Geology and Paragenesis of the Boseto Copper Deposits, Kalahari

Home , Ghanzi, Ghanzi District, Kupferschiefer

GEOLOGY AND PARAGENESIS OF THE BOSETO COPPER DEPOSITS,

KALAHARI COPPERBELT, NORTHWEST BOTSWANA

Wesley S. Hall

A thesis submitted to the Faculty and the Board of Trustees of the Colorado School of

Mines in partial fulfillment of the requirements for the degree of Master of Science (Geology)

Golden, Colorado

Date ______Signed: ______Wesley S. Hall

Signed: ______Dr. Murray W. Hitzman Thesis Advisor

Golden, Colorado

Date ______

Signed: ______Dr. John D. Humphrey Associate Professor and Head Department of Geology & Geological Engineering

ABSTRACT

Detailed lithostratigraphic, structural, and petrographic studies coupled with fluid inclusion and stable isotopic analyses and geochronological studies indicate that the Boseto copper deposits formed initially during diagenesis as metalliferous brines ascended along basin faults and moved along a stratigraphic redox boundary between continental red beds and an overlying reduced marine siliciclastic sequence. The hanging wall rocks to copper-silver ore zones comprise comprises a series of at least three stacked coarsening upwards cycles deposited in a deltaic depositional setting. Early copper mineralization may have been accompanied by regionally extensive albitization. Later multiple pulses of faulting and hydrothermal fluid flow associated with a southeast-vergent folding event in the Ghanzi-Chobe belt resulted in extensive networks of bedding-parallel and discordant quartz-carbonate-(Cu-Fe-sulfide) veins. This contractional deformation-related vein and shear system was responsible for significant remobilization of pre-existing vertically and laterally zoned copper sulfide minerals into high- grade zones by hot (250-300˚C), syn-orogenic, metamorphic-derived hydrothermal fluids.

Orientation analysis indicates that the mineralized veins probably formed in association with a flexural slip folding processes. Mineralized vein systems display intense carbonate- chlorite-Cu-Fe-sulfide replacement of wall rock slivers within veins and clasts within shear zones, potassic alteration of the surrounding wall rock, and significant remobilization of early diagenetic disseminated copper sulfide minerals. Sulfur isotopic analyses indicate copper sulfides were probably both mechanically and chemically remobilized.

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TABLE OF CONTENTS

ABSTRACT ...... iii

LIST OF FIGURES ...... vii

LIST OF TABLES ...... x

ACKNOWLEDGEMENTS ...... xi

CHAPTER 1 INTRODUCTION ...... 1

1.1 Objectives ...... 1

1.2 Location and Exploration History ...... 2

1.3 Sedimentary Rock-Hosted Copper Deposits ...... 4

CHAPTER 2 GEOLOGIC BACKGROUND ...... 6

2.1 Regional Basement ...... 6

2.2 Ghanzi-Chobe Belt...... 9

CHAPTER 3 REGIONAL STRUCTURAL GEOLOGY OF THE GHANZI RIDGE ...... 16

3.1 Regional Geophysical Data...... 16

3.2 Regional Structural Geology...... 18

CHAPTER 4 STRATIGRAPHY OF THE BOSETO AREA ...... 23

4.1 Introduction ...... 23

4.2 Ngwako Pan Formation ...... 24

4.3 D’Kar Formation – Lower Member ...... 27

4.4 Stratigraphic Architecture ...... 31

4.5 Paleo-Environmental Reconstruction of the Boseto Area ...... 36

CHAPTER 5 GEOLOGY OF THE BOSETO COPPER DEPOSITS ...... 45

5.1 Introduction ...... 45

5.2 Structural Setting ...... 46

5.3 District-Scale Features of Stratiform Copper-Silver Mineralized Zones ...... 49

5.4 Hydrothermal Alteration Associated with Mineralization ...... 55

5.5 Macroscopic Structures in the Boseto area ...... 62

5.6 The Plutus Deposit ...... 68

5.7 The Zeta Deposit ...... 73

CHAPTER 6 FLUID INCLUSIONS ...... 93

6.1 Introduction ...... 93

6.2 Microthermometry ...... 95

6.3 Crush Leach Analysis ...... 99

CHAPTER 7 STABLE ISOTOPIC ANALYSES ...... 105

7.1 Introduction ...... 105

7.2 Results for Carbon and Oxygen Isotopic Analyses ...... 106

7.3 Sulfur Isotopic Analyses ...... 110

CHAPTER 8 RHENIUM-OSMIUM CHRONOMETRY ...... 114

8.1 Introduction ...... 114

8.2 Re-Os Chronometry Results ...... 115

CHAPTER 9 DISCUSSION ...... 120

9.1 Sedimentary Architecture ...... 120

9.2 Early to Late Diagenetic Stratiform Copper Mineralization ...... 121

9.3 Basin Inversion, Metamorphism, and Structurally Controlled Mineralization ...... 122

9.4 Comparison to Other Sedimentary Rock-Hosted Stratiform Copper Deposits ...... 126

REFERENCES CITED ...... 128

APPENDIX A: LITHOLOGY PHOTOGRAPHS ...... 134

APPENDIX B: STABLE ISOTOPE DATA ...... 139

LIST OF FIGURES

Figure 1.1: Location of the Boseto copper deposits...... 3

Figure 2.1: Lithostratigraphy of the Ghanzi-Chobe Belt ...... 7

Figure 2.2: Regional basement rocks of Botswana ...... 8

Figure 2.3: Precambrian rock exposures in northern Botswana ...... 10

Figure 3.1: Aeromagnetic map of the Ghanzi Ridge area...... 17

Figure 3.2: Geologic map of the Ghanzi Ridge area interpreted from aeromagnetic data...... 21

Figure 3.3: Schematic regional cross section of the Ghanzi Ridge area ...... 22

Figure 4.1: Geologic map of the Boseto area...... 25

Figure 4.2: Representative stratigraphic columns ...... 26

Figure 4.3: Photomicrograph of a pebble conglomerate, Ngwako Pan Formation ...... 28

Figure 4.4: Strike-parallel stratigraphic section of the Plutus deposit...... 34

Figure 4.5: Strike-parallel stratigraphic section of the Zeta deposit...... 37

Figure 4.6: Depositional environments in the Boseto area ...... 44

Figure 5.1: Fold orientation analysis ...... 48

Figure 5.2: Fold orientation analyses ...... 48

Figure 5.3: Typical stratigraphic log depicting up-section ore mineral zonation, Plutus deposit. 52

Figure 5.4: Lateral mineral zonation map, Boseto Copper Deposits...... 53

Figure 5.5: Long section of copper grade distribution, Zeta Deposit...... 54

Figure 5.6: Long section of sulfide-oxide distribution, Zeta Deposit...... 54

Figure 5.7: Diagenetic features, Plutus deposit...... 58

Figure 5.8: Metamorphic minerals and fabrics, Plutus deposit ...... 59

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Figure 5.9: Hydrothermal alteration related to veins and shear zones, Plutus deposit ...... 60

Figure 5.10: Hydrothermal alteration replacement textures ...... 61

Figure 5.11: Axial plane cleavage in the Boseto area ...... 65

Figure 5.12: Veins in the Boseto area ...... 66

Figure 5.13: Shear zones in the Boseto area ...... 67

Figure 5.14: Disseminated sulfide minerals; Plutus deposit and Nexus prospect ...... 74

Figure 5.15: Mineralized nodules, cleavage-parallel lenticles, and stringers, Plutus deposit ...... 75

Figure 5.16: Vein sulfide textures, Plutus deposit ...... 76

Figure 5.17: Orientation analysis of veins, Plutus deposit...... 77

Figure 5.18: Vein orientations at Plutus ...... 78

Figure 5.19: Paragenetic table of structural features observed at Boseto...... 79

Figure 5.20: Foliation and open folds within the ore zone at the Zeta open pit ...... 86

Figure 5.21: Ductile fabrics related to shearing, Zeta deposit ...... 87

Figure 5.22: Foliation at Zeta...... 87

Figure 5.23: Disseminated sulfides and mineralized nodules, Zeta deposit ...... 88

Figure 5.24: Mineralized vein orientations, Zeta deposit ...... 89

Figure 5.25: Classification of boudins ...... 89

Figure 5.26: Examples of boudinage from the Zeta deposit...... 90

Figure 5.27: Wall rock boudinage, Zeta deposit ...... 91

Figure 5.28: Mineralized foliation fabric, Zeta deposit ...... 91

Figure 5.29: Brittle fault system in the hanging wall of the Zeta deposit...... 92

Figure 6.1: Bedding-parallel vein used in fluid inclusion analyses...... 99

Figure 6.2: Primary fluid inclusions used in microthermometry...... 99

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Figure 6.3: Na-Cl-Br systematic of crush-leach data from mineralized veins from Boseto...... 103

Figure 6.4: Na-Cl-Cl-Br systematics of crush-leach data from mineralized veins from Boseto. 104

Figure 7.1: δ18O versus δ13C plot for carbonates, Boseto Cu deposits...... 108

Figure 7.2: Frequency plot of δ34S values by sulfide/sulfate species, Boseto Cu deposits...... 112

Figure 7.4: Stratigraphic variation in δ34S values...... 113

Figure 8.1: Samples used in Re-Os chronometry...... 118

Figure 9.1: Schematic evolution diagram of the Boseto copper deposits...... 125

LIST OF TABLES

Table 6.1: Microthermometry data for primary fluid inclusions analyzed in this study...... 98

Table 6.2: Results of crush-leach extraction analyses, with atomic ratio normalized to Cl...... 102

Table 7.1: Typical carbonate producing processes and accompanying δ13C values...... 109

Table 8.1: Re-Os chronometry data...... 119

ACKNOWLEDGEMENTS

The author thanks his advisor, Dr. Murray Hitzman, and committee members, Dr. Eric

Nelson and D. Thomas Monecke, for their support and guidance is completing this manuscript.

The author acknowledges Dr. Katharina Plaff, director of the Electron Microscopy Laboratory at

Colorado School of Mines, and Dr. Richard Wendlandt, director of the XRD laboratory for their help and time for critical mineralogical analyses. The author would like to thank Dr. Jim

Reynolds of Fluid Inc. (Denver) for his time and efforts regarding fluid inclusion work. Special thanks to Dr. Poul Emsbo of the United States Geological Survey for carrying out crush-leach analyses and providing useful insight for interpretation of the results. Stable isotope studies and interpretations of the data would not have been possible without the help of Dr. John Humphrey, director of the Stable Isotope Laboratory at the Colorado School of Mines and Dr. Craig Johnson of the United States Geological Survey. The author would like to thank Dr. Holy Stein and her assistant Aaron Zimmerman for their work on Re-Os chronometry. The author would like to thank John Skok, laboratory director for the Colorado School of Mines Department of Geology and Geological Engineering, for preparation of thin sections, staining, and tutoring in the use of the scanning electron machine housed at Colorado School of Mines. The preparation this manuscript would not have been possible without the support of Discovery Metals Limited, Ltd., especially Dr. Wallace MacKay, and the exploration team. Their guidance and knowledge of the rocks was especially helpful as a base for this study. Finally, a special thanks to the Society of

Economic Geologist for funding this study through the Hugh E. McKinstry Fund grant.

CHAPTER 1

INTRODUCTION

1.1 Objectives

The Boseto copper deposits are the first deposits within the Kalahari Copperbelt to go into commercial production since the closure of the Klein Aub Mine in Namibia in 1986. Schwartz et al. (1995) conducted the first in depth study at Boseto (previously referred to as the Ngwako

Pan copper deposits). Other published studies conducted in the Ghanzi-Chobe belt described the lithostratigraphy, tectonic environment, and the mineralized horizon in a broad sense (Borg,

1988; Borg and Maiden, 1989; Modie, 1996, 2000; Sillitoe et al., 2010). Previous studies in the

Kalahari Copperbelt have linked stratabound copper-silver ore to early to late diagenetic mineralizing events (Schwartz et al., 1995) as well as epigenetic mineralizing events (Sillitoe et al, 2010, Maiden and Borg, 2011). The Kalahari Copperbelt has received renewed interest since development of the Boseto copper deposits, thus demonstrating a need for a better understanding of how these deposits were formed in order to help focus future exploration programs.

This study investigates the host rock stratigraphy, diagenesis, metamorphism, deformation, alteration, and mineralization at Boseto utilizing data primarily from drill core at the Plutus and

Zeta deposits, which are approximately nine kilometers apart. In addition to standard core logging and petrographic techniques, this study employed fluid inclusion and stable isotope analyses as well as Re-Os chronometry data to place constraints on the mineralizing system and help develop a metallogenic model for the Boseto copper deposits.

1.2 Location and Exploration History

The Boseto copper deposits are located in the Northwest district of Botswana, 80 kilometers southwest of the city of Maun and 20 kilometers southwest of Lake Ngami (Figure

1.1). The deposits lie within the Kalahari Desert, a semi-arid sandy savannah. The majority of the Boseto area displays low relief; an exception is a linear chain of small hills comprising the

Ghanzi Ridge.

Johannesburg Consolidated Investments conducted geological mapping in the Boseto area in the early 1960’s to explore for copper (Van Der Heever et al., 2010). From 1967 to 1970,

Anglovaal South West Africa worked around Cu soil anomalies at what is now Discovery

Metals’ Zeta deposit. Subsequent drilling by U.S. Steel from 1970 to 1980 defined a deposit

(Zeta) with potential reserves of 20 Mt of ore with a grade of 1.74 percent Cu and 39 ppm Ag and discovered the mineralized zone at the Plutus prospect (Schwartz et al., 1995). From 1989 to1994, Anglo American Corporation renewed exploration efforts in the region, discovering the

Banana prospect (102.9 Mt at 1.46 percent Cu and 16.6 g/t Ag) 60 km to the south-west of the

Boseto Copper Project. Further exploration was conducted from 1996 to 2000 by a joint venture involving Delta Gold of Zimbabwe (Delta), Kalahari Gold & Copper (Pty) Ltd of Namibia and

Gencor/BHP Billiton (Van Der Heever et al., 2010).

In 2005, Discovery Metals (Botswana) Limited acquired the property and began a drilling program that defined a mineral resource (measured, indicated, and inferred) of 131.0 Mt at 1.3%

Cu and 16.2 g/t Ag with a cut-off grade of 0.6% Cu and ore reserves of 29.1 Mt at 1.4% Cu and

19.8 g/t Ag (Discovery Metals Limited Report, 2012). Open-pit mining at the Zeta deposit commenced in 2012. Commercial production at the Plutus deposit will start in 2013. Several smaller resources within a 40-km radius of the Boseto copper deposits are currently known and

Figure 1.1: Location of the Boseto copper deposits.

3 will likely supplement the Boseto concentrator plant. The Boseto deposits are the first to have been developed and produced within the Kalahari Copperbelt since the closure of the Klein Aub mine in Namibia in 1987. Success at Boseto has sparked exploration activity throughout the

Kalahari Copperbelt.

1.3 Sedimentary Rock-Hosted Copper Deposits

The Boseto copper deposits are classified as sedimentary rock-hosted stratiform copper deposits (Hitzman et al., 2005). They share many similarities with the world-class

Kupferschiefer deposits of Poland and Germany as well as deposits in the Central African

Copperbelt in Zambia and the Democratic Republic of Congo. Copper sulfides at Boseto occur at the base of a shallow marine mixed siliciclastic-carbonate sequence deposited above continental red beds and bi-modal volcanic rocks. Ore zones contain spatially zoned disseminated and, more importantly, structurally controlled copper sulfide minerals.

Sedimentary rock-hosted stratiform copper deposits comprise disseminated to veinlet Cu and Cu-Fe sulfides in siliciclastic or dolomitic sedimentary rocks. These deposits are the products of evolving basin- or sub-basin-scale fluid-flow systems that include source(s) of metal and sulfur, source(s) of metal- and S-transporting fluids, the transport paths of these fluids, a thermal or hydraulic pump to collect and drive the fluids, and the chemical and physical processes which result in precipitation (trapping) of the sulfides (Hitzman et al., 2005).

Mineralization in these systems is widely accepted to have occurred at any time between diagenesis to basin inversion and metamorphism, or can be a protracted process during this interval. The style, morphology, and mineralogy of most sediment-hosted stratiform copper deposits are remarkably similar despite the large number of variables in the basinal settings of

4 these ore systems (Kirkham, 1989; McGowan et al., 2003; Hitzman et al., 2005; Selley et al.,

2005; Brown 2009).

Host rock stratigraphy in sediment-hosted stratiform copper deposits generally consists of oxidized, hematite-bearing continental red-bed sequences that act as a metal source overlain by reduced marine to lacustrine shale siltstone, sandstone, and dolomite deposited in continental rift settings that act as a trap for metal precipitation (Hitzman et al., 2005). The host sedimentary sequences often contain or preserve evidence of abundant evaporite minerals, indicative of near shore facies in arid or semi-arid environments. Sulfide precipitation was typically controlled by stratigraphic oxidation-reduction boundaries. The actual reductants in the systems were variable ranging from in situ carbonaceous material, pre-existing sulfide minerals, and mobile hydrocarbons, to a combination of these (Hitzman et al., 2005).

Typically, sediment-hosted stratiform copper deposits occur near basement highs adjacent to syn-sedimentary normal faults that control the location of separate sub-basins. The faults provided conduits for ore fluids escaping during basin compaction, inversion, or possibly both.

CHAPTER 2

GEOLOGIC BACKGROUND

2.1 Regional Basement

The northern portion of Botswana can be divided into three Proterozoic belts. From oldest to youngest, these are: 1) the Paleoproterozoic Eburnian Belt consisting of the Kheis and

Magondi Belts and the Okwa Basement Complex, 2) the Mesoproterozoic Kibaran Belt including the Kwando Complex and the late Mesoproterozoic Kgwebe volcanic complex, and 3) the Neoproterozoic to early Paleozoic Pan-African Ghanzi-Chobe Belt (Figure 2.1).

Paleoproterozoic gneissic granitoids and related rocks of the Kheis-Okwa-Magondi belt comprise the southeast portion of the region (Aldiss and Carney, 1992; Ramokate et al., 2000) while Mesoproterozoic gneisses and granite gneiss of the Kwando Complex occupy the northwest portion of the region (Singletary et al., 2003; Figure 2.2). The Boseto Cu deposits are hosted within the Ghanzi-Chobe Belt, a 500-km-long by 100-km-wide deformed volcano- sedimentary basin comprising the basal Kgwebe volcanic complex and the unconformably overlying Ghanzi-Chobe Supergroup sedimentary successions (Figure 2.2).

Rocks of the Ghanzi Group (Figure 2.1) comprise a continental rift sedimentary succession (Modie, 1996). The Boseto Cu deposits are hosted at the base of a marine siliciclastic rock package within the Ghanzi Group. Rocks of the Ghanzi Group were deformed during the

Pan-African Damara Orogen. The unconformably overlying Okwa Group consists of molasse deposited in foreland basins during and after northwest-southeast directed contractional deformation (Ramokate et al., 2000). The Carboniferous to Permian Karoo Supergroup

Figure 2.1: Lithostratigraphy of the Ghanzi-Chobe Belt (modified from Ramokate et al., 2000).

Figure 2.2: Regional basement rocks of Botswana (modified after Singletary et al., 2003).

8 unconformably overlies the deformed host rock package (Figure 2.1). Erosion has stripped

Paleozoic cover off much of the Ghanzi-Chobe Belt. Thus, Cenozoic calcrete and sandstone of the Kalahari Group unconformably overlying the deformed host rock package in most places.

2.2 Ghanzi-Chobe Belt

The Kgwebe volcanic complex is considered the base of the Ghanzi-Chobe volcano- sedimentary basin (Modie, 1996). It is exposed in several inliers containing felsic intrusive rocks with spatially associated mafic igneous rocks that have been dated by the U-Pb zircon method to approximately 1106 Ma (Schwartz et al., 1996; Singletary et al., 2003). Exposures of rhyolitic volcanic rocks (Figure 2.3) include the Kgwebe Hills, Mabeleapodi Hills, and the Groote Laagte area (Lüdtke et al., 1986; Schwartz et al., 1995) in the Ghanzi Ridge area, and the Goha,

Gubatsha, and Chinamba Hills areas in the northeastern portion of the Ghanzi-Chobe Belt

(Carney et al., 1994; Key and Ayres, 2000). Intrusive rocks within the Ghanzi-Chobe Belt include the Kavimba granite and granite near the Chinamba Hills (U-Pb zircon age of 1107 ± 2.1

Ma; U-Pb zircon age of 1107 ± 0.5 Ma, respectively; Singletary et al., 2003). Mafic intrusions within the Kwando Complex have a U-Pb zircon age of 1107 ± 0.8 Ma (Singletary et al., 2003).

In eastern Namibia, the Oorlogsende Porphyry has been dated by U-Pb methods on zircon at

1094 ± 20 Ma (Hegenberger and Burger, 1985).

Within the Ghanzi Ridge area, the Kgwebe volcanic complex attains a maximum known thickness of 2500 m near the Kgwebe Hills east of Boseto and gradually thins to near zero thickness in the southwest based on geophysical interpretations of Schwartz et al. (1995). The lower member of the Kgwebe volcanic complex consists of a bimodal volcanic suite composed of porphyritic rhyolite-dacite flows and tuffs with minor ignimbrites and basaltic flows

Figure 2.3: Precambrian rock exposures in northern Botswana, with locations of Kgwebe volcanic complex rocks and basement borehole locations (modified from Singletary et al., 2003).

10 intercalated with minor arenites (Modie, 1996). The chemical composition of the igneous rocks indicates they are within-plate low titanium-phosphorus (LTP) continental tholeiites and post- orogenic high-K rhyolites (Kampunzu et al., 1998). The chemical composition and field relations suggest that these volcanic rocks were emplaced during a collision-related extensional collapse event (Kampunzu et al., 1998). Volcanism appears to have been concentrated within northeast-southwest elongate sub-basins developed during early phases of extension. The middle and upper members of the complex contain fluvial arkosic sedimentary rocks with paleocurrent data indicating a northeasterly transport direction (Modie, 1996). U-Pb ages from detrital zircons within the arkosic rocks indicate two sediment sources: 1) local erosion of the Kgwebe volcanic complex volcanic rocks, and 2) Paleoproterozoic basement rocks (Kampunzu et al.,

2000).

The Ghanzi-Chobe Supergroup represents a basin-fill package that unconformably overlies the Kgwebe volcanic complex. The sequence attains maximum stratigraphic thickness of 13,500 meters thick near the Namibian border (Litherland, 1982; Modie, 1996) and thins to

5000 meters thick in the study area based on geophysical interpretations (Schwartz et al., 1995;

Modie, 2000). The sedimentary rocks of the Ghanzi Group were probably deposited during renewed rifting, marine incursion, and basin infilling; the sequence is capped by progradational fluvial sediments (Modie, 1996). The Ghanzi Group is formally divided into the Kuke, Ngwako

Pan, D’Kar, and Mamuno Formations, in ascending stratigraphic order (Modie et al., 1998). A major unconformity exists between the Ghanzi and overlying Okwa groups, indicating deposition of the Ghanzi Group likely ceased much earlier. Deposition of the Ghanzi Group is bracketed between 1104 ± 16 Ma and 627 ± 6 Ma by ion microprobe U-Pb ages from detrital zircons within the Ghanzi Group and the unconformably overlying Okwa Group, respectively

(Kampunzu et al., 2000). The majority of the sedimentary rocks in the Ghanzi Group were sourced locally from the Kgwebe volcanic complex, while exotic zircons may have been sourced from the Kibaran-aged (ca. 1400-1300 Ma) Choma-Kalomo Block in Zambia and granitoids exposed in northern Namibia and southern Angola (Singletary et al., 2003). The Ghanzi Group is correlated with the Klein Aub, Doornport, and Eskadron formations of the Lower Damara

Sequence in Namibia, all of which host stratiform copper occurrences (Figure 1.1; Watters,

1977; Maiden et al., 1984; Borg, 1988a, b; Borg and Maiden, 1989; Modie, 2000).

The basal Kuke Formation of the Ghanzi Group consists of a basal conglomerate containing clasts derived from the underlying volcanic sequence and sandstone that contains grains that suggest an input from extra-basinal sources (Modie, 2000). Continental red beds of the Ngwako Pan Formation overlie the Kuke Formation. The Ngwako Pan Formation varies in thickness from near zero adjacent to basement highs, to over 4,500 meters thick in the southwest portion of the area, based on aeromagnetic interpretations (Litherland, 1982; Modie, 1996). The lower member of the Ngwako Pan Formation consists of poorly sorted, silty grey sandstone with intercalated purplish-red mudstone. The middle and upper members consists of moderate to well-sorted, red, arkosic sandstone, and minor siltstone that displays ripple marks, thin, parallel and locally graded lamination, and planar cross bedding (Modie, 1996; Ramokate et al, 2000).

Outcrops of well-sorted, trough cross-bedded sandstone with bedding defined by heavy mineral concentrations occur near the center of the Plutus anticline between the Plutus deposit and Nexus prospect. The upper member of the Ngwako Pan Formation consists primarily of well-sorted, red, planar laminated sandstone (Ramokate et al., 2000). The uppermost 10-40 meters of the

Ngwako Pan Formation contains abundant pebble-rich beds throughout the Ghanzi Ridge area; these beds have been interpreted as high-energy fluvial deposits formed in response to renewed

12 basin subsidence (Modie, 1996). Rocks of the Ngwako Pan Formation are interpreted to have been deposited in an axial fluvial system (Modie, 1996) or within lower, middle, and upper shoreface environments adjacent to alluvial and fluvial sources (Master, 2010; Caterall, 2012).

The overlying D’Kar Formation marks a significant transgressive event in the Ghanzi-

Chobe Belt. The D’Kar Formation varies in stratigraphic thickness from 1,500 to 3,000 meters

(Modie, 1996; Ramokate et al., 1998). The contact between the Ngwako Pan Formation and the overlying D’Kar Formation is commonly transitional over a few meters. The lower member of the D’Kar Formation consists of reduced grey-green siltstone, subarkose, arkose, sandstone, and claystone with subordinate limestone, marlstone, and volcaniclastic rocks. These rocks predominantly display a planar parallel lamination that is interpreted to indicate suspension deposition below storm wave base (Modie, 1996). A distinct rhythmite facies consisting of fining upward siltstone and mudstone beds is present in many sections of the lower D’Kar

Formation. This facies is suggestive of deposition in a tidally influenced regime. Periodically intercalated thin and laterally extensive sandstones probably formed during high-energy storm events (Modie, 1996). Carbonate rocks occur in the lower D’Kar Formation throughout the area and attain thicknesses of up to 40 meters to the southwest of Boseto. The carbonates were likely deposited in shallow, warm waters of restricted lagoons or playa lakes (Modie, 1996). The upper member of the D’Kar Formation consists of interstratified reduced and oxidized subarkose, sandstone, and siltstone. Only in the Bothatogo area west of Boseto does the D’Kar Formation consist entirely of reduced siliciclastic sedimentary rocks that are interpreted to have been deposited entirely below wave base (Schwartz et al., 1995).

The overlying Mamuno Formation in Botswana is correlated with the Kamtsas Formation in Namibia (Schwartz et al., 1996). The Mamuno Formation is not present in the study area, but

13 outcrops near the Namibian border. It varies in stratigraphic thickness from 1,500 meters thick to over 6,000 m thick (Ramokate et al., 2000). The formation consists exclusively of purple to red sandstone and mudstone with minor intercalations of limestone (Modie, 1996). Sedimentary structures include planar parallel laminations, planar cross-bedding, reactivation surfaces, oscillatory ripples, and straight-crested symmetrical ripples. These features indicate deposition in the nearshore to shoreline environment (Modie, 1996). Detrital grains are predominantly quartz, K-feldspar, albite, muscovite, epidote, and opaque minerals; heavy minerals define bedding in some well-bedded sandstone. Sandstones are typically cemented by calcite.

The Roibok Complex (Figure 2.2) is an elongate mafic body situated on the northern margin of the Ghanzi-Chobe Belt and the Kwando Complex. The relationship between the

Roibok Complex and rocks of the Ghanzi Group is still poorly understood due to lack of exposure. The Roibok Complex is interpreted to have been emplaced at 717 ± 2 Ma based on U-

Pb zircon age dates (Singletary et al., 2003) and is widely thought to be equivalent to the

Matchless Amphibolite in Namibia and may represent proto-oceanic crust developed during rifting related to the break-up of Rodinia (Singletary et al., 2003).

Syn- to post-orogenic sedimentary rocks of the Okwa Group unconformably overlie the

Ghanzi Group. These rocks were deposited in sub-basins in the southern foreland of the Damara

Orogen in response to contractional deformation, uplift, and erosion (Figure 2.1). In Botswana, thick accumulations of Okwa Group sedimentary rocks were deposited in the Passarge Basin to the southeast of the Ghanzi-Chobe Belt (Figure 2.2; Ramokate et al., 2000). In Namibia, correlative rocks of the Nama Group were deposited in the Nosop Basin. The basal Takatswaane

Formation (Kacgae Subgroup) of the Okwa Group is disconformably overlain by non-deformed but tilted rocks of the Tswaane Formation (Kacgae Subgroup) and Boitsevango Subgroup

(Figure 2.1). The maximum depositional age of these rocks is considered to be 579 ± 12 Ma based on the youngest detrital zircon U-Pb age dates obtained from the Takatswaane Formation

(Ramokate et al., 2000).

Rocks of the Ghanzi Group underwent fold-and-thrust style deformation in the southern foreland of the Damara Orogen during Pan-African assembly of Gondwana that led to the suturing of the Congo and Kalahari cratons (Modie, 1996). Deformation in the southern foreland is broadly bracketed between 580 and 500 Ma based on Ar-Ar recrystallization ages from white micas, hornblende, and whole rock (Gray et al., 2006). Peak metamorphism and deformation occurred at roughly 530 Ma based on K-Ar ages from detrital white micas within the Nama

Group in Namibia (equivalent to the Okwa Group; Horstmann et al., 1990). Post-peak deformation continued along major shear zones in the Kaoko and Damara Belts in Namibia through 460 Ma based on mica blocking temperatures and discordant Ar-Ar age spectra (Gray et al., 2006). Late deformation in the Ghanzi-Chobe Belt may have been related to sinistral movement on the Mwembeshi Shear Zone, the suture between the Kalahari and Congo cratons.

CHAPTER 3

REGIONAL STRUCTURAL GEOLOGY OF THE GHANZI RIDGE

3.1 Regional Geophysical Data

Discovery Metals Limited sponsored a regional airborne magnetic and radiometric survey conducted by New Resolution Geophysics to help delineate major lithological units and structures within the Ghanzi Ridge area. A combination of reduction to the pole (RTP) and tilt derivative (TDR) filters were utilized so that anomalies could be traced for long distances along strike. A downward continuation filter was used to sharpen the anomalies (Figure 3.1). The results of the aeromagnetic survey, together with available geologic information, were utilized to construct a regional geologic map of the Ghanzi Ridge area (Figure 3.2). The geophysical data was combined with regional aeromagnetic geophysical data from northern Botswana (Reeves,

1978).

Both the Kgwebe volcanic complex and the D’Kar Formation contain strong magnetic anomalies (29150 – 29250 nT) due to the presence of disseminated magnetite in the rocks. The two formations are distinguished from each other by correlation with known exposures and exploration drill holes. The magnetic response from the Ngwako Pan Formation is subdued and more uniform (28950 – 29050 nT) compared to the response from the Kgwebe volcanic complex. Up to three distinct laterally persistent magnetic anomalies occur within the D’Kar

Formation presumably due to the presence of detrital magnetite in specific beds. Isolated anomalies with similar magnetic responses to the Kgwebe volcanic complex occur to the northwest of the Ghanzi Ridge. These anomalies are correlated with volcanic rocks of the Karoo

Figure 3.1: Aeromagnetic map of the Ghanzi Ridge area. Reduction to the pole, tilt derivative, and downward continuation filters applied. Star indicates location of the Boseto Copper deposits. Aeromagnetic data courtesy of Discovery Metals (Botswana) Ltd. 17

Supergroup, which occupy fault-bounded grabens interpreted from geophysical data, limited surface exposure, and drill core intercepts (Karoo Graben #2). The magnetic survey highlights several WNW-trending dikes of interpreted Karoo age that crosscut the Ghanzi-Ridge. Dike spacing ranges between 20 and 50 km in the southwest and 10 to 15 km in the northeast. The edge of a wide dike swarm with 1-2 km dike spacing is present in the northeastern most portion of the survey area.

3.2 Regional Structural Geology

The mineralized horizon at the base of the D’Kar Formation forms subcrop below

Kalahari sands in the Ghanzi Ridge area in a series of northeast trending close to tight folds

(Figure 3.2). Major anticline and syncline axial surface traces can be traced over distances of 10 to 50 km with traces spaced 2 to 8 km apart. A regional cross section of the Ghanzi Ridge area suggests that fold amplitudes are approximately 4-6 kilometers (Figure 3.3). Fold limbs range in dip from 45˚ to vertical, and fold axial planes strike 220˚ and 235˚ (right-hand-rule format) and dip between 80˚ to the northwest and vertical. Fold asymmetry defines southeast vergence.

Many folds in the Ghanzi Ridge area have a cuspate shape with an interlimb angle between 50˚ and 20˚, although some folds, including the Plutus anticline, have a box geometry with limb dips abruptly changing from 60˚ to 45˚-30˚ closer to fold crests. With the exception of the region to the northeast of (and including) the Boseto copper deposits, anticlines and synclines plunge at shallow angles to both the northeast and southwest between 0˚ and 15˚ (Schwartz et al.,

1995) creating doubly-plunging folds. To the northeast, plunging folds are not recognized in aeromagnetic data. Overturned limbs and/or nappe complexes have not been recognized in this part of the Damara foreland unlike central Namibia (Ahrendt et al., 1978). However, bedding

18 may be locally overturned on steeply dipping limbs and parasitic folds.

Although no known large-scale reverse faults have been documented in the Ghanzi-Ridge area as is the case to the southwest in Namibia (Kasch, 1983; Miller, 1983, Schwartz et al.,

1995), interpretation of the geophysical dataset suggests that the Ghanzi-Chobe region is dissected by several laterally extensive southwest-striking reverse faults with a component of sinistral displacement. Schwartz and Akanyang (1994b) documented a 15-km-long northeast- striking dextral strike-slip fault on the northwest flank of the Ngunaekau Hills and interpreted thick north- to north-northeast-striking quartz veins as pinnate tension fractures related to movement along the fault plane. They suggested the Ngunaekau Hills fault was a regional feature related to the dextral displacement of the Kaapvaal craton with respect to the Congo craton.

This study indicated that several faults with sinistral strike separation and thrust dip separation of folds cut the D’Kar Formation in the southwestern portion of the study area (Figure

3.2). This suggests that the thick north-northeast-trending quartz veins described by Schwartz and Akanyang (1994b) could be Reidel fractures related to sinistral displacement. As these faults displace folded structures, they must have been developed during a late stage of Damara orogenesis. In addition to the southwest-striking fault system, several north-northeast-striking faults are visible in the aeromagnetic dataset. The timing relationship between the two fault systems is poorly understood although the north-northeast-striking faults show dextral separation of the southwest-striking faults, suggesting the southwest-striking faults are earlier.

Satellite imagery and regional aeromagnetic data reveal closely-spaced second-order parasitic folds within the D’Kar Formation throughout the belt. A syncline to the west of the

Plutus-Petra deposit contains several parasitic folds that formed in response to buckling of the

D’Kar Formation in the hinge of the syncline and resulted in repetition of magnetostratigraphic units within the D’Kar Formation. Broadly spaced, open parasitic folds in the Ngwako Pan

Formation may exist on the limbs of some major folds.

The southwestern nose of the Plutus anticline is cut by a poorly defined 5-km-wide west- northwest-trending graben structure filled with Karoo Supergroup strata. A series of larger grabens occur roughly 40 km west-southwest of the Kgwebe Hills and 10 km north of the

Ngunaekau Hills. These grabens are bounded by north-northwest and southwest striking normal faults and filled by basalts of the Stormberg Member of the Karoo Supergroup. A similar graben occurs roughly 20 km north of the Kgwebe Hills near the town of Toteng. The traces of the southwest-striking faults that bound the Karoo grabens are coincident with the regional southwest-striking faults dissecting the Ghanzi Ridge area.

The region to the north and west of the Boseto Cu project was down-dropped along the southwest-striking Thamalakane and Kunyere faults that form the southeastern margin of the nascent Okavango rift. The southwestern margin of the Okavango rift (including associated normal faults) terminates along the west-striking dextral Sekaka shear zone (Modisi et al., 2000).

Figure 3.2: Geologic map of the Ghanzi Ridge area interpreted from aeromagnetic data.

Figure 3.3: Schematic regional cross section of the Ghanzi Ridge area. No vertical exaggeration.

CHAPTER 4

STRATIGRAPHY OF THE BOSETO AREA

4.1 Introduction

The rocks of the Ghanzi-Chobe belt underwent regional lower greenschist grade metamorphism during the Damaran orogenic event. In the Boseto area (Figure 4.1), primary depositional textures are widely preserved at the Plutus deposit. Deformation resulted in locally intense recrystallization of the stratigraphically equivalent host rock package at the Zeta deposit.

Forty-nine near surface, intermediate depth, and deep inclined diamond drill cores comprising 26 holes from the Plutus deposit, 19 holes from the Zeta deposit, and 4 holes from the Nexus prospect were logged for a total of 7,087 meters of core. Lithologies, sedimentary structures, where present, and sulfide minerals were logged in order to construct stratigraphic columns for each drill hole (Figure 4.2). Several fences of drill holes were selected in order to correlate stratigraphy both down dip and along strike. Due to alluvial cover and sparse outcrop, no geological mapping was conducted for the study. Open pit observations were made, however, when exposures became available. Previously published geological maps and proprietary aeromagnetic data (see Section 3.1) were utilized to interpret bedrock geology and construct regional and local geologic maps.

Lateral and vertical variations in sedimentary facies and stratigraphy were studied to determine any possible controls on the location of ore zones. A detailed sedimentological analysis was performed along ~10 km of strike length in the area of the Plutus deposits as well as

~2.5 km of strike in the Zeta deposit area. Six drill holes from Plutus were spaced between 0.5

23 and 3.0 km apart while drill holes at Zeta were spaced 1.0 and 1.5 km apart. Information on sedimentary structures, grain-size trends, lithological contacts, and bed thicknesses data were collected and utilized to determine sedimentary facies and constrain the depositional environments at Plutus and Zeta. Northeast-southwest long-sections were constructed from the drill hole logs to examine two-dimensional facies and sedimentary body geometry variations.

The top of the Ngwako Pan Formation was chosen as the long-section datum because it marks a major transgressive surface. A regionally extensive black shale bed, limestone beds, and a tuffaceous siltstone/epi-volcaniclastics bed were also utilized as marker horizons.

4.2 Ngwako Pan Formation

Generally, only the uppermost 10 meters of the Ngwako Pan Formation was available for study as drill holes were ended soon after intersecting this unit (Appendix A). The upper five to ten meters of the Ngwako Pan Formation consists of buff to red, planar-parallel to planar cross- bedded, fine-grained sandstone. At Plutus, the rocks are fine-grained, well-sorted, and grain supported with angular to sub-rounded framework grains composed primarily of fine-grained quartz with lesser plagioclase, lithic fragments, and potassium feldspar set in a muscovite-rich matrix that may contain minor biotite and chlorite. The grains are cemented by quartz and/or calcite. Optically continuous authigenic quartz overgrowths on some detrital quartz grains in the sandstones contain fine-grained hematite stained rims, which gives rise to their red color.

Sandstones at Zeta display more metamorphic recrystallization and are generally buff colored, with specular hematite occurring as larger disseminated grains and within veins. In zones of well-developed foliation, coarsely recrystallized muscovite wraps strained and rotated detrital grains, including larger pebbles.

Figure 4.1: Geologic map of the Boseto area.

Figure 4.2: Representative stratigraphic columns compiled from drill core logging in the Boseto area. GDRD1127 represents the Zeta deposit and PSRD1257 represents the Plutus deposit. The columns to the right of the stratigraphic columns represent the occurrence of sulfide minerals within the stratigraphic section.

Pebble-rich beds are common in the upper beds of the Ngwako Pan Formation. Pebbles range in size from 0.75-2 mm in diameter and are moderately to well rounded, indicating short to moderate transport distances. Pebble lithologies include large quartz and feldspar grains and lithic fragments of volcanic rock, granitic material, foliated micro-granite, and phyllitic and schistose metamorphic rocks; pebbles may also include and locally derived sandstone and mudstone rip-up clasts (Figure 4.3). The section also includes thick, laterally discontinuous pebble-conglomerate beds with minor phyllosilicate minerals cemented by calcite and lesser quartz. The pebble beds probably represent channel lag deposits. At Plutus, pebbles are often scattered along bedding planes and are set within finer-grained quartz-rich sand. At Zeta, similar pebbles are slightly flattened with axes stretched parallel to the foliation.

Maroon colored mudstone beds were intersected within the Ngwako Pan Formation at both prospects in deeper geotechnical holes and were observed in pit-wall exposures at Zeta.

The mudstone units vary in thickness between 10’s of cm to 1-2 meters and are laterally discontinuous. Planar laminations, chaotic bedding, and minor coarse-grained pebbles and/or rip up clasts of previously deposited sandstone and mudstone indicate high-energy planar flow with a high sediment load. Within the Zeta pit, this unit is strongly deformed in contrast to the surrounding sandstone, indicating strain was localized within the weaker mudstone unit.

4.3 D’Kar Formation – Lower Member

The lowermost coarsening-upward assemblage at the base of the D’Kar Formation that hosts copper sulfide minerals is informally referred to as the ore zone package (Appendix A).

The ore zone package consists of green to grey mudstone and siltstone with minor marlstone and sandstone. The package varies in thickness across the Boseto area, ranging from over 130

Figure 4.3: Photomicrograph of a pebble conglomerate, Ngwako Pan Formation. Plutus deposit, PSRD310 156.0 m.

meters thick at the Plutus deposit to 20-50 meters thick at the Zeta deposit. At Plutus, the ore zone package displays a variable thickness along strike from 130 meters thick in the southwest to roughly 80 meters thick in the northeast. At Zeta, the ore zone package averages approximately

30 meters thick in the southwest and it is dominated by siltstone and minor mudstone. To the northeast the package thickens to approximately 50 meters, is dominated by mudstone, and grades upwards to mudstone intercalated with minor siltstone, limestone, and sandstone.

The lowermost facies of the D’Kar Formation at both Plutus and Zeta is a laminated calcareous rock that ranges from less than a meter to over 5 meters in thickness. The rock is composed of 45-65% calcite, indicating it is a marlstone. The marlstone is subdivided into a lower red, tan, or yellow sandy laminated marlstone, an intermediate gray argillaceous marlstone to calcareous sandstone/siltstone, and an upper green to grey laminated marlstone. The intermediate marlstone is not always present, in which case the marlstone transitions from red/yellow to green-grey in color over 5-10 cm. At the Plutus deposit, the basal laminated

28 marlstone of the formation consists of planar, 0.1-0.5 mm thick phyllosilicate-rich laminae composed of muscovite, chlorite, quartz, biotite, and minor potassium feldspar that alternate with calcite-rich laminae that enclose apparently detrital plagioclase and quartz grains. These marlstone beds were probably deposited in quiet shallow waters with a local sediment source.

The marlstones at Zeta are highly recrystallized.

The marlstone beds transition into the texturally similar overlying veined mudstone unit.

The veined mudstone unit ranges from five to 20 meters in thickness and hosts the majority of copper sulfides at Boseto. At Plutus, the veined mudstone displays alternating dark and light laminae with variable calcite content. In general, calcite content decreases with increasing silt content. Coarser-grained silt-sized laminae within this sequence contain angular to sub-rounded, apparently detrital grains of quartz, plagioclase, and potassium feldspar with irregular grains of chlorite, muscovite, and biotite. Detrital grains of potassium feldspar (orthoclase, microcline) comprise one to two percent of the whole rock in the mineralized zone. Accessory minerals are typically concentrated in coarser laminae and include mafic minerals (amphibole, pyroxene, garnet), apatite, epidote, tourmaline, rutile-anatase, ilmenite, titanite, and magnetite. These grains are cemented by quartz and/or calcite. Mud-sized laminae are muscovite-rich with lesser amounts of chlorite, biotite, and very fine-grained quartz and potassium feldspar. The veined mudstone unit at Zeta contains less abundant silt-sized material and is usually well-foliated. It consists of grey to green chlorite-rich laminae with flattened detrital grains that alternate with buff colored muscovite-rich laminae. The veined mudstone typically contains abundant fine- grained pyrite that occurs as sub- to euhedral grains. This unit has been referred to as the rhythmite unit in previous studies (Schwartz et al., 1995, Modie, 1996). The laminated texture of the rocks is indicative of suspension sedimentation below storm wave-base.

The veined mudstone unit transitions upwards into a 15-20 meters thick unit characterized by normally graded siltstone-mudstone beds. Individual beds range in thickness from a few centimeters near the base of the unit to tens of centimeters near the top of the unit.

The beds are mud-rich near the base of the section and display increased silt content up section.

The mud-rich tops of beds often display ripples, wispy fluid escape structures, load structures, and small-scale slump features. Beds display sharp and/or erosional bases indicating episodic deposition by density currents during high-energy storm events that transported material below fair-weather wave-base.

At Plutus, the siltstone-mudstone beds unit grades upwards into a thick succession (up to

100 meters in the southwest) of medium- to thick-bedded, normally graded, grey siltstone with occasional interbedded mudstone and lesser fine-grained sandstone. These thicker beds suggest a more proximal sediment source than underlying beds. At Zeta and Nexus, the mudstone- siltstone beds are capped by a 1-2 meter thick, poorly sorted, green to brown, siltstone to fine- grained sandstone with 5-10% randomly distributed coarse grains. The unit is commonly calcareous and locally grades into argillaceous marlstone. The chaotic texture of this unit indicates it was deposited from high-energy events, possibly involving re-working of the underlying sediments.

At Zeta the poorly-sorted siltstone beds are transitional upwards to a 15-20 meter thick tan to brown, massive to thickly bedded sandstone that extends across the deposit area. This sandstone is medium to fine-grained, moderately-sorted, and contains angular to sub-rounded detrital grains of quartz, plagioclase, and potassium feldspar in a muscovite-rich matrix cemented by quartz and/or calcite. Detrital potassium feldspar grains are more common within this sandstone unit than in the underlying siltstones and mudstones.

A 0.2-3.0 meter thick interval of bedded volcaniclastic material is present at Zeta just above or just below the top of the massive to thickly bedded sandstone. This unit is also present at the Nexus prospect slightly higher in the stratigraphic section. However, it does not occur at the Plutus deposit, suggesting the horizon has a limited aerial extent. The origin of the volcaniclastic material is unknown, although microscopic textures indicate they may be tuffaceous siltstones or possibly epi-volcaniclastics. This volcaniclastic unit contains thin beds composed predominantly of lapilli- to ash-size epiclasts of sedimentary clastic material set in a groundmass of very fine-grained potassium feldspar. Individual beds can be clast rich or clast poor.

Throughout the Boseto area, the sandstone unit is usually capped by interbedded limestone or marlstone and siltstone. The limestone beds probably formed during periods of sediment starvation. At Zeta, this limestone is commonly capped by thinly bedded to laminated, commonly calcareous black shale. A lithologically similar black shale occurs higher in the stratigraphic section at Plutus. It overlies 250-300 meters of rocks comprising five stacked coarsening upward sequences similar to the ore zone package. This black shale unit varies from mud- to silt-rich and is locally calcareous. These black shales were likely deposited through suspension sedimentation.

4.4 Stratigraphic Architecture

Sedimentary facies in the lower D’Kar Formation in the Boseto area form a coarsening upward sequence. Core logging of deep exploration drill holes at both Plutus and Zeta indicate several stacked, partial to complete coarsening upwards cycles with similar facies assemblages overlying this lower sequence. Lithologies within these thicker sections (600 m at Plutus, 250

31 meters at Zeta) were grouped into mudstone-, siltstone-, or sandstone-dominated facies assemblages as well as limestone-marlstone and black shale facies. The mudstone-dominated facies assemblage comprises laminated mudstone-siltstone and massive mudstone. The siltstone-dominated facies assemblage comprises mudstone-siltstone beds and thin- to thick- bedded graded siltstone with minor mudstone and/or sandstone. The sandstone-dominated facies assemblage comprises thin to thick-bedded sandstone and massive amalgamated sandstone.

The stratigraphic section at Plutus contains three major coarsening upward cycles (Figure

4.4). Each cycle (1-3, base to top) is comprised of coarsening upward sub-cycles (a, b, and c) that are 30-60 meters thick. The sub-cycles are in turn composed of individual small-scale coarsening-upwards cycles approximately that are 5-20 meters thick. Sandstone-dominated facies become predominant up section within each sub-cycle. Each individual cycle is capped by limestone and/or black shale.

Cycle 1a (ore zone package) at Plutus displays an overall thinning to the northeast with lateral pinch outs of sandstone-dominated facies, thinning of siltstone-dominated facies, and increasing mudstone-dominated facies. Sub-cycles 1b and 1c are both sandstone-dominated and have relatively consistent thicknesses of 50-60 meters along strike; they also appear to thin or pinch-out to the northeast. The sandstone-dominated facies in sub-cycle 1b occur as lens-shaped bodies within siltstone-dominated facies. Sub-cycle 1b is capped by one to two thin limestone horizons while the top of sub-cycle 1c is marked by a return of thick mudstone- and siltstone- dominated facies.

Cycle 2 at Plutus varies between 110 meters thick in the southwest and 160 meters thick in the northeast. Sub-cycle 2a consists of siltstone-dominated facies with minor sandstone-facies beds in the southwest and transitions to mudstone- and lesser siltstone-dominated facies in the

32 northeast. A thick limestone bed caps the unit. In the southwest, sub-cycle 2b is 25 meters thick.

It is comprised of a number of individual sandstone bodies separated by siltstone and minor mudstone beds. In the northeast, sub-cycle 2b increases to 50 meters in thickness and is comprised of sandstone-dominated facies that pinch out to the northeast separated by thick siltstone and mudstone beds. Sub-cycle 2b is capped by a thin limestone in the southwest and a thin calcareous black shale in the southwest. The overlying sub-cycle 2c contains predominantly sandstone-dominated facies to the southwest and a thicker package of mudstone- and siltstone- dominated facies to the northeast.

The base of the uppermost cycle intersected in drill holes at Plutus (cycle 3) contains a thick, regionally extensive black shale bed. This black shale bed thins to the northeast. The lowermost sub-cycle (3a) grades from sandstone-dominated to siltstone-dominated facies to the northeast. Sub-cycle 3b contains a laterally extensive basal black shale bed that grades upwards into mudstone- and then siltstone-dominated facies with minor sandstone-dominated facies. The unit thickens from 20 meters in the southwest to 40 meters in the northeast. A 25 to 30 meter thick sandstone with a strike length of over 3 kilometers occupies the top of sub-cycle 3b in the northeastern portion of the area. The most northeastern drill holes at Plutus contain a number of other apparently discontinuous sandstone beds that may represent channels. The uppermost sub- cycle (3c) displays a thick basal black shale bed that grades upwards to mudstone-dominated facies with lesser siltstone-dominated facies above.

Correlation of stratigraphic units and facies assemblages at Zeta is more problematic due to structural complexity. However, limestone/marlstone and black shale beds, as well as the volcaniclastic horizon serve as marker beds. The lower 300 meters of the D’Kar Formation at the Zeta deposit contains three coarsening upwards cycles that vary in thickness from 50 meters

Figure 4.4: Strike-parallel stratigraphic section of the Plutus deposit.

34 to over 150 meters thick (Figure 4.5). Most sub-cycles comprise incomplete coarsening upward sequences with abrupt facies changes and/or pinch-outs over short distances.

The lowermost sub-cycle (1a, the ore zone package) at Zeta is approximately 50 meters thick in the southwest and increases to 60 meters in the northeast with an accompanying increase in mudstone-dominated facies. In the northeastern portion of the Zeta area, a thin limestone bed occurs at the base of laterally persistent, thick, amalgamated sandstone-dominated facies that caps the sub-cycle. Sub-cycle 1b, comprised of siltstone-dominated facies with minor sandstone beds, pinches out to the southwest. Laterally persistent limestone and black shale beds that comprise the base of the sub-cycle 1c overlie a 0.5-3.0 meter thick horizon of volcaniclastic material. This sub-cycle is capped by thin mudstone-dominated facies that transitions into sandstone-dominated facies. The mudstone-dominated facies appear to pinch out towards the northeast

The second coarsening upward cycle (sub-cycle 2a) contains an incomplete sub-cycle of basal mudstone- and siltstone-dominated facies overlain by a thin discontinuous black shale bed in the central area that grades upwards into a laterally continuous limestone beds. Sub-cycle 2b comprises mudstone- and siltstone-dominated facies with a thin intercalated sandstone bed in the southwest. Sub-cycle 2c is a 15-meter-thick coarsening upward sequence with laterally continuous facies assemblages consisting of basal mudstone-dominated facies overlain by siltstone- and then sandstone-dominated facies. Facies assemblages in sub-cycle 2d vary from siltstone- and sandstone-dominated in the southwest to mudstone-dominated in the northeast; in both areas the sub-cycle is capped by a thin black shale bed. Sub-cycle 2e contains sandstone- dominated facies. It is 40 meters thick in the southwest and pinches out to the northeast.

The uppermost cycle (3) is composed of four sub-cycles, each of which is composed primarily of mudstone- and siltstone-dominated facies, with mudstone-dominated facies prevalent in the southwest and siltstone-dominated facies in the northeast. A thin limestone bed that comprises the base of the third cycle drapes sub-cycle 2e. The sub-cycles in cycle 3contain minor intercalated sandstone beds in the southwest. A thick, laterally discontinuous black shale bed is present at the base of sub-cycle 3d.

Schwartz et al. (1995) and Modie (1996) described the upper member of the D’Kar

Formation as containing mixed oxidized sandstone-dominated and reduced siltstone- and mudstone-dominated facies. These facies were not encountered at the Plutus deposit. However, buff to reddish colored magnetite-bearing sandstones was intersected in deep drill holes at Zeta

300-400 meters above the footwall contact. These sandstone beds may represent the upper member of the D’Kar Formation. This would imply that the lower member of the D’Kar

Formation is roughly 350 meters thick at Zeta and increases to greater than 600 meters thick at

Plutus.

4.5 Paleo-Environmental Reconstruction of the Boseto Area

The Ngwako Pan Formation comprises a thick sequence of arkosic to sub-arkosic sandstone. Modie (1996) described the Ngwako Pan Formation as being deposited in an active rift basin with basal alluvial sedimentation giving way upwards to an axial-trough fluvial system.

Modie (1996) considered the pebble/grit beds in the uppermost Ngwako Pan Formation to represent a period of renewed rifting. Recent reinterpretations suggest much of the Ngwako Pan

Formation was deposited, from base to top, in lower, middle, and upper shoreface environments

(Master, 2010; Caterall, 2012).

Figure 4.5: Strike-parallel stratigraphic section of the Zeta deposit.

Observations from the Zeta open pit indicate that the sandstone beds of the upper

Ngwako Pan Formation are dominantly planar parallel and are laterally continuous. The presence of planar parallel to poorly sorted and massive mudstone beds suggests periodic high- energy events. In addition, the discontinuous pebble-conglomerate beds at Plutus suggest relatively short sediment transport distances within narrow channels. These observations suggest that these are channelized mud-, sheet-, and/or debris-flow deposits and fluvial channel lag deposits that were probably deposited on a distal alluvial fan (Figure 4.6 a). The better-sorted sandstones in the uppermost Ngwako Pan Formation probably formed in an upper shoreface environment that was fed by local alluvial to fluvial material derived from the Kgwebe volcanic complex. The pebble-rich plane-parallel beds in the uppermost Ngwako Pan Formation probably represent wave or tidal re-working of fluvial material entering a body of water, resulting in longshore sand bars and/or strandplain deposits (Figure 4.6 a).

The D’Kar Formation marks a transition to a marine or lacustrine environment and indicates significant sea level rise or tectonic subsidence. The laminated texture of the basal marlstone indicates cyclic suspension deposition of distally transported clays alternating with precipitation of carbonate material during periods of diminished sediment transport. The overlying laminated mudstone represents similar suspension sedimentation with cyclic deposition of siltstone and mudstone. An increase in laminae thickness coupled with a decrease in carbonate cement up section suggest increased sediment input with time. The overlying mudstone- and siltstone-dominated facies and overlying sandstone- dominated facies display increased grain size relative to underlying units and indicates shallowing and/or continued increased sediment input.

The overlying coarsening upward cycles of the D’Kar Formation are characteristic of deltaic deposition within offshore and silty to sandy prodelta sub-environments (Figure 4.6 b).

Mudstone and siltstone beds deposited predominantly by suspension sedimentation probably represent offshore deposits. The mudstone-siltstone beds and the thick normally graded siltstone-dominated beds probably represent offshore to prodelta silts and clays deposited by density currents progressively overlain by prodelta silts and sands. Cycle capping siltstone beds and thin- to thick-bedded and massive amalgamated sandstone beds represent delta-front clinothem sets deposited in delta-front sheets and subaqueous distributary channels.

The lowermost sub-cycle (1a, the ore zone package) with a marlstone at the base overlain by mudstone- and siltstone-dominated facies, probably represent distally deposited sediments in offshore and prodelta zones. The siltstone- and sandstone-dominated facies likely represent prodelta deposits and delta-front clinothem sets deposited through progradation of the delta.

Limestone beds, including those at the top of the first sub-cycle, may have formed during periods of depositional lobe switching. Sub-cycles 1b and 1c consist predominantly of prodelta to delta- front clinothem sets that show progressive southwest to northeast progradation of proximal prodelta to delta-front sediments over distal prodelta deposits interrupted by a period of lobe switching between sub-cycles 1b and 1c.

The thick accumulations of prodelta and/or offshore sedimentary rocks at the base of the second coarsening upward cycle reflect formation of additional accommodation space within the depositional system. The lowermost sub-cycle (2a) consists primarily of prodelta deposits and thin delta-front clinothem sets in the southwest and prodelta to offshore deposits to the northeast that contain thin limestone/marlstone beds that represent periods of lobe-switching and non- deposition. The intermediate sub-cycle (2b) displays southwest to northeast thickening. The

39 southwestern package contains thin sandy prodelta to delta-front clinothem sets that are laterally equivalent to a thicker sequence of prodelta and offshore deposits in the northwest that are capped by prograding delta-front clinothem sets. The dramatic thickening of the sequence to the northeast, especially compared to relatively constant thickness of the underlying sequences, may indicate syn-sedimentary faulting in the area. The uppermost sub-cycle (2c) contains a thin a basal limestone bed in the southwest suggesting a hiatus in terrigenous sedimentation. To the northeast, a black shale bed at the base of the sub-cycle indicates suspension sedimentation in a deeper, more distal position. The black shale bed is overlain by prodelta deposits, which are in turn capped by the laterally continuous sandy prodelta to delta-front deposits that also overlie the limestone to the southwest. These relationships suggest creation of accommodation space to the northeast that was filled with sediment during the same time interval of limestone deposition in the southwest.

Thick accumulations of laterally continuous offshore sedimentary rocks represented by the regionally extensive black shale bed at the base of the third cycle represent renewed deepening of the depositional system. The lower and middle sub-cycles (3a and 3b) contain thin sandy prodelta deposits in the southwest and silty prodelta deposits towards the northeast. This suggests the sediment source stepped back to the southwest relative to the stratigraphic section at

Plutus. The base of sub-sequence (3c) contains a black shale bed indicating further deepening of the depositional system.

An interfingering lobe of sandy prodelta deposits (sub-cycle 4a) is present in the northeast. The sediments were probably sourced from an adjacent depositional lobe transporting sediment towards the southwest or oblique to the section. The northeastern most drill hole in the sequence examined contains several non-correlative sandstone beds within the stratigraphic

40 package. These sandstone beds are probably related to adjacent depositional lobes (Figure 4.6 b).

The laterally continuous geometry of the coarsening upward cycles at Plutus is characteristic of river-dominated delta systems. The along-strike section constructed for the

Plutus sequence appears to be oriented semi-parallel to the original northwest-directed sediment transport direction. Each of the observed cycles may represent progressive progradation episodes, followed by step-backs of the sediment source to the southwest.

Correlation of stratigraphy at the Zeta deposit is more problematic due to structural disruption and lack of continuous marker units. The basal sub-cycle at Zeta (1a, the ore zone package) contains a very similar, but thinner, coarsening upward cycle as Plutus. The thick sandy prodelta to delta-front deposits that cap the cycle are composed primarily of massive amalgamated sands, indicating high deposition rates on the proximal delta-front. Thick-bedded prodelta deposits in sub-cycle 1b abruptly pinch-out towards the southwest and may have been sourced from an adjacent depositional lobe. The sub-cycle is capped by the volcaniclastic horizon as well as limestone and black shale beds. In sub-cycle 1c, sandy prodelta to delta-front deposits with an apparent northeast to southwest transport direction progressively overlie silty prodelta deposits. The base of cycle 1 appears to prograde from southwest to northeast, while the upper part of cycle 1 appears to have had sediment transport from the northeast.

Cycle 2 contains five sub-cycles with relatively constant thicknesses. Sub-cycles 2a and

2b are composed predominantly of laterally continuous offshore and silty prodelta deposits capped by thin limestone beds. The marked difference in grain size from the underlying cycle indicates a deepening of the depositional system. Sub-cycles 2c and 2d represent thin progradational coarsening upward cycles. Sub-cycle 2e contains thick-bedded sandy prodelta to

41 delta-front deposits that pinch out abruptly to the northeast. A thin limestone bed occurring above the prodelta to delta-front deposits appears to cut down from a stratigraphic position above the delta-front to offshore deposits of sub-cycle 2d and may represent a surface of strong sediment reworking. A component of sediment transport direction at Zeta in cycle 2 appears to have been from southwest to northeast.

The uppermost cycle at Zeta consists primarily of offshore and silty prodelta deposits containing thick laterally discontinuous black shale and limestone beds. The substantial change to a finer grain size from the preceding cycle indicates a deepening of the depositional system, similar to that observed in the lithostratigraphically equivalent section at Plutus. The predominantly offshore character of these rocks suggests the main depositional sites may have migrated laterally away from the Zeta area, while the fringes of an adjacent depositional lobe to the northeast migrated laterally over the Zeta.

The overall geometry of cycles at Zeta indicates deposition within closely spaced (2-4 km) distributary lobes. Limestone deposition may have occurred during lobe abandonment.

Thick amalgamated sandy prodelta to delta-front sand deposits suggest possible re-working by major storm events or the presence of offshore sand bars. Coarsening upward cycles at Zeta are generally thinner than those at Plutus and sometimes incomplete, suggesting that the host rocks at Zeta may have been deposited closer to the margin of the delta (Figure 4.6 b).

The extensive black shales within the lower D’Kar Formation sequence throughout the

Boseto area serve as important marker beds and aid significantly in interpretation of the configuration of the basin fill. At Plutus, the most extensive black shale appears 350 meters above contact with the Ngwako Pan Formation. Towards the southeast at Nexus, this basal black shale is located 120-140 meters above the contact with the Ngwako Pan Formation while at Zeta

42 the same unit occurs 55-80 meters above the contact. Farther to the southeast in the Mango area

(Figure 4.1), a similar black shale bed overlies a 20-40 meter thick limestone succession that directly overlies the Ngwako Pan Formation. If this is the same regionally extensive black shale bed, then the basin deepened from the southeast to the northwest. Schwartz et al. (1995) demonstrated that the D’Kar Formation in the Bothatogo area to the west of the Boseto area was comprised entirely of reduced facies marine mudstone and siltstone and suggested this area represented a deeper portion of the basin. The available stratigraphic data suggest the basin deepened to the northwest and west.

Figure 4.6: Depositional environments in the Boseto area. A) Upper member of the Ngwako Pan Formation. Alluvial and fluvial sediment derived from the Kgwebe volcanic complex is transported to a shallow lacustrine or marine environment. Pebbles from channel deposits entering the body of water are re-worked by wave or tidal processes, resulting in strandplain and long shore sand bar deposits with widely distributed pebbles. B) Lower D’Kar Formation. A major marine transgression results in back-stepping of fluvial and deltaic systems. The distal deltas supply sediment to a shallow shelf environment that now characterizes the Boseto area. Subsequent relative sea-level rises and back-stepping of sediment sources indicate syn-sedimentary faulting may have occurred in the area. Faulting probably led to creation of accommodation space that was later filled by stacked coarsening upward progradational delta sequences.

CHAPTER 5

GEOLOGY OF THE BOSETO COPPER DEPOSITS

5.1 Introduction

Large-scale features related to stratiform copper-silver mineralization were investigated through drill core logging and Cu-grade distribution modeling from exploration and grade- control assay data. In contrast to many sedimentary rock-hosted copper deposits, copper sulfide minerals at Boseto occur predominantly within veins and structural fabrics that formed during metamorphism and deformation. Structural features and crosscutting relationships were described through logging of drill core. Oriented-core data were utilized to determine the orientation of regional fold axes and trend and plunge of local parasitic folds as well as to determine vein orientations. These results were utilized to characterize the deformation mechanisms relevant to ore at Boseto.

Samples from the deposit were analyzed utilizing standard transmitted and reflected light petrographic techniques, scanning electron microscope (SEM) analyses, and automated quantitative mineralogical analysis in order to delineate the paragenetic sequence of alteration and mineralization. The QEMSCAN® instrument at the Colorado School of Mines is an automated quantitative mineralogy tool that utilizes a Carl Zeiss EVO50 SEM platform, four

Bruker energy dispersive (EDS) detectors, and proprietary software to produce false-colored mineral maps from backscatter electron signals and EDS (energy dispersive spectrometer) spectra. A PC-based software suite, iDiscover™, allows automated data acquisition and interactive data analysis.

5.2 Structural Setting

Rocks of the Ghanzi Group, which host the Boseto copper deposits, underwent lower greenschist-facies regional metamorphism and contractional deformation in the southern foreland fold and thrust belt of the Damara orogen. The Boseto copper deposits are located on the northwestern limbs of the Plutus and Kgwebe Hills anticlines (Figure 4.1). Most folds are inclined to the southwest (northwest-dipping axial surfaces), with the southeast anticlinal limbs having steeper dips. However, the folds steepen with proximity to the Kgwebe Hills, and the

Zeta syncline is slightly inclined to the northwest.

The northwest limb of the Plutus anticline, which contains the Plutus deposit, strikes 225˚ and dips 50-60˚, while the Petra prospect, located near the nose of the Plutus anticline, has an orientation of 219˚/47˚. The southeast limb of the Plutus anticline, where the Nexus prospect is located, dips 70˚ to the southeast. The northwest limb of the Kgwebe Hills anticline, hosting the

Zeta deposit, dips between 80˚ and 85˚ to the northwest; bedding is locally overturned (dipping

85˚ to the southeast).

Fold axes of both the Plutus anticline and Zeta syncline were modeled by pi-analysis

(Figure 5.1) with data collected from drill core. The average orientation of the Plutus anticline fold axis is 12˚/234˚, indicating the fold plunges to the southwest. The axial plane to the Plutus anticline is oriented 232˚/78˚. Modeling of the axial plane from cleavage measurements resulted in an orientation of 227˚/87˚. The Zeta syncline has similar trend but lower plunge with a fold axis orientation of only 05˚/232˚. The axial plane of the Zeta syncline has an orientation of

231˚/82˚. The general northwest dip of the axial planes indicates southeasterly-directed vergence in the Boseto area. The limb dips and orientation of the axial plane indicate and interlimb angle of approximately 50˚ for the Plutus anticline, indicating it is a close fold. The Zeta syncline has

46 an interlimb angle of 30˚, indicating it is a close to tight fold.

The Plutus and Kgwebe Hills anticlines plunge approximately 15˚ to the southwest, although locally plunge can vary from 10˚ to 25˚. The trend and plunge of fold axes based on bedding-cleavage intersections were measured in several drill holes along strike at each deposit.

Analysis of these data indicates smaller-scale parasitic folds are common features within the

Boseto area, with fold axes plunging to the northwest and southeast between 0˚ and 25˚ (Figure

5.2).

Modeling of the folds based on the Pi-analysis, limb dips, interpretations of aeromagnetic data, and stratigraphic reconstructions suggests the Plutus deposit was buried to a depth of at least 3 and 5 kilometers near the center of the fold limb while the Zeta deposit, closer to the hinge of the Zeta syncline, was buried to a depth of at least 5 to 7 kilometers. This difference in burial depth may account for the differences in deformation intensity observed between the deposits. In addition, syn-sedimentary normal faults near either of the deposits could have acted as a buttresses during folding, resulting in tighter folds with more intense deformation.

A fault interpreted from geophysical data cuts the nose of the Plutus anticline and strikes sub-parallel to the southeast limb of the anticline (Figure 4.1). The mapped trace of the fault ends abruptly near the northeastern limit of the Nexus prospect. The exact nature and timing of the fault is unknown, however, it appears to post-date the main folding as it cuts folded strata. In satellite and aerial imagery, the Kgwebe anticline is cut by several north- to north-northeast- trending lineaments that are inferred to be faults.

Figure 5.1: Fold orientation analysis. Lower hemisphere equal area projections of poles to beds and cleavage including Pi-analysis models of fold axes in the Boseto area. A) Fold axis and axial plane orientation of the Plutus anticline based on bedding measurements. B) Average orientation of the axial plane of the Plutus anticline based on cleavage measurements. C) Fold axis and axial plane orientations of the Zeta syncline based on bedding measurements.

Figure 5.2: Fold orientation analyses. Lower hemisphere equal area projections of bedding and cleavage (poles to planes) and bedding-cleavage intersection (lines). Bedding-cleavage intersections indicate folds plunge at shallow angles to the northeast and southwest. The data indicate that the axial planes of folds in the Boseto area plunge shallowly to both the southwest and northeast.

5.3 District-Scale Features of Stratiform Copper-Silver Mineralized Zones

Zones of stratiform disseminated and vein- and shear-band-hosted Cu-Ag mineralized rock occur throughout the Boseto area at the base of the D’Kar Formation. These zones contain pyrite, chalcopyrite, bornite, and chalcocite as well as lesser galena and sphalerite. Rare marcasite is present within pyrite. Mineralized zones may also include minor disseminated molybdenite. Pyrrhotite is locally present as inclusions in pyrite and chalcopyrite. Tennantite- occurs sporadically with chalcopyrite and as inclusions in large pyrite crystals. Arsenopyrite occurs in trace amounts in most mineralized zones, usually intergrown with galena. Idaite has been noted as a replacement of both chalcocite and chalcopyrite grains (Schwartz et al., 1995).

Though Schwartz et al. (1995) states that silver is concentrated in chalcocite, recent assay data indicate that silver displays a positive correlation with lead, suggesting that galena at Boseto is somewhat argentiferous. Parts per billion concentrations of gold and platinum group elements have been noted in mineralized zones throughout the Boseto area. Anomalous gold concentrations (20 to ≥ 300 ppb) appear to be associated with bornite-rich mineralized zones

(Figure 5.2). However, gold and PGE analyses have only been conducted on a subset of assayed samples. Thus, the existing data do not allow for a robust analysis of gold and PGE zonation patterns or common mineral associations.

The Zeta and Plutus deposits each contain a five- to fifteen-meter-thick high-grade (>1.4%

Cu) mineralized lower zone at the base of the D’Kar Formation overlain by a low-grade (0.6 to

1.4%) to very low grade (<0.6% Cu) zone that ranges from fifteen to thirty meters thick that is in turn capped by a one- to two-meter-thick, upper chalcopyrite-pyrite ore horizon within normally graded siltstone and mudstone beds. Significant, but non-economic Pb-Zn mineralization often occurs stratigraphically above the ore zone package, commonly in limestone beds. Generally

49 similar, but less well-mineralized rocks are present at the intervening Nexus prospect.

Copper sulfide minerals within the lower ore zones at deposits in the Boseto area show a up-section zonation from chalcocite to bornite to chalcopyrite to pyrite-galena-sphalerite to a broad zone of pyrite-only moving upward from the Ngwako Pan-D’Kar Formation contact

(Figure 5.3). Boundaries between the different assemblages of sulfide minerals are generally gradational with overlap of adjacent sulfide assemblages. Sulfides in the deposits occur both as disseminations and within veins and shear zones. In mineralized zones containing both disseminated and vein- or shear-related sulfides each of the different styles contain the same sulfide mineral assemblage and display similar sulfide mineral zonation. There does not appear to be a regular progression of replacement of one sulfide by another in overlapping zones.

Inverse replacement relationships among sulfide species, such as bornite replacing chalcopyrite and vice versa, are common (Schwartz et al., 1995).

Disseminated pyrite occurs within the basal marlstone in the low-grade to barren gap to the northeast of the Plutus deposit, which suggests copper and copper-iron sulfides may have replaced pyrite in the ore zones. Schwartz et al. (1995) observed replacement of pyrite by chalcocite only within zones containing disseminated mineralization.

In both Plutus and Zeta, sulfide assemblages change laterally along strike from northeast to southwest from chalcocite to bornite to chalcopyrite to pyrite ± galena ± sphalerite to pyrite- only (Figure 5.4). Northeast to southwest lateral zonation of copper sulfides is also recognized at other prospects in the Boseto area, including Selene, Zeta northeast, Mango northeast, and

Mango southwest (Figure 5.4). Repetition of the northeast to southwest lateral zonation is also observed at the Ophion prospect 30 km southwest of the Zeta Deposit, which contains narrow chalcocite- and bornite-rich zones that pass laterally into a broad, low-grade chalcopyrite-rich

50 zone and finally into a pyrite-only zone that continues for at least 50 kilometers along strike.

Thus, it appears that Cu-bearing fluids throughout the Boseto area moved laterally along the base of the D’Kar Formation from northeast to southwest suggesting the presence of northwest- southeast-trending normal faults; such faults have not been located to date.

Copper grades are unevenly distributed along strike and down dip at both the Plutus and

Zeta deposits and are a function of both hypogene and supergene processes. Hypogene grade appears to be largely structurally controlled. Grade distribution modeling of the Zeta deposit indicates it contains five high-grade (greater than 2% Cu) zones that form pods raking slightly to the northeast (Figure 5.5). Each pod has a narrow shell of 1.5% copper that grades out to 1% copper. The Plutus deposit displays a similar grade distribution with four high-grade pods occurring along eight kilometers of strike length. Higher-grade zones within the deposits contain both disseminated and vein-hosted sulfide while lower-grade zones contain only disseminated sulfide minerals; the highest-grade zones also contain shear-band-hosted sulfide minerals.

Drilling indicates that supergene processes leached hypogene sulfide minerals in rocks adjacent to high angle faults. Supergene earthy hematite and goethite replace sulfides in these zone and copper oxide minerals such as malachite and chrysocolla are generally present at depth beneath leached zones (Figure 5.6). These copper oxide minerals replaced disseminated, vein-, and shear-band-hosted hypogene sulfide minerals. Additional minerals present in minor amounts within the oxide mineralized zones include azurite, lepidocrocite, tenorite, cuprite, hemimorphite, smithsonite, cerussite, copper, and silver; covellite is often intergrown with oxide minerals (Schwartz et al., 1995). Below the zone containing copper oxide minerals there is commonly a twenty to thirty-meter-thick zone containing both copper oxide minerals and chalcocite. Such zones do not have increased copper values relative to underlying hypogene

Figure 5.3: Typical stratigraphic log depicting up-section ore mineral zonation, Plutus deposit.

Figure 5.4: Lateral mineral zonation map, Boseto Copper Deposits. The zonation suggests that syn-sedimentary northwest-southeast-trending normal faults, possibly transfer faults between a major array of northeast-southwest-trending normal faults, may have been present within the area but are difficult to discern due to later deformation.

Figure 5.5: Long section of copper grade distribution, Zeta Deposit.

Figure 5.6: Long section of sulfide-oxide distribution, Zeta Deposit.

54 mineralized zones indicating that supergene enriched chalcocite blankets were not formed.

5.4 Hydrothermal Alteration Associated with Mineralization

Detrital and diagenetic fabrics are locally preserved within the weakly metamorphosed rocks of the Ghanzi-Chobe belt, particularly in coarser-grained siliciclastic rocks of the Ngwako

Pan Formation. However, most of the finer-grained silty to carbonate-bearing sediments of the

D’Kar Formation, particularly those in the mineralized zones, were recrystallized during the

Damaran metamorphic event. It is difficult to separate possible hydrothermal alteration assemblages related to early to late diagenetic copper mineralization from metamorphic mineral assemblages and late deformation-related hydrothermal alteration assemblages. Mineral assemblages that previous workers (Borg, 1988b; Borg and Maiden, 1989; Schwartz et al., 1995;

Modie, 2000) ascribed to hydrothermal alteration have mineralogies and textures that suggest they are dominantly metamorphic assemblages.

Rocks at the Plutus deposit display depositional textures and retain detrital and diagenetic fabrics, except locally in intensely deformed zones, making them ideal to study the effects of metamorphism and deformation as well as any relationships to copper-silver mineralization.

Common diagenetic features observed at Plutus include authigenic quartz overgrowths on detrital quartz grains, quartz and calcite cements, chlorite that replaced lath-shaped detrital biotite and irregularly shaped mafic minerals in coarser-grained laminae, and authigenic pyrite

(Figure 5.7). The dominant plagioclase mineral in both the footwall and hanging wall sedimentary rocks at Plutus is albite; calcic plagioclase occurs only as relicts in detrital feldspar.

Albite occurs as a replacement of detrital plagioclase and as overgrowths on detrital potassium feldspar grains (Figure 5.7). Albitization of detrital plagioclases is a common diagenetic process

55 in pelitic rocks that have been buried to depths of two to five kilometers and heated to temperatures between 100-150˚C (the “albitization window”; Ramseyer et al., 1992). Such albitization may not be related to mineralization.

The main metamorphic mineral assemblage present at Plutus consists of muscovite, chlorite, quartz, and albite with minor biotite (Figure 5.8). This lower greenschist-facies mineral assemblage is typical of regionally metamorphosed pelitic rocks (Winter, 2009). Metamorphic minerals are distinguished from detrital and diagenetic minerals by their greater grain size, lack of strain texture (quartz), and growth within foliation planes (Figure 5.8). At Plutus, foliation is moderately- to well-developed within mudstone protoliths and poorly developed within siltstone and sandstone protoliths (Figure 5.8). This relationship is manifested both megascopically and microscopically, even within individual laminae. Micas define a foliation that is commonly oriented parallel to bedding in weakly deformed rocks but may be inclined up to 25˚ to 30˚ to bedding in more deformed rocks (Figure 5.8). Muscovite is the most abundant mineral in mudstone protoliths. Original mudstone rocks now contain interlayered muscovite, potassium feldspar, quartz, and chlorite with variable amounts of biotite (Figure 5.8). Where metamorphic foliation cross cuts coarser-grained beds, diagenetic or early metamorphic chlorite on the foliation plane may be replaced by biotite and may have intergrown muscovite (Figure 5.8).

Very fine-grained potassium feldspar is commonly intergrown with biotite.

Hydrothermal alteration at Plutus is recognized by a selvage of bleached wall rock around some quartz-calcite veinlets and more commonly adjacent to layer-parallel shear zones; both bedding- and cleavage-parallel veins may have alteration selvages (Figure 5.9). The bleached alteration envelopes are 0.5-2 millimeter thick. In these alteration selvages, chlorite, biotite, and detrital mafic minerals were replaced by potassium feldspar, quartz, and lesser

56 muscovite (Figure 5.9). The absence of chlorite and biotite in the alteration selvages and together with the presence of chlorite and ankerite in the veins suggests that Mg and Fe may have been leached from the wall rock and moved into the veins (Figure 5.9).

Sulfides are generally absent in the alteration selvages of cleavage-related veins but present as disseminated grains in the wall rock outside the selvages. Many of the quartz-calcite veinlets with alteration selvages contain the same sulfide assemblage that is observed outside of the alteration selvages around veins; the sulfides in the veins are commonly intergrown with chlorite, dolomite, and ankerite as well as quartz and calcite. The relationships suggest that the vein-hosted sulfide minerals may have been derived from the alteration envelope and re- precipitated within veinlets.

At Plutus, alteration spatially associated with layer-parallel shear-bands is more intense than that displayed by veinlets. Within the potassic alteration envelopes, very fine-grained quartz flooding is concentrated nearest to the margin of shear zones (Figures 5.9 and 5.10).

Outward from this, chlorite, muscovite, and detrital plagioclase are largely replaced by very fine- grained potassium feldspar, quartz, and biotite (Figures 5.9 and5.10). Disseminated sulfide minerals are largely absent from the potassic alteration envelope but are commonly present in unaltered wall rock outside the alteration selvage. The intensity of potassic alteration decreases away from the margin of the shear zone with thicker zones having thicker potassic alteration selvages. Alteration intensity is best developed along bedding planes.

Within the cores of shear zones, muscovite, biotite, and potassium feldspar are overgrown and replaced by skeletal calcite (Figures 5.9 and 5.10). Less commonly albite, chlorite, and quartz may also be replaced by calcite in this zone. In areas where thin, laminae- scale shear zones overprint bedding-parallel quartz-albite veinlets, skeletal albite may be

Figure 5.7: Diagenetic features, Plutus deposit. A) A cross-polarized light photomicrograph showing diagenetic quartz overgrowth on a detrital quartz grain, an albitized detrital plagioclase grain, and calcite cement. PSRD292 91.0 m. B) A cross-polarized light photomicrograph showing quartz and calcite cements in siltstone. PSRD292 91.0 m. C) Reflected light photomicrograph of an authigenic pyrite euhedra replaced by chalcopyrite within a siltstone. PSRD310 141.0 m. D) Reflected light photomicrograph of an authigenic magnetite octahedral that is partially replaced by chalcopyrite in siltstone. PSRD310 134.0 m. E) QEMSCAN® false-colored mineral map that depicts an albitized detrital plagioclase grains and lath-like to irregularly shaped chlorite that replaced detrital phyllosilicate and mafic minerals in a marlstone. Micron-size sulfides are commonly intergrown with diagenetic chlorite. PSRD1188 480.0 m. F) QEMSCAN® false-colored mineral map showing feldspar grains only. Image depicts both detrital plagioclase and potassium feldspar. Calcic plagioclase is only present as relicts in detrital grains. Both albite and potassium feldspar form rims to one another. Note abundance of very fine-grained potassium feldspar throughout fine-grained laminae. PSRD1251 454.0 m.

Figure 5.8: Metamorphic minerals and fabrics, Plutus deposit. A) A cross-polarized light photomicrograph showing well-sorted siltstone that is cut by a discordant quartz-calcite veinlet. PSRD1187 452.0 m. B) QEMSCAN® false- colored mineral map depicting the typical quartz-albite-muscovite-chlorite-biotite metamorphic mineral assemblage developed in a weakly foliated siltstone. PSDD310 148.9 m. C) A cross-polarized light photomicrograph of a siltstone with minor recrystallized muscovite laths oriented parallel to stratification. PSDD310 141.0 m. D) QEMSCAN® false-colored mineral map depicting moderately-developed foliation within alternating silt- and mud- sized laminae. Note the interlayered texture of muscovite- and biotite-rich zones in mud-sized laminae. PSRD1251 454.7 m. E) Cross-polarized light photomicrograph of a muscovite-rich foliation developed at 25˚ to stratification. PSRD1251 454.7 m. F) QEMSCAN® false-colored mineral map depicting metamorphic muscovite, biotite, and potassium feldspar that replaced diagenetic chlorite grains along a spaced crenulation. PSRD1251 454.7 m.

Figure 5.9: Hydrothermal alteration related to veins and shear zones, Plutus deposit. A) The photograph shows a 1 mm wide bedding-parallel veinlet that contains a bleached selvage. QEMSCAN® false-colored mineral maps indicate that the alteration selvage is depleted in calcite, chlorite, biotite, and mafic minerals (left) and enriched in quartz, potassium feldspar, and muscovite (right). PSDD310 148.9 m. B) QEMSCAN® false-colored mineral maps showing the distribution of individual minerals and sulfides within a laminae-scale shear zone. Zone 1 corresponds to a pre-existing quartz-albite-sulfide veinlet. Zone 2a corresponds to strong carbonate alteration within the shear zone where skeletal texture calcite pervasively replaced wall rock muscovite, biotite, and potassium feldspar and lesser albite and chlorite. Carbonate alteration is less pervasive outward of the vein within the potassic alteration zone (2b). Zone 2b corresponds to the potassic alteration selvage where potassium feldspar and quartz replaced muscovite, chlorite, and lesser albite. Zone three shows a syn-deformation quartz- calcite-sulfide veinlets that crosscuts some alteration but is in turn displaced along the shear zone. Zone 4 corresponds to a late calcite-only veinlet. PSRD1251 454.7 m. 60

Figure 5.10: Hydrothermal alteration replacement textures. QEMSCAN® false-colored mineral maps from a ~1 cm wide layer-parallel shear zone. Plutus deposit, PSRD1251 454.7 m. A) Image depicts very fine-grained quartz flooding within the potassic alteration selvage immediately adjacent to the shear zone. B) Image depicts a detrital albitized plagioclase grain partially replaced by calcite within the potassic alteration selvage. Mud-sixed laminae were replaced by very fine-grained potassium feldspar, quartz, and biotite. C) Image depicts albitized detrital plagioclase grains replaced by minor potassium feldspar and aluminum silicate minerals within the potassic alteration selvage. D) Image depicts pervasive skeletal texture calcite that replaced muscovite, biotite, and potassium feldspar and lesser quartz, albite, and chlorite. E) Image depicts calcite pseudomorphing skeletal albite within a pre-existing quartz-albite-chlorite-sulfide veinlet. Coarser-grained chlorite contains micron-scale sulfide minerals. F) Image depicts a coarse-grained clot of sulfide and chlorite that appear to have replaced muscovite, biotite, and potassium feldspar and enclosed metamorphic skeletal albite.

61 replaced by skeletal calcite (Figure 5.10). Coarser-grained chlorite and sulfide minerals appear to have replaced potassic wall rock minerals encased by the skeletal albite (Figure 5.10). Within these zones sulfides may also occur as microscopic inclusions within chlorite grains (Figures 5.7 and 5.10). Carbonate alteration locally extended into the potassic alteration envelopes outboard of shear zones where it replaced albitized detrital feldspars (Figure 5.10).

5.5 Macroscopic Structures in the Boseto area

Axial plane cleavage is well developed within sandstones of the Ngwako Pan Formation, but is variably developed and generally restricted to competent sandstone and siltstone beds within the D’Kar Formation. Cleavage orientation is sub-parallel to strike and typically dips between 80˚ to the northwest and vertical. Cleavage spacing is typically 0.5- 1 cm in the

Ngwako Pan Formation sandstones and 1-5 mm within finer-grained rocks of the D’Kar

Formation. It is also manifested as axial plane cleavage to small-scale layer-parallel folds

(Figure 5.11). A similarly oriented spaced crenulation associated with asymmetrically folded laminae is present within weakly deformed marlstones and variably developed in mudstones

(Figure 5.11).

Quartz veins occur throughout the stratigraphic package in the Boseto area. These veins consist of massive, milky quartz and may be several meters thick and laterally continuous for hundreds of meters in the hinges of the Plutus and Kgwebe anticlines (Schwartz et al., 1995).

Centimeter-scale metamorphic quartz veins are more common outward from fold hinges and within finer grained host rocks. Larger veins within the Kgwebe volcanic rocks may contain specular hematite. These veins do not contain sulfide minerals within rocks of the D’Kar

Formation, although they may be overprinted by later quartz-calcite-sulfide veins. The veins

62 may display pseudomorphic textures of pre-existing cleavage (Figure 5.11).

Most quartz-calcite veins at Boseto are bedding-parallel. Such veins are most common in the lowermost 5-15 meters of the D’Kar Formation where the vein density reaches 5-10 veins per meter. Above this zone, vein density decreases to 1-2 veins per meter and vein thickness decreases as well. Bedding-parallel veins range in thickness from 1 mm to greater than 5 cm and average 2 cm. Veins typically have sharp contacts with wall rock and commonly form at bedding contacts (Figure 5.12). Bedding-parallel veins commonly display a syntaxial texture with quartz near the vein margins and calcite in the center. The margins of bedding-parallel veins commonly contain dark green stylolitic chlorite. Veins may contain slivers of chloritized and sericitized wall-rock indicating antitaxial growth and suggesting that multiple ‘crack-seal’ events occurred (Ramsay, 1980; Figure 5.12). A ‘crack-seal’ texture is manifested in many veins by early precipitation of quartz followed by precipitation of intergrown calcite, quartz, and sulfide that split early quartz veins, formed on the margins of early quartz veins, or precipitated in micro-fractures within the quartz crystals.

Discordant veins are less common than bedding-parallel veins. Their orientations vary widely. Discordant vein density at Boseto is difficult to determine since only drill core was available for study. At least some discordant veins and veinlets form stockworks surrounding bedding-parallel veins, suggesting contemporaneous formation (Figure 5.12). Discordant veinlets are generally narrower than bedding-parallel veins, averaging 2-3 mm in thickness, although some veins up to 5 cm in thickness have been observed (Figure 5.12). Thin discordant veinlets often display a single stage of syntaxial growth with mineral fibers growing either perpendicular or at high angles to vein walls (Figure 5.12). These geometries indicate either pure extension or oblique extension, respectively, during vein formation. Thicker discordant

63 veins display blocky vein fill and syntaxial growth textures with a secondary stage of centerline calcite. In zones of well-developed cleavage and/or folded beds, discordant veinlets are commonly oriented parallel to axial plane cleavage. Cleavage-parallel veinlets are generally 0.5-

5 mm in thickness and are characterized by syntaxial fibrous quartz and calcite growth (Figure

5.12). En echelon veinlets are variably developed in the Boseto area and are commonly associated with more deformed rocks (Figure 5.12). Their orientations vary by deposit.

In a broad sense, the contact between the dominantly siltstone-rich D’Kar Formation and the underlying more competent sandstones of the Ngwako Pan Formation forms a shear zone.

Layer-parallel shear zones are important hosts to copper mineralization in the Boseto area and are most common at the base of the D’Kar Formation. These shear zones commonly display scaled morphologies including 1) individual ductile movement horizons along bedding, 2) internally brecciated veins with discordant stockworks, 3) one to two-meter-wide shear-bands containing open to isoclinal folded beds and /or brecciated veins and wall rock, and 4) up to 20- meter-wide shear zones of pervasively foliated and open folded rock containing boudinaged veins and wall rock (Figure 5.13). Textural evidence indicates that at least some quartz-calcite veining was contemporaneous with shear zone development.

Late structural features that overprint strongly deformed rocks are common in the Boseto area. Calcite veinlets with highly variable orientations occur throughout the stratigraphic package. Kink folds occur, usually within sandstones, and show both hanging wall up and footwall up displacement. The kink angle between fold limbs ranges from 30˚ in competent rocks to nearly 90˚ in some mudstones and siltstone. In general, the axial plane of kink folds is parallel to the overall strike of the wall rock and dips 45-60˚ to the northwest or southeast.

Brittle faults occur in the Boseto area, but their orientations are difficult to determine.

Figure 5.11: Axial plane cleavage in the Boseto area. A) Drill core sample displaying bornite mineralized axial plane cleavage related to a small layer-parallel open fold. Core is NQ size, 47.6 mm in diameter. PSRD1187 512.3 m. B) Cross-polarized light scan of a thin section showing a marlstone that displays tight asymmetrical folds associated with a spaced crenulation. Width of view is two cm. GD080-07 205.0 m. C) A close-up cross-polarized light photomicrograph (with the gypsum plate inserted) of image B that shows an individual tight asymmetric fold in a marlstone. Note the alignment of phyllosilicate minerals parallel to the axial plane of the microfold. D) A metamorphic quartz vein a sandstone bed located in the hanging wall to the ore zone. The margins of the vein display a pseudomorphic texture of the pre-existing cleavage. Zeta deposit, GDRD1144 260.0 m.

Figure 5.12: Veins in the Boseto area. All drill core is NQ size, 47.6 mm in diameter. A) A bedding-parallel quartz-calcite-chlorite-chalcopyrite vein that contains chloritized and sericitized wall rock slivers indicating ‘crack-seal’ textures. Note the discordant stockwork vein. Plutus deposit, PSRD1188 460.5 m. B) Two bedding-parallel quartz-calcite-chalcopyrite veins that display ‘crack-seal’ textures. The veins occur at contacts between graded beds. Zeta deposit, GDRD1110 187.0 m. C) A 1.5 mm wide discordant quartz-calcite-chalcopyrite veinlet that displays syntaxial growth fibers and is displaced along a bedding- parallel quartz-calcite-veinlet. Plutus deposit, PSRD1250 499.5 m. D) A 3 cm wide discordant quartz-calcite-bornite vein within laminated mudstone and siltstone that is displaced along a 1.5 mm wide bedding-parallel shear zone mineralized with bornite. Plutus deposit, PSRD1251 454.0 m. E) A cleavage-parallel quartz-calcite-bornite veinlet that occurs within thinly bedded siltstone and mudstone. The host rock displays bornite-mineralized cleavage. Plutus deposit, PSRD1250 506.5 m. F) An en echelon quartz-calcite-bornite veinlet array that displays sigmoidal geometries indicating several growth stages during shearing. Plutus deposit, PSDD1187 509.2 m.

Figure 5.13: Shear zones in the Boseto area. Sample shown are from the Plutus deposit. All drill core is NQ size, , 47.6 mm in diameter. A) A 1.5 mm wide bornite mineralized layer-parallel shear zone with synchronous discordant quartz-calcite-bornite vein. PSRD1254 454.7 m. B) A bedding-parallel quartz-calcite-chalcocite breccia vein that is cemented by chalcocite. Note the discordant stockwork veins. PD022-07 69.3 m. C) A 15 cm wide layer-parallel brittle shear zone that contains brecciated and rotated, foliated, and sericite altered wall rock clasts set within quartz, calcite, chlorite, and bornite. PSRD1187 513.4 m. D) A 15 cm wide brittle-ductile shear zone that displays anastomosing bornite stringers concentrated around pre-existing vein clasts. Note the intense carbonate and sericite alteration (bleached area) of isoclinal folded wall rock with the shear zone. PSRD1187 512.7 m. 67

5.6 The Plutus Deposit

The Plutus deposit contains low-grade disseminated sulfides and high-grade structurally controlled sulfides. Sulfide minerals occur as: (1) disseminated grains, (2) nodules, lenses, and stringers commonly intergrown with gangue minerals, (3) cleavage-parallel lenses rimmed by quartz and/or calcite, (4) grains within veinlets and veins, and (5) grains in shear bands.

Disseminated sulfide minerals most commonly occur as irregularly shaped, inter-granular pore and/or dissolution cavity fillings between 20 and 200 μm wide (Figure 5.14). The shapes of these pore and/or dissolution cavity fillings are similar to those of quartz and/or calcite cements in poorly mineralized rocks. Disseminated sulfide grains may be randomly oriented or weakly oriented parallel to bedding. They may also replace authigenic anhedral to euhedral pyrite up to

1 mm in diameter as well as authigenic magnetite crystals (Figure 5.7 c and d, respectively).

Minor disseminated sulfide grains also occur as tiny inclusions only a few microns in diameter within quartz, calcite, and chlorite mineral assemblages that replace detrital mafic minerals

(Figure 5.7 e).

Sulfide nodules and lenses comprise aggregates of grains or single coarse grains along bedding planes and less frequently within coarse-grained laminae. Nodules are typically lenticular to oval, with their long axes parallel to bedding (Figure 5.15). Individual sulfide grains in nodules and lenses range from 0.1 mm up to a few centimeters in diameter. Sulfide minerals are intergrown with, or enclosed by, quartz and/or calcite (Figure 5.15); the dominant mineral generally reflects the composition of the host rock. Nodules and lenses may also contain chlorite, albite, muscovite, dolomite, ankerite, and rutile (Figure 5.15). Sulfide minerals in nodules and lenses are often encased by or intergrown with very fine-grained skeletal albite

(Figure 5.15).

Sulfide-gangue mineral assemblages similar to those in nodules and lenses occur in cleavage-parallel lenses (Figure 5.15). These lenses are best developed within the lowermost chalcocite and bornite zones hosted by laminated marlstone and calcareous mudstone (Figure

5.11a). The lenses are commonly between one and 10 mm long, cross cut sedimentary laminae, and are associated with a spaced crenulation. Bedding planes with mineralized cleavage-related lenses are commonly interlayered with bedding planes that lack mineralized cleavage domains

(Figure 5.11 a). Some sulfide lenses are elongate parallel to cleavage domains and show folded and flattened and/or thinned wall rock laminae at their margins. These textures suggest they formed prior to deformation and were rotated into cleavage domains.

Two distinct mineralized vein compositions are recognized at Plutus: bedding-parallel quartz-albite and quartz-calcite veinlets and veins. Thin bedding-parallel veinlets are common throughout the Plutus deposit. Such veinlets are typically poorly formed and may be discontinuous over several cm. They are composed primarily of fine- to coarse-grained clots/nodules of recrystallized quartz, chlorite, and sulfide with lesser muscovite, biotite, potassium feldspar, calcite, and accessory minerals encased within very fine-grained layer- parallel skeletal albite (Figure 5.15 d). They are commonly difficult to discern from the wall rock and are distinguished primarily by their slightly darker grey color in thin section (Figure

5.15 d and e).

Quartz-calcite-sulfide veins and veinlets (less than 0.5 cm in width) at Plutus are categorized as: 1) bedding-parallel, 2) discordant, 3) cleavage-parallel, 4) en echelon, and 5) brecciated (Figure 5.12). These veins are composed predominantly of translucent quartz and white to pink calcite. Mineralized quartz-calcite veinlets and veins may contain chlorite, dolomite and/or ankerite, as well as occasional clasts of wall rock along the contact between

69 centerline calcite and marginal quartz (Figure 5.12). Quartz-calcite veinlets and veins have been observed to cut metamorphic foliation, cleavage-parallel lenses, and quartz-albite veinlets. The mineralogy in fibrous discordant quartz-calcite veinlets and veins appears to be controlled by wall rock mineralogy. Veins contain calcite primarily where the veins crosscut coarser-grained calcite-bearing beds. Similarly, sulfide minerals are usually present only where the veinlets and veins crosscut rock containing disseminated sulfide minerals. Many quartz-calcite veins display early quartz precipitation followed by later precipitation of calcite > quartz ± chlorite ± muscovite ± dolomite-ankerite ± sulfide, sometimes along fractures in quartz (Figure 5.16).

Sulfide minerals within quartz-calcite veins occur as small interstitial clots. Sulfide clots may grow up to a few cm in width. Sulfide minerals are typically associated with calcite and may be intergrown with coarse-grained chlorite (Figure 5.16). Dolomite and ankerite also have a strong spatial relationship with sulfides. Sulfides may also form stringers along vein margins or along centerlines.

The orientation of mineralized discordant veins varies significantly based on measurements from oriented core. Four distinct discordant vein sets occur at Plutus based on strike direction (Figure 5.17):

1. cleavage-parallel veinlets with dips between 80˚ to 90˚ with a mean principal orientation

of 230˚/87˚;

2. north-northeast south-southwest-striking vein set with dips between 10˚ and 80˚ with a

mean principal orientation of 007˚/72˚;

3. north-northwest and south-southeast-striking vein set with dips between 30˚ and 50˚ with

a mean principal direction of 158˚/64˚; and

4. west-northwest and east-southeast-striking vein set with dips between 30˚ to 75˚ with a

mean principal orientation of 286˚/75˚.

The reasons for the large variability in strike and dip of discordant veins remains unresolved.

However, most of the discordant veins probably formed as extensional and/or oblique extensional veins from regional compressive stresses during northwest to southeast-directed shortening. Some of the discordant veins display acute bi-sector orientations roughly parallel to

σ1, assuming σ1 is directed perpendicular to the fold belt, while others could have formed as conjugate shear fractures (Figure 5.18). Progressive folding in the area may have rotated these veins to their current orientations (Figure 5.18). Alternatively, the high variability in dip angle for the discordant veins may have resulted from outer-arc extension processes during parasitic fold development as indicated by the variability in fold axes orientations (Figure 5.18).

En echelon veins are developed along narrow brittle-ductile shear planes at Plutus. The shear planes strike parallel to bedding and dip between 0˚ and 20 degrees to the southeast and cuts bedding at high angles (Figure 5.12 e). The en echelon veinlet arrays commonly inclined between 30˚ and 60˚ to shear zone boundaries, resulting in semi-vertical veinlets. The veinlets in en echelon arrays are either straight or sigmoidal, indicating progressive movement along the shear zone during vein growth. The shear sense of shear planes that contain en echelon arrays is commonly hanging-wall-down. En echelon veinlet arrays typically crosscut foliated wall rock and macroscopically folded beds. A distinct set of calcite-only en echelon veinlet arrays developed coeval with weakly developed wall rock boudinage in strongly deformed zones, suggesting they during brittle-ductile shearing.

Brecciated veinlets and veins occur locally at Plutus in zones of stronger deformation.

These veins are always oriented parallel to bedding. Such veinlets and veins may be internally

71 brecciated and contain small clasts of pre-existing vein quartz and calcite. Macroscopically, sulfides are observed to cement brecciated veins and pervasively replace some wall rock clasts.

Locally such breccias and breccia veins may have developed contemporaneously with discordant quartz-calcite-sulfide stockworks (Figure 5.13 b).

Layer-parallel shear zones, developed on various scales at Plutus, are concentrated in finer-grained rocks along zones of rheological contrast (Figure 5.13 a-d). They commonly form thin zones restricted to single bedding-parallel movement horizons (Figure 5.13 a). Laminae in the basal marlstone often display tightly spaced microfolds related to a spaced crenulation. This spaced crenulation has wider spacing within overlying calcareous mudstones (Figure 5.11 b-c).

Asymmetry of microfolds indicates a reverse shear sense. In deformation zones with thick veins, one- to two-meter-wide zones of intervening wall rock may be folded, resulting in small-scale, rootless, open to isoclinal folds. Sulfides minerals are commonly precipitated in the cores of such rootless folds and appear to replace adjacent altered wall rock. Fold asymmetry indicates hanging-wall-up shear sense. Bedding to either side of these zones may display weak to no deformation. These zones appear to be broadly coeval with the formation of mesoscopic-scale parasitic folds.

At Plutus, these veins and breccia zones grade into shear-bands with progressive deformation. Shear-bands form narrow, high-grade (>2% Cu) zones containing anatomizing networks of semi-massive sulfides (Figure 5.13 d). Large fragments of wall rock and veins may be brecciated and rotated within the shear band and/or intensely folded (Figure 5.13 c). Semi- massive sulfides are concentrated around folded and brecciated quartz-calcite veins. Within wall rock fragments, sulfide grains are attenuated parallel to bedding in the limbs of folds. Sulfide minerals also grew along axial planar cleavage developed within small-scale tight to isoclinal

72 folds, and more commonly within the crests of folds (i.e., saddle reefs). Within mudstones and siltstone at Plutus, such shear zones are restricted to narrow intervals, generally 10’s of centimeters and up to two meters in thickness, that are separated by weakly to non-deformed rocks. Between one and three such shear bands occur within the Plutus deposit ore zone. The shear sense from folded wall rock in these zones indicates a large component of reverse movement occurred parallel to bedding planes. This suggests the shear bands formed during fold tightening and parasitic folds growth, subsequent to the main vein-forming phases. Similar shear bands may form within limestone and black shale beds in the hanging wall stratigraphy.

The crosscutting relationships amongst veins and shear zones suggests that early calcite and quartz nodules, lenses, and veinlets formed during early northwest to southeast directed compression. Cleavage-parallel lenticles were subsequently developed at the onset of fold amplification, which was followed by the initiation flexural-slip processes and generation of quartz-calcite veins. Parasitic folds and shear zones were developed during fold tightening.

Thus vein and shear zones at Boseto were formed from the onset of deformation and through final phases of fold tightening (Figure 5.19).

5.7 The Zeta Deposit

The host rock package at the Zeta deposit underwent more intense deformation than at the Plutus deposit. Bedding at Zeta is largely transposed within the ore zone and the uppermost portions of the footwall, particularly the basal marlstone. Diagenetic textures at Zeta were largely destroyed except in less deformed rocks on the northeastern and southwestern ends of the deposit and within hanging wall rocks. Despite these differences, the Plutus and Zeta deposits display broadly similar vertical and later zonation of sulfide minerals. Both contain

Figure 5.14: Disseminated sulfide minerals; Plutus deposit and Nexus prospect. A) A reflected light microphotograph that displays irregularly shaped bornite grains (partially replaced by supergene covellite) that fill and/or replace intergranular pore and/or dissolution cavities. PSRD1254 462.0 m. B) Reflected light microphotograph of chalcopyrite that filled intergranular pore space or a dissolution cavity. PSDD310 141.0 m. C) Reflected light photomicrograph of a bedding plane that contains abundant disseminated chalcopyrite that is crosscut by a discordant calcite-quartz-chalcopyrite veinlet. GD080-07 191.0 m. D) Drill core (NQ size, 47.6 mm diameter) that displays blebby fine- to coarse-grained disseminated bornite grains and bornite rims on nodule features. The mineralized rock is cut by a discordant quartz-calcite-bornite veinlet. PSRD1187 511.5 m.

Figure 5.15: Mineralized nodules, cleavage-parallel lenticles, and stringers, Plutus deposit. A) Bornite mineralized clots and nodules concentrated along a bedding plane adjacent to a bedding-parallel quartz-albite-bornite stringer/veinlet. PSRD1254 459.0 m. B) A cross-polarized light photomicrograph of a mineralized nodule that shows very fine-grained skeletal albite that encloses quartz and chalcopyrite grains. PSDD310 148.9 m. C) QEMSCAN® false colored mineral map depicting mineralized cleavage lenticles. The composition of the lenticles is mainly quartz and bornite with minor albite, chlorite, and wall rock mineral fragments. PSRD1188 477.0 m. D) QEMSCAN® false colored mineral map depicting a bedding-parallel quartz-albite-chlorite-sulfide stringer. Note the irregular to diffuse veinlet margins. PSDD310 148.9 m. E) A cross-polarized light photomicrograph that shows the same veinlet as image D. Note the difficulty in discerning the veinlet due to the presence of very fine-grained albite as the main mineral.

Figure 5.16: Vein sulfide textures, Plutus deposit. A) Cross-polarized light scan of a bedding-parallel vein. The image shows blocky quartz that is cracked by finer-grained quartz, coarse-grained calcite, and chalcocite. The vein contains a chloritized wall rock sliver that indicates a ‘crack-seal’ event. PSDD310 150.1 m. B) A cross-polarized light photomicrograph that shows a chalcopyrite clot situated between quartz grains. The quartz veins are cut by calcite filled microfractures that intersect the chalcopyrite. PSDD310 150.1 m. C) A cross-polarized light photomicrograph that shows clots and microfractures in quartz that are filled with chalcopyrite in a discordant quartz-calcite-chalcopyrite veinlet. PSDD310 148.9 m. D) A cross-polarized light photomicrograph that shows chalcopyrite clots encased by calcite, dolomite, and ankerite. The assemblage filled intergranular space between quartz grains in a discordant quartz vein-calcite- sulfide vein. E) A cross-polarized light photomicrograph that shows a bedding-parallel veinlet that contains thin quartz-(chlorite) margins and a centerline of calcite-bornite-chalcopyrite. PSDD310 148.9 m.

Figure 5.17: Orientation analysis of veins, Plutus deposit. Equal area lower hemisphere stereographic projection that displays the poles to planes of bedding- and cleavage-parallel as well as discordant vein orientations and their corresponding mean principal directions. Bedding-parallel veins are the most common mineralized vein orientation. All vein sets contain some mineralized veins, although the west-northwest-trending vein system contains more abundant mineralized discordant veins.

Figure 5.18: Vein orientations at Plutus. Schematic diagrams illustrating possible mechanisms for the variable orientation of veins at Plutus. Deep burial and initial compression results in bedding-parallel calcite and/or quartz-albite nodules, lenses, and veinlets. Continued compression generated discordant extensional veins as well as discordant conjugate shear veins. As folding progressed, bedding-parallel as well as new discordant extensional quartz-calcite veins are generated through flexural-slip folding processes. Shear zones are also developed during this period. As the limb is steepened, flexural-slip processes cease. Since the limb is nearly perpendicular to the regional compressive stresses, low angle veins are developed, which crosscut deformed rock and pre-existing veins and shear zones.

Figure 5.19: Paragenetic table of structural features observed at Boseto.

79 abundant bedding-parallel quartz-calcite-sulfide veins and both have high-grade ore in shear zones and bands.

Fold amplification and tightening at Zeta resulted in a 10- to 20-meter-wide shear zone along the Ngwako Pan-D’Kar Formation contact (Figure 5.20). Shearing was concentrated within rocks of the D’Kar Formation and to a lesser extent in the upper few meters of the

Ngwako Pan Formation. Shearing along the contact was probably due to the strong rheological contrasts between the two rock types. The basal marlstone of the D’Kar Formation at Zeta is invariably transposed and recrystallized. It contains darker bands that are stretched parallel to foliation and commonly display isoclinal folds that resemble winged mini-boudins (Figure 5.21 a). Intensely deformed zones commonly contain mesoscopic isoclinal folds with fold axes parallel to foliation (transposed bedding, Figure 5.21 b-c). A well-developed, penetrative foliation formed preferentially in finer-grained portions of the stratigraphic section. Foliation orientation is broadly parallel to the overall strike and dip of the host rock package. Coarsely recrystallized (relative to wall rock) well-oriented muscovite-sericite and chlorite grains impart a phyllitic texture to the rocks at Zeta (Figure 5.22). Foliation development was accompanied by mesoscopic-scale open folding within shear zones, resulting in a broadly warped appearance

(Figure 5.20). Fold axes mostly trend parallel to the strike of the host rocks and plunge shallowly to both the northeast and southwest. A gently northeast plunging second-order parasitic fold affecting 10’s of meters of stratigraphic section has been documented within the

Zeta pit. Parasitic folding appears to have been coeval with shearing and open folding within the ore zone.

Metamorphic mineral assemblages at Zeta are similar to those at Plutus. However, growth of phyllosilicate minerals is much more pervasive at Zeta. Metamorphic chlorite is better

80 developed, imparting a distinctive green color to the rocks. In intensely deformed zones, mudstone protoliths contain abundant very fine-grained sericite. Footwall sandstones at Zeta display white mica foliation planes that wrap flattened and elongate framework grains (Figure

5.22).

The Zeta deposit contains less disseminated sulfide minerals than Plutus. The majority of the sulfides at Zeta are structurally controlled and occur as: 1) grains and clots within the necks of boudinaged veins, 2) grains centered along the inter-boudin planes of boudinaged wall rocks, and 3) aggregates of grains stretched into foliation.

Disseminated sulfide grains at Zeta are largely absent from intensely deformed rocks.

Sulfide minerals in intergranular pore space and/or dissolution cavities are commonly recrystallized and stretched parallel to foliation (Figure 5.23). Sulfide minerals are also concentrated along transposed bedding planes and may form discontinuous sulfide stringers.

More commonly, disseminated sulfides occupy the strain shadows to flattened, strained, and sometimes recrystallized framework grains (Figure 5.23). In some instances, quartz and other metamorphic minerals grow in strain shadows of disseminated subhedral to euhedral pyrite that has been partially to completely replaced by copper sulfides and sulfide nodules (Figure 5.23).

The gangue mineral assemblages in nodules and lenses in less deformed rocks at Zeta are comparable to those at Plutus. However, nodules and lenses are commonly flattened and stretched along foliation (Figure 5.23). Weakly deformed mudstone-siltstone rocks exposed in the hanging wall strata display a tightly spaced cleavage on the order of less than 1 mm that is transposed, demonstrating that cleavage was developed prior to the penetrative foliation.

Cleavage-parallel lenticles and quartz-calcite-sulfide veinlets are largely absent from Zeta.

However, foliation fabrics are commonly mineralized.

Generally, bedding-parallel quartz-albite veinlets are present only in weakly deformed rocks at Zeta (Figure 5.12 b). Quartz veins are more widely developed at Zeta than at Plutus.

Blobby, highly irregular, low-angle quartz veins were observed within the footwall sandstones in the Zeta pit. As at Plutus, these veins do not contain sulfide minerals unless they are overprinted by later quartz-calcite vein phases. Some veins may contain specular hematite along fractures.

Quartz-calcite veinlets and veins are widespread throughout the Zeta deposit (Figure

5.24). The most common orientation of quartz-calcite veinlets and veins is parallel to bedding.

Within the high-grade ore zone, bedding-parallel vein densities may be as high as ten per meter.

Such veins commonly contain abundant coarse coarse-grained chlorite. Discordant mineralized veins with orientations similar to those encountered at Plutus rarely occur in the ore zone at Zeta.

However, discordant veinlets and veins containing specular hematite are encountered within the footwall sandstone and in less deformed, oxidized rocks within the hanging-wall. Discordant veinlet and vein density in the footwall sandstone of the Zeta pit varies irregularly along strike from 3-4 veins per meter to less than one veinlet per meter. Generally, such veins are broadly perpendicular to fold axes with dips that vary from near vertical to as low as 25˚. The presence of discordant veinlets below and above the strongly deformed ore zone suggests the lack of such vein orientations in the ore zone may be due to transposition. Mineralized veins that are wrapped by, and terminate within, strongly foliated rocks were observed in the Zeta pit suggesting that these veins may have been transposed during the foliation event. Shear sense indicators suggest such veins may have been rotated during subsequent hanging-wall-up layer- parallel shearing. A distinct set of mineralized discordant veins within the ore zone at Zeta is oriented strike-parallel with shallow dips between 10˚ and 50˚. These veins are spatially associated with boudinaged wall rock (Figure 5.24).

At Zeta, brittle-ductile shearing during late phases of deformation probably resulted from fold tightening and the development of parasitic folds. Brittle-ductile shearing and open folding of the ore zone at Zeta is inferred to have taken place predominantly after the main phases of vein emplacement based on the abundance of boudinage veins. Early quartz-calcite veinlets and veins often display signs of post-formational strain at Zeta, including the presence of twinned calcite and sutured quartz that are cross cut by coarse-grained calcite and/or quartz. Boudinage of veins was common at Zeta. Boudins are sub-divided based on characteristics of the inter- boudin plane (Sib) and include drawn and/or torn boudins, domino boudins, shearband boudins, and sigmoidal-gash boudins (Goscombe et al., 2004; Figure 5.25). Boudinaged veins typically contain stretched lensoidal quartz with the necks between quartz filled with calcite, chlorite, and sulfide minerals (Figure 5.26). Sulfide minerals may also form mantles on boudin blocks.

Boudinage due to movement along bedding planes affected both bedding-parallel and discordant veins (Figure 5.26). Shear sense indicators indicate hanging-wall up displacement.

The boudinage of pre-existing veins and the development of a penetrative foliation at

Zeta was accompanied by locally intense sericitization and chloritization of adjacent wall rocks.

Sericitized wall rock is interlayered with chloritized wall rock, imparting an alternating dark green-grey and buff banding (Figure 5.26 c). This mineralogical banding probably reflects original compositional difference in the layers. Sericite appears to replace muscovite-rich protoliths while chlorite probably reflects replacement of more iron- and magnesium-rich protoliths. Boudins of wall rock foliation are commonly present adjacent to boudinaged veins, where flanking folds and/or shear bands occur on the inter-boudin plane that extends into the wall rock (Figure 5.27). Small quartz ± calcite ± chlorite ± sulfide veinlets precipitated along the inter-boudin planes in wall rock that underwent boudinage. Calcite-only en echelon veinlet

83 arrays typically formed at angles between 45˚ and 60˚ to the inter-boudin plane within boudinaged wall rock outboard of veins.

Cross cutting relationships suggest that veining coincided in part with deformation. Both boudinaged layer-parallel and discordant veins are cross cut by non-deformed discordant veins with identical sulfide and gangue mineralogy (Figure 5.27 c). Mineralized discordant veinlets associated with wall rock boudinage are predominantly oriented parallel to strike and dip shallowly to the southeast between 10˚ and 50˚. These veinlets commonly cross cut ductile deformation fabrics (Figures 5.18, 5.19).

Quartz-calcite veins show internal breccia with sulfide minerals as cement around breccia clasts. Some of these brecciated zones display adjacent discordant stockwork zones, similar to those observed at Plutus. Some brecciated veins contain angular, rotated fragments of cleaved wall rock, demonstrating post-cleavage formation. Shear zones at Zeta commonly display weak folds and/or step-over zones filled with quartz, calcite, and sulfide minerals. Within intensely deformed rock, sulfide minerals may be concentrated within foliation planes (Figure 5.28).

Disseminated sulfide minerals are largely absent from the deformed wall rock. Sulfide-rich fabrics commonly occur adjacent to and/or between brecciated and sheared layer-parallel veins where the intervening wall rock was subject to intense shearing. Sulfide minerals within these fabrics sometimes crosscut the margins of veins. Similar to Plutus, a one to two-meter-wide shear-band is usually present between 20 and 30 meters above the footwall contact at Zeta

(Figure 5.20). Narrow shear zones are developed in fine-grained hanging wall strata, especially within black shale and limestone beds. Hydrothermal alteration related to intense deformation at

Zeta is similar to that observed at Plutus, with shear zones displaying carbonate alteration of adjacent wall rock and weak to strong potassic alteration outboard of the shear zone. Potassic

84 alteration is typically manifested by strong sericitization of mudrock protoliths. The presence of alternating sericite- and chlorite-rich laminae is in stark contrast to Plutus where chlorite is commonly leached from the selvages.

At the Zeta deposit, a steeply dipping, 10-20 meter wide, strike-parallel anatomizing brittle fault system occurs in the hanging wall massive sandstone overlying the ore zone package

(Figure 5.29). The fault is mostly confined to the massive sandstone unit, although it appears to cut into the upper five meters of the underlying ore zone package at depth. The fault breccia consists primarily of clasts of wall rock sandstone, siltstone, and minor mudstone and vein material in an earthy hematitic clay matrix. Petrography shows that the brecciated and rotated clasts consist of well-foliated siltstone and sandstone, suggesting movement along the fault postdates ductile deformation and metamorphism. Exposed slip planes at the base of the fault system in the Zeta pit average 237˚/84˚ in orientation with slickenlines raking 38˚ to 40˚ to the north.

Reidel steps within the fault documented within the Zeta pit indicate reverse shear sense with minor sinistral strike slip.

Late-stage brittle kink folds overprint mineralized ductile deformation structures at the

Zeta deposit. Kink folds occur most commonly within sandy beds. Kink folds show both hanging wall up and footwall up displacement. The angle between fold limbs ranges from 30˚ in competent rocks to nearly 90˚ in some mudstones and siltstone. In general, the axial plane of kink folds is parallel to the overall strike of the wall rock and dips 45-60˚ to the northwest or southeast. The acute bisector between the axial planes of kink folds indicates σ1 was vertical at the time of kinking, suggesting these features formed in response to tectonic loading and partial collapse of the over-steepened limb.

Figure 5.20: Foliation and open folds within the ore zone at the Zeta open pit. Upper image is a photo-mosaic of the ore zone including adjacent hanging wall and footwall rocks. Lower photographs (a-d) correspond to locations (a-d) or examples similar to that location in the upper photo-mosaic. A) One meter wide zone of foliated rock between relatively undeformed siltstone beds in hanging wall low-grade ore zone. B) Open folds in the low-grade ore zone. C) Foliation wrapping a boudinaged quartz-calcite-chalcocite vein within the high-grade ore zone. D) Foliated Ngwako Pan Formation sandstone immediately below the D’Kar Formation.

Figure 5.21: Ductile fabrics related to shearing, Zeta deposit. All drill core is NQ size, 47.6 mm in diameter. A) A transposed and recrystallized marlstone that display tight isoclinal folds that resemble winged mini-boudins. GDDD1009 206.6 m. B) Isoclinal folded laminated mudstone. Fold asymmetry indicates reverse shear sense. GDRD1142 328.0 m. C) Isoclinal fold in a calcareous siltstone. GDRD1142 329.0 m.

Figure 5.22: Foliation at Zeta. A) Ngwako Pan Formation sandstone with sericitic foliation (GDDD1009, 680.9 meters). B) Plane polarized light photomicrograph of sericitic foliation wrapping stretched framework grains, indicating transposition of bedding (GDDD1009, 680.9 meters). C) Cross-polarized light photomicrograph of sericitic foliation in Ngwako Pan Formation sandstone cut by a hematite-(calcite) veinlet (GDDD1009, 680.9 meters).

Figure 5.23: Disseminated sulfides and mineralized nodules, Zeta deposit. A) A reflected light photomicrograph that shows disseminated pyrite and magnetite euhedra that are replaced by bornite. GDRD1135 267.0 m. B) A reflected light photomicrograph showing flattened/stretched pore and/dissolution cavities that are filled by bornite within a foliated siltstone. GDRD1135 267.0 m. C) A reflected light photomicrograph that shows bornite (replaced by supergene covellite) that precipitated within pressure shadows of a flattened quartz grain within a strongly foliated siltstone. GDRD1135 267.0 m. D) A cross-polarized light photomicrograph showing quartz filled pressure shadows on a euhedral disseminated pyrite grain that is replaced by bornite. GDRD1135 267.0 m. E) A reflected light photomicrograph of a quartz-albite- muscovite-chlorite-bornite nodule stretched parallel to foliation. GDRD1135 267.0 m. F) A cross-polarized light photomicrograph of a quartz-albite-muscovite-calcite-chalcocite nodule that is stretched parallel to foliation. GD080-07 202.0 m. ±

Figure 5.24: Mineralized vein orientations, Zeta deposit. Equal are stereonet, lower hemisphere projection of poles to planes of bedding-parallel and discordant veins at Zeta. The majority of discordant veins are oriented parallel to strike and crosscut bedding at low angles.

Figure 5.25: Classification of boudins (modified from Goscombe et al., 2004).

Figure 5.26: Examples of boudinage from the Zeta deposit. All drill core is NQ size, 47.6 mm in diameter. A) Torn and/or gash boudins with quartz-calcite-chlorite inter-boudin vein infill. Boudinage affects both the quartz-calcite vein and wall rock. GDDD1008, 553.0 m. B) Quartz vein displaying torn boudins with calcite inter-boudin vein infill. GDDD1008 487.0 m. C) Bedding-parallel quartz-calcite-chlorite veins displaying drawn boudins. GDRD1159 148.0 m. D) A bedding-parallel quartz-calcite-chalcopyrite vein that displays torn boudins with inter- boudin vein infill consists of calcite and chalcopyrite. GDRD1109 195.2 m. E) Torn and drawn boudins within a bedding-parallel quartz-calcite-chalcocite vein. The vein is wrapped by foliation and terminates within foliated wall rock. Zeta pit. F) A quartz-calcite-(ankerite)-pyrite vein displaying shearband boudins. A later extensional vein of the same mineralogy cuts the boudinaged vein. GDDD1008 582.0 m.

Figure 5.27: Wall rock boudinage, Zeta deposit, GDDD1008. All drill core is NQ size, 47.6 mm in diameter. A) Shearband boudin in wall rock. The pyrite mineralized shearband is parallel to the inter-boudin plane. 593.0 m. B) Domino boudins developed in the wall rock with weakly developed inter-boudin planes; the inter-boudin surface terminates a short distance away from boudinaged wall rock. 477.0 m. C) Shearband boudinage of a quartz vein that has produced a foliation boudinage in the adjacent wall rock. 625.0 m.

Figure 5.28: Mineralized foliation fabric, Zeta deposit. All drill core is NQ size, 47.6 mm in diameter. A) Bornite mineralized foliation fabric adjacent to a layer-parallel quartz-calcite-bornite vein. Note the wall rock on the opposite side of the vein is relatively undeformed. GDRD1143 230.5 m. B) Bornite mineralized foliation fabric that is adjacent to a layer-parallel quartz-calcite-bornite vein. GDRD1109 – 202.0 m. 91

Figure 5.29: Brittle fault system in the hanging wall of the Zeta deposit

CHAPTER 6

FLUID INCLUSIONS

6.1 Introduction

Fluid inclusions are tiny cavities within mineral grains filled with fluids present when the minerals grew or with fluids introduced along fractures after mineral growth (Roedder, 1984).

They may contain liquid, vapor, and/or solid minerals. Primary fluid inclusions form on growth planes and represent fluids trapped as the host minerals grew. Secondary inclusions represent fluid trapped in healed fractures at some time after the growth of the mineral. Pseudo-secondary inclusions are inclusions that are found along healed fractures that formed during growth of the crystal. Microthermometry can be utilized to constrain the homogenization temperature, which is not necessarily the temperature of formation, and salinity of the fluid from which a mineral precipitated.

Fluid inclusions can be heated in the laboratory until liquid and vapor phases are homogenized. The temperature at which this occurs is known as the homogenization temperature. However, for this temperature to be meaningful, it must be shown that the inclusion was originally trapped as a homogenous phase, that the inclusion remained a closed system, and that the inclusion has maintained a constant volume (Roedder, 1984). The only way to show that such requirements have been met for any inclusion of interest is that the inclusion must be observed along with other inclusions of different sizes and shapes trapped at the same time within the same assemblage of inclusions (Goldstein and Reynolds, 1994). If all the inclusions in the same assemblage yield consistent homogenization temperatures then the

93 inclusion assemblage probably meets the requirements. Homogenization temperatures provide only a minimum estimate of entrapment temperatures in most cases. If the depth of the formation of an inclusion assemblage is known and if the pressure was lithostatic or hydrostatic, then a pressure correction can be applied to homogenization temperatures to yield the actual entrapment temperature.

The salinity of fluids in inclusions may be estimated from the melting temperature of ice, although some uncertainty remains due to possible variation in cation chemistry; the eutectic temperature can be determined to check whether this is the case. Salinities are commonly reported as ‘equivalent’ concentrations of NaCl, i.e. the salinity of NaCl solution that would yield the same melting temperature as the natural fluid measured (Goldstein and Reynolds, 1994;

Yardley and Graham, 2002).

The crush-leach process involves crushing a sample to release the fluid in all primary, pseudo-secondary, and secondary fluid inclusions present. The resulting fluid is geochemically analyzed to determine its composition. This technique does not allow determination of fluid compositions related to the formation of a specific mineral assemblage as it samples all generations of fluid inclusions. Crush-leach data represent mixing of different fluid compositions, thus it is impossible to make a quantitative interpretation concerning the contribution of each inclusion population to the bulk fluid. However, these data may allow qualitative hypothesis concerning fluid history. Cl-Br systematics in particular can provide important information as these elements generally behave conservatively in sedimentary and metamorphic settings (Yardley and Banks, 1995).

6.2 Microthermometry

This study utilized a fluid inclusion assemblage (FIA) approach (Goldstein and Reynolds,

1994) to place limitations on fluid composition and homogenization temperatures. Six samples of coarse vein-hosted quartz were cut into thick (~100 microns) sections and inspected under a petrographic microscope. Microthermometry was performed on vein quartz using an Olympus

BX51 microscope and a FLUID INC. adapted USGS Gas-Flow Heating/Freezing stage.

Microthermometric analysis was conducted on sample GDRD1110 187.0 m (Figure 6.1) that contained a vein of quartz with intergrown calcite and chalcopyrite and a chlorite-sericite alteration selvage; the vein also contained fragments of the wall rock. Detailed petrography established that cores of vein quartz grains contained pseudo-secondary fluid inclusions. Clear growth rims around the core of the crystals contained two distinct primary fluid inclusion assemblages each with multiple inclusions of similar shapes and sizes (Figure 6.2). Seven assemblages of primary fluid inclusions within seven separate quartz grains were examined.

Within these seven assemblages, sixteen useable primary fluid inclusions were analyzed (Table

6.1). Only primary inclusions that occurred in assemblages of four or more inclusions were measured for this study.

Fluid inclusion measurements were taken from the clear quartz growth rims adjacent to coarse chalcopyrite grains (Table 6.1). Primary fluid inclusions from these rims vary in size from 2.5 to 5.0 μm. Two types of fluid inclusions were identified and measured. One type of assemblage contained equant to negative crystal shaped inclusions while the second type of assemblage contained irregularly shaped inclusions (Figure 6.2). The equant to negative crystal shaped inclusion were typically 5-10 μm in size and contained liquid and vapor; vapor bubbles typically constituted 25-50% of the inclusion. Irregularly shaped fluid inclusions were typically

5-10 μm in size and contained liquid and vapor. Solid carbonate inclusions were observed adjacent to or in close proximity to the irregularly shaped inclusions (Figure 6.2).

Homogenization temperatures for 11 equant to negative crystal shaped inclusions (that occur within five different quartz grains that contained similar fluid inclusion assemblages) ranged from 165-190˚C. Since all 11 inclusions yielded similar homogenization temperatures, the fluid inclusions were determined to meet the requirements of a fluid inclusion assemblage.

Equant to negative crystal shaped inclusions are thought to form above 260˚C due to faster healing rates (Goldstein and Reynolds, 1994). This suggested that the equant shaped inclusions measured in this study probably require a substantial pressure correction. Five irregularly shaped primary fluid inclusions (from two quartz grains that contained similar fluid inclusion assemblages) that occurred within a different growth zone from the equant to negative crystal shaped inclusions yielded homogenization temperatures of 225-235˚C, indicating they meet the requirements for a fluid inclusion assemblage. This temperature is within the range thought to be responsible for the formation of irregularly shaped fluid inclusions (Goldstein and Reynolds,

1994). Thus the measured homogenization temperatures for the irregularly shaped fluid inclusions are probably close to the actual entrapment temperatures.

The presence of both negative crystal shaped and irregular primary fluid inclusions that yield different homogenization temperatures within adjacent quartz growth zones suggests a complex history of quartz precipitation. The different temperatures could reflect highly variable fluid temperatures through time or a more likely scenario of fluctuating pressure. It is widely assumed that lithostatic fluid pressures are the norm in metamorphism (Yardley, 1996). The intimate relationship between the irregularly shaped fluid inclusions and solid carbonate inclusions suggests trapping may have occurred during vein opening and precipitation of quartz,

96 calcite, and chalcopyrite. A pressure drop is expected upon vein opening. This suggests that the equant to negative crystal shaped inclusions, which require a substantial pressure correction based on fluid inclusion shape and homogenization temperatures, formed while the fluid was under lithostatic pressures. As growth of the rim continued, deformation-induced movement along the vein resulted in the opening of voids that caused a pressure drop at which time the irregularly shaped inclusion and carbonate solid inclusions were entrapped in the growth rim and calcite and chalcopyrite were precipitated within the voids.

Final melting temperatures of seven selected inclusions (two equant to negative crystal shaped and five irregularly shaped inclusions) from within the fluid inclusion assemblage yielded

TmICE in the range of -11 to -17˚C. These temperatures correspond to a fluid with 15-20 weight percent NaCl equivalent (Goldstein and Reynolds, 1994), comparable with known basinal metamorphic fluids (Yardley and Graham, 2002).

6.3 Crush Leach Analysis

The bulk composition of vein-forming fluids was also investigated. Crush-leach analysis was performed on quartz, calcite, bornite, chalcopyrite, and pyrite mineral separates. Samples from the Plutus deposit consisted of a quartz-calcite-bornite vein (sample ID PSRD1127 268.39 m) and a brecciated vein containing quartz and calcite clasts cemented by bornite (sample ID

PSRD1187 505.75 m). Samples from the Zeta deposit included a quartz-calcite-chalcopyrite vein (sample ID GDRD1127 256.6 m) and a quartz-pyrite vein that lacked appreciable carbonate

(sample ID GDRD1180 154.4 m).

Fluids from fluid inclusions in vein-hosted quartz, calcite, chalcopyrite, bornite, and pyrite were extracted from ~0.5 gram separates and analyzed at The Fluid Inclusion Solute

Table 6.1: Microthermometry data for primary fluid inclusions analyzed in this study. Primary Fluid Primary Fluid Inclusion Associated minerals Number of primary Th (˚C) TmICE Salinity Inclusion Shape fluid inclusions (wt % NaCl) Assemblage analyzed within the assemblage

G1110-187a equant to negative crystal 2 165-170 -11.5 to -11.0 15-20 G1110-187b irregular Carbonate, sulfide 2 230-235 -11.2 15-20 G1110-187c equant to negative crystal 4 175-180 G1110-187d equant to negative crystal 2 170-175 G1110-187e equant to negative crystal 2 180-185 G1110-187f equant to negative crystal 1 190 G1110-187g irregular Carbonate, sulfide 3 225-235 -17 15-20

Figure 6.1: Bedding-parallel vein used in fluid inclusion analyses. The vein consists of a quartz, calcite, minor chlorite, and chalcopyrite with thin sericitized wall rock slivers. Chalcopyrite is strongly associated with calcite and grey translucent quartz that overgrows light grey, milky quartz with abundant secondary inclusions. Vein is hosted by dark grey meta-siltstone. Zeta deposit, GDRD1110 187.0 meters.

Figure 6.2: Primary fluid inclusions used in microthermometry. Image shows primary fluid inclusions hosted within a quartz growth rim adjacent to chalcopyrite. Primary negative crystal shaped and irregularly shaped inclusions are two-phase liquid-vapor inclusions, carbonate solid inclusions are associated with the latter. Scale bar = 10.0 μm. Zeta deposit, GDRD1110 187.0 meters.

Chemistry Laboratory, USGS, Denver, Colorado. A stream of pure water (>18 Mega ohm resistance) was added to the fluid during crushing of decrepitated fluid inclusions. The mixed fluid was then split into two separate injection valves, one with a cation trap and one with an anion trap. A USGS Dual Ion Chromatograph was used to identify major and minor cations (e.g.

+ 2+ + 2+ 2+ 2+ + + - - - 2- 2- 2- Na , Ca , K , Mg , Sr , Ba , NH4 , Li ) and anions (e.g. Cl , Br ,CH3OH , SO4 , S2O3 ,CO3 ,

2- - - - PO4 F , I , HS ) in the mixed fluid. Calibration curves were generated from standard samples that follow the same process. Data are reported as atomic ratios to chloride because the amount of water used through the process was not determined (Table 6.2).

Sodium (0.564 – 1.207) and calcium (0.044-1.413) were the most abundant cations identified in the leachates. Potassium (0.017 – 0.074) and magnesium (0.002 – 0.220) were present in lesser amounts. SO4 content ranged from 0.006 – 1.570 and CO3 content ranged from zero to 1.566. Trace amounts of Sr (zero to 0.023) were identified in most leachate samples while Ba (zero to 0.002) was identified in measurable amounts in two leachate samples. NH4

- (zero to 0.033) and CH3OH (acetate; zero to 0.009) was present in some leachate samples.

Because of the small distribution coefficient of Br in halite, brine formation by evaporation of seawater decreases the Cl/Br and Na/Br ratio of the residual brine as indicated by the Seawater Evaporation Trajectory (SET: McCaffrey et al., 1987; Figure 6.4). In contrast, an increase in salinity due to dissolution of halite causes an increase in Cl/Br and Na/Br. The molar

Cl/Br ratios of all leachates (211-626) are similar to the seawater value (Cl/Br ~640: McCaffrey et al., 1987; P. Emsbo, pers. comm, 2012). The Cl/Br ratios from the Boseto crush-leach data suggest that the crustal fluids that generated the syn- to post-deformational quartz-calcite-sulfide veins inherited their Cl/Br signature from evaporated seawater. Metamorphic (crustal) fluids evolve from sedimentary pore waters via a complex series of processes during compaction,

100 dewatering, and devolatilization of minerals (Yardley, 1996). The data indicate that the fluids at

Boseto were enriched in both Ca and Na in respect to most sedimentary brines and are most similar in composition to fluids associated with metamorphic base metal deposits such as those in the Coeur D’Alene district in Idaho, USA (P. Emsbo, pers. comm., 2012).

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Table 6.2: Results of crush-leach extraction analyses, with atomic ratio normalized to Cl.

Sample Mineral Na NH4 K Rb Mg Ca Sr Ba F Acet Cl Br NO3 CO3 SO4 PO4 I GDRD1127 Qtz 0.763 0.000 0.020 0.000 0.004 0.347 0.008 0.000 0.000 0.000 1.000 0.004 0.001 0.550 0.013 0.000 0.000 256.6 m

GDRD1127 Cal 0.783 0.021 0.065 0.000 0.020 0.713 0.023 0.000 0.000 0.000 1.000 0.003 0.006 1.566 0.152 0.000 0.000 256.6 m GDRD1127 Cpy 0.721 0.000 0.069 0.001 0.002 0.130 0.001 0.001 0.000 0.000 1.000 0.003 0.001 0.109 0.020 0.000 0.000 256.6 m GDRD1127 Cpy 0.713 0.000 0.074 0.001 0.002 0.161 0.001 0.000 0.000 0.000 1.000 0.003 0.000 0.191 0.014 0.000 0.000 256.6 m GDRD1127 Qtz 0.776 0.000 0.057 0.000 0.003 0.181 0.000 0.000 0.000 0.000 1.000 0.005 0.000 0.540 0.008 0.000 0.000 268.4 GDRD1127 Cal 0.737 0.019 0.072 0.000 0.015 0.553 0.019 0.000 0.001 0.002 1.000 0.004 0.002 0.843 0.044 0.001 0.000 268.4 m GDRD1180 Qtz 0.743 0.000 0.023 0.000 0.004 0.044 0.001 0.000 0.000 0.002 1.000 0.005 0.001 0.000 0.011 0.000 0.000 154.4 m GDRD1180 Py 0.823 0.033 0.018 0.000 0.220 1.413 0.000 0.000 0.000 0.000 1.000 0.002 0.117 0.072 1.570 0.000 0.000 154.5 m GDRD1180 Py 1.207 0.000 0.017 0.000 0.096 0.389 0.000 0.000 0.005 0.005 1.000 0.002 0.059 0.000 0.738 0.000 0.000 154.5 m PSRD1187 Qtz 0.708 0.000 0.025 0.000 0.005 0.098 0.003 0.002 0.000 0.000 1.000 0.004 0.000 0.186 0.006 0.000 0.000 505.75 PSRD1187 Bo 0.642 0.017 0.027 0.000 0.031 0.187 0.014 0.000 0.005 0.009 1.000 0.003 0.003 1.183 0.061 0.000 0.000 505.75 m PSRD1187 Bo 0.564 0.024 0.032 0.000 0.010 0.057 0.005 0.000 0.001 0.005 1.000 0.003 0.002 0.000 0.029 0.000 0.000 505.75m

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Figure 6.3: Na-Cl-Br systematic of crush-leach data from mineralized veins from Boseto. Seawater evaporation trajectory (SET) is indicated by dashed arrow based on the data of McCaffrey et al. (1987). The data indicate that the vein-forming fluids at Boseto inherited their Cl/Br ratios from evaporated seawater.

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Figure 6.4: Na-Cl-Cl-Br systematics of crush-leach data from mineralized veins from Boseto. Seawater evaporation trajectory (SET) is indicated by dashed arrow based on the data of McCaffrey et al. (1987). The data indicate the fluids at Boseto were significantly enriched in Na+ in respect to modern seawater and sedimentary brines as well as most sedimentary rock-hosted base metal deposits, which plot to the left of the SET (P. Emsbo, pers. comm., 2013).

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CHAPTER 7

STABLE ISOTOPIC ANALYSES

7.1 Introduction

The stable isotopic compositions of carbon and oxygen in calcite from host-rock marlstone and vein-hosted calcite as well as the isotopic composition of sulfur from disseminated and vein-hosted sulfides (chalcocite, bornite, chalcopyrite, pyrite, sphalerite, and galena), and barite from the Boseto area were determined. The main purpose of conducting isotopic analyses of carbonate minerals was to determine the isotopic composition of depositional limestone/marlstone across the area and investigate changes from depositional values due to hydrothermal alteration and/or metamorphism. Initial sulfur isotopic analysis of sulfide minerals was directed at determining the source of sulfur for both disseminated and vein-hosted sulfides.

Further sulfur isotopic analysis focused on determining vertical and lateral variation in sulfur isotopic composition of sulfides across the deposits. In total, 49 calcite, 41 pyrite, 27 chalcocite,

21 chalcopyrite, 18 bornite, six galena, five sphalerite, and one barite samples were analyzed.

These analyses were conducted on a GV Instruments IsoPrime gas-source stable isotope ratio mass spectrometer at the Colorado School of Mines Department of Geology and Geological

Engineering’s Stable Isotope Laboratory

Carbon and oxygen stable isotope analyses were performed with traditional dual-inlet techniques. Calcite samples were collected with a dental drill from carbonate-rich laminae in marlstones and coarse crystalline calcite within veins. For each analysis, a sample weighing approximately 100 μg was reacted with 100% phosphoric acid at 90˚C in separate reaction

105 vessel. The resulting carbon dioxide was cryogenically purified and then analyzed with the mass spectrometer. All carbon values are reported as a per mil difference from the VPDB international reference standard via standard reference material and laboratory working standards, while all oxygen values are reported relative to the VSMOW international reference standard (Coplen, 1994; Table B-1; Figure 7.1). Repeated analysis of an internal laboratory standard yielded a precision of 0.04‰ for carbon and 0.08‰ for oxygen.

Sulfide and sulfate samples were collected using a dental drill. Approximately 25-100 μg of an individual sample was combusted in a Eurovector 3000 elemental analyzer, yielding sulfur dioxide that was delivered to the mass spectrometer using continuous-flow techniques with helium as the carrier gas. Values of δ34S are expressed relative to the Vienna Cañon Diablo

Troilite (VCDT) standard, using the NBS-127 standard reference material obtained from the

National Institute of Standards and Technology. Repeat analysis of a lab-working standard (also barium sulfate) yielded a precision of 0.5‰. The mean half range for three duplicates is 0.25‰.

7.2 Results for Carbon and Oxygen Isotopic Analyses

Previous work by Schwartz et al. (1995) on a thick (up to 30 meters) massive, and weakly deformed limestone from the Mango area approximately 20 km to the east-southeast of the Zeta deposit yielded δ13C values ranging from 0.4‰ to 2.3‰ and δ18O values ranging from 18.5‰ to

19.6‰. Analyses of the more deformed marlstone and limestone at Plutus and Zeta by Schwartz et al. (1995) yielded δ13C values ranging from -1.7‰ to -6.4‰ and a broad range in δ18O values

(13.7 to 20.3‰), with two distinct groups between 13-15‰ and 17-20‰. In the present study, laminated marlstone from the Plutus deposit yielded δ13C values that range from 0.1 to -3.6‰ and δ18O values ranging from 15.1 to 16.2‰ (Figure 7.1). More deformed and recrystallized

106 limestone/marlstone at Zeta displays δ13C values range from -1.1 to -1.3‰ and δ18O values of

17.5 and 18.2‰.

The available isotopic data from the Boseto region were compared with data for early- and mid-Neoproterozoic marine carbonates (Shields and Veizer, 2002; Bull et al., 2010, respectively; Figure 7.1). Data from the Mango area (Schwartz et al., 1995) plots well within the range of values reported for the early Neoproterozoic, whereas none of the data from the Plutus and Zeta deposits plot within the range of values reported from early- or mid-Neoproterozoic marine carbonates. This suggests that if these carbonate rocks originally had values similar to those of other early to mid-Neoproterozoic carbonate rocks, recrystallization during metamorphism resulted in significant isotopic exchange with metamorphic fluids leading to lower carbon and oxygen isotopic values.

Vein-hosted minerals from Plutus and Zeta yielded δ13C values ranging from -0.1‰ to

-13.2‰ and δ18O values from 11.6‰ to 22.1‰ (Table B-1; Figure 7.1). The carbon isotopic compositions are similar to those reported for diagenetic calcite (Ohmoto and Goldhaber, 1997;

Table 7.1). The δ18O values show two distinct groupings between 11.6‰ - 15.3‰ (average

14.2‰) and 16.4‰ and 22.1‰ (average 19.0‰). These values indicate that vein carbonates have even more depleted isotopic values relative to marine Neoproterozoic carbonates than the metamorphosed and recrystallized carbonate rocks. The isotopic data for the Boseto carbonates indicates that all host rock carbonates in the Boseto area were recrystallized during metamorphism and reaction with metamorphic fluids.

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Figure 7.1: δ18O versus δ13C plot for carbonates, Boseto Cu deposits. 1 Data from Schwartz et al (1995) - Mango area: presumed unaltered Ghanzi Ridge carbonate values. Square-circle and diamond-triangle symbols represent separate drill holes in corresponding area. 2 Data from this study.

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Table 7.1: Typical carbonate producing processes and accompanying δ13C values. Process Depth T (˚C) Product δ13C values Evaporation subaerial 15-40˚C Normal marine limestone 0‰

Upper few centimeters to ~ 2 meters below -25‰ Decomposition of organic matter 15-40˚C CO (aq) and HCO - the seafloor (early diagenetic regime) 2 3

-5‰ to -15‰ - Upper few centimeters to ~ 2 meters of Mixing of HCO3 and pore fluids 15-40˚C Diagenetic calcite seafloor (early diagenetic regime) Methanogenic bacteria decomposition of > 2 meters below seafloor (diagenetic 60-80˚ Methane -60‰ to -110‰ residual organic matter regime) 2-5 km below surface (thermocatalytic Various hydrocarbon gases, Catagenesis > 80˚C -40‰ (for CH4) regime) CO2, H2S Metamorphism: at high metamorphic grades, kerogen  graphite; isotopic > 5 km > 200˚C graphite -15‰ exchange with carbonate in metasediments Metamorphism: at high metamorphic CO2; -2‰ to -8‰; grades, presence of H2O results in > 5 km > 200˚C decomposition of graphite CH4 -20‰ to -30‰ Modified text from Ohmoto and Goldhaber (1997).

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7.3 Sulfur Isotopic Analyses

Sulfides at Boseto display a wide range of generally depleted values from 3.5‰ to

-43.1‰ (Table B-2; Figure 7.2). Bornite generally contains the lightest δ34S values ranging between -18 to -40‰ (Figure 7.2). δ34S values for chalcocite are only slightly heavier at -11 to -

33‰. The isotopic composition of chalcopyrite varies widely from -5 to -43‰. Pyrite also has a wide range of δ34S values (3 to -37‰), however, the bulk of the pyrite analyzed have values between -5 and -29‰. Sphalerite and galena also display a wide range of values, although most are concentrated between -9 to -23‰. One sample of barite in a vein about 30 meters into the footwall from the Zeta deposit yielded a δ34S value of 5.4‰. The sulfur isotopic compositions obtained in this study are similar to those reported for other deposits within the Kalahari

Copperbelt (Ruxton, 1986; Ruxton and Clemmey, 1986; Table B-2; Figure 7.3).

The δ34S values obtained for vein- and shear-hosted sulfide overlap those obtained for disseminated sulfides. The data support petrographic observations suggesting that the vein sulfides were derived from dissolution of disseminated sulfides in the wall rock during deformation.

To investigate spatial variations in sulfur isotopic compositions, sulfide minerals were sampled at regular intervals (1, 5, 10, 20, and 30-50 meters) above the footwall contact (Table B-

2; Figure7.4). The data indicate that the sulfur isotopic compositions of the sulfides display distinctive trends with an initial shift to more depleted values followed by a trend to heavier values upwards in the stratigraphy. The δ34S values of sulfides occurring at or near the footwall contact range from -10 to -32.9‰ while δ34S values for samples from roughly five meters above the footwall in the high-grade ore zone are isotopically lighter, ranging between -15 and -40‰.

The sulfide minerals sampled within this zone are coincident with increased vein density and

110 shear zones. With increased stratigraphic height above the footwall contact sulfur isotopic values increase with disseminated pyrite approximately 40-60 meters above the footwall having

δ34S values averaging -12.5‰ at Plutus and -14.8‰ at Zeta. At even higher stratigraphic levels

350-570 meters above the footwall, pyrite displays values from -5.5‰ to 3.5‰. The stratigraphic variation and strong depletion of δ34S values within densely veined and sheared rocks suggests that remobilization and/or significant redistribution of sulfur was restricted to localized shear zones during metamorphism.

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Figure 7.2: Frequency plot of δ34S values by sulfide/sulfate species, Boseto Cu deposits.

Figure 7.3: δ34S values reported from Boseto and other deposits in the Kalahari Copperbelt. Data from other deposits include the Klein Aub Mine (Ruxton, 1986) and Witvlei prospect (Ruxton and Clemmey, 1986) in Namibia.

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Figure 7.4: Stratigraphic variation in δ34S values. A) Plutus deposit. B) Zeta deposit. The red boxes show the position of densely veined and sheared rocks.

113

CHAPTER 8

RHENIUM-OSMIUM CHRONOMETRY

8.1 Introduction

Re-Os Chronometry (Chelsey and Ruiz, 1998; Stein et al, 1998a) was carried out on three sulfide samples from Boseto at the AIRIE Program, Colorado State University, Fort Collins,

Colorado. Two samples from the Zeta deposit (bornite - GDRD‐1127 279.9 m, AIRIE run #

LL‐610; chalcopyrite - GDRD‐ 1127 262.0 m, AIRIE run #LL‐611) and one sample of chalcopyrite from the Plutus deposit (PSRD‐1188 460.85 m, AIRIE run #LL‐612) were analyzed using the method outlined by Stein et al. (2001). The Re-Os chronometer is based on the β-decay of parent 187Re (62.6% of total Re) to daughter 187Os. Molybdenite and other sulfide minerals

(pyrite, chalcopyrite, bornite) are a sink for Re (substituting for Mo, Fe, Cu). Molybdenite generally has Re concentrations well into the parts-per-million range while other sulfides contain lesser Re concentrations. Osmium is essentially excluded from the sulfide structure upon formation. Therefore, the lack of initial or common Os coupled with parts-per-million levels of

Re (Re/Os > 106) yields readily measurable radiogenic 187Os in a geologic sample. The sulfide age is obtained by applying the equation:

187Os = 187Re(eλt – 1) where t is the age and λ is the decay constant for 187Re. The data and results for this study were reported by Dr. Holly Stein (H. Stein, pers. comm. 2012).

Sample LL‐610 consisted of bornite with an estimated 3% native silver ± chalcocite (in the drill mixture) from a bedding-parallel quartz-calcite-chlorite-bornite vein. Sample LL‐611

114 consisted of nearly pure chalcopyrite from a bedding-parallel quartz-calcite-chalcopyrite-chlorite vein with minor hematite. Sample LL-612 consisted of nearly pure chalcopyrite from a bedding- parallel quartz-calcite-chalcopyrite vein. The Re and Os values reported are for sulfide with minimal dilution by silicate minerals during the mineral separation process.

All three samples have accompanying analytical blanks (Table 8.1). For runs LL‐611 and

LL‐612 the blank comprised less than a percent of Os and less than 0.25% of Re. For run

LL‐610 the Re and Os blank was slightly higher (1.1% and 3.8%, respectively, of total measured) as the Re and Os concentrations for the bornite‐silver mixture were lower than for the chalcopyrite samples; LL‐610 was somewhat less than optimally spiked, yet the isotopic ratio measurements are precise for this sample.

The 2σ uncertainties in the calculated ages (Table 8.1) primarily reflect the initial assumed 187Os/188Os value of 0.2 or 1.0, corresponding to Os dominated by a mafic or primitive component in the ore‐forming system or representative of present day eroding continental crust.

Most 187Os/188Os for bulk continental crust in the past would fall between these two values (0.2 to 1.0; e.g., Georgiev et al., 2012). Re-Os age dates are reported as the median age obtained from calculating the model age using initial 187Os/188Os values of 0.2 and 1.0.

8.2 Re-Os Chronometry Results

Sample LL-610 (Zeta, bornite) contained sub-ppb concentrations of Re (0.798 ppb), with a radiogenic Os concentration of 0.0111 ppb. The corresponding 187Re/188Os and 187Os/188Os ratios of 1391 and 24.3, respectively, yield a calculated model age of 1029 Ma or 995 Ma

(dependant on the initial assumed 187Os/188Os values of 0.2 or 1.0). Sample LL-611 (Zeta, chalcopyrite) contained significant Re (7.2 ppb) and had a high 187Re/188Os ratio (5792), with

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90% of total Os as radiogenic daughter 187Os. This composition made this an ideal sample in that regardless of what initial 187Os/188Os value is assumed, the calculated model age changes little (H. Stein, pers. comm. 2012). Using an initial 187Os/188Os ratio of 0.2 yielded a calculated model age of 918 Ma while using an initial 187Os/188Os ratio of 1.0 yielded a calculated model age of 910 Ma. Sample LL-612 (Plutus, chalcopyrite) contained significant Re (3.289 ppb), and a high concentration of common 188Os (0.25 ppb). This composition resulted in a less well- constrained age. Assuming an initial 187Os/188Os value of 0.2, the calculated model age is 496

Ma, while using an initial 187Os/188Os value of 1.0 yields a calculated model age of 442 Ma.

These results are counterintuitive to what would have been expected as the rocks at Zeta are more deformed, presumably during the Damara orogeny, than those at Plutus. The average

Re-Os model age dates of 1012 ± 17 Ma and 914 ± 4 Ma for vein-hosted bornite and chalcopyrite, respectively, from the Zeta deposits are broadly contemporaneous with inferred timing of sedimentation and burial diagenesis in the Ghanzi-Chobe Belt. These results were not expected for sulfide minerals hosted by veins that were apparently formed during deformation.

It is possible that the ages could reflect vein formation during diagenesis with later deformation.

The age from the Plutus sample probably reflects mineralization during late stages of the Damara orogeny, the 496 Ma age is preferred as it corresponds to the age of Damara orogeny while the age of 442 Ma postdates any known Damara deformation (Gray et al., 2006). The Re-Os chronometer has been shown to be extraordinarily robust and lacking disturbance, even through granulite facies metamorphism and intense deformation (e.g., Stein et al., 1999; Raith and Stein,

2000; Bingen and Stein, 2001). With molybdenite, pyrite, and chalcopyrite, solid-state recrystallization does not result in the loss of Re or Os. Both elements are preferentially retained in the sulfide, substituting for Mo, Cu, and Fe, relative to the surrounding silicate or aqueous

116 phases (Stein et al., 2001). This suggests that diagenetic-aged disseminated sulfides could have been locally recrystallized by mechanical processes such as dislocation slide during penetrative foliation development and then incorporated within veins during deformation and metamorphism without resetting the Re-Os chronometer. However, the chalcopyrite in the vein from Plutus with a Re-Os age date of 459 ± 37 Ma probably reflects complete dissolution of sulfide minerals and transport to sites of precipitation in an aqueous phase during vein formation. Stein et al.

(2001) indicated that only during complete chemical dissolution of the sulfide crystal will Re and

Os be liberated and the radiometric clock reset.

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Figure 8.1: Samples used in Re-Os chronometry. A) Re-Os sample LL-610 consisting of a bedding-parallel quartz-calcite-chlorite-bornite vein. Zeta deposit, GDRD1127 279.9 meters. B) Re-Os sample LL-611 consisting of a bedding-parallel quartz-calcite-chlorite-chalcopyrite vein with minor hematite. Zeta deposit, GDRD1127 262.0 meters. C) Re-Os sample LL-612 consisting of a bedding-parallel quartz-calcite-chlorite-chalcopyrite vein. Plutus deposit, PSRD1188 460.85 meters. Note the similarity in composition and textures of the veins.

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Table 8.1: Re-Os chronometry data. Model Age, Model Average AIRIE Sulfide Re, 2σ Total 2σ Common 187Re/ 2σ 187Os/ 2σ rho Ma Age, Ma Model Age, Run # ppb Os, Os, ppb 188Os 188Os assumed assumed reflecting 2σ ppb 187Re/188Os 187Re/188Os uncertainties initial 2σ = initial 2σ = 0.2 1.0

LL-610 bn 0.798 0.002 0.0115 0.0001 0.004 1391 6 24.3 0.1 0.610 1029 995 1012 ± 17 Ma

LL-611 cpy 7.19 0.01 0.0757 0.0010 0.009 5792 17 89.4 0.3 0.636 918 910 914 ± 4 Ma

LL-612 cpy 3.289 0.006 0.0419 0.0005 0.025 646 2 5.56 0.02 0.559 496 442 469 ± 27 Ma

Sample-spike equilibration by Carius tube dissolution single 185Re and 190Os spikes; isotopic ratios measured by NTIMS, AIRIE Program, Colorado State University. Re blanks are 2.02 ± 0.04 picograms. Os = 0.104 ± 0.001 picograms with 187Os/188Os = 0.440 ± 0.009. Sample weights were 230-150 milligrams.

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CHAPTER 9

DISCUSSION

9.1 Sedimentary Architecture

Sedimentation of the Ghanzi Group began after eruption of the ~1106 Ma Kgwebe volcanic complex (Kampunzu et al., 2000; Singletary et al., 2003). Rocks of the Ghanzi Group were deposited in a northeast-trending continental rift with axial-trough drainage (Modie, 1996).

The earliest phases of rifting resulted in deposition of oxidized continental red bed sediments of the Ngwako Pan Formation. The sedimentary rocks contain detrital grains of quartz and feldspar, as well as lithic fragments derived from the unconformably underlying Kgwebe volcanic complex. The Ngwako Pan Formation displays large thickness variations and lateral pinch outs near paleo-topographic highs. The uppermost portions of the sequence represent upper shoreface deposits.

Subsequent extension and enlargement of the depositional basin resulted in marine transgression and deposition of reduced marine siliciclastic and minor carbonate rocks of the

D’Kar Formation. Rocks of the D’Kar Formation were deposited as deltaic sediments on a shallow continental shelf. Syn-sedimentary faulting appears to have accompanied D’Kar

Formation deposition as indicated by the presence of several stacked progradational coarsening upward cycles with finer-grained offshore deposits abruptly overlying proximal prodelta deposits. The Boseto area contains an extensive black shale unit within the D’Kar Formation that thickens to the northwest. Nearly 350 meters of deltaic sediments were deposited below this black shale unit at Plutus while only 50 meters of deltaic sediments underlies the black shale at

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Zeta. Farther to the southeast at the Mango prospect, this shale occurs directly above 20-35 meters of massive limestone that was probably deposited in shallow water above a basement high. These relationships suggest an active tectonic environment with major subsidence and/or down-throw of northeast-southwest elongate fault blocks to the northwest in the Boseto area.

The lack of diamictites within the Ghanzi Group indicates sedimentation ceased prior to about 750 Ma, the maximum age constraint for Sturtian-aged glacial diamictites found within correlative overlying rocks in Namibia (Frimmel at al., 1996). In the Boseto area syn- to post- orogenic deposits of the Okwa Group unconformably overly the Ghanzi Group (Ramokate et al.,

2000), their age is poorly constrained.

9.2 Early to Late Diagenetic Stratiform Copper Mineralization

Sulfur stable isotopic data indicate that early diagenetic pyrite in the Ghanzi Group sedimentary rocks formed from bacteriogenic reduction of seawater sulfate. Mineral textures indicate that copper sulfides replaced diagenetic pyrite and authigenic mineral cements. The

1012 ± 17 Ma Re-Os model age for bornite at Zeta may reflect the age of diagenesis in the lowermost D’Kar Formation if this bornite replaced diagenetic pyrite (Figure 9.1). The Re-Os model age date of 914 ± 4 Ma for chalcopyrite at Zeta may reflect a late diagenetic mineralizing event (Figure 9.1). If these dates represent diagenetic ages they suggest deposition of the underlying Ngwako Pan and Kuke formations between 1106 and 1012 Ma.

The stratigraphic section at Plutus indicates creation of accommodation space along a northeast striking zone coincident with high-grade mineralized zones containing chalcocite and bornite. The vertical and lateral zonation of presumably early disseminated copper sulfide minerals at Plutus and in other portions of the Boseto area suggest migration of mineralizing

121 fluids outward from northeast-striking transfer faults during diagenesis of the Ghanzi Group sequence. The transfer faults were apparently oriented sub-perpendicular to basin-bounding faults.

Mineralizing fluids were oxidized. It is unclear what types of hydrothermal alteration may have occurred during early copper mineralization. Diagenetic albitization of the host rocks was probably not related to the mineralizing event, as it appears to be a regional phenomenon in the

Ghanzi-Chobe Belt (Figure 9.1). It is possible that the very fine-grained potassium feldspar observed in weakly deformed rocks could have formed during hydrothermal alteration related to diagenetic copper mineralization as is observed in the Zambian Copperbelt (Selley et al., 2005;

Figure 9.1). Such early potassium feldspar was later largely replaced by metamorphic biotite during subsequent deformation.

9.3 Basin Inversion, Metamorphism, and Structurally Controlled Mineralization

During the Damara orogeny, the Ghanzi-Chobe basin underwent basin inversion with the development of predominantly northeast-southwest-trending southeast-vergent folds (Kasch,

1983; Miller, 1983; Schwartz et al., 1995; Modie, 2000; Figure 9.1) probably nucleated along northeast-trending normal syn-sedimentary faults. Fold and thrust belt style deformation in the southern foreland of the Damara Orogen is bracketed between 580 and 480 Ma (Gray et al.,

2006). Regional metamorphism resulted in lower greenschist-facies metamorphism (Figure 9.1).

The formation of mineralized calcite and quartz-albite nodules, lenses, and bedding- parallel veinlets at Boseto may have been related to an early phase of compressive tectonism

(Figure 9.1). Further compression could have resulted in sulfide mineral recrystallization along cleavage planes (Figure 9.1). A similar model of sulfide recrystallization has been proposed for

122 the copper sulfide-bearing veins at the Nkana South Orebody mine in the Zambian Copperbelt

(Brems et al., 2009). The Re-Os age date of 459 ± 37 Ma obtained from vein-hosted chalcopyrite at the Plutus deposit suggests formation of quartz-calcite-sulfide veinlets and veins in the Boseto area during the late phases of Damara folding (Figure 9.1).

The bedding-parallel orientations of the mineralized veins suggest they formed by a flexural slip folding process (Ramsay, 1974; Tanner, 1989). Multiple periods of slip along bedding planes is inferred by widespread crack-seal vein textures (Ramsay, 1980). Flexural slip dominates deformation until folds ‘lock up’ at an interlimb angle of about 30˚ (Ramsay, 1974).

Bedding planes that contained abundant disseminated sulfide minerals together with previously formed mineralized quartz-albite veinlets appear to have focused flexural slip movements once folding was initiated. Quartz-calcite vein formation appears to have continued during fold tightening (Figure 9.1). Folding was accompanied by the development of parasitic folds that are spatially associated with reverse slip-sense layer-parallel shear zones at the base of the D’Kar

Formation (Figure 9.1). Shear zone development at Plutus was limited to narrow shear-bands.

At Zeta, the entire ore zone package at the base of the D’Kar Formation underwent intense shearing characterized by the development of a penetrative foliation, open folds, and boudinage of previously formed veins. Deformation in the Boseto copper deposits is inferred to have developed at depths of between 5 and 7 km based on simple fold models. The presence of relatively ductile carbonate minerals in the Ghanzi Group sequence may have facilitated folding during shearing. Steepening of the fold limb at Zeta to near-perpendicular to the horizontal σ1

(regionally consistent in the fold belt) resulted in cessation of or decrease in the amount of flexural-slip folding and the formation of low-dip veins.

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Many veins and shear zones at Boseto display a selvage depleted in chlorite and sulfide and carbonate minerals relative to the surrounding wall rock (Figure 9.1). However, the sulfide and carbonate mineral assemblages within the veins mirror those of wall rock outside of the leached vein selvages, suggesting that the vein material was largely sourced from the immediately adjacent wall rock. Both disseminated and structurally controlled sulfide minerals contain isotopically similar sulfur isotopic values suggesting that the vein related sulfides inherited sulfur from the earlier precipitated disseminated sulfides. Re-Os chronometry giving both early and late ages for vein-hosted copper sulfide minerals provides support for a model of diagenetic-aged disseminated copper sulfides that were dissolved and then re-precipitated into layer-parallel veins and shear zones (Figure 9.1). Similarly, wall rock calcite and calcite in veins have similar isotopic values indicating possible dissolution of calcite in alteration selvages and re-precipitation of calcite in veins. Thus, there is textural, isotopic, and geochronological evidence for a diagenetic mineralizing event with subsequent local dissolution of both sulfides and carbonates adjacent to veins and shear zones during deformation and re-precipitation of these minerals within veins and shear zones. This process appears to have been responsible for the formation of high-grade, structurally controlled ore zones.

The fluids responsible for local remobilization of sulfide minerals are poorly understood.

The Cl/Br ratios of fluids trapped within vein-hosted sulfides from the crushed leach experiments indicate the Cl/Br ratios of the fluids were inherited from evolved seawater. This seawater was probably trapped as pore fluid, which then underwent exchange with the wall rocks during burial.

Crustal/metamorphic fluids are widely assumed to be at lithostatic pressures at metamorphic conditions (Yardley 1996). Transient development of enhanced permeability in

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Figure 9.1: Schematic evolution diagram of the Boseto copper deposits.

125 veins and shear zones can lead to highly focused fluid fluxes with attendant metasomatism

(Yardley, 1996). Fluid inclusion evidence from Boseto suggests that structurally controlled ore and gangue mineral precipitation was driven by episodic pressure changes induced during layer- parallel shearing. However, the mechanisms responsible for the dissolution of calcite and sulfide minerals from the wall rock and subsequent re-precipitation within adjacent veins during shearing are not understood.

9.4 Comparison to Other Sedimentary Rock-Hosted Stratiform Copper Deposits

The Boseto copper deposits share many similarities to other sedimentary rock-hosted stratiform copper districts around the world. The contained metals (dominantly Cu-Ag with minor Zn-Pb), deposit- to district-scale metal zonation patterns, and importance a regional redox horizon at Boseto are similar what is observed in the world-class Kupferschiefer deposits in southern Germany and southwestern Poland (Hitzman et al., 2005). Most workers agree that the

Kupferschiefer records multiple stages of Cu mineralization (Vaughn et al., 1989; Wodzicki and

Piestrzynski, 1994; Kucha, 2003; Hitzman 2005). Re-Os chronometry results from this study indicate that copper sulfides were initially precipitated during early to late diagenesis with a later period of sulfide dissolution and re-precipitation within veins and shear zones during metamorphism and deformation. In the Kupferschiefer, late-stage, structurally controlled high- grade copper zones (“Rucken” veins) are poorly developed. However, in the Boseto deposits zones of closely spaced veins are more common and form the highest-grade copper zones.

Most sedimentary rock-hosted stratiform copper districts around the world, including world-class and giant deposits, contained significant evaporitic sequences, commonly as a cap to the mineralized interval (Hitzman et al., 2005, 2010). The sedimentary sequence hosting the

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Boseto copper deposits does not appear to have contained significant evaporites in either the footwall or hanging wall to the mineralized horizon. The Boseto deposits appear to be more similar in stratigraphic setting to the deposits of the White Pine (Michigan, USA; Mauck et al.,

2002) and Redstone districts (Northwest Territories, Canada; Jefferson and Ruelle, 1986).

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APPENDIX A: LITHOLOGY PHOTOGRAPHS

Figure A-1: Photographs of representative samples of the Ngwako Pan Formation , Boseto Copper deposits. A) Grey fine-grained sandstone with transposed pebble-rich horizons (grit layers), sericitic foliation, Zeta. Discordant calcite ± quartz vein. B) Pebbles within weak phyllitic foliation defined by micas and sericite, Zeta. C) Grey to buff fine-grained sandstone with sericitic foliation, Zeta. Discordant quartz-calcite vein. D) Grey to buff-pink fine grained sandstone. Disseminated specularite and quartz-specularite ± carbonate veins, Zeta. E) Grey fine-grained sandstone with pebble-rich horizons, Plutus. F) Pebble conglomerate containing abundant rounded grains, Plutus. G) Grey to red fine-grained sandstone and pebble-rich horizons. Discordant quartz-calcite vein, Plutus. H) Red fine-grained sandstone with dark minerals defining a weakly developed foliation, Plutus.

134

Figure A-2: Photographs of basal marlstone and limestone, D’Kar Formation, Boseto Cu deposits. A) Upper marlstone, grey to green, transposed and recrystallized, small scale isoclinal folding of mud-rich laminae, Zeta. B) Intermediate interbedded marlstone and mudstone, light to dark grey, transposed and recrystallized, isoclinal folds, Zeta. C) Lower marlstone to limestone, pink to red, transposed and recrystallized, Zeta. D) Lower marlstone, light to dark grey, transposed and recrystallized, small scale isoclinal folding of mud-rich laminae, Zeta. E) Upper marlstone, medium to light grey, finely laminated, mineralized with bornite clots parallel to cleavage F) Upper marlstone, light grey to green, finely laminated to weakly recrystallized, Plutus. G) Lower marlstone, yellowish grey, finely laminated and mineralized with chalcocite clots parallel to laminations with alteration selvages, Plutus. H) Lower marlstone, light brown to pink (stained for calcite), finely laminated, Nexus.

135

Figure A-3: Photographs of representative samples of the ore zone member D’Kar Formation, Plutus deposit. A) Uppermost mudstone. Grey to brown mudstone with very thin siltstone laminae. Clots of pyrite-sphalerite parallel to bedding. B) Uppermost siltstone-mudstone. Grey to dark brown, thin-bedded siltstone-mudstone beds, mud dominated. C) Siltstone-mudstone. Light grey to dark grey-brown, medium to thin bedded siltstone mudstone beds. Erosional siltstone bases overlying mudstone with flame structures. Small-scale cross-bedding present. D) Grey-brown massive siltstone, beds up to few meters thick. usually have a thin mudstone cap layers. Cross cut by discordant quartz-calcite-chalcopyrite veinlets. E) Siltstone, similar to above except medium to dark grey. Cut by quartz-calcite-chalcopyrite –bornite veinlets. F) Siltstone-mudstone. Dark grey to brown, thinly bedded siltstone-mudstone beds. Bedding parallel chalcopyrite clots and disseminated chalcopyrite. G) Lower mudstone rhythmite. Medium to dark grey, occasionally green in color. H) Lower mudstone rhythmite similar to above. Chalcocite-bornite-chalcopyrite aggregates parallel to bedding and cleavage. Generally strongly folded along discrete planes.

136

Figure A-4: Photographs of representative samples of the ore zone member of the D’Kar Formation, Zeta Deposit. A) Uppermost mudstone brecciated by the hanging-wall fault in this sample. B) Grey, two centimeter thick normally graded siltstone-mudstone beds with bedding-parallel quartz-calcite-chalcopyrite veins at bedding contacts. C) Medium grey and/or green massive to faintly bedded siltstone with disseminated bornite. D) Green-grey normally graded mudstone- siltstone. Darker grey siltstone laminae with bornite define bedding. E) Isoclinal folds in green-grey and brown mudstone separated by thin dark grey siltstone laminae. F) Isoclinal folds in grey to dark grey mudstone-siltstone rhythmite overlying the basal marlstones, calcareous in places.

137

Figure A-5: Photographs of representative samples of hanging-wall stratigraphy, Zeta Deposit (top to bottom of hole). A) Light pink to green, moderate to poorly sorted sandstone with chloritic foliation overprinted by light colored alteration selvage to quartz vein. B) Green to pink, moderate to poorly sorted sandstone with chloritic foliation. C) Dark grey to black mudstone with minor siltstone laminae. D) Variable colored, poorly sorted, bedded volcaniclastic material with lithic fragments and lapilli size crystal fragments set in an aphanitic to microcrystalline potassium feldspar-rich groundmass in places. The sample contains abundant quartz-calcite veinlets and has a silicified or weakly hornfelsed appearance. E) Light pink to grey-green sandstone with stockwork quartz micro-veinlets. It has a silicified or hornfelsed appearance. F) Brown to tan, moderately sorted, fine to medium grained sandstone located within the hanging- wall fault zone. Sample contains abundant quartz ± calcite veinlets that predate faulting. The fault breccia is cemented by iron oxides and carbonate.

138

APPENDIX B: STABLE ISOTOPE DATA

Table B-1: Results for carbon and oxygen stable isotope analyses. Sample ID Drill Hole Depth Description – mineral sampled δ13C (VPDB) δ18O (SMOW) (m)

Zeta CO-19 GDDD1009 668.8 Laminated marlstone -1.4 17.5 CO-27 GD157 50.3 Qtz-cal-cc vein, discordant – calcite, white (Ngwako Pan Fm) -7.6 18.2 CO-46 GDDD508 131.3 Qtz-cal-chl-cc vein, BP, BX, calcite, pink to buff -6.9 21.4 CO-08 GDDD1009 660.3 Qtz-cal-bn-cc vein, BP, BD -2.4 14.7 CO-11 GDDD1009 680.9 Qtz-cal vein, discordant – calcite, white (Ngwako Pan Fm) -4.1 11.6 CO-28 GDDD1110 141.3 Qtz-cal-chl vein, discordant – calcite, white -8.7 19.1 CO-34 GDDD1110 187.7 Qtz-cal-chl-bn-cpy vein, BP, BD – calcite, buff -2.1 14.7 CO-45 GDRD1110 197.6 Qtz-cal-chl-cc vein, BP, BX – calcite, buff -4.4 16.4 CO-32 GDDD1110 209.0 Cal vein, discordant – calcite, white (Ngwako Pan Fm) -1.1 14.1 CO-03 GDRD1113 268.4 Qtz-cal-chl-bn-(cc) vein, BP – calcite -3.3 14.7 CO-29 GDRD1113 271.4 Qtz-cal-spec vein, discordant – calcite, pink -0.8 14.1 CO-35 GDRD1127 256.6 Qtz-cal-py vein, discordant – calcite, pink -3.3 14.8 CO-02 GDRD1127 262.0 Qtz-cal-chl-cpy-hem vein, BP – calcite, pink -2.4 14.3 CO-01 GDRD1127 279.8 Qtz-cal-chl-bn vein, BP – calcite -3.0 14.6 CO-04 GDRD1144 292.8 Qtz-cal-chl-cpy vein, BP, BX – calcite, pink -3.9 - CO-44 GDRD1180 259.2 Qtz-cal-chl-cpy vein, BP, BD – calcite, pink to white -5.6 14.3 CO-47 GDRD1181 139 Qtz-cal-chl-py vein, BP – calcite, buff -3.1 11.6 CO-33 GDRD1181 158.5 Qtz-cal-hem vein, BP – calcite, white to pink -4.7 18.8 Nexus CO-18 GD080-07 205.0 Laminated marlstone -1.2 18.2 CO-26 GD078-07 145.3 Qtz-cal-py-hem vein, discordant – calcite – yellowish -8.9 17.9 CO-30 GD078-07 149.5 Cal-py-hem vein, discordant – calcite, white -7.9 18.2 CO-31 GD081-07 50.5 Qtz-cal-chl vein, discordant – calcite, white (cuts dolerite dike) -6.9 14.8

Plutus CO-21 PD014-06 178.7 Marlstone, brown to red -1.8 14.2 CO-20 PSRD270 67.5 Laminated marlstone -1.3 15.3 CO-23 PSRD1188 470.1 Marlstone -1.0 15.0 CO-22 PSRD1252 519.95 Laminated marlstone, recrystallized 0.1 14.4 CO-12 PD014-06 181.0 Cal-hem vein, BP – calcite, pink to white (Ngwako Pan Fm) -2.6 12.3 CO-15 PSRD265 128.0 Qtz-cal-bn vein, discordant – calcite – pink to buff -3.2 14.8 CO-09 PSDD310 148.9 Qtz-cal-cc vein, discordant – calcite, pink -3.0 19.3 CO-10 PSDD310 150.1 Qtz-cal-chl-cc vein, BP – calcite, white to pink -4.1 18.2 CO-16 PSRD1187 464.5 Qtz-cal-cpy vein, discordant – calcite, white -3.6 14.5 CO-13 PSRD1188 451.0 Qtz-cal-hem vein, BP – calcite, white to pink -7.5 18.8

139

Table B-1: (con’t) Sample ID Drill Hole Depth Description – mineral sampled δ13C (VPDB) δ18O (SMOW) (m)

CO-17 PSRD1188 459.3 Qtz-cal-cpy-(bn) vein, discordant – calcite, white -6.3 14.5 CO-05 PSRD1188 460.9 Qtz-cal-chl-cpy vein, BP – calcite, pink to buff -3.6 14.9

CO-07 PSRD1188 462.9 Qtz-cal-cpy vein, discordant – calcite -6.1 14.8

CO-06 PSRD1188 466.5 Qtz-cal-cpy vein, discordant – calcite -2.0 14.7

CO-14 PSDD1251 454.7 Qtz-cal-bn vein, discordant – calcite, white -3.2 14.9

CO-24 PSDD1251 454.7 Qtz-cal-bn vein, discordant – calcite, white -3.6 14.0

CO-25 PSDD1251 455.5 Qtz-cal-bn vein, discordant – calcite, white -4.5 15.0

Petra CO-39 PTDD809 57.0 Qtz-cal-spec vein, discordant – calcite, brown to buff (possible -6.9 12.8 ankerite) CO-42 PTDD874 27.0 Qtz-cal-chl vein, discordant – calcite, white -13.2 14.9 CO-49 PTDD874 38.6 Cal-hem vein, discordant – calcite, pink to buff -12.3 14.7 CO-43 PTDD874 44.0 Qtz-cal-chl vein, BP, BX – calcite, pink to buff -11.4 14.6 CO-36 PTDD874 93.7 Qtz-cal-cc vein, BP – calcite, pink to buff -0.1 14.2 CO-40 PTDD874 95.4 Qtz-cal-cc-hem vein, BP – calcite, white -0.6 14.2 CO-41 PTDD875 45.2 Qtz-cal-chl-cc vein, BP, BX – calcite, orange -2.4 22.1

CO-38 PTDD875 45.5 Qtz-cal vein, BP – calcite, white to pink -3.3 14.4 CO-48 PTDD876 37.9 Qtz-cal-chl-cc-hem vein, BP – calcite, buff -3.5 15.0 CO-37 PTDD876 44.4 Cal vein, discordant – calcite, white (Ngwako Pan Fm) -1.2 13.7

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Table B-2: Results from sulfur stable isotope analyses.

Sample # Drill Hole ID Meters above Sample Description Mineral δ34S Reference footwall Sampled (‰) Bedding-parallel sulfide-bearing veins, D'Kar Formation, Boseto S029 GD078-07 - Qtz-cal-sph-py Sph -30.9 This study S020 PTDD874 - Qtz-cal-cc Cc -15.6 This study S005 PSRD1188 - Qtz-cal-chl-cpy Cpy -19.3 This study S010 PSRD271 - Bn-cpy Bn -37.6 This study S001 GDRD1127 - Qtz-cal-(chl)-bn Bn -29.9 This study S002 GDRD1127 - Qtz-cal-chl-hem-cpy Cpy -9.6 This study S003 GDRD1127 - Qtz-cal-chl-bn-(cc) Bn -18.4 This study S011 GDDD1009 - Qtz-cal-cc - BD Cc -25.8 This study S018 GDDD1110 - Qtz-cal-bn-cpy-(chl) - BD Cpy -10.3 This study S022 GDRD1180 - Qtz-cal-chl-ser-cpy - BD Cpy -37 This study S025 GDDD1110 - Qtz-cal-chl-ser-cc - BD Cc -15.9 This study S028 GDDD512 - Qtz-cal-chl-cc - BD Cc -29 This study

Discordant sulfide-bearing veins, D'Kar Formation, Boseto S016 GD078-07 - Qtz-cal-hem-py Py -10.2 This study S017 GD078-07 - Cal-hem-py Py -31.2 This study S008 PSRD1251 - Qtz-cal-bn Bn -29.6 This study S009 PSRD1251 - Qtz-cal-bn Bn -29.2 This study S012 PSDD310 - Qtz-cal-chl-cc-(bn) Cc -17.2 This study S014 PSRD271 - Qtz-cal-cc Cc -29.9 This study S019 GDRD1127 - Qtz-cal-py Py -5.5 This study S024 GDRD1180 - Qtz-cal-py Py -12 This study

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Table B-2: Results from sulfur stable isotope analyses (con’t).

Sample # Drill Hole ID Meters above Sample Description Mineral δ34S Reference footwall Sampled (‰)

Shear-hosted sulfides, D'Kar Formation, Boseto S006 PSRD1187 - Qtz-cal-bn - SH BX Bn -35.9 This study S007 PSRD1188 - Bn-rich SH Bn -27.3 This study S013 PSRD270 - Qtz-cal-cc - SH BX Cc -31.2 This study S004 GRDR1144 - Qtz-cal-chl-cpy - SH BX Cpy -23 This study S026 GDDD508 - Qtz-cal-chl-ser-cc Cc -22.5 This study S027 GDDD1110 - Bn-(cc) Bn -21.7 This study

Disseminated sulfides, D'Kar Formation, Boseto S015 PD014-06 Cleavage-parallel lenticles Cc -3.8 This study S021 GDRD1181 Bleb - patchy disseminated Py -18.2 This study S023 GDRD1188 Bleb - patchy disseminated Py -12.3 This study

Klein Aub mine, Namibia Kobos Massive Sulphides Py, Cpy, Sph, 0.2 Ruxton (1986) Gal Galena Qtz-ga vein Gal -0.8 Ruxton (1986) B2 HM M Disseminated cc Cc -35.2 Ruxton (1986) B3 TS M Disseminated cc Cc -34.5 Ruxton (1986) 40980 Disseminated cc Cc -21.4 Ruxton (1986) 4051 Qtz-cal-cc-bn vein Cc, Bn -28.9 Ruxton (1986) Bornite Bn after py cubes Bn, Py -20.9 Ruxton (1986) B4Pyrite Disseminated py cubes Py -30.4 Ruxton (1986) AUNPyrite Disseminated py cubes Py -27.4 Ruxton (1986)

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Table B-2: Results from sulfur stable isotope analyses (con’t).

Sample # Drill Hole ID Meters above Sample Description Mineral δ34S Reference footwall Sampled (‰)

Witvlei, Namibia ESK 9 Disseminated bn Bn -22.2 Ruxton & Clemmey (1986) ESK 54 Disseminated bn Bn -11.8 Ruxton & Clemmey (1986) OKW 5 Chalcopyrite/alb/cal nodule Cpy -8.9 Ruxton & Clemmey (1986) ESK60 Chalcopyrite vein - stratiform Cpy -17.4 Ruxton & Clemmey (1986) Gypsum Gypsum nodules Gyp -16 Ruxton & Clemmey (1986)

Stratigraphic Interval Sampling, Zeta Deposit

S031 GDRD1182 1 S1 parallel qtz-cal-cc vein Cc -25.6 This study

S032 GDRD1182 4.5 S1 parallel qtz-cal-py-cc vein Cc -29.3 This study S033 GDRD1182 10 Disseminated py bleb Py -23.4 This study

S034 GDRD1182 18.6 S1 parallel qtz-cal-cpy-bn vein Cpy -9.6 This study S035 GDRD1182 42.5 Disseminated py cube Py -16 This study

S036 GDRD1149 1 S1 parallel qtz-cal-chl-hem-cc vein Cc -23.6 This study S037 GDRD1149 5.8 S1 parallel qtz-cal-cpy-(py-bn-chl) vein Cpy -28.7 This study

S038 GDRD1149 9.1 S1 parallel py-cal stringer Py -34.4 This study

S039 GDRD1149 20 S1 parallel py-cal stringer Py -18.1 This study S040 GDRD1149 49.1 Disseminated py Py -15.9 This study

S041 GDRD1137 2 S1 parallel qtz-cal-cc vein Cc -30.2 This study S042 GDRD1137 6.1 S1 parallel bn-cc stringer Bn -33.8 This study

S043 GDRD1137 10.2 S1 parallel qtz-cal-chl-bn vein Bn -30.7 This study

S044 GDRD1137 19.4 S1 parallel qtz-cal-cpy vein Cpy -13.5 This study S045 GDRD1137 40.1 Disseminated py cube Py -13.3 This study

S046 GDRD1128 1.5 S1 parallel qtz-cal-cc-(chl) vein Cc -15.7 This study S047 GDRD1128 5.5 S1 parallel qtz-cal-bn-chl vein, BX Bn -33.3 This study

S048 GDRD1128 11.2 S1 parallel qtz-cal-chl-bn vein Bn -24.4 This study

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Table B-2: Results from sulfur stable isotope analyses (con’t).

Sample # Drill Hole ID Meters above Sample Description Mineral δ34S Reference footwall Sampled (‰)

S049 GDRD1128 22.5 S1 parallel qtz-cal-chl-cpy-bn vein Cpy -9.3 This study S050 GDRD1128 60.4 S0 parallel qtz-cal-py vein Py -8.3 This study

S051 GDDD1121 0.9 S1 parallel qtz-cal-chl-cc-(bn) vein Cc -10 This study S052 GDDD1121 5.4 S1 parallel qtz-cal-chl-cc vein Cc -23.6 This study

S053 GDDD1121 13.3 S1 parallel qtz-cal-chl-bn-(cc) vein Bn -11.9 This study S054 GDDD1121 19.7 Discordant qtz-cal-py vein Py -13.3 This study S055 GDDD1121 45.1 Disseminated py cube Py -15 This study

S056 GDRD1171 1.1 S1 parallel qtz-cal-chl-cc-(hem) vein Cc -27.1 This study S057 GDRD1171 5.2 S1 parallel qtz-cal-chl-bn-(cc) vein, BX Bn -30.9 This study

S058 GDRD1171 16 S1 parallel qtz-cal-cpy vein Cpy -33.2 This study

S059 GDRD1171 34.9 S1 parallel qtz-cal-py vein Py -22.9 This study

S060 GD075-07 0.8 Shear breccia-hosted bn Bn -32.9 This study S061 GD075-07 5 S1 parallel qtz-cal-cpy-(bn) vein Cpy -37.6 This study

S062 GD075-07 8.3 S1 parallel qtz-cal-chl-cpy-(bn) vein, BX Cpy -19.5 This study S063 GD075-07 19.7 Qtz-cal-chl-cc vein, BX Cc -31 This study

S065 GD075-07 37.5 Disseminated py, parallel to S1 Py -8.9 This study S064 GD075-07 38.8 Disseminated py cube Py -14 This study

S066 GD083-07 2.8 S1 parallel cpy stringer Cpy -30.1 This study S067 GD083-07 8 S1 parallel qtz-cal-cpy stringer Cpy -43.1 This study

S068 GD083-07 31.2 S1 parallel qtz-cal-cpy vein Cpy -21.2 This study

Stratigraphic Interval Sampling, Plutus Deposit

S069 PD014-06 1 S1 parallel cc lenticles Cc -3.9 This study

S070 PD014-06 5.1 S1 parallel cc lenticles Cc -32.7 This study S071 PD014-06 14.7 Disseminated ga blebs Ga -23.1 This study S072 PD014-06 23.4 Disseminated py cubes Py -19 This study

144

Table B-2: Results from sulfur stable isotope analyses (con’t).

Sample # Drill Hole ID Meters above Sample Description Mineral δ34S Reference footwall Sampled (‰) S073 PD014-06 45.9 Disseminated py cubes Py -13.8 This study

S074 PSRD1260 1.4 S1 parallel cc stringer Cc -13 This study S075 PSRD1260 4.6 S0 parallel qtz-cal-cpy-bn vein Bn -35 This study

S076 PSRD1260 9 S0 parallel qtz-cal-cpy vein Cpy -32.8 This study S077 PSRD1260 19.6 Disseminated ga blebs, carb alt Ga -17 This study S078 PSRD1260 36.3 Disseminated py cube Py -9 This study

S079 PSRD1256 2.4 S1 parallel cc lenticles Cc -25.3 This study S080 PSRD1256 6.8 S1 parallel cpy stringer Cpy -30.2 This study

S081 PSRD1256 9.4 So parallel qtz-cal-py vein Py -28.3 This study S082 PSRD1256 21.1 Disseminated ga blebs Ga -11 This study S083 PSRD1256 27.4 Disseminated py cubes Py -28 This study

S084 PSRD1256 32.9 S0 parallel qtz-cal-chl-ga vein Ga -14.2 This study

S085 PSDD1254 1.5 Cc lenticles Cc -15.3 This study S086 PSDD1254 5.8 S0 parallel cpy lenticles Cpy -36.6 This study

S087 PSDD1254 11 S0 parallel cal-py vein Py -24.5 This study S088 PSDD1254 22.5 Disseminated py cube Py -25.8 This study

S089 PSRD1257 0.8 S0 parallel qtz-cal-chl-cc vein, BX Cc -22.9 This study S090 PSRD1257 5 S0 parallel qtz-cal-cpy vein Cpy -31.4 This study S091 PSRD1257 10.9 Disseminated py cubes Py -14.2 This study S092 PSRD1257 21.3 Disseminated py cubes Py -22.6 This study S093 PSRD1257 40.3 Disseminated py cubes Py -16.5 This study

S094 PSRD1250 0.6 S0 parallel qtz-cal-cc-(hem) vein, BX Cc -27.4 This study S095 PSRD1250 5.6 S1 parallel cpy lenticles Cpy -32.1 This study S096 PSRD1250 11.7 Disseminated py cubes Py -22.8 This study S097 PSRD1250 26.5 Disseminated py patches Py -11.1 This study

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Table B-2: Results from sulfur stable isotope analyses (con’t).

Sample # Drill Hole ID Meters above Sample Description Mineral δ34S Reference footwall Sampled (‰) S098 PSRD1250 39.5 Discordant qtz-cal-py vein Py -15 This study

S099 PSRD1188 1.3 S0 parallel qtz-cal-cc vein Cc -19 This study S100 PSRD1188 6 S0 parallel bn stringer Bn -30 This study S101 PSRD1188 11.1 Disseminated py cubes Py -23.2 This study S102 PSRD1188 41.7 Disseminated py cubes Py -8.2 This study

S103 PSRD1252 2.4 S0 parallel cc stringer Cc -10.3 This study S104 PSRD1252 4.8 S1 parallel bn lenticles Bn -39.4 This study S105 PSRD1252 11.8 Py-cal nodules Py -10.2 This study

S106 PSRD1252 21.5 S1 parallel ga lenticles Ga -19.7 This study

Disseminated pyrite cubes from upper D’Kar Formation – Zeta Deposit S107 GDDD1008 357 Disseminated pyrite cube Py -5.5 This study S108 GDDD1008 417 Disseminated pyrite cube Py -4.7 This study S109 GDDD1008 429 Disseminated pyrite cube Py -1.5 This study S110 GDDD1008 563 Disseminated pyrite cube Py 3.5 This study

Vein-hosted barite – Zeta Deposit S111 GDRD1143 NPF Discordant qtz-cal-bar-hem vein Bar 5.4 This study

Abbreviations: S0 = bedding parallel, S1 = cleavage parallel, BD = boudinage, BX = brecciated, SH = sheared, alb = albite, bar = barite, bn = bornite, cal = calcite, cc = chalcocite, chl = chlorite, cpy = chalcopyrite, py = pyrite, qtz = quartz, sph = sphalerite, ga = galena, gyp = gypsum, NPF = Ngwako Pan Formation.

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