Download Geology and paragenesis of the Boseto copper deposits, Kalahari

GEOLOGY AND PARAGENESIS OF THE BOSETO COPPER DEPOSITS,
KALAHARI COPPERBELT, NORTHWEST BOTSWANA
by
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
ii
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 highgrade 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 carbonatechlorite-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
iv
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
v
REFERENCES CITED ............................................................................................................... 128
APPENDIX A: LITHOLOGY PHOTOGRAPHS ..................................................................... 134
APPENDIX B: STABLE ISOTOPE DATA .............................................................................. 139
vi
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
vii
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
viii
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
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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
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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.
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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
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
2
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.
5
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 volcanosedimentary 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
6
Figure 2.1: Lithostratigraphy of the Ghanzi-Chobe Belt (modified from Ramokate et al., 2000).
7
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
9
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 postorogenic 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
11
(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 GhanziChobe 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 UPb 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
14
(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.
15
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
16
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 northeaststriking 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
19
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 westnorthwest-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).
20
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. No vertical exaggeration.
22
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 crossbedded, 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.
24
Figure 4.1: Geologic map of the Boseto area.
25
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.
26
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
27
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 finegrained 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.
29
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 mudstonesiltstone beds are capped by a 1-2 meter thick, poorly sorted, green to brown, siltstone to finegrained 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.
30
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 thickbedded 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 siltstonedominated 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 siltstonedominated 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 subcycle (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
33
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 sandstonedominated facies. It is 40 meters thick in the southwest and pinches out to the northeast.
35
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).
36
Figure 4.5: Strike-parallel stratigraphic section of the Zeta deposit.
37
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 highenergy 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.
38
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 deltafront 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 nondeposition. 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.
43
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.
44
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 gradecontrol 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.
45
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-northeasttrending lineaments that are inferred to be faults.
47
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.
48
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. Tennantiteoccurs 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 pyriteonly (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 northwestsoutheast-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
51
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. 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.
53
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 reprecipitated 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 finegrained 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, laminaescale shear zones overprint bedding-parallel quartz-albite veinlets, skeletal albite may be
57
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.
58
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® falsecolored 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 mudsized 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.
59
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 quartzcalcite-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
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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 0.55 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 20meter-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.
64
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.
65
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 beddingparallel 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.
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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).
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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 layerparallel 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
70
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). Semimassive 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
73
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.
74
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.
75
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-calcitesulfide 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.
76
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.
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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.
78
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.
81
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 layerparallel 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).
82
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 interboudin 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.
85
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.
86
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).
87
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. ±
88
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).
89
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 interboudin 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.
90
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
92
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).
94
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
95
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
97
Table 6.1: Microthermometry data for primary fluid inclusions analyzed in this study.
Number of primary
fluid inclusions
analyzed within the
assemblage
Th (˚C)
2
165-170
-11.5 to -11.0
15-20
2
230-235
-11.2
15-20
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
3
225-235
-17
15-20
Primary Fluid
Inclusion
Assemblage
Primary Fluid Inclusion
Shape
G1110-187a
equant to negative crystal
G1110-187b
irregular
G1110-187c
Associated minerals
Carbonate, sulfide
Carbonate, sulfide
98
TmICE
Salinity
(wt % NaCl)
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.
99
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.
Na+, Ca2+, K+, Mg2+, Sr2+, Ba2+, NH4+, Li+) and anions (e.g. Cl-, Br-,CH3OH-, SO42-, S2O32-,CO32-,
PO42- 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).
101
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
256.6 m
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
GDRD1127
256.6 m
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
GDRD1127
256.6 m
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
GDRD1127
256.6 m
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
GDRD1127
268.4
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
GDRD1127
268.4 m
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
GDRD1180
154.4 m
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
GDRD1180
154.5 m
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
GDRD1180
154.5 m
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
PSRD1187
505.75
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
PSRD1187
505.75 m
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
PSRD1187
505.75m
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
102
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.
103
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).
104
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 earlyand 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.
107
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.
108
Table 7.1: Typical carbonate producing processes and accompanying δ13C values.
T (˚C)
δ13C values
Process
Depth
Evaporation
subaerial
15-40˚C
Normal marine limestone
Decomposition of organic matter
Upper few centimeters to ~ 2 meters below
the seafloor (early diagenetic regime)
15-40˚C
CO2 (aq) and HCO3-
Mixing of HCO3- and pore fluids
Upper few centimeters to ~ 2 meters of
seafloor (early diagenetic regime)
15-40˚C
Diagenetic calcite
Methanogenic bacteria decomposition of
residual organic matter
> 2 meters below seafloor (diagenetic
regime)
60-80˚
Methane
-60‰ to -110‰
Catagenesis
2-5 km below surface (thermocatalytic
regime)
> 80˚C
Various hydrocarbon gases,
CO2, H2S
-40‰ (for CH4)
Metamorphism: at high metamorphic
grades, kerogen  graphite; isotopic
exchange with carbonate in metasediments
> 5 km
> 200˚C
graphite
-15‰
Metamorphism: at high metamorphic
grades, presence of H2O results in
decomposition of graphite
> 5 km
> 200˚C
CO2;
-2‰ to -8‰;
CH4
-20‰ to -30‰
Modified text from Ohmoto and Goldhaber (1997).
109
Product
0‰
-25‰
-5‰ to -15‰
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 B2; 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.
111
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.
112
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:
187
Os = 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 beddingparallel 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
8.2
187
Os/188Os values of 0.2 and 1.0.
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
115
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 wellconstrained 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.
117
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.
118
Table 8.1: Re-Os chronometry data.
Re,
ppb
2σ
bn
0.798
LL-611
cpy
LL-612
cpy
AIRIE
Run #
Sulfide
LL-610
Total
Os,
ppb
2σ
0.002
0.0115
0.0001
7.19
0.01
0.0757
3.289
0.006
0.0419
187
Re/
Os
2σ
0.004
1391
0.0010
0.009
0.0005
0.025
Common
Os, ppb
187
Model Age,
Ma
assumed
187
Re/188Os
initial 2σ =
0.2
Model
Age, Ma
assumed
187
Re/188Os
initial 2σ =
1.0
Average
Model Age,
reflecting 2σ
uncertainties
0.610
1029
995
1012 ± 17 Ma
0.3
0.636
918
910
914 ± 4 Ma
0.02
0.559
496
442
469 ± 27 Ma
Os/
Os
2σ
rho
6
24.3
0.1
5792
17
89.4
646
2
5.56
188
188
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.
119
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
120
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 postorogenic 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 beddingparallel 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.
123
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
124
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 layerparallel 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 highgrade 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
126
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).
127
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Bingen, B., and H. Stein. “Re-Os dating of the Ørsdalen W-Mo district in Rogaland, S Norway,
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Borg, G. “The Koras-Sinclair-Ghanzi rift in southern Africa, volcanism, sedimentation, age
relationships, and geophysical signature of a late Middle Proterozoic rift system.”
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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 mudstonesiltstone. 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 hangingwall 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
(m)
Description – mineral sampled
δ13C (VPDB)
δ18O (SMOW)
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
(m)
Description – mineral sampled
δ13C (VPDB)
δ18O (SMOW)
CO-17
PSRD1188
459.3
CO-05
PSRD1188
460.9
Qtz-cal-cpy-(bn) vein, discordant – calcite, white
-6.3
14.5
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
ankerite)
-6.9
12.8
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
140
Table B-2: Results from sulfur stable isotope analyses.
Sample #
Drill Hole ID
Meters above
footwall
Sample Description
δ34S
(‰)
Reference
Sph
-30.9
This study
Mineral
Sampled
Bedding-parallel sulfide-bearing veins, D'Kar Formation, Boseto
S029
GD078-07
-
Qtz-cal-sph-py
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
141
Table B-2: Results from sulfur stable isotope analyses (con’t).
Sample #
Drill Hole ID
Meters above
footwall
Sample Description
Mineral
Sampled
δ34S
(‰)
Reference
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
Py, Cpy, Sph,
Gal
Gal
0.2
Ruxton (1986)
Klein Aub mine, Namibia
Kobos
Massive Sulphides
Galena
Qtz-ga vein
-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)
142
Table B-2: Results from sulfur stable isotope analyses (con’t).
Sample #
Drill Hole ID
Meters above
footwall
Sample Description
Mineral
Sampled
δ34S
(‰)
Reference
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)
S1 parallel qtz-cal-cc vein
Cc
-25.6
This study
Cc
-29.3
This study
Stratigraphic Interval Sampling, Zeta Deposit
S031
GDRD1182
1
S032
GDRD1182
4.5
S1 parallel qtz-cal-py-cc vein
S033
GDRD1182
10
Disseminated py bleb
S034
GDRD1182
18.6
S1 parallel qtz-cal-cpy-bn vein
S035
GDRD1182
42.5
Disseminated py cube
S036
S037
GDRD1149
GDRD1149
1
5.8
S1 parallel qtz-cal-chl-hem-cc vein
S1 parallel qtz-cal-cpy-(py-bn-chl) vein
S038
GDRD1149
9.1
S039
GDRD1149
20
S040
GDRD1149
S041
S042
Py
-23.4
This study
Cpy
-9.6
This study
Py
-16
This study
Cc
Cpy
-23.6
-28.7
This study
This study
S1 parallel py-cal stringer
Py
-34.4
This study
S1 parallel py-cal stringer
Py
-18.1
This study
49.1
Disseminated py
Py
-15.9
This study
GDRD1137
GDRD1137
2
6.1
S1 parallel qtz-cal-cc vein
S1 parallel bn-cc stringer
Cc
Bn
-30.2
-33.8
This study
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
S047
GDRD1128
GDRD1128
1.5
5.5
S1 parallel qtz-cal-cc-(chl) vein
S1 parallel qtz-cal-bn-chl vein, BX
Cc
Bn
-15.7
-33.3
This study
This study
S048
GDRD1128
11.2
S1 parallel qtz-cal-chl-bn vein
Bn
-24.4
This study
143
Table B-2: Results from sulfur stable isotope analyses (con’t).
Mineral
Sampled
δ34S
(‰)
Reference
S1 parallel qtz-cal-chl-cpy-bn vein
S0 parallel qtz-cal-py vein
Cpy
Py
-9.3
-8.3
This study
This study
0.9
5.4
S1 parallel qtz-cal-chl-cc-(bn) vein
S1 parallel qtz-cal-chl-cc vein
Cc
Cc
-10
-23.6
This study
This study
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
S057
GDRD1171
GDRD1171
1.1
5.2
S1 parallel qtz-cal-chl-cc-(hem) vein
S1 parallel qtz-cal-chl-bn-(cc) vein, BX
Cc
Bn
-27.1
-30.9
This study
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
S061
GD075-07
GD075-07
0.8
5
Shear breccia-hosted bn
S1 parallel qtz-cal-cpy-(bn) vein
Bn
Cpy
-32.9
-37.6
This study
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
S067
GD083-07
GD083-07
2.8
8
S1 parallel cpy stringer
S1 parallel qtz-cal-cpy stringer
Cpy
Cpy
-30.1
-43.1
This study
This study
S068
GD083-07
31.2
S1 parallel qtz-cal-cpy vein
Cpy
-21.2
This study
Sample #
Drill Hole ID
Meters above
footwall
S049
S050
GDRD1128
GDRD1128
22.5
60.4
S051
S052
GDDD1121
GDDD1121
S053
Sample Description
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).
Mineral
Sampled
δ34S
(‰)
Reference
Disseminated py cubes
Py
-13.8
This study
S1 parallel cc stringer
S0 parallel qtz-cal-cpy-bn vein
Cc
Bn
-13
-35
This study
This study
S0 parallel qtz-cal-cpy vein
Cpy
-32.8
This study
Disseminated ga blebs, carb alt
Ga
-17
This study
36.3
Disseminated py cube
Py
-9
This study
PSRD1256
PSRD1256
2.4
6.8
S1 parallel cc lenticles
S1 parallel cpy stringer
Cc
Cpy
-25.3
-30.2
This study
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
S086
PSDD1254
PSDD1254
1.5
5.8
Cc lenticles
S0 parallel cpy lenticles
Cc
Cpy
-15.3
-36.6
This study
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
S090
PSRD1257
PSRD1257
0.8
5
S0 parallel qtz-cal-chl-cc vein, BX
S0 parallel qtz-cal-cpy vein
Cc
Cpy
-22.9
-31.4
This study
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
S095
PSRD1250
PSRD1250
0.6
5.6
S0 parallel qtz-cal-cc-(hem) vein, BX
S1 parallel cpy lenticles
Cc
Cpy
-27.4
-32.1
This study
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
Sample #
Drill Hole ID
Meters above
footwall
S073
PD014-06
45.9
S074
S075
PSRD1260
PSRD1260
1.4
4.6
S076
PSRD1260
9
S077
PSRD1260
19.6
S078
PSRD1260
S079
S080
Sample Description
145
Table B-2: Results from sulfur stable isotope analyses (con’t).
Mineral
Sampled
δ34S
(‰)
Reference
Discordant qtz-cal-py vein
Py
-15
This study
1.3
6
S0 parallel qtz-cal-cc vein
S0 parallel bn stringer
Cc
Bn
-19
-30
This study
This study
PSRD1188
11.1
Disseminated py cubes
Py
-23.2
This study
S102
PSRD1188
41.7
Disseminated py cubes
Py
-8.2
This study
S103
S104
PSRD1252
PSRD1252
2.4
4.8
S0 parallel cc stringer
S1 parallel bn lenticles
Cc
Bn
-10.3
-39.4
This study
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
Sample #
Drill Hole ID
Meters above
footwall
S098
PSRD1250
39.5
S099
S100
PSRD1188
PSRD1188
S101
Sample Description
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
Bar
5.4
This study
Vein-hosted barite – Zeta Deposit
S111
GDRD1143
NPF
Discordant qtz-cal-bar-hem vein
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.
146