests the accuracy of the
normative algorithm. The average relative error for the normative algorithm is 30% (Figure 2.8). At low
mineral concentrations, less than 4 wt. %, the relative error becomes more variable reflecting the
influence of element substitutions, mineral exsolution, clay absorption and poorly crystallized mineraloid
phases. On Figure 2.8, the highest error shown, 246 % relative error for chalcocite/bornite, represents a
MLA measured concentration of 1.7 wt. % and a predicted concentration of 5.7 wt. %. Chalcopyrite and
bornite are prone to exsolution textures and for mining and milling differentiating the Cu-phase is of
lesser importance and this degree of relative error is still considered acceptable.
35
Figure 2.7: Flow chart demonstrating the steps taken to calculate normative mineralogy from the
whole-rock geochemistry: 1) The normative calculates the sphalerite concentration using Zn wt. %
and the determined sphalerite formula: (Zn0.96Fe0.04)S. 2) The arsenide and sulpharsenides are
calculated using the Ni+Co M%/As M% ratio. 3) The remaining S is used to discern the Cu-phase
minerals using the Cu M%/S M% ratio. 4) Galena is calculated using the remaining S and available
Pb. 5) Pyrite/pyrrhotite are calculated with the remaining S.
36
Figure 2.8: Relative error (%) is calculated for the normative mineral proportions relative to MLA
quantification. Relative error (%) is plotted against MLA wt. %. Below 4 wt. % the relative error
increases. The average relative error is 30% in samples with >4 wt. % of a mineral. Note that the
highest relative error shown, 246% for bornite/chalcopyrite represents a predicted 5.7 wt. % vs
MLA measured 1.7 wt. %.
Due to the linear design, the error in quantification increases sequentially through the normative
algorithm. For mining and milling, accurate reporting of arsenides and sulpharsenides is more critical than
the subsequently calculated sulphides and the normative algorithm has been designed accordingly. Figure
2.9 demonstrates that sphalerite (R2 = 0.99), gersdorffite (R2 = 0.91), cobaltite (R2 = 0.92),
rammelsbergite (R2 = 0.94), and niccolite (R2 = 0.80) have a better correlation and less relative error than
chalcopyrite (R2 = 0.74), bornite/chalcocite (R2 = 0.78), pyrite (R2 = 0.61), and galena (R2 = 0.35) due to
the sequential design of the algorithm.
The success of the normative algorithm is fundamentally dependent on the consistent mineralogy
identified throughout the Phase 1 pods. The normative algorithm is dependent on the molar element ratios
of the identified mineral phases. Therefore, a major shift in mineralogy, or significant element
substitutions, would require refinement of the algorithm.
37
R2 = 0.99
n = 6
R2 = 0.91
n = 22
R2 = 0.92
n = 16
R2 = 0.94
n = 8
R2 = 0.80
n = 15
R2 = 0.74
n = 22
R2 = 0.78
n = 14
R2=0.35
n = 22
R2=0.61
n = 20
38
Figure 2.9: Correlation between normative mineral proportions (wt. %) and MLA (wt. %) results.
Minerals calculated early in the normative algorithm (sphalerite, gersdorffite, rammelsbergite,
niccolite) have a better correlation and higher R2 values than minerals calculated in the later
calculations (chalcopyrite, bornite, chalcocite, galena, pyrite). All R2 values are reported to the 100th
confidence interval.
2.4 Results
2.4.1 Mineralogy
The results represent the culmination of XRD, MLA and petrographic interpretation used to
identify and quantify mineralogy and ascertain textural relationships. Here, minerals with EOCs were
targeted and incorporated into a paragenesis for the Cigar Lake deposit. The 3 stages of uranium
crystallization and alteration previously identified at Cigar Lake (e.g. Bruneton, 1987; Reyx and
Ruhlmann, 1993; Fayek and Kyser, 1993) have been expanded to 4 stages to reflect mineralogical,
textural and chemical changes in the orebody through the evolution (Figure 2.10).
2.4.1.1 Uranium Ore Mineralogy
At Cigar Lake, uranium occurs primarily as reduced tetravalent (IV) oxides and silicate minerals.
Uraninite (UO2) is the dominant uranium oxide mineral in high-grade (>50% U3O8) zones occurring
primarily as botryoidal masses, and to a lesser extent massive aggregates, veins and disseminated
subhedral crystals. The dominant millimeter to centimeter-scaled botryoids coalesce to form radiating
globular aggregates. Uniformly distributed radial and polygonal shrinkage cracks, resembling desiccation
cracks, occur within the primary uraninite crystals suggesting the crystals underwent dehydration during
precipitation of uraniferous gels (Figure 2.11A, C; Figure 2.14A). Primary uraninite (U1) is typically
overgrown and intergrown by sulphides and arsenides within a chlorite, illite matrix. Some U1 crystals
are brecciated and overgrown by coeval Cu-sulphides and Ni-sulpharsenides reflecting syngenetic
faulting (Figure 2.11B).
39
Ore StageMinerals Stage 1 Stage 2 Stage 3 Stage 4
UraniniteCoffiniteBoltwooditeUranophaneNiccoliteRammelsbergiteSkutteruditeGersdorffiteCobaltiteGlaucodotSe, Bi SulpharsenidesBravoiteChalcopyritePyrite/MarcasitePyrrhotiteBornite ChalcociteSphaleriteGalenaErytheriteAnnabergiteAerugiteQuartzIlliteSideriteCalciteKaoliniteChloriteLimoniteHematiteRutileHydrocarbons
Gan
gue
Post-Ore Alteration
Ura
niu
m
min
eral
s
Ars
enid
e an
d
Sulp
har
sen
ide
Sulp
hid
eA
rsen
ate
U1 U2 U3 U4 U5 U6
CA1 CA2
CPY1 CPY2 CPY3
GER1 GER2
PY1 PY2 PY3 PY4
SPH
CPY4
Figure 2.10: Mineral paragenesis for the Cigar Lake ore body. Red U1 denotes primary
mineralization whereas blue (U2–U6) indicates predominantly alteration and Pb-loss rather than
complete recrystallization.
Under BSE, even pristine botryoidal or subhedral U1 crystals display some grey-scale mottling
indicative of chemical heterogeneity and alteration. Alteration of the initially emplaced U1 was
substantial during the subsequent stages of the paragenesis (Stages 2–4). U1 crystals were strongly and
almost ubiquitously altered resulting in enhanced grey-scale mottling, observable under BSE. Uraninite
dissolution and alteration results in irregular embayed crystal boundaries (Figure 2.11C). Structure ranges
from microfracturing to local cataclastic brecciation responsible for the fragmentation of U1 crystals.
40
100 µm
UR
COFF
A B
C D
E F
GER
GER
CLCY
41
Figure 2.11: BSE images of uraninite (UR) and coffinite (COFF). A) Botryoidal uraninite crystals
with symmetrical shrinkage cracks. Desiccation cracks are filled with chalcopyrite (CPY) and
gersdorffite (GER). Sample has a chlorite clay (CLCY) matrix. B) Brecciated uraninite with
hairline microfractures of coffinite. Fragments are filled with gersdorffite. Gersdorffite is
overgrown with subsequent generation of coffinite. C) Altered uraninite with ribbon texture and
embayed crystal boundaries. D) Uraninite showing extensive coffinitization. E) Uraninite with
direct coffinitization but also primary remobilization and subsequent recrystallization of coffinite
crosscutting chalcopyrite. F) Uraninite with coffinitization and crosscutting uranophane.
Highly altered uraninite crystals occur as remnant irregular bands with a ribbon-like texture (Figure 2.11
C). In less altered samples, vugs occur between uraninite crystals indicating dissolution and chemical
buffering from sulphides, sulpharsenides and arsenides.
Coffinite (U(SiO4)1-x(OH)4x), a tetravalent uranium silicate, is prevalent throughout the orebody.
Coffinite typically forms irregular anhedral crystal aggregates, sooty disseminations and feathery, slightly
fibrous masses. Direct coffinitization of the uraninite is prevalent along uraninite microfractures, around
crystal boundaries or as complete replacement (Figure 2.11D). Coffinite occurs with sulphide,
sulpharsenide and arsenide overgrowths resulting from the remobilization of U initially emplaced as
uraninite (Figure 2.11E). Coffinite is the dominant uranium mineral in lower-grade ore resulting from
remobilization and alteration of the initially deposited uraninite.
Oxidation of the deposit has resulted in local remobilization and subsequent re-precipitation as
neoform uranyl minerals, observed in drillcore as an argillaceous yellow to orange overprint. In
petrographic sections, uranyl minerals occur as microveinlets cross-cutting uraninite and coffinite (Figure
2.11F). The only uranyl minerals identified are boltwoodite (HK(UO2)SiO4 . 1.5H2O) and uranophane
(Ca(UO2)2[HSiO4]2 . 5H2O). Uranyl minerals are rare, with samples containing less than 8 % (RIR) uranyl
phases, highlighting the overall reduced state of the deposit.
Trace amounts of brannerite (U(Ti,Fe)2O6), a tetravalent oxide mineral containing REE, Ti and
Fe-oxides, were identified with MLA/SEM but could not be confirmed with XRD because of its low
abundance. Brannerite occurs in association with coffinite as neoform, irregular anhedral aggregates,
within strongly clay altered ores. In agreement with Bruneton (1987), only a minor proportion of the TiO2
42
concentration forms U–Ti minerals with rutile, anatase and leucoxene as the main TiO2 phases (Bruneton,
1987).
2.4.1.1.1 Uranium Mineral Chemistry
U-bearing mineral phases, including uraninite and coffinite, have been shown to contain elevated
concentrations of Se, Mo and Zr (Janeczek and Ewing, 1992; Fayek et al. 1997; Heinrich et al., 2010).
Here, EMPA analysis was used to quantify the EOC content within uraninite and coffinite (Table 2.1).
The EOCs are compared with UO2, which is stable within the U-bearing mineral structure, to show their
variability within the minerals (Figure 2.12). Within uraninite, variable MoO3 contents occur with
concentrations up to 0.46 wt. % MoO3 (Figure 2.12C). Uraninite does not typically contain elevated SeO2
(nil–0.02 wt. %) or ZrO2 (nil–0.02 wt. %) contents, however one anomalous analysis from an altered
sample contains 0.23 wt. % SeO2 (Figure 2.12D, E). In contrast, coffinite contains elevated ZrO2 (nil–1.36
wt. %) and SeO2 (nil–0.34 wt. %) contents but minimal MoO3 (nil–0.18 wt. %).
The concentration of PbO, SiO2, CaO, MnO, and FeO within the U-bearing mineral phase is
reflective of alteration or U-recrystallization and can vary within the mineral structure (Figure 2.12A, B;
Fayek and Kyser, 1993; Fayek and Kyser, 1997). To constrain the EOCs within the paragenesis, their
compositional variations are illustrated in a SiO2–CaO–MnO–FeO vs. chemical U/Pb age diagram (Figure
2.13A). The most unaltered and earliest generations of uraninite, are characterized by high UO2 (78.65–
82.16 wt. %) contents, high PbO (13.73–15.81 wt. %) contents, and low CaO (0.66–1.35 wt. %), FeO
(0.04–0.22 wt. %), MnO (nil–0.08 wt. %), and SiO2 (0.12–0.23 wt. %) contents (Figure 2.13B). The U–
Pb chemical ages for relatively unaltered uraninite range from 1353 to 1150 Ma (Figure 2.13A). Elevated
MoO3 (>0.2 wt. %) concentrations in the uraninite coincide with high PbO contents (14.30–15.81 wt. %),
and Stage 1 crystallization (Figure 2.13C). Within the studied samples, primary uraninite does not contain
anomalous SeO2 (nil–0.02 wt. %) or ZrO2 (nil–0.02 wt.).
43
Table 2.1: Averaged EMPA results for uraninite and coffinite.
Mineral Uraninite Coffinite
Oxide wt.
(%) (n=90) DL (n=74) DL
UO2 78.65-85.21 0.2 63.97-78.15 0.3
ThO2 <DL-0.026 0.03 <DL-0.42 0.09
PbO 9.11-15.81 0.05 <DL-2.28 0.1
Y2O3 <DL-0.32 0.05 <DL-1.89 0.2
Ce2O3 <DL 0.05 <DL-0.96 0.2
Gd2O3 <DL-0.11 0.05 <DL-0.50 0.2
Dy2O3 <DL-0.17 0.1 <DL-0.58 0.3
Yb2O3 <DL 0.06 <DL 0.2
SiO2 0.12-0.65 0.02 8.97-17.92 0.08
TiO2 <DL-0.88 0.04 <DL-1.91 0.1
ZrO2 <DL 0.04 <DL-1.36 0.2
MoO3 <DL-0.46 0.03 <DL-0.18 0.09
FeO 0.041-0.50 0.03 <DL-0.82 0.09
MnO <DL-0.16 0.03 <DL-0.14 0.09
CaO 0.66-1.93 0.03 0.45-3.6 0.07
SeO2 <DL-0.23 0.02 <DL-0.34 0.06
Total (%) 96.38-98.46 85.01-92.12
Samples (n) refers to the number of spots analyzed on uraninite and coffinite crystals.
Corresponding detection limits (DL) are listed adjacent to the range. <DL indicates
that a given oxide was below detection limit.
44
Figure 2.12: Bivariate diagram showing the linear relationship between UO2 and the EOCs. A)
Uraninite alteration elements: SiO2, CaO, MnO and FeO plotted vs UO2 . The concentration of UO2
decreases with increased alteration. B) Lead is a mobile element particularly within U-oxide
minerals. The concentration of Pb within the U oxide is a function of time and reflects the
mineralogy and stage of crystallization/resetting. C) D) E) Concentration of deleterious elements:
Mo, Se, and Zr are shown in standard EMPA oxide notation and plotted against UO2. The
concentration of UO2 is stable within the U oxide mineral however changes in wt. % reflect
incorporation of other elements into the structure.
A B
C D
E
End Member Uraninite
45
Brecciation, alteration and to a lesser extent recrystallization occurs in association with a
continuum of Pb-depletion and concomitant Ca, Mn, Fe, Si enrichment. These uraninite generations are
characterized by high UO2 (81.19–85.21 wt. %) and intermediate PbO (9.11–13.75 wt. %), CaO (0.81–
Figure 2.13: Bivariate diagram showing the linear relationship between chemical U–Pb ages (A)
and PbO (B) to uraninite alteration elements: SiO2, CaO, MnO and FeO. The U–Pb chemical
ages are calculated using the method of Bowles (1990). C)D)E) Concentration of deleterious
elements: Mo, Se, and Zr are shown in standard EMPA oxide notation.
C B
D E
A
46
1.93 wt. %), FeO (0.18–0.50 wt. %), MnO (0.05–0.16 wt. %) and SiO2 (0.17–0.65 wt. %) contents.
Altered uraninite crystals have younger U–Pb chemical ages ranging from 1143 to 761 Ma. These
generations of uraninite do not contain significant MoO3 (nil–0.16 wt. %) or ZrO2 (nil). Although
uraninite does not typically contain elevated SeO2 (nil–0.02 wt. %), one anomalous analysis yielded 0.23
wt. % SeO2 suggesting Se can reside within Stage 2 altered uraninite (Figure 2.13D).
Coffinite is characterized by relatively low UO2 (63.97–78.15 wt. %) and PbO (nil–2.28 wt. %)
contents, and high CaO (0.45–3.60 wt. %), FeO (nil–0.82 wt. %), MnO (nil–0.14 wt. %), and SiO2 (8.79–
17.92 wt. %) contents. The EMPA-analyzed coffinite yield young chemical ages ranging between 242 to
0 Ma (Figure 2.13A). These Stage 4 crystals contain elevated ZrO2 (nil–1.36 wt. %) and SeO2 (nil–0.34
wt. %) contents but minimal MoO3 (nil–0.18 wt. %) contents (Figure 2.13 C, D, E).
2.4.1.2 Arsenides and Sulpharsenides
Cigar Lake mineralization occurs in association with Ni–Co arsenides and sulpharsenides
(Bruneton, 1987; Reyx and Ruhlmann, 1993). Petrography, XRD and MLA identified gersdorffite
(NiAsS), cobaltite (CoAsS), niccolite (NiAs), rammelsbergite (NiAs2), skutterudite ((Ni,Co,Fe)As3),
glaucodot (Ni,Fe)AsS), erythrite (Co3(AsO4)2.8H2O), annabergite (Ni3(AsO4)2
.8H2O) and aerugite
(Ni9(AsO4)2AsO6), listed in order of decreasing overall abundance. The arsenides and sulpharsenides
occur as prismatic euhedral to subhedral disseminations, crystal aggregates, or botryoidal and colloform
masses (Figure 2.14). Consistent with the mineralogy, the whole-rock geochemistry confirms that the As-
phase is dominated by gersdorffite, niccolite and cobaltite (Figure 2.15). Stoichiometrically these
arsenides and sulpharsenides have a 1:1 molar ratio of Ni–Co:As. Furthermore, Ni-rich mineral end-
members predominate over their Co-rich varieties throughout most of the deposit.
47
Figure 2.14: BSE images of arsenides and sulpharsenides. A) Uraninite with GER1 shrinkage crack
fill. B) GER1 nucleating on niccolite. Niccolite is partially consumed by the parasitic overgrowth.
C) Subhedral GER1 overgrowing bravoite (BR) as a parasitic overgrowth. D) Colloform
sulpharsenides with BR core and concentric growths of niccolite (NICC) and GER1. Sulphur
content decreases towards the crystal margins. E) Gersdorffite (GER2) overgrowing chalcopyrite
(CPY1) in association with coffinite. F) GER2 overgrowing pyrite and galena. Gersdorffite is
rimmed by coffinite.
A B
C D
E F
48
Figure 2.15: Ni+Co molar % vs. As molar % from whole-rock geochemistry. Molar % values are
normalized to 100% with SiO2. Corresponding mineral slopes are shown for gersdorffite (GER),
cobaltite (COB), niccolite (NICC), rammelsbergite (RAM) and Ni-Co skutterudite (SKUTT). Cigar
Lake Phase 1 pods are dominated by 1:1 molar ratio (Ni+Co:As) minerals: gersdorffite, cobaltite
and niccolite.
At Cigar Lake, Ni–Co–As ores have complex textures, and mineral relationships suggest several
stages of crystallization. Stage 1 arsenides and sulpharsenides, comprising predominantly gersdorffite
(GER1), occur overgrowing botryoidal U1 crystals or intergrown within U1 shrinkage cracks (Figure
2.14A). In contrast, Stage 2 arsenides and sulpharsenides, occur as overgrowths on uraninite, arsenides
and sulphides or disseminated within the chlorite, illite matrix (Figure 2.14E, F). Stage 2 arsenides and
sulpharsenides have a strong association with coffinite. Variable reaction alteration textures are prominent
throughout the arsenide and sulpharsenide phases with crystal nucleation, parasitic overgrowths,
concentric zoning, skeletal dissolution and replacement textures. Gersdorffite, the most prevalent As-
bearing mineral in the Phase 1 pods, is commonly observed nucleating on niccolite and pyrite with
subsequent parasitic overgrowth forming concentric crystals with vuggy and clay filled cores (Figure
2.14B, C, D).
49
Cobaltite is the dominant Co-bearing mineral within the deposit. Texturally, cobaltite is similar to
gersdorffite occurring as euhedral to subhedral crystal disseminations, or crystal aggregates within the
chlorite, illite matrix. However, unlike gersdorffite, cobaltite was not observed in proximity to U1
uraninite. Cobaltite occurs as late Stage 1 and Stage 2 overgrowth on sulphides and arsenides often in
association with coffinite.
Only minor, local occurrences of arsenates such as erythrite, annabergite and aerugite were
identified indicating that the deposit has largely remained in a reduced state since formation. Arsenates
occur along the crystal boundaries as alteration of arsenides and sulpharsenides, and typically contribute
<3% of the total arsenic concentration within a given sample. Texturally, arsenates demonstrate late
mineral relationships formed during Stages 3 and 4 of the paragenetic sequence resulting from late
incursion by oxidized fluids along structures (Figure 2.10).
2.4.1.2.1 Arsenide and Sulpharsenide Mineral Chemistry
Gersdorffite is the most dominant As-bearing mineral within the Phase 1 pods, having a variable
and complex chemistry with both cation and anion element substitution. Gersdorffite contains up to 5.13
wt. % Co (Table 2.2; Figure 2.16A). Cobalt has a strong negative correlation with Ni (Pearsons
correlation coefficient, PCC = -0.998), implying that Co readily substitutes for Ni, forming a nearly
complete solid solution between gersdorffite and cobaltite. The substitution of Ni by Fe is limited within
the gersdorffite samples, with Fe contents ranging from nil to 0.68 wt. %. Bismuth and Se occur locally
within gersdorffite crystals with concentrations up to 12.44 wt. % Bi and 2.56 wt. % Se (Figure 2.16B, C,
D). Elevated Bi and Se contents occur predominantly in association with a concentric, zoned, band
proximal to the perimeter of the gersdorffite crystals, observed as brighter illumination in BSE (Figure
2.17). The S content is highly variable ranging from 11.21 to 19.60 wt. %, and there is a strong negative S
correlation with Bi (PCC = -0.98) and Se (PCC = -0.97). Arsenic contents remain more consistent and
less depleted, ranging from 39.74 to 45.99 wt. %. Bismuth and Se enrichment is primarily a function of S-
depletion suggesting that Se2- and Bi2- substitute for S2- within the sulpharsenide (Figure 2.16B, C, D;
50
Figure 2.17). Therefore, changes in the concentrations of these elements within the mineralizing solutions
may control their uptake, with concentrations of Se and Bi increasing when S2- becomes deficient.
Table 2.2: Summarized EMPA results for arsenides and sulpharsenides.
Mineral Gersdorffite Niccolite Cobaltite
Wt. % DL (n=56) (n=8) (n=29)
Ni 0.03 24.7-36.17 42.84-44.83 3.96-14.05
Co 0.02 0.34-5.13 0.13-0.59 21.91-30.81
Fe 0.02 <DL-0.68 0.011-0.80 0.17-1.43
Cu 0.02 <DL-2.83 <DL-0.27 <DL-0.43
Ag 0.02 <DL-0.18 <DL <DL-0.056
As 0.1 39.74-45.99 52.21-55.31 44.10-45.38
Sb 0.02 <DL-1.00 0.022-1.2 <DL-0.16
Bi 0.05 <DL-12.44 <DL-3.12 <DL-0.79
S 0.03 11.21-19.60 0.14-1.06 18.83-19.75
Se 0.04 <DL-2.56 <DL-0.47 <DL-0.66
Samples (n) refers to the number of spots analyzed on uraninite and coffinite crystals.
Corresponding detection limits (DL) are listed adjacent to the range. <DL indicates
that a given oxide was below detection limit
Gersdorffite is the only As-bearing mineral showing elevated Cu and Ag with concentrations up
to 2.83 wt. % and 0.18 wt. % respectively. Copper and Ag have a strong correlation (PCC = 0.97) and
typically occur in association with elevated Bi and Se content suggesting a coupled substitution may
facilitate slight changes in the ionic size, charge or strength responsible for Cu and Ag uptake. The
average chemical formula for Cigar Lake gersdorffite in this study is (Ni0.98Co0.04)As1.00(S0.97Bi0.02Se0.01).
With its simple chemical structure, niccolite is generally homogenous showing minimal chemical
variability and element substitutions (Table 2.2; Figure 2.16). Minimal cation substitution of Ni by Co
and Fe is observed with maximum concentrations of 0.59 wt. % and 0.80 wt. % respectively. Locally, Bi,
Sb and Se appear to substitute for As with up to 3.12 wt. % Bi, 1.20 wt. % Sb, and 0.47 wt. % Se. Only
minor S is incorporated into the mineral with concentrations ranging from 0.14 to 1.06 wt. %. The
average chemical formula for Cigar Lake niccolite in this study is (Ni0.98Co0.01Fe0.01)As0.97S0.02Bi0.01.
51
Cobaltite is a Co-sulpharsenide end-member that forms a solid solution with gersdorffite. In the
studied samples, cobaltite contains up to 14.05 wt. % Ni, forming a nearly complete solid solution with
gersdorffite (Figure 2.16). In contrast to the other As-bearing minerals analyzed, cobaltite contains
relatively high Fe contents ranging from 0.17 wt. % to 1.43 wt. %. Bismuth, Se, and Sb content is low
with maximum concentrations of 0.79 wt. %, 0.66 wt. % and 0.16 wt. %, respectively. The average
chemical formula for Cigar Lake cobaltite in this study is (Co0.74Ni0.24Fe0.02)As1.00S1.00.
Figure 2.16: A) Molar % (M %) proportions of S, Ni and Co showing the composition of the
main arsenide and sulpharsenides: gersdorffite, cobaltite and niccolite. B) Bivariate diagram
with S and As M% showing stoichiometric gersdorffite, cobaltite and niccolite. Some
gersdorffite crystals are S deficient. C) Selenium and Bi substituting for S within gersdorffite
crystals. D) Selenium and Bi showing a strong correlation in niccolite and gersdorffite.
A
C
B
D
Cobaltite Gersdorffite
Niccolite
52
Figure 2.17: A) Subhedral gersdorffite crystals in BSE. B) Heat map showing elevated Bi as
brighter colours around the margins of the gersdorffite crystals. C) Heat map showing elevated Se
content as brighter colours at the core and along the crystal margins in association with anomalous
Bi content. D) Heat map showing elevated Co as brighter colours around the margins of the
gersdorffite crystals.
2.4.1.3 Sulphides
At Cigar Lake, the dominant sulphide minerals identified are chalcopyrite (CuFeS2), pyrite
(FeS2), bornite (Cu5FeS4), sphalerite (ZnS), galena (PbS), chalcocite (Cu2S), pyrrhotite (Fe1-xS) and
molybdenum (MoS2) listed in order of abundance (Figure 2.18).
The Cu phase is predominantly chalcopyrite with Stage 1 chalcopyrite (CPY1) occurring
intimately with uraninite as shrinkage crack intergrowths, overgrowths on uraninite and anhedral masses
(Figure 2.18A, B, C). Within the samples analyzed, a strong association between CPY1 and high-grade
uraninite ores is observed. Stage 2 and Stage 3 chalcopyrite (CPY2, CPY3) occur as anhedral aggregates,
A B
C D
Selenium
Bismuth
Cobalt
53
blebs, veinlets and disseminated crystals within the chlorite, illite matrix in association with coffinite.
Chalcopyrite is prone to pyrite replacement, particularly common on the north and northwest ends of both
Phase 1 pods suggesting increasing S and decreasing O and Cu activities (Figure 2.18A, B, C, D).
Relative to the other Cu-sulphides, chalcopyrite occurs more commonly with Ni-arsenides and
sulpharsenides.
Bornite occurs in association with relatively monometallic, high-grade U ores. Bornite forms
anhedral aggregates and laths overgrowing uraninite or disseminated within the chlorite, illite matrix
commonly in association with galena (Figure 2.18E, F). Bornite forms a solid solution with chalcopyrite
demonstrating bornite–chalcopyrite exsolution textures. Chalcocite occurs in samples containing elevated
concentrations of Cu (> 7 wt. %), commonly in association with bornite. Chalcocite and bornite are
interpreted to be the result of Cu remobilization and re-crystallization reflecting epigenetic fluid incursion
during Stage 3 of the mineral paragenesis. The following crystallization and alteration series is proposed
for the Cu-phase as a function of increasing Cu and decreasing S activities: chalcopyrite (CuFeS2) –
bornite (Cu5FeS4) – chalcocite (Cu2S).
Pyrite is a common mineral phase within the deposit with Stage 1 pyrite (PY1) occurring as
subhedral to euhedral disseminated crystals and anhedral crystal aggregates within a strongly chlorite,
illite matrix. Subsequent stages of pyrite (PY2, PY3, and PY4) occur as euhedral to anhedral
disseminations, sooty disseminations, overgrowing aggregates, and as a common replacement of
chalcopyrite. Pyrrhotite occurs locally within the deposit as blades and laths with boxwork-like texture,
and as chalcopyrite-pyrite exsolution and overgrowths (Figure 2.18D). The following crystallization and
alteration series is observed with increasing S and decreasing O fugacity: chalcopyrite (CuFeS2) – pyrite
(FeS2) – pyrrhotite (Fe1-xS).
54
Figure 2.18: A) Chalcopyrite CPY1 replaced and overgrown by pyrite (PY1) with late laths of
pyrrhotite (PYR). Texture demonstrates increasing S and decreasing O activities. Petrographic
image in plane polar (PP) reflected light (RL). B) CPY1 replacement by pyrite with cobaltite COB2
overgrowth (PP and RL). C) CPY1 replacement by pyrite. Atoll replacement of pyrite by chalcocite
(CC) (PP and RL). D) Subhedral pyrite and anhedral sphalerite (SPH) with laths of pyrrhotite.
Sphalerite appears to be older than pyrrhotite (PP and RL). E) Wispy almost fibrous coffinite
(COFF) occurring in association with Stage 3 minerals bornite (BO) and galena (GN). F) Extensive
coffinitization of uraninite occurring in association with Stage 3 bornite. Bornite shows an
exsolution texture with minor chalcopyrite.
55
Galena typically occurs as subhedral to euhedral crystals and is more prevalent in association
with higher-grade U samples (Figure 2.18E). Galena commonly occurs as overgrowths on uraninite or
intergrown confined crystals within uraninite microfractures and shrinkage cracks. The main source of Pb
within the mineralization is from the decay of uranium since formation.
Sphalerite was the only Zn-mineral identified within the deposit. Sphalerite occurs as anhedral
aggregates and blebs commonly disseminated throughout the chlorite, illite matrix (Figure 2.18D).
Textural mineral relationships and associations suggest sphalerite crystallized during the waning stages of
the main mineralizing event (Stage 1). LA-ICP-MS was used to determine the concentration of Fe within
the Cigar Lake sphalerite and the formula was calculated as (Zn0.96Fe0.04)S.
Molybdenite was identified occurring in minor concentrations as anhedral disseminated crystals
within the chloritic matrix (Figure 2.19). No crosscutting relationships or mineral associations were
observed to constrain the relative timing of molybdenite within the deposit.
Figure 2.19: Molybdenite shown in BSE (A) and with MLA interpretation (B). Anhedral
molybdenite crystals (Blue) are disseminated within the chloritic matrix (Green).
A B
56
2.4.1.4 Zirconium
Zirconium was identified within detrital zircons (Figure 2.20) and within coffinite (Figure 2.13E).
The zircon crystals display varying degrees of alteration ranging from completely altered to pristine
(Figure 2.20). Detrital zircons were even identified within intensely altered samples immersed in a
massive chlorite matrix suggesting resistance to some clay forming hydrothermal fluids (Figure 2.20B).
In contrast to uraninite, which yielded below detection limit ZrO2 contents, coffinite crystals contain up to
1.36 wt. % ZrO2. This suggests conditions suitable for coffinitization, occurring predominantly within the
later stages of the paragenesis (Figure 2.10: Stage 2–4), were more amenable to zircon destruction and Zr
mobility.
Figure 2.20: Zircon (ZR) crystals identified and shown with BSE. A) Zircon crystal is overgrown
and intergrown with galena (GN). B) Well-rounded, detrital zircon crystal within strongly clay
altered and chloritized matrix.
2.4.1.5 Clay Mineralogy
Clay characterization of the surrounding alteration zone at Cigar Lake has been extensive (e.g.
Percival and Kodama, 1989; Pacquet and Weber, 1993; Percival et al., 1993; Mosser et al., 1996; Percival
et al., 2000). However, characterization of clays within the ore itself has been limited due to the
availability of uranium ore samples. The clay mineralogy was characterized here using XRD and a
A B
57
portable shortwave infrared spectrometer (SWIR). The purpose was to help characterize the clay
mineralogy within the ore zone.
Results from on-going drilling programs indicate that the clay mineralogy is dominated by white-
mica mixtures of illite and muscovite. Local patches of phengite, paragonite, Fe-chlorite, Mg-chlorite,
kaolinite and montmorillonite are observed throughout the ore body (Appendix I). Fe-chlorite is more
common distal to the unconformity whereas Mg-chlorite occurs proximal to the unconformity and within
the underlying basement. The white-mica composition also shifts with distance to the unconformity, with
paragonite more proximal to the unconformity and muscovite, illite and phengite occurring distally.
Metallurgically, swelling smectite clays are the most problematic for settling in the slurry comminution
circuit but have only rarely been documented within the Athabasca Basin (e.g. Percival et al., 1993, Ito
and Azam, 2017). Here, minor patches of montmorillonite are measured with SWIR and typically
correspond with geological structures. No glycol testing was performed in the current study to assess clay
swelling. Percival et al. (1993) reported illite–smectite mixed layers in the alteration zone at Cigar Lake
with 5–10% expandability. Alternatively, illite–chlorite mixed clay layers at Close Lake and McArthur
River have XRD patterns consistent with smectite mixed-layer clays, however no swelling was observed
with glycolation suggesting sepiolite, illite–chlorite, illite–vermiculite or hydrobiotite (Quirt, 1999).
2.4.2 Three Dimensional Modelling of Element and Ore Distribution
The overall geochemical characterization of the orebody was initially described by Bruneton
(1987). With 30 years of delineation and operational drilling, the geochemical dataset has been vastly
expanded with improved spatial coverage. Implicit modelling with Leap Frog 3D software was used to
create geochemical grade shell interpolants for prominent ore forming elements of interest: U, Ni, Co, As,
Zn, Pb, Cu, K20, Al2O3, MgO and ZrO2.
The orebody contains a cumulative ~11,000 m3 high-grade (>40% U3O8) core, enveloped within a
lower grade shell (Bruneton, 1987). The magnitude and extent of the encapsulating clay is proportional to
58
the grade and thickness of the orebody. The Phase 1, East-Pod contains the highest-grade and most
continuous high-grade mineralization overlying the main basement, east-west oriented, and strongly
graphitic, semi-brittle fault and occurs in association with a deposit scale redox front (Figure 2.21A). On
the west-end of the East-Pod, blowout high-grade mineralization coincides roughly with crosscutting, or
potentially Riedel northwest oriented faults. High-grade lenses extending out from the high-grade core
display an east-northeast orientation coincident with the regional basement foliation, major regional
faults, and local crosscutting or Riedel east-northeast faults (Bruneton, 1993). The lower-grade West-Pod
has an overall east-northeast orientation with high-grade ores focused locally on the far west-side of the
pod. The mineralogy of higher-grade U ores is dominated by uraninite whereas the lower-grade ores
contains a higher proportion of coffinite.
The high-grade U-ores have a strong geochemical and empirical spatial correlation with Cu, Mo,
Se, and Pb. The correlation with Pb is expected as most Pb is radiogenic (Bruneton, 1987). Molybdenum
and Se, at elevated concentrations (Mo >5000 ppm, Se > 300 ppm), display an inverse relationship along
the high-grade U3O8 corridor (Figure 2.21B). Molybdenum occurs as molybdenite and within uraninite
with concentrations to 0.31 wt. % Mo in association with elevated Pb suggesting coeval crystallization
within U1. In contrast, Se2- substitutes for S2- predominantly within Stage 2 sulphides and sulpharsenides
and was also observed within Stage 2 altered uraninite. The mineralogy, paragenesis and empirical spatial
distribution, suggest that elevated Se along the high-grade U3O8 corridor may reflect localities particularly
effected by Stage 2 fluid incursion.
59
Figure 2.21: Leap Frog 3D implicit geochemical grade shells for the Phase 1 pods. A) High-grade U
mineralization (>40% U3O8) showing more continuous high-grade ore in the East-Pod. High-grade
ore is more continuous above the main east-west oriented graphitic fault and occurs along the redox
front between more oxidized ore to the southwest and more reduced polymetallic ore to the
northeast. B) Se and Mo showing a strong spatial correlation with high-grade ore. Selenium and
Mo have an inverse relationship along the high-grade corridor. C) Copper and Ni showing spatial
zoning in the deposit with Cu typically occurring to the southwest and distal to the unconformity.
D) Strong inverse relationship between Fe2O3 (total) and As. Hematite is the dominant Fe2O3
controlling mineral highlighting the transition between oxidized ores to the southwest and more
reduced polymetallic ores to the northeast. High-grade mineralization is concentrated along the
redox boundary. Also shown is mine grid section 10898 from Figure 2.4.
60
In contrast, As, Ni, Co and Zn geochemically and spatially show a limited correlation with
high-grade U3O8 and are offset towards the north, east and northeast ends of both Phase 1 pods (Figure
2.21C,D). These elements also demonstrate local upgraded concentrations in association with east-west
oriented semi-brittle, graphitic basement faults. Consistent with the mineralogy, arsenic has a strong
correlation with Ni and Co. Particularly well developed on the East-Pod, transitional metals are zoned
from the southwest to northeast, with a crystallization series of: Cu – to Ni – to Co – to Zn. This zonation
is also observed vertically with Cu typically occurring further from the unconformity than Ni, Co and Zn.
The concentration of total Fe2O3 can be attributed predominantly to siderite and hematite
(Bruneton, 1987). Fe2O3 shows a strong inverse spatial relationship with As (Figure 2.21D), occurring
predominantly along the south and southwest ends of both Phase 1 pods. High-grade U3O8 is concentrated
directly between the more oxidized monometallic ores to the south and southwest and polymetallic ores to
the north and northeast.
Historically, Zr has not been a routinely analyzed element at Cigar Lake. Therefore, the spatial
distribution of Zr data is limited, restricting the ability for empirical spatial modelling. Zirconium does
not have a strong spatial or geochemical correlation with U. The preliminary distribution suggests that Zr
appears to occur preferentially along the flanks of the deposit particularly on the southern side. Several
isolated pods (>2000 ppm) occur in association with breaks in the high-grade U (>50% U3O8) corridor.
These Zr enriched zones still contain strong to intense clay alteration. Therefore, it is unclear whether
these observed enrichments represent a lithological control or lack of geochemical conditions for zircon
dissolution.
2.4.3 Three Dimensional Modelling of the Mineral Distribution
As previously discussed in Section 2.3, mineral stoichiometry and rock geochemistry were used
for mineral quantification throughout the deposit to extend the spatial coverage of mineralogical data.
Implicit modelling with Leap Frog 3D software was used to create mineral grade shell interpolants for the
dominant sulphide and arsenide minerals: gersdorffite, cobaltite, niccolite, rammelsbergite, chalcopyrite,
61
bornite, chalcocite and pyrite. Normatively-derived mineral grade shells, with composited data at 0.5 m,
exhibit good correlation with the measured semi-quantitative XRD and MLA results (Figure 2.22).
As expected, the grade-shells for sulpharsenides and arsenides show a spatial zonation with As-
bearing minerals occurring in the north and northeast ends of both pods, with the highest concentrations
being offset from high-grade (>40% U3O8) ores (Figure 2.21 and 2.23). Gersdorffite is the most pervasive
As-bearing mineral and is the only arsenide to occur in association with the high-grade (>40% U3O8)
uranium ore corridor. This empirical spatial correlation is consistent with mineral textures showing only
GER1 in association with primary uraninite (U1). Cobaltite, niccolite and rammelsbergite all show a
strong inverse relationship to the high-grade U3O8 ore corridor, reflecting higher Ni–Co–As activities
along the north and northeast ends of both Phase 1 pods.
In contrast, the Cu-phases, dominated by chalcopyrite, have a strong spatial association with the
high-grade U3O8 ore corridor (Figure 2.24). Consistent with the mineral observations, chalcopyrite is the
most pervasive Cu-bearing mineral within the deposit. Particularly well developed on the East-Pod,
bornite and chalcocite occur in higher proportions in the south and southwest ends of both pods. This
zonation highlights the overall redox gradient observed across both the east and west pods responsible for
mineral stability.
Pyrite is a fairly common mineral in the deposit. The Leap Frog 3D interpolant demonstrates a
pervasive spatial distribution. The empirical spatial distribution of pyrite does not appear to be
constrained by the redox gradient.
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Figure 2.22:10930 mine grid cross-section showing normative mineral proportions as implicitly modelled grade shells. Pie charts show
MLA/RIR-XRD proportions. Gersdorffite occurring with chalcopyrite in the south (hanging wall) and cobaltite in the north (footwall).
Chalcocite occurring distal to the unconformity resulting from Cu remobilization during Stage 3 meteoric water incusion.
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Figure 2.23: Leap frog 3D interpolant grade shells of As-bearing minerals: gersdorffite, cobaltite,
niccolite and rammelsbergite. Interpolants have been generated from normative mineral
proportions calculated from whole-rock geochemistry. Mineralogical concentrations have been
composited at 0.5 m. Also shown are the mine grid section locations 10898 and 10930 from Figure
2.4 and Figure 2.22 respectively.
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Figure 2.24: Leap frog 3D grade shell interpolants of major sulphides within the Cigar Lake
deposit: chalcopyrite, bornite, chalcocite and pyrite. Interpolants have been generated from
normative mineral proportions calculated from whole-rock geochemistry. Mineralogical
concentrations have been composited at 0.5 m. Also shown are the mine grid section locations 10898
and 10930 from Figure 2.4 and Figure 2.22 respectively.
2.5 Discussion
Overall, the mineralogy of the Cigar Lake ores suggests that the abundance is
uraninite>coffinite>>uranyl minerals>>>brannerite, as described in previous studies (Bruneton, 1987;
Reyx and Ruhlmann, 1993). From a hydrometallurgical perspective, the overall distribution of the
uranium ore minerals is conducive to acid leach milling using oxidants for conversion of tetravalent to
more soluble hexavalent U complexes (IAEA, 2001; Bowell et al., 2011). Uraninite has a relatively
65
simple structure and readily dissolves when an oxidant is used to convert from tetravalent to hexavalent U
(Bowell et al., 2011). Coffinite, although more complex, is still readily processed under oxidized acid
leach conditions. Uranyl uranium minerals are not a concern for processing because they readily dissolve
in acid (Bowell et al., 2011). The occurrence of brannerite can be more challenging in acid milling
(Bowell et al., 2011), however it is a trace constituent within the Cigar Lake uranium ores.
Within the Phase 1 pods, As, Ni, and Co occur primarily in a reduced state as arsenides and
sulpharsenides. The arsenides and sulpharsenides are dominated by 1:1 ratio, Ni–Co:As minerals such as
gersdorffite, niccolite and cobaltite. Although thermodynamic data is not readily available for the
arsenides and sulpharsenides that occur at Cigar Lake, milling experience in the Athabasca Basin over the
last forty years has shown that Ni–Co arsenides and sulpharsenides are typically less exothermic during
oxidation than Ni–Co biarsenide and triarsenide and are therefore less problematic during milling
(McClean Lake Metallurgists – Areva Resources, personal communication, October 2015).
The complex Ni–Co–As phase textures suggest fluctuating S, As, Ni and Co activities during
formation that likely changed with redox conditions. Increasing S and decreasing As activities result in
the following sequence of crystallization: skutterudite (NiAs3) – rammelsbergite (NiAs2) – niccolite
(NiAs) – gersdorffite (NiAsS) – bravoite (NiS2). This sequence has been described extensively by
Bruneton (1987) and Reyx and Ruhlmann (1993) and reflects the overall evolution of the hydrothermal
fluids (Figure 2.14B). In contrast, decreasing S activity and increasing As activity is observed locally and
results in alteration and crystallization in the opposite direction along the aforementioned series (Figure
2.14C, D). Concurrent with the evolving and local fluctuation in As and S activities, Ni and Co readily
substitute for one another forming solid solutions. The Ni-rich mineral end-members dominate over Co-
rich varieties throughout most of the deposit. Textural relationships suggest a transition from Ni to Co,
during the waning stages of successive hydrothermal events.
Molybdenum was identified in the mineral phase molybdenite and within Stage 1 uraninite with
concentrations up to 0.46 wt. % MoO3. Elevated Mo concentrations in the uraninite coincides with
66
elevated Pb levels suggesting that Mo is syngenetic with primary uranium mineralization. Alternatively,
the correlation between Pb and Mo may suggest the formation of Pb-molybdates within uraninite,
however this could not be confirmed by EMPA or SEM. The spatial distribution of Mo coincides with
high-grade ore (>40% U3O8) confirming its mode of occurrence within the uraninite (Figure 2.21A, B).
Selenium was found to occur in sulphides and sulpharsenides with Se2- substituting for S2-.
Selenium appears to be paragenetically late (Stage 2) and is observed with increasing concentrations
towards the boundaries of GER2 crystals. Galena, another paragenetically late mineral was also observed
to be prone to Se uptake as galena forms a solid solution with clausthalite (PbSe). Selenium typically
coincides with anomalous Bi within the sulpharsenide crystals, where Bi appears to substitute for S. This
suggests that S fugacity may control the spatial distribution of Se, with Se concentrations increasing with
decreasing S activity (e.g. Huston et al., 1995; Layton-Matthews et al., 2008). Some coffinite crystals
contain anomalous SeO2 content with up to 0.34 wt. % SeO2. The spatial distribution of Se within the
orebody shows a strong correlation with the high-grade U3O8 corridor but an inverse relationship with
elevated Mo (>5000 ppm) suggesting whole-rock Se content may reflect localities within the deposit
particularly effected by Stage 2 fluid incursion.
The mode of occurrence for Zr was determined to be within detrital zircon crystals and within
coffinite with some crystals yielding up to 1.36 wt. % ZrO2. Uraninite samples contained negligible ZrO2
content, although Cigar Lake uraninite has been reported to contain up to 1.4 wt. % ZrO2 (Janeczek and
Ewing, 1992; Fayek et al., 1997). Preliminary modelling of the empirical spatial distribution suggests that
Zr appears to occur preferentially along the flanks of the deposit particularly on the southern side. This
distribution is consistent with the primary mode of occurrence as detrital zircons.
2.6 Conclusions
The holistic geometallurgical approach of integrating geochemical, mineralogical, geological, and
geospatial characterization proved effective with the Cigar Lake uranium deposit. The U-bearing
mineralogy was identified as predominately uraninite and coffinite conducive with acid leach milling. The
67
dominant deleterious element phases were identified including: 1) As, 1:1 molar ratio Ni–Co:As arsenides
and sulpharsenides; 2) Se, substituting for S2- in sulphides and sulpharsenides; 3) Mo, occurring as
molybdenite and within uraninite; and 4) Zr occurring as zircon and within coffinite. Geochemistry
correlates with the mineralogy and mineral proportions allowing for more extensive modelling using the
extensive historic multi-element whole-rock geochemical database facilitating the refinement of
metallurgical protocols. Using these techniques, the geometallurgy at Cigar Lake can now be used to
optimize and reduce risk during mining, milling and mine tailings management.
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Chapter 3
EVOLUTION OF THE HIGH-GRADE POLYMETALLIC
UNCONFORMITY-RELATED URANIUM CIGAR LAKE ORE BODY
3.1 Abstract
Cigar Lake is a high grade polymetallic uranium orebody with complex geochemistry and
mineralogy. It is located in the Proterozoic Athabasca Basin in Northern Saskatchewan, Canada.
Delineation and ongoing operational drilling have provided improved exposure of the Phase I pods for
mining. Mineralogical and geochemical characterization has been undertaken on the uranium minerals,
and arsenide and sulphide gangue minerals, using semi-quantitative techniques including XRD, mineral
liberation analysis (MLA) and optical petrography. Electron microprobe analysis and laser ablation ICP-
MS were used to measure the chemical compositions and element deportment of selected minerals. Stable
and radiogenic isotopes were used to ascertain the age, source and evolution of fluids that formed and
altered the deposit and geostatistical implicit modelling was used to evaluate the empirical spatial
distribution of whole-rock geochemistry. These mineralogical, geological, geochemical and geospatial
characteristics are integrated here to update the evolutionary model for the Cigar Lake uranium deposit.
Four major evolutionary stages are discernable, representing transitions in mineralogy, mineral
association and textural relationships, and U-Pb isotope chemistry. The main mineralization event
occurred before 1468±93 Ma and was likely syngenetic with the basin wide ca. 1590 Ma U mineralization
event. Stage 1 sulphides and sulpharsenides have relatively low 206Pb/204Pb (~75) and high 207Pb/206Pb
(~0.5) values, confirming that primary mineralization was polymetallic containing Ni, Co, As, S, Zn and
Mo. Relatively high δ34S values up to 14.6‰ from the ore zone indicate a significant basinal and marine
contribution of SO42- relative to basement-derived S and an open-system with a well-mixed source. Syn-
ore calcite equilibrated with a fluid having δ18O values between +1.8 to +7.2‰, typical of diagenetic
basinal brine in the Athabasca Basin, and low δ13C values ranging from -22.4 to -21.8‰, consistent with a
69
source from underlying graphite. Stage 2 alteration with U/Pb ages of 1270–1163 Ma reflect the impact of
fluid incursion from the Grenville Orogeny, the Mackenzie dyke swarm and the Moore Lake olivine
diabase. Crystallization, recrystallization and Pb absorption resulted in radiogenic (206Pb/204Pb ~1000)
sulphides and sulpharsenides and enhanced clay alteration. Stage 3 alteration, constrained by U/Pb ages at
947–755 Ma, resulted from meteoric water incursion during the breakup of the supercontinent of Rodinia.
Coeval carbonates equilibrated with fluids having δ18O values between -12.7 to -3.1‰ and δ 13C values
from -18.8 to -11.1‰, suggesting a distal C source from an oxidized organic-rich environment. These
fluids were responsible for extensive hematization and crystallization of highly radiogenic (206Pb/204Pb
~3000–20,000) bornite, chalcocite and galena. Stage 4 alteration manifests as extensive coffinitization of
the main orebody and re-mobilization of U into brittle structures overlying the deposit, resulting in
perched mineralization. These oxidized fluids had low δ2H values (< -151‰) and were responsible for re-
mobilized C from underlying graphite forming bituminous aggregates and hydrocarbon buttons with δ13C
ranging from -31.2 to -27.3‰.
The paragenesis, geochronology and stable isotopes reveal a protracted history, with ongoing
episodic fluid incursion, driven by far-field tectonics that resulted in alteration and re-mobilization of
selected elements. Preservation of the deposit is a function of geochemical stabilities with sulphides and
arsenides showing a tendency to buffer penetrating fluids.
70
3.2 Introduction
Unconformity-related uranium deposits consist of massive pods or veins of uranium that occur
proximal to an unconformity between siliciclastic Paleoproterozoic-Mesoproterozoic sedimentary basins
and the underlying metasedimentary basement (Kyser and Cuney, 2015). Unconformity-related uranium
deposits in the Athabasca Basin of northern Saskatchewan are unrivalled as the highest-grade uranium
deposits in the world (IAEA, 2009). Uranium mines in Saskatchewan produce approximately 16% of the
total current global production of U3O8 (World-nuclearorg, 2016). Cigar Lake, with the highest mined
uranium grades in the world, is the newest operation in the mining district, achieving commercial
production status in 2015 (Cameco, 2015).
The Cigar Lake uranium deposit is located approximately 650 km north of the city of Saskatoon
and approximately 40 km west of the eastern margin of the Athabasca Basin (Figure 3.1) (Bishop et al.,
2016). The narrow, flat-lying, cigar-shaped unconformity deposit is polymetallic with a complex highly-
variable geochemistry, mineralogy and structural geology. Elements such as As, Mo, Se and Zr can be
problematic during mining, milling and tailings management, and have been identified as elements of
concern (EOC; Bishop et al., 2016). Ongoing delineation and operational drilling of the Phase 1 pods has
improved the spatial coverage of the deposit allowing for renewed insight into the ore-forming system
(Figure 3.2). A robust understanding of the chemical, structural and temporal controls on the spatial
distributions of the EOCs is desired to improve forecasting in support of mining and milling operations.
Since the initial discovery of the Cigar Lake orebody in 1981, many geological, mineralogical,
geochemical and geochronological aspects of the deposit have been studied in detail (e.g. Bruneton, 1987,
1993; Percival and Kodama, 1989; Cumming and Krstic, 1992; Landais et al., 1993; Pacquet and Weber,
1993; Pagel et al., 1993; Percival et al., 1993; Philippe et al., 1993; Reyx and Rulmann, 1993; Toulhoat
and Beaucaire, 1993; Cramer and Smellie, 1994; Janeczek and Ewing, 1992, 1994; Cramer, 1995; Fayek
and Kyser, 1997; Fayek et al., 1997, 2000, 2002). However, research focused directly on the uranium ore
and associated metals, and the underlying paragenetic model is sparse (e.g. Bruneton, 1987; Reyx and
71
Rulmann, 1993), with studies typically lacking access to high-grade uranium ore samples. Quality
polymetallic samples are particularly challenging to obtain due to the heterogeneity of the ores and high
clay content that masks the sulphides and arsenides in drill core. Establishing the age of initial U
emplacement for the deposit has also been problematic because of significant disruption of the U–Pb
system. Significant Pb-loss from U-bearing minerals resulted from several extensive alteration events.
Various ages have been documented for the deposit using U–Pb and K–Ar geochronology on the uranium
ores and clay minerals respectively (e.g. Cumming and Krstic, 1992; Philippe et al., 1993; Fayek et al.,
1997, 2000, 2002). However, research on the mineralogy and paragenesis of the deposit (e.g. Bruneton,
1987; Reyx and Rulmann, 1993) was undertaken separately from the geochronological studies (e.g.
Cumming and Krstic, 1992; Philippe et al., 1993; Fayek et al., 1997, 2000, 2002) resulting in speculative
integration for the metallic gangue minerals.
Figure 3.1: Location of the Athabasca Basin and Cigar Lake (yellow star). Also shown are the
locations of several other high-grade unconformity-type uranium deposits (black squares) and
northern communities (white circles).
72
Figure 3.2: Air photograph of the Cigar Lake mine site with outlined study area, the Phase 1
Cigar Lake ore body. The Phase 1 deposit is divided into the East Pod and the West Pod.
In this study, a textural paragenetic interpretation is integrated with semi-quantitative mineral
characterization and empirical geospatial modelling of the phase 1 Cigar lake deposit to reevaluate the
evolution of the ore-forming system and subsequent alteration events. Geochronological characterization
using U–Pb and Pb–Pb systematics provides constraints on the absolute ages of various uranium ores,
arsenides, sulphides and non-metallic gangue minerals. Stable isotope chemistry is used to characterize
fluids responsible for the formation and subsequent alteration of the orebody.
3.3 Geological Setting
3.3.1 Regional Geology
The Cigar Lake uranium deposit occurs in the Athabasca Basin near the unconformity between
the underlying crystalline basement rocks and the Athabasca group sediments (Figure 3.3). This Paleo to
Mesoproterozoic intracratonic basin unconformably overlies the remnants of two orogenic belts, the ca.
1.9 Ga Taltson Magmatic Zone to the West and the younger, ca. 1.8 Ga Trans-Hudson to the East
(Ramaekers, 1980).
73
Figure 3.3: Geological map of northern Saskatchewan with the stratigraphic divisions of the
Athabasca Group and basement geology. Major unconformity-related U deposits (squares),
including the Cigar Lake deposit (star), are indicated (Modified from Card et al., 2007, Ramaekers
et al., 2007).
The initial accommodation and subsequent exhumation of the basal Athabasca Group (Manitou
Falls and Fair Point Formations) occurred in NE-SW trending Hudsonian (1.7 Ga) basement faults
(Armstrong and Ramaekers, 1985; Kyser et al., 2000). Rapid uplift during the Trans-Hudson Orogeny
provided the siliciclastic input for the Athabasca Basin beginning at 1.75–1.7 Ga (Armstrong and
Ramaekers, 1985). Basin fill consists predominantly of unmetamorphosed quartz arenitic sandstone and
conglomerate overlain by siltstone, mudstone and dolostone (Ramaekers, 1990). The depositional
environment of the flat-lying, upward-fining red-bed succession is interpreted as major river systems and
near-shore shallow marine shelf environments (Ramaekers, 1990). The end of sediment deposition, based
on Re–Os dating of the Douglas Formation, occurred after 1,540 Ma (Creaser & Stasiuk, 2007).
The crystalline basement rocks underlying the Athabasca Basin can be divided into three
lithotectonic zones: (1) the Taltson magmatic zone that underlies the westernmost side of the basin, (2)
74
the Rae Province that underlies the central basin and (3) the Hearne Province on the easternmost side
(Card et al, 2007). The Hearne Province, which underlies Cigar Lake, typically comprises granites and
granitoid gneisses unconformably overlain by and folded with supracrustal, upper amphibolite facies,
metasedimentary gneisses (Tran and Smith, 1999; Card et al., 2007). The metasedimentary gneisses range
compositionally from psammite to pelite with discontinuous calc-silicate, calc-arkose, arkose and
granitoid segregations (Tran and Smith, 1999; Card et al., 2007). Cigar Lake is situated along the
gradational contact between the Wollaston and Mudjadik Domains (Tran and Smith, 1999). The
Wollaston Domain comprises steeply dipping, doubly plunging north-easterly oriented folds, in contrast
to the Mudjatik Domain with non-linear dome and basin refolded deformational features (Tran and Smith,
1999). The granitoid gneiss basement to the Wollaston Supergroup has ages between 2.59–2.56 Ga
(Annesley et al., 1999), whereas the Wollaston Supergroup is constrained by the Wathaman Batholith
dated at 1.865–1.850 Ga (Ray and Wanless, 1980; Van Schmus et al., 1987).
3.3.2 The Cigar Lake Deposit
The Cigar Lake uranium deposit occurs 410 to 450 m below surface within the Athabasca Basin
along the unconformity between the Helikian Athabasca Group sediments and the underlying Aphebian
graphitic metasediments of the Wollaston Domain (Bruneton, 1987; Bishop et al., 2016) (Figure 3.4, 3.5).
The narrow, flat-lying, cigar-shaped deposit is approximately 1,950 m long, 20 to 100 m wide and has a
maximum thickness of 13.5 m, with an average thickness of approximately 5.4 m (Bishop et al., 2016).
Basement-hosted root-mineralization and perched-mineralization are lower in grade and spatially
confined to structures resulting in limited mining potential (Bishop et al., 2016). As of December 31,
2015, both Phase 1 and Phase 2 of the Cigar Lake deposit have a combined reserve of 100,501 tonnes
(221.6 M lbs.) U3O8 and a total resource (measured, indicated and inferred) of 48,262 tonnes (106.7 M
lbs.) U3O8 (Bishop et al., 2016).
In the Cigar Lake area, the basin fill is unmetamorphosed quartz arenitic sandstone and
conglomerate of the Manitou Falls Formation (MF). Only the MFd, MFc and MFb are observed proximal
75
to the deposit (Bruneton, 1987). Basal conglomerates of the MFb are observed locally. The sandstone
units represent a finning upward, transgressive succession.
Directly underlying the deposit, the rocks are moderately graphitic (3–10%), locally anatectic,
cordieritic protomylonitic pelites that have undergone extensive shearing and local semi-brittle fault
reactivation (Bruneton, 1987; Andrade, 2002). The regional foliation in the area strikes northeast,
however the shear zone underlying the deposit is oriented east-west (Bruneton, 1993). The local, roughly
10 km long, reverse dextral shear zone has been interpreted to be Hudsonian (Bruneton, 1987, 1993;
Andrade, 2002). Local foliation-concordant discontinuous lenses of amphibole and pyroxene bearing
calcic-magnesium rich gneisses and granulites occur adjacent to the shear-zone (Bruneton, 1993). The
Cigar Lake deposit is situated directly on top of an unconformity structure-contour high interpreted as a
pre-Athabasca paleo-topographic ridge (Bruneton, 1993).
Figure 3.4: Schematic illustration of the Cigar Lake deposit and surrounding alteration. Modified
from Jefferson et al., (2007) and Cameco (2015) with drill core data and field observations.
Secondary dispersion exploration pathfinder elements from Holk et al. (2003) and Drever (2010).
76
Figure 3.5: East Pod section along line 10731 (mine grid) showing orebody facies and structural
interpretation. Orebody outline at 1% U3O8 cutoff highlighted in red.
The Cigar Lake deposit is located within an extensive hydrothermal alteration zone characterized
by interstitial illitization, in contrast to the regional dickite, forming a sub-cropping chimney around the
deposit (Figure 3.4; Wasyliuk, 2002). Alteration intensifies 100–200 m above the unconformity with
intense bleaching, patches of fine-grained interstitial and fracture controlled sulphides, and structurally
controlled quartz dissolution and clay alteration (Bruneton, 1987; Andrade, 2002). Intense structure in the
basal sandstone (~100 m) and sagging sedimentary marker horizons suggest extensive volume loss and
the development of collapse structures from the mineralizing system (Andrade, 2002). Proximal to the
mineralization, the clay alteration becomes intense around the periphery of the deposit (Figure, 3.5). The
orebody is commonly capped by hematite-rich massive mixtures of illite, muscovite and kaolinite clays
77
with local Fe–Mg chlorite (Bruneton, 1987; Percival and Kodama 1989; Philippe et al., 1993). Local
paragenetically late induration of clay by calcite and siderite is common (Bruneton, 1987). An extensive
argillitized basement alteration halo of Mg-chlorite (sudoite and chlinochlore) and Mg- and Fe-rich illite
extends more than 50 m below the deposit, masking the pre-Athabasca paleoweathering (Bruneton, 1987;
Percival and Kodama, 1989). Graphite destruction directly below the deposit is extensive with traces of
remobilized carbonaceous material occurring proximal to the mineralization as irregular aggregates of
bituminous carbon or 1–5 mm black flakes that form hydrocarbon buttons (Bruneton, 1987; Landais et
al., 1993).
The mineralization at Cigar Lake predominantly contains the uranium oxide uraninite and the
uranium silicate coffinite (Bruneton, 1987; Janeczek and Ewing, 1992; Reyx and Ruhlmann, 1993;
Cramer and Smellie, 1994). Uraninite forms euhedral, radiating, botryoidal and massive aggregates and
occurs in association with Ni–Co arsenides, sulpharsenides and sulphides (Bruneton, 1987; Reyx and
Ruhlmann, 1992; Cramer and Smellie, 1994). Reyx and Ruhlmann (1993) suggested that the first major
stage of mineralization, responsible for the unconformity-hosted uranium was a polyphased hydrothermal
system that deposited U–Ni–Co–As–S–Bi–Cu–Zn and Pb. Two subsequent stages of U crystallization
have been identified reflecting mobilization of Stage 1, the primary mineralizing event. Stage 2 uranium
oxide is associated with secondary Ni–Co arsenides, sulpharsenides and Fe–Cu-rich sulphides (Bruneton,
1987; Phillipe et al, 1992). The third stage occurs with extensive Fe-oxides replacement and is
responsible for coffinitization and the redistribution of U as perched mineralization (Bruneton, 1987;
Reyx and Ruhlmann, 1993).
Geochronology of uranium deposits in the Athabasca Basin indicates that there have been three
major fluid events at ca. 1590 Ma (the initial and main basin-wide mineralizing event), ca. 950 Ma and ca.
300 Ma that coincide with the three stages of mineralization, the latter two being dominantly mobilization
of the primary mineralization (Bruneton, 1987; Reyx and Ruhlmann, 1993; Philippe et al., 1993; Fayek
and Kyser, 1993; Fayek et al., 1997; Alexandre et al., 2009). The oldest reported age for the major
78
mineralizing event (Stage 1) at Cigar Lake is 1468 Ma, but this is interpreted to be a minimum age for
mineralization (Fayek et al., 2000). Numerous younger ages have been reported for the deposit (e.g.
Cumming and Krstic, 1992; Janeczek and Edwing, 1992; Philippe et al., 1993; Fayek et al., 1997; Fayek
et al., 2002) and Pb-loss has consistently been reported for the deposit and likely resulted from episodic
hydrothermally-enhanced volume diffusion (e.g. Janeczek and Edwing, 1992; Fayek et al., 1997). Clay
mineral dating using K–Ar has yielded similar results for illite (1255–1148 Ma) and sudoite (850 Ma) due
to the episodic hydrothermal fluids that have accessed the deposit along structures (Percival et al., 1993).
3.4 Methods
Extensive on-going operational and delineation drilling at Cigar Lake has vastly improved the
spatial coverage of geochemical data. Prior to this study, Cameco Corporation (Cameco), 50% owner and
operator of the Cigar Lake mine, possessed an extensive multi-element whole-rock geochemical dataset
containing over 10,000 samples (Appendix A). The geochemical dataset was utilized to evaluate the
empirical spatial distribution of selected elements and stoichiometric mineralogy. Leapfrog Geo software
was utilized to determine the empirical spatial distribution of Al2O3, As, Co, Mo, Ni, Se, Fe2O3, K2O,
MgO, Pb, U3O8 and Zn. Geostatistics based variography was utilized to identify geochemical domaining
and structural controls.
The mineralogy of the ores was initially characterized by analyzing 53 spatially and
geochemically representative samples with X-ray diffraction (XRD; Appendix B, D). Samples were
coarsely crushed with greater than 80% of the material passing through a 2 mm sieve. A split subsample
was collected and hand ground with a mortar and pestle. The analysis was undertaken at the Queen’s
Facility for Isotope Research (QFIR) at Queen’s University in Canada with an Xpert Pro Philips powder
diffractometer equipped with a cobalt X-ray tube and an X’celerator area detector. The X-ray beam was
in Bragg-Brentano configuration. To minimize the effects of preferred mineral orientation, samples were
loaded into jacket-style holders and spun during the procedure. Mineral identification was performed by
pattern-matching using PANalytical HighScore software. Semi-quantitative mineral concentrations were
79
determined using the reference intensity ratio method (RIR) provided with PANalytical HighScore
software.
Petrographic analyses (reflected and transmitted light microscopy) were performed on a select
suite of samples (50) collected from active on-going drilling to enhance the coverage of mineralogical
data. Petrographic sections provided in-situ mineral relationships and textures to assess the paragenesis
and ore forming processes of the Cigar Lake ore deposit.
X-ray diffraction (25) and petrographic samples (12) with mineral phases and textural
relationships of interest were scanned with a Mineral Liberation Analyses (MLA) equipped Scanning
Electron Microscope (SEM) for further mineral confirmation and textures, and improved mineral
quantification (Appendix E). Coarsely crushed sample material (>80% passing through 2 mm sieve) was
mounted into epoxy, polished, carbon coated and scanned using a MLA 650 FEG ESEM at QFIR. Back-
scattered electron (BSE) images and energy-dispersive (EDS) spectra facilitated mineral identification
and were used to establish a customized Cigar Lake EDS mineral library for MLA.
On a refined sample set, electron microprobe analysis (EMPA) and laser ablation inductive
coupled plasma mass spectrometry (LA-ICP-MS) were used to measure the chemical compositions and
element deportment within selected mineral phases (Appendix C, L). The EMPA work was undertaken on
select minerals including uraninite, coffinite, gersdorffite, cobaltite and niccolite using a JEOL JXA-8230
equipped with five wavelength dispersive spectrometers (WDS). Uraninite was analyzed using 15 kV
accelerating potential, 100 nA beam current and a 7 μm beam diameter. Acquisitions of coffinite were
collected with a 15 kV accelerating potential, 10 nA beam current and a 3.5–7 μm beam diameter.
Arsenides were analyzed using a 20 kV accelerating potential, 30 nA beam current with a focused beam.
A Thermo Scientific ELEMENT XR LA-ICP-MS was used to analyze for trace elements on
uraninite and coffinite (Appendix L). Laser ablation ICP-MS element concentrations were quantified
using external glass standards NIST610, NIST612 and an in-house galena calibrated to NIST610. A
typical sample set started with the NIST glasses and calibrated galena, followed by ten sample analyses.
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Internal standardization was performed by normalizing measured intensities to the average U
concentration obtained with the EMPA.
Uranium minerals were dated by U/Pb isotope ratios acquired by LA on polished thin sections
and epoxy grain mounts using a NWR-193 Laser Ablation Platform system (Appendix J). In-situ ablation
was performed on the samples using a 15- to 50-μm spot size with 40 to 50 % laser output power at a
frequency of 3 Hz. For each sample, 204Pb, 206Pb, 207Pb, 235U, and 238U counts were determined with a
ThermoFinnigan NEPTUNE high-resolution multi-collector inductively-coupled plasma mass
spectrometer (HR-MC-ICP-MS), at QFIR following the method of Chipley et al. (2007). The analyses
were bracketed by gas blanks and an in-house uraninite standard. All intensities and resulting isotopes
ratios were blank, Hg and standard corrected.
A Thermo Scientific ELEMENT XR LA-ICP-MS was used to analyze Pb isotopes for minerals
including gersdorffite, cobaltite, chalcopyrite, pyrite, sphalerite, bornite, chalcocite and hematized clay
(Appendix K). The analyses were bracketed by gas blanks and an in-house galena calibrated to NIST610.
A typical sample set started with the NIST610 calibrated galena, followed by ten sample analyses. All
intensities and resulting isotope ratios were blank and Hg isotope corrected.
Two uraninite samples were mechanically drilled and the O-isotopes were analyzed. Oxygen was
extracted from 5 mg samples at 550–600°C according to the conventional BrF5 procedure of Clayton and
Mayeda (1963) and analyzed via dual inlet on a Thermo-Finnigan DeltaPlus XP Isotope-Ratio Mass
Spectrometer (IRMS). δ18O values are reported using the delta (δ) notation in units of permil (‰) relative
to Vienna Standard Mean Ocean Water (VSMOW) international standard, with a precision of 0.1‰. The
δ18O values for the fluids were calculated using the uraninite-water fractionation factor proposed by
Fayek and Kyser (2000).
Stable isotopic compositions were determined for S-bearing mineral phases. Fifteen S samples
targeting petrographic and XRD identified mineral phases were collected from within the ore zone (12)
and underlying basement rock (3). Samples were either whole-rock pulps (10) or mechanically drilled
81
extractions from thin-section offcuts (5). The mineralogy from whole-rock samples was quantified with
the XRD reference intensity ratio method (RIR; Hubbard and Snyder, 1988), MLA and stoichiometric
normative calculations. Sulphur samples were weighed into tin capsules and the sulphur isotopic
composition was measured using a MAT 253 Stable Isotope Ratio Mass Spectrometer coupled to a
Costech ECS 4010 Elemental Analyzer. δ34S values were calculated by normalizing the 34S/32S ratios in
the sample to the Vienna Canyon Diablo Troilite (VCDT) international standard. Values are reported
using the delta (δ) notation in units of permil (‰) and are reproducible to 0.3‰.
Stable isotopic compositions were determined for C-bearing mineral phases. Eleven carbonate
samples targeting petrographic and XRD identified mineral phases calcite and siderite collected within
the ore zone were analyzed for carbon isotopes. Except for one mechanically drilled sample extraction
(CAM085024), all siderite and calcite samples were extracted from whole-rock pulps. Sample
CAM085024 was drilled out from cross-cutting veins within uraninite crystals. The δ18O and δ13C values
of calcite were determined by reacting approximately 1 mg of powdered material with 100% anhydrous
phosphoric acid at 72°C for 4 hours. The CO2 released was analyzed using a Thermo-Finnigan Gas Bench
coupled to a Thermo-Finnigan DeltaPlus XP Continuous-Flow Isotope-Ratio Mass Spectrometer (CF-
IRMS). δ18O and δ13C values are reported using the delta (δ) notation in permil (‰), relative to Vienna
Pee Dee Belemnite (VPDB) and Vienna Standard Mean Ocean Water (VSMOW) respectively, with
precisions of 0.2‰. The δ18O values for the fluids were calculated using the calcite-water and siderite-
water fractionation factors proposed by Zheng (1999).
Two hydrocarbon buttons sampled from above (CAM085001) and below (CAM085002) the
East-Pod ore zone were analyzed for carbon and hydrogen isotopes. Sample CAM085001 was collected
within the intensely clay altered, indurated hematized facies, whereas CAM085002 was sampled from the
strongly clay altered basement regolith occurring in association with chalcopyrite and pyrite. The
hydrocarbon buttons were identified by petrography and confirmed with SEM. Samples were extracted
from thin section offcuts by mechanical drilling and weighed into tin and silver capsules for the respective
82
carbon and hydrogen analyses. The C isotopic composition was measured using a Costech ECS 4010
Elemental Analyzer coupled to a Thermo-Finnigan DeltaPlus XP Continuous-Flow Isotope Ratio Mass
Spectrometer (CF-IRMS). δ13C values are reported using the delta (δ) notation in units of permil (‰)
relative to Vienna Pee Dee Belemnite (VPDB) international standard, with a precision of 0.2‰.
Hydrogen samples were degassed for 1 hour at 100°C then crushed and loaded into a zero-blank auto
sampler. The hydrogen isotopic composition was measured using a Thermo-Finnigan thermo-combustion
elemental analyzer (TC/EA) coupled to a Thermo-Finnigan DeltaPlus XP Continuous-Flow Isotope-Ratio
Mass Spectrometer (CF-IRMS). δ2H values are reported using delta (δ) notation in permil (‰), relative to
Vienna Standard Mean Ocean Water (VSMOW), with a precision of 3‰.
3.4.1 U/Pb and Pb Isotope Systematics
Lead has four naturally occurring isotopes, 204Pb, 206Pb, 207Pb and 208Pb. The radiogenic Pb
isotopes, 206Pb, 207Pb and 208Pb form from the complex decay series of 235U, 238U and 232Th respectively. In
contrast, 204Pb (common Pb), is isotopically stable resulting in a constant concentration throughout
geological time in the absence of fractionation (Stacey and Kramers, 1975).
The Cigar Lake deposit is comprised predominantly of U-oxide minerals containing 238U and 235U
isotopes and the resulting 206Pb and 207Pb respectively, at a known decay rate. Due to the large ionic
radius difference between Pb2+ (1.29 Å) and U4+ (1.00 Å), Pb is incompatible within the crystallographic
structure of U-oxide minerals (Janeczek and Ewing, 1995). This results in rapid Pb diffusion from the U-
oxide minerals, three to six orders of magnitude faster than diffusion rates from silicate minerals
(Janeczek and Ewing, 1995). Episodic incursion by hydrothermal fluids can significantly enhance
radiogenic Pb loss and subsequent dispersion by U-oxide dissolution, recrystallization or coffinitization.
At Cigar Lake, late fluid incursion is responsible for kilometer-scale dispersion of radiogenic Pb away
from the deposit along regional structures and lithostratigraphic conduits (Holk et al., 2003).
There are two ways in which a mineral can acquire radiogenic Pb. In a closed system, the only
source of additional radiogenic Pb, is through the internal decay of U. The 206Pb/204Pb and 207Pb/204Pb
83
values increase as a function of their decay rate and time making a closed geochronometer. In contrast, in
open systems, Pb can be leached and mobilized from the U source. Furthermore, a rock can obtain Pb
from hydrothermal fluids where Pb has been mobilized. In these open systems, the isotopic proportions
provide details on the source and passage of the fluid, but compromise the geochronometer properties of
the U–Pb system.
3.5 Results
3.5.1 Mineralogy and Textural Paragenesis
The paragenesis is interpreted from textural relationships made from 50 petrographic sections,
25 epoxy grain mounts and from drillcore observations. The 3 stages of uranium crystallization and
alteration previously identified at Cigar Lake (e.g. Bruneton, 1987; Reyx and Ruhlmann, 1993; Fayek and
Kyser, 1993) has been expanded to 4 stages to reflect mineralogical, textural and chemical changes in the
orebody through the evolution (Figure 3.6).
3.5.1.1 Uranium Ore Mineralogy
At Cigar Lake, U occurs primarily as reduced tetravalent (IV) oxides and silicate minerals.
Uraninite (UO2) is the sole U-bearing mineral precipitated during the main Stage 1 mineralizing event,
displaying the highest reflectivity and occurring predominantly as botryoidal masses, and to a lesser
extent massive aggregates, veins and disseminated subhedral crystals (Figure 3.7). The millimeter to
centimeter-scaled botryoids coalesce to form radiating globular aggregates. Uniformly distributed radial
and polygonal shrinkage cracks, resembling desiccation cracks, occur within the primary uraninite
crystals suggesting U1 underwent dehydration during precipitation forming from the crystallization of
uraniferous gels (Figure 3.7A). Under BSE, even pristine botryoidal or subhedral U1 crystals display
some grey-scale mottling and chemical heterogeneity, indicative of alteration. Primary uraninite (U1) is
typically overgrown and intergrown by sulphides, arsenides and carbonates within a chlorite, illite matrix
(Figure 3.7A, B). Some U1 crystals are brecciated and overgrown by coeval, Stage 1, Cu-sulphides, Ni-
84
sulpharsenides and calcite indicating active tectonic faulting, dissolution induced collapse brecciation, or
hydraulic fracturing are syngenetic with the main mineralizing event. Native Cu has been documented
coating uraninite crystals suggesting cathodic-like crystallization of U and Cu (Fayek et al., 1997). Within
the samples analyzed, a strong association between Stage 1 chalcopyrite (CPY1), gersdorffite (GER1) and
U1 uraninite is observed.
Alteration of the initially emplaced U1 was substantial during the subsequent stages of the
paragenesis (Stage 2–4). Brecciated and altered uraninite crystals still contain traces of remnant U1
mineralization observed as spots with higher BSE reflectance, but these are sparse and have been variably
altered (Figure 3.7C, D). U1 crystals were strongly and almost ubiquitously altered resulting in enhanced
grey-scale mottling, observable under BSE. Uraninite dissolution and alteration results in irregular
embayed crystal boundaries (Figure 3.7C). Structure ranges from microfracturing to local cataclastic
brecciation responsible for the fragmentation of U1 crystals. Highly altered uraninite crystals occur as
remnant irregular bands with a ribbon-like texture. In less altered samples, vugs occur between uraninite
crystals indicating dissolution and chemical buffering from sulphides, sulpharsenides and arsenides
helped to preserve the uraninite by reducing U remobilization. Extensive Stage 3 alteration resulted in
recrystallization and hydration of crackle-brecciated uraninite crystals with strong BSE and optical
heterogeneity.
85
Ore StageMinerals Basement Basin Stage 1 Stage 2 Stage 3 Stage 4QuartzPlagioclaseK-FeldsparBiotiteGraphiteCordieriteGarnetZirconIllmeniteActinoliteChloritePyriteChalcopyriteCalciteHematiteQuartz (Detrital)IlmeniteMonaziteZirconHematiteQuartz overgrowthsKaolinite (Dickite)IliteChloriteUraniniteCoffiniteBoltwooditeUranophaneNiccoliteRammelsbergiteSkutteruditeGersdorffiteCobaltiteGlaucodotSe, Bi SulpharsenidesBravoiteChalcopyritePyrite/MarcasitePyrrhotiteBornite ChalcociteSphaleriteGalenaErytheriteAnnabergiteAerugiteQuartzIlliteSideriteCalciteKaoliniteChloriteLimoniteHematiteRutileHydrocarbons
Cig
ar L
ake
Ore
Bo
dy
Ura
niu
m
min
eral
s
Ars
enid
e an
d
Sulp
har
sen
ide
Sulp
hid
eA
rsen
ate
Gan
gue
Pre-Ore Post-Ore Alteration
Gra
ph
itic
Pel
ite
MFb
San
dst
on
e
Det
rita
lD
iage
nes
is
U1 U2 U3 U4 U5 U6
CA1 CA2
CPY1 CPY2 CPY3
GER1 GER2
PY1 PY2 PY3 PY4
SPH1
CPY4
PY0
BO1
CC1
COB1 COB2
86
Figure 3.6: Mineral paragenesis summarizing the relative timing of the major minerals within the
Cigar Lake Deposit. Dotted lines indicate minor occurrence. Red U1 denotes primary
mineralization whereas blue (U2–U6) indicates predominantly alteration and Pb-loss rather than
complete recrystallization.
Coffinite (U(SiO4)1-x(OH)4x), a tetravalent uranium silicate, is prevalent throughout the orebody.
Textural mineral relationships indicate that coffinite has formed in all stages of the paragenetic sequence,
although later than U1 in Stage 1. Coffinitization and coffinite crystallization intensified throughout the
evolution of the deposit, becoming dominant in Stage 4 of the mineral paragenesis. Coffinite typically
forms irregular anhedral crystal aggregates, sooty disseminations and feathery, slightly fibrous masses.
Direct coffinitization of the uraninite is prevalent along microfractures, around crystal boundaries or as
complete replacement (Figure 3.7D). Crystallization of coffinite occurs with sulphide, sulpharsenide and
arsenide overgrowths resulting from mobilization of U initially emplaced as uraninite. Coffinite is the
dominant U mineral in lower-grade ore demonstrating the mobilization, alteration and dispersion of
uranium from the high-grade core of the deposit. Trace amounts of brannerite (U(Ti, Fe)2O6) a tetravalent
oxide mineral containing REE, Ti and Fe-oxides, were identified with MLA/SEM but could not be
confirmed with XRD because of its low abundance. Brannerite occurs in association with coffinite as
neoform, irregular anhedral aggregates, within strongly clay altered ores. In agreement with Bruneton
(1987), only trace amounts of the total TiO2 concentration forms U–Ti minerals with most TiO2 forming
rutile, anatase and leucoxene.
87
Figure 3.7: BSE images of uraninite and coffinite. A) Botryoidal uraninite crystals with
symmetrical shrinkage cracks indicating U1 underwent dehydration and Ostwald ripening during
crystallization. Desiccation cracks are filled with CPY1 chalcopyrite and GER1 gersdorffite. B)
Brecciated uraninite with CA1 calcite microfracture fill. C) Altered uraninite with ribbon texture
and embayed crystal boundaries. D) Uraninite showing extensive coffinitization. E) Coffinitization
of uraninite with euhedral cubic galena grown within a vug indicating Pb-loss from the alteration.
F) Coffinite altered uraninite with a galena overgrowth resulting from extensive Pb-loss.
B A
F
C D
E
88
Oxidization of the deposit has resulted in local remobilization and subsequent precipitation as
neoform uranyl minerals observed in drillcore as argillaceous yellow to orange overprint. In thin-section,
uranyl minerals occur as microveinlets cross-cutting uraninite and coffinite. The only hexavalent uranium
minerals identified are boltwoodite (HK(UO2)(SiO4) .1.5H2O) and uranophane
(Ca(UO2)2(SiO3OH)2.5H2O). Uranyl minerals are rare in the Cigar Lake deposit, highlighting the overall
reduced state of the deposit.
3.5.1.2 Arsenides and Sulpharsenides
Cigar Lake mineralization occurs in association with Ni–Co arsenides and sulpharsenides
(Bruneton, 1987; Reyx and Ruhlmann, 1993). Petrography, XRD and MLA identified the following in
order of decreasing overall abundance: gersdorffite (NiAsS), cobaltite (CoAsS), niccolite (NiAs),
rammelsbergite (NiAs2), skutterudite ((Ni,Co,Fe)As3), glaucodot (Ni,Fe)AsS), erythrite
(Co3(AsO4)2.8H2O), annabergite (Ni3(AsO4)2
.8H2O) and aerugite (Ni9(AsO4)2AsO6). The arsenides and
sulpharsenides occur as prismatic euhedral to subhedral disseminations, crystal aggregates, or botryoidal
and colloform masses (Figure 3.8). Consistent with the mineral determination, the whole-rock
geochemistry confirms that the Phase 1 pod is dominated by 1:1 molar ratio, Ni–Co:As arsenides and
sulpharsenides: gersdorffite (NiAsS), cobaltite (CoAsS) and niccolite (NiAs). Furthermore, Ni-rich
mineral end-members dominate over their Co-rich varieties throughout most of the deposit.
At Cigar Lake, Ni–Co–As ores have complex textures and mineral relationships suggesting
several stages of crystallization. Stage 1 arsenides and sulpharsenides, comprising predominantly
gersdorffite (GER1) occur overgrowing botryoidal U1 crystals or intergrown within U1 shrinkage cracks
(Figure 3.7A). In contrast, Stage 2 arsenides and sulpharsenides occur as uraninite, arsenide and sulphide
overgrowths or disseminated within the chlorite, illite matrix (Figure 3.8B, C, D, E, F). Stage 2 arsenides
and sulpharsenides are also associated with coffinite. Variable reaction alteration textures are prominent
throughout the arsenide and sulpharsenide phases with crystal nucleation, parasitic overgrowths,
concentric zoning, skeletal dissolution and replacement textures resulting from element-
89
Figure 3.8: A) Gersdorffite GER1 nucleating on niccolite. Niccolite is partially consumed by the
parasitic overgrowth. B) Gersdorffite is overgrown with cobaltite which is a common texture
suggesting a transition from Ni–Co through the paragenesis C) Gersdorffite GER2 overgrowing
CPY1. CPY infills shrinkage cracks within botryoidal uraninite D) Gersdorffite GER overgrowing
chalcopyrite in association with coffinite. E) Gersdorffite GER2 overgrowing pyrite in with galena
inclusions. Gersdorffite is rimmed by coffinite. F) Skeletal and atoll textured cobaltite disseminated
throughout the chlorite clay matrix. Cobaltite COB2 crystals are overgrown by wispy coffinite.
A B
C D
E F
90
bonding competition in a dynamic environment with fluctuating S, As, Ni and Co activities. Gersdorffite,
the most prevalent As-bearing mineral in the Phase 1 deposit, is frequently observed nucleating on
niccolite and pyrite with subsequent parasitic overgrowth forming concentric crystals with vuggy cores
(Figure 3.8A, E).
Increasing S and decreasing As activity resulted in the following crystallization and alteration
series: skutterudite (NiAs3) – into rammelsbergite (NiAs2) – into niccolite (NiAs) – into gersdorffite
(NiAsS) – into bravoite (NiS2). This series has been described extensively by Bruneton (1987) and Reyx
and Ruhlmann (1993) and reflects the overall evolution of the hydrothermal fluids. In contrast, decreasing
S activity and increasing As activity is also observed and resulted in the opposite crystallization and
alteration series: bravoite (NiS2) – into gersdorffite (NiAsS) – into niccolite (NiAs) – into rammelsbergite
(NiAs2) – into skutterudite (NiAs3). This reaction series variation reflects a dynamic multi-fluid mixing
environment with highly variable hydraulic conductivity, resulting in fluctuating S, As, Ni, and Co
activities and varying redox conditions.
Concurrent with the evolving and locally fluctuating As and S activities, Ni and Co readily
substituted for one another forming a solid-solution. Cobaltite is the dominant Co-bearing mineral within
the deposit. Texturally, cobaltite is similar to gersdorffite occurring as euhedral to subhedral crystal
dissemination, or crystal aggregates within the chlorite, illite matrix. However, unlike gersdorffite,
cobaltite was not observed in proximity to U1 uraninite. Cobaltite occurs as late Stage 1 and Stage 2
overgrowth on sulphides and arsenides often in association with coffinite. Textural relationships suggest a
transition from Ni to Co, during the waning stages of successive hydrothermal events.
Only minor, local occurrences of arsenates, including erythrite, annabergite and aerugite, were
identified, highlighting the overall reduced state of the deposit. Arsenates occur along the crystal
boundaries as alteration products of arsenides and sulpharsenides, typically contributing <3% of the total
arsenic concentration in each sample. Texturally, arsenates demonstrate late crystallization forming
91
during Stages 3 and 4 of the paragenetic sequence as a result of late incursion by oxidized fluids along
structures.
3.5.1.3 Sulphides
The dominant sulphide minerals identified are chalcopyrite (CuFeS2), pyrite (FeS2), bornite
(Cu5FeS4), sphalerite (ZnS), galena (PbS), chalcocite (Cu2S), pyrrhotite (Fe1-xS) and bravoite (NiS2)
listed in order of decreasing overall abundance (Figure 3.9). The main Cu phase is chalcopyrite with
primary, Stage 1, chalcopyrite (CPY1) occurring as uraninite shrinkage crack intergrowths, uraninite
overgrowths and anhedral masses (Figure 3.7A). Chalcopyrite CPY1 is observed with gersdorffite
overgrowths, confirming the relative timing as an early Stage 1 mineral in contrast to the paragenetic
sequences reported by Bruneton (1987) and Reyx and Ruhlmann (1993) (Figure 3.8D; Figure 3.9A,B,C).
Subsequent Stage 2 and Stage 3 chalcopyrite (CPY2, CPY3), occurs as anhedral aggregates, blebs,
overgrowths, veinlets and disseminated crystals within the chlorite, illite matrix in association with
coffinite (Figure 3.9D, E). Chalcopyrite is prone to pyrite replacement, particularly common on the north
and northwest ends of both Phase 1 pods, suggesting increasing S and decreasing O2 fugacities (Figure
3.9A, B).
Pyrite is a common mineral phase within the deposit with Stage 1 pyrite (PY1) occurring as
subhedral to euhedral disseminated crystals and anhedral crystal aggregates within a chlorite, illite matrix.
Subsequent stages of pyrite (PY2, PY3, and PY4) occur as euhedral to anhedral disseminations, sooty
disseminations, overgrowing aggregates, and as a common replacement of chalcopyrite (Figure 3.9B).
Pyrrhotite occurs locally within the deposit as blades and laths with boxwork-like texture, and as
chalcopyrite-pyrite exsolutions and overgrowths (Figure 3.9A). The following crystallization and
alteration series is observed and reflects increasing S and decreasing O fugacities: chalcopyrite (CuFeS2)
– into pyrite (FeS2) – into pyrrhotite (Fe1-xS).
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Figure 3.9: A) Chalcopyrite replacement by pyrite and overgrowth by pyrite PY1 and pyrrhotite
suggesting increasing S and decreasing O fugacities. B) Chalcopyrite replacement by pyrite with
cobaltite COB2 overgrowth. C) Altered chalcopyrite CPY1 with more pristine gersdorffite GER2.
D) Remobilized and recrystallized pristine chalcopyrite CPY2 overgrowing gersdorffite within
strongly altered and coffinitized uraninite. E) Chalcopyrite CPY3 overgrowing calcite CA2. F)
Wispy coffinite occurring in association with Stage 3 minerals bornite and galena.
93
Sphalerite is the only Zn bearing mineral identified within the deposit. Sphalerite occurs as
anhedral aggregates and blebs often disseminated throughout the chlorite, illite matrix. Textural mineral
relationships and associations suggest sphalerite crystallized during the waning stages of the main
mineralizing event (Stage 1).
Galena typically occurs as subhedral to euhedral crystals and is more prevalent in association
with higher-grade U samples (Figure 3.9F). The main source of Pb within the deposit is from the decay of
U, remobilized from uraninite since formation. Galena commonly occurs as overgrowths on uraninite or
intergrown as confined crystals within uraninite microfractures and shrinkage cracks.
Bornite occurs in association with relatively monometallic, high-grade U ores. Bornite forms
anhedral aggregates and laths commonly overgrowing uraninite or disseminated within the chlorite, illite
matrix often in association with galena (Figure 3.9F). Bornite forms a solid solution with chalcopyrite
demonstrating bornite-chalcopyrite exsolution textures. Chalcocite occurs in samples containing relatively
high concentrations of Cu (> 7 wt. %) commonly in association with bornite. Chalcocite and bornite are
interpreted to be from the incursion of basinal fluids responsible for the dissolution of initially emplaced
chalcopyrite and recrystallization during Stage 3 of the mineral paragenesis. The following crystallization
and alteration series is purposed for the Cu-phase as a function of increasing Cu and decreasing S
activities: chalcopyrite (CuFeS2) – into bornite (Cu5FeS4) – into chalcocite (Cu2S).
3.5.2 Mineral Geochemistry
3.5.2.1 Uranium mineral chemistry
The chemical compositions of U minerals are reported in Table 3.1 and their compositional
variations illustrated in a SiO2–CaO–MnO–FeO vs. chemical U/Pb age diagram (Figure 3.10A). The CaO
and SiO2 concentrations of uraninite crystals is an alteration index and can be used to differentiate
paragenetic stages of uraninite crystallization (e.g. Fayek and Kyser, 1993; Fayek and Kyser, 1997; Fayek
et al., 1997).
94
The most unaltered and earliest generations of uraninite, typically U1 and U2 uraninite, are
characterized by high UO2 (78.65–82.16 wt. %), high PbO (13.73–15.81 wt. %), and low CaO (0.66–1.35
wt. %), FeO (0.04–0.22 wt. %), MnO (nil–0.08 wt. %), and SiO2 (0.12–0.23 wt. %) contents. Variable
MoO3 contents occur in uraninite with concentrations up to 0.46 wt. %. Elevated MoO3 (~ >0.2 wt. %)
concentrations in the uraninite coincide with high Pb contents (~>14.88 wt. %) suggesting the entrapment
of Mo was syngenetic with the main U mineralizing event. The U–Pb chemical ages calculated using the
method of Bowles (1990) for these relatively unaltered, early uraninite crystals range from 1353 to 1150
Ma (Figure 3.10A).
Table 3.1: Average EMPA results for uraninite and coffinite.
Mineral Uraninite Coffinite
Oxide wt.
(%) (n=90) DL (n=74) DL
UO2 78.65-85.21 0.2 63.97-78.15 0.3
ThO2 <DL-0.026 0.03 <DL-0.42 0.09
PbO 9.11-15.81 0.05 <DL-2.28 0.1
Y2O3 <DL-0.32 0.05 <DL-1.89 0.2
Ce2O3 <DL 0.05 <DL-0.96 0.2
Gd2O3 <DL-0.11 0.05 <DL-0.50 0.2
Dy2O3 <DL-0.17 0.1 <DL-0.58 0.3
Yb2O3 <DL 0.06 <DL 0.2
SiO2 0.12-0.65 0.02 8.97-17.92 0.08
TiO2 <DL-0.88 0.04 <DL-1.91 0.1
ZrO2 <DL 0.04 <DL-1.36 0.2
MoO3 <DL-0.46 0.03 <DL-0.18 0.09
FeO 0.041-0.50 0.03 <DL-0.82 0.09
MnO <DL-0.16 0.03 <DL-0.14 0.09
CaO 0.66-1.93 0.03 0.45-3.6 0.07
SeO2 <DL-0.23 0.02 <DL-0.34 0.06
Total (%) 96.38-98.46 85.01-92.12
Samples (n) refers to the number of spots analyzed on uraninite and coffinite crystals.
Corresponding detection limits (DL) are listed adjacent to the range. <DL indicates
that a given oxide was below detection limit
95
Figure 3.10: A) Bivariate diagram of EMPA data showing the linear relationship between chemical
U–Pb ages and uraninite alteration elements: SiO2, CaO, MnO and FeO. The U–Pb chemical ages
are calculated using the method of Bowles (1990). B) Chondrite normalized REE in uraninite
showing LREE depletion and HREE enrichment. C) Chondrite normalized REE in coffinite
showing a simillar REE pattern as uraninite. All REE contents were collected with LA-ICP-MS and
are normalized to C1 chondrite from McDonough and Sun (1989).
Brecciation, alteration and to a lesser extent recrystallization of the uraninite occurs in association
with Pb-depletion and Ca, Mn, Fe, Si enrichment. A continuum of Ca, Si, Mn and Fe absorption and
concomitant Pb-loss reflect variable alteration intensities during multiple Pb-loss and alteration events.
Subsequent uraninite generations (U3 and U4), are characterized by high UO2 (81.19–85.21 wt. %) and
intermediate PbO (9.11–13.75 wt. %), CaO (0.81–1.93 wt. %), FeO (0.18–0.50 wt. %), MnO (0.05–0.16
A
C B
96
wt. %) and SiO2 (0.17–0.65 wt. %) contents. Recrystallized, brecciated and altered uraninite crystals have
younger U–Pb chemical ages ranging from 1143 to 761 Ma.
Coffinite is characterized by relatively low UO2 (63.97–78.15 wt. %) and PbO (nil–2.28 wt. %)
contents, and high CaO (0.45–3.60 wt. %), FeO (nil–0.82 wt. %), MnO (nil–0.14 wt. %), and SiO2 (8.79–
17.92 wt. %) contents. All the analyzed coffinite give very young chemical ages ranging between 242 to 0
Ma (Figure 3.10A).
REE contents from the various stages of uranium minerals were analyzed by LA-ICP-MS and a
select subset of the REE, including Ce, Gd, Dy, and Yb, were analyzed with EMPA (Figure 3.10 B, C).
REE concentrations within coffinite range from nil–424 ppm Ce, nil–987 ppm Gd, nil–1492 ppm Dy, nil–
342 ppm Yb, and within uraninite range from nil–8161 ppm Ce, nil–4329 ppm Gd, nil–5030 ppm Dy,
nil–1155 ppm Yb. Chondrite normalized REE patterns from all stages of uraninite show a similar pattern
with HREE-enrichment and LREE-depletion. The REE results from uraninite are similar to previously
reported patterns at Cigar Lake and elsewhere within the Athabasca basin (e.g. Fayek and Kyser, 1993,
1997; Mercadier et al., 2013). The similarity between all stages of uranium minerals indicates that there is
no appreciable fractionation of REE during brecciation, alteration and recrystallization of uraninite. Late
hydrothermal fluids responsible for uraninite modification either have a similar REE pattern as the initial
ore forming fluids, or uraninite is altered and recrystallized with minimal REE-fluid interaction.
In contrast, some generations of coffinitization and primary coffinite crystallization contain
higher overall REE contents and are less LREE depleted compared to uraninite. This REE pattern is more
typical of regional Athabasca Basin samples and barren alteration halo samples (Fayek and Kyser, 1997).
Therefore, late hydrothermal fluids responsible for coffinitization and primary coffinite crystallization had
a different REE composition resulting in modification of the U-mineral REE pattern.
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3.5.2.2 Arsenide and Sulpharsenide Mineral Chemistry
The most dominant As-bearing mineral within the Phase 1 Cigar Lake deposit, gersdorffite, has a
variable and complex chemistry with both cation and anion element substitution. Gersdorffite contains up
to 5.13 wt. % Co (Table 3.2; Figure 3.11A). Cobalt has a strong negative correlation with Ni (Pearson
correlation coefficient, PCC = -0.998), implying that Co readily substitutes for Ni forming a nearly
complete solid solution between gersdorffite and cobaltite. The substitution of Ni by Fe is limited in the
gersdorffite, with Fe contents ranging from nil to 0.68 wt. %. Bismuth and Se occur locally in gersdorffite
crystals with concentrations up to 12.44 wt. % Bi and 2.56 wt. % Se. Elevated Bi and Se occur
predominantly in association with a concentric, zoned, band proximal to the perimeter of the gersdorffite
crystals observed as brighter illumination in BSE. The S content is highly variable ranging from 11.21 to
19.60 wt. %, and there is an inverse S correlation with Bi (PCC = -0.98) and Se (PCC = -0.97). Arsenic
remains more consistent and is less depleted ranging from 39.74 to 45.99 wt. %. Bismuth and Se
enrichment is primarily a function of S-depletion suggesting that Se2- and Bi2- substitute for S2- within the
sulpharsenide (Figure 3.11B, C, D). Therefore, S fugacity within the mineralizing fluid may control the
uptake of Se and Bi (e.g. Huston et al., 1995; Layton-Matthews et al., 2008). Gersdorffite was the only
As-bearing mineral with elevated Cu and Ag contents, with up to 2.83 wt. % and 0.18 wt. %, respectively.
Copper and Ag have a strong correlation (PCC = 0.97) and typically occur in association with elevated Bi
and Se contents suggesting a coupled substitution may facilitate Cu and Ag uptake. The average chemical
formula from gersdorffite in this study is (Ni0.98Co0.04)As1.00(S0.97Bi0.02Se0.01).
Niccolite is relatively homogenous, showing minimal chemical variability and element
substitutions. Cation substitution by Co and Fe are only up to 0.59 wt. % and 0.80 wt. % respectively
(Table 3.2; Figure 3.11A). Locally, Bi, Sb and Se appear to substitute for As, with up to 3.12 wt. % Bi,
1.20 wt. % Sb, and 0.47 wt. % Se. Only minor S is incorporated into the mineral with concentrations
ranging from 0.14 to 1.06 wt. %. The average chemical formula for Cigar Lake niccolite in this study is
(Ni0.98Co0.01Fe0.01)As0.97S0.02Bi0.01.
98
Cobaltite is a Co-sulpharsenide end-member that forms a solid solution with gersdorffite. In the
deposit, cobaltite contains up to 14.05 wt. % Ni forming a nearly complete solid solution with gersdorffite
(Table 3.2; Figure 3.11A). In contrast to the other As-bearing minerals analyzed, cobaltite contains
relatively high Fe contents ranging from 0.17 wt. % to 1.43 wt. %. Bismuth, Se, and Sb are low with
maximum concentrations of 0.79 wt. %, 0.66 wt. % and 0.16 wt. % respectively. The average chemical
formula for Cigar Lake cobaltite in this study is (Co0.74Ni0.24Fe0.02)As1.00S1.00.
Table 3.2: Average EMPA results for arsenides and sulpharsenides.
Mineral Gersdorffite Niccolite Cobaltite
Wt. % DL (n=56) (n=8) (n=29)
Ni 0.03 24.7-36.17 42.84-44.83 3.96-14.05
Co 0.02 0.34-5.13 0.13-0.59 21.91-30.81
Fe 0.02 <DL-0.68 0.011-0.80 0.17-1.43
Cu 0.02 <DL-2.83 <DL-0.27 <DL-0.43
Ag 0.02 <DL-0.18 <DL <DL-0.056
As 0.1 39.74-45.99 52.21-55.31 44.10-45.38
Sb 0.02 <DL-1.00 0.022-1.2 <DL-0.16
Bi 0.05 <DL-12.44 <DL-3.12 <DL-0.79
S 0.03 11.21-19.60 0.14-1.06 18.83-19.75
Se 0.04 <DL-2.56 <DL-0.47 <DL-0.66
Samples (n) refers to the number of spots analyzed on uraninite and coffinite crystals.
Corresponding detection limits (DL) are listed adjacent to the range. <DL indicates
that a given oxide was below detection limit
99
Figure 3.11: A) Molar % (M %) proportions of S, Ni and Co showing the composition of the main
arsenide and sulpharsenides: gersdorffite, cobaltite and niccolite. B) Bivariate diagram with S and
As M% showing stoichiometric gersdorffite, cobaltite and niccolite. Some gersdorffite crystals are S
deficient. C) Selenium and Bi substituting for S within gersdorffite crystals. D) Selenium and Bi
showing a strong correlation in niccolite and gersdorffite.
3.5.3 Geochemical and Mineral Modelling
The overall geochemical characterization of the orebody was initially described by Bruneton (1987).
With 30 years of delineation and operational drilling the geochemical dataset has been vastly expanded
with improved spatial coverage. Implicit modelling with Leap Frog 3D software was used to create
geochemical grade shell interpolants for prominent ore forming elements, including U, Ni, Co, As, Zn,
Pb, Cu, K2O, Al2O3, MgO and ZrO2.
The orebody contains a massive high-grade (>40% U3O8) core, enveloped within a lower grade
shell (Bruneton, 1987). The magnitude and extent of the encapsulating clay is proportional to the grade
and thickness of the orebody. The Phase 1, East-Pod contains the highest-grade and most continuous
A
C
B
D
Cobaltite Gersdorffite
Niccolite
100
high-grade mineralization overlying the main basement, east-west oriented, strongly graphitic, semi-
brittle fault. On the west-end of the East-Pod, blowout high-grade mineralization coincides roughly with
crosscutting, or potentially Riedel northwest oriented faults (Figure 3.12A). High-grade lenses extending
out from the high-grade core display an east-northeast orientation coincident with the regional basement
foliation, major regional faults, and local crosscutting or Riedel east-northeast faults (Bruneton, 1993).
The lower grade West-Pod has an overall east-northeast orientation with high-grade ores focused locally
on the far west-side of the pod. The mineralogy of higher-grade U ores is dominated by uraninite whereas
the lower-grade ores contain a higher proportion of coffinite. The high-grade corridor contains the most
continuous and well preserved uraninite ores (Stage 1-2) in the deposit along the redox front (Figure
3.12B).
The high-grade U-ores have a strong spatial correlation with Cu, Mo, Se, and Pb contents. The
correlation with Pb is not surprising, reaffirming that most Pb is radiogenic (Bruneton, 1987).
Molybdenum and Se, at high concentrations (Mo > 5000 ppm, Se > 300 ppm), have an inverse
relationship along the high-grade U3O8 corridor (Figure 3.12C). Mo occurs as molybdenite and within
uraninite with concentrations up to 0.31 wt. % Mo in association with elevated Pb, suggesting coeval
crystallization within U1. In contrast, Se2- substitutes for S2- predominantly within Stage 2 sulphides and
sulpharsenides. The mineralogy, paragenesis and empirical spatial distribution suggest that elevated Se
along the high grade U3O8 corridor may reflect localities particularly effected by Stage 2 fluid incursion.
101
102
Figure 3.12: Leap Frog 3D implicit geochemical grade shells for the Phase 1 Cigar Lake orebody.
A) High-grade U mineralization (>40% U3O8) showing more continuous high-grade ore on the
East-pod. High-grade ore is more continuous above main east-west oriented graphitic fault and
occurs along the redox front between more oxidized ore to the southwest and more reduced
polymetallic ore to the northeast. B) Grade shell illustrating whole-rock Pb concentrations
normalized to U3O8. At concentrations between 10 and 16%, the grade shell is interpreted to
represent the most preserved Stage 1-2 ores within the deposit. These preserved ores show a strong
correlation with the high grade U3O8 corridor. C) Se and Mo showing a strong spatial correlation
with high-grade ore. Selenium and Mo have an inverse relationship along the high-grade corridor.
D) Copper and Ni showing spatial zoning in the deposit with Cu typically occurring to the
southwest and distal to the unconformity. E) Inverse relationship between Fe2O3 (total Fe) and As.
Hematite is the dominant mineral highlighting the transition between oxidized ores to the southwest
and more reduced polymetallic ores to the northeast. High-grade mineralization is concentrated
along the redox boundary.
In contrast, As, Ni, Co and Zn spatially show a limited correlation with U3O8 and are offset
towards the north, east and northeast ends of both Phase 1 pods (Figure 3.12D). These elements
demonstrate upgraded concentrations in association with east-west oriented semi-brittle, graphitic
basement faults. Consistent with the mineralogy, As has a strong correlation with Ni and Co. Particularly
well developed in the East-Pod, transitional metals are zoned from the southwest to northeast, with a
crystallization series of: Cu – to Ni – to Co – to Zn. This zonation is also observed vertically with Cu
typically occurring further from the unconformity than Ni, Co and Zn resulting from increasing S fugacity
and decreasing O fugacity.
The concentration of Fe2O3 (Total Fe) can be attributed predominantly to siderite and hematite
(Bruneton, 1987). Fe2O3 shows a strong inverse spatial relationship with As (Figure 3.12E), occurring
predominantly along the south and southwest ends of both Phase 1 pods. High-grade U3O8 is concentrated
directly between the more oxidized monometallic ores to the south and southwest and polymetallic ores to
the north and northeast.
Mineral stoichiometry and rock geochemistry were used to facilitate mineral quantification
throughout the deposit extending spatial coverage of mineralogical data (Chapter 2). Modelled normative
mineralogy of the Cu-phase highlights the effect of Stage 3 fluid incursion into the deposit (Figure 3.13).
Zoning is observed from southwest to northeast reflecting the mineral stabilities of the Cu phases and
103
corresponding S and Cu activities. Chalcocite and bornite occur in high concentrations on the south and
southwest sides of both the Phase 1 deposits.
Figure 3.13: Normative mineral proportions of Cu-bearing mineral phase showing a geochemical
zonation with bornite and chalcocite occurring to the southwest of chalcopyrite suggesting Stage 3
fluid flowed towards the northeast
3.5.4 Geochronology of U-Bearing Mineral Phase
3.5.4.1 207Pb/206Pb Systematics of U-Bearing Minerals
207Pb/206Pb dates obtained from uraninite typically range from 1350–1100 Ma (Figure 3.14). The
dates display a bivariate distribution with two main clusters from 1200–1130 Ma and 1350–1250 Ma.
Both clusters contain a range of uraninite crystal forms, and display subsequent brecciation and alteration.
Laser ablation transects across uraninite crystals reveal a high degree of variability within a single crystal.
On some crystals, the oldest 207Pb/206Pb ages are recorded along the crystal margins suggesting Pb
absorption from fluids with distally sourced 207Pb/206Pb values. The oldest 207Pb/206Pb date of 1352±2 Ma
104
was obtained on a low grade (0.792% U3O8) clay-rich sample with brecciated uraninite located on the east
end of the deposit. High clay concentration, resulting in lower permeability, facilitated radiogenic Pb
preservation within the Pb-incompatible uraninite structure. Uraninite and coffinite crystals with
207Pb/206Pb dates below 1100 Ma are poorly resolved and display a high degree of variability. Coffinite
has 207Pb/206Pb dates ranging from 1338±2 to 818±7 Ma. The dates obtained here are similar to previously
reported results from uraninite and coffinite within the Athabasca Basin and at Cigar Lake (e.g. Cumming
and Krstic, 1992; Philippe et al., 1993; Fayek et al., 2002).
Figure 3.14: A) Dates obtained using 207Pb/206Pb values showing two clusters between 1200–1130
Ma and 1350–1250 ma. The oldest 207Pb/206Pb date of obtained is 1352± 2 Ma. B) Chemical ages
from the EMPA U/Pb values calculated using the method outlined in Bowles (1990). Three
generations of U mineralization can be chemically discerned with the EMPA results.
Geochemically calculated dates appear to cluster at roughly at 100–0 Ma, 1020–900 Ma and 1210–
1150 Ma. These geochemically derived dates are younger than their corresponding U–Pb and 207Pb/206Pb dates providing further evidence for ubiquitous radiogenic Pb loss from the U-bearing
crystals.
3.5.4.2 U–Pb Systematics of U-Bearing Minerals
Uraninite and coffinite crystals analyzed with LA-MC-ICP-MS for U–Pb isotopes reveal
significant disruption of the U–Pb system coincident with alteration (Figure 3.15). Textural observations,
mineral associations, 207Pb/206Pb values and U/Pb values were used to discern four timings of systematic
resetting. The few points exhibiting Pb gain were filtered to prevent bias from ex-situ Pb sources.
A B
105
Uraninite crystals with pristine botryoidal crystal form, and no detectable alteration were plotted
on a concordia diagram (Figure 3.15A). The 11 points regressed together have an upper intercept of
1468±68 Ma and lower intercept of 755±93 Ma with a mean square weighted deviation (MSWD) of 0.65.
All 11 data points are highly discordant with respect to the concordia line indicating extensive Pb-loss.
The dates obtained from the regression provide a minimum age for the deposit and are similar to dates
reported by Fayek et al. (2000) at 1467±63 Ma and Fayek et al. (2002) at 1461±47 Ma. The lower
intercept is higher than 443±96 Ma and 437±71 Ma reported by Fayek et al. (2000) and Fayek et al.
(2002) respectively.
Uraninite crystals defined by weak alteration and brecciation and 207Pb/206Pb dates >1250 Ma
were plotted on the concordia diagram (Figure 3.15B). The 36 data points are regressed together and have
an upper intercept of 1270±10 Ma, lower intercept of 51±52 Ma and an MSWD of 1.4. This date
corresponds with the Cigar Lake U–Pb date of 1287±16 Ma reported by Cumming and Krstic (1992).
This date does not represent additional mineralization in the deposit, but rather the timing of a resetting
event of U1. The upper intercept date corresponds to timing of intrusion of the Mackenzie dykes
throughout the basin (LeCheminant and Heaman, 1989).
Uraninite that is associated with prominent alteration, brecciation, ribboned textured and has
207Pb/206Pb dates between 1130–1200 Ma has an upper intercept of 1163±25 Ma, lower intercept of
52±120 Ma and an MSWD of 2.7 (Figure 3.15C). This date corresponds with a previously reported age
by Fayek et al. (2002) with 1176±9 Ma and a lower intercept of 14±10 Ma.
Sample points with 207Pb/206Pb values below 7.59E-02 (1100 Ma) are poorly resolved showing
more variable U/Pb values. The 16 data points have an upper intercept of 947±57 Ma, lower intercept of
23±150 Ma and an MSWD of 22 (Figure 3.15D). This date corresponds with a previously reported
uraninite age of 876±14 Ma (Fayek et al., 2002) and K/Ar age for Fe-illite and Fe-kaolinite of 900±50 Ma
(Philippe et al., 1993).
106
Figure 3.15: Dates interpreted from U–Pb ICP-MS data: A) Concordia plot using only the most
unaltered botryoidal uraninite crystals. The sample points are discordant and the upper intercept is
interpreted as the youngest possible age of the deposit. B) U2 generation uraninite forming 33 data
point regression. C) U3 generation uraninite forming a 21 sample regression of the concordia plot.
D) Sample points with low Pb (207Pb/235U < 1.4 and 206Pb/238U < 1.4E-01) and low 207Pb/206Pb values
(<7.59E-02).
3.5.5 Pb Isotopes of Sulpharsenides, Sulphides and Non-Metallic Gangue Minerals
The Pb isotope values of selected minerals were analyzed by LA-ICP-MS to quantitatively
confirm paragenetic relationships (Table 3.3 and Figure 3.16). Stage 1 minerals include: chalcopyrite
(CPY1), pyrite (PY1), sphalerite (SPH) and cobaltite (COB1). Of the Stage 1 minerals, CPY1
chalcopyrite has 206Pb/204Pb (17–35), 207Pb/204Pb (13–23), and 208Pb/204Pb (31–49) values that are
anomalously non-radiogenic and indicative of common Pb (Figure 3.16A). These chalcopyrite crystals
demonstrate that some minerals within the orebody, even overgrowing botryoidal uraninite or within
D
B
C
A U2
U3
107
uraninite shrinkage cracks, may not have absorbed any radiogenic Pb preserving the source rock Pb
isotopic fingerprint. The 206Pb/204Pb values from CPY1 (206Pb/204Pb = 17–35) crystals are consistent with
proximal samples of granitic gneiss (206Pb/204Pb = 16.51–18.04), considered remobilized Archean
basement (Bruneton, 1993; Pagel et al., 1993), and silicified sandstone (206Pb/204Pb = 19.58–36.75) that
obtained Pb-closure prior to secondary dispersion (Holk et al., 2003).
Stage 1 PY1 (206Pb/204Pb = 34–86), SPH (206Pb/204Pb = 39–120), COB1 (206Pb/204Pb = 110–220)
and GER1 (206Pb/204Pb = 38–320) all show variable, but generally low 206Pb/204Pb values consistent with
distally sourced regional common Pb derived from unaltered pelite (206Pb/204Pb = 19.85–99.41) (Pagel et
al., 1993), but also with local radiogenic Pb from uraninite. Similarly, the 207Pb/206Pb values for PY1
(207Pb/206Pb = 0.21–0.50), SPH (207Pb/206Pb = 0.17–0.44), COB1 (207Pb/206Pb = 0.11–0.18) and GER1
(207Pb/206Pb = 0.09–0.60) are variable, but generally high and consistent with unaltered pelites (207Pb/206Pb
= 0.25–0.79) (Pagel et al., 1993), but also with local absorption of radiogenic Pb from the uraninite.
Within the Stage 1 minerals, elevated radiogenic Pb typically coincides with alteration suggesting
absorption as the primary mode of radiogenic Pb uptake.
108
Figure 3.16: A)207Pb/206Pb vs 206Pb/204Pb diagram of sulphide minerals from the orebody. B)
207Pb/206Pb vs 206Pb/204Pb diagram of arsenide minerals from the orebody. C) 207Pb/204Pb vs 206Pb/204Pb diagram of sulphides from various stages. The denoted stages correspond with the
mineral paragenetic sequence. D) 207Pb/204Pb vs 206Pb/204Pb of arsenides from the orebody. The
denoted stages correspond with the paragenetic sequence.
Stage 2 minerals include chalcopyrite (CPY2; 206Pb/204Pb = 79–2200), pyrite (PY2; 206Pb/204Pb =
100–2400) and gersdorffite (GER2; 206Pb/204Pb = 470–11000). All have relatively high concentrations of
radiogenic Pb (Figure 3.16). The 207Pb/206Pb values for CPY2 (207Pb/206Pb = 0.054–0.23), PY2 (207Pb/206Pb
=0.083–0.25) and GER2 (207Pb/206Pb = 0.050–0.11) give younger relative ages resulting from
predominantly in-situ uranium supported radiogenic Pb and only a trace contribution from distally
sourced common Pb.
B
C
A
D
109
Stage 3 minerals from the paragenetic sequence include chalcocite (206Pb/204Pb = 1400–2300), Pb
adsorbed to hematized clay (206Pb/204Pb = 1000–3700), galena (206Pb/204Pb = 4100–21000) and bornite
(206Pb/204Pb = 3400–88000). All have anomalously high radiogenic Pb in their structures. The 207Pb/206Pb
values from chalcocite (207Pb/206Pb = 0.045–0.067), hematized clay (207Pb/206Pb = 0.058–0.069), galena
(207Pb/206Pb = 0.048–0.095), and bornite (207Pb/206Pb =0.055–0.079) further demonstrate Pb-isotopes
dominated by in-situ uranium supported Pb.
Common Pb (204Pb) and Th derived Pb (208Pb) concentrations are all relatively consistent
displaying no discernable trend throughout the paragenesis. The 208Pb/204Pb values range from 22 to 210,
comparable with granitic gneiss (208Pb/204Pb = 36.62–40.32), unaltered pelites (208Pb/204Pb = 39.15–
125.42) and altered pelites (208Pb/204Pb = 44.77–190.86) proximal to Cigar Lake (Pagel et al., 1993).
Because the Pb systematics at Cigar Lake suggest a relatively open Pb system with both in-situ
Pb from the decay of U, and ex-situ Pb from regional source rocks, caution must be taken when assessing
absolute ages. Here, only minerals late in the paragenesis are used for Pb–Pb dating because they are so
highly radiogenic that externally sourced Pb is negligible (Figure 3.17). The least squares regression of
207Pb/204Pb vs. 206Pb/204Pb from bornite crystals correspond to a model age of 745±110 Ma (MSWD =
2.3). Lead adsorbed to hematized clay, with a 207Pb/204Pb vs. 206Pb/204Pb least squares regression, gives a
model age of 844±72 Ma (MSWD = 5.4). These model ages represent the closure of the Pb-system within
the mineral and therefore correspond with the time between initial emplacement of U1 and crystallization
of the minerals from a fluid during secondary dispersion of Pb from U1.
110
0
40
80
120
160
200
240
280
320
0 1000 2000 3000 4000 5000206Pb/204Pb
207P
b/2
04P
b
Age = 844±72 MaMSWD = 5.4
A
0
2000
4000
6000
8000
10000
0 50000 100000206Pb/204Pb
207P
b/2
04P
b
Age = 745±110 MaMSWD = 2.3
B
Figure 3.17: A) Pb–Pb model ages for hematized illitic clay. B) Pb–Pb model ages for bornite.
3.5.6 Stable Isotopes
The U1 uraninite crystals from the deposit have δ18O values of -17.5 and -8.6‰ (Table 3.4),
within the range reported by Fayek et al. (1997, 2002). Using the oxygen isotopic fractionation factor for
uraninite proposed by Fayek and Kyser (2000), and 200°C, the approximate temperature of mineralization
(Wilson and Kyser, 1987; Fayek and Kyser, 2000; Fayek et al., 2002), the uraninite crystals would have
been in equilibrium with fluids having δ18O values of -5.8 and +3.7‰.
3.5.6.1 Sulphur Isotope Systemic of Sulphides and Sulpharsenides
Samples collected within the ore body have δ34S values ranging from +1.4 to +14.6‰ (Table 3.4).
Overall, no discernable mineralogical trends were observed within the dataset, with all mineral phases and
their respective paragenetic stages displaying a similar range in δ34S values. However, sulphides from the
basement are significantly depleted in 34S with δ34S values of -22.7 and -20.0‰.
3.5.6.2 Carbon and Oxygen Isotopes in Carbonates
Calcite 1 (CA1) occurs as micro-fracture fill within U1 uraninite crystals (Table 3.4, Figure
3.18B). The δ13C and δ18O values range from -22.4 to -21.8‰ and +12.3 to +13.4‰, respectively. Wang
(2010) analyzed single phase fluid inclusions at Cigar Lake from euhedral quartz entrapped in a calcite
vein and obtained a homogenization temperatures of 187 and 285°C. Based on these temperatures, the
calculated δ18O values for a fluid in equilibrium with CA1 would range from +1.8 to +7.2‰, similar to
111
those reported for basinal fluids that formed the deposits elsewhere (e.g. Wilson and Kyser, 1987; Kotzer
and Kyser, 1990, 1992, 1993; Rees 1992; Alexandre et al., 2005; Cloutier et al., 2011).
Calcite 2 (CA2) and siderite are typically disseminated within intensely altered, commonly
hematite indurated clay or aggregates within vugs (Figure 3.18A). In contrast to CA1, these samples were
collected from lower grade ore (<20% U3O8). The δ13C and δ18O values for CA2 range from -14.0
to -11.5‰ and +21.1 to +22.6‰, respectively, whereas the δ13C and δ18O for siderite range from -18.8 to
-16.9‰ and +22.0 to +22.6‰, respectively. Siderite and CA2 appear co-precipitated and because both
minerals have similar isotopic compositions they are interpreted to have been deposited during the same
hydrothermal event. Microthermometry on fluid inclusions in siderite at McArthur River analyzed by
Kotzer and Kyser (1995) have low homogenization temperatures between 25 and 50°C. Based on these
temperatures the calculated δ18O values for a fluid in equilibrium with the siderite range from -12.7
to -5.5 ‰ and for calcite between -10.1 to -3.1‰.
Hydrocarbon buttons analyzed have low δ13C values between -31.0 and -28.4‰ (Figure 3.18 C,
D, E, F; Table 3.4). The δ2H values from the hydrocarbon buttons are low and highly variable at -257 and
-151‰ (Table 3.4). The hydrocarbon buttons are isotopically similar to bitumen samples previously
analyzed at Cigar Lake, which have δ13C and δ2H values of -31.2 to -27.3‰ and -237.2 to -229.5 ‰,
respectively (Landais et al., 1993).
112
Sample ID Hole Depth U-Mineral Fluid* C-Mineral Sulphide Stage
U3O8 δ18O δ18O δ13C δ18O δ2H δ18O δ34S
(wt.%) (‰) (‰) (‰) (‰) (‰) (‰) (‰)
200°C 25°C 50°C 187°C 285°C
CAM085043 SF766_12 426.7 U Massive Pitchblende 65.2 Uraninite (U1) -8.6 3.7
CAM085024 SF781_04A 434.3 U Hematite Massive Pitchblende 47.5 Uraninite (U1) -17.8 -5.8 Calcite (CA1) -21.8 12.3 1.8 6.1
80243 SF766_05 435.2 U Massive Pitchblende 21.9 Calcite (CA1) -22.4 13.4 2.9 7.2
34756 349 417.3 U Chloritized MFb Sandstone 1.11 Calcite (CA2) -11.5 21.9 -9.3 -3.8 Gersdorffite GER1 3.2
34758 349 418 Massive Sulphides 0.27 Calcite (CA2) -13.8 22.6 -8.6 -3.1 Gersdorffite GER1 12.4
59843 348 430.8 MFb Sandstone 0.04 Calcite (CA2) -14.0 21.1 -10.1 -4.6 Sphalerite SPH1 8.2
82311 369 430.9 U Hematized Clay 3.86 Siderite -18.8 22.0 -12.0 -6.1
80064 353 433 U Hematized Clay 1.96 Siderite -18.8 22.6 -11.4 -5.5
80067 353 433.7 Hematized Clay 0.05 Siderite -16.9 21.3 -12.7 -6.8
80075 353 440.2 U Chloritized MFb Sandstone 13.2 Chalcopyrite CPY2 7.4
50879 337 431.8 U Clay 3.19 Chalcocite CC1 1.4
80148 365 442.8 U Chloritized MFb Sandstone 4.25 Gersdorffite GER2 11.1
81673 363 426.5 Chloritized Clay 0.16 Chalcopyrite CPY2 7.0
80786 SF826_10 429.3 U Massive Pitchblende 26.4 Chalcopyrite CPY2 7.0
83235 SF766_13 422.9 U Chloritized Clay 3.7 Bornite BO1 12.7
83405 SF802_13 429.8 U Chloritized Clay 1.81 Gersdorffite GER2 7.6
CAM085014 SF725_12 434 U Chloritized Pelite 14.6 Chalcopyrite CPY1 12.0
CAM085015 SF731_16 434.6 Massive Sulphides 0.02 Chalcopyrite CPY1 8.0
CAM085033 U385 268.1 Graphitic Pelite 0.002 Pyrite PY0 -22.7
CAM085022 480-167 XC Graphitic Pelite 0.004 Pyrite PY0 -20.0
CAM085009 SF725_16 459 Graphitic Pelite 0.002 Pyrite PY3 14.6
CAM085001 SF725_16 427.3 MFb Sandstone 0.242 Hydrocarbon Button -31.0 -151
CAM085002 SF725_16 436.2 Altered Pelite 0.009 Hydrocarbon Button -28.4 -271
*The δ18O values for uraninite fluids were calculated using the uraninite-water fractionation factors proposed by Fayek and Kyser (2000).
**The δ18O fluid values for carbonate fluids were calculated using the calcite-water and siderite-water fractionation factors proposed by Zheng (1999).
Variation of the individual analyses is ±0.2 ‰ for δ18O and δ13C and ±0.3 ‰ for δ34S and δ2H.
See text for more details.
Fluid+Lithology/ Ore Facies
δ18O and δ2H are calculated to the VSMOW standard, whereas δ13C and δ34S are calculated to VPDB and VCDT standards respectively.
Table 3.4: Stable Isotopes
113
Figure 3.18: A) Vug filled with paragenetically late platy columnar aggregates of siderite (SI)
occurring in association with hematite. B) Uraninite crystals are crosscut by veinlets of calcite
(CA1) occurring in association with hematite and illitic clay. C) Hydrocarbon buttons (HB) within
strongly clay altered basement regolith occurring in association with chalcopyrite CPY4 and pyrite
PY4 (CAM085002). D) Hydrocarbon buttons within the intensely clay altered, indurated hematized
illitic clay overlying the orebody (CAM085001). E) Petrographic image (X-polar, reflective light) of
CAM085002 showing cockscomb texture and disseminated sulphides. F) Petrographic image
(transparent light) of opaque carbon buttons displaying cockscomb crystal structure. Reduced
alteration rim (dirty yellow) surrounds the carbon buttons. Goethite and hematite are generally
pervasive throughout the sample.
A B
C D
E F
2 mm 2 mm
1 cm 1 cm
1 cm 3 cm
MF
SST
VUG
SI
HECY
UR
CA
HB
PY4
CPY4
HB
HB
CLCY
PY4
HB
114
3.6 Discussion
3.6.1 Geochronology and Far-Field Tectonics
The oldest date obtained for the Cigar Lake deposit is 1468±93 Ma (U1), which represents the
minimum possible age of the deposit. This date is similar to the 1461±47 Ma reported for Cigar Lake by
Fayek et al. (2002) and those from various deposits throughout the Athabasca Basin (e.g. Beshears, 2010;
Sheahan et al., 2016). Significant discordance presented here and by Fayek et al. (2002) and the high error
in all of these analyses suggests the ca. 1468±93 Ma date reflects a partial resetting event rather than the
true age of initial mineralization. The Cigar Lake deposit is likely syngenetic with the older ca. 1590 Ma
event reported for McArthur and Millennium (Alexandre et al., 2009; Cloutier et al., 2009). However, the
deposit is located at the unconformity within extensive post-mineralization structures and alteration,
thereby inhibiting adequate preservation and resulting in ubiquitous Pb-loss.
The 1468±93 Ma partial resetting likely resulted from far-field tectonics responsible for faulting,
fault reactivation and fluid movement in the Athabasca Basin. Between ca. 1550–1350 Ma, the juvenile
volcanic arc terrane of the Granite-Rhyolite province, was accreted along the Southeast margins of
Laurentia, extending mid-continent along a Northeast trend from Mexico to Ontario (Whitmeyer and
Karlstrom, 2007). The Granite-Rhyolite province and Paleoproterozoic crust to the west were intruded by
A-type granites and anorthesite between ca. 1480 and 1350 Ma (Whitmeyer and Karlstrom, 2007).
Although the intrusions were previously interpreted as anorogenic, more recent interpretations of the
Grenville orogeny overprinted rocks, suggest an orogenic link (McLelland et al., 1996; Corrigan and
Hanmer, 1997; Whitmeyer and Karlstrom, 2007). Within the Athabasca Basin, this age range coincides
with a major fluid flow regime captured by north-northeast magnetic polarization of peak-diagenetic
hematite occurring between ca. 1500 to 1400 Ma (Kotzer et al., 1992).
The second and most dominate U-mineral generation has a date of 1270±10 Ma. Uraninite
generations at similar dates have been reported at Cigar Lake and at various deposits throughout the
Athabasca Basin (e.g. Cumming and Krstic, 1992; Cuney et al., 2002; Laverret et al., 2010; Sheahan et
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al., 2016). This age coincides with the onset of the Grenville Orogeny, a protracted period of tectonism
responsible for the assemblage of the supercontinent Rodinia with continent-continent accretion spanning
between ca. 1300 to 900 Ma (Whitmeyer and Karlstrom, 2007). During the initial stages of the Grenville
Orogeny between ca. 1.3–1.2 Ga, the Elzevir and Frontenac blocks amalgamated along eastern margin of
Laurentia (Moore and Thompson, 1980). Compressive northwest contraction along the southern margin
of Laurentia during the Grenvillian Orogeny was accompanied by intracratonic extension and mafic
magmatism (Whitmeyer and Karlstrom, 2007). In the Athabasca Basin, the ca. 1267 Ma Mackenzie dyke
swarm (LeCheminant and Heaman, 1989) was particularly important for reactivation of structures and the
fluids along them and temporally coincides with U2 Pb-resetting and U recrystallization.
Some uraninite crystals have dates of 1163±25 Ma. Similar dates have been widely reported at
various deposits throughout the Athabasca Basin (e.g. Fayek et al., 2002; Cloutier et al. 2009; Boulanger,
2012). Far-field tectonics from the Grenville Orogeny likely contributed to another U–Pb resetting event
at Cigar Lake. During the mid-to-late stages of the Grenville Orogeny, orogenic collapse and over
thickened crust resulted in widespread plutonic magmatism throughout Laurentia (McLelland, 1996;
Whitmeyer and Karlstrom, 2007). The emplacement of roughly coeval magmatism proximal to the Cigar
Lake deposit, the Moore Lake olivine diabase lopolith, in the southeastern Athabasca Basin at ca.
1100±25 Ma (MacDougall and Williams, 1993; French et al., 2002) may have contributed to this Pb-loss
event.
An additional common date for the uraninite crystals is 947±57 Ma. Dates at ca. 900 Ma have
been previously reported for various deposits throughout the Athabasca Basin (Kotzer et al., 1992;
Philippe et al., 1993; Fayek et al., 2002). This date represents the waning stages of the Grenville Orogeny
(Whitmeyer and Karlstrom, 2007). Within the Athabasca Basin, this date coincides with a major basin
fluid flow event responsible for magnetic polarization of hematite, coeval with the Rb–Sr date of ca. 970
Ma for illite observed at Key Lake, Rabbit Lake and the Midwest Lake deposits (Kotzer et al., 1992).
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The relatively unaltered botryoidal uraninite crystals regressed on a concordia plot have a lower
intersection age of 755±85 Ma coincident with the breakup of the supercontinent Rodinia. Current models
for the breakup of Rodinia indicate that rifting began in the western margin of Laurentia between ca. 780
and 680 Ma (Moores, 1991; Dalziel, 1991; Li et al., 2007; Whitmeyer and Karlstrom, 2007). The far-field
effects appear to have reactivated the fault systems at Cigar Lake, resulting in yet another Pb-loss event.
Although poorly resolved, younger dates are evident from lower U–Pb concordia intersection
points and EMPA chemical ages at 51±52 Ma and 200 Ma and 100–0 Ma, respectively. Due to the
multitude of Pb-loss events it is unclear whether these ages reflect distinct hydrothermal events, a mixture
of several hydrothermal events, or outward migration of Pb from U-bearing phases. However, what is
clear is that the process of protracted, episodic fluid incursion and subsequent Pb-loss from U-bearing
minerals has continued to modern times, as evident from the near zero (within error) U–Pb dates.
3.6.2 Pb Isotopes of Sulpharsenides, Sulphides and Non-Metallic Gangue Minerals
The Pb isotope ratios of sulphides, arsenides and non-metallic gangue illustrate a complex open
Pb-system that transitions from a common Pb isotope dominated regime to a uranium dominated Pb
isotope regime. A diverse range of sulphides and sulpharsenides occur with limited radiogenic Pb and
high 207Pb/206Pb values including: CPY1, PY1, SPH, COB1 and GER1 yielding 207Pb/206Pb, 206Pb/204Pb
and 207Pb/204Pb values similar to regional granitic gneiss, unaltered pelites and silicified sandstones
proximal to the Cigar Lake deposit (Pagel et al., 1993; Holk et al., 2003). Lower concentrations of
radiogenic Pb suggests Stage 1 minerals predominantly incorporated common Pb from the hydrothermal
system. This confirms paragenetic observations that the majority of metals were nearly syngenetic with
the main mineralizing event including Cu–Fe–Zn sulphides and Ni–Co sulpharsenides.
At a model age of ca. 1600 Ma (Figure 3.16 A, B), U-supported radiogenic Pb from the deposit
begins to dominate the Pb during precipitation of the sulphide minerals in Stage 2 CPY2, PY2 and GER2.
The 207Pb/206Pb values result from the mixing between distally sourced Pb and in-situ uranium supported
Pb.
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Stage 3 minerals from the paragenetic sequence include chalcocite, Pb adsorbed to hematized
clay, galena, and bornite. All these phases are anomalously radiogenic with uranium supported Pb
isotopes. The 207Pb/206Pb values of chalcocite, galena and bornite are low (0.05–0.09) and demonstrate a
young relative age. Model ages obtained from bornite and Pb adsorbed on hematized clay are 745±110
Ma and 844±72 Ma respectively. With an initial age of mineralization estimated at ca. 1590 Ma (Cloutier
et al., 2009; Alexandre et al., 2009) within the Athabasca Basin, the young date from bornite and
hematized clay coincident with the breakup of the supercontinent Rodinia.
3.6.3 Characterization of Fluids
3.6.3.1 Oxygen Isotopes
The U1 uraninite crystals have δ18O values of -17.5‰ and -8.6‰, indicating formation from a
fluid having δ18O values of -5.8‰ and +3.7‰, respectively, at 200°C (Wilson and Kyser, 1987; Fayek
and Kyser, 2000; Fayek et al., 2002). Previous isotopic and microthermometric studies on gangue
minerals in textural equilibrium with uraninite from the Athabasca Basin indicate that the mineralizing
fluid had a δ18O value between +2 and +8‰ (Wilson and Kyser, 1987; Kotzer and Kyser, 1990, 1992,
1993; Rees 1992; Alexandre et al., 2005; Cloutier et al., 2011). Therefore, the uraninite samples analyzed
are within the expected range for primary Stage 1 mineralization or are slightly 18O-depleted.
Unaltered primary uraninite analyzed from unconformity deposits throughout the Athabasca Basin
have δ18O between -34‰ to -15‰ resulting from total recrystallization (e.g. Kotzer and Kyser, 1993;
Fayek and Kyser, 1993; Fayek et al., 2002; Sheahan et al., 2016). In effect, U1 primary uraninite samples
from Cigar Lake are within the range reported for the Athabasca Basin (δ18O = -33.9 to -20.5‰; Fayek et
al., 2002). The low δ18O value of -5.8‰ for the mineralizing fluids and the significant depletion reported
by Fayek et al, (2002) reflect fluid incursion by recent low-temperature meteoric waters that exchange O
with uraninite with only minor disturbance to the texture (Kotzer and Kyser, 1993; Fayek et al., 2002).
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3.6.3.2 S Isotopes of Sulphides and Sulpharsenides
The δ34S values of the Cigar Lake minerals are compared with common sulfur reservoirs in Figure
3.19. The sulphides from the orebody at Cigar Lake have high δ34S values of +1.4 to +14.6‰ with no
discernable mineralogical or paragenetic trends. These results are comparable to δ34S values reported for
sulpharsenides at Key Lake, sulphides at McArthur and some stages of sulphide crystallization at Kianna
(Kotzer and Kyser, 1992; Emberley, 2014; Sheahan et al., 2016). In contrast, basement pyrite (PY0)
shows significantly lower δ34S values (-22.7 and -20.0‰) indicative of an isotopically distinct S source
that predates the Athabasca Basin. Because no igneous δ34S sources are found proximal to Cigar Lake
deposit, the stark contrast between basement and ore zone sulphides suggests a significant S contribution
from oxidized basin fluids that include Proterozoic seawater or evaporites (Kotzer and Kyser, 1992;
Sheahan et al., 2016). In a closed-system, if significant sulphate is present during sulphide crystallization,
the minerals will become increasingly enriched in 34S as the system evolves (Kotzer and Kyser, 1992).
Because the orebody has relatively constant and high δ34S values, demonstrating no evidence of 34S
fractionation, it can be concluded that the main mineralizing event was an open-system with sulphur from
a relatively homogenous oxidized basin brine source.
Kotzer and Kyser (1992) proposed that δ34S values obtained from the Athabasca Basin unconformity
deposits suggest the mixing of two isotopically distinct fluids: (1) a basement fluid with δ34S near 0‰,
and (2) a basinal fluid containing a component of Proterozoic seawater having a δ34S value near +12‰.
Such an interpretation is consistent with the Cigar Lake data, however the significant difference between
basement sulphides and orebody sulphides suggests only a minor contribution from the basement source.
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Figure 3.19: δ34S values from Cigar Lake samples. Ore body samples show consistent δ34S values
regardless of mineralogy. Basement pyrite (PY3) has a significantly lower δ34S values. Also shown
are δ34S from various sources with data from Kaplan and Hulston (1966), Rees et al. (1978) and
Krouse (1980). Modified from Seal et al. (2000).
3.6.3.3 C and O Isotopes in Carbonates and Hydrocarbons
The δ18O values of CA1 range from +12.3 to +13.4‰, suggesting equilibration with a fluid with
values between +1.8 to +7.2‰. Previous isotopic and microthermometric studies on gangue minerals
from the Athabasca Basin suggest that the primary mineralizing fluid had a δ18O value between +2 and
+8‰ (e.g. Wilson and Kyser, 1987; Kotzer and Kyser, 1990, 1992, 1993; Rees 1992; Alexandre et al.,
2005; Cloutier et al., 2011). The strong correlation between the δ18O values for CA1 and the mineralizing
fluids are consistent with the interpretation that CA1 precipitated during Stage 1 of the mineral
paragenesis. The textural cross-cutting relationship observed with CA1 filling microfractures of crackle-
brecciated uraninite crystals suggests CA1 precipitated at the end of Stage 1, during the waning
hydrothermal system. Measured δ13C values for CA1 range from -22.4 to -21.8‰. Isotopic analyses on
Cigar Lake graphite has been reported with δ13C values ranging from -28.2 to -27.5‰ (Landais et al.,
1993), whereas more extensive graphite sampling at Key Lake obtained δ13C values between -29.2
to -19‰ (Kyser et al., 1989). Similarities between the δ13C values of basement graphite and CA1 suggest
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the underlying graphite as the dominant CA1 calcite C source and that oxidizing basinal brines liberated
C during Stage 1 mineralization.
Measured δ18O values from CA2 and siderite suggest equilibration with a fluid having a δ18O
of -12.7 to -3.1‰. The high δ18O values in the carbonates are indicative of fluid incursion by recent low-
temperature meteoric waters (Kotzer and Kyser, 1993). Measured δ13C values from CA2 and siderite
are -18.8 to -11.5‰, which are higher than basement graphite, suggesting limited contribution from a
basement graphite source. Instead the δ13C values reflect a contribution from an organic-rich, oxidized
source of C, consistent with low-temperature Neoproterozoic (1000 to 543 Ma) meteoric waters and the
δ18O results (Johnston et al., 2012).
The δ13C values obtained for the hydrocarbon buttons (-31.0 and -28.4‰) are similar to those of
basement graphite at Cigar Lake (-28.2‰ to -27.5; Landais et al., 1993), suggesting that the graphite is
the likely C source. The degradation and displacement of graphite from the basement rocks immediately
underlying the deposit likely resulted from alteration caused by oxidized basinal fluids (Kyser et al., 1989,
Landais et al., 1993). Whole-rock geochemistry reveals that the carbon button samples contain
anomalously high concentrations of Fe2O3 (total Fe) at 21.5 (CAM85001) and 27.5 wt. % (CAM85002).
These high Fe contents are typically attributed to paragenetically late hematite and siderite at Cigar Lake
(Bruneton, 1987).
The δ2H values from the hydrocarbon buttons are extremely low and variable at -257 and -151‰.
Recent meteoric waters have δ2H values as low as -170‰ and glacial melt waters from the last glaciation
event would have even been lower (Wilson and Kyser, 1987; Kotzer and Kyser, 1990, 1991). Landais et
al. (1993) proposed that radiolysis of water forming hydrogen free radicals could result in extremely low
δ2H values. However, there is no evidence from the δ18O values of any mineral phase that this process
was widespread (Kotzer and Kyser, 1991). The variability in δ2H in the hydrocarbon buttons suggests that
both may be correct, with the original values having low δ2H values from meteoric and glacial water or
radiolysis on the water contributing to the lowest and varied δ2H values.
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3.6.4 Genetic Model
The genetic deposit model that best explains the data obtained expands on the diagenetic-
hydrothermal model first purposed by Hoeve and Sibbald (1976, 1978). The diagenetic-hydrothermal
model postulates that oxidized U-bearing diagenetic basinal brines were focused by reactivated structures
and reacted with basement rocks or reduced basement-sourced fluids at the unconformity to produce
mineralization (Hoeve and Sibbald, 1976, 1978). Geochemical, geochronological and stable isotopic
characterization of the deposits and host-rocks over the last forty years has led to a refinement of the
diagenetic-hydrothermal model and characterization of the fluids involved in their formation (e.g. Kotzer
and Kyser, 1995; Fayek and Kyser, 1997; Alexandre et al., 2005; Cloutier et al., 2011). Here we
incorporate deposit scale observations and interpretations to explain the evolution of the Cigar Lake
deposit (Figure 3.20).
Stage 1 Event
The oldest age obtained from the Cigar Lake deposit, ca. 1468±93 Ma (U1), is discordant and
represents the minimum possible age of the deposit, but the discordance indicates it is more likely a
resetting event. The Cigar Lake deposit is likely syngenetic with the basin-wide uranium event at ca. 1590
Ma (Cloutier et al., 2009; Alexandre et al., 2009). During the mineralization Stage 1, uraninite was
initially precipitated as uraniferous gels and underwent Ostwald ripening and dehydration (e.g. Schindler
et al., 2017). Primary mineralization was crystallized as uraninite rather than coffinite suggesting silica
undersaturation. Pre-mineralization and syngenetic faulting created preferential pathways focusing
regional fluid flow. Syngenetic brecciation of the initially emplaced uraninite resulted from tectonic
faulting, dissolution induced collapse brecciation, and over pressured hydraulic fracturing. Within the
deposit, ubiquitous structures resulted in structural and geochemical control of the distribution of primary
metals (Figure 3.20). Empirical spatial modelling of the deposit demonstrates a zonation of metals from
Cu to Ni to Co to Zn as a function of O and S fugacities.
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123
Figure 3.20: Genetic and evolutionary model for the Cigar Lake deposit. Stage 1 mineralization
(top) resulted from the mixing of oxidized U+Cu+SO42-+Mo-bearing basin fluid with an evolved
basement fluid containing Ni+Co+As+Fe+Mg. Distally sourced Pb-isotopes were crystallized within
Stage 1 sulphides and sulpharsenides. Stage 2 fluid incursion resulted in Pb-loss from primary
uraninite and absorption into Stage 2 sulphides and sulpharsenides. Stage 3 and 4 resulted in
further Pb-loss and Pb absorption in highly radiogenic sulphides. Parenthesis denote element minor
contribution.
The initial mineralizing event was polymetallic with the precipitation of redox active elements
such as U, S, As, Ni, Co and Mo. Syn-mineralization sulphides and sulpharsenides incorporated
regionally sourced common Pb (206Pb/204Pb = ~75) with high 207Pb/206Pb values (~0.5). Later mobilization
of radiogenic Pb from U1 was also incorporated into these sulphides primarily by coprecipitation or
substitution of Pb into syn-mineralization sulphides, although this process was limited. The high δ34S
values suggest a significant basinal contribution of sulphate with minimal fractionation indicating an
open-system involving a relatively homogenous oxidized basin brine. The anomalously low 207Pb/206Pb
values (0.56–0.86) imprinted on co-precipitated CPY1, the cathodic-like crystallization of Cu on U, and
the deposit-scale spatial zonation of Cu intimately associated with high-grade primary uraninite, all
suggest that Cu and U were transported in the same basinal fluid and sourced from Archean minerals.
The Pb–Pb and stable isotopes along with the mineral textures and deposit-scale geochemical
zonation all suggest that a U+Cu+SO42- basinal brine was focused along major regional northeast and
east-northeast oriented structures coincident with the regional basement foliation. Mineralization occurred
when oxidized basin fluids mixed with egress Ni+Co+As+Fe+Mg-bearing evolved basement brines that
ascended up the reactivated local east-west corridor.
CA1 calcite crystallized at the end of Stage 1, within the waning hydrothermal system from a
fluid having a δ18O value between +1.8 to +7.2‰, consistent with previous reported ranges for primary
mineralization (e.g. Wilson and Kyser, 1987; Kotzer and Kyser, 1990, 1992, 1993; Rees 1992; Alexandre
et al., 2005; Cloutier et al., 2011). CA1 has similar δ13C values (-22.4 to -21.8‰) as graphite suggesting
the carbon source for CA1 calcite was the oxidization of the underlying reduced graphite. The absence of
12C enrichment suggests CH4 was minimal during primary mineralization and that most of the C in CA1
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was mobilized as CO2. Calcite CA1 content is limited within the deposit, suggestion Fe2+ liberated from
mafic minerals in the basement contributed to U reduction (Quirt, 1989; Hetch & Cuney, 2000; Derome et
al., 2003; Alexandre et al., 2005; Acevedo and Kyser, 2015).
Stage 2 Alteration Event
The second stage in the mineral paragenesis reflects the major far-field tectonic events recorded
by U–Pb resetting of U1 uraninite at 1270±10 Ma and 1163±25 Ma. These dates coincide with the
Grenville Orogeny, a protracted period of tectonism responsible for the assemblage of the supercontinent
Rodinia (Whitmeyer and Karlstrom, 2007). In the Athabasca Basin, the regional effects of the orogeny
include the Mackenzie dyke swarm at ca. 1267 Ma and the Moore Lake olivine diabase lopolith at ca.
1100 Ma (LeCheminant and Heaman, 1989; MacDougall and Williams, 1993; French et al., 2002). This
tectonically induced alteration event resulted in crystallization, recrystallization and Pb absorption
forming radiogenic sulphides and sulpharsenides (CPY2, PY2 and GER2) with higher 206Pb/204Pb (79–
11000) and lower 207Pb/206Pb values (0.050–0.25) in contrast to the initial mineralizing event. The high
δ34S values of +1.4 to +14.6‰ from all generations of sulphides and sulpharsenides suggest chemical
buffering by dissolution of primary sulphide minerals and subsequent recrystallization as the dominant S
source. Consistent 208Pb/204Pb values reflect limited basement sourced egress fluid contribution. Selenium
and Bi are crystallized within these sulphides and sulpharsenides formed during the Grenville stage. Clay
alteration was substantial with K–Ar dating giving 1255 to 1148 Ma ages for illite (Percival et al., 1993).
Stage 3 Alteration Event
The third stage in the mineral paragenesis reflects further alteration resulting from far-field
tectonics recorded in the uraninite crystals as U–Pb resetting at 947±57 Ma and 755±85 Ma. Furthermore,
Stage 3 alteration coincides with the crystallization of bornite and the loading of radiogenic Pb on
hematized clay. These events involved very radiogenic Pb and have Pb–Pb dates of ca. 845 Ma and ca.
746 Ma respectively. These dates coincide with a major tectonic transition in the Trans Hudson from the
waning stages of the Grenville Orogeny to the breakup of the supercontinent of Rodinia (Whitmeyer and
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Karlstrom, 2007). Basin incursion by low temperature (~50°C), oxidized, meteoric waters resulted from
rifting of the supercontinent (Sheaghan et al., 2016).
Stage 3 resulted in significant oxidization of the deposit with local dissolution of primary
chalcopyrite and subsequent crystallization of bornite and chalcocite. The spatial zonation of Cu-bearing
minerals (Figure 3.13) suggests an east-northeast fluid flow enhanced the geochemical zonation
developed during primary mineralization. This fluid flow direction is consistent with modern groundwater
movement at Cigar Lake suggesting a preferential fluid anisotropy likely related to regional structures
(Cramer and Smellie, 1994). The Stage 3 alteration event resulted in sulphides (BO, GN, HECY) with
very high 206Pb/204Pb (1000–88000) and low 207Pb/206Pb (0.045–0.095) values. Consistently high δ34S
values confirms chemical buffering, dissolution and subsequent recrystallization as the dominant S
source. Low and relatively consistent 208Pb/204Pb values reflect limited basement sourced egress fluid
contribution.
CA2 and siderite from Stage 3 impregnated and indurated clays, crystallized from a fluid having a
range in δ18O of -12.7 to -3.1‰. δ13C values from CA2 and siderite suggest a distal C source from an
oxidizing organic-rich C environment consistent with low-temperature meteoric waters. The K–Ar dates
(Percival et al., 1993; Philippe et al., 1993) of coeval clay minerals sudoite (850 Ma), Fe-illite and Fe-
kaolinite (900 Ma) are consistent with the timing of Stage 3 deduced from the U–Pb and Pb–Pb dates of
altered uraninite and sulphides in this study.
Stage 4 Alteration Event
Stage 4 alteration resulted in extensive coffinitization and U remobilization. The event is
responsible for the remobilization of U into brittle structures in the overlying sandstone as uneconomic
perched mineralization (Cumming and Krstic, 1992). This event is poorly constrained but likely reflects
fluid incursion from meteoric waters following unloading during deglaciation. The 18O-depleted meteoric
waters flushed through the deposit and were responsible for overprinting the δ18O isotopes within primary
uraninite during re-crystallization (Kotzer and Kyser, 1993; Fayek et al., 2002). The δ13C values obtained
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for paragenetically late hydrocarbon buttons are similar to the basement graphite at Cigar Lake and have
extremely low δ2H values consistent with meteoric and glacial meltwaters. The degradation and
displacement of C from the basement graphite immediately underlying the deposit likely resulted from
alteration caused by penetrating oxidized basinal fluids (Kyser et al., 1989; Landais et al., 1993).
3.7 Exploration Implications
Previous studies have revealed that Pb isotopes can be used as a pathfinder for uranium
mineralization (Holk et. al., 2003). Regional Pb isotopes surrounding the Cigar Lake deposit have been
particularly suitable for these techniques due to the extensive dispersion of leachable radiogenic Pb (Holk
et. al., 2003). It has long been understood that some of the radiogenic-Pb lost from the U minerals is
entrapped within the deposit particularly as extremely uranogenic galena crystals (Kister et al., 2003).
Here, we identify other mineral phases that have absorbed and adsorbed radiogenic Pb including
sulphides, sulpharsenides, arsenides, hematite and clay minerals. Therefore, only a portion of the
radiogenic Pb that has been removed from U-bearing mineral phases during post-mineralizing fluid
incursion events is dispersed from the deposit. This makes Pb isotope anomalies in the overlying
sandstones or along fractures more significant for exploration targeting because the mass of remobilized
Pb is less than previously thought (e.g. Kister et al., 2003).
3.8 Conclusions
Uranium at the Cigar Lake deposit was affected by at least six major fluid events. Concordia plots of
U–Pb data obtained through LA-ICP-MS on uraninite and coffinite indicate discordia-lines with upper
intercepts at 1468±93 Ma, 1270±10 Ma, 1163±25Ma and ca. 947±57Ma and one discrete lower intercept
from pristine botryoidal uraninite at 755±81 Ma. A continuum of younger ages from 0 to 242 Ma are also
evident from the lower intercepts of the U–Pb system and from EMPA chemical dating of coffinite.
Mineral associations, textural relationships, Pb isotopes and stable isotopes reveal four major stages of
evolution for the deposit (Figure 3.20):
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1) Stage 1, the main mineralization event, occurred before 1468±93 Ma in response to far-field
tectonics with the accretion of the Granite-Rhyolite province along the Southeast margins of
Laurentia. Fault reactivation in the Athabasca Basin resulted in mixing of oxidized U+Cu+SO42--
bearing basin fluid with an evolved basement fluids containing Ni+Co+As+Fe+Mg. Stage 1
mineralization was polymetallic with structural permabilities and O and S fugacity dictating the
precipitation and spatial distribution of metallic minerals along the redox boundary. Stage 1
minerals equilibrated with fluids having high δ18O up to +7.2‰ and δ34S values up to +15‰,
which are typical of basinal brines in the basin and marine sulphate.
2) Stage 2 alteration at 1270±10 Ma, 1163±25 Ma reflects the impact of fluid incursion from the
Grenville Orogeny, the Mackenzie dyke swarm and the Moore Lake olivine diabase. Dissolution
of Stage 1 primary minerals and subsequent recrystallization resulted in radiogenic, often Se and
Bi bearing sulphides and sulpharsenides and enhanced clay alteration.
3) Stage 3, with alteration at 947±57 and 755±81 Ma, reflects the major tectonic transition from the
breakup of the supercontinent of Rodinia resulting in basin-wide incursion by low temperature
(~50°C), oxidized, meteoric waters, with low δ18O values as low as -12.7‰. These fluids were
responsible for extensive hematization, induration of clays by siderite and CA2 calcite, and
crystallization of bornite and chalcocite with highly radiogenic Pb mobilized from primary
uraninite.
4) Stage 4 alteration is responsible for extensive coffinitization of the main orebody and the
mobilization of U into brittle structures overlying the deposit as perched mineralization. These
oxidized fluids had very low δ2H values (< -151‰) and were responsible for the mobilizing C
from the underlying graphite in the basement forming bituminous C and hydrocarbon buttons.
The paragenesis, geochronology and stable isotopes reveal a protracted history, with ongoing episodic
fluid incursion, driven by far-field tectonics and resulting in alteration and re-mobilization of susceptible
elements. Preservation of the Cigar Lake deposit over approximately 1500 Ma years is largely thought to
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be due to: 1) the large mass of sulphides and arsenides maintaining low redox values in present day fluids
and 2) a substantial clay cap which physically reduces hydraulic conductivity, all of which favors long
term geochemical stability of uraninite.
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Chapter 4
GENERAL DISCUSSION, SUMMARY OF CONTRIBUTIONS AND
RECOMMENDATIONS FOR FUTURE WORK
4.1 General Discussion
The overall objective of this research is to integrate mineralogical, geochemical and geospatial
characteristics within geometallurgy for the sandstone-hosted, high-grade, polymetallic, unconformity-
related, Cigar Lake U deposit. This thesis is a reevaluation of the ore-forming system and evolution of the
deposit in support of exploration, mining, milling and mine tailings management. Since the initial
discovery of the Cigar Lake orebody in 1981, many geological, mineralogical, geochemical and
geochronological aspects of the deposit have been studied in detail (e.g. Bruneton, 1987, 1993; Percival
and Kodama, 1989; Cumming and Krstic, 1992; Landais et al., 1993; Pacquet and Weber, 1993; Pagel et
al., 1993; Percival et al., 1993; Philippe et al., 1993; Reyx and Rulmann, 1993; Toulhoat and Beaucaire,
1993; Cramer and Smellie, 1994; Janeczek and Ewing, 1992, 1994; Cramer, 1995; Fayek and Kyser,
1997; Fayek et al., 1997, 2000, 2002). However, research focused directly on the U ore and associated
metals, and the underlying paragenetic model is sparse (e.g. Bruneton, 1987; Reyx and Rulmann, 1993),
with studies typically lacking access to high-grade U ore samples. Recent delineation and ongoing
operational drilling has improved access and spatial coverage of high-grade U ores. Technological
advancements have also been made since the early works (e.g. Bruneton, 1987; Reyx and Rulmann, 1993)
providing a new opportunity to reevaluate the deposit using quantitative mineralogical and geochemical
techniques (XRD, SEM-MLA, EMPA, LA-ICP MS), as well as stable and radiogenic isotopes (LA-ICP
MS, IRMS), to develop a robust empirical model and understand the evolution of the Cigar Lake deposit.
The results from this thesis are presented as two manuscripts, each dealing with key aspects of the
overall objective. The first examines and characterizes the empirical spatial distribution of minerals and
their crystal chemistry with a focus on elements of concern (EOCs) as it pertains to mining, milling and
130
mine tailings management. Within the complex polymetallic ores of Cigar Lake, elements such as As,
Mo, Se and Zr, can be problematic during mining, milling and tailings management and have been
identified as EOCs (Bishop et al., 2016). Minerals and mineraloid phases represent significant elemental
controls and also have properties that affect mineral processing and mobility of EOCs in process waters
and long-term tailings management facilities. Normative algorithms have been designed to utilize an
extensive historical whole-rock geochemical dataset to predict minerals and quantify mineral proportions.
Integration of geochemical, mineralogical, geological and geospatial characterization of the high-grade U
and Ni–Co–As–S ores has been done to support current geologic modelling by providing predictive ore
characteristics for mining and milling.
The second manuscript addresses the underlying genetic and evolutionary model for the Cigar
Lake deposit. Excellent early works by Bruneton (1987), and Reyx and Rulmann (1993) developed a
paragenetic model for the deposit. This work was only qualitative and was based solely on textural
mineral observations. In this study, a textural paragenetic interpretation is integrated with semi-
quantitative mineral characterization and empirical geospatial modelling of the Phase 1 Cigar Lake
deposit. Geochronological characterization using U–Pb and Pb–Pb systematics provides constraints on
the absolute ages of various uranium ores, arsenides, sulphides and non-metallic gangue minerals. Stable
isotope chemistry is used to characterize fluids responsible for the formation and subsequent alteration of
the orebody. The various analytical methods are combined to reevaluate the initial ore-forming system
and protracted evolution of the deposit, through episodic fluid incursion, resulting in alteration and re-
mobilization of susceptible elements.
131
4.2 Significant Contributions
4.2.1 Geometallurgical Contributions for Mining, Milling and Tailings Management
Within the Phase 1 pods, As, Ni, and Co occur primarily in a reduced state as arsenides and
sulpharsenides. The arsenides and sulpharsenides are dominated by 1:1 molar ratios of Ni–Co:As
in minerals such as gersdorffite (NiAsS), niccolite (NiAs) and cobaltite (CoAsS). Ni-rich mineral
end-members predominate over their Co-rich varieties throughout most of the deposit. Milling
experience in the Athabasca Basin over the last forty years has shown that 1:1 molar ratio Ni–
Co:As minerals are typically less exothermic during oxidation than their Ni–Co biarsenide and
triarsenide counterparts and are therefore less problematic during milling (Areva Resources,
personal communication). Arsenides and sulpharsenides are spatially zoned within the deposit
increasing in prominence along the north and northeastern side of the deposit on both the West
and East Pods in association with the reduced side of the deposit scale redox front.
Molybdenum was identified occurring in the mineral phase molybdenite and within Stage 1
uraninite with concentrations up to 0.46 wt. % MoO3. Elevated Mo concentrations in the uraninite
coincides with elevated Pb levels suggesting that Mo is syngenetic with primary U
mineralization. The spatial distribution of Mo coincides with high-grade U (>40% U3O8) along
the deposit scale redox front confirming its mode of occurrence within the uraninite.
Selenium was found to occur in sulphides and sulpharsenides with Se2- substituting for S2-.
Selenium appears to be paragenetically late (Stage 2) and is observed with increasing
concentrations towards the boundaries of GER2 crystals. Galena, another paragenetically late
mineral was also observed to be prone to Se uptake. This suggests that S fugacity may control the
spatial distribution of Se, with Se concentrations increasing with decreasing S activity. Some
coffinite crystals contained anomalous SeO2 content with up to 0.34 wt. % SeO2. The spatial
distribution of Se within the orebody shows a strong correlation with the high-grade U3O8
132
corridor but an inverse relationship with elevated Mo (>5000 ppm) suggesting whole-rock Se
content may reflect localities within the deposit particularly effected by Stage 2 fluid incursion.
The mode of occurrence for Zr was determined to be within detrital zircon crystals and within
coffinite with some crystals yielding up to 1.36 wt. % ZrO2. Preliminary modelling of the
empirical spatial distribution suggests that Zr appears to occur preferentially along the flanks of
the deposit particularly on the southern side. This distribution is consistent with the primary mode
of occurrence as detrital zircons.
Clay characterization from the ores indicate that the clay mineralogy is dominated by white-mica
clay mixtures of illite and muscovite. Local patches of phengite, paragonite, Fe-chlorite, Mg-
chlorite, kaolinite and montmorillonite are observed throughout the ore body.
Utilizing whole-rock geochemistry to develop normative algorithms for mineral quantification
was successful in expanding the spatial coverage of mineralogical data at a significantly lower
cost than conventional mineral analyses. In combination with implicit geospatial modelling, this
geometallurgical paradigm offers guidance for mining, milling and mine tailings management.
4.2.2 Cigar Lake Deposit Evolutionary Model Contributions
U–Pb data obtained through LA-ICP-MS on uraninite and coffinite indicate U at the Cigar Lake
deposit was affected by at least six major fluid alteration events. Concordia plots demonstrate
discordia-lines with upper intercepts at 1468±93 Ma, 1270±10 Ma, 1163±25 Ma and 947±57 Ma
and one discrete lower intercept from pristine botryoidal uraninite at 755±81 Ma. A continuum of
younger ages from 242 to 0 Ma are also evident from the lower intercepts of the U–Pb system
and from EMPA chemical dating of coffinite. The paragenesis, geochronology and stable isotopes
reveal a protracted history, with ongoing episodic fluid incursion, coincident with far-field
tectonics and resulting in alteration and re-mobilization of susceptible elements. The main
mineralization event occurred before 1468±93 Ma and was likely syngenetic with the basin wide
133
ca. 1590 Ma U mineralization event (Cloutier et al., 2009; Alexandre et al., 2009). The first major
alteration event occurred at 1468±93 Ma in response to the accretion of the Granite-Rhyolite
province along the Southeast margins of Laurentia. Stage 2 with alteration at 1270±10 Ma and
1163±25 Ma reflects the impact of fluid incursion from the Mackenzie dyke swarm, the Grenville
Orogeny, and the Moore Lake olivine diabase. Stage 3, with alteration at 947±57 and 755±81 Ma,
reflects the major tectonic transition from the breakup of the supercontinent of Rodinia resulting
in basin-wide incursion by low temperature (~50°C), oxidized, meteoric waters. Stage 4 is poorly
constrained (~242 to 0 Ma), but likely reflects fluid incursion from meteoric waters during
deglaciation.
The low 207Pb/206Pb values (0.56–0.86) imprinted on co-precipitated CPY1, the cathodic-like
crystallization of Cu on U, and the deposit-scale spatial zonation of Cu intimately associated with
high-grade primary uraninite, all suggest that Cu and U were transported in the same basinal
fluid. The low 207Pb/206Pb values (0.56–0.86) confirms that the U at Cigar Lake was not sourced
from zircons and monzonites directly from the Wollaston Supergroup or pre-Athabasca protores,
but was more likely transported in a basinal brine and sourced from Archean minerals.
A diverse range of sulphides and sulpharsenides occur with limited radiogenic Pb and high
207Pb/206Pb values including: CPY1, PY1, SPH, COB1 and GER1 yielding 207Pb/206Pb, 206Pb/204Pb
and 207Pb/204Pb values similar to regional Archean granites, unaltered pelites and silicified
sandstones proximal to the Cigar Lake deposit (Pagel et al., 1993; Holk et al., 2003). Lower
concentrations of radiogenic Pb suggest Stage 1 minerals predominantly incorporated common
Pb from the hydrothermal system. This confirms paragenetic observations that the majority of
metals were nearly syngenetic with the main mineralizing event including Cu–Fe–Zn sulphides
and Ni–Co sulpharsenides.
Empirical spatial modelling of the geochemistry reveals a deposit scale redox front with more
oxidized monometallic ores occurring to the south and southwest and Ni-Co-As-bearing
134
polymetallic ores occurring to the north and northeast ends of both the Phase 1 pods. A corridor
of high-grade U (>40% U3O8) coincides with the redox boundary. This spatial zonation together
with the Pb–Pb isotopes from coeval sulphides suggests that Stage 1 mineralization was
polymetallic with structural permabilities, fluid source, and O and S fugacities dictating the
precipitation and spatial distribution of metallic minerals along the redox boundary. The zonation
indicates that basinal brines were focused along major regional northeast and east-northeast
oriented structures coincident with the regional basement foliation. This fluid flow direction is
consistent with modern groundwater movement at Cigar Lake suggesting a preferential fluid
anisotropy likely related to major regional structures (Cramer and Smellie, 1994).
The sulphides from the orebody at Cigar Lake have high δ34S values up to 15‰, which are typical
of basinal brines and marine sulphate. In contrast, pre-Athabasca basement pyrite showed
significantly lower δ34S values (-20.0 and -22.7‰). The high δ34S values suggest a significant
basinal contribution of sulphate with minimal fractionation indicating an open-system involving a
relatively homogenous oxidized basin brine that was consumed during reduction associated with
formation of the deposit. The consistent δ34S values throughout all stages of the paragenesis
indicate that preservation of the deposit is a function of kinetics and geochemical stabilities, with
sulphides and arsenides showing a tendency to buffer penetrating fluids resulting in
remobilization and subsequent recrystallization within the deposit.
Syngenetic Stage 1 minerals equilibrated with fluids having high δ18O ranging between +1.8 to
+7.2‰ typical of basinal brines in the Athabasca Basin. In contrast, the δ18O for fluids in
equilibrium with Stage 3 minerals were low with values ranging from -12.7 to -3.1‰ consistent
with low temperature meteoric waters. During Stage 4 oxidized fluids had very low δ2H values
(< -151‰) suggesting incursion by meteoric glacial meltwaters.
135
The δ13C values from Stage 1 coeval calcite are consistent with underlying graphite suggesting C
was oxidized during the main mineralizing event. Stage 3 carbonates have higher δ 13C values
(-18.8 to -11.1‰) suggesting a C source from an oxidized organic-rich environment.
4.2.3 Contributions to Mineral Exploration
Lead isotopes can be used as a pathfinder for U mineralization and the regional Pb isotopes
surrounding the Cigar Lake deposit have been particularly suitable for these techniques due to the
extensive dispersion of leachable radiogenic Pb (Holk et. al., 2003). Within this study numerous
mineral phases have been shown to absorb radiogenic Pb including sulphides, sulpharsenides,
arsenides, and clay. Therefore, only a portion of the radiogenic Pb that has been removed from U-
bearing mineral phases during post-mineralizing fluid incursion events is dispersed from the
deposit. This makes Pb isotope anomalies in the overlying sandstones or along fractures even
more significant.
4.3 Recommendations for Future Work
High clay content and the presence of poorly crystallized phases made Cigar Lake ores unsuitable
for mineral quantification by Rietveld refinement. Within this study the Reference Intensity Ratio
(RIR) method was used to provide semi-quantification and Mineral Liberation Analysis (MLA)
was utilized for quantification of the mineralogy. The MLA results did show a reasonable
correlation with normatively derived mineral proportions. However, using a crystal scanning
technique such as MLA, assumes that the samples represent a well-constrained standard analysis.
Alternatively, using flotation techniques to remove clay prior to analysis by XRD would facilitate
Rietveld refinement and could be attempted to verify the normative.
Minor patches of montmorillonite have been measured with SWIR and typically correspond with
geological structures. Percival et al. (1993) has reported illite–smectite mixed layers in the
alteration zone at Cigar Lake with 5–10% expandability, however limited details were provided
136
in the report. Alternatively, illite–chlorite mixed clay layers at Close Lake and McArthur River
have XRD patterns consistent with smectite mixed-layer clays, however no swelling was
observed with glycolation suggesting sepiolite, illite–chlorite, illite–vermiculite or hydrobiotite
(Quirt, 1999). No glycol testing was performed in the current study to assess clay swelling. It is
recommended that glycol testing be undertaken to confirm the SWIR mineral interpretation and
quantify the degree of expandability.
The success of the normative algorithm is fundamentally dependent on the consistent mineralogy
identified throughout the Phase 1 pods. The normative is dependent on the molar element ratios
of the identified mineral phases. Therefore, a major shift in mineralogy, or significant element
substitutions, would require refinement of the algorithm. The current study utilized ongoing
operational drilling that provided improved exposure of the Phase 1 pods with a focus on near- to
medium-term mining. Expanding the work to Phase 2 will be required to confirm that the
mineralogy remains consistent as production progresses.
This study has shown that technological advancements since the early works by Bruneton (1987),
and Reyx and Rulmann (1993) can be used to reevaluate even high quality genetic and
evolutionary models. Within this study, quantitative mineralogical and geochemical techniques
(XRD, SEM-MLA, EMPA, LA-ICP-MS), as well as stable and radiogenic isotopes (LA-ICP-MS,
IRMS), were used to refine the Cigar Lake evolutionary model. Applying this paradigm to other
deposits within the Athabasca Basin may help in our understanding of unconformity-related U
deposits and how they can be exploited.
137
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Appendix A
Whole-Rock ICP-OES and ICP-MS Data Summary
U3O8 S* As Ni Co Mo Se Cu Pb Zn Zr Bi Al2O3 MgO K2O Fe2O3 CaO
(%) (%) (%) (%) (%) (%) (ppm) (%) (%) (%) (ppm) (ppm) (%) (%) (%) (%) (%)
Samples (n)** 9038 3527 8131 8452 8188 8239 4018 7982 8169 7818 3876 2556 8024 4279 4279 4279 4279
Mean 6.52 1.65 0.88 0.55 0.12 0.10 65 0.49 0.61 0.03 835 276 12.3 2.82 2.65 10.2 0.81
Maximum 82.9 35.2 37.7 28.5 9.79 3.94 1590 27.0 11.5 8.20 11098 10300 37.3 20.1 10.1 73.2 26.4
Minimum <DL <DL <DL <DL <DL <DL <DL <DL <DL <DL 5 0.1 0.01 <DL <DL 0.06 <DL
<DL denotes below lower detection limit
*Analyzed with leco induction furnace
** Samples typically range in downhole length from 0.1 to 0.5 m
152
Ho
leSa
mp
leFr
om
ToU
3O8
PbS*
As
Ni
CoCu
ZnM
oSe
**B
i**
ZrC*
Org
C*
(%)
(%)
(%)
(%)
(%)
(%)
(%)
(pp
m)
(pp
m)
(pp
m)
(pp
m)
(pp
m)
(%)
(%)
342A
3460
643
8.2
438.
610
.80.
240.
020.
030.
010.
010.
0129
390.
50.
513
700.
130.
68
342A
3461
344
0.3
440.
91.
180.
030.
780.
090.
030.
020.
045
220.
515
231
0.19
0.05
342A
3461
544
1.4
441.
66.
650.
120.
110.
240.
210.
030.
0420
230.
511
467
40.
190.
20
342A
3461
944
2.7
442.
82.
320.
050.
110.
030.
050.
000.
005
50.
50.
540
30.
330.
04
349
3475
641
7.3
417.
91.
310.
1511
.420
.10
10.8
06.
690.
2924
2017
4061
872
847
2.03
1.14
349
3475
841
841
8.4
0.30
0.10
14.0
32.6
027
.60
2.08
0.15
1260
2130
4799
017
81.
550.
38
349
3476
141
9.5
420.
236
.53.
274.
820.
610.
260.
153.
1611
900
3370
390
655
1190
0.26
0.09
337
5087
042
8.1
428.
40.
200.
041.
840.
130.
050.
048.
958
2934
263
1050
0.23
0.19
337
5087
943
1.8
432.
24.
010.
767.
041.
891.
670.
1623
.60
8493
6067
757
941
0.18
0.15
337
5089
043
6.3
436.
871
.99.
891.
170.
330.
140.
100.
1012
600
225
462
508
385
0.19
0.03
348
5984
343
0.8
431.
10.
040.
094.
130.
060.
010.
030.
8024
000
778
7321
322
0.87
0.09
348
5984
443
1.1
431.
529
.03.
383.
140.
220.
020.
061.
1813
029
5037
247
693
0.17
0.06
348
5984
743
2.2
432.
44.
341.
271.
800.
160.
060.
070.
0942
2040
667
361
890
0.08
0.07
353
8006
443
343
3.3
2.12
0.06
0.01
0.34
0.04
0.10
0.02
1814
0.5
146
1180
3.91
2.14
353
8006
743
3.7
434.
10.
060.
030.
030.
080.
010.
020.
0112
512
190
884
3.50
0.20
353
8007
544
0.2
440.
515
.21.
265.
182.
091.
210.
823.
3130
1070
800.
527
1026
800.
120.
12
361
8010
642
8.44
429.
13.
170.
090.
140.
050.
030.
010.
018
1222
719
800.
480.
30
361
8010
742
9.1
429.
652.
600.
090.
040.
160.
060.
040.
0211
1616
2818
800.
110.
05
365
8014
444
1.6
441.
82.
420.
534.
282.
812.
220.
171.
0090
1740
054
1490
1610
0.13
0.12
365
8014
844
2.8
443.
14.
920.
435.
546.
815.
880.
240.
7217
111
500
123
2720
1810
0.10
0.09
365
8015
244
3.8
444.
20.
920.
073.
386.
124.
940.
380.
8116
311
6011
340
418
0.25
0.22
SF76
6_05
8023
943
3.8
434.
28.
650.
260.
030.
010.
020.
000.
0113
9284
270
2320
0.53
0.07
SF76
6_05
8024
343
5.2
435.
626
.40.
880.
010.
010.
010.
000.
066
6027
729
711
901.
540.
13
SF76
6_05
8024
643
6.4
436.
62.
570.
132.
090.
150.
050.
027.
250.
543
8075
849
123
100.
170.
15
SF82
6_10
8077
842
6.8
427.
118
.82.
352.
310.
140.
080.
035.
0172
1010
383
912
300.
180.
07
SF82
6_10
8078
642
9.3
429.
731
.62.
916.
153.
022.
160.
214.
4013
546
8010
112
6091
60.
170.
06
SF82
6_10
8079
143
0.9
431.
357
.84.
835.
402.
221.
510.
552.
7296
8020
110
1740
371
0.37
0.03
SF82
6_10
8079
643
2.7
433
0.13
0.02
0.70
0.13
0.04
0.04
0.00
2710
0013
8634
50.
670.
64
363
8166
142
1.8
422.
42.
110.
153.
551.
460.
990.
1712
.20
131
3512
0.5
1740
0.09
0.06
363
8166
842
4.5
424.
875
.79.
611.
090.
460.
360.
051.
4449
413
7058
281
822
40.
170.
04
363
8167
042
5.3
425.
54.
740.
735.
845.
984.
930.
165.
5123
372
9010
953
375
90.
380.
34
ICP
-OES
an
d L
ECO
Wh
ole
-Ro
ck G
eo
che
mis
try
*Ele
me
nt
an
aly
zed
by
LECO
**Pa
rtia
l D
ige
stio
n
Appendix B
153
Ho
leSa
mp
leFr
om
ToU
3O
8Pb
S*A
sN
iCo
CuZn
Mo
Se**
Bi*
*Zr
C*O
rg C
*
(%)
(%)
(%)
(%)
(%)
(%)
(%)
(pp
m)
(pp
m)
(pp
m)
(pp
m)
(pp
m)
(%)
(%)
363
8167
342
6.5
427
0.19
0.18
8.28
1.92
1.27
0.12
10.8
010
534
104
2811
400.
460.
46
364
8168
243
3.3
433.
457
.07.
183.
182.
731.
980.
212.
4453
347
3065
985
774
00.
310.
19
364
8168
543
4.1
434.
536
.13.
793.
321.
000.
700.
201.
8946
1790
052
524
7015
600.
170.
06
364
8169
243
7.1
437.
559
.76.
941.
420.
550.
340.
171.
1727
738
491
477
561
0.23
0.09
364
8169
543
8.2
438.
71.
540.
140.
940.
290.
130.
050.
6417
1980
1914
533
60.
730.
67
369
8231
143
0.9
431.
24.
550.
120.
600.
580.
180.
150.
015
7357
178
1610
0.85
0.40
369
8232
143
5.5
436.
10.
050.
152.
000.
170.
130.
061.
7252
4400
5789
131
200.
440.
35
369
8232
443
6.7
437.
266
.25.
812.
792.
662.
040.
246.
4823
216
5093
120
0030
20.
160.
06
370
8235
344
2.5
442.
90.
010.
021.
041.
691.
770.
010.
0211
591
0.5
7421
600.
220.
11
370
8235
944
4.7
445.
14.
010.
200.
446.
976.
090.
020.
0315
1730
7860
317
000.
180.
17
370
8236
344
6.2
446.
641
.34.
452.
7313
.60
10.8
00.
060.
3133
421
5059
615
1035
20.
220.
10
SF76
6_13
8323
542
2.9
423.
34.
370.
413.
870.
700.
370.
2410
.90
5429
471
1490
1750
0.23
0.17
SF76
6_13
8324
342
6.4
426.
967
.57.
481.
680.
020.
070.
070.
864
9590
928
1250
461
0.17
0.15
SF76
6_13
8324
642
7.45
428
37.5
4.43
3.41
0.91
0.80
0.14
2.20
8313
000
589
744
1500
0.35
0.27
SF76
6_13
8325
042
9.1
429.
50.
130.
050.
590.
550.
460.
030.
1325
716
1418
238
00.
580.
54
SF81
4_04
8326
243
1.7
432
6.38
0.48
1.14
0.07
0.09
0.01
0.02
2123
9452
826
100.
170.
16
SF81
4_04
8326
643
3.25
433.
664
.17.
791.
330.
030.
040.
024.
099
604
813
885
313
0.21
0.07
SF81
4_04
8326
843
443
4.4
0.03
0.07
0.50
0.04
0.02
0.02
0.12
944
7033
388
2540
0.24
0.22
SF80
2_13
8340
342
9.2
429.
40.
600.
050.
752.
281.
140.
430.
4669
233
4932
428
200.
730.
73
SF80
2_13
8340
542
9.8
429.
92.
070.
123.
4418
.30
9.36
3.37
0.72
321
760
9013
3012
801.
341.
24
SF80
2_13
8341
043
0.6
431
5.44
0.28
1.68
5.11
1.79
1.58
0.09
7118
8048
562
1820
0.28
0.25
SF80
2_13
8341
543
2.1
432.
316
.21.
710.
580.
440.
320.
040.
1631
7039
011
252
322
800.
120.
04
**P
art
ial
Dig
est
ion
*Ele
me
nt
an
aly
zed
by
LECO
ICP
-OES
an
d L
ECO
Wh
ole
-Ro
ck G
eo
che
mis
try
154
ICP-OES Whole-Rock Geochemistry of Major OxidesHole Sample From To Al 2O3 CaO Fe2O3 FeO* K2O MgO Na 2O TiO2
(%) (%) (%) (%) (%) (%) (%) (%)
342A 34606 438.2 438.6 10.8 0.43 22.8 0.21 1.67 1.30 0.08 1.33
342A 34613 440.3 440.9 3.23 0.15 10.2 0.14 0.32 0.57 0.02 0.46
342A 34615 441.4 441.6 6.62 0.40 10.2 0.07 0.82 1.20 0.04 0.95
342A 34619 442.7 442.8 4.22 0.17 4.79 0.14 0.35 0.92 0.02 0.43
349 34756 417.3 417.9 8.72 6.70 8.42 1.89 0.45 4.83 0.06 1.24
349 34758 418 418.4 1.57 7.97 2.28 1.68 0.10 0.82 0.02 0.32
349 34761 419.5 420.2 9.86 1.93 9.79 2.48 0.68 1.98 0.19 2.60
337 50870 428.1 428.4 27.3 0.24 1.89 1.75 8.83 1.42 0.19 0.83
337 50879 431.8 432.2 15.1 0.19 3.90 5.10 3.28 1.23 0.09 0.82
337 50890 436.3 436.8 0.81 1.76 1.51 3.79 0.09 1.39 0.08 0.99
348 59843 430.8 431.1 2.40 4.34 4.8 1.39 0.16 1.29 0.02 0.90
348 59844 431.1 431.5 7.10 0.96 9.53 2.77 0.62 3.55 0.08 1.47
348 59847 432.2 432.4 6.87 0.37 6.57 3.06 0.40 3.39 0.05 2.32
353 80064 433 433.3 13.6 0.50 38.9 23.35 3.93 2.02 0.10 0.85
353 80067 433.7 434.1 18.3 0.42 28.0 19.03 5.67 1.62 0.11 0.74
353 80075 440.2 440.5 13.9 0.58 12.3 3.21 1.28 4.54 0.06 4.80
361 80106 428.44 429.1 20.8 0.42 27.0 7.73 0.94 6.50 0.08 3.99
361 80107 429.1 429.65 18.8 0.26 35.1 6.42 0.82 5.65 0.08 3.81
365 80144 441.6 441.8 19.7 0.44 8.59 3.21 0.82 9.93 0.08 1.95
365 80148 442.8 443.1 16.1 0.50 6.42 3.06 0.81 7.41 0.07 7.35
365 80152 443.8 444.2 22.8 0.25 3.18 1.02 5.09 6.25 0.14 0.85
SF766_05 80239 433.8 434.2 15.6 2.51 25.9 0.87 2.71 5.63 0.16 3.14
SF766_05 80243 435.2 435.6 10.5 7.68 16.0 3.65 1.12 4.71 0.23 2.20
SF766_05 80246 436.4 436.6 27.8 0.37 0.88 1.60 1.24 1.19 0.11 7.38
SF826_10 80778 426.8 427.1 12.4 1.29 20.3 6.71 2.86 1.79 0.14 1.62
SF826_10 80786 429.3 429.7 9.20 0.54 13.8 3.35 0.23 1.68 0.03 1.95
SF826_10 80791 430.9 431.3 2.81 1.08 7.56 6.86 0.10 1.04 0.04 0.94
SF826_10 80796 432.7 433 29.4 0.20 3.40 1.02 8.50 5.28 0.18 0.84
363 81661 421.8 422.4 20.0 0.24 12.3 8.32 4.00 4.55 0.10 1.21
363 81668 424.5 424.8 1.75 1.25 2.34 4.67 0.34 0.91 0.04 0.28
363 81670 425.3 425.5 19.9 0.33 8.18 2.77 1.64 5.43 0.07 0.56
*Analyzed by ti tration
155
ICP-OES Whole-Rock Geochemistry of Major OxidesHole Sample From To Al 2O3 CaO Fe2O3 FeO* K2O MgO Na 2O TiO2
(%) (%) (%) (%) (%) (%) (%) (%)
363 81673 426.5 427 19.1 0.25 14.1 2.78 1.30 7.20 0.08 2.09
364 81682 433.3 433.4 4.46 1.16 6.27 4.23 0.61 1.17 0.05 0.88
364 81685 434.1 434.5 10.2 0.91 9.38 3.50 0.29 3.04 0.05 2.32
364 81692 437.1 437.5 4.6 1.83 6.27 4.08 0.11 3.40 0.07 1.14
364 81695 438.2 438.7 27.2 0.30 3.05 1.31 7.19 5.60 0.21 0.93
369 82311 430.9 431.2 22.3 0.35 16.2 8.03 6.17 1.52 0.11 0.96
369 82321 435.5 436.1 27.1 0.37 5.16 2.33 2.28 5.97 0.07 6.93
369 82324 436.7 437.2 1.70 0.53 1.51 7.15 0.15 0.07 0.01 1.13
370 82353 442.5 442.9 18.6 0.87 6.97 3.80 0.31 19.20 0.06 1.97
370 82359 444.7 445.1 18.0 0.69 8.47 6.71 0.32 13.80 0.06 2.68
370 82363 446.2 446.6 4.59 0.98 2.35 6.42 0.15 1.62 0.03 0.60
SF766_13 83235 422.9 423.3 18.2 0.46 12.1 6.86 1.28 6.98 0.13 1.09
SF766_13 83243 426.4 426.9 3.37 1.84 2.15 3.94 0.67 0.80 0.10 0.80
SF766_13 83246 427.45 428 10.6 1.06 5.73 3.20 1.09 3.76 0.09 3.62
SF766_13 83250 429.1 429.5 28.4 0.28 1.93 0.87 7.09 5.60 0.22 0.94
SF814_04 83262 431.7 432 18.2 0.39 19.3 5.69 3.08 5.17 0.11 3.18
SF814_04 83266 433.25 433.6 2.18 2.31 5.30 5.98 0.37 1.41 0.12 0.97
SF814_04 83268 434 434.4 26.9 0.38 6.01 3.06 1.76 8.94 0.10 5.38
SF802_13 83403 429.2 429.4 25.5 0.28 5.42 3.07 3.77 6.61 0.12 3.08
SF802_13 83405 429.8 429.9 15.2 0.26 4.14 3.65 1.72 3.09 0.07 1.77
SF802_13 83410 430.6 431 21.2 0.40 10.3 5.25 1.97 4.38 0.11 3.13
SF802_13 83415 432.1 432.3 13.9 0.60 21.4 5.40 1.09 3.60 0.10 4.09
*Analyzed by ti tration
156
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
1C
AM
08
50
40
_1_1
Ura
nin
ite
84
.27
<DL
10
.37
0.1
7<D
L<D
L<D
L<D
L0
.49
0.2
0<D
L<D
L0
.34
0.1
11
.51
<DL
97
.54
86
2
2C
AM
08
50
40
_1_2
Ura
nin
ite
83
.66
<DL
10
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0.2
3<D
L<D
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L<D
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0.1
11
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97
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2
3C
AM
08
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83
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0.2
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97
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1
4C
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08
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ite
84
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97
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08
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84
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84
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83
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ite
83
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85
5
10
CA
M0
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ran
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1<D
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.52
0.2
6<D
L<D
L0
.42
0.1
21
.64
<DL
97
.58
83
2
14
CA
M0
85
04
0_2
_2U
ran
init
e8
4.9
1<D
L9
.88
0.1
8<D
L<D
L<D
L<D
L0
.54
0.2
3<D
L<D
L0
.32
0.1
01
.55
<DL
97
.78
81
8
15
CA
M0
85
04
0_3
_1U
ran
init
e8
4.3
2<D
L1
0.0
80
.23
<DL
0.0
5<D
L<D
L0
.54
0.2
3<D
L<D
L0
.48
0.1
21
.70
<DL
97
.86
83
9
16
CA
M0
85
04
0_3
_2U
ran
init
e8
4.3
5<D
L1
1.0
90
.21
<DL
<DL
<DL
<DL
0.4
20
.22
<DL
<DL
0.3
30
.10
1.3
5<D
L9
8.1
39
17
17
CA
M0
85
04
0_1
_1U
ran
init
e8
3.7
6<D
L1
1.5
20
.18
<DL
<DL
<DL
<DL
0.3
50
.24
<DL
<DL
0.3
00
.10
1.2
0<D
L9
7.7
49
56
18
CA
M0
85
04
0_1
_2U
ran
init
e8
4.1
0<D
L1
0.8
60
.19
<DL
<DL
<DL
<DL
0.4
90
.26
<DL
<DL
0.4
10
.14
1.4
9<D
L9
8.1
09
02
19
CA
M0
85
04
5_1
_1U
ran
init
e8
3.0
0<D
L1
2.1
00
.17
<DL
0.0
60
.14
<DL
0.2
90
.41
<DL
<DL
0.2
70
.12
1.2
4<D
L9
7.8
11
00
9
20
CA
M0
85
04
5_1
_2U
ran
init
e8
2.8
8<D
L1
1.7
80
.24
<DL
<DL
<DL
<DL
0.3
80
.24
<DL
<DL
0.3
30
.11
1.3
5<D
L9
7.4
29
86
21
CA
M0
85
04
5_1
_3U
ran
init
e8
2.7
7<D
L1
2.6
20
.24
<DL
<DL
0.1
7<D
L0
.26
0.2
0<D
L<D
L0
.28
0.0
81
.06
<DL
97
.80
10
52
22
CA
M0
85
04
5_1
_4U
ran
init
e8
3.1
4<D
L1
2.1
00
.24
<DL
<DL
0.1
1<D
L0
.33
0.1
3<D
L<D
L0
.34
0.1
21
.38
<DL
97
.97
10
08
23
CA
M0
85
04
5_1
_5U
ran
init
e8
4.5
2<D
L9
.11
0.1
9<D
L<D
L<D
L<D
L0
.49
0.0
5<D
L<D
L0
.50
0.1
61
.93
<DL
97
.17
76
1
24
CA
M0
85
04
5_1
_6U
ran
init
e8
3.1
7<D
L1
1.4
10
.30
<DL
0.0
5<D
L<D
L0
.34
0.1
7<D
L<D
L0
.34
0.1
01
.37
<DL
97
.39
95
4
25
CA
M0
85
04
5_1
_7U
ran
init
e8
2.5
7<D
L1
2.4
40
.21
<DL
<DL
<DL
<DL
0.3
20
.37
<DL
<DL
0.2
80
.08
1.1
9<D
L9
7.5
71
04
0
26
CA
M0
85
04
5_1
_8U
ran
init
e8
3.0
0<D
L1
2.0
00
.24
<DL
<DL
<DL
<DL
0.2
90
.37
<DL
<DL
0.2
70
.10
1.1
3<D
L9
7.5
61
00
2
27
CA
M0
85
04
5_1
_9U
ran
init
e8
3.0
2<D
L1
2.2
60
.24
<DL
<DL
<DL
<DL
0.2
70
.36
<DL
<DL
0.2
70
.08
1.0
8<D
L9
7.7
11
02
1
28
CA
M0
85
04
5_1
_10
Ura
nin
ite
82
.23
<DL
13
.57
0.1
7<D
L<D
L<D
L<D
L0
.18
0.2
5<D
L<D
L0
.22
0.0
70
.85
<DL
97
.71
11
31
29
CA
M0
85
04
5_1
_11
Ura
nin
ite
81
.37
<DL
14
.34
0.1
60
.05
<DL
<DL
<DL
0.2
00
.29
<DL
<DL
0.1
80
.07
0.9
1<D
L9
7.7
71
20
1
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
Appendix C
157
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
30
CA
M0
85
04
5_1
_1U
ran
init
e8
1.2
0<D
L1
4.3
20
.13
<DL
<DL
<DL
<DL
0.2
30
.32
<DL
<DL
0.2
10
.06
0.9
9<D
L9
7.6
31
20
2
31
CA
M0
85
04
5_1
_2U
ran
init
e8
2.7
2<D
L1
2.2
90
.18
<DL
<DL
<DL
<DL
0.2
90
.14
<DL
<DL
0.2
90
.07
1.2
2<D
L9
7.4
41
02
7
32
CA
M0
85
04
5_1
_3U
ran
init
e8
3.4
8<D
L1
1.0
90
.24
<DL
<DL
<DL
<DL
0.3
80
.16
<DL
<DL
0.3
80
.12
1.4
8<D
L9
7.4
39
26
33
CA
M0
85
04
5_1
_4U
ran
init
e7
8.6
5<D
L1
5.8
10
.04
<DL
<DL
<DL
<DL
0.1
20
.04
<DL
0.4
60
.04
<DL
1.2
1<D
L9
6.3
81
35
4
34
CA
M0
85
04
5_1
_5U
ran
init
e7
9.4
9<D
L1
5.1
10
.07
<DL
<DL
<DL
<DL
0.1
60
.05
<DL
0.3
60
.14
0.0
41
.35
<DL
96
.89
12
87
35
CA
M0
85
04
5_1
_6U
ran
init
e8
0.0
6<D
L1
4.3
20
.14
<DL
<DL
<DL
<DL
0.1
40
.18
<DL
0.4
10
.22
0.0
50
.97
<DL
96
.61
12
18
36
CA
M0
85
04
5_1
_7U
ran
init
e8
1.1
2<D
L1
4.3
00
.14
<DL
<DL
<DL
<DL
0.1
50
.19
<DL
<DL
0.2
00
.06
0.8
8<D
L9
7.4
81
20
2
37
CA
M0
85
04
5_1
_8U
ran
init
e8
0.7
0<D
L1
4.8
80
.10
<DL
<DL
<DL
<DL
0.1
70
.41
<DL
<DL
0.2
00
.05
0.6
6<D
L9
7.4
51
25
2
38
CA
M0
85
03
9_1
_1U
ran
init
e8
1.9
2<D
L1
3.6
70
.20
<DL
<DL
<DL
<DL
0.2
20
.44
<DL
<DL
0.2
10
.07
0.9
2<D
L9
7.7
31
14
3
39
CA
M0
85
03
9_1
_2U
ran
init
e8
3.1
1<D
L1
1.4
40
.26
<DL
<DL
<DL
<DL
0.3
60
.38
<DL
<DL
0.3
10
.09
1.2
8<D
L9
7.4
69
57
40
CA
M0
85
03
9_1
_3U
ran
init
e8
3.0
1<D
L1
2.2
60
.19
<DL
<DL
<DL
<DL
0.2
50
.40
<DL
<DL
0.2
60
.08
0.9
8<D
L9
7.6
41
02
1
41
CA
M0
85
03
9_1
_4U
ran
init
e8
3.0
5<D
L1
1.0
80
.32
<DL
<DL
<DL
<DL
0.3
40
.43
<DL
<DL
0.3
50
.11
1.2
7<D
L9
7.1
89
29
42
CA
M0
85
03
9_1
_5U
ran
init
e8
2.5
1<D
L1
2.9
00
.26
<DL
<DL
<DL
<DL
0.2
50
.40
<DL
<DL
0.2
50
.07
1.0
3<D
L9
7.8
51
07
7
43
CA
M0
85
03
9_1
_1U
ran
init
e8
3.3
0<D
L1
1.3
50
.30
<DL
0.0
6<D
L<D
L0
.36
0.4
4<D
L<D
L0
.34
0.0
91
.32
<DL
97
.69
94
8
44
CA
M0
85
03
9_1
_2U
ran
init
e8
3.3
1<D
L1
1.3
50
.30
<DL
0.0
6<D
L<D
L0
.38
0.5
0<D
L<D
L0
.32
0.0
91
.35
<DL
97
.69
94
8
45
CA
M0
85
03
9_1
_3U
ran
init
e8
2.1
6<D
L1
3.9
10
.18
<DL
<DL
<DL
<DL
0.1
90
.36
<DL
<DL
0.1
70
.06
0.8
0<D
L9
7.9
61
15
8
46
CA
M0
85
03
9_1
_4U
ran
init
e8
2.4
2<D
L1
3.7
50
.15
<DL
<DL
0.1
2<D
L0
.17
0.4
4<D
L<D
L0
.18
0.0
50
.81
<DL
98
.12
11
43
47
CA
M0
85
03
9_1
_5U
ran
init
e8
2.5
9<D
L1
3.5
10
.20
<DL
<DL
<DL
<DL
0.1
90
.39
<DL
<DL
0.1
90
.05
0.8
4<D
L9
8.0
81
12
2
48
CA
M0
85
03
9_1
_6U
ran
init
e8
2.5
3<D
L1
3.2
90
.18
<DL
<DL
<DL
<DL
0.2
20
.38
<DL
<DL
0.1
90
.06
0.8
8<D
L9
7.8
31
10
6
49
CA
M0
85
03
9_1
_1U
ran
init
e8
2.0
4<D
L1
3.3
10
.19
<DL
<DL
0.1
6<D
L0
.20
0.2
7<D
L<D
L0
.19
0.0
70
.91
<DL
97
.54
11
14
50
CA
M0
85
03
9_1
_2U
ran
init
e8
3.3
4<D
L1
2.2
20
.21
<DL
<DL
<DL
<DL
0.2
70
.20
<DL
<DL
0.2
50
.07
1.1
0<D
L9
7.8
11
01
5
51
CA
M0
85
03
9_1
_3U
ran
init
e8
2.3
1<D
L1
3.4
30
.19
<DL
<DL
<DL
<DL
0.2
30
.27
<DL
<DL
0.2
40
.07
0.9
9<D
L9
7.8
61
12
0
52
CA
M0
85
03
9_1
_4U
ran
init
e8
3.1
4<D
L1
1.3
60
.22
<DL
<DL
<DL
<DL
0.3
50
.26
<DL
<DL
0.3
40
.11
1.3
4<D
L9
7.2
39
50
53
CA
M0
85
03
9_1
_5U
ran
init
e8
2.8
9<D
L1
2.4
60
.23
<DL
<DL
<DL
<DL
0.3
00
.45
<DL
<DL
0.2
60
.08
1.0
8<D
L9
7.8
81
03
8
54
CA
M0
85
03
9_1
_6U
ran
init
e8
0.9
4<D
L1
4.9
10
.18
<DL
<DL
<DL
<DL
0.1
20
.29
<DL
<DL
0.1
20
.03
0.7
3<D
L9
7.4
61
25
1
55
CA
M0
85
01
2_1
_1U
ran
init
e8
2.6
4<D
L1
1.9
20
.27
<DL
0.1
10
.11
<DL
0.3
10
.88
<DL
<DL
0.2
40
.11
1.1
2<D
L9
7.7
69
99
56
CA
M0
85
01
2_1
_2U
ran
init
e8
2.4
5<D
L1
1.9
70
.28
<DL
0.1
1<D
L<D
L0
.34
0.8
7<D
L<D
L0
.28
0.1
01
.13
<DL
97
.69
10
05
57
CA
M0
85
01
2_1
_3U
ran
init
e8
2.3
2<D
L1
3.6
40
.23
<DL
0.0
60
.11
<DL
0.2
00
.62
<DL
<DL
0.1
80
.05
0.8
4<D
L9
8.2
71
13
6
58
CA
M0
85
01
2_1
_4U
ran
init
e8
1.6
6<D
L1
3.7
30
.17
<DL
<DL
<DL
<DL
0.2
00
.65
<DL
<DL
0.1
80
.08
0.8
2<D
L9
7.6
51
15
1
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
158
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
59
CA
M0
85
01
2_1
_5U
ran
init
e8
1.6
2<D
L1
4.4
70
.11
<DL
<DL
<DL
<DL
0.1
90
.55
<DL
<DL
0.1
80
.04
0.6
8<D
L9
7.9
11
20
8
60
CA
M0
85
01
2_1
_1U
ran
init
e8
1.6
4<D
L1
4.1
60
.20
<DL
<DL
<DL
<DL
0.1
90
.54
<DL
<DL
0.1
5<D
L0
.75
<DL
97
.80
11
84
61
CA
M0
85
01
2_1
_2U
ran
init
e8
1.2
1<D
L1
4.4
00
.13
<DL
<DL
0.1
3<D
L0
.18
0.5
1<D
L<D
L0
.15
0.0
40
.72
<DL
97
.54
12
08
62
CA
M0
85
01
2_1
_3U
ran
init
e8
1.9
2<D
L1
4.0
50
.13
<DL
0.0
5<D
L<D
L0
.21
0.5
8<D
L<D
L0
.18
0.0
60
.75
<DL
98
.07
11
72
63
CA
M0
85
01
2_1
_4U
ran
init
e8
1.5
1<D
L1
4.1
50
.17
<DL
<DL
<DL
<DL
0.2
00
.59
<DL
<DL
0.1
80
.04
0.7
8<D
L9
7.7
91
18
5
64
CA
M0
85
01
2_1
_5U
ran
init
e8
1.5
7<D
L1
4.4
30
.12
<DL
<DL
<DL
<DL
0.1
80
.53
<DL
<DL
0.1
50
.05
0.7
0<D
L9
7.8
81
20
6
65
CA
M0
85
01
2_1
_1U
ran
init
e8
1.7
9<D
L1
3.5
10
.19
<DL
0.0
8<D
L<D
L0
.22
0.7
8<D
L<D
L0
.23
0.0
70
.89
<DL
97
.86
11
32
66
CA
M0
85
01
2_1
_2U
ran
init
e8
2.5
7<D
L1
2.9
60
.22
<DL
0.0
5<D
L<D
L0
.26
0.7
6<D
L<D
L0
.21
0.0
70
.94
<DL
98
.11
10
80
67
CA
M0
85
01
2_1
_3U
ran
init
e8
2.0
7<D
L1
3.5
00
.23
<DL
<DL
<DL
<DL
0.2
20
.75
<DL
<DL
0.1
90
.09
0.8
8<D
L9
8.1
31
12
8
68
CA
M0
85
01
2_1
_4U
ran
init
e8
2.3
0<D
L1
3.1
30
.21
<DL
0.0
6<D
L<D
L0
.27
0.7
9<D
L<D
L0
.20
0.0
60
.93
<DL
98
.05
10
97
69
CA
M0
85
01
2_1
_5U
ran
init
e8
1.1
6<D
L1
3.9
00
.22
<DL
0.0
70
.12
<DL
0.2
20
.58
<DL
<DL
0.1
90
.08
0.8
4<D
L9
7.3
91
17
1
70
CA
M0
85
01
2_1
_6U
ran
init
e8
1.4
9<D
L1
4.6
60
.15
<DL
<DL
<DL
<DL
0.1
80
.52
<DL
<DL
0.1
90
.07
0.6
8<D
L9
8.0
91
22
4
71
CA
M0
85
03
5_1
_1U
ran
init
e8
2.3
4<D
L1
3.5
80
.19
<DL
<DL
<DL
<DL
0.2
00
.45
<DL
<DL
0.2
00
.06
0.9
3<D
L9
8.0
31
13
1
72
CA
M0
85
03
5_1
_2U
ran
init
e8
2.9
4<D
L1
2.5
80
.15
<DL
<DL
0.1
1<D
L0
.28
0.4
7<D
L<D
L0
.27
0.0
91
.14
<DL
98
.10
10
47
73
CA
M0
85
03
5_1
_3U
ran
init
e8
2.9
6<D
L1
1.8
30
.20
<DL
<DL
<DL
<DL
0.3
70
.49
<DL
<DL
0.3
00
.10
1.2
1<D
L9
7.5
49
89
74
CA
M0
85
03
5_1
_4U
ran
init
e8
3.6
4<D
L1
1.8
90
.22
<DL
<DL
0.1
2<D
L0
.33
0.4
9<D
L<D
L0
.28
0.0
81
.26
<DL
98
.38
98
6
75
CA
M0
85
03
5_1
_5U
ran
init
e8
3.0
6<D
L1
2.2
60
.18
<DL
<DL
0.1
5<D
L0
.32
0.4
9<D
L<D
L0
.28
0.0
81
.21
<DL
98
.05
10
21
76
CA
M0
85
03
5_1
_6U
ran
init
e8
3.6
6<D
L1
1.6
70
.18
<DL
<DL
<DL
<DL
0.3
40
.32
<DL
<DL
0.2
60
.08
1.1
5<D
L9
7.7
39
69
77
CA
M0
85
03
5_1
_1U
ran
init
e8
2.0
2<D
L1
4.4
10
.15
<DL
<DL
<DL
<DL
0.2
00
.40
<DL
<DL
0.1
70
.07
0.8
4<D
L9
8.4
61
19
8
78
CA
M0
85
03
5_1
_2U
ran
init
e8
0.9
5<D
L1
4.4
30
.14
<DL
<DL
<DL
<DL
0.2
20
.38
<DL
<DL
0.1
80
.08
0.9
5<D
L9
7.4
51
21
4
79
CA
M0
85
03
5_1
_3U
ran
init
e8
2.1
4<D
L1
3.1
90
.17
<DL
<DL
0.1
2<D
L0
.29
0.4
3<D
L<D
L0
.21
0.1
11
.08
<DL
97
.79
11
03
80
CA
M0
85
03
5_1
_4U
ran
init
e8
1.1
9<D
L1
3.4
10
.15
<DL
<DL
<DL
<DL
0.2
90
.47
<DL
<DL
0.2
00
.11
1.0
50
.23
97
.22
11
32
81
CA
M0
85
03
5_1
_5U
ran
init
e8
2.0
6<D
L1
3.6
50
.14
<DL
<DL
<DL
<DL
0.2
50
.44
<DL
<DL
0.2
00
.09
0.9
4<D
L9
7.9
41
14
0
82
CA
M0
85
03
5_1
_6U
ran
init
e8
1.3
0<D
L1
4.6
60
.19
<DL
<DL
0.1
0<D
L0
.19
0.3
8<D
L<D
L0
.17
0.0
70
.87
<DL
98
.09
12
27
83
CA
M0
85
03
5_1
_7U
ran
init
e8
1.7
0<D
L1
4.9
40
.16
<DL
<DL
<DL
<DL
0.1
40
.36
<DL
<DL
0.1
20
.04
0.6
8<D
L9
8.2
91
24
3
84
CA
M0
85
03
5_1
_1U
ran
init
e8
3.0
2<D
L1
1.8
10
.20
0.0
4<D
L<D
L<D
L0
.36
0.5
1<D
L<D
L0
.32
0.1
01
.21
<DL
97
.68
98
7
85
CA
M0
85
03
5_1
_2U
ran
init
e8
3.8
9<D
L1
0.8
60
.23
<DL
<DL
0.1
2<D
L0
.35
0.5
9<D
L<D
L0
.35
0.1
01
.42
<DL
98
.01
90
4
86
CA
M0
85
03
5_1
_3U
ran
init
e8
3.1
2<D
L1
1.8
20
.20
<DL
<DL
<DL
<DL
0.2
90
.58
<DL
<DL
0.2
80
.09
1.2
3<D
L9
7.6
99
86
87
CA
M0
85
03
5_1
_4U
ran
init
e8
3.6
2<D
L1
2.0
10
.22
<DL
<DL
<DL
<DL
0.2
80
.53
<DL
<DL
0.2
90
.09
1.1
7<D
L9
8.3
69
95
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
159
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
88
CA
M0
85
03
5_1
_5U
ran
init
e8
3.8
3<D
L1
0.9
70
.22
<DL
0.0
6<D
L<D
L0
.36
0.6
0<D
L<D
L0
.30
0.1
01
.33
<DL
97
.92
91
3
89
CA
M0
85
03
5_1
_6U
ran
init
e8
3.8
7<D
L1
0.6
60
.21
<DL
0.0
6<D
L<D
L0
.42
0.5
3<D
L<D
L0
.33
0.1
11
.43
<DL
97
.67
88
8
90
CA
M0
85
03
5_1
_7U
ran
init
e8
1.8
0<D
L1
4.3
40
.14
<DL
<DL
<DL
<DL
0.1
80
.48
<DL
<DL
0.1
90
.04
0.8
5<D
L9
8.1
21
19
6
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
160
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
1C
AM
08
50
40
_A_1
Co
ffin
ite
69
.59
<DL
1.8
80
.88
<DL
0.2
00
.47
<DL
14
.29
<DL
<DL
<DL
0.1
9<D
L2
.32
0.3
49
0.2
31
99
2C
AM
08
50
40
_A_2
Co
ffin
ite
72
.27
<DL
1.2
50
.85
<DL
0.2
1<D
L<D
L1
2.7
9<D
L<D
L<D
L0
.32
<DL
1.9
20
.13
90
.14
12
8
3C
AM
08
50
40
_A_3
Co
ffin
ite
71
.70
<DL
1.2
30
.89
<DL
0.2
2<D
L<D
L1
3.4
2<D
L<D
L<D
L0
.20
<DL
2.6
2<D
L9
0.7
41
27
4C
AM
08
50
40
_A_4
Co
ffin
ite
69
.35
<DL
2.2
80
.65
<DL
<DL
<DL
<DL
14
.65
<DL
<DL
<DL
0.1
9<D
L1
.95
0.1
08
9.4
82
42
5C
AM
08
50
40
_B_1
Co
ffin
ite
75
.85
<DL
<DL
0.4
7<D
L<D
L0
.32
<DL
11
.26
<DL
<DL
<DL
0.4
2<D
L2
.67
<DL
91
.31
7
6C
AM
08
50
40
_B_2
Co
ffin
ite
70
.98
<DL
<DL
0.7
0<D
L<D
L<D
L<D
L1
5.3
6<D
L<D
L<D
L0
.30
<DL
2.2
3<D
L9
0.1
20
7C
AM
08
50
40
_B_3
Co
ffin
ite
74
.39
<DL
<DL
0.5
7<D
L<D
L<D
L<D
L1
2.5
00
.17
<DL
<DL
0.3
8<D
L2
.88
<DL
91
.28
3
8C
AM
08
50
40
_B_4
Co
ffin
ite
71
.95
<DL
0.2
80
.73
<DL
0.1
7<D
L<D
L1
3.0
5<D
L<D
L<D
L0
.52
<DL
1.9
6<D
L8
9.1
62
9
9C
AM
08
50
40
_B_5
Co
ffin
ite
74
.30
<DL
0.7
70
.37
<DL
<DL
<DL
<DL
12
.06
<DL
0.3
1<D
L0
.48
<DL
2.9
9<D
L9
1.5
27
7
10
CA
M0
85
04
0_C
_1C
off
init
e6
6.9
5<D
L0
.79
0.7
0<D
L<D
L<D
L<D
L1
4.0
4<D
L0
.52
<DL
0.2
5<D
L2
.30
<DL
85
.84
88
11
CA
M0
85
04
0_C
_2C
off
init
e7
0.0
0<D
L0
.83
1.4
3<D
L0
.15
0.5
8<D
L1
2.7
4<D
L<D
L<D
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.26
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1.8
6<D
L8
8.2
38
9
12
CA
M0
85
04
0_C
_3C
off
init
e7
2.1
6<D
L0
.38
1.3
5<D
L0
.28
0.4
5<D
L1
2.6
2<D
L<D
L<D
L0
.28
<DL
1.9
0<D
L8
9.8
94
0
13
CA
M0
85
04
0_C
_4C
off
init
e6
8.3
6<D
L0
.33
1.0
4<D
L<D
L0
.39
<DL
13
.24
<DL
0.1
7<D
L0
.21
<DL
3.6
0<D
L8
7.7
63
6
14
CA
M0
85
03
4_A
_1C
off
init
e7
5.1
6<D
L0
.12
0.1
8<D
L<D
L<D
L<D
L1
2.7
81
.01
0.2
9<D
L0
.21
<DL
1.0
7<D
L9
0.9
71
2
15
CA
M0
85
03
4_A
_2C
off
init
e7
5.0
0<D
L0
.23
0.3
5<D
L<D
L<D
L<D
L1
2.3
20
.31
0.1
5<D
L0
.29
<DL
1.0
8<D
L9
0.0
72
3
16
CA
M0
85
03
4_A
_3C
off
init
e7
5.1
20
.12
<DL
0.3
4<D
L<D
L<D
L<D
L1
2.2
80
.18
<DL
<DL
0.3
1<D
L1
.08
<DL
89
.65
2
17
CA
M0
85
03
4_A
_4C
off
init
e7
7.3
5<D
L0
.17
<DL
<DL
<DL
<DL
<DL
10
.46
1.9
1<D
L<D
L0
.35
<DL
1.0
6<D
L9
1.6
31
6
18
CA
M0
85
03
4_A
_5C
off
init
e7
3.3
4<D
L0
.48
0.5
2<D
L<D
L<D
L<D
L1
3.0
20
.20
<DL
<DL
0.2
8<D
L1
.07
<DL
89
.19
49
19
CA
M0
85
03
4_A
_6C
off
init
e7
3.1
6<D
L<D
L0
.36
<DL
<DL
<DL
<DL
14
.18
0.2
2<D
L<D
L0
.19
<DL
1.2
5<D
L8
9.9
72
20
CA
M0
85
03
4_B
_1C
off
init
e7
3.7
0<D
L0
.12
0.3
8<D
L<D
L<D
L<D
L1
3.0
90
.35
0.1
6<D
L0
.23
<DL
1.0
6<D
L8
9.3
71
2
21
CA
M0
85
03
4_B
_2C
off
init
e7
3.1
00
.18
0.1
40
.31
<DL
<DL
<DL
<DL
13
.22
0.2
50
.38
<DL
0.2
9<D
L1
.03
<DL
89
.22
14
22
CA
M0
85
03
4_B
_3C
off
init
e7
4.0
20
.38
0.2
10
.19
<DL
<DL
<DL
<DL
13
.00
0.3
20
.39
<DL
0.3
0<D
L1
.18
<DL
90
.29
22
23
CA
M0
85
03
4_B
_4C
off
init
e7
3.7
50
.11
<DL
0.2
2<D
L<D
L<D
L<D
L1
4.0
90
.28
0.1
2<D
L0
.17
<DL
1.2
1<D
L9
0.2
71
1
24
CA
M0
85
03
4_C
_1C
off
init
e7
1.8
6<D
L0
.49
0.5
6<D
L0
.16
<DL
<DL
14
.80
0.3
9<D
L<D
L0
.32
<DL
1.2
0<D
L9
0.2
35
1
25
CA
M0
85
03
4_C
_2C
off
init
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7.1
5<D
L0
.29
<DL
<DL
<DL
<DL
<DL
11
.71
0.1
90
.13
<DL
0.3
3<D
L1
.29
<DL
91
.38
28
26
CA
M0
85
03
4_C
_3C
off
init
e7
4.1
9<D
L0
.13
<DL
<DL
<DL
<DL
<DL
13
.87
0.3
7<D
L<D
L0
.26
<DL
1.0
3<D
L9
0.2
71
3
27
CA
M0
85
03
4_C
_4C
off
init
e7
5.9
0<D
L0
.16
0.1
6<D
L<D
L<D
L<D
L1
2.7
00
.68
<DL
<DL
0.2
8<D
L1
.09
<DL
91
.18
16
28
CA
M0
85
00
4_A
_1C
off
init
e7
7.0
80
.24
0.2
30
.48
0.4
40
.18
<DL
<DL
9.3
00
.89
0.2
3<D
L<D
L<D
L0
.46
<DL
89
.73
22
29
CA
M0
85
00
4_A
_2C
off
init
e6
9.6
20
.26
0.3
71
.06
0.7
80
.21
<DL
<DL
12
.34
1.2
40
.42
<DL
0.1
0<D
L0
.45
<DL
87
.13
40
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
161
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
30
CA
M0
85
00
4_A
_3C
off
init
e7
0.4
50
.25
0.1
70
.80
0.6
40
.29
<DL
<DL
11
.56
0.2
70
.15
<DL
0.1
3<D
L0
.50
<DL
85
.31
18
31
CA
M0
85
00
4_B
_1C
off
init
e7
5.0
90
.35
0.4
50
.74
0.4
80
.16
<DL
<DL
10
.22
0.1
6<D
L<D
L<D
L<D
L0
.81
<DL
88
.82
45
32
CA
M0
85
00
4_B
_2C
off
init
e7
4.3
90
.42
0.8
00
.82
0.5
90
.17
<DL
<DL
10
.24
0.2
1<D
L<D
L<D
L<D
L0
.77
<DL
88
.77
80
33
CA
M0
85
00
4_B
_3C
off
init
e7
0.5
70
.37
<DL
1.5
40
.48
0.2
70
.39
<DL
12
.96
0.3
70
.20
<DL
<DL
<DL
0.7
0<D
L8
8.0
04
34
CA
M0
85
00
4_B
_4C
off
init
e7
4.2
90
.33
0.5
90
.74
0.3
90
.25
<DL
<DL
10
.74
0.3
90
.47
<DL
<DL
<DL
0.7
7<D
L8
9.2
25
9
35
CA
M0
85
00
4_B
_5C
off
init
e7
0.4
40
.39
0.5
70
.95
0.4
90
.30
<DL
<DL
12
.64
0.9
80
.36
<DL
<DL
<DL
0.5
8<D
L8
7.7
86
1
36
CA
M0
85
00
4_C
_1C
off
init
e7
2.5
8<D
L<D
L0
.94
0.3
80
.28
<DL
<DL
12
.19
0.2
90
.18
<DL
<DL
<DL
0.6
3<D
L8
7.7
31
0
37
CA
M0
85
00
4_C
_2C
off
init
e6
7.2
7<D
L<D
L1
.43
0.7
30
.30
0.3
6<D
L1
2.7
20
.52
1.3
60
.18
<DL
<DL
1.1
4<D
L8
6.2
30
38
CA
M0
85
00
4_C
_3C
off
init
e6
5.0
40
.12
<DL
1.7
50
.92
0.4
3<D
L<D
L1
5.0
30
.29
0.5
6<D
L<D
L<D
L1
.09
<DL
85
.69
1
39
CA
M0
85
00
4_C
_4C
off
init
e6
3.9
70
.15
<DL
1.8
90
.89
0.5
0<D
L<D
L1
5.2
60
.18
0.8
2<D
L<D
L<D
L0
.96
<DL
85
.01
0
40
CA
M0
85
00
4_C
_5C
off
init
e7
2.9
80
.18
<DL
0.7
00
.37
0.1
8<D
L<D
L1
1.8
10
.80
0.1
7<D
L<D
L<D
L0
.62
<DL
88
.26
9
41
CA
M0
85
01
2_A
_1C
off
init
e6
9.5
9<D
L0
.54
1.0
70
.63
0.3
40
.44
<DL
13
.18
0.2
1<D
L<D
L0
.23
<DL
1.5
0<D
L8
7.7
75
8
42
CA
M0
85
01
2_A
_2C
off
init
e7
4.4
7<D
L0
.89
0.4
90
.45
<DL
<DL
<DL
11
.59
0.2
1<D
L<D
L0
.82
0.1
00
.96
<DL
90
.43
89
43
CA
M0
85
01
2_A
_3C
off
init
e6
8.9
1<D
L<D
L0
.82
0.7
80
.21
<DL
<DL
14
.90
0.3
6<D
L<D
L0
.26
<DL
1.1
6<D
L8
7.8
36
44
CA
M0
85
01
2_A
_4C
off
init
e6
8.4
4<D
L0
.19
0.9
70
.77
0.2
6<D
L<D
L1
4.6
20
.23
<DL
<DL
0.4
7<D
L1
.25
<DL
87
.40
21
45
CA
M0
85
01
2_A
_5C
off
init
e7
2.4
2<D
L0
.47
0.4
60
.87
<DL
<DL
<DL
13
.52
0.4
5<D
L<D
L0
.36
<DL
1.0
3<D
L8
9.8
74
9
46
CA
M0
85
01
2_A
_6C
off
init
e7
6.9
5<D
L1
.93
0.5
80
.39
<DL
<DL
<DL
10
.41
0.1
5<D
L<D
L0
.50
<DL
0.7
2<D
L9
2.1
21
85
47
CA
M0
85
01
2_B
_1C
off
init
e6
9.0
5<D
L0
.26
0.8
10
.72
<DL
<DL
<DL
13
.05
0.2
40
.26
<DL
0.2
5<D
L1
.67
<DL
86
.69
28
48
CA
M0
85
01
2_B
_2C
off
init
e7
0.1
8<D
L0
.48
0.5
70
.56
0.2
20
.36
<DL
11
.82
0.2
3<D
L<D
L0
.47
<DL
1.8
4<D
L8
7.0
45
1
49
CA
M0
85
01
2_B
_3C
off
init
e7
2.4
80
.09
0.5
30
.63
0.6
5<D
L<D
L<D
L1
1.6
60
.49
<DL
<DL
0.2
4<D
L1
.60
<DL
88
.61
55
50
CA
M0
85
01
2_B
_4C
off
init
e6
8.2
7<D
L0
.45
0.8
60
.63
0.3
00
.34
<DL
15
.03
0.2
4<D
L<D
L0
.57
<DL
1.1
3<D
L8
7.9
84
9
51
CA
M0
85
01
2_B
_5C
off
init
e7
8.1
5<D
L0
.55
0.4
20
.42
<DL
<DL
<DL
8.9
70
.30
<DL
<DL
0.5
9<D
L1
.15
<DL
91
.01
52
52
CA
M0
85
01
2_B
_6C
off
init
e7
3.5
4<D
L0
.41
0.4
00
.54
<DL
<DL
<DL
10
.97
1.1
3<D
L<D
L0
.49
<DL
1.0
1<D
L8
8.7
64
2
53
CA
M0
85
01
2_C
_1C
off
init
e7
7.1
0<D
L0
.40
0.3
80
.42
<DL
<DL
<DL
11
.16
0.1
5<D
L<D
L0
.46
0.1
21
.02
<DL
91
.57
39
54
CA
M0
85
01
2_C
_2C
off
init
e7
2.6
4<D
L0
.48
0.7
60
.73
0.3
5<D
L<D
L1
3.4
10
.22
<DL
<DL
0.3
3<D
L0
.94
<DL
90
.07
49
55
CA
M0
85
01
2_C
_3C
off
init
e7
0.1
2<D
L0
.77
0.9
30
.87
0.2
60
.49
<DL
14
.11
0.2
7<D
L<D
L0
.57
<DL
0.8
7<D
L8
9.3
48
2
56
CA
M0
85
01
2_C
_4C
off
init
e7
6.4
2<D
L0
.85
0.3
80
.52
<DL
<DL
<DL
11
.30
0.1
6<D
L<D
L0
.75
<DL
0.6
4<D
L9
1.2
78
3
57
CA
M0
85
01
2_C
_5C
off
init
e7
0.4
3<D
L0
.61
1.0
00
.77
0.3
30
.33
<DL
14
.33
0.3
4<D
L<D
L0
.24
<DL
1.0
0<D
L8
9.5
36
5
58
CA
M0
85
03
8_A
_1C
off
init
e6
9.8
50
.20
0.3
40
.47
0.7
70
.17
<DL
<DL
16
.18
0.3
1<D
L<D
L0
.09
<DL
1.6
8<D
L9
0.0
63
7
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
162
Po
int
Sam
ple
IdM
iner
alU
O2
ThO
2P
bO
Y 2O
3C
e 2O
3G
d2O
3D
y 2O
3Yb
2O
3Si
O2
TiO
2Zr
O2
Mo
O3
FeO
Mn
OC
aOSe
O2
Tota
l
59
CA
M0
85
03
8_A
_2C
off
init
e6
9.0
50
.20
0.4
70
.57
0.8
00
.17
<DL
<DL
15
.96
0.2
5<D
L<D
L<D
L<D
L1
.70
<DL
89
.44
50
60
CA
M0
85
03
8_A
_3C
off
init
e6
8.7
40
.11
0.2
30
.55
0.6
70
.17
<DL
<DL
15
.36
0.1
9<D
L<D
L0
.10
<DL
1.1
1<D
L8
7.4
72
5
61
CA
M0
85
03
8_A
_4C
off
init
e7
3.6
2<D
L0
.16
0.5
00
.56
0.2
0<D
L<D
L1
2.3
80
.22
<DL
<DL
0.2
70
.14
1.1
1<D
L8
9.4
31
7
62
CA
M0
85
03
8_A
_5C
off
init
e6
8.3
10
.12
0.1
90
.54
0.7
50
.28
<DL
<DL
16
.70
0.2
3<D
L<D
L<D
L<D
L1
.01
<DL
88
.33
21
63
CA
M0
85
03
8_A
_6C
off
init
e7
3.1
9<D
L0
.53
0.2
50
.68
<DL
<DL
<DL
12
.93
0.6
3<D
L<D
L0
.20
0.1
10
.97
<DL
89
.66
54
64
CA
M0
85
03
8_A
_7C
off
init
e6
8.7
0<D
L0
.59
0.2
20
.96
<DL
<DL
<DL
16
.24
0.3
1<D
L<D
L0
.10
<DL
0.9
7<D
L8
8.3
26
4
65
CA
M0
85
03
8_A
_8C
off
init
e7
2.9
80
.21
0.7
00
.36
0.8
3<D
L<D
L<D
L1
2.5
20
.37
0.1
2<D
L0
.10
<DL
1.0
6<D
L8
9.5
37
1
66
CA
M0
85
03
8_A
_9C
off
init
e6
9.4
30
.09
<DL
0.4
70
.65
0.1
5<D
L<D
L1
5.8
90
.29
<DL
<DL
<DL
<DL
1.3
0<D
L8
8.5
21
0
67
CA
M0
85
03
8_B
_1C
off
init
e6
7.5
3<D
L0
.79
0.7
40
.69
0.2
1<D
L<D
L1
6.4
30
.29
<DL
<DL
<DL
<DL
1.7
3<D
L8
8.7
78
8
68
CA
M0
85
03
8_B
_2C
off
init
e6
9.9
40
.13
<DL
0.7
10
.64
0.2
2<D
L<D
L1
3.6
70
.21
<DL
<DL
0.1
5<D
L1
.00
<DL
87
.09
12
69
CA
M0
85
03
8_B
_3C
off
init
e7
1.3
0<D
L0
.49
0.7
40
.69
0.3
1<D
L<D
L1
3.1
30
.30
<DL
<DL
0.1
0<D
L1
.04
<DL
88
.43
51
70
CA
M0
85
03
8_B
_4C
off
init
e6
8.5
0<D
L0
.72
0.6
70
.76
0.1
9<D
L<D
L1
6.3
80
.19
<DL
<DL
<DL
<DL
1.5
5<D
L8
9.4
47
8
71
CA
M0
85
03
8_B
_5C
off
init
e6
8.6
0<D
L0
.66
0.8
30
.69
0.2
5<D
L<D
L1
5.9
90
.25
<DL
<DL
<DL
<DL
1.7
2<D
L8
9.1
97
2
72
CA
M0
85
03
8_B
_6C
off
init
e6
8.1
6<D
L0
.11
0.7
50
.67
0.3
0<D
L<D
L1
6.5
80
.23
<DL
<DL
<DL
0.1
11
.48
<DL
88
.49
13
73
CA
M0
85
03
8_B
_7C
off
init
e7
0.0
5<D
L<D
L0
.49
0.5
80
.27
<DL
<DL
15
.81
0.2
8<D
L<D
L<D
L<D
L0
.91
<DL
88
.84
12
74
CA
M0
85
03
8_B
_8C
off
init
e6
4.3
80
.13
0.9
90
.65
0.9
60
.31
<DL
<DL
17
.92
0.2
1<D
L<D
L<D
L<D
L1
.10
<DL
86
.90
11
5
Mic
rop
rob
e d
ata
- o
xid
e w
t. %
Ch
emic
al
Age
(M
a)
163
Point Sample ID Mineral Ni Co Fe Cu Ag As Sb Bi S Se Total
1 CAM085040_1_1 Gersdorffite 35.23 0.23 0.05 <DL <DL 44.73 <DL 0.08 18.79 1.36 100.51
2 CAM085040_1_2 Gersdorffite 35.33 0.61 0.06 <DL <DL 45.32 <DL 0.06 19.49 0.07 100.96
3 CAM085040_1_3 Gersdorffite 27.71 1.05 0.20 1.09 0.05 41.02 1.00 11.01 12.53 2.23 97.90
4 CAM085040_1_4 Gersdorffite 27.01 1.60 0.36 1.44 0.09 39.74 0.93 11.92 13.16 2.43 98.68
5 CAM085040_1_5 Gersdorffite 33.13 1.93 0.45 0.25 <DL 44.61 0.22 0.37 19.02 0.49 100.48
6 CAM085040_1_6 Gersdorffite 31.93 2.96 0.49 0.47 <DL 44.61 0.21 0.63 18.89 0.47 100.66
7 CAM085040_2_1 Gersdorffite 34.00 1.80 0.18 0.04 <DL 45.12 <DL 0.10 19.33 0.06 100.63
8 CAM085040_2_2 Gersdorffite 33.82 1.86 0.16 0.05 <DL 45.20 <DL 0.15 19.43 0.06 100.75
9 CAM085040_2_3 Gersdorffite 26.02 2.44 0.36 1.26 0.10 44.08 0.38 11.55 11.21 2.44 99.83
10 CAM085040_2_4 Gersdorffite 26.20 3.26 0.68 1.53 0.11 42.56 0.40 10.34 13.36 2.16 100.59
11 CAM085040_2_5 Gersdorffite 33.08 1.28 0.54 0.50 <DL 44.36 0.22 0.81 19.10 0.49 100.40
12 CAM085040_1_1 Gersdorffite 35.04 0.80 0.04 <DL <DL 44.96 <DL <DL 19.53 0.07 100.51
13 CAM085040_1_2 Gersdorffite 24.70 4.78 0.12 1.41 0.11 43.99 0.30 10.11 11.89 2.13 99.54
14 CAM085040_2_1 Gersdorffite 35.24 0.38 <DL <DL <DL 45.07 0.02 0.12 19.21 0.60 100.66
15 CAM085040_2_2 Gersdorffite 24.86 4.66 0.33 2.34 0.17 40.70 0.44 10.32 14.31 1.83 99.96
16 CAM085040_3_1 Gersdorffite 34.79 0.87 0.03 <DL <DL 45.04 <DL <DL 19.57 0.05 100.42
17 CAM085040_3_2 Gersdorffite 27.42 2.15 0.13 1.14 0.08 43.21 0.82 9.44 13.20 1.93 99.52
18 CAM085040_3_3 Gersdorffite 32.17 2.76 0.26 0.40 <DL 44.62 0.20 0.56 19.13 0.46 100.56
19 CAM085040_1_1 Gersdorffite 33.69 2.20 0.11 0.04 <DL 45.30 <DL <DL 19.50 <DL 100.94
20 CAM085040_1_2 Gersdorffite 34.04 1.69 0.16 0.07 <DL 45.45 0.02 0.13 19.47 0.09 101.12
21 CAM085040_1_3 Gersdorffite 24.78 5.13 0.24 2.35 0.18 41.55 0.43 9.90 13.88 1.89 100.32
22 CAM085040_1_4 Gersdorffite 26.50 2.27 0.25 1.36 0.09 40.85 0.98 12.44 12.35 2.56 99.65
23 CAM085040_1_5 Gersdorffite 33.52 1.33 0.51 0.29 <DL 44.38 0.22 0.27 19.01 0.46 99.98
24 CAM085040_1_6 Gersdorffite 26.84 4.89 0.59 2.83 0.15 42.88 0.27 4.72 16.99 0.56 100.72
25 CAM085040_1_1 Gersdorffite 34.17 1.68 0.10 0.02 <DL 44.98 <DL 0.07 19.31 0.05 100.40
26 CAM085040_1_2 Gersdorffite 25.47 3.54 0.21 1.57 0.10 41.61 0.58 11.71 12.84 2.25 99.88
27 CAM085040_1_3 Gersdorffite 27.79 4.87 0.47 2.16 0.09 43.83 0.22 3.03 17.85 0.48 100.79
28 CAM085040_1_1 Niccolite 44.45 0.16 0.05 <DL <DL 54.81 0.04 0.81 0.37 0.09 100.79
29 CAM085040_1_2 Niccolite 43.59 0.14 <DL <DL <DL 52.99 0.03 2.59 0.77 0.26 100.38
30 CAM085040_1_3 Niccolite 43.54 0.59 0.20 <DL <DL 55.25 0.56 0.31 0.14 <DL 100.62
31 CAM085040_2_1 Niccolite 43.64 0.49 0.04 <DL <DL 54.71 0.08 1.64 0.25 0.13 100.98
32 CAM085040_2_2 Niccolite 44.83 0.51 0.07 <DL <DL 53.33 1.20 <DL 1.06 <DL 101.06
33 CAM085040_2_3 Niccolite 43.84 0.58 0.45 0.11 <DL 55.31 0.67 0.26 0.20 <DL 101.45
34 CAM085040_3_1 Niccolite 43.18 0.21 0.42 0.14 <DL 52.21 0.02 3.12 0.81 0.46 100.58
35 CAM085040_3_2 Niccolite 42.84 0.13 0.80 0.27 <DL 52.65 0.02 2.72 0.93 0.47 100.83
36 CAM085034_1_1 Gersdorffite 35.64 0.22 0.05 <DL <DL 45.13 0.02 0.07 19.39 <DL 100.52
37 CAM085034_1_2 Gersdorffite 35.70 0.05 <DL <DL <DL 45.30 <DL 0.10 19.29 <DL 100.49
38 CAM085034_1_3 Gersdorffite 35.75 0.03 <DL <DL <DL 45.22 <DL <DL 19.36 <DL 100.46
39 CAM085034_1_4 Gersdorffite 35.46 0.08 0.02 <DL <DL 45.30 <DL 0.07 19.31 <DL 100.28
40 CAM085034_1_5 Gersdorffite 35.66 0.10 <DL <DL <DL 45.27 <DL 0.07 19.40 <DL 100.56
Microprobe data - element weight %
164
Point Sample ID Mineral Ni Co Fe Cu Ag As Sb Bi S Se Total
41 CAM085034_1_6 Gersdorffite 35.28 0.13 0.13 <DL <DL 45.34 <DL 0.08 19.36 <DL 100.34
42 CAM085034_1_7 Gersdorffite 35.55 0.24 0.05 <DL <DL 45.62 <DL <DL 19.16 <DL 100.68
43 CAM085034_1_8 Gersdorffite 35.68 0.09 0.02 <DL <DL 45.32 <DL 0.08 19.26 0.04 100.49
44 CAM085034_1_9 Gersdorffite 35.63 0.07 0.04 <DL <DL 45.15 <DL 0.07 19.47 <DL 100.45
45 CAM085004_1_1 Cobaltite 5.79 29.26 0.75 <DL <DL 45.09 <DL <DL 19.25 0.53 100.73
46 CAM085004_1_2 Cobaltite 6.01 29.04 0.97 0.02 <DL 44.58 <DL <DL 19.37 0.34 100.38
47 CAM085004_2_1 Cobaltite 8.74 27.23 0.33 <DL <DL 45.38 0.16 0.07 18.83 0.33 101.08
48 CAM085004_2_2 Cobaltite 8.33 26.83 0.84 <DL <DL 44.88 <DL 0.08 19.09 0.61 100.68
49 CAM085004_3_1 Cobaltite 4.96 29.86 1.05 0.03 <DL 45.15 <DL 0.07 19.28 0.29 100.70
50 CAM085004_3_2 Cobaltite 7.95 25.79 1.35 0.02 <DL 44.47 <DL 0.05 19.09 0.66 99.38
51 CAM085004_1_1 Cobaltite 6.18 27.65 1.43 0.43 0.06 44.73 <DL 0.79 19.04 0.49 100.80
52 CAM085004_1_2 Cobaltite 7.09 27.87 0.70 0.12 <DL 44.51 <DL 0.20 19.35 0.34 100.21
53 CAM085004_1_3 Cobaltite 6.40 28.77 0.33 0.05 <DL 44.96 <DL 0.06 19.25 0.52 100.34
54 CAM085004_1_1 Cobaltite 5.87 29.46 1.02 <DL <DL 44.95 <DL <DL 19.32 0.42 101.07
55 CAM085004_1_2 Cobaltite 6.20 28.76 1.01 0.06 <DL 45.27 <DL 0.07 19.32 0.41 101.10
56 CAM085004_2_1 Cobaltite 8.51 26.62 0.69 0.03 <DL 44.92 <DL <DL 19.05 0.56 100.42
57 CAM085004_2_2 Cobaltite 3.96 30.81 0.96 0.03 <DL 44.99 <DL <DL 19.18 0.19 100.16
58 CAM085004_2_3 Cobaltite 5.64 30.03 0.42 <DL <DL 45.06 <DL 0.06 19.34 0.32 100.89
59 CAM085012_1_1 Gersdorffite 35.31 0.24 0.26 0.11 <DL 45.65 <DL 0.16 19.21 0.04 100.99
60 CAM085012_1_2 Gersdorffite 35.52 0.08 0.06 0.04 <DL 45.52 <DL 0.38 19.60 <DL 101.22
61 CAM085012_1_3 Gersdorffite 35.60 0.06 <DL <DL <DL 45.51 <DL 0.14 19.40 <DL 100.72
62 CAM085012_1_4 Gersdorffite 35.56 0.07 <DL <DL <DL 45.06 <DL 0.47 19.40 0.05 100.62
63 CAM085012_1_5 Gersdorffite 35.15 0.41 <DL 0.09 <DL 45.49 <DL 0.73 19.07 0.06 101.03
64 CAM085012_1_1 Gersdorffite 35.48 0.12 0.14 0.07 <DL 45.09 <DL 0.14 19.44 0.05 100.54
65 CAM085012_1_2 Gersdorffite 35.67 0.06 <DL <DL <DL 45.45 <DL 0.07 19.48 0.04 100.80
66 CAM085012_1_3 Gersdorffite 35.78 0.08 <DL <DL <DL 44.88 <DL <DL 19.29 0.04 100.16
67 CAM085012_1_4 Gersdorffite 36.17 0.07 <DL <DL <DL 45.94 <DL 0.10 19.48 <DL 101.80
68 CAM085012_1_5 Gersdorffite 35.19 0.16 0.12 0.15 <DL 45.59 <DL 0.12 19.28 0.14 100.76
69 CAM085012_1_1 Gersdorffite 35.65 0.28 0.08 <DL <DL 45.24 <DL <DL 19.58 <DL 100.92
70 CAM085012_1_2 Gersdorffite 35.86 0.06 <DL <DL <DL 45.30 <DL 0.21 19.31 0.05 100.81
71 CAM085012_1_3 Gersdorffite 35.77 0.15 <DL <DL <DL 45.63 0.02 0.09 19.32 <DL 100.98
72 CAM085012_1_4 Gersdorffite 35.49 0.42 <DL <DL <DL 45.44 <DL 0.15 19.02 <DL 100.55
73 CAM085012_1_5 Gersdorffite 35.81 0.07 <DL <DL <DL 45.99 <DL 0.07 19.39 <DL 101.37
74 CAM085012_1_6 Gersdorffite 35.59 0.10 0.06 <DL <DL 45.64 <DL 0.07 19.42 <DL 100.92
75 CAM085012_2_1 Gersdorffite 35.58 0.08 0.03 <DL <DL 45.56 <DL 0.08 19.35 <DL 100.72
76 CAM085012_2_2 Gersdorffite 35.61 0.07 <DL <DL <DL 45.39 <DL 0.12 19.32 <DL 100.55
77 CAM085012_2_3 Gersdorffite 35.62 0.18 <DL <DL <DL 45.27 <DL 0.22 19.24 <DL 100.59
78 CAM085012_2_4 Gersdorffite 35.13 0.06 0.02 <DL <DL 45.38 0.03 0.12 19.31 0.11 100.17
79 CAM085038_1_1 Cobaltite 9.11 26.16 0.36 0.03 <DL 44.67 0.03 0.34 19.55 0.08 100.32
80 CAM085038_1_2 Cobaltite 12.93 22.56 0.32 <DL <DL 44.59 <DL 0.28 19.56 <DL 100.27
Microprobe data - element weight %
165
Point Sample ID Mineral Ni Co Fe Cu Ag As Sb Bi S Se Total
81 CAM085038_1_3 Cobaltite 9.56 26.80 0.33 <DL <DL 44.95 <DL 0.05 19.50 <DL 101.20
82 CAM085038_1_4 Cobaltite 7.11 28.04 0.38 0.02 <DL 44.96 <DL <DL 19.46 <DL 100.03
83 CAM085038_2 1 Cobaltite 12.95 22.89 0.25 <DL <DL 44.80 <DL 0.07 19.75 <DL 100.73
84 CAM085038_2_2 Cobaltite 8.86 27.56 0.43 <DL <DL 45.01 <DL <DL 19.65 <DL 101.55
85 CAM085038_2_3 Cobaltite 12.94 23.51 0.17 <DL <DL 45.38 <DL 0.07 19.28 0.07 101.43
86 CAM085038_2_4 Cobaltite 10.64 25.67 0.30 <DL <DL 44.85 <DL <DL 19.22 0.06 100.79
87 CAM085038_2_5 Cobaltite 11.63 23.82 0.75 <DL <DL 44.10 <DL 0.29 19.62 0.04 100.28
88 CAM085038_2_6 Cobaltite 14.05 22.06 0.29 <DL <DL 44.74 <DL 0.10 19.41 0.06 100.72
89 CAM085038_2_7 Cobaltite 7.71 26.95 0.34 0.03 <DL 45.02 <DL 0.10 19.39 <DL 99.58
90 CAM085038_3_1 Cobaltite 7.56 28.20 0.43 <DL <DL 44.76 <DL <DL 19.51 <DL 100.50
91 CAM085038_3_2 Cobaltite 8.14 28.00 0.35 <DL <DL 45.13 <DL <DL 19.39 <DL 101.06
92 CAM085038_3_3 Cobaltite 13.19 22.72 0.54 <DL <DL 44.76 <DL 0.07 19.46 0.08 100.84
93 CAM085038_3_4 Cobaltite 12.61 21.91 0.36 <DL <DL 44.81 <DL 0.10 19.52 0.08 99.41
Microprobe data - element weight %
166
Ho
leD
ep
thSa
mp
le
Pyr
ite
FeS2
Ge
rdo
rffi
te
NiA
sS
Nic
coli
te
NiA
s
Co
bal
tite
Co
AsS
Ram
me
lsb
erg
ite
NiA
s2
Ch
alco
pyr
ite
Cu
FeS2
Ch
alco
cite
Cu
2S
Bo
rnit
e
Cu
5Fe
S4
Sph
ale
rite
(Zn
,Fe
)S
Gal
en
a
Pb
S
Ura
nin
ite
Co
ffin
ite
342A
438.
234
606
++
342A
440.
334
613
342A
441.
434
615
++
342A
442.
734
619
349
417.
334
756
730
1010
1
349
418
3475
841
2
349
419.
534
761
2
337
428.
150
870
4
337
431.
850
879
1410
337
436.
350
890
38.4
449
.58
348
430.
859
843
22
3
348
431.
159
844
65
14
348
432.
259
847
++
353
433
8006
4
353
433.
780
067
353
440.
280
075
322
12
361
428.
480
106
361
429.
180
107
365
441.
680
144
205
365
442.
880
148
11.1
4
365
443.
880
152
22.8
4
SF76
6_05
433.
880
239
SF76
6_05
435.
280
243
11
SF76
6_05
436.
480
246
2
SF82
6_10
426.
880
778
54
3
SF82
6_10
429.
380
786
314
2
SF82
6_10
430.
980
791
434
3213
SF82
6_10
432.
780
796
363
421.
881
661
11
363
424.
581
668
21
63
363
425.
381
670
3412
XR
D R
IR -
Su
lph
ide
s, A
rse
nid
es
and
Ars
en
ate
s
Appendix D
167
Ho
leD
ep
thSa
mp
le
Pyr
ite
FeS2
Ge
rdo
rffi
te
NiA
sS
Nic
coli
te
NiA
s
Co
bal
tite
Co
AsS
Ram
me
lsb
erg
ite
NiA
s2
Ch
alco
pyr
ite
Cu
FeS2
Ch
alco
cite
Cu
2S
Bo
rnit
e
Cu
5Fe
S4
Sph
ale
rite
(Zn
,Fe
)S
Gal
en
a
Pb
S
Ura
nin
ite
Co
ffin
ite
363
426.
581
673
36
364
433.
381
682
158
364
434.
181
685
149
114
364
437.
181
692
2424
364
438.
281
695
6
369
430.
982
311
1
369
435.
582
321
2
369
436.
782
324
56
14
370
442.
582
353
1.3
0.6
370
444.
782
359
1614
370
446.
282
363
2127
14
SF76
6_13
422.
983
235
9
SF76
6_13
426.
483
243
292
69
SF76
6_13
427.
583
246
++
++
SF76
6_13
429.
183
250
SF81
4_04
431.
783
262
SF81
4_04
433.
383
266
1613
665
SF81
4_04
434
8326
8
SF80
2_13
429.
283
403
2
SF80
2_13
429.
883
405
185
SF80
2_13
430.
683
410
48.
1
SF80
2_13
432.
183
415
23
XR
D R
IR -
Su
lph
ide
s, A
rse
nid
es
and
Ars
en
ate
s
168
Hole Depth Sample Ill ite Chlinochlore Kaolinite Hematite Siderite Calcite Rutile Quartz Goethite Gypsum Boltwoodite
342A 438.2 34606 + + ++ +
342A 440.3 34613 7.2 4.6 8.7 79.5
342A 441.4 34615 73.7 + + 26.3
342A 442.7 34619 6 3.2 90.8
349 417.3 34756 17 27
349 418 34758 57
349 419.5 34761 38.6 25.7 29.7
337 428.1 50870 82 3 8
337 431.8 50879 67 9
337 436.3 50890
348 430.8 59843 88 3
348 431.1 59844 21 64
348 432.2 59847 2 98
353 433 80064 38 8 8 45
353 433.7 80067 42 13 3 43
353 440.2 80075 33 40
361 428.4 80106 40 31 23 6
361 429.1 80107 51 21 24 4
365 441.6 80144 52 5
365 442.8 80148 34.3 20.2
365 443.8 80152 56.4 16.8
SF766_05 433.8 80239 46.5 19 8.9 9 12.9 4
SF766_05 435.2 80243 32 12 38 8
SF766_05 436.4 80246 38 12 42 6
SF826_10 426.8 80778 61 17 10
SF826_10 429.3 80786 63
SF826_10 430.9 80791
SF826_10 432.7 80796 88 11
363 421.8 81661 71 19 7
363 424.5 81668 34
363 425.3 81670 31 18
363 426.5 81673 32 27 3
XRD RIR - Silicates and Oxides
169
Hole Depth Sample Id Ill ite Chlinochlore Kaolinite Hematite Siderite Calcite Rutile Quartz Goethite Gypsum Boltwoodite
364 433.3 81682 49 28
364 434.1 81685 62
364 437.1 81692 51
364 438.2 81695 72 22
369 430.9 82311 66 17 16
369 435.5 82321 29 24 26 19
369 436.7 82324 65 19
370 442.5 82353 3.9 82.2 3 9
370 444.7 82359 70
370 446.2 82363 35
SF766_13 422.9 83235 44 21 26
SF766_13 426.4 83243
SF766_13 427.5 83246 + ++ + +
SF766_13 429.1 83250 84 16
SF814_04 431.7 83262 73 27
SF814_04 433.3 83266
SF814_04 434 83268 35 48 19
SF802_13 429.2 83403 67 22 8
SF802_13 429.8 83405 45 26 6
SF802_13 430.6 83410 52.5 28.3 7.1
SF802_13 432.1 83415 40 37 18
XRD with RIR - Silicates and Oxides
170
Ho
leD
ep
thSa
mp
le
Ura
nin
ite
Co
ffin
ite
Nic
coli
teG
ers
do
rffi
teC
ob
alti
teR
amm
els
be
rgit
eB
ravo
ite
Ch
alco
cite
Ch
alco
pyr
ite
Bo
rnit
eG
ale
na
Sph
ale
rite
Pyr
ite
349
417.
334
756
0.27
1.41
0.00
23.4
117
.27
0.00
1.33
0.00
0.78
0.00
0.21
0.38
6.77
349
418
3475
80.
350.
170.
1569
.32
6.07
0.06
0.23
0.00
0.19
0.00
0.12
0.02
1.36
349
419.
534
761
7.35
36.9
80.
000.
240.
030.
010.
110.
003.
690.
011.
002.
401.
52
337
428.
150
870
0.06
0.18
0.00
0.15
0.30
0.00
0.00
4.67
0.30
3.64
0.01
0.00
0.00
337
431.
850
879
3.09
5.83
0.04
1.72
0.05
0.01
0.00
18.4
80.
4613
.43
0.42
0.00
0.05
348
430.
859
843
0.00
0.01
0.00
0.01
0.00
0.00
0.36
0.00
3.14
0.00
0.21
4.47
3.96
353
440.
280
075
0.73
15.1
40.
002.
281.
590.
040.
430.
007.
090.
010.
780.
561.
12
365
441.
680
144
0.37
2.30
0.01
5.52
0.04
0.00
1.05
0.00
0.68
0.00
1.13
0.00
0.78
365
442.
880
148
1.25
6.82
0.03
15.9
00.
300.
001.
480.
001.
150.
002.
920.
001.
36
365
443.
880
152
4.44
2.19
0.24
14.7
20.
510.
000.
390.
002.
450.
000.
020.
000.
52
SF76
6_05
436.
480
246
0.09
1.58
0.00
0.01
0.00
0.00
0.08
19.2
70.
032.
480.
780.
000.
11
SF82
6_10
430.
980
791
23.3
549
.56
0.00
2.32
1.26
0.00
0.56
0.01
6.12
0.00
0.95
0.00
1.80
SF82
6_10
426.
880
778
19.5
012
.76
0.00
0.01
0.07
0.02
0.00
0.75
2.33
2.93
0.49
0.00
0.00
SF82
6_10
429.
380
786
9.41
18.9
10.
414.
380.
040.
000.
470.
0012
.03
0.01
0.56
0.00
1.40
363
421.
881
661
0.41
3.06
0.22
1.81
0.04
0.00
0.00
3.93
1.22
6.92
0.06
0.00
0.00
363
425.
381
670
0.84
2.01
1.35
18.7
20.
030.
000.
050.
015.
201.
650.
420.
000.
51
363
426.
581
673
0.01
0.07
0.02
2.14
0.07
0.00
0.04
0.00
27.7
80.
010.
110.
000.
07
364
434.
181
685
11.0
623
.00
0.01
0.99
0.13
0.00
0.13
0.00
2.73
0.02
0.86
0.00
0.63
370
444.
782
359
0.06
5.12
7.46
2.21
0.00
0.02
0.08
0.00
0.02
0.00
0.10
0.00
0.06
370
446.
282
363
23.7
829
.22
9.40
16.2
70.
000.
000.
070.
010.
150.
060.
800.
030.
02
SF76
6_13
422.
983
235
0.58
5.47
0.05
0.05
2.61
0.00
0.13
2.53
3.55
4.95
0.17
0.00
0.54
SF80
2_13
429.
883
405
0.07
1.34
1.04
28.5
55.
3011
.13
0.40
0.00
0.92
0.00
0.02
0.00
0.15
SF80
2_13
430.
683
410
0.72
6.00
0.01
2.70
5.77
3.71
0.32
0.00
0.07
0.00
0.08
0.00
0.06
SEM
-MLA
Min
eral
Qu
anti
fica
tio
n
Appendix E
171
Hole Depth Sample Quartz Hematite Chlorite Muscovite Illite Kaolinite Calcite Annabergite/
Erythrite
Rutile
349 417.3 34756 0.05 0.01 36.38 0.01 0.58 0.00 7.72 0.27 0.80
349 418 34758 0.03 0.00 8.03 0.00 0.14 0.00 10.61 1.14 0.11
349 419.5 34761 0.13 1.38 40.95 0.01 0.67 0.16 1.05 0.02 0.91
337 428.1 50870 0.06 1.64 2.24 15.52 69.87 0.00 0.00 0.00 0.99
337 431.8 50879 0.03 1.11 14.25 0.39 38.35 0.02 0.00 0.10 0.93
348 430.8 59843 61.75 0.02 13.14 0.02 0.32 0.03 8.97 0.00 0.93
353 440.2 80075 0.09 0.04 59.56 0.05 3.93 0.07 0.02 0.01 2.78
365 441.6 80144 0.04 0.00 84.35 0.00 1.57 0.01 0.00 0.02 0.90
365 442.8 80148 0.02 0.01 59.23 0.00 1.17 0.02 0.00 0.03 5.31
365 443.8 80152 0.01 0.00 36.00 1.40 35.69 0.00 0.00 0.13 0.93
SF766_05 436.4 80246 2.05 0.01 56.34 0.82 8.17 1.20 0.00 0.00 5.80
SF826_10 430.9 80791 0.05 0.00 12.18 0.00 0.02 0.00 0.05 0.03 0.28
SF826_10 426.8 80778 0.06 8.19 35.08 0.01 14.14 0.07 0.18 0.01 0.87
SF826_10 429.3 80786 0.01 0.00 49.68 0.00 0.01 0.00 0.01 0.04 0.98
363 421.8 81661 0.04 5.02 47.56 0.92 26.59 0.01 0.00 0.11 1.11
363 425.3 81670 0.03 0.01 67.33 0.01 0.94 0.01 0.00 0.18 0.21
363 426.5 81673 0.03 0.00 65.95 0.04 1.44 0.02 0.01 0.01 1.35
364 434.1 81685 0.01 0.01 56.46 0.00 0.05 0.00 0.00 0.01 1.91
370 444.7 82359 0.13 0.68 77.54 0.00 0.13 0.00 0.31 1.17 2.40
370 446.2 82363 0.03 0.33 17.37 0.00 0.01 0.01 0.10 0.53 0.24
SF766_13 422.9 83235 0.02 1.80 74.93 0.00 0.49 0.02 0.19 0.09 1.13
SF802_13 429.8 83405 0.02 0.01 30.13 0.11 8.92 0.16 0.00 6.39 0.68
SF802_13 430.6 83410 0.06 1.16 68.14 0.11 7.25 0.08 0.02 0.87 2.38
SEM-MLA Mineral Quantification
172
Appendix F
173
+
Appendix G
Mineral Normative
The normative algorithm has been created to calculate inferred mineral proportions based on
geochemical stoichiometry. Mineral proportions are calculated for the dominant sulphide and arsenide
minerals identified within the Phase 1 Cigar Lake ore: sphalerite, gersdorffite, niccolite, rammelsbergite,
chalcopyrite, bornite, chalcocite, pyrite and galena. Several geochemical stoichiometric techniques are
utilized in the calculations. Molar element ratios are used to differentiate element control by differing
mineral phases. Minerals with element-constrained ratios, exhibiting the sole control over an element (e.g.
Zn in sphalerite) or tri-element minerals with more than one element-ratio (e.g. Ni:As and Ni:S in
gersdorffite) are calculated first in the linear algorithm. A subtractive method, of calculating the element
consumption by element-constrained minerals before calculating the concentration of non-element
constrained minerals is used to help differentiate between mineral phases. An overview of the normative
algorithm is provided:
Step 1: Calculating Sphalerite using Zn
The concentration of sphalerite is calculated using the whole-rock Zn wt. %. Allotting the Zn to
sphalerite was substantiated by XRD, SEM-MLA and petrographic interpretations that identified
sphalerite as the only controlling mineral phase of Zn. The sphalerite concentration is calculated using the
formula (Zn0.96Fe0.04)S determined from LA-ICP-MS analysis. The sphalerite molecular mass (MM) / Zn
MM ratio was used to quantify the concentration:
Zn wt% * ((MM Zn+ MM S)/ MM Zn) = % sphalerite (1)
Step 2: Calculating concentration of sulpharsenides and sulphadiarsenides
Based on XRD, SEM-MLA and petrographic interpretations the dominant control on Ni-Co-As
throughout the deposit was determined to be: gersdorffite, cobaltite and niccolite (listed in descending
174
order of abundances). Whole-rock geochemistry confirmed these analyses showing that the Ni M% + Co
M%/ As M% ratio is dominated by 1:1 molar ratio of Ni:As. Idealized formulas for niccolite, gersdorffite
and cobaltite were determined reasonable using EMPA and LA-ICP-MS. However, it is noted that solid
solution substitutions exist between the mineral phases. For the normative algorithm the Co M% + Ni
M%: As M% ratio was used to differentiate Ni-As elemental control by sulpharsenides from
sulphadiarsenides. Samples with Co M% +Ni M%: As M% ratios greater than one were calculated as
cobaltite and gersdorffite using Co wt. % and Ni wt. % respectively. Mineral proportion were calculated
using the mineral MM/ element MM ratios as illustrated in Step 1 (equation 1). Overestimation of mineral
proportions, and overconsumption of As is avoided by multiplying the initial Co wt. % and Ni wt. % with
the As M%/ (Ni M% + Co M%) ratio to balance the equation.
Based on XRD, SEM-MLA and petrographic interpretations samples with Co M% + Ni M%: As
M% ratio lower than one contained the sulphadiarsenide rammelsbergite (NiAs2). Ni-Co Skutterrudite
((Ni;Co)As3) was identified but only in minor concentrations and is therefore omitted from the normative.
The concentration of rammelsbergite is calculated using Ni:As ratio linear equations:
Sulphadiarsenides: (2)
= y=mx+b
= Ni M% + Co M% = (1/2) As M% + b
Sulpharsenides:
=y= mx+b
= Ni M% + Co M% = (1/1) As M% + b
175
The intersection point of the sulpharsenide-sulphadiarsenide Ni:As ratios is used to quantify the
whole-rock element control from the minerals. First, the steeper slope (equation 1: sulpharsenides) was
aligned on the axis origin (0, 0). The Y intercept is then calculated for the sample using the linear
equation of the shallower slope (equation 1: sulphadiarsenides):
Solve Y intercept (b) for sulphadiarsenides: (3)
y=mx+b
y=1/2x+b
b=y-1/2x
b= (NiM% + CoM%) – (1/2)(AsM%)
Using the sulphadiarsenide y-intercept, the intersection point between the sulpharsenides and
sulphadiarsenides linear equations can be calculated algebraically providing the stoichiometric control on
the whole-rock geochemistry:
Solving intersection point (X) of the linear equations (Where y=NiM%+CoM% and x = AsM%):
(4)
1/1x+0(sulpharsenide) = 1/2x+b(sulphadiarsenide)
1x(sulpharsenide) -.5x(sulphadiarsenide)=b
.5x=b
X=2b
Therefore:
x(intersection)=2(y-1/2x)
= 2((NiM% + CoM%) – (1/2)AsM%)
176
Since sulpharsenides have a 1:1 NiM%+CoM%:AsM% ratio, the Ni M% + Co M% is equal to
the As M% at the intersection point. All the As above the As M% intersection point is interpreted to be
controlled by the sulphadiarsenide rammelsbergite, whereas all the As below the intersection point is
distributed to the remaining Ni and Co to quantifying the concentration of gersdorffite and cobaltite
respectively. Mineral proportion were calculated using the mineral MM/ element MM ratios as illustrated
in Step 1 (equation 1).
Step 3: Calculating concentration of arsenides
Niccolite with the idealized chemical formula NiAs, could not be initially differentiated based
solely on the NiM%+CoM%/AsM% ratio. Therefore it is allocated if the sample becomes S-deficient
after calculating the normative concentration of sphalerite, preliminary-gersdorffite and cobaltite. Sulphur
consumption is calculated by subtracting the allotted S (mineral wt. % * S MM/ mineral MM) from the
whole-rock S wt. %
Calculating S consumption: (5)
S wt. % remaining = S wt. %-(( S wt. % sphalerite) + (S wt. % gersdorffite) + (S wt. % cobaltite)
This assumes that the only non-sulphide control on Ni or As is niccolite. Based on XRD, SEM-MLA and
petrographic interpretations this has been deemed reasonable. Only minor concentrations of arsenate
minerals annabergite, erythrite and aerugite have been identified. The concentration of niccolite is
calculated by balancing the S deficiency against the Ni content within niccolite.
Calculating niccolite concentration:
Niccolite wt. % =
ABS(S deficiency)* (MM Ni + MM As + MM S)/ MM S) *(MM Ni/ (MM Ni + MM As + MM S) *
(MM Ni + MMAs)/ MM Ni)
The concentration of gersdorffite is recalculated for S-deficient samples after accounting for the
allotment of Ni to niccolite. Mineral proportion were calculated using the mineral MM/ element MM
ratios as illustrated in Step 1 (equation 1). The normative derived niccolite concentration provides a
177
conservative estimate for niccolite since it is based on the S-deficiency prior to the allotment of S to Cu-
bearing sulphides, pyrite and galena.
Step 4: Calculating concentration of Cu Sulphides
Based on XRD, SEM-MLA and petrographic interpretations the dominant control on Cu
throughout the deposit was determined to be: chalcopyrite (CuFeS2), bornite (Cu5FeS4) and chalcocite
(Cu2S) (listed in descending order of abundances). The Cu-bearing mineral phase is derived using the Cu
M%/ S remaining M% after accounting for the S consumption by Zn and the Ni-As phases. If the sample
contains a Cu M%/ S remaining M% ratio of greater than two the mineralogical control on Cu is
interpreted to be from chalcocite (Cu2S). If the sample contains a Cu M%/ S remaining M% ratio of less
than two but greater than 5/4 the mineralogical control on Cu is interpreted to be from chalcocite and
bornite. If the sample contains a Cu M%/ S remaining M% ratio of less than 5/4 but greater than 1/2 the
mineralogical control on Cu is interpreted to be from bornite and chalcopyrite. If the sample contains a
Cu M%/ S remaining M% ratio of less than 1/2 the mineralogical control on Cu is interpreted to be from
solely chalcopyrite.
Relative proportions of the Cu-bearing minerals were quantified using the linear algebraic method
described for in Step 3 equations 2-4. Mineral proportion were again calculated using the mineral MM/
allotted element MM ratios as illustrated in Step 1 (equation 1). Since the Cu-bearing phase is being
calculated prior to the allotment of S to pyrite, pyrrhotite, and galena, there is the potential to overestimate
chalcopyrite at the expense of bornite and chalcocite. This compromise is deemed the most reasonable
since chalcopyrite is the most abundant Cu-bearing mineral.
Step 5: Calculating concentration of Galena
Quantifying the concentration of galena (PbS) results in a lower level of confidence due to the
high and often variable Pb contents within U-bearing and gangue mineral phases. The concentration of
common Pb within the deposit typically accounts for less than 0.5% of the overall concentration of Pb.
Within the deposit, the vast majority of Pb is radiogenic occurring as 206Pb, 207Pb resulting from the decay
178
of U. To account for the concentration of Pb within the U-bearing and gangue mineral phases the bulk-
rock geochemical U/Pb ratio is used to identify Pb-oversaturation indicative of the Pb-bearing sulphide
galena.
Step 6: Calculating Concentration of Pyrite
Pyrite has to be calculated at the end of the normative algorithm since it contains only Fe (II) and
S. Pyrite cannot be differentiated based on the Fe (II) concentration since Fe-chlorite is abundant
throughout the deposit. Pyrite is calculated based on the concentration of remaining S after the metals
have been allotted to sulphides. Therefore the pyrite concentration derived from the normative contains
more error relative to early minerals calculated with the normative.
Step 7: Calculating remaining element proportion
The remaining element concentration are calculated for As, Ni, Co, S, Cu after the elements have
been balanced and the predicted minerals generated. Remaining element concentrations can be used to
evaluate missing mineral phases, incorrect mineral identification and element clay absorption. Remaining
element concentrations are calculated by adding all the consumed element concentration allotted to the
predicted mineral. For example, the remaining Co is calculated:
Co wt. % - (Cobaltite % * Co MM / (Co MM +As MM +S MM))
179
Ho
leD
epth
Sam
ple
Sph
aler
ite
Ram
mel
sber
gite
Co
bal
tite
Nic
colit
eG
ersd
orf
fite
Ch
alco
pyr
ite
Ch
alco
cite
Bo
rnit
eG
alen
aP
yrit
e
ZnS
NiA
s 2C
oA
sSN
iAs
NiA
sSC
uFe
S 2C
u2S
Cu
5Fe
S 4P
bS
FeS 2
34
2A
43
8.2
34
60
60
.00
.00
.00
.00
.00
.00
.00
.00
.10
.0
34
2A
44
0.3
34
61
30
.00
.00
.00
.00
.00
.00
.00
.00
.01
.5
34
2A
44
1.4
34
61
50
.00
.00
.00
.00
.00
.00
.00
.00
.10
.2
34
2A
44
2.7
34
61
90
.00
.00
.00
.00
.00
.00
.00
.00
.00
.2
34
94
17
.33
47
56
0.0
0.0
17
.00
.02
7.5
0.8
0.0
0.0
0.2
4.7
34
94
18
34
75
80
.00
.05
.00
.06
7.1
0.0
0.0
0.0
0.0
0.1
34
94
19
.53
47
61
1.8
0.2
0.0
0.0
0.6
9.1
0.0
0.0
0.5
1.6
33
74
28
.15
08
70
0.0
0.0
0.0
0.0
0.0
0.0
9.2
0.0
0.0
0.0
33
74
31
.85
08
79
0.0
0.0
0.0
0.0
3.8
0.0
26
.73
.60
.10
.0
33
74
36
.35
08
90
1.9
0.0
0.0
0.0
0.0
0.0
0.0
0.0
4.1
0.0
34
84
30
.85
98
43
3.6
0.0
0.0
0.0
0.0
2.3
0.0
0.0
0.0
4.0
34
84
31
.15
98
44
0.0
0.0
0.0
0.0
0.0
3.4
0.0
0.0
1.5
3.3
34
84
32
.25
98
47
0.6
0.0
0.0
0.0
0.0
0.0
0.0
0.0
1.5
2.6
35
34
33
80
06
40
.00
.00
.00
.00
.00
.00
.00
.00
.00
.0
35
34
33
.78
00
67
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.1
35
34
40
.28
00
75
0.4
0.0
1.9
0.0
2.8
9.6
0.0
0.0
1.5
1.1
36
14
28
.44
80
10
60
.00
.00
.00
.00
.00
.00
.00
.00
.00
.3
36
14
29
.18
01
07
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.1
36
54
41
.68
01
44
0.0
0.0
0.0
0.0
5.8
2.9
0.0
0.0
0.6
3.9
36
54
42
.88
01
48
0.0
0.0
0.0
0.0
14
.62
.10
.00
.00
.53
.6
36
54
43
.88
01
52
0.0
0.0
1.0
0.0
12
.62
.10
.00
.10
.00
.0
SF7
66
_05
43
3.8
80
23
90
.00
.00
.00
.00
.00
.00
.00
.00
.20
.0
SF7
66
_05
43
5.2
80
24
30
.00
.00
.00
.00
.00
.00
.00
.00
.00
.0
SF7
66
_05
43
6.4
80
24
60
.00
.00
.00
.00
.00
.06
.92
.70
.00
.0
SF8
26
_10
42
6.8
80
77
80
.00
.00
.00
.00
.01
.40
.07
.20
.00
.0
SF8
26
_10
42
9.3
80
78
60
.00
.00
.00
.06
.11
2.7
0.0
0.0
0.6
0.8
SF8
26
_10
43
0.9
80
79
10
.00
.01
.30
.03
.67
.90
.00
.00
.53
.1
SF8
26
_10
43
2.7
80
79
60
.00
.00
.00
.00
.00
.00
.00
.00
.01
.3
36
34
21
.88
16
61
0.0
0.0
0.0
0.0
2.8
0.0
15
.00
.00
.00
.0
36
34
24
.58
16
68
0.0
0.0
0.0
0.0
0.9
1.6
0.0
1.4
0.0
0.0
36
34
25
.38
16
70
0.0
0.0
0.0
0.0
12
.85
.40
.05
.70
.00
.0
Cal
cula
ted
No
rmat
ive
Min
era
l Pro
po
rtio
ns
fro
m W
ho
le-R
ock
Ge
och
em
istr
yAppendix H
180
Ho
leD
epth
Sam
ple
Sph
aler
ite
Ram
mel
sber
gite
Co
bal
tite
Nic
colit
eG
ersd
orf
fite
Ch
alco
pyr
ite
Ch
alco
cite
Bo
rnit
eG
alen
aP
yrit
e
ZnS
NiA
s 2C
oA
sSN
iAs
NiA
sSC
uFe
S 2C
u2S
Cu
5Fe
S 4P
bS
FeS 2
36
34
26
.58
16
73
0.0
0.4
0.0
0.0
3.3
15
.70
.08
.50
.00
.0
36
44
33
.38
16
82
0.0
0.0
0.0
0.0
5.5
5.5
0.0
0.9
0.0
0.0
36
44
34
.18
16
85
0.0
0.0
0.0
0.0
1.7
5.5
0.0
0.0
1.2
1.7
36
44
37
.18
16
92
0.0
0.0
0.0
0.0
0.8
3.4
0.0
0.0
0.6
0.0
36
44
38
.28
16
95
0.0
0.0
0.0
0.0
0.0
1.8
0.0
0.0
0.2
0.5
36
94
30
.98
23
11
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.1
1.1
36
94
35
.58
23
21
0.0
0.0
0.0
0.0
0.0
5.0
0.0
0.0
0.2
0.5
36
94
36
.78
23
24
0.0
0.0
0.0
0.0
5.3
0.0
7.0
1.4
0.0
0.0
37
04
42
.58
23
53
0.0
0.0
0.0
0.0
3.7
0.0
0.0
0.0
0.0
0.6
37
04
44
.78
23
59
0.0
0.0
0.0
10
.62
.30
.00
.00
.00
.00
.0
37
04
46
.28
23
63
0.0
0.0
0.0
12
.71
4.3
0.0
0.0
0.0
0.0
0.0
SF7
66
_13
42
2.9
83
23
50
.00
.00
.00
.00
.90
.05
.99
.80
.00
.0
SF7
66
_13
42
6.4
83
24
30
.00
.00
.00
.00
.02
.50
.00
.02
.70
.9
SF7
66
_13
42
7.4
58
32
46
0.0
0.0
0.0
0.0
1.7
6.4
0.0
0.0
1.8
1.2
SF7
66
_13
42
9.1
83
25
00
.00
.00
.00
.01
.10
.00
.00
.00
.00
.7
SF8
14
_04
43
1.7
83
26
20
.00
.00
.00
.00
.00
.00
.00
.00
.52
.0
SF8
14
_04
43
3.2
58
32
66
0.0
0.0
0.0
0.0
0.0
0.0
2.7
3.1
0.0
0.0
SF8
14
_04
43
48
32
68
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.0
0.9
SF8
02
_13
42
9.2
83
40
30
.00
.81
.20
.02
.60
.00
.10
.00
.00
.0
SF8
02
_13
42
9.8
83
40
50
.08
.69
.59
.18
.30
.00
.10
.00
.00
.0
SF8
02
_13
43
0.6
83
41
00
.02
.34
.50
.03
.30
.00
.00
.00
.30
.3
SF8
02
_13
43
2.1
83
41
50
.50
.00
.00
.00
.90
.00
.00
.01
.90
.0
Cal
cula
ted
No
rmat
ive
Min
era
l Pro
po
rtio
ns
fro
m W
ho
le-R
ock
Ge
och
em
istr
y
181
Hole Depth Alteration SPP2 SWIR Mineral Interpretation
(CPS) Primary Mineral Secondary Mineral
SF731_01 444.5 MFB4 4CY 3HE CL 450 Paragonite Mg Chlorite
SF731_01 445 MFB4 3CY 3BH CL 200 Paragonite Mg Chlorite
SF731_01 447.5 PELT 3CY 1 CL 1 SE 150 Illitic Paragonite Mg Chlorite
SF731_01 443 MFB4 3BH 2CY 150 Illitic Muscovite
SF731_11 428.7 MFB4 3CL 3CY 800 Muscovite
SF731_11 429.5 MFB4 4CY 950 Muscovite
SF731_11 430.4 MFB4 4APY 1CY 15000 Muscovite Montmorrilonite
SF731_11 431.2 MFB4 4CY 10000 Muscovite Montmorrilonite
SF731_11 432.6 MFB4 4CY 3CL 15000 Muscovite
SF731_11 434.4 PELT 3CY 2SE 1CL 3500 Muscovite
SF731_11 435.5 PELT 3CY 1CL 1SE 400 Muscovite
SF731_19 434.9 MFB4 2CY 1 CL 250 Montmorrilonite
SF731_19 436.5 MFB4 2CY 2CL 900 Muscovite
SF731_19 437.2 MFB4 2CY 3CL 2000 Muscovite Chlorite
SF731_19 438.7 PELT 3CL 2CY 2000 Muscovite Chlorite
SF731_19 439.5 PELT 2CY 3CL 250 Muscovite Chlorite
SF731_19 440.3 PELT 2CY 2CL 150 Muscovite Chlorite
SF737_03 422.8 MFb4 3BH 1CY 300 Muscovite
SF737_03 423.8 MFb4 4CY 2SE CL 300 Muscovite
SF737_03 427.9 MFb4 1CL 1HE 1CY 250 Muscovite
SF737_03 429.5 MFb4 3HE 2CL 1CY 180 Muscovite
SF737_03 430.4 MFb4 4CY 3 CL 3HE 3SE 700 Aspectral Hematite
SF737_03 432 U_Clay 2SIL 1CY 2CL 2APY 3200 Muscovite
SF737_03 432.8 U_Clay 2CL 1CY 3APY 7500 Aspectral
SF737_03 434.5 U_Clay 4CY 3CL 2APY 15000 Aspectral
SF737_03 437.7 U_Clay 3CL 3CY 15000 Illitic Muscovite Montmorrilonite
SF737_03 441.3 U_CLMP 2CY 3CL 2000 Muscovite
SF737_03 441.9 PELT 3BH 3CL 1CY 500 Illitic Muscovite Mg Chlorite
SF737_03 444.8 U_ClMP 3CL 3CY 2APY 3000 Illitic Muscovite Mg Chlorite
SF737_03 445.3 U_CLMP 2SIL 3APY 1CY 2CL 15000 Illitic Muscovite Mg Chlorite
SF737_03 446.8 PELT 1HE 3CL 1BH 2CY 400 Paragonite Mg Chlorite
SF737_03 449.3 GFMP 2CY 3CL 480 Paragonite Mg Chlorite
SF737_07 245.5 MFB4 1APY 1CL 360 Montmorillonite
SF737_07 409.4 MFB4 1CY 3BH 150 Muscovite
SF737_07 416.9 MFB4 3BH 1SE 1CL 4CY 160 Muscovite
SF737_07 419.4 MFB4 2CL 2BH 4CY 1SE 200 Muscovite
SF737_07 421.4 MFB4 3CL 2HE 3CY 750 Muscovite Fe Chlorite
SF737_07 423.8 MFB4 2CL 3SE 4CY 320 Muscovite
SF737_07 424.4 MFB4 2CL 3HE 2APY 130 Muscovite Fe Chlorite
SF737_07 425.9 MFB4 3CL 2HE 3CY 2APY 575 Muscovite Fe Chlorite
SF737_07 428.6 U_MFCL 2CL 3APY 1CL 9250 Muscovite
Lithology/
Facies
Appendix I
182
Hole Depth Alteration SPP2 SWIR Mineral Interpretation
(CPS) Primary Mineral Secondary Mineral
SF737_07 430.1 U_CLAY 3CL 2SE 2BH 3CY 15000 Muscovite Montmorrilonite
SF737_07 432.4 U_CLAY 3HE 2CL 1LI 3CY 1SIL 12500 Muscovite Montmorrilonite
SF737_07 433.4 U_CLAY 3CL 3APY 3CY 2400 Muscovite Aspectral
SF737_07 435.4 PELT 3CL 3CY 325 Muscovite Mg Chlorite
SF737_07 437.7 PEG 2CL 2CY 1SE 2BH 300 Mg Chlorite Muscovite
SF737_07 443.4 GFMP 2CL 3CY 260 Muscovite Mg Chlorite
SF742_10 428 MFB4 3CL 3CY 370 Muscovite Fe Chlorite
SF742_10 428.7 MFB4 3CY 3HE 2950 Aspectral
SF742_10 430.3 MFB4 3CL 3CY 5500 Muscovite Fe Chlorite
SF742_10 431.9 MFB4 4CY 2CL 15000 Illite
SF742_10 432.6 MFB4 3HE 3CY 2CL 15000 Illite Montmorrilonite
SF742_10 433.8 MFB4 3CY 3CL 5500 Illite Montmorrilonite
SF742_10 434.8 PELT 3BH 3CY 2CL 900 Muscovite
SF742_12 425 MFB4 1CL 3CY 2BH 300 Muscovite
SF742_12 428 MFB4 1CL 1HE 4CY 350 Muscovite
SF742_12 428.1 MFB4 2CY 2CL 0HE 350 Muscovite Fe Chlorite
SF742_12 429.7 MFB4 3HE 4CY 3CL 280 Muscovite
SF742_12 431.8 MFB4 3CL 3CY 2SIL 2000 Muscovite Fe Chlorite
SF742_12 432.9 MFB4 2CY 3HE 350 Muscovite
SF742_12 434.2 U_MFCL 3CL 2HE 2CY 2000 Aspectral
SF742_12 434.8 U_MFCL 4CL 2HE 2CY 15000 Aspectral
SF742_12 435.3 MSP 3SIL 4CL 15000 Calcite Fe Chlorite
SF742_12 436 U_MFCL 3CY 3CL 5000 Muscovite Chlorite
SF742_12 437 PELT 4CY 3CL 1SE 700 Muscovite Mg Chlorite
SF742_12 440 PELT 2CY 2BH 2CL 150 Muscovite
SF853_18 439.5 MFb4 3BH 3CY 250 Illite Mg Chlorite
SF853_18 441.5 MFb4 4CY 3CY 3APY 300 Muscovite Mg Chlorite
SF853_18 442.2 CDMP 3CY 3CL 250 Mg Chlorite Paragonite
SF853_18 443 CDMP 3CY 3CL 250 Mg Chlorite Paragonite
SF886_02 440.4 MFB3/4 2-3CL 3CY 420 Paragonite Chlorite
SF886_02 442.5 MFB3 4CY 3CL/APY 800 Paragonite Mg Chlorite
SF886_02 443.5 MFB3/UC 4CY 3CL/APY 250 Illite Mg Chlorite
SF886_02 445.2 PELT 4CY 3SE 200 Illite Montmorrilonite
SF886_10 417.9 MFB4 3BH 2CY 275 Muscovite
SF886_10 420.9 MFB4 3BH 2CY 200 Muscovite
SF886_10 422 MFB4 4CY 3CL 3HE 325 Muscovite
SF886_10 424.3 MFB4 4CY 3SE 1CL SK 750 Muscovite Montmorrilonite
SF886_10 425 MFB4/UCY 4CY 3SE 2CL 2500 Muscovite Montmorrilonite
SF886_10 425.6 MFB4/UCY 4CY 3SE 2CL 3500 Muscovite Montmorrilonite
SF886_10 426.9 U_CLAY 4CY 3CL 15000 Muscovite
SF886_10 428 U_CLAY 4CY 3CL 15000 Muscovite
SF886_10 430.4 U_CLAY 4CY 3CL 15000 Muscovite
Lithology/
Facies
183
Hole Depth Alteration SPP2 SWIR Mineral Interpretation
(CPS) Primary Mineral Secondary Mineral
SF886_10 432.8 PELT 3CY 2CL 700 Muscovite Chlorite
SF886_10 433.5 PELT 2CY 3SE 7500 Muscovite
SF886_10 437 GFPL 3CY 2SE 1CL 300 Muscovite
SF892_04 409.6 MFB4 3BH 1CY 80 Illitic Muscovite
SF892_04 411.3 MFB4 3BH 4CY 80 Illitic Muscovite
SF892_04 417 MFB4 1BH 1CY OAPY 80 Illitic Muscovite
SF892_04 423.4 MFB4 2HE 3 BH 1 CY 100 Illitic Muscovite
SF892_04 426.6 MFB4 2 CY 2HE 2CL 200 Illitic Muscovite
SF892_04 430.7 MFB4 2CL 1APY 2CY 200 Illitic Muscovite Fe Chlorite
SF892_04 434.3 MFB4 1CL 1CY 2 HE 125 Muscovite
SF892_04 436.6 MFB4 4CY 3HE 2SE 250 Muscovite
SF892_04 437.1 UCY 4CY 3CL 0HE 2APY 5000 Illitic Muscovite Fe Chlorite
SF892_04 438.4 UCY 3APY 2CL 2CY 2SIL 12000 Aspectral
SF892_04 439.3 PELT 3CL 3CY 1500 Illite Mg Chlorite
SF892_04 440.6 PELT 3CL 3CY 200 Illite Mg Chlorite
SF892_04 451.2 PELT 2CY 3CL 2SE 100 Illite
SF892_04 459 GFMP 2CY 3CL 2SE 100 Muscovite Mg Chlorite
SF892_04 458.1 GFMP 3CL 2CY 3SE 100 Illite Montmorrilonite
SF892_06 419.1 MFB4 2BH 1CL 1HE 200 Muscovite
SF892_06 421.3 MFB4 2CY 1HE 200 Muscovite
SF892_06 422 U_CLAY 4CY 2CL 1700 Muscovite
SF892_06 424.8 U_CLAY 4CY 4HE 450 Muscovite Montmorrilonite
SF892_06 427.8 MFB4 2CL 3CY 400 Muscovite Fe Chlorite
SF892_06 428.9 U_CLAY 4CY 2CL 2BH 1200 Muscovite
SF892_06 431 U_MSP 3UR 3SI 2CY 15000 Aspectral
SF892_06 433 U_CLAY 4CY 3BH 2CL 15000 Montmorillonite
SF892_06 434 U_CLAY 3CY 2BH 2CL 900 Monmorillonite Mg Chlorite
SF892_06 434.5 U_CHLMP 3CY 3CL 5000 Muscovite
SF892_06 435.7 PELT 3CY 3BH 0CL 200 Illite Montmorrilonite
SF892_06 437 PELT 3CY 3BH 0CL 200 Illitic Muscovite Montmorrilonite
SF892_06 438.6 PELT 2CY 2CL 300 Illitic Muscovite
SF892_06 443.9 PELT 2CY 2CL 250 Paragonite Mg Chlorite
SF892_08 413 MFB4 3BH 1CY 200 Illitic muscovite Montmorrilonite
SF892_08 419 MFB4 2SHE 2APY 1CY 350 Muscovite Aspectral
SF892_08 420.7 MFB4 3CL 3APY 2IN CY 3HE 500 Muscovite Aspectral
SF892_08 422 MFB4 4CY 4HE 3CL 500 Muscovite
SF892_08 423 MFB4 4CY 3HE 3CL 3SE 500 Muscovite Fe Chlorite
SF892_08 424 MFB4 3CL 1SHE 1300 Muscovite Fe Chlorite
SF892_08 425 U_CLAY 3CL 3SE 4CY 9000 Muscovite Fe CHlorite
SF892_08 426.8 U_HMCLAY 4CL 4CY 2SE 10500 Muscovite
SF892_08 427.4 U_MSP 2SIL 4CL 3CY 15000 Muscovite
SF892_08 428.2 U_MSP 3CL 3CL 15000 Muscovite
Lithology/
Facies
184
Hole Depth Alteration SPP2 SWIR Mineral Interpretation
(CPS) Primary Mineral Secondary Mineral
SF892_08 429.7 U_MPC 2CY 3CL 15000 Muscovite
SF892_08 430.2 U_MPC 4CY 3CL 15000 Muscovite
SF892_08 432.1 U_MSP 4CL 15000 Aspectral
SF892_08 432.8 CDPE 3CY 3CL 550 Muscovite Chlorite
SF892_08 434 GFPL 3CL 3CY 250 Muscovite
SF892_12 406.9 MFB4 2BH 3QZD 3CY 150 Illitic Muscovite
SF892_12 415.4 MFB4 2BH 1CY 1QZD 150 Illitic Muscovite
SF892_12 421 MFB4 4CY 200 Illitic Muscovite
SF892_12 427.2 MFB4 4CY 2CL 1HE 200 Illitic Muscovite
SF892_12 429 U_CLAY 4CY 2CL 1HE 1200 Illitic Muscovite Montmorrilonite
SF892_12 430.3 U_MFCHL 3CY 2CL 900 Montmorillonite Fe Chlorite
SF892_12 433.3 U_CHLMP 3CL 2BH 1CY 3000 Muscovite
SF892_12 434.4 PELT 2BH 2CY 1CL 200 Ilitic Muscovite
SF892_12 437 U_CHLMP 3BH 2CY 1CL 1100 Muscovite
SF892_12 439 PELT 3BH 2CY 2CL 200 Illitic Muscovite
SF892_12 441.6 U_CHLMP 3BH 2CY 2CL 3000 Muscovite
SF892_12 444.1 ANAT 3BH 2CY 2CL 150 Muscovite
SF892_12 451.8 GFPL 2BH 2CL 2SAUS 100 Ilitic paragonite
SF892_16 428.2 Mfb4 3CY 3BH 150 Illitic Muscovite
SF892_16 430.7 Mfb4 2ACL 2RHE 150 Muscovite
SF892_16 434.0 Mfb4 3-4 CY STRT 150 Muscovite
SF892_16 435.4 Mfb4 2BH 3CY 200 Illitic Muscovite
SF892_16 437.0 Mfb4 2QZD 1HE 2APY/CL 200 Paragonite Mg Chlorite
SF892_16 438.2 Mfb4 4CY 3CL 900 Paragonite Mg Chlorite
SF892_16 439.7 Mfb4 4CY 3CL 1300 Muscovite Chlorite
SF892_16 440.3 Mfb4 3ACL 400 Paragonite Fe Chlorite
SF892_16 441.8 PELT 4CL 3CY 1100 Muscovite Chlorite
SF892_16 442.5 PELT 2CY 3GCL 500 Illitic Muscovite
SF892_16 444.7 GFPL 2 CY 2CL 400 MgChlorite Paragonite
SF892_16 449.0 GFPL 2CY 2 CL 170 Illitic Muscovite Mg Chlorite
SF892_16 449.9 GFPL 3CL 3CY 170 Illitic Muscovite
SF719_06 414 MFb4 3BH 1CY 150 Muscovite
SF719_06 418.5 MFb4 4BH 2 CY 150 Muscovite
SF719_06 422.8 MFb4 2CY 4 BH 150 Muscovite
SF719_06 426.3 MFb4 2BH 3CY 150 Muscovite
SF719_06 430.8 MFb4 2HE 2CY 150 Muscovite
SF719_06 434 U_Hm 3BRHE 2CL 1100 Muscovite
SF719_06 434.6 Uchl_MF 3CL 15000 Apsectral
SF719_06 438.5 MFb4 2Sil 2CL 2HE 200 Muscovite
SF719_06 440.2 MFb4 2CY 3BH 250 Muscovite
SF719_06 440.5 MFb4 3CY 3QZD 3CL 700 Muscovite
SF719_06 442.9 Pelite 3BH 2CY 2CL 500 Illitic Muscovite Mg Chlorite
Lithology/
Facies
185
Hole Depth Alteration SPP2 SWIR Mineral Interpretation
(CPS) Primary Mineral Secondary Mineral
SF719_06 449 GRMP 1CY <250 Mg Chlorite Paragonite
SF735_16 424.3 2CY 3 BH Muscovite
SF735_16 426 U_Clay 2CY 3 BH 1500 Muscovite
SF735_16 429.5 Mfb4 1CL 2CY 3BH 200 Muscovite Fe Chlorite
SF735_16 430.2 Mfb4 1CL 2CY 3BH 700 Muscovite Aspectral
SF735_16 432.8 Mfb4 1CL 2CY 3BH 200 Muscovite Aspectral
SF735_16 435.7 Mfb4-UCY 4CY 1 CL 2000 Muscovite Mg Chlorite
SF735_16 437 Pelite 2CY 1CL <250 Muscovite Mg Chlorite
SF735_16 438.7 Pelite 1CL 3CY <250 Illitic Paragonite Mg Chlorite
SF735_16 441.9 Pelite 2CY <160 Illitic Muscovite Mg Chlorite
SF735_16 447 Pelite 3CY <100 Illitic Muscovite Mg Chlorite
SF735_16 453.7 GRMP CY <100 Illitic Muscovite Mg Chlorite
SF898_07 423 Mfb4 4CY 4BH 200 Illitic Muscovite
SF898_07 426.9 Mfb4 4CY 1CL 160 Muscovite Fe Chlorite
SF898_07 434 Mfb4 3HE 3CL 3CY 2500 Muscovite Fe Chlorite
SF898_07 435.5 Pelite 2CL 3CY 200 Muscovite
SF898_07 438 Pelite 2CL 3CY 200 Illitic Muscovite
SF898_07 432.5 Pelite 2CL 3CY 1500 Muscovite Chlorite
SF898_07 443.7 Pelite 3CL 3CY 2800 Muscovite Chlorite
SF898_07 444.7 Pelite 3CL 3CY 3400 Muscovite Chlorite
SF898_11 418.8 Mfb4 3CY 1CL 350 Muscovite
SF898_11 420.9 Mfb4 3CY 1CL 550 Muscovite
SF898_11 425.3 UCHl 3CY 3CL 1CB 1000 Muscovite Fe Chlorite
SF898_11 426.2 UCHl 3CY 2 CL 1800 Muscovite Montmorrillonite
SF898_11 430.5 UCHl 3CY 2 CL CB 10000 Muscovite
SF898_11 432.7 Pelite 3CY 2 CL 600 Muscovite Chlorite
SF898_11 437 Pelite 3CY 1 CL 200 Muscovite
SF898_11 443 Pelite 4CY 1CL 200 Muscovite
SF898_11 432.1 UC 3CY 2CL 3500 Muscovite
SF731_13 417.4 3BH 2CY Muscovite
SF731_13 418.5 3BH 2CY 1CL Illitic Muscovite
SF731_13 422.4 1CL 1HE 2CY Aspectral
SF731_13 424.1 3HE 3CY 3CL Illitic Muscovite
SF731_13 426.2 3CL 3CY Muscovite
SF731_13 429 3CL 3CY Muscovite Mg Chlorite
SF731_13 431.3 3CL 3CY Aspectral
SF731_13 433.9 3CL 2CY Phengite
SF731_13 433.4 3CY 2CL Phengite
Lithology/
Facies
186
207P
b/2
06P
b
Hole
Depth
Sam
ple
ID
Min
era
lC
rysta
l Form
207P
b/2
06P
bE
rror
349
419.5
CA
M034761_1
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.4
7E
+00
9.9
9E
-02
1.2
9E
-01
8.5
7E
-03
8.4
0E
-02
1.7
3E
-04
1292
±4
349
419.5
CA
M034761_2
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.1
8E
+00
4.0
0E
-02
1.0
4E
-01
3.3
6E
-03
8.3
2E
-02
1.9
5E
-04
1275
±5
349
419.5
CA
M034761_4
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.2
8E
+00
1.4
4E
-01
1.1
3E
-01
1.2
3E
-02
8.3
1E
-02
3.4
4E
-04
1272
±8
365
443.8
CA
M080152_1
Ura
nin
ite
Bre
ccia
ted
1.9
5E
+00
1.6
8E
-01
1.6
7E
-01
1.4
3E
-02
8.6
3E
-02
9.7
1E
-05
1345
±2
365
443.8
CA
M080152_2
Ura
nin
ite
Bre
ccia
ted
2.1
7E
+00
8.6
6E
-02
1.8
5E
-01
7.4
3E
-03
8.6
6E
-02
9.6
2E
-05
1352
±2
365
443.8
CA
M080152_3
Ura
nin
ite
Bre
ccia
ted
2.3
0E
+00
1.0
2E
-01
1.9
6E
-01
8.7
9E
-03
8.6
3E
-02
5.5
0E
-05
1344
±1
365
443.8
CA
M080152_4
Ura
nin
ite
Bre
ccia
ted
2.1
7E
+00
8.6
2E
-02
1.8
6E
-01
7.4
2E
-03
8.6
0E
-02
8.1
5E
-05
1337
±2
365
443.8
CA
M080152_7
Ura
nin
ite
Bre
ccia
ted
2.1
9E
+00
5.7
3E
-02
1.9
4E
-01
5.1
5E
-03
8.3
3E
-02
6.5
7E
-05
1275
±2
365
443.8
CA
M080152_8
Ura
nin
ite
Bre
ccia
ted
2.3
7E
+00
7.8
9E
-02
2.0
4E
-01
6.8
5E
-03
8.5
8E
-02
8.2
4E
-05
1335
±2
365
443.8
CA
M080152_9
Ura
nin
ite
Bre
ccia
ted
2.3
1E
+00
6.7
6E
-02
2.0
1E
-01
6.0
1E
-03
8.4
6E
-02
6.4
1E
-05
1307
±1
365
443.8
CA
M080152_10
Ura
nin
ite
Bre
ccia
ted
1.9
7E
+00
6.3
2E
-02
1.7
1E
-01
5.5
9E
-03
8.4
7E
-02
6.7
4E
-05
1308
±2
365
443.8
CA
M080152_11
Ura
nin
ite
Bre
ccia
ted
2.5
2E
+00
6.9
7E
-02
2.1
6E
-01
5.7
1E
-03
8.6
5E
-02
3.8
4E
-04
1350
±9
364
434.1
CA
M081685_1
Ura
nin
ite
Bre
ccia
ted
2.2
7E
+00
7.9
4E
-02
2.0
5E
-01
6.9
8E
-03
8.1
6E
-02
9.5
9E
-05
1236
±2
364
434.1
CA
M081685_2
Ura
nin
ite
Bre
ccia
ted
1.7
5E
+00
4.2
8E
-02
1.6
2E
-01
3.7
9E
-03
7.9
4E
-02
1.9
5E
-04
1182
±5
364
434.1
CA
M081685_3
Ura
nin
ite
Bre
ccia
ted
1.9
1E
+00
6.1
7E
-02
1.7
6E
-01
5.5
4E
-03
7.9
8E
-02
1.2
4E
-04
1192
±3
364
434.1
CA
M081685_4
Ura
nin
ite
Bre
ccia
ted
2.2
1E
+00
3.7
1E
-02
1.9
8E
-01
3.6
0E
-03
8.2
5E
-02
1.9
9E
-04
1258
±5
364
434.1
CA
M081685_5
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.6
1E
+00
4.3
8E
-02
1.4
2E
-01
3.7
7E
-03
8.3
7E
-02
9.3
9E
-05
1287
±2
364
434.1
CA
M081685_6
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.7
7E
+00
1.6
6E
-01
1.5
6E
-01
1.4
8E
-02
8.4
3E
-02
2.6
8E
-04
1298
±6
364
434.1
CA
M081685_7
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.7
9E
+00
1.1
2E
-01
1.5
9E
-01
1.1
0E
-02
8.3
7E
-02
3.8
3E
-04
1285
±9
370
446.2
CA
M082363_1
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.8
8E
+00
3.2
1E
-02
1.8
1E
-01
2.8
6E
-03
7.6
8E
-02
1.7
6E
-04
1116
±5
370
446.2
CA
M082363_2
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.6
4E
+00
5.8
1E
-02
1.5
2E
-01
5.3
5E
-03
7.9
6E
-02
1.7
8E
-04
1188
±4
370
446.2
CA
M082363_3
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.3
6E
+00
2.8
4E
-02
1.3
0E
-01
2.7
9E
-03
7.7
3E
-02
1.0
1E
-04
1130
±3
370
446.2
CA
M082363_4
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.6
7E
+00
8.5
5E
-02
1.6
0E
-01
8.0
6E
-03
7.6
7E
-02
7.1
5E
-05
1112
±2
370
446.2
CA
M082363_5
Ura
nin
ite
Bre
ccia
ted/A
ltere
d2.3
7E
+00
2.9
3E
-02
2.1
4E
-01
2.5
5E
-03
8.1
8E
-02
1.4
6E
-04
1240
±3
370
446.2
CA
M082363_6
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.6
8E
+00
4.5
6E
-02
1.5
9E
-01
4.0
9E
-03
7.8
0E
-02
1.5
6E
-04
1148
±4
370
446.2
CA
M082363_7
Coff
inite
Bre
ccia
ted/A
ltere
d1.4
9E
+00
3.2
8E
-02
1.5
7E
-01
3.7
3E
-03
7.0
2E
-02
2.6
5E
-04
936
±8
370
446.2
CA
M082363_8
Coff
inite
Bre
ccia
ted/A
ltere
d2.5
9E
+00
7.5
9E
-02
2.2
7E
-01
6.7
9E
-03
8.4
5E
-02
1.0
9E
-04
1304
±2
370
446.2
CA
M082363_9
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.8
3E
+00
5.2
8E
-02
1.6
8E
-01
5.2
0E
-03
8.1
5E
-02
8.7
7E
-05
1233
±2
370
446.2
CA
M082363_10
Ura
nin
ite
Bre
ccia
ted/A
ltere
d2.3
9E
+00
6.0
4E
-02
2.0
8E
-01
5.3
4E
-03
8.4
4E
-02
7.8
0E
-05
1302
±2
370
446.2
CA
M082363_11
Coff
inite
Bre
ccia
ted/A
ltere
d2.0
6E
+00
7.9
7E
-02
1.8
1E
-01
6.9
1E
-03
8.4
3E
-02
6.7
1E
-05
1298
±2
SF
731_15
433.4
CA
M085012_4
Ura
nin
ite
Botr
yoid
al
2.1
3E
+00
6.2
5E
-02
1.8
9E
-01
5.7
1E
-03
8.3
0E
-02
7.7
3E
-05
1269
±2
SF
731_15
433.4
CA
M085012_5
Ura
nin
ite
Botr
yoid
al/A
ltere
d1.5
4E
+00
1.2
9E
-01
1.4
0E
-01
1.2
0E
-02
8.1
1E
-02
3.4
3E
-04
1224
±8
SF
731_15
433.4
CA
M085012_6
Ura
nin
ite
Botr
yoid
al/A
ltere
d1.6
1E
+00
1.1
1E
-01
1.4
5E
-01
1.0
1E
-02
8.2
1E
-02
2.5
4E
-04
1248
±6
SF
731_15
433.4
CA
M085012_7
Ura
nin
ite
Botr
yoid
al/A
ltere
d1.4
3E
+00
1.5
1E
-01
1.2
7E
-01
1.3
5E
-02
8.3
1E
-02
4.1
3E
-04
1271
±10
LA
-IC
P-M
S U
-Beari
ng
Ph
ase A
naly
sis
Appare
nt ages
207P
b/2
35U
206P
b/2
38U
Std
Err
or
(abs)
Std
Err
or
(abs)
Std
Err
or
(abs)
Appendix J
187
207P
b/2
06P
b
Hole
Depth
Sam
ple
ID
Min
era
lC
rysta
l F
orm
207P
b/2
06P
bE
rror
SF
731_15
433.4
CA
M085012_8
Ura
nin
ite
Botr
yoid
al/A
ltere
d1.7
7E
+00
1.1
0E
-01
1.5
7E
-01
9.9
6E
-03
8.3
5E
-02
2.4
7E
-04
1280
±6
SF
731_15
433.4
CA
M085012_9
Ura
nin
ite
Botr
yoid
al/A
ltere
d1.6
4E
+00
9.7
6E
-02
1.4
6E
-01
8.5
6E
-03
8.3
2E
-02
2.5
1E
-04
1275
±6
SF
731_15
433.4
CA
M085012_10
Coff
inite
Botr
yoid
al/A
ltere
d1.3
4E
+00
6.4
1E
-02
1.1
5E
-01
5.5
5E
-03
8.6
0E
-02
6.7
2E
-05
1338
±2
SF
719_09
436.7
CA
M085024_1
Ura
nin
ite
Alte
red
1.8
2E
+00
3.1
2E
-02
1.6
0E
-01
2.6
3E
-03
8.3
7E
-02
7.3
8E
-05
1286
±2
SF
719_09
436.7
CA
M085024_2
Ura
nin
ite
Alte
red
2.2
6E
+00
5.5
0E
-02
2.0
0E
-01
4.9
8E
-03
8.3
5E
-02
9.0
5E
-05
1282
±2
SF
719_09
436.7
CA
M085024_3
Ura
nin
ite
Alte
red
1.6
3E
+00
3.9
0E
-02
1.6
1E
-01
3.8
8E
-03
7.4
6E
-02
1.0
3E
-04
1058
±3
SF
719_09
436.7
CA
M085024_4
Ura
nin
ite
Alte
red
1.8
6E
+00
5.5
3E
-02
1.7
4E
-01
4.9
3E
-03
7.8
4E
-02
2.7
4E
-04
1158
±7
SF
719_09
436.7
CA
M085024_5
Ura
nin
ite
Alte
red
1.6
6E
+00
4.6
1E
-02
1.6
1E
-01
4.6
1E
-03
7.5
9E
-02
1.1
8E
-04
1092
±3
SF
719_09
436.7
CA
M085024_6
Ura
nin
ite
Alte
red
2.0
9E
+00
8.0
8E
-02
1.8
5E
-01
7.2
6E
-03
8.3
1E
-02
1.0
6E
-04
1271
±2
SF
719_09
436.7
CA
M085024_7
Ura
nin
ite
Alte
red
1.5
8E
+00
5.4
4E
-02
1.5
2E
-01
5.6
5E
-03
7.7
7E
-02
8.5
0E
-05
1139
±2
SF
719_09
436.7
CA
M085024_8
Ura
nin
ite
Alte
red
2.0
8E
+00
3.3
0E
-02
1.9
7E
-01
2.6
1E
-03
7.8
2E
-02
1.7
2E
-04
1152
±4
SF
719_09
436.7
CA
M085024_9
Ura
nin
ite
Alte
red
1.5
3E
+00
4.3
7E
-02
1.3
7E
-01
3.7
4E
-03
8.2
2E
-02
7.2
6E
-05
1250
±2
SF
719_09
436.7
CA
M085024_10
Ura
nin
ite
Alte
red
1.5
1E
+00
6.6
1E
-02
1.3
6E
-01
5.3
2E
-03
8.2
9E
-02
1.9
0E
-04
1266
±4
SF
719_09
436.7
CA
M085024_11
Ura
nin
ite
Alte
red
2.0
8E
+00
8.2
9E
-02
1.8
2E
-01
7.3
9E
-03
8.4
1E
-02
8.3
8E
-05
1295
±2
SF
719_09
436.7
CA
M085024_12
Ura
nin
ite
Alte
red
2.4
1E
+00
5.0
4E
-02
2.1
1E
-01
4.4
8E
-03
8.4
2E
-02
6.0
0E
-05
1297
±1
SF
898_11
428.5
CA
M085035_01
Ura
nin
ite
Botr
yoid
al
1.9
1E
+00
3.1
5E
-02
1.7
4E
-01
2.9
5E
-03
7.9
7E
-02
3.0
1E
-05
1190
±1
SF
898_11
428.5
CA
M085035_02
Ura
nin
ite
Botr
yoid
al
1.7
9E
+00
2.3
3E
-02
1.6
7E
-01
2.1
7E
-03
7.7
5E
-02
4.4
9E
-05
1135
±1
SF
898_11
428.5
CA
M085035_03
Coff
inite
Alte
red
1.7
7E
+00
2.7
0E
-02
1.6
1E
-01
2.4
6E
-03
7.9
5E
-02
4.1
1E
-05
1185
±1
SF
898_11
428.5
CA
M085035_04
Ura
nin
ite
Botr
yoid
al
2.1
1E
+00
2.5
6E
-02
1.8
5E
-01
2.3
1E
-03
8.2
4E
-02
4.5
0E
-05
1254
±1
SF
898_11
428.5
CA
M085035_05
Coff
inite
Alte
red
1.4
1E
+00
2.1
1E
-02
1.2
8E
-01
1.9
1E
-03
7.9
6E
-02
8.2
4E
-05
1187
±2
SF
898_09
430.2
CA
M085039_01
Ura
nin
ite
Botr
yoid
al
1.9
5E
+00
3.9
9E
-02
1.7
6E
-01
3.5
7E
-03
8.0
3E
-02
2.0
1E
-04
1205
±5
SF
898_09
430.2
CA
M085039_02
Ura
nin
ite
Botr
yoid
al
2.0
2E
+00
3.2
9E
-02
1.8
0E
-01
2.8
2E
-03
8.1
5E
-02
6.3
6E
-05
1233
±2
SF
898_09
430.2
CA
M085039_03
Ura
nin
ite
Botr
yoid
al
1.6
5E
+00
4.2
9E
-02
1.5
6E
-01
4.0
5E
-03
7.6
3E
-02
7.8
5E
-05
1103
±2
SF
898_09
430.2
CA
M085039_04
Ura
nin
ite
Botr
yoid
al
2.0
5E
+00
4.0
3E
-02
1.8
5E
-01
3.8
3E
-03
8.0
4E
-02
5.6
1E
-05
1207
±1
SF
898_09
430.2
CA
M085039_05
Ura
nin
ite
Botr
yoid
al/A
ltere
d1.7
0E
+00
5.5
7E
-02
1.5
7E
-01
4.9
2E
-03
7.8
5E
-02
1.2
0E
-04
1160
±3
SF
898_09
430.2
CA
M085039_06
Ura
nin
ite
Botr
yoid
al
2.3
2E
+00
5.6
6E
-02
2.0
1E
-01
5.1
3E
-03
8.3
8E
-02
1.2
6E
-04
1289
±3
SF
898_09
430.2
CA
M085039_07
Ura
nin
ite
Botr
yoid
al
1.8
4E
+00
4.9
9E
-02
1.7
0E
-01
4.3
0E
-03
7.8
2E
-02
2.1
3E
-04
1152
±5
SF
766_10
429.1
CA
M085041_01
Ura
nin
ite
Bre
ccia
ted
1.3
6E
+00
5.9
9E
-02
1.2
6E
-01
5.5
6E
-03
7.8
3E
-02
1.6
1E
-04
1155
±4
SF
766_10
429.1
CA
M085041_02
Ura
nin
ite
Bre
ccia
ted
1.3
1E
+00
5.0
0E
-02
1.2
3E
-01
4.1
3E
-03
7.9
2E
-02
1.2
3E
-04
1178
±3
SF
766_10
429.1
CA
M085041_03
Ura
nin
ite
Bre
ccia
ted
1.7
9E
+00
4.7
1E
-02
1.5
6E
-01
4.2
8E
-03
8.2
9E
-02
6.6
0E
-05
1267
±2
SF
766_10
429.1
CA
M085041_04
Ura
nin
ite
Bre
ccia
ted
1.3
7E
+00
3.9
0E
-02
1.2
5E
-01
3.7
1E
-03
7.9
2E
-02
6.4
6E
-05
1176
±2
SF
766_10
429.1
CA
M085041_06
Ura
nin
ite
Bre
ccia
ted
1.7
3E
+00
6.0
8E
-02
1.5
3E
-01
5.5
2E
-03
8.1
8E
-02
4.6
8E
-05
1240
±1
SF
766_10
429.1
CA
M085041_07
Ura
nin
ite
Bre
ccia
ted
1.6
3E
+00
4.1
1E
-02
1.5
2E
-01
3.2
7E
-03
7.7
5E
-02
2.6
7E
-04
1134
±7
SF
766_10
429.1
CA
M085041_08
Ura
nin
ite
Bre
ccia
ted
1.7
8E
+00
7.0
0E
-02
1.5
7E
-01
6.3
3E
-03
8.2
2E
-02
9.6
9E
-05
1250
±2
SF
776_12
426.7
CA
M085043_01
Ura
nin
ite
Bre
ccia
ted/A
ltere
d2.0
8E
+00
3.0
4E
-02
1.8
1E
-01
2.5
8E
-03
8.3
3E
-02
3.5
6E
-05
1276
±1
LA
-IC
P-M
S U
-Beari
ng
Ph
ase A
naly
sis
207P
b/2
35U
Err
or
Std
(abs)
206P
b/2
38U
Err
or
Std
(abs)
Err
or
Std
(abs)
Appare
nt ages
188
207P
b/2
06P
b
Hole
Depth
Sam
ple
ID
Min
era
lC
rysta
l Form
207P
b/2
06P
bE
rror
SF
776_12
426.7
CA
M085043_02
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.1
1E
+00
2.2
9E
-02
1.1
8E
-01
2.1
8E
-03
6.7
9E
-02
1.3
7E
-04
867
±4
SF
776_12
426.7
CA
M085043_03
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.2
8E
+00
1.1
8E
-02
1.2
7E
-01
1.1
1E
-03
7.2
9E
-02
7.3
3E
-05
1011
±2
SF
776_12
426.7
CA
M085043_04
Ura
nin
ite
Bre
ccia
ted/A
ltere
d1.1
3E
+00
3.4
4E
-02
1.1
6E
-01
3.3
1E
-03
7.0
0E
-02
9.7
4E
-05
929
±3
SF
904_11
427.8
CA
M085045_01
Ura
nin
ite
Vein
/Alte
red
1.3
2E
+00
3.1
2E
-02
1.3
1E
-01
3.1
1E
-03
7.3
1E
-02
1.2
2E
-04
1015
±3
SF
904_11
427.8
CA
M085045_02
Ura
nin
ite
Vein
/Alte
red
1.5
4E
+00
2.7
5E
-02
1.4
5E
-01
2.4
9E
-03
7.7
0E
-02
1.6
5E
-04
1120
±4
SF
904_11
427.8
CA
M085045_03
Ura
nin
ite
Vein
/Alte
red
1.3
5E
+00
3.0
2E
-02
1.2
8E
-01
3.0
5E
-03
7.5
9E
-02
8.5
3E
-05
1092
±2
SF
904_11
427.8
CA
M085045_04
Ura
nin
ite
Vein
/Alte
red
1.3
1E
+00
2.7
0E
-02
1.3
1E
-01
2.9
4E
-03
7.2
1E
-02
1.1
5E
-04
988
±3
SF
904_11
427.8
CA
M085045_05
Ura
nin
ite
Vein
/Alte
red
1.2
6E
+00
1.9
8E
-02
1.3
6E
-01
2.1
9E
-03
6.6
9E
-02
5.0
8E
-05
835
±2
SF
904_11
427.8
CA
M085045_06
Ura
nin
ite
Vein
/Alte
red
1.4
0E
+00
4.0
9E
-02
1.4
4E
-01
3.9
2E
-03
7.0
3E
-02
1.1
6E
-04
936
±3
SF
904_11
427.8
CA
M085045_07
Ura
nin
ite
Vein
Alte
red
1.7
6E
+00
2.9
0E
-02
1.6
2E
-01
2.5
9E
-03
7.8
7E
-02
8.9
3E
-05
1165
±2
SF
904_11
427.8
CA
M085045_08
Ura
nin
ite
Vein
/Alte
red
1.9
2E
+00
3.0
7E
-02
1.7
8E
-01
2.7
3E
-03
7.8
1E
-02
6.4
5E
-05
1149
±2
SF
904_11
427.8
CA
M085045_09
Ura
nin
ite
Vein
/Alte
red
2.1
0E
+00
4.3
4E
-02
1.9
0E
-01
3.5
1E
-03
8.0
0E
-02
1.6
1E
-04
1196
±4
SF
904_11
427.8
CA
M085045_10
Ura
nin
ite
Vein
2.2
4E
+00
5.4
8E
-02
1.9
7E
-01
4.3
9E
-03
8.2
3E
-02
1.7
0E
-04
1254
±4
SF
904_11
427.8
CA
M085045_11
Coff
inite
Vein
/Alte
red
1.7
6E
+00
7.4
2E
-02
1.6
5E
-01
7.1
8E
-03
7.7
5E
-02
1.2
1E
-04
1134
±3
SF
904_11
427.8
CA
M085045_12
Coff
inite
Vein
2.3
1E
+00
3.6
2E
-02
2.0
3E
-01
3.0
8E
-03
8.2
4E
-02
5.6
7E
-05
1255
±1
SF
904_11
427.8
CA
M085045_13
Ura
nin
ite
Vein
2.1
8E
+00
3.7
9E
-02
1.8
8E
-01
3.4
0E
-03
8.4
0E
-02
3.3
6E
-05
1292
±1
SF
904_11
427.8
CA
M085045_15
Ura
nin
ite
Vein
/Alte
red
1.3
8E
+00
4.5
3E
-02
1.3
1E
-01
4.5
0E
-03
7.6
2E
-02
8.1
1E
-05
1099
±2
SF
904_11
427.8
CA
M085045_16
Ura
nin
ite
Vein
/Alte
red
1.4
6E
+00
4.4
3E
-02
1.3
8E
-01
4.3
8E
-03
7.7
5E
-02
1.2
8E
-04
1134
±3
SF
904_11
427.8
CA
M085045_17
Ura
nin
ite
Vein
/Alte
red
1.2
8E
+00
3.4
1E
-02
1.2
9E
-01
3.8
0E
-03
7.2
3E
-02
6.1
6E
-05
995
±2
SF
904_11
427.8
CA
M085045_18
Ura
nin
ite
Vein
/Alte
red
1.0
6E
+00
3.0
1E
-02
1.1
4E
-01
3.4
1E
-03
6.7
1E
-02
8.4
9E
-05
840
±3
SF
904_11
427.8
CA
M085045_19
Ura
nin
ite
Vein
/Alte
red
1.3
5E
+00
4.1
3E
-02
1.2
5E
-01
3.9
7E
-03
7.8
1E
-02
1.4
1E
-04
1150
±4
SF
904_11
427.8
CA
M085045_20
Ura
nin
ite
Vein
/Alte
red
1.8
6E
+00
5.4
7E
-02
1.6
6E
-01
4.6
2E
-03
8.1
1E
-02
1.8
9E
-04
1224
±5
SF
904_11
427.8
CA
M085045_21
Ura
nin
ite
Vein
1.7
4E
+00
5.0
8E
-02
1.5
1E
-01
4.5
6E
-03
8.2
9E
-02
7.0
7E
-05
1268
±2
LA
-IC
P-M
S U
-Beari
ng
Ph
ase A
naly
sis
Appare
nt ages
207P
b/2
35U
Err
or
Std
(abs)
206P
b/2
38U
Err
or
Std
(abs)
Err
or
Std
(abs)
189
Hole Depth Sample ID Analysis Mineral Stage
SF904_14 435.2 CAM052911 1 Sphalerite 0.38 0.99 59.65 22.66 58.88
CAM052911 2 Sphalerite 0.31 0.99 56.05 17.44 55.37
CAM052911 3 Sphalerite 0.33 0.92 57.68 19.30 53.31
CAM052911 4 Sphalerite 0.35 0.89 38.72 13.60 34.52
CAM052911 5 Sphalerite 0.35 1.00 50.52 17.77 50.55
CAM052911 6 Sphalerite 0.35 1.06 50.30 17.62 53.54
CAM052911 7 Sphalerite 0.19 0.41 112.56 20.99 46.46
CAM052911 8 Pyrite PY2 0.08 0.09 2387.23 198.37 212.49
CAM052911 9 Pyrite PY2 0.14 0.23 468.69 63.75 106.27
CAM052911 10 Pyrite PY2 0.19 0.25 220.53 41.93 54.34
CAM052911 11 Pyrite PY2 0.14 0.24 283.45 40.75 66.80
CAM052911 12 Sphalerite 0.21 0.53 97.97 20.12 51.62
CAM052911 13 Sphalerite 0.17 0.32 122.98 21.44 39.63
CAM052911 14 Gersdorffite GER2 0.06 0.00 11181.47 628.32 26.71
CAM052911 15 Gersdorffite GER2 0.07 0.08 1531.69 112.85 125.26
CAM052911 16 Gersdorffite GER2 0.09 0.06 785.99 72.48 49.98
CAM052911 17 Gersdorffite GER2 0.11 0.09 494.22 52.94 46.51
CAM052911 18 Gersdorffite GER2 0.09 0.07 472.68 44.35 33.35
CAM052911 19 Gersdorffite GER2 0.09 0.06 724.72 67.94 46.91
CAM052911 20 Gersdorffite GER2 0.09 0.06 1080.09 93.02 61.85
CAM052911 21 Gersdorffite GER1 0.60 1.37 38.36 22.93 52.69
CAM052911 22 Gersdorffite GER1 0.42 0.51 52.22 21.76 26.44
CAM052911 23 Gersdorffite GER1 0.13 0.36 239.20 31.47 86.19
CAM052911 24 Gersdorffite GER1 0.22 0.53 114.56 24.70 60.52
CAM052911 25 Gersdorffite GER1 0.24 0.79 81.34 19.93 64.51
CAM052911 26 Gersdorffite GER1 0.26 0.59 84.04 21.73 49.80
CAM052911 27 Sphalerite 0.44 1.53 68.78 30.39 105.35
CAM052911 28 Sphalerite 0.34 0.96 49.36 16.65 47.37
CAM052911 29 Sphalerite 0.41 1.37 67.91 27.64 92.77
CAM052911 30 Sphalerite 0.30 1.01 54.62 16.35 54.89
CAM052911 31 Pyrite PY1 0.36 0.98 75.49 27.20 73.92
CAM052911 32 Pyrite PY1 0.34 0.85 61.10 20.62 52.03
CAM052911 33 Pyrite PY1 0.26 0.70 60.13 15.42 42.20
CAM052911 34 Pyrite PY1 0.21 0.55 72.07 14.87 39.62
CAM052911 35 Pyrite PY1 0.32 0.84 85.80 27.80 72.00
CAM052911 36 Pyrite PY2 0.13 0.13 675.62 89.63 89.49
CAM052911 37 Pyrite PY2 0.12 0.22 335.31 40.19 74.60
CAM052911 38 Pyrite PY2 0.10 0.17 223.44 23.40 38.25
CAM052911 39 Pyrite PY2 0.12 0.11 621.25 77.25 67.44
CAM052911 40 Pyrite PY2 0.16 0.27 144.87 22.51 39.48
CAM052911 41 Pyrite PY2 0.20 0.40 154.73 31.66 62.29
CAM052911 42 Pyrite PY2 0.25 0.47 183.26 45.47 86.78
CAM052911 43 Pyrite PY2 0.19 0.40 103.86 19.71 41.69
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
Appendix K
190
Hole Depth Sample ID Analysis Mineral Stage
SF731_15 432 CAM085010 1 Chalcopyrite CPY2 0.12 0.30 137.36 16.77 40.68
CAM085010 2 Chalcopyrite CPY2 0.20 0.32 154.90 30.95 49.62
CAM085010 3 Chalcopyrite CPY2 0.12 0.26 248.37 29.48 64.62
CAM085010 4 Chalcopyrite CPY2 0.12 0.21 113.34 13.83 24.10
CAM085010 5 Chalcopyrite CPY2 0.11 0.23 157.68 17.31 36.56
CAM085010 6 Chalcopyrite CPY2 0.13 0.24 711.93 90.57 174.38
CAM085010 7 Chalcopyrite CPY2 0.10 0.22 344.08 32.82 75.95
CAM085010 8 Chalcopyrite CPY2 0.16 0.25 310.55 48.20 78.62
CAM085010 9 Chalcopyrite CPY2 0.17 0.34 157.40 26.69 54.10
CAM085010 10 Gersdorffite GER1 0.08 0.61 173.99 14.79 105.71
CAM085010 11 Gersdorffite GER1 0.11 0.22 153.27 17.18 34.01
CAM085010 12 Gersdorffite GER2 0.07 0.03 1410.35 94.75 44.45
CAM085010 13 Gersdorffite GER2 0.07 0.03 1503.90 102.78 49.55
CAM085010 14 Gersdorffite GER2 0.07 0.03 1423.35 104.53 46.87
CAM085010 15 Gersdorffite GER2 0.08 0.03 1383.48 104.19 43.04
CAM085010 16 Gersdorffite GER2 0.11 0.11 775.12 85.18 86.21
CAM085010 17 Gersdorffite GER2 0.06 0.07 488.91 31.36 34.78
CAM085010 18 Gersdorffite GER2 0.07 0.19 601.72 43.49 117.17
CAM085010 19 Gersdorffite GER2 0.06 0.07 776.73 50.06 55.68
CAM085010 20 Gersdorffite GER2 0.07 0.08 565.15 41.35 42.97
CAM085010 21 Gersdorffite GER2 0.06 0.07 492.98 30.16 33.14
CAM085010 22 Gersdorffite GER2 0.10 0.13 847.03 85.18 111.80
CAM085010 23 Gersdorffite GER2 0.09 0.10 588.81 51.96 57.44
CAM085010 24 Chalcopyrite CPY2 0.11 0.32 153.44 16.93 48.57
CAM085010 25 Chalcopyrite CPY2 0.12 0.32 221.45 27.09 71.31
CAM085010 26 Chalcopyrite CPY2 0.17 0.38 136.53 22.89 52.50
CAM085010 27 Chalcopyrite CPY2 0.09 0.16 337.67 30.54 53.66
CAM085010 28 Chalcopyrite CPY2 0.06 0.09 764.67 46.06 67.53
CAM085010 29 Chalcopyrite CPY2 0.06 0.13 457.17 25.74 60.56
CAM085010 30 Chalcopyrite CPY2 0.13 0.28 229.85 30.19 65.34
CAM085010 31 Chalcopyrite CPY2 0.09 0.18 218.94 20.09 39.53
SF731_15 432.4 CAM085012 1 Chalcopyrite CPY2 0.13 0.16 211.71 28.53 33.11
CAM085012 2 Chalcopyrite CPY2 0.09 0.03 1760.18 154.35 47.19
CAM085012 3 Chalcopyrite CPY1 0.63 1.29 29.16 18.40 37.55
CAM085012 4 Chalcopyrite CPY1 0.56 1.28 33.87 19.04 43.30
CAM085012 5 Chalcopyrite CPY1 0.66 1.39 35.07 23.17 48.65
CAM085012 6 Chalcopyrite CPY1 0.76 1.62 25.57 19.35 41.48
CAM085012 7 Chalcopyrite CPY1 0.78 1.73 22.62 17.61 39.16
CAM085012 8 Chalcopyrite CPY1 0.79 1.83 19.13 15.08 35.07
CAM085012 9 Chalcopyrite CPY2 0.05 0.02 2199.86 120.87 40.70
CAM085012 10 Chalcopyrite CPY2 0.19 0.33 155.68 29.28 50.87
CAM085012 11 Chalcopyrite CPY2 0.18 0.36 182.39 32.91 66.42
CAM085012 12 Chalcopyrite CPY2 0.11 0.11 439.64 50.10 48.05
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
191
Hole Depth Sample ID Analysis Mineral Stage
SF731_15 432.4 CAM085012 13 Chalcopyrite CPY2 0.14 0.34 96.25 13.89 33.00
CAM085012 14 Chalcopyrite CPY2 0.18 0.38 149.70 27.32 56.61
CAM085012 15 Chalcopyrite CPY2 0.10 0.13 398.97 39.72 50.83
CAM085012 16 Gersdorffite GER1 0.14 0.35 159.68 22.81 56.10
CAM085012 17 Gersdorffite GER1 0.14 0.27 198.76 28.18 53.00
CAM085012 18 Gersdorffite GER1 0.14 0.25 92.60 13.05 23.06
CAM085012 19 Chalcopyrite CPY2 0.18 0.46 105.50 18.57 48.06
CAM085012 20 Chalcopyrite CPY2 0.21 0.44 148.36 30.76 65.85
CAM085012 21 Chalcopyrite CPY2 0.20 0.36 149.44 30.20 53.96
CAM085012 22 Chalcopyrite CPY2 0.12 0.18 231.09 27.09 40.58
CAM085012 23 Chalcopyrite CPY2 0.15 0.21 292.75 44.84 60.46
CAM085012 24 Chalcopyrite CPY2 0.08 0.06 445.54 35.36 25.78
CAM085012 25 Gersdorffite GER2 0.08 0.09 518.25 43.46 48.92
CAM085012 26 Gersdorffite GER1 0.11 0.21 176.00 18.76 37.33
CAM085012 27 Gersdorffite GER1 0.16 0.31 319.04 51.28 99.42
CAM085012 28 Gersdorffite GER1 0.20 0.39 163.66 33.31 63.53
CAM085012 29 Gersdorffite GER2 0.08 0.06 820.97 63.37 47.75
CAM085012 30 Gersdorffite GER2 0.07 0.04 788.12 53.57 35.16
CAM085012 31 Gersdorffite GER2 0.07 0.06 920.54 68.56 59.24
CAM085012 32 Gersdorffite GER2 0.07 0.05 992.30 68.41 48.40
CAM085012 33 Gersdorffite GER2 0.07 0.06 980.67 72.22 58.70
CAM085012 34 Gersdorffite GER2 0.07 0.08 620.19 42.88 48.48
CAM085012 35 Chalcopyrite CPY2 0.08 0.13 311.55 24.78 42.05
CAM085012 36 Chalcopyrite CPY2 0.08 0.12 452.93 35.22 55.26
CAM085012 37 Chalcopyrite CPY2 0.11 0.09 473.91 53.79 44.79
CAM085012 38 Chalcopyrite CPY2 0.08 0.12 636.94 51.33 78.87
CAM085012 39 Chalcopyrite CPY2 0.09 0.13 463.56 43.94 62.21
CAM085012 40 Chalcopyrite CPY2 0.16 0.36 126.10 20.07 45.35
CAM085012 41 Chalcopyrite CPY2 0.18 0.47 125.64 23.17 58.54
CAM085012 42 Chalcopyrite CPY2 0.19 0.55 78.92 14.77 43.02
CAM085012 43 Chalcopyrite CPY2 0.23 0.52 132.50 30.94 68.65
CAM085012 44 Chalcopyrite CPY2 0.19 0.52 144.80 27.08 75.54
CAM085012 45 Chalcopyrite CPY2 0.19 0.39 159.91 30.46 61.76
CAM085012 46 Chalcopyrite CPY2 0.18 0.44 162.30 28.53 71.88
CAM085012 47 Gersdorffite GER2 0.07 0.02 1973.94 143.01 43.81
CAM085012 48 Gersdorffite GER2 0.08 0.05 2230.43 183.52 101.89
CAM085012 49 Gersdorffite GER2 0.07 0.03 1477.22 102.05 46.81
CAM085012 50 Gersdorffite GER2 0.06 0.02 2356.75 151.47 43.39
CAM085012 51 Gersdorffite GER2 0.06 0.01 3651.57 204.84 54.75
CAM085012 52 Gersdorffite GER2 0.05 0.02 2654.17 132.03 40.22
CAM085012 53 Gersdorffite GER2 0.05 0.02 1807.77 92.26 30.40
CAM085012 54 Gersdorffite GER2 0.08 0.05 1123.76 84.92 56.77
CAM085012 55 Gersdorffite GER2 0.08 0.05 803.42 67.15 42.81
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
192
Hole Depth Sample ID Analysis Mineral Stage
SF731_15 432.4 CAM085012 56 Gersdorffite GER2 0.08 0.05 1090.15 86.03 52.66
CAM085012 57 Gersdorffite GER2 0.07 0.05 881.90 61.40 41.14
CAM085012 58 Gersdorffite GER2 0.09 0.06 776.32 71.88 48.22
CAM085012 59 Gersdorffite GER2 0.07 0.06 938.12 70.30 57.89
CAM085012 60 Chalcopyrite CPY2 0.10 0.20 253.10 25.06 50.76
CAM085012 61 Chalcopyrite CPY2 0.11 0.24 212.01 24.02 50.01
CAM085012 62 Chalcopyrite CPY2 0.12 0.28 162.98 19.95 45.09
CAM085012 63 Chalcopyrite CPY2 0.11 0.25 180.26 20.25 45.35
CAM085012 64 Chalcopyrite CPY2 0.09 0.20 163.95 15.05 33.30
CAM085012 65 Chalcopyrite CPY1 0.72 1.52 21.73 15.73 33.08
CAM085012 66 Chalcopyrite CPY1 0.72 1.58 20.19 14.62 31.94
CAM085012 67 Chalcopyrite CPY1 0.76 1.77 17.31 13.16 30.71
CAM085012 68 Chalcopyrite CPY1 0.71 1.81 22.41 15.83 40.55
CAM085012 69 Chalcopyrite CPY1 0.73 2.30 18.69 13.56 42.97
CAM085012 70 Chalcopyrite CPY1 0.69 1.95 19.18 13.15 37.43
CAM085012 71 Chalcopyrite CPY1 0.60 1.41 22.99 13.77 32.46
CAM085012 72 Chalcopyrite CPY1 0.75 1.68 21.48 16.01 36.01
CAM085012 73 Chalcopyrite CPY1 0.86 1.95 21.88 18.71 42.57
CAM085012 74 Chalcopyrite CPY1 0.69 1.75 24.72 16.98 43.30
SF731_16 434.6 CAM085015 1 Pyrite PY1 0.39 1.03 42.09 16.48 43.24
CAM085015 2 Pyrite PY1 0.42 1.21 42.99 18.00 52.14
CAM085015 3 Pyrite PY1 0.48 1.16 42.72 20.69 49.76
CAM085015 4 Pyrite PY1 0.45 1.19 38.23 17.09 45.37
CAM085015 5 Pyrite PY1 0.46 1.22 40.62 18.59 49.62
CAM085015 6 Pyrite PY1 0.43 1.18 39.75 17.24 46.84
CAM085015 7 Pyrite PY1 0.50 1.04 34.12 16.94 35.62
CAM085015 8 Pyrite PY1 0.43 1.01 38.13 16.38 38.63
CAM085015 9 Pyrite PY1 0.44 1.02 43.98 19.47 44.87
CAM085015 10 Pyrite PY1 0.43 1.00 38.81 16.87 38.67
CAM085015 11 Pyrite PY1 0.44 1.11 39.51 17.22 43.94
CAM085015 12 Pyrite PY1 0.43 1.06 39.78 17.07 42.16
CAM085015 13 Pyrite PY1 0.44 1.03 42.09 18.61 43.20
CAM085015 14 Pyrite PY1 0.47 1.06 43.04 20.38 45.66
CAM085015 15 Pyrite PY1 0.44 1.10 38.55 16.78 42.37
CAM085015 16 Pyrite PY1 0.45 1.12 39.29 17.69 44.07
CAM085015 17 Pyrite PY1 0.42 1.02 41.05 17.13 41.90
CAM085015 18 Pyrite PY1 0.41 0.99 45.62 18.67 45.06
CAM085015 19 Pyrite PY1 0.39 0.99 49.24 19.37 48.67
CAM085015 20 Pyrite PY1 0.43 1.06 41.25 17.80 43.78
CAM085015 21 Pyrite PY1 0.50 1.18 42.60 21.21 50.10
CAM085015 22 Pyrite PY1 0.43 1.06 41.91 17.83 44.60
SF719_09 436.7 CAM085024 1 HeCy 0.06 0.04 1060.04 62.00 43.61
CAM085024 2 HeCy 0.07 0.05 1093.48 74.35 59.38
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
193
Hole Depth Sample ID Analysis Mineral Stage
SF719_09 436.7 CAM085024 3 HeCy 0.06 0.05 1182.86 74.97 57.33
CAM085024 4 HeCy 0.06 0.05 1263.49 80.82 56.98
CAM085024 5 HeCy 0.07 0.05 1147.76 76.00 53.40
CAM085024 6 HeCy 0.06 0.05 1118.79 69.63 56.35
CAM085024 7 HeCy 0.06 0.04 1284.25 78.56 51.26
CAM085024 8 HeCy 0.07 0.05 1008.63 66.04 46.69
CAM085024 9 HeCy 0.07 0.06 1203.65 79.35 68.39
CAM085024 10 HeCy 0.07 0.05 1175.05 76.85 58.52
CAM085024 11 HeCy 0.06 0.04 1212.05 76.27 53.40
CAM085024 12 HeCy 0.06 0.05 1221.69 78.62 57.66
CAM085024 13 HeCy 0.06 0.05 1198.71 72.14 54.05
CAM085024 14 HeCy 0.06 0.03 1367.16 80.23 42.60
CAM085024 15 HeCy 0.06 0.03 1963.47 124.67 62.07
CAM085024 16 HeCy 0.06 0.03 1381.79 82.82 40.58
CAM085024 17 HeCy 0.06 0.03 1938.49 121.83 55.57
CAM085024 18 HeCy 0.06 0.03 1896.05 123.18 57.88
CAM085024 19 HeCy 0.06 0.03 1951.64 123.04 61.11
CAM085024 20 HeCy 0.07 0.02 1994.07 133.93 41.93
CAM085024 21 HeCy 0.06 0.02 3186.92 201.44 63.51
CAM085024 22 HeCy 0.07 0.02 2539.88 165.63 45.40
CAM085024 23 HeCy 0.06 0.02 2866.17 183.21 60.46
CAM085024 24 HeCy 0.06 0.02 1994.40 124.76 43.04
CAM085024 25 HeCy 0.07 0.02 2032.84 136.15 44.45
CAM085024 26 HeCy 0.06 0.02 2107.22 135.95 40.68
CAM085024 27 HeCy 0.07 0.02 2207.13 151.56 53.63
CAM085024 28 HeCy 0.06 0.02 1892.70 121.72 43.13
CAM085024 29 HeCy 0.07 0.02 3659.01 249.38 77.22
CAM085024 30 HeCy 0.06 0.02 3049.27 193.66 64.76
CAM085024 31 HeCy 0.07 0.02 2437.33 163.59 47.97
CAM085024 32 Chalcocite 0.05 0.03 2310.37 104.53 74.09
CAM085024 33 Chalcocite 0.05 0.03 1811.50 92.35 48.28
CAM085024 34 Chalcocite 0.05 0.03 1807.36 89.24 48.50
CAM085024 35 Chalcocite 0.05 0.03 1686.44 86.32 52.29
CAM085024 36 Chalcocite 0.05 0.02 2348.24 108.37 58.02
CAM085024 37 Chalcocite 0.05 0.03 1833.06 96.29 55.11
CAM085024 38 Chalcocite 0.05 0.03 1693.20 85.44 48.37
CAM085024 39 Chalcocite 0.05 0.03 1976.57 103.43 61.77
CAM085024 40 Chalcocite 0.06 0.03 1720.21 101.09 60.04
CAM085024 41 Chalcocite 0.05 0.03 1752.03 94.96 53.94
CAM085024 42 Chalcocite 0.07 0.04 1432.33 93.43 52.00
CAM085024 43 Chalcocite 0.06 0.03 1751.40 99.91 59.85
CAM085024 44 Chalcocite 0.06 0.03 1860.54 112.85 60.55
CAM085024 45 Chalcocite 0.07 0.03 1820.29 122.77 59.88
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
194
Hole Depth Sample ID Analysis Mineral Stage
SF719_15 433 CAM085028 1 Cobaltite COB1 0.16 0.39 216.64 34.04 83.55
CAM085028 2 Cobaltite COB1 0.15 0.36 162.35 23.76 58.77
CAM085028 3 Cobaltite COB1 0.15 0.37 174.93 25.99 64.69
CAM085028 4 Cobaltite COB1 0.17 0.44 196.27 32.49 86.44
CAM085028 5 Cobaltite COB1 0.16 0.46 164.70 26.55 76.42
CAM085028 6 Cobaltite COB1 0.15 0.36 161.68 24.09 58.23
CAM085028 7 Cobaltite COB1 0.11 0.37 147.23 16.78 54.61
CAM085028 8 Cobaltite COB1 0.16 0.37 136.87 21.50 50.68
CAM085028 9 Cobaltite COB1 0.16 0.38 144.36 23.26 54.73
CAM085028 10 Cobaltite COB1 0.12 0.31 113.18 13.89 35.12
CAM085028 11 Cobaltite COB1 0.15 0.42 152.90 23.01 63.52
CAM085028 12 Cobaltite COB1 0.16 0.43 158.36 25.03 67.60
CAM085028 13 Cobaltite COB1 0.18 0.58 123.40 22.00 71.45
CAM085028 14 Cobaltite COB1 0.17 0.47 158.97 27.64 74.00
CAM085028 15 Cobaltite COB1 0.17 0.43 168.77 29.03 73.08
CAM085028 16 Cobaltite COB1 0.17 0.49 135.06 23.07 65.82
CAM085028 17 Cobaltite COB1 0.15 0.39 175.81 25.55 67.82
CAM085028 18 Cobaltite COB1 0.18 0.46 149.78 26.60 68.87
CAM085028 19 Cobaltite COB1 0.15 0.38 138.50 21.25 52.27
CAM085028 20 Cobaltite COB1 0.12 0.38 110.30 13.70 42.07
CAM085028 21 Cobaltite COB1 0.16 0.48 137.05 22.04 65.69
CAM085028 22 Cobaltite COB1 0.16 0.45 134.88 21.27 61.01
CAM085028 23 Cobaltite COB1 0.14 0.42 105.62 15.00 44.37
CAM085028 24 Cobaltite COB1 0.16 0.40 170.59 27.87 68.31
CAM085028 25 Cobaltite COB1 0.17 0.47 120.24 21.00 56.09
CAM085028 26 Cobaltite COB1 0.13 0.44 112.98 14.84 49.47
CAM085028 27 Cobaltite COB1 0.15 0.41 151.78 22.73 61.57
CAM085028 28 Cobaltite COB1 0.13 0.30 115.12 14.70 35.06
CAM085028 29 Cobaltite COB1 0.18 0.44 160.59 29.08 70.80
CAM085028 30 Cobaltite COB1 0.17 0.43 108.43 18.35 47.13
CAM085028 31 Cobaltite COB1 0.15 0.43 112.36 16.45 48.50
CAM085028 32 Cobaltite COB1 0.13 0.39 120.82 16.26 47.25
SF776_12 426.7 CAM085043 1 Bornite 0.06 0.00 36741.59 2277.06 32.18
CAM085043 2 Bornite 0.06 0.02 3426.10 193.42 54.90
CAM085043 3 Bornite 0.06 0.00 20042.20 1104.15 28.53
CAM085043 4 Bornite 0.08 0.00 27002.08 2140.20 128.62
CAM085043 5 Bornite 0.08 0.00 8932.92 698.88 26.54
CAM085043 6 Bornite 0.06 0.00 47223.98 2883.37 70.01
CAM085043 7 Bornite 0.06 0.00 18708.48 1052.58 24.99
CAM085043 8 Bornite 0.07 0.00 27657.48 1921.76 44.99
CAM085043 9 Galena 0.07 0.01 5220.01 352.73 40.77
CAM085043 10 Galena 0.06 0.01 6406.10 358.89 44.17
CAM085043 11 Galena 0.05 0.01 6725.85 365.26 48.80
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
195
Hole Depth Sample ID Analysis Mineral Stage
SF776_12 426.7 CAM085043 12 Galena 0.05 0.01 5833.73 315.90 41.15
CAM085043 13 Galena 0.07 0.01 4745.38 340.03 37.61
CAM085043 14 Galena 0.08 0.01 8333.83 634.07 70.92
CAM085043 15 Galena 0.05 0.00 4691.50 225.99 21.53
CAM085043 16 Galena 0.09 0.01 8692.90 822.15 77.73
CAM085043 17 Galena 0.08 0.01 6291.14 485.35 52.14
CAM085043 18 Galena 0.07 0.01 6473.26 468.97 51.89
CAM085043 19 Galena 0.07 0.01 6398.07 424.55 47.03
CAM085043 20 Galena 0.06 0.01 4046.96 262.99 26.11
CAM085043 21 Galena 0.07 0.01 20917.65 1454.66 193.47
CAM085043 22 Bornite 0.06 0.00 72183.68 4235.86 97.73
CAM085043 23 Bornite 0.07 0.00 49439.75 3371.38 65.32
CAM085043 24 Bornite 0.07 0.00 80094.79 5260.78 81.36
CAM085043 25 Bornite 0.07 0.00 88206.43 5853.37 103.09
CAM085043 26 Galena 0.06 0.01 5444.18 352.92 49.49
CAM085043 27 Galena 0.07 0.01 5730.16 372.52 64.96
CAM085043 28 Galena 0.07 0.01 5418.23 387.56 60.53
CAM085043 29 Galena 0.06 0.01 5682.60 355.85 64.19
CAM085043 30 Galena 0.06 0.01 6721.80 426.91 58.03
CAM085043 31 Galena 0.08 0.01 9286.64 705.16 88.72
CAM085043 32 Galena 0.08 0.01 5445.83 427.51 52.39
CAM085043 33 Galena 0.07 0.01 5303.13 395.46 48.38
CAM085043 34 Galena 0.07 0.01 5714.88 425.35 55.41
CAM085043 35 Galena 0.07 0.01 5939.90 431.02 52.08
CAM085043 36 Galena 0.07 0.01 6172.01 445.46 65.13
CAM085043 37 Galena 0.07 0.01 6132.73 438.59 52.02
CAM085043 38 Bornite 0.06 0.00 47617.88 2756.47 208.93
CAM085043 39 Bornite 0.07 0.01 6097.67 414.70 41.58
CAM085043 40 Bornite 0.06 0.00 27427.31 1656.83 125.48
CAM085043 41 Bornite 0.06 0.01 21743.41 1271.28 109.54
CAM085043 42 Bornite 0.06 0.00 23416.09 1339.72 116.52
CAM085043 43 Bornite 0.07 0.01 8251.54 598.48 53.86
CAM085043 44 Bornite 0.07 0.01 11029.43 808.31 65.94
CAM085043 45 Bornite 0.07 0.01 26212.11 1720.05 140.05
CAM085043 46 Bornite 0.07 0.01 10909.66 717.09 58.23
LA-ICP-MS Pb-Isotopes207Pb/ 206Pb
208Pb/ 206Pb
206Pb/ 204Pb
207Pb/ 204Pb
208Pb/ 204Pb
196
Laser-Ablation ICP-MS Data
Sample Mineral La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
(ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm)
CAM085010_06 Coffinite 82 319 25 82 64 24 206 76 438 54 103 12 92 6
CAM085024_04 Coffinite 16 19 6 58 111 35 363 116 715 104 198 29 177 16
CAM085035_02 Coffinite 31 46 9 65 48 21 231 78 549 79 147 19 80 7
CAM085035_03 Coffinite 25 38 8 58 40 15 196 68 457 68 124 16 91 7
CAM085035_04 Coffinite 15 28 6 37 30 9 143 54 369 55 102 14 79 7
CAM085035_05 Coffinite 7 31 6 41 38 12 178 66 470 68 127 15 91 9
CAM085039_06 Coffinite 28 46 11 86 81 23 316 101 690 98 193 26 129 11
CAM085039_08 Coffinite 8 26 8 60 56 16 205 66 447 65 123 16 86 7
CAM085045_01 Coffinite 160 308 29 111 51 18 196 73 460 66 134 17 116 10
CAM085035_01 Uraninite 5 50 11 70 57 23 267 91 580 84 161 22 120 11
CAM085039_01 Uraninite 11 29 8 57 77 22 302 80 550 70 130 15 72 7
CAM085039_04 Uraninite 15 38 11 75 77 23 250 92 603 81 166 21 133 9
CAM085039_05 Uraninite 3 33 9 80 76 22 291 96 653 93 176 21 124 10
CAM085039_09 Uraninite 5 27 9 66 76 23 325 100 679 89 160 17 88 8
CAM085043_01 Uraninite 14 53 12 66 47 18 190 62 391 61 123 18 105 12
CAM085043_02 Uraninite 17 65 15 73 56 17 208 72 498 71 142 20 119 11
CAM085043_04 Uraninite 23 53 11 71 49 16 210 66 422 60 119 17 108 10
CAM085043_05 Uraninite 28 84 14 76 50 21 224 67 455 64 125 16 106 9
CAM085043_06 Uraninite 26 51 9 45 47 16 194 68 439 71 139 18 122 12
CAM085045_02 Uraninite 19 145 42 267 132 42 482 154 1002 151 328 41 273 24
CAM085045_03 Uraninite 21 148 37 220 125 43 478 147 940 140 300 38 258 23
CAM085045_04 Uraninite 25 141 35 220 121 37 444 143 953 147 304 40 264 23
CAM085045_06 Uraninite 21 117 26 162 84 30 378 135 969 144 298 40 262 22
Appendix L
