Delia the Dalmatian on ‘Dalmatian Rock’€¦ · Delia the Dalmatian on ‘Dalmatian Rock’ (spotted anorthosite), Eastern Bushveld. Photo courtesy of Lew Ashwal. Acknowledgements

- Economic Geology Field Trip 2012 - 1

Delia the Dalmatian on ‘Dalmatian Rock’ (spotted anorthosite), Eastern Bushveld. Photo courtesy of Lew Ashwal.

Acknowledgements This trip would not have been possible without the time and expertise donated by John Mavrogenes,

and the financial assistance of our sponsors.

Thank you for making such an incredible experience possible!

Our Platinum Sponsors Our Gold Sponsors


ECONOMIC GEOLOGY FIELD TRIP TO

SOUTH AFRICA AND NAMIBIA

CLASSES OF 2011 AND 2012

Research School of Earth Science, Australian National University

“Tristan da Cunha Matata”


Essential Information: Emergency Numbers in South Africa

112 (from a mobile phone only) 10111 (for police/fire emergencies) 10177 (for medical emergencies)

Emergency Numbers in Namibia

1011(1) (for police emergencies) See Dom’s sheet for local emergency numbers in Namibia

24-hour Australian Consular Emergency Centre: +61 2 6261 3305 Insurance Policy: Call the ACE Assistance number (+612 8907 5995) this can be done on reverse charge

Advise them of • Your name • Your policy number (ours is 01PP528643) • Name of insured (Australian National University) • Phone number, or where you are calling from • Nature of assistance required

Flight Information:

Space to exchange mobile phone numbers: (remember you must have our numbers, and we must have yours if you want to leave the group at any stage) Dom: +27 _____________ (African SIM) or 0402 496 664 in Australia Oliver: +61 424 911 338

British Airways Ticket 1 JNB-WDH Flight BA6275 WDH-JNB Flight BA6274

Ref: 722JL2

British Airways Ticket 2 JNB-WDH Flight BA6275 WDH-JNB Flight BA6274

Ref: 72UYJG

Air Namibia Ticket JNB-WDH Flight SW701 WDH-JNB Flight SW702

Ref: ITD2858806IBE or RH856 Mavro Temma Andy S-J

Dean Snowy

Tiff Chris Sean

Oliver Michaela

Tim (Caleb) Alex Elle

Dom Tazz

Dylan Vikraman

Matt Heather Sarah

Morgan

9 April 12:00-13:00 19 April 13:55-16:50

9 April 12:00-13:00 19 April 13:55-16:50

9 April 1:00-2:00 19 April 16:55-19:40

Area Codes South Africa

+27 Namibia

+264 Australia +61 (2)

dominique

Rectangle


Table of Contents

Cover illustration by Chris Harris-Pascal

5 Geology Maps 7 Itinerary

9 The Vredefort Dome Sarah Buckerfield

13 The Distinguishing Characteristics of Impact Craters Tim Hobern

17 The Witwatersrand Basin Eleanor Peterson

19 A Brief History of South African Mining Andy Clark

21 Overview of Namibian Geology Vikraman Selvaraja

25 The Navachab Gold Mine Matthew Peacock

29 The Rössing Uranium Mine Tiffany Halcon

32 Road Map: Navachab to the Coast Dominique Tanner

33 Snowy’s Road Guides Snowy Haiblen

35 Grosse and Klein Spitzkuppe Michaela Flanigan

39 The Uis Sn-Ta Pegmatites Morgan Williams

43 The Damara Orogen Temma Carruthers-Taylor

47 Gondwana Break-up and Large Igneous Provinces Oliver Nebel

49 Snowball Earth and Namibia Dean Erasmus

51 Road Guide: Uis-Otjiwarongo-Tsumeb Dominique Tanner

53 The Tsumeb Ore Body Alex Moody

56 Pegmatites: Did you know…? Dominique Tanner

57 The Okorusu Fluorspar Mine S-J Collum

61 The Otjihaenamaparero Dinosaur Footprints Dylan Singh

65 The Windhoek Meteorites Sean Jefferson

67 The Bushveld Complex Heather Marman

71 Ore Mineral Help Guides Tarun (Tazz) Whan

74 Appendix of Useful Figures Dominique Tanner and Tarun (Tazz) Whan

Field guide compiled by Dominique Tanner




Trip Itinerary: Friday 6th April (Public Holiday) 2pm: Meet at Johannesburg Airport. Drive to Vredefort (2 hours), stopping for grocery shopping en route. If time, we will stop to look at a coesite-stishovite melt. Stay at Thwane Bushcamp in tents for the next few days. Cook our own dinner. After dinner, Sarah will present her talk. Saturday 7th April (Easter Weekend) Breakfast, lunch and dinner are self-prepared in the lapa/braai area. Drive around the Vredefort Dome, looking at the world-class outcrops including shatter cones and a brecciated pseudotachylite mine. We may potentially get to look at an old Wits Au mine and komatiites. After dinner, Elle will present her talk. Sunday 8th April (Easter Weekend) Breakfast, lunch and dinner are self-prepared in the lapa/braai area. Drive to Habula Lodge, to do a 1-hour hike around the Vredefort Dome. This walk will include possible impact folds and inverted stratigraphy. After dinner, Vikraman will present his talk. Monday 9th April (Public Holiday) Drive to Johannesburg after breakfast (2 hours). 12.00pm/1.00pm flights: Fly to Windhoek, Namibia. Take a shuttle bus from Windhoek Airport to the Cardboard Box Hostel, Windhoek. Talk to be given by Matt, before dinner. DIY dinner in Windhoek. Tuesday 10th April Breakfast (pancakes and coffee) provided by the hostel. Meet the coach and coach driver, and drive to Navachab Au Mine (2 hrs). 10am or earlier: Tour of Navachab Au mine. If spare time, drive 45 mins (each way) to see the Rubicon Pegmatite Mine. Drive to Swakopmund (2 hrs 15 mins). Stay in Villa Weise Backpackers Accommodation. 7.30pm: Dinner at Swakopmund Brauhaus. After dinner, Tiff will give her talk.

Wednesday 11th April Breakfast provided by the hostel. Day-trip from Swakopmund to Rossing (1 hr drive each way). 10am: Tour of Rossing uranium mine. BYO lunch and water. 2pm: Possible trip to the Naukluft National Park to see the Welwitschia Plant. Then free time in Swakopmund. Stay in Villa Weise Backpackers Accommodation. Before dinner, Michaela will give her talk. 7.30pm: Dinner at Swakopmund Brauhaus. Thursday 12th April Breakfast provided by the hostel. 2 hour drive from Swakopmund to look at Grosse and Klein Spitzkuppe in the Spitzkuppe Nature Reserve (short hiking trails to lookouts). Drive to Uis (1 hr 35 mins). Look at some road outcrops en route? Stay in Uis at the Brandberg Rest Camp. Dinner provided by the Rest Camp. Talk given by Morgan after dinner. Friday 13th April Breakfast provided by the rest camp. Spend the day looking at the pegmatites around Uis, including a private tour to the abandoned Uis mine and road cuttings. Possible trip to West Brandberg? (>1 hr drive). Stay in Uis at the Brandberg Rest Camp. Dinner provided by the rest camp. Talk given by Temma after dinner. Saturday 14th April Breakfast provided by the rest camp. Spend the morning fossicking around Uis, and/or looking at mineral stalls. Drive to Otjiwarongo (3 hrs 30 mins), stopping to look at road outcrops including the Erongo ring dyke. Stay in Bushpillow Guesthouse (some of the group will stay in another nearby guesthouse). A local restaurant will provide a set menu for dinner. Talks to be given by Alex and Chris after dinner.


Sunday 15th April Breakfast provided by the guesthouse. Day trip from Otjiwarongo to Tsumeb (2 hrs 20 mins each way). 11am: Tsumeb Museum (opened just for us!). Spend the rest of the day looking around Tsumeb (and its mineral dealers?). Possible Snowball Earth outcrop to see on the way back. A local restaurant will provide a set menu for dinner. Talk to be given in Otjiwarongo by S-J. Monday 16th April Breakfast provided by the guesthouse. Day trip from Otjiwarongo to Okorusu (2 hrs each way). 9.30am: Tour of Okorusu Fluorspar Mine. Look at outcrops on the way back? A local restaurant will provide a set menu for dinner. Talk to be given after dinner by Dylan. Tuesday 17th April Breakfast provided by the guesthouse. Leave Otjiwarongo in the morning and drive to Otjiheanamaperero (1 hr 30 mins). Walk to see the dinosaur footprints (500m). Find lunch in Okahandja (2 hr drive). Drive to Windhoek (1 hr), stopping to look at road outcrops along the way. Stay the night in Windhoek, at the Cardboard Box hostel. Talk to be given by Sean, before dinner. DIY dinner in Windhoek. Wednesday 18th April Breakfast (pancakes and coffee) provided by the hostel. Day-trip to see the Matchless Cu Mine (1 hr drive each way). If time in the afternoon, we will go and see the museum at the Namibian Geological Survey. Stay the night in Windhoek, at the Cardboard Box hostel. Talk to be given by Heather, before dinner. DIY dinner in Windhoek.

Thursday 19th April Breakfast (pancakes and coffee) provided by the hostel. People flying British Airways get a shuttle bus to Windhoek Airport, and fly to Johannesburg (1.55pm flight arr. 4.50pm). Air Namibia people get to spend the morning in Windhoek, including going to the local shopping centre to observe the meteorites. Time to by gifts, souvenirs etc. 1.15pm People flying Air Namibia get a shuttle bus to Windhoek Airport, and fly to Johannesburg (4.55pm flight arr. 7.40pm). Dinner at the airport, if not provided by the plane. Drive to Rustenburg (3 hrs). No time for talks, as we will arrive late at night and have an early morning rise. Friday 20th April Breakfast provided by the hostel. 7am: Tour of Rustenburg Pt Mines. Drive back to Johannesburg (3 hrs). Stay at the Backpacker’s Ritz of Johannesburg. Mav, Dom, Tazz and Olli leave. Potkje and rice for dinner provided by the hostel. Saturday 21st April Cooked breakfast provided by the hostel. Shuttle buses available from the hostel to Johannesburg Airport. End of the trip. Note: Place names in italics are shown on the Namibian geology map.


The Vredefort Impact Structure Written by Sarah Buckerfield

Introduction The Vredefort Dome is possibly the world’s largest impact structure, with an estimated diameter of 180-300 km and current levels of exhumation exceeding 8 km. Located in the Witwatersrand Basin of South Africa, it was formed at 2023±4 Ma in Archaean and Palaeoproterozoic rocks of the Kaapvaal craton. The impact shockwave and subsequent readjustment of the crust below the crater have resulted in impact melt rocks, micro deformation features as well as large-scale fractures and faults, high-pressure silica polymorphs, high strain rate folds, and exhumation exposing Archaean basement rocks. The dome is best exposed in the North and North-west; the South-eastern sector is mainly hidden under shales and dolerite of the Phanerozoic. The 40 km wide core of the dome comprises Archaean basement gneisses of amphibolite to granulite grade, up to 3.6 Ga in age. The core provides evidence of crustal processes in the Kaapvaal Craton over 3 billion years ago; four deformation events are believed to have resulted in the greenschist to granulite facies metamorphism. The collar comprises sedimentary and volcanic sequences deposited between 3.07 and 2.1 Ga in a series of basins in the Kaapvaal Craton. The units are overturned (the youngest located on the rim), and include andesite and felsic lavas and rift-generated clastic sediments. From 2.97-2.71 Ga, approximately 7 km of clastic sediments were deposited, known as the Witwatersrand Supergroup. Almost half the world’s gold has been mined from this sequence, with depositional environments believed to range from fluvial to subtidal. The dominant ore mineral, pyrite, is found in both rounded and crystalline form indicating both concentration by sedimentation and in situ growth. Eruption of flood basalts followed by rift sediments at 2.714 Ga terminated the Witwatersrand sedimentation. Sedimentation of the Transvaal Supergroup commenced at approximately 2.6 Ga, when much of the crated was covered by a shallow sea producing dolomite and iron formations. From 2.35-2.1Ga, the 3 km Pretoria Group of Argillaceous-arenaceous sediments was deposited. The remaining geological history is boring. Two distinct classes of melt rocks have been recognised. The Vredefort Granophyre is an impact melt rock comprising impact melt dikes up to kilometres in length, found at the core-collar contact and in the western core. The dikes occur in radial and tangential fractures related to the centre of impact, have a uniform bulk composition, and contain abundant clasts and geochemical tracers of the impactor. These characteristics indicate the Vredefort Granophyre originated from a homogenized melt body within the crater. Breccias in impact events can be produced by shock compression and later crater modification processes. Pseudotachylite, or pseudotachylitic breccia, is defined as a friction melt, believed to result from brittle deformation. Pseudotachylite has been identified throughout the Vredefort Dome but care must be taken as the term is used loosely to describe an array of brecciated rocks. Figure 1 below is a cross section of the Vredefort Dome, indicating the structure and distribution of melt rocks after the impact event and readjustment of the crust below had occurred, as well as the current level of erosion.


Figure 1: Schematic cross-section through a peak-ring impact structure illustrating how the Vredefort impact structure would have appeared in section after the modification stage. Dashed line represents approximate level of erosion of the Vredefort impact structure. Impact melt breccia intruded downwards into the fractured basement in the central peak. Shock pressure zones in the Vredefort Dome A series of pressure zones are recognised in impact events. Distances are not highly specified and zones are recognised by observation.

1. Point of impact: approximately 100 GPa, causing vaporisation of projectile and adjacent rocks

2. Pressures of 10’s of GPa, causing melting of the crust 3. Pressures of 10-50 GPa causing mineralogical and grain/outcrop scale changes covering an

area at least the diameter of the crater. 4. Down to a few GPa, the shock wave is replaced by an elastic wave. This is the zone in

which shatter cones may form. 5. Below a few GPa, only megascopic fracturing and brecciation occurs, which can be difficult

to differentiate from normal tectonic disturbances.

Figure 2 below is a map of the distribution of these shock pressure zones in the Vredefort Structure (taken from Gibson and Reimold, Vredefort Impact Structure: A Guide to Sites of Interest, 2008).

Figure 2: Shock pressure zones in the Vredefort Structure.


Figure 3 below is a time sequence illustrating the central uplift and its subsequent collapse, causing formation of the rim syncline. The degree of overturning at 8-10 km depth in the Vredefort Structure indicates greater uplift and a more ductile response occurred than indicated by this model.

Figure 3: (from Gibson and Reimold, Vredefort Impact Structure: A Guide to Sites of Interest, 2008).


A brief guide to the outcrops we will be visiting is given below:

Sites are numbered in the order we will be visiting them. The outcrop numbers are as listed in Geology of the Vredefort Impact Structure: A guide to sites of interest (Gibson and Reimold, 2008).

1. Outcrop 12: Coesite-Stishovite Melt

Coesite and Stishovite are formed at pressures between 5-10 GPa and 12-15 GPa respectively. In this outcrop, they occur as acicular aggregates and single crystals in breccia veins and quartz, in rocks with no evidence of burial deeper than 10-20 km. Kyanite inclusions in stishovite crystals are interpreted as indicating crystallisation from a hot, highly coordinated melt phase.

2. Outcrop 14: Leeukop Quarry: Pseudotachylite and Impact Breccia

The quarry was mined for 11 years producing decorative stone, closing in 1998. The quarry provides a 3-dimensional view of structures in the Archaean Basement Complex.

3. Outcrop 13: Salvamento Quarry

The quarry was mined for pink granite/granodiorite and comprises predominantly grey trondhjemitic gneiss intruded by coarse granite and pegmatite veins.

4. Outcrop 5: Amazon Reef (Old Wits Gold Mine) The reef is part of the Kimberly Formation of the Central Rand Group (Witwatersrand Basin). Gold found in this 90m thick conglomerate unit has been concentrated by sedimentary processes, and is thought to have been deposited at the junction between a large river system draining mountains and a large inland sea. Like all gold in the Wits, both rounded and crystalline pyrite growth morphologies are observed indicating detrital and in situ growth.

5. Outcrop 8: Shatter Cones Shatter cones are fractured, conical fragments of rock with striations radiating outward from an apex. They are believed to form when rocks are subjected to shock waves associated with meteorite impact. Reconstruction of the original bed orientation has led to the suggestion that the shatter cones are all oriented towards a central point of explosion, although this hypothesis is fairly speculative. The cones are found in immature sericitic quartzites (containing quartz, sericite, chloritoid, magnetite and apatite).

6. Outcrop 24: Greenstone Complex-Komatiites, Pillow Basalts, and shatter cones The greenstone complex comprises primarily basalt of mid-greenshist grade, displaying some of the original volcanic features-including pillow basalts, komatiitic lava, and flow banding. Banded iron formations are also found with the basalt. The formation has been tentatively dated at 3.3 Ga.

7. Outcrop 11: Impact Fold Kilometre-scale quartzite folds in the North western and eastern collars are believed to have been generated during central uplift after impact, before central uplift collapse. The folds are kilometre scale, with brecciated hinges and pseudotachylitic breccia veins. The axial planes dip mostly radially inwards towards the centre of the dome, with the exception of the Parsons Rus fold which is believed to have been rotated during collapse.


The Distinguishing Characteristics of Impact Craters Written by Caleb (Tim) Hobern

This section discusses the characteristics of impact craters that distinguish them from volcanic craters. Impact craters are the basins formed when projectiles collide with a larger body, and the resulting force excavates a depression. Impact craters mark all solid planets and moons in the solar system, and since its formation, the Earth has been subjected to continual bombardment. While there are craters worldwide, marking impacts responsible for major biological and geological events, tectonic and sedimentary processes erode and bury those on Earth and other geologically active planets. The largest surviving impact crater is the 300km wide Vredefort crater, in the Free State Province, South Africa (see Fig 1.). Formed 2 billion years ago, it was initially considered a volcanic feature. However, research over the last few decades has shown that the crater was formed by the impact of a 10km wide asteroid. Volcanic craters and calderas have similar, basic structures, but are the result of eruptions. The explosive force of a volcanic eruption can result in circular craters. Similarly, calderas can form circular basins; during intense eruptions, magma chambers are emptied of their contents, leaving hollow interiors. The surface is incapable of supporting its own weight, and circular 'ring fault' forms, leading to a collapse and the formation of a crater. While both impacts and eruptions form surface depressions, impact structures can be distinguished by geographical, structural and mineralogical characteristics. The certainty of these characteristics, however, is often unrelated to their usefulness to the field geologist. Although chemical and microscopic analyses are the best indicators of an impact, general observations of topography and local geology are often more practical indicators in the field.

Figure 1: Vredefort Crater in Free State Province, South Africa. Despite the enormity of the impact, it still exhibits the parabolic-shape typical of impacts. Excavation has gouged a deep basin, and deposited an encircling ring of unearthed material. Although it was initially believed to be the result of a super-volcano, research over the past two decades has shown that is was formed by a impact that may have been powerful enough to have initiated the nearby Bushveld Igneous Formation.


Locality of Crater The location of crater is a crude way to determine its origin, and often the location will not be indicative. The value of locality as an identifying characteristic, however, is that it is immensely simple, and relies only on basic observations of terrain and local geology. The most obvious sign that an impact is extra-terrestrial is where the projectile, or fragments of it, remain in and around the crater. However, the absence of meteorite fragments does not mean that an impact did not occur; at the Chicxulub crater in Mexico, no material from the bolide has been found, despite the obvious size of the projectile. Another simple indicator of impact is the location of the crater. A crater in an area that has been volcanically active is not indicative; however, a crater in a volcanically inactive area is anomalous and suggests an impact. Similarly, a crater found at a volcanic peak is an almost certain sign that it was formed by an eruption rather than an impact. However, a crater's position in relation to the local terrain is not always helpful. Maars consist of craters alone, and are typically formed on flat terrain, meaning that a crater on level ground could have been formed by impact or eruption.

Figure 2. The geology of the Vredefort crater, showing formations intruded by various volcanics. Taken from Gibson and Reimold, 2008.


Geological Structures The tremendous force of an impact is enough to substantially affect the lithology as well as the terrain. Shock waves propagate from the impact site, shattering and melting the impacted rock, and deforming rocks at the macro and micro scale. Although not the only processes to do so, impacts will form large amounts of brecciated material at the crater site. This is concentrated as 'breccia lenses', buried beneath a layer of metamorphosed rock. The problem, however, with using these structural features to identify impact sites, is the lack of clear exposures; shock-metamorphosed rocks are formed at the centre of the impact, and will often be submerged. Generally the presence of volcanic flows will be a simple indicator that a structure was formed by volcanic activity, rather than impact. At smaller impact sites, the presence of basalt flows would indicate volcanic formation, but with larger impacts can trigger volcanism. The impact that formed the Vredefort crater is suspected to have been responsible for the formation of the contemporaneous Bushveld Igneous Formation nearby. Furthermore, the area was permeated by intrusive upwellings and associated volcanics, meaning that nearby igneous rocks do prove a crater volcanic. The immediate force of an impact will excavate an area significantly larger than the projectile itself. Although both eruptions and impacts will form a basin structure, impacts will form raised rims around the edge of the impact, generally not formed by volcanic activity. Rim structures are also useful indicators, as those formed by impacts will be composed of the older rocks than the stratigraphy below, as material is ejected and upturned by the impact. Conversely, the outer material of a volcanic flow will be the youngest material in the local stratigraphy, as it cooled after eruption at the surface. The 30km wide rim that surrounds the Vredefort crater is marked by a series of faults and folds that resulted from the impact. The brecciated hinges on these gently dipping synforms and antiforms are aligned with the impact structure as result of the stress.

Figure 3: Diagram showing the post-impact temperature of the Vredefort crater, which resulted in temperatures of over 1,000°c at its centre. These high temperatures were responsible for the formation of high-temperature rock types, strong indicators of an impact. Taken from Gibson and Reimold, 2008.


Mineralogical Characteristics Several different structures are formed at impact sites by the propagation of shock-waves, and the high pressures of the strike, the most useful of which are crystal deformation at the microscopic scale, and the formation of high-temperature compounds. The 'shock-metamorphic effects' are the surest signs of impact, but have limited use in the field as they require microscopic examination, and are often buried at great depth. High-temperature compounds are typically good indicators of an impact, as they form under intense pressures and temperatures not normally produced on Earth. Compounds such as tektites and other natural glasses, such as fused sandstones, show an environment not normally found on Earth. Similarly, shocked quartz is a clear sign of impact, and was used to identify the Chicxulub crater's origin. Its microscopic structure mimics the macroscopic brecciation and deformation, and varies from normal quartz where intense pressure deforms the crystal structure along planes, producing 'shock lamellae'. This phenomenon is not limited to quartz also observed in other crystals. Recent work has found zircons at the Vredefort crater that have undergone similar deformation.

Figure 4: Shocked quartz showing planes of deformation in plane polarised light. The faults have propagated along the crystal in response to intense pressure at low temperatures. Taken from the USGS.


The Witwatersrand Basin Written by Eleanor Peterson

The Witwatersrand basin (or the ‘Rand’) is an enormous cratonic sedimentary basin covering an elliptical area of approximately 42 000 square kilometres. It is the world’s largest quartz pebble conglomerate hosted gold deposit with associated goldfields having provided over 35% of all gold produced in the history of mankind! ‘Witwatersrand’ (translated literally from Afrikaans) means ‘ridge of white waters’ – named for the waterfalls cascading over large white quartzite cliffs that demarcate the basin boundary.

Figure 1: Basic geologic map and stratigraphic column of the Witwatersrand Basin, South Africa. (Source: Kirk et al., 2001). The basin’s stratigraphy (Figure 1) is anchored by a granite-greenstone basement, overlain by the 2250m thick Dominion unit of bimodal lavas. Above sits the Witwatersrand sedimentary complex (~7600m thick) comprised of interbedded quartzites, shales, conglomerates (frequently auriferous reefs) and sandstones. The Witwatersrand basin is sealed by the Ventersdorp volcanics (~3600m thick) and the sequence is topped by dolomites, lavas and clastic sediments of the Transvaal Supergroup (~4000m thick). The age of the Witwatersrand basin is constrained by the relative ages of the underlying Dominion and overlying Ventersdorp stratagraphic units which depict a period of sedimentation lasting from 3.07 Ga to 2.71 Ga. Detrital uraninite indicates sedimentation occurred prior to the evolution of an oxidised atmosphere and hydrosphere. Numerous models have been proposed to explain the origin and mineralisation of the more than 48000 tonnes of gold contained within the Witwatersrand Basin – the topic is contentious and a summary of three leading hypotheses is provided here (Figure 2):


(1) Placer Model (unmodified) – gold is deposited and concentrated as heavy detrital particles during sedimentation with minimal hydrothermal contribution or modification. Observations that favour this scenario include gold mineralisation confined to conglomerate layers and association of gold with inferred fluvial channels.

(2) Placer Model (modified) – similar to the previous model but with subsequent remobilisation and/or mineralisation of gold by hydrothermal/metamorphic fluids. Remobilisation may include the dissolution of detrital gold that is later re-precipitated in a more confined/concentrated form or; detrital gold grains may have acted to nucleate the precipitation of secondary gold carried by fluids.

(3) Hydrothermal/metamorphic – gold is deposited purely from hydrothermal and/or metamorphic fluids percolating through sediments after initial sediment deposition. In this scenario gold is carried in fluids from outside the basin and deposited via interaction with sediments. Fluid flow is structurally controlled through faults, fractures, stratagraphic boundaries and permeable conglomerates.

There is significant evidence to suggest the Witwatersrand basin sediments have been subject to burial and regional metamorphism and associated fluids have contributed to an epigenetic component of gold mineralisation. The extent of this is unconfirmed. Of further consideration is the huge Vredefort impact structure in the centre of the basin. It indicates the basin was hit by a sizable meteorite (~2025 Ma) and associated fluid circulation likely had an impact on gold mineralisation. Irrespective of the debate surrounding the exact mechanisms of gold mineralisation, the Witwatersrand basin is a spectacular and bountiful geologic formation. Gold reserves within it were/are so extensive that despite having been commercially mined since the 1890s, it still produces most of South Africa’s gold and provides a significant portion of total world output. Kirk, J., Ruiz, J., Chesley, J., Titley, S. and Walshe, J. (2001) A detrital model for the origin of gold and sulphides in the Witwatersrand basin based on Re-Os isotopes, Geochimica et Cosmochimica Acta, v.65(13), p. 2149-2159 – and references therein. Reimold, W.U. and Gibson, R.L. (2010) Meteorite Impact: the danger from space and South Africa’s mega-impact: Berlin, Springer-Verlag, 337 p. Robb, L. and Meyer, M. (1995) The Witwatersrand basin, South Africa: Geological framework and mineralisation processes, Ore Geology Reviews, v. 10, p. 67-94 Robb, L. (2005) Introduction to ore-forming processes: Oxford, Blackwell Publishing, 373 p. Simanovich, I.M. (2009) Auriferous Precambrian conglomerates of Witwatersrand, Lithology and Mineral Resources, v.44(5), p.543-558.

Figure 2: Simple schematic representation of three proposed models for the Au mineralisation of the Witwatersrand Basin. (1) Placer (unmodified) (2) Placer (modified) (3) Hydrothermal-metamorphic. TRD are ‘Re depletion ages’ for Au and give an indication of the minimum age of separation from the mantle. Note the different age constraints for each model (i.e. hydrothermal = younger Au). (Source: Kirk et al., 2001).


A Brief History of South African Mining Written by Andrew (Andy) Clark

Whilst South Africa produces a large variety of minerals, its extraordinary mining history has developed as a result of the discovery and subsequent exhumation of the huge diamond and gold deposits of the Kimberley and Witwatersrand basin. It would appear that South Africa has been occupied for a significant period, with evidence from numerous sites such as Sterkfontein and Blombos caves suggesting modern humans have inhabited the South African region for over 100 000 years. Small mobile hunter gatherer tribes began to adopt a less nomadic lifestyle between Namibia and the Eastern Cape along the coast approximately 2000 years ago as a result of the arrival of Northern agro-pastoralists. Cattle farming settlements became established, along with Iron Age technology brought from the North including the mining and processing of iron, copper, tin and gold. Portuguese sailors visited the area frequently during the early 1500’s. Other Europeans followed later in the 16th century however it was not until the 1800’s that the mineral wealth of the region began to have a large influence on its development. The discovery of alluvial diamonds in the Vaal river in the late 1860’s, followed by dry deposits at the soon to be city of Kimberley, brought huge numbers of black and white settlers to what would become the first industrial hub of Africa. Whilst the British Empire had established a colony on the Southern cape of the continent, it was not until the discovery of significant mineral deposits in the region that tensions between the African inhabitants and the British colonists became increasingly inflamed. Unfortunately, despite the British outlawing of slavery, local farming communities became increasingly wage based due to British demand for workers, leaving workers open to manipulation and abuse from their British overlords, a trend which would continue throughout South Africa’s industrialisation. As a result of the discovery of extensive diamond deposits, the British annexed the areas around Kimberley into the existing colony at the cape of the continent, and vast networks of railway were developed that linked the interior of the region with coastal ports such as modern day Cape Town, Port Elizabeth, East London and Durban.

In the 15 years since the discovery of the Kimberley deposits, South Africa yielded more diamonds than Indian mines from the last 2000 years, with the Kimberley supplying 95% of the world’s diamonds at the time. Despite this huge increase in the global diamond supply, prices did not fall due to the co-incidental depletion of Brazilian deposits and the rapidly increasing wealth of the United States, allowing the continued development of mining interests in the region. It was during this period that English immigrants Cecil John Rhodes and Barney Barnato held a fierce rivalry in Kimberley prospects, resulting in an eventual merge into De Beers Consolidated Mines Ltd, the biggest player in the diamond industry today.

The Kimberley mine in 1873. Square claims at the site had lengths of 31 feet, resulting in a landscape of deep pits and high walls, making mining operations extremely hazardous (Williams, 1906).


The increasing discovery of mineral wealth in the region led to increased imperialist (British) efforts to gain absolute control over the states surrounding their colony containing lucrative mineral deposits. Local republican forces (Particularly the Boers) resisted the British expansion, leading to British withdrawals from previously annexed states such as Transvaal 1881. This is known as the First Boer War. With the subsequent discovery of the Witwatersrand deposits in 1886, British interest in control of the mineral rich areas of the African subcontinent was once again renewed. This was further influenced by mining magnates who lacked confidence in the ability of the Boer government to provide an efficient and stable environment in which the mining industry could be developed. British raids into the Boer territories resulted in the declaration of the Second Boer War in 1899, lasting until 1902, resulting in the death of tens of thousands of African women and children in British camps and the defeat of the Boers at the hands of imperialist forces. Following their victory, the British set about establishing white control of newly occupied land and increasing supplies of wage labour to satisfy the increasing demands of the mines. In 1904, due to the insatiable demand for workers, 60 000 Chinese were brought in to work in the South African mining industry. Interestingly the usual influx of freelance diggers that accompany a gold rush did not occur. The nature of the South African goldfields required the sinking of deep shafts, requiring a large amount of capital. As a result, the entrepreneurs who had developed the lucrative diamond industry a few years earlier were the major players in the new gold industry. The continued growth of the mining industry in South Africa under the control of the British had the effect of depriving locals of the chance to increase their standing in society whilst the mining companies grew increasingly wealthy. More and more black peasants were forced off their farms due to high taxes to work for wages in the mines. The continued development of isolated communal labour camps resulted in the abuse of workers’ rights and a lowering of wages. In addition, black workers were forbidden to move on to skilled labour, resulting in the creation of a permanent black lower class. Whilst enormously rich in natural resources, it could easily be argued that significant inequality still exists in South Africa today, with just one illustration being that only 8.9% of the mineral assets in the country are owned by black people.

Plan of claims at the Kimberley diamond mine from June 30, 1883 (Williams, 1906).


Overview of Namibian Geology Written by Vikraman Selvaraja

Introduction Namibia (formerly South-West Africa) is a coastal nation located in Southern Africa bounded on the south by the Orange River and South Africa, Botswana to the east and Angola to the north as shown in the map of Africa below. It is a large nation, covering an area of 825,418km2 (about the size of NSW) with a very varied geology encompassing a period of nearly 2.6Ga. Most of the country is covered by the Kalahari and Namib deserts making the geological history of the country only visible where there is bedrock exposure to any extent (about half the country.) There are five major episodes in the geological history of Namibia. (Martin, 1965) Firstly, the formation of metamorphic complexes between 2500Ma (Vaalian) to 1800Ma (Early Mokolian). These rocks form the basement on which it is suspected that most of Namibia is built on. The next major rock forming period was between the Middle to the Late Mokolian periods (1800Ma to 1000Ma) in which sedimentary and igneous rocks of various families such as the Gariep System, the Kunene Igneous Complex, Opdam-Kapok-Sinclair Formations, Namaqualand Granite-Gneiss Complex and the Khoabendus-Skumok Formations were emplaced. The final events in this sequence culminated in the formation of the supercontinent Rodinia at about 1Ga. The next major event was the formation of the Damara Orogen which started with intracontinental rifting and sedimentation about 900Ma and culminated in a large orogenic body formed in a continental collision setting between 650-450Ma. This marks the formation of the supercontinent Gondwana. The Gondwanan supercontinent was to remain stable for the next 350 Ma before the final breakup to represent the continental structures we see today. The deposition of the Karoo sequence from sediments sourced from the erosion of the Damara orogen between 300Ma and 135Ma during the largely tectonically stable period was then followed by large amounts of volcanism associated with the breakup of Gondwana and this is fourth phase of rock formation in Namibia. The fifth and final phase is the formation of Cretaceous and Quaternary deposits which cover many of the older rocks (Miller, 1992). Basement Rocks The Namibian basement is comprised of several metamorphic complexes which have undergone several Wilson cycles and are part of the Proterozoic Congo and Kalahari cratons. The oldest rocks in Namibia are from the Hoarusib Valley in the NW of Namibia which have been dated to 2645 Ma (Seth et al. 1998). This is likely to be part of the Epupa gneisses which also lie across the border in Angola. It is largely intruded by the massive anorthosites of the Kunene Complex which comprise of interlayered troctolite and norite. Other major basement groups are the gneisses and amphibolites of the Huab Complex west of Fransfontein and granites with amphibolites dykes make up the Grootfontein Complex in the NE. The Khoabendus-Skumok formation occurs separately in the NW and SW parts of the Kamanjab Inlier and is comprised of a variety of acid and basic volcanic rocks in the lower portion and a sedimentary upper portion which contains quartzite, conglomerate, limestone, dolomite, chert and BIFs. In the Karibib-Usakos area the Abbabis Complex which has banded gneiss as its main rock type dominates. Further south the Hohewarthe Complex in Rehoboth consists largely of a migmatised schist, quartzite and amphibolites with some granite and basalt as part of the overlying Neuhof Formation. Finally in the far south, the Orange River Group represented mainly by the Vioolsdrif Suite of granodiorite and adamellite dominates.


Figure 1: A graphical reconstruction of Namibia’s known tectonic history (modified from Truswell, 1977).


Middle Mokolian to Rodinia The Rehoboth Sequence (1800Ma) unconformably overlies the Elim sequence and is thought to have formed in the back-arc basin of a magmatic arc which extended from South America through Namibia, Zambia, Zimbabwe and Tanzania to Uganda. It is represented mainly by interlayered quartzite, phyllites, conglomerate and volcanics. The Namaqua metamorphic complex in the south covers large areas and contains a range of pre-tectonic meta-sediments, gneisses and amphibolites, syn-tectonic granites and post-tectonic granites. Much of this is directly associated with the magmatic arc lying along the Rodinian coast. The Sinclair Sequence is also associated with this same active margin and is comprised of volcanic rocks interlayered with sandstones and shales. The last part of this is the intrusion of the Gamsberg, Piksteel, Weener and Nubib granites which intruded into much of the Sinclair Sequence and the youngest of which were dated at about 1000Ma. Namibian Most Namibian rocks are from this era where both the Damara Orogen and the Gariep Belt were formed during a complete Wilson cycle. The Nosib group at the base of the Damara Sequence was formed in an intracontinental rift setting and comprises of arkose, quartzite, conglomerate, phyllites and limestone. They have been dated to about 820Ma from association with nearby ignimbrites. The formation of the Swakop group of lithologies comes with continental breakup of Rodinia and contains interbedded schists, BIF’s and basic lavas. The top of this group is the Chuos Formation which is geo-stratigraphically very important as it contains the glacial mixtites which are found throughout the world some 700Ma ago and are seen as the markers for Snowball Earth (Hoffmann et al. 1998). In the NE of the Damara Orogen, the Otavi group was laid down on what appears to be a stable carbonate platform which has evolved to be dolomites, quartzite, phyllite and other minor volcanics. The subduction zone and the formation of the Damara Orogen developed paired metamorphic belts in the NW (high temperature) and in the SE (high pressure). This marks the beginning of the Gondwanan supercontinent and the erosion of the Damara orogen creates the Nama sedimentary group. Karoo Group to Present Long term tectonic stability resulted in the erosion of the Damara and Gariep mountain belts. The resulting peneplains became the sedimentary basins where the Karoo sequence was deposited. It comprises mainly of basal glaciogenic rocks (which correspond to Southern Africa’s position close to the south pole.), which is overlain by shales, mudstones, limestones and other sedimentary rocks which formed as a result of plate tectonics bringing southern Africa into more moderate climatic situations which created a large release of freshwater as the glaciers melted. The last major piece of the puzzle is the Etendeka Basalt in the NW which have an age of 132Ma and can be directly correlated with the Parana Basalts of Brazil and mark the separation of Africa from South America and the start of the breakup of Gondwana. A series of complex intrusions with ages of 137-132Ma occur in a zone extending from the coast near Swakopmund in a NE direction. They are extremely mineralogically complex and contain rhyolite, granophyres, granite, syenite, foyaite, dunite, pyroxenite and carbonatite. Over 60 kimberlite pipes have been recognised in the Gibeon area and they are interpreted to be the result of a hot mantle plume. Isostatic uplift has contributed to the creation of the Kalahari Basin and desert which covers much of the country in sandy deserts.


Geological Map of Namibia illustrating major formations.


The Navachab Gold Mine Written by Matthew Peacock

The Navachab Gold Mine is located approximately 120 km northwest of Windhoek, see Figure 1. The deposit is characterized by a polymetallic Au-Bi-As-Cu-Ag ore assemblage, consisting of pyrrhotite, chalcopyrite, sphalerite, arsenopyrite, bismuth, gold, bismuthinite and bismuth-tellurides. Geology of the Navachab Deposit

Figure 1: Geological map of area around Navachab deposit after Driggel, 2009. The deposit is situated on the northwestern limb of the Karibib dome, a northwest-verging, non-cylindrical anticline. This part of the anticline is near vertical and steeply dipping to the northwest or southeast. The lithologies exposed in the Navachab pit consist of rocks from the Spes Bona, Okawayo and Oberwasser Formations. The regional metamorphic environment is interpreted to be a medium to high temperature low-pressure terrane (Wulff et al. 2010).

Table 1: Lithologies after Driggel, 2009.

Lithologies Description Spes Bona Biotite Schists and interlayered calc-silicate rocks.

Okawayo Base of formation made up of banded calc-silicate rocks of the marble calc-silicate unit (MC unit), which holds the highest-grade gold mineralization. This is overlain by a banded marble and breccia marble.

Oberwasser Biotite schists and calc-silicate rocks. Igneous rocks The metasedimentary rocks have been intruded by lamprophyre, pegmatite and

aplite dykes, which cross cut mineralization but are themselves cut by mineralized veins. Magmatic activity overlaps with mineralization.


Figure 2: Cross-section of Navachab deposit after Driggel, 2009. Mineralization Gold mineralization occurs in two styles. The richest style of mineralization is associated with bedding parallel cigar shaped sulfide lenses in the banded calc-silicate rocks within the MC unit these lenses are approximately 5m high 1.5m wide bodies that form down plunge tens of meters in length. The second style of mineralization is abundant quartz-sulfide veins that crosscut all units and form a 1x3 km wide system. These quartz sulfide veins are northwesterly dipping and southeasterly dipping in orientation. The poles to these indicate that mineralization was coeval with deformation (Dziggel, 2009).


Genetic Model The unusual setting of the deposit within carbonate bearing meta-sediments means that the deposit mineralogically could be classified as a skarn deposit however the Navachab shares many characteristics with orogenic gold deposits including the epigenetic nature of mineralization, the strong structural controls of mineralization, the reducing redox state of the ore fluid, the process of fault-valve action and the relatively low salinity of the ore fluid (Dziggel, 2009). Some features of Navachab that are not typical to orogenic Au deposits are; the polymetallic ore assemblage of Au-Bi-As-Cu-Ag, the geological setting of Navachab being part of the continental arc which is characterized by extremely high geothermal gradients and abundant syn to post tectonic granites (Wulff 2010, Dziggel 2009). The polymetallic ore assemblage is usually related to intrusion related gold systems and is interpreted to be the result of progressive enrichment during crystal fractionation. Haack et al. (1984) found that metapelitic rocks are able to release large amounts of Cu and Bi during higher grade metamorphism and as such, igneous activity is not essential for the enrichment of these elements. Wulff et al (2010) found that there was little evidence at Navachab to support a magmatic fluid source however the overlap of magmatic water and metamorphic water from D and 18O means that a magmatic influence over the ore fluid cannot be completely out ruled, see Figure 3.

Figure 3: Fluid composition in equilibrium with altered and unaltered metasedimentary rocks from Navachab after Wulff et al., 2009. The source of the mineralizing fluid is believed to be a result of the dehydration of a sedimentary protolith. Partial melting resulted in the emplacement of s-type granite plutons that may have had the potential to concentrate metals and affect fluid flow (Wulff et al. 2009). It is hard to truly identify the role of magmatic activity on the Navachab deposit with so little evidence, however the presence of igneous plutons and granitic dykes would have most definitely affected the geothermal gradient of the terrane at the time of deposition 550-540Ma (Wulff et al. 2009) as well as increasing the fluid pressures as well as modifying the regional fluid flow patterns and as such magmatic activity must have played a role to an extent with the formation of the Navachab deposit.


Figure 4: Aerial image of the Navachab operation, Google 2012.

Figure 5: Cross-section through the Navachab deposit after Attridge (1991).

Figure 6: Neo-Proterozoic Au deposits and Au-prospective terranes after Steven (2010).


The Rössing Uranium Mine Written by Amy (Tiffany) Halcon

Introduction The Rössing uranium mine in Namibia is the only currently operating uranium mine in the country. However, Namibia is a uranium province and primary igneous occurrences have been discovered at Rössing, Rössing mountain, Goanikontes, Valencia and Ida Dome. The uranium was founded in 1928 in the Namib Desert and is now being exploited by Rio Tinto since 1978. Rössing is now a world-class uranium deposit, which has both primary and secondary uranium minerals. The mine is located in the Namib Desert near the town of Arandis, approximately 70km northeast from the coastal town of Swakopmund. The licensed area of the mine site is approximately 180km2 and 25km2 of which is used for waste disposal.

The above map indicates the location of Rössing. Map taken from the Rössing website (www.rossing.com) The Rössing formation lies over the Khan formation, which is part of the Nosib group. In the area of the Rössing mine the units are divided into a lower serpentinitic marble, which is a metapelitic gneiss and an upper siliceous and serpentinitic metacarbonate unit, that is interbedded with granofelsic/schistose layers, these units are then followed by metapelitic gneiss subunits.


Table taken from Basson and Greenway, 2004. The Rössing uranium deposit is an example of an intrusive alaskite. In the main pit of the Rössing mine is composed of leucogranite sheets. The deposit has been highly deformed thus forming sheeted leucogranites. Nex et al. (2001) subdivided the leucogranite sheets into six distinct types and were named type A-F. The subdivision was done purely on detailed fieldwork and their observations. The subdivision is based on observable field characteristics of structural setting, form of intrusion, colour, grain size, texture, radiometric measurements and macro-scale mineralogy. The following table outlines each type:

Table taken from Nex et al., 2001

The Etusis Formation locally consists of a LowerBiotite Schist Member, a Lower Metaquartzite Member,an Upper Biotite Schist Member and an Upper Meta-quartzite Member (Nash, 1971; Lehtonen et al., 1996;Oliver and Kinnaird, 1996; Table 1). The Upper BiotiteSchist and Upper Metaquartzite Members are truncatedand overlain by the intraformational R!ossing SJ ShearZone (Oliver, 1994, 1995). The gneisses of the KhanFormation (Nosib Group), comprising locally bandedmigmatitic lithotypes that reflect a protracted tectono-thermal history, gradationally overlie (Nash, 1971;Berning et al., 1976) or interfinger (Downing, 1983) withthe Etusis Formation. The Khan Formation consistsmainly of amphibole–clinopyroxene gneisses, inter-preted to reflect the change to a more calcareous and lessclastic sedimentary protolith (Downing, 1983; Martin,1983). The uppermost portion of the Khan Formationshows increased amphibole in place of pyroxene, cul-minating in a thin unit of amphibole–biotite schist withminor, discontinuous pebble horizons (Nash, 1971;Table 1).

The R!ossing Formation paraconformably and dis-conformably overlies the Khan Formation. In thevicinity of the R!ossing Mine it is divided into a lowerserpentinitic marble, a metapelitic gneiss and an uppersiliceous and serpentinitic metacarbonate unit inter-bedded with granofelsic/schistose layers, which issucceeded by metapelitic gneiss subunits (Nash, 1971).

4. Regional geophysical lineaments

Lineaments in the Damara Orogen are defined by thealignment of prominent negative magnetic anomaliesrelated to layers within shallow suboutcropping oroutcropping Nosib Group metasedimentary units(Corner, 1982, 1983, 2000; Eberle et al., 1995), changesin the orientation of kinematic fabrics and by structuralcorridors (c.f. O’Driscoll, 1981, 1986). Geophysicalmodeling by Corner (1982, 1983, 2000) found that manymagnetic lineaments are late-kinematic geanticlinal rid-ges dividing the NE-trending intracontinental branch of

Table 1Lithostratigraphy of the Etusis, Khan and R!ossing Formations in the vicinity of R!ossing

Group and nature Formation Member/unit Thickness (m) Description

Swakop(pelitic and calcareous)

R!ossing Metaquartzite >100 Medium-grained quartzite, coarseningtowards base

Upper metapelitic gneiss 40–50 Diopside-biotite (-scapolite) gneiss; thin lensesof coarse-grained metaquartzite and marble;grading upwards into cordieritegneiss

Upper marble 50–70 Serpentinitic and diopside-quartz-bearingmarble; lenses of biotite-diopside granofelsand cordierite-biotite gneiss. Basal con-glomerate

Lower metapelitic gneiss 30–40 Cordierite-biotite-sillimanite gneiss; gradingupwards into biotite-hornblende schist

Lower marble 20–50 Dominant serpentinitic and minor graphiticmarble; lenses of biotite-diopside granofels andpelitic schist

Nosib (fluviatile) Khan Amphibolite 10–20 Amphibole–biotite schist with relict pebblebands or conglomerate at its base

Upper gneiss 70–100 Amphibole- and pyroxene-gneissPyroxene-garnet ortho-amphibolite gneiss

<120 Pyroxene-garnet gneiss and pods of horn-blende-oligoclase ortho-amphibolite, oftenmigmatized

Lower gneiss 70–150 Strongly banded clinopyroxene-amphibolegneiss; migmatized in high-strain zones

Etusis Upper metaquartzite >300 Arkosic and micaceous gneisses, feldspathicquartzites and biotite schist; extensively in-truded by late- to post-kinematic granites andmigmatized in high-strain zones

Upper biotite schistLower metaquartziteLower biotite schist

Abbabis metamorphiccomplex/basement

– – Augen-, migmatitic-, biotite-, silimanite- gran-ite-gneiss; biotite schist and amphibolite

The quoted thicknesses are approximate. Many of the lithologies are partially or pervasively migmatized. Summarized after Nash (1971), SACS(1980), Coward (1983), Downing (1983), Martin (1983), Lehtonen et al. (1996) and Nex (1997).

416 I.J. Basson, G. Greenway / Journal of African Earth Sciences 38 (2004) 413–435

Type B SLGs are inhomogeneous with a variablegrain size from fine- to pegmatitic and form parallelsided tabular intrusions. They are white in colour andare often boudinaged and folded by D3 folds (Fig. 5).The presence of garnet is diagnostic and they frequentlycontain biotite and tourmaline. They have been a!ectedby the high-temperature post-peak metamorphic ther-mal overprint as they contain sillimanite which in placesreplaces biotite.

Type C SLGs vary in colour from pink to white, areirregular in shape and occur in flexures associated withD3 folds. They are the most abundant sheet type withinthe Goanikontes area and are 0.5–10 m in width withtwo feldspars and interstitial quartz.

Type D SLGs are irregular and anastomosing andhave a medium to coarse grain size. They are white incolour, always contain smoky quartz with minor betafitewhich is readily recognisable by its dodecahedral form,surrounded by radial cracks and often have secondarybeta-uranophane on joint surfaces. In addition theyfrequently contain topaz that is absent from other types.

Type E’s are characterised by the presence of ubiq-uitous ‘‘oxidation haloes’’ (Corner and Henthorn, 1978)(Fig. 6). These haloes are irregular and frequently ex-hibit a reddened rim with grey cores although di!erenttypes may occur within the same sheet (Fig. 7). Adjacentor cross-cutting sheets of di!erent SLG type, less than ametre away do not contain these haloes. Within thesheets there is a wide range in modal mineralogy (Fig. 4)and the only generalisations that can be made are thatred cores contain abundant opaque minerals and littlebiotite whereas the grey cores contain abundant biotite,

few opaques and sericitised plagioclase with albitic rims.There was no correlation between mineralogy eitherwithin or between sheets relative to corona type andcolour. In the past these haloes have been interpreted asskialiths of Khan Formation metasediments (Barnesand Hambleton-Jones, 1978) however all parts of a typeE sheet are granitic and there is no petrological evidenceof remnant country rock.

Type F SLGs are narrow (0.5–3 m) tabular straight-sided bodies which cross-cut all other types of SLG.They are brick-red in colour and contain large perthiticK-feldspars up to 30 cm in size, milky quartz and in-terstitial biotite; they are almost albite-free.

Two types of these sheeted leucogranites containappreciable uranium mineralisation, namely D and E.It is significant that D3 deformation, the main domeforming event, also sub-divides the SLGs into twogroups, those that are a!ected by D3 (A, B, and C) andthose that intruded post-D3 (types D, E, and F). Inaddition a fundamental observation is that these sheetedgranites cross-cut all lithologies including basementgneisses and therefore they cannot be derived frompartial melting of the metasedimentary Damaran Se-quence.

4. Geochemistry

Sampling procedures of the granitic rocks involvedcollection of at least three individual samples from eachsheet or outcrop each of at least 10 kg of material inorder to produce representative analyses. Samples were

Table 3Field characteristics of SLGs from Goanikontes with average scintillometer counts per second in the first column. Numbers in brackets are themaximum scintillometer counts observed for the U-enhanced sheets

Type Width (m) Diagnostic structural features Diagnostic mineralogical features

A, <20 cps <0.75 Infrequent occurrence, irregular form, weakfoliation, boundinaged and folded by D3,only occurs within the high strain zone

Pale pink, fine–medium grain size,homogeneous saccharoidal texture, weakfoliation

B, <20 cps 1–4 Common outside the high strain zone, finegrain size sheets are weakly foliated, fre-quently boundinaged and occasionallyfolded by D3

White, fine-pegmatitic grain size, typicallygarnetiferous, infrequent abundant biotiteand tourmaline

C, 10–20 (200) cps 0.5–10 Most frequent type of SLG within the typicalcover sequence, occasionally boudinaged,occurs in F3 fold flexures

Pale pink-cream, medium-pegmatitic grainsize, hypersolvus with interstitial clearquartz, magnetite, ilmenite and tourmaline

D, 100 (400) cps 1–7 Irregular and anastomosing, restricted to thehigh strain zone and the Khan–R!oossingboundary

White, medium-coarse grain size, granulartexture, white feldspar with characteristicsmoky quartz, frequently visiblebeta-uranophane and occasional betafite

E, 30 (300) cps 1–10 The dominant type of SLG within the highstrain zone. Generally tabular, occasionallybifurcating generally emplaced parallel to theprominent gneissosity

Extremely variable colour and grain size,contains ‘‘oxidation haloes’’ (Corner andHenthorn, 1978)

F, <20 cps 0.5–3 Tabular with straight parallel sides,occurring throughout the area, cross-cuts allstructural features

Distinctive red colour, coarse-pegmatiticgrain size, pink perthitic feldspar and milkycoloured quartz

486 P.A.M. Nex et al. / African Earth Sciences 33 (2001) 481–502


Mineralisation The high concentration of uranium mineralisation is limited to the sheeted leucogranites of the D type. The main primary uranium mineral from the Rössing formation is magmatic uraninite, this provides 55% of the ore mineral. The uraninite is within the quartz and feldspar. Due to hydrothermal or weathering processes, the secondary uranium mineralisation is in the form of uranophane, beta-uranophane, gummite, torbernite/metatorbenite, carnotite, metahawaiweeite and thorogummite. The most abundant of the secondary minerals is beta-uranophane which accounts for 40% of mineralisation. Such a high degree of secondary enrichment has not been found to occur in uranium occurrences in the Central Zone. This enrichment thus distinguishes Rössing form its surrounding areas and contributes to its economic value. It is thought that the concentration of the uranium in the sheeted leucogranites is due to magmatic processes rather than hydrothermal activity or even fluid flow on a regional scale. The reaction of the magma with the marbles of the Rössing formation is thought to have lead to the boiling of the magma and thus was important in the mineralisation process of the uranium. However at the Rössing mine, in the main pit, the mineralisation is associated with high amounts of fluids and high amounts of water from fluid extraction studies. This may be the reason for the area reaching economic proportions of uranium mineralisation and the surrounding areas unable to attain such high concentrations of uranium.


Road Map of the Southern Damara Orogen: Navachab to the Coast Compiled by Dominique Tanner


Snowyʼs Road Guides Written by Anna (Snowy) Haiblen

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Figure 1: The Paraná-Etendeka LIP, including Spitzkuppe (lower right). The rifting of South America and Africa opened the south Atlantic ocean and was closely associated with the Tristan mantle plume. Image from Haapala et al. (2007).

Figure 2: The local geology of the Spitzkuppe formations, showing the prevalence of dykes across many kilometers. Image from Thompson et al. (2007).

Grosse and Klein Spitzkuppe Written by Michaela Flanigan

Introduction The highly-evolved, approximately 125 Ma Grosse and Klein Spitzkuppe granites are the youngest rocks in the Paraná-Etendeka Large Igneous Province (LIP) that spans the Atlantic Ocean between Namibia and Brazil (Thompson et al., 2007). This article will describe the rocks found at Grosse and Klein Spitzkuppe and present likely models of formation, linking these to the larger processes occurring in southern Africa at the time of emplacement. The Paraná-Etendeka LIP lies partially in the Damara orogen (see Figure 1), though it formed some 350 Ma later (around 130 Ma). At this time, the southern Atlantic ocean was being formed by the rifting of the Western part of Gondwanaland, and the Tristan mantle plume was upwelling beneath areas of Brazil and Namibia that are now separated by the ocean. These processes formed the Paraná-Etendeka LIP, which contains substantial bi-modal volcanism, with both mafic and felsic dykes and magma emplacements occurring. The area surrounding Spitzkuppe contains large swarms of both mafic and felsic dykes (see Figure 2), with some mafic dykes cutting the granite formations (Frindt et al., 2004). The Spitzkuppe granites, like many of the other felsic rocks associated with the Paraná-Etendeka LIP, are A-type granites, and have been the focus of studies

into the origin of A-type granites in general. Models for the formation of A-type granites range from pure fractionation from a more mafic magma, to partial melting of a previously-melted crustal source, or melting of a relatively felsic source rock, such as tonalite (Frindt et al. 2004).


The Grosse Spitzkuppe granite contains miarolitic cavities (cavities formed by the exclusion of volatiles from a fluid-saturated crystallising magma), and from these, gem-quality topaz and beryl have been mined by local populations for some time (Haapala et al. 2007). These minerals appear as accessory phases in the granite bodies, but are found in large, euhedral crystals in pegmatitic zones. Description The Grosse Spitzkuppe stock outcrops around 700m above the surrounding ground, while Klein Spitzkuppe stands around 600m higher than the surrounding area (Haapala et al., 2007). Both have been subject to prolonged exposure, causing jointing along planes defined by the ancient cooling of the magma body from outside to inside. This causes the distinctive ‘onion-skin’ weathering pattern that results in rounded granite boulders (Thompson et al., 2007).

Figure 3: Structural and textural features in the Grosse Spitzkuppe and Klein Spitzkuppe granite stocks. (a) Mingling and hybridization textures between porphyritic granite and mafic enclaves (upper margin of the picture) at the northeastern margin of the Grosse Spitzkuppe stock. The alkali feldspar phenocrysts are corroded and mantled by plagioclase shells. (b) Large miarolitic cavity in porphyritic granite, Grosse Spitzkuppe stock. Such druses contain gem-quality topaz and beryl crystals. (c) Layered aplite stockscheider at the margin of the Grosse Spitzkuppe stock. Growth direction downward (see Frindt and Haapala, 2004). (d) Layered pegmatite stockscheider dike at the southeastern margin of the Klein Spitzkuppe stock.


Grosse Spitzkuppe is divided into three main textural zones which are geochemically very similar (see Figure 3 – image from Haapala et al. (2007)). There is an outer medium-grained zone containing globules and globule swarms of mafic enclaves, followed by the most prevalent coarse-grained granite, which contains biotite schlieren that illustrate flow and fractionation in the cooling magma chamber. The central and eastern portion is porphyritic, and contains abundant miarolitic cavities. Klein Spitzkuppe contains only two textural zones; a coarse-grained area that forms the mountain, and a surrounding medium-grained area. Contacts between all zones vary between gradational and sharp, and there are pegmatite zones in both outcrops (Haapala et al. 2007). The mineralogy of the granites is dominated by K-feldspar, with plagioclase and quartz making up the other main minerals. Biotite is common, with up to 8wt.% in parts of Grosse Spitzkuppe. Topaz and fluorite are major accessory minerals, and the stocks also include minor zircon, columbite, magnetite, monazite, apatite, Nb-rich rutile and allanite (ibid; Frindt et al. 2004). Geochemically, the Grosse and Klein Spitzkuppe stocks are similar. Both are slightly peraluminous, with an average A/CNK of around 1.03. They are clearly highly evolved, with Klein Spitzkuppe marginally more evolved than Grosse Spitzkuppe; both have high SiO2 contents (around 76 wt.%) and Na2O+K2O (around 8.4wt.%, with K2O dominant). Their high degree of evolution can also be seen in their high Fe/Mg ratio and enrichment in incompatible elements such as Rb, Ga and F, and their paucity of TiO2 and Sr. The relatively more evolved status of Klein Spitzkuppe is evident in its depletion of REEs and compared to Grosse Spitzkuppe, and its larger Eu anomaly (Haapala et al., 2007). Formation The formation of the Spitzkuppe granites is an unresolved question, with several plausible theories existing. Haapala et al. (2007) used isotopic studies to demonstrate that the granites could be produced from fractionation of the material found in the felsic dyke swarms in the area, but the question of where that material originated remains. The granites, like others in the Paraná-Etendeka LIP, appear to have both crustal and mantle input; the bulk rock composition is extremely felsic and could not have evolved from mantle material on its own, yet the presence of mafic enclaves and geochemical considerations demonstrate that the granite is not purely a crustal melt. The geochemical evidence in question is the low εNd (around -6) and 87Sr/86Sr (around 0.711), which are closer to mantle values than could be believed for a crustal source (Thompson et al., 2007). For illustration of Spitzkuppe’s relative Nd and Sr ratios, see Figure 4 (opposite). The most likely origin of the melt that formed the Spitzkuppe granites is the melting of a deep crustal source by magmatic underplating. In this model, a heat source is provided by the increased geothermal gradient beneath the thinning crust, and by the Tristan mantle plume pooling beneath the crust. Mantle plumes can have temperatures far higher than the solidus temperature of crustal rocks (consider MORBs, generally cooler than plumes, at over 1000ºC, compared to the granite solidus around 700ºC), and a relatively high degree of melting can thus be achieved when they come in contact.


Figure 4: The mantle-like characteristics of the Spitzkuppe granites (top left) in relation to other rocks in the region. The underplating could have melted:

• lower or mid-crustal rocks that had previously had melt extracted; • an undepleted mafic zone which subsequently incorporated other crustal rocks when new

injections of hot magma arrived, or • lithospheric mantle heavily metasomatised by fluids from a degassing mantle under the rift

zone, and associated with the plume. The melt produced in this manner subsequently underwent fractionation as it migrated upwards, as evidenced by the discrepancy between the feeder dyke composition and the large granite bodies. Distinguishing between these models for the formation of Spitzkuppe is difficult; there is clearly mixing of magmas occurring, but in order to determine how much mixing occurred, the magmas need to be identified, and identifying the magmas without any end-members relies on knowing the degree of mixing. The plume magmas may be one candidate for an end-member, but unlike other formations in the Paraná-Etendeka LIP, the Spitzkuppe granites do not show any trace element evidence of plume contribution. Perhaps we will solve this dilemma on our trip – 20 undergrads is a lot of brainpower, after all. References Frindt, S., Trumbull, R. and Romer, R. (2004) Petrogenesis of the Grosse Spitzkuppe topaz granite,

central western Namibia. Chemical Geology 206, 43-71. Haapala, I., Frindt, S. and Kandara, J. (2007) Cretaceous Grosse Spitzkuppe and Klein Spitzkuppe

stocks in Namibia. Lithos 97, 174-192. Schneider, G. (2008) The roadside geology of Namibia. Gebruder Borntraeger, Stuttgart. Thompson, R. et al. (2007) Origin of CFB Magmatism: Multi-tiered intracrustal picrite-rhyolite

magmatic plumbing at Spitzkuppe, western Namibia Journal of Petrology 48, 1119-1154.


The Uis Sn-Ta Pegmatites Written by Morgan Williams

Locality and Regional Geology Uis is a small village (downgraded to a 'declared settlement area' in 2010) located in the Erongo Region of central Namibia. The village of approximately 3,600 people owns 10km2 of land, and is known almost solely for its tin mine. The town is situated on the Cape Cross-Uis pegmatite belt. The 120km long Cape Cross-Uis belt trends north-east from Cape Cross, and represents a 24km wide half-graben system. The belt is primarily composed of foliated metasedimentary rocks of Pan-African age (Damara Sequence, late Precambrian), into which the granites and pegmatites intrude. The Pan-African granites of the region consist of syn-tectonic porphyritic granodiorites, syn to late-tectonic biotite granites and late to post-tectonic mica granites. The Nosib Group represents the base of the Damara Sequence in this region. The cordierite-amphibolite facies unit is comprised of primarily quartz-feldspar gneisses, amphibolites, biotite schists, calc-silicates and meta-rhyolites. The Swakop Group (600-700 Ma) turbiditic meta-greywackes, biotite schists, quartzites, calc-silicates and minor tourmalinite overly the Nosib Group. To the south-east of Uis, the Autseib Fault downthrows the greenschist-facies Amis River Formation (Upper Swakop Group) against the marbles, calc-silicates and biotite schists of the Rossing, Overwasser, Karibib and Kuiseb Formations (Lower Swakop Group). Post-tectonic rare metal pegmatites can contain economic concentrations of tin, niobium, tantalum and lithium. These pegmatites are generally limited to schistose meta-sediments towards the centre of the half-graben system, where the east-west striking and northerly dipping shear zones acted as traps for pegmatitic melts and fluids. Mineralogy and Geochemistry The group of pegmatites can be divided into three swarms: the Uis, Karlowa and Strathmore swarms. Furthermore, the pegmatites of the region can be broadly grouped into two categories:

undifferentiated, homogeneous quatzo-feldspathic pegmatites that lack rare metal mineralisation;

the rare metal pegmatites. The rare metal pegmatites of the Cape Cross-Uis pegmatite belt have been subdivided by Diehl (1993) as follows:

Unzoned tin-bearing (cassiterite) Uis-type pegmatites; Zoned Li, Nb, Ta and Be-bearing pegmatites;

o These occur within the Karlowa and Strathmore swarms. o Characteristic zones of metasomatic replacement of pre-existing pegmatitic material

are present. Minerals introduced by the influx of volatile-rich fluids (either with the fluid or though breakdown of primary assemblages) include: Nb-Ta oxides, cassiterite, Li-silicates and phosphates, as well as minor REE minerals.

Gem-quality tourmaline and beryl bearing pegmatites; o These are relatively rare in the Cross Cape-Uis belt, in contrast to other Namibian

pegmatite belts. Diehl suggests that mineralisation is partly controlled by the nearby post-tectonic alaskites (K-feldspar-quartz leucogranites).


The Uis-Type Pegmatites These pegmatites are arguably the most economically significant within the belt, and commonly contain concentrations of tin greater than 800ppm (varying from 200-1600 ppm, compared to the average crustal abundance of ~2.2ppm, and around 20ppm in granites) alongside up to 440ppm niobium-tantalum. Uis-type pegmatites are relatively unzoned coarse-grained quartzo-feldspathic deposits, are mica-bearing and host cassiterite as the primary tin ore. Under Cerny's classification (Cerny, 2005 and Singh, 2007) the pegmatites of the Uis belt are within the peraluminous LCT family, with prominent accumulation of Li, Cs and Ta, which are commonly derived from S-type granites. The Uis swarm contains approximately 120 pegmatites, striking north-easterly to easterly, dipping between 30 and 70 degrees north-west. At it's centre (close to the Uis tin mine), some of the pegmatites reach dimensions of up to 1km in length by 50m wide. The pegmatites and aplites of the region intruded biotite and knotted schists between 550Ma and 460Ma (Diehl suggests a value of 490Ma). Cassiterite is concentrated in replacement zones (albite-muscovite rich), and is also present with fractures in albite-quartz-muscovite rich areas. Petrogenesis and Placement The pegmatite placement is likely to be controlled by the presence of 'tension gashes' (sinistral Riedel shears, as s-shaped en echelon fractures) within the Amis River Formation schists. In the Uis region, aplite and pegmatites are found to be mainly parallel to the strike of the meta-sediments, but can also be found to cut across them. Within the Damara pegmatite province, increasing muscovite ages towards the coast (interpreted as higher metamorphic grade) can be correlated with increasing Li, Nb, Ta concentration within the rare metal pegmatites. The pegmatite ages obtained indicate that they seem to be intruding during the peak of regional metamorphism of the north-western section of the Damara Orogen.

Rb/Sr fractionation diagram for Cape Cross-Uis belt granites (left) and rare metal pegmatites (from Diehl, 1993).

The Rb/Sr data from the rare metal pegmatites throughout the belt collected by Diehl seems to indicate that the original magma from which the pegmatites were generated formed during the peak of regional metamorphism, and hence pre-dates the intrusion of the post-tectonic leucogranites. This refutes earlier hypotheses for pegmatite genetic relation to post-tectonic granites (of the same or similar emplacement ages).


Fig. 1 from Diehl (1993). The three swarms of pegmatites (Strathmore, Karlowa and Uis) are shown, along with the relevant geology of the region.

Tin-bearing pegmatite belts of central Namibia. (From the Geological Survey of Namibia's Rare Metals Map).


The granites within the Cape Cross-Uis belt follow a fractionation trend from early granodioritic material towards highly fractionated post-tectonic leucogranites. However, the Uis-type pegmatites follow a differing trend, from less-fractionated pegmatites towards highly fractionated (high Sn, Rb with low Sr, Ba) rare metal pegmatites. Diehl has also investigated strontium isotope ratios, and suggests an anatectic origin for the magma, potentially from re-melting of basement or early Damaran meta-sediments. Deep-seated lineaments along early graben faults would favour the rise of melts along the distinct belts in which they are found, trapped within the large Riedel shears of the Central graben area. Rare metal pegmatites in Nigeria (occurring in a similar geological environment) have been found to be derived from basement material, rather than the Pan-African granites. References A. Hartman (2010). Town regrading a sad move. The Namibian, 27th August 2010. http://www.namibian.com.na/news-articles/national/full-story/archive/2010/august/article/town-regrading-a-sad-move/ P. Cerny and T Ercit (2005). The Classification of Granitic Pegmatites Revisited. The Canadian Mineralogist, Vol. 43, pp 2005-2026. M. Diehl (1993). Rare metal pegmatites of the Cape Cross-Uis pegmatite belt, Namibia: geology, mineralisation, rubidium-strontium characteristics and petrogenesis. Journal of African Earth Sciences, Vol. 17 No. 2, pp. 167-181. Namibia-1on1.com (2009). “Uis”. http://www.namibia-1on1.com/uis.html Singh (2007). Tantalite exploration in “Block-A” of Uis region, Namibia. Trabajos de Geologia, Univ. de Oviedo, Vol. 27, pp 41-69.


The Damara Orogen Written by Temma Carruthers-Taylor

Introduction The Damara Orogen is a Neoproterozoic (1,000 to 542 +/- 1.0 million years old) fold and thrust belt that is a triple collision junction, with two to three coastal arms north and south extensions (the Kaoko and Gariep Belts) and 3) an Inland belt (Damara Belt). The Inland Branch covers the central part of Namibia, falling slightly short of the Botswana-Angola border, therefore this Belt and particularly the orogeny processes and what evidence encountered on the field trip will be the focus. The Damara Orogen formed during the assembly of Gondwana, and now links two surrounding cratons (Congo and Kalahari Cratons), intra-continental rifting happening from 780 to 746 Ma. The majority of the Orogeny spanned from 550 to 490 Ma (Pan-African age), but the events before this time are important. Damara Orogeny formation timeline (Gray et al, 2008):

• ~560-550 Ma – Subduction and oceans present o Syn-tectonic syenites, diorites and granites. Turbidite sedimentation and

subduction in the Khomas Ocean. o Magmatism in the Central zone from 570-540 Ma.

• 540-505 Ma – closure of ocean basins o High-grade metamorphism (530-505 Ma), peak deformation, granitic magmatism o continent-continent collision of the Kalahari and Congo Cratons happening in the

Lower to Middle Cambrian (540 to 505 Ma), forming thrust faults in the Southern Zone and granitoid domes in the Central Zones.

o SE-verging deformation between 520-508 Ma in the southern Central Zone o Magmatism is apparent in the Northern Zone from 520-509 Ma and in the Central

Zone form 535- 500 Ma • ~500-495 Ma – cooling and isostatic adjustment, activity on major shear zones and thrusts

continue o Continued thrusting, emplacement of Naukluft nappes (Southern Zone), A-type

granites. o Cooling and exhumation happened in the Southern and Southern Margin Zones from

502-470 Ma and in the Northern and Central Zones from 490-470 Ma.

Figure 1: Cross section of the Damara Belt, Gray et al, 2008.

,


Table 1: Zones and the related metamorphism and features

The Matchless Amphibolite The Matchless Amphibolite is situated around Windhoek, which we may cross as the trip heads to Navachab Au mine, although different maps show slight variations and infrastructure may interfere. The Matchless Amphibolite is amphibolite interlayered with quartz-mica schist and extends from the Kuiseb River (60km from the coast) to approximately 300km inland so there are other chances on this trip where we could see it, like at the Matchless Mine toward the end of the trip. The amphibolites are metamorphosed tholeiitic basaltic rocks which formed in a time of crustal extension; the east and western regions differ from basalts having oceanic qualities and within-plate features (Breitkoft et al, 1988). The Southern Zone As we drive over the Southern Zone of the Damara Belt, starting in the south going to the northern area (while heading to Navachab) the road cuttings (and hopefully some stops) will prove that the metamorphic grade increases as we head north, seeing a facies change from greenschist to amphibolite. We may also see the Okahandja Shear or Lineament Zone (OL) as we have to travel through Okahandja before heading to the coast. This zone is a significant feature of the Southern Zone representing 750 and 520 Ma southern margin of a magmatic arc produced by the northerly subduction of the Damaran Ocean. It composes of granitoid intrusives and sediments of the Orogen (Downing et al, 1981). As we head from Okahandja we cross over the Pan-African granites, the post-tectonic and syn-post tectonic granites that have an age range from 494 to 484 Ma, interlayered with a turbidite facies, the Swakop Group. These may be studied in road cuttings.

Zone Metamorphic conditions

Major features Facies/rock types

Southern 600oC and 8 kbar. Low-T, high-P

Metamorphic grade increasing from south to north. Contains the Matchless Amphibolite and thrust faults.

Greenschist to amphibolite facies

Central 700-750oC and 5-6 kbar. High-T, low-P Medium to high grade

Increased metamorphic grade from east to west, high-grade conditions reached from partial melting (Jung, 2003). Intense deformation, pre-Damaran basement exposed. Dome structures.

Granite dominated. Late stage aplite dykes present. Lower amphibolite facies in the south to upper-amphibolite and granulite facies south west of S CZ

Northern margin

Intense deformation without inclusion of basement (Goscombe et al, 2003).

Distal turbidite sediments – greenschist facies (Goscombe et al, 2003)

Northern 430-530oC and 2-3 kbar (Miller, 1983 in Jung, 2003). Eastern peak: 635oC and 8.7 kbar

Low P contact metamorphism. Anticlockwise P-T paths in the west, higher P (Barrovian) metamorphism. Clockwise P-T paths – in the east where deep high-P/moderate-T Barrovian metamorphism burial.

Mid-amphibolite facies (Goscombe et al, 2003) to greenschist facies. Post-kinematic granites. Late stage aplite dikes.

N


Figure 2: Points of interest. 1 = partial melting due to: muscovite+plagioclase+quartz+H2O <=> melt+sillimanite. 2= K-feldspar+cordierite inset. Ol = Okahandja Lineament, OML = Omaruru

Lineament, AF = Auteseib Fault, ST = Sesfontain Thrust. Central Zone On the road to Swakopmund we cross the Southern Central Zone where the SE trending deformation may be evident. Exiting Karibib there are opportunities to see huge (kilometre-scale) dome structures, in particular the Karibib Dome which is just south west of Karibib (probably on the way to the Navachab Mine), where the basement gneisses and lower formations of the Damara sequence is seen. The Central Zone (CZ) is highly deformed with striking dome structures (formed in the Neoproterozoic/Cambrian due to collision) which are NW-trending, kilometre scale and often have a core of gneiss basements (Kisters et al, 2004). The Karibib Dome and surrounding domes are NW-verging instead of SW-verging suggested throughout the south-western CZ (Longridge et al, 2011). The response variation of crustal shortening explains differences in the dome lithologies, where ductile mid-crust in highest grade SW CZ went through tectonic escape and thrusting and crustal thickening – lower grade- is seen near Karibib.

Navachab

Uis

Otjiheanamaperero

Matchless Matchless Amphibolite

South-‐north increase of metamorphic grade

OL

Central Zone

Sheath folds prominent

AF

East to west increase of metamorphic grade (central zone)

Karibib Dome

Northern Zone

Southern Zone

OML

Partial melting

ST


N

As we travel west, still in the Central Zone, there might be hints of an increase in metamorphic grade, from lower amphibolite facies in the east to upper-amphibolite and granulite facies to the west near Swakopmund. On this same road we cross two isograds. The right purple line (1) represents (approximate position) partial melting due to: muscovite+plagioclase+quartz+H2O <=> melt+sillimanite. The second (2) is a K-feldspar+cordierite inset (Jung, 2005). We may not be able to see this as it is uncertain that it runs near the road. The partial melting line due to biotite+K-feldspar+plagioclase+quartz+cordierite <=> melt+garnet is crossed less than 50 km from Swakopmund (see the left purple line). Hopefully this is represented in road cuttings (Jung, 2005). Sheath folds are also prominent in the northern CZ and the southern Northern Zone (Downing et al, 1981) so there may be exposed sites along the road to Uis. References Weckmann, U., Olivier, R., and Haak, V., 2003, A magnetotelluric study of the Damara Belt in Namibia: 2. MT phases over 90° reveal the internal structure of the Waterberg Fault/Omaruru Lineament, Physics of the Earth and Planetary Interiors, v. 138, issue 3, pp. 91-112. Breitkoft, J.H., and Maiden, K.J., 1988, Tectonic setting of the Matchless Belt pyritic copper deposits, Namibia, Economic Geology, v. 83, no 4, pp 710-723. Gray, D.R., Foster., D.A., Meert J.G., Goscombe, B.D., Armstrong R., Trouw, R.A.J., and Passchier, C.W., 2008., A Damara orogen perspective on the assembly of southwestern Gondwana, Geological Society, London, Special Publications, v. 294; pp 257-278. Goscombe, B.D., Gray, D., and Hand, M., 2004, Variation in Metamorphic Style along the Northern Margin of the Damara Orogen, Namibia, Journal of Petrology, v. 45, no 6, pp 1261-1295. Longridge, L., Gibson, R.L., Kinnaird, J.A., and Armstrong R., 2011, Constraining the timing if deformation in the southwestern Central Zone of the Damara Belt, Namibia, , Geological Society, London, Special Publications, v. 357, pp 107-135. Jung, S., 2005, Isotopic equilibrium/disequilibrium in granites, metasedimentary rocks and migmatites (Damara orogen, Namibia) – a consequence of polymetamorphism and melting, Lithos, v. 84, pp 168-184. Downing, K.N., and Coward, M.P., 1981, The Okahandja Lineament and its Significance for Damaran Tectonics in Namibia, Geologishe Rundschau, v. 70, no 3, pp 972-1000. Kisters, A.F.M., Jordaan, S.L., and Neumaier, K., 2004, Thrust-related dome structures in the Karibib district and the origin of orthogonal fabric domains in the south Central Zone of the Pan-African Damara belt, Namibia, Precambrian Research, v. 133, pp 283-303.

Figure 3: The eastern Northern Zone, from Goscombe et al., 2004 and Miller and Grote, 1988.


Gondwana Break-up and Large Igneous Provinces (LIPs) Written by Oliver Nebel

In paleogeography, Gondwana (originally Gondwanaland) was the southernmost of two supercontinents (the other being Laurasia) that have been part of the Pangaea supercontinent. Gondwana survived independently after the breakup and existed from ~510 to ~180 Ma. Gondwana is believed to have sutured between ca. 570 and 510 Ma, thus joining East Gondwana to West Gondwana, comprising e.g., Antartica, Africa, Australia and India (Fig. 1). Gondwana reconstruction is based on numerous lines of evidence, such as faunal distributions on present day separated continents; contemporaneous magmatism, either plume-related or along a common active continental margin; sea floor spreading between continents and associated reversals of the planet’s magnetic field (with a mirroring image on both sides of the ridge), and plate tectonic reconstructions based on pole wander paths and seafloor spreading (relative to plume reference frames). Fig. 1: present day continents that formed Gondwana; http://sawhalecentre.com.au/evolution_of_whales. html It is believed by some that super-continental assembly and breakup is a continuous processes throughout Earth’s history that repeatedly occurs every 200-500 Ma or so. The driving force for supercontinent assembly is plate tectonics and associated rifts of continents. With the subduction of seafloor and the closure of oceans, large cratonic masses get amalgamated together, often associated with Mountain building processes (Himalaya, Alps) and fold belts (Lachlan Fold Belt). As for the break-up of supercontinents, plate tectonic forces alone cannot account for the dismantling of large continental masses. Old continents have deep cratonic keels with relatively cold, dry and refractory

lithosphere. Such lithosphere is not easily destroyed. Associated with the break-up of Gondwana are continental flood basalts (CFB), often referred to as large igneous provinces (LIPs). Within Gondwana, numerous LIP are recognized, the most important one for the break up being the Karoo-Ferrar-Tasman LIP (188-178 Ma) with magmatism characteristic for destructive plate margins (more silicic) at 172-162 and 157-153 Ma, respectively. Fig. 2: present day outcrop of Mesozoic CFB & oceanic plateaus. Hotspot traces as dashed lines. From Hawkesworth et al., 1999, JAES.

Modes of Supercontinent break-up

The number of CFB preserved over the lifetime of Gondwana is considerably lower than can be expected by distributions in other continental settings

Plume heads interact with lithosphere and not always result in CFB/LIP Continuous plume activity underneath the Gondwana continent caused long-term thermal

incubation of the lithosphere and associated weakening Supercontinents effectively self-destruct in response to the build-up of heat and can be

triggered by any new plume event. The lithosphere is susceptible to regional extensional tectonics and will respond to

mantle/plate movement (present day examples: East African Rift / Rhine Graben)


The Paraná-Etendeka-flood basalts and the Tristan Da Cunha hot spot

Basalts forming the island of Tristan da Cunha in the central South Atlantic Ocean are currently sampling the Tristan da Cunha hot spot. Its paleo-trail is preserved in the Walvis ridge that points towards the NW coast of Namibia; a mirroring image is pointing to Brazil. On both sides of the continents, the trail coincides with the Paraná (Brazil) and Etendeka (Namibia) flood

basalts (134-129 Ma). There are considerable geochemical variations between Etendeka and Tristan da Cunha basalts, which reflect the relative lithosphere-crust-plume interaction. The time of Paraná-Etendeka CFB eruption coincides with the opening of the Atlantic and the break-up of Africa and South America, leading to the assumption that the plume head triggered the rifting of the African and South-American plate. Figures 3+4: Paleo and present day position of the Tristan da Cunha hot spot; www.tristandc.com/

*** Fact Sheet *** Supercontinent A landmass that generally consists of all modern continents. Supercontinents

merged and broke up from ~ 3.1 Ga to 0.2 Ga, starting with proto-continents (cratons). Cycles of supercontinent formation and endurance and times between supercontinents spread between 200 to 500 Ma. Gondwana and Laurentia are the only supercontinent that existed independently and contemporaneously. Which supercontinents do you know?

Sub-continental lithosphere

Buoyantly, refractory mantle sections under continents. Relatively cold, dry and non-flexible (as compared to the underlying convecting asthenosphere). Forms by melt extraction of basaltic liquid from a fertile peridotite assemblage early in Earth’s history. Usually old, especially under cratons with ages up to 3.5 Ga. How can lithosphere be dated?

Oceanic lithosphere

Lithosphere forming at Mid-Ocean Ridges (MOR) by the extraction of basaltic melt, which forms the oceanic crust (layers 2&3). What happens to oceanic lithosphere when oceanic crust subducts?

Mantle plume Uprising mantle material, usually 100-200°C hotter than ambient mantle. Source in the mantle transition zone (600 km discontinuity) or core-mantle boundary (D’’ layer). Often associated with the return of subducted material. How can mantle plumes and its components be traced?

Hot Spot (-track)

A thermal anomaly in the mantle associated with intra-plate volcanism, generally considered to reflect uprising mantle plumes. The latter are ±fix in the mantle so penetration of the overriding plate records relative plate motion (e.g., Emperor Seamount Chain). What’s the difference between OIB and MORB?

CFB Continental Flood basalts – sudden and massive accumulation of basaltic lavas in excess of contemporary volcanic processes. Large flood basalt events mark the onset of major hot spots (also oceanic plateau). The massive extrusion is considered to reflect the plume head, whereas later hot spot activity is considered to sample the plume tail. CFB coincide with rifting events and are thus believed to be active triggers of rifting. CFB vs. layered intrusion: discuss!


Snowball Earth and Namibia Written by Dean Erasmus

Negative carbon isotope signatures in carbonate units succeeding the glacial deposits in Namibia indicate that biological productivity collapsed in the surface ocean for millions of years (Hoffman, 1998). Following glaciation, gasses expelled from volcanism resulted in greenhouse gas concentrations reaching significant levels - approximately 350 times modern concentration levels (Hoffman, 1998). These gases were taken up by the earth’s oceans, and consequently resulted in rapid deposition of calcium carbonates. Evidence of such extensive glaciation

can be found in correlative glacial and carbonate deposits on almost every continent (Hoffman, 1998). This period has been termed Snowball Earth and occurred between 900-600 Ma - during the Neoproterozoic (Hoffman, 1998).

Figure 2: Geological map of the Otavi Fold Belt showing the Neoproterozoic carbonate-dominated succession (Otavi Group) and the foreslope-platform facies change.

The Otavi Group (Figures 2 and 3) is a carbonate platform that covers the southern Congo Craton in northern Namibia (Hoffman, 1998). This group contains two distinct units where evidence of Snowball Earth can be found. This includes the Sturtian (~760-700 Ma) Chous and Gaub Formations (Hoffman, 1998). Both units contain thick carbonate successions with high δ13C values, overlain by cap carbonates, which record higher atmospheric greenhouse gas concentrations, as demonstrated in Figure 4.

Maieberg Fm.

Oombaaitjie Fm. A

Elandshoek Fm.

Figure 1: Maieberg Wall – Tweelingskop Farm 676. Alternation of glacial diamictite and cap carbonates.


The youngest of these (Gaub Formation) consists of debris flow deposits, mixed sediments, and glacial dropstones – amongst other things. It is relatively thin and discontinuous. The relationship between units within the Otavi Group as well as the corresponding δ13C throughout the succession is indicated in Figure 2 above. Figures 3a-c below present evidence of glaciation having occurred during the deposition of the Gaub Formation.

Figure 3: (a) Ice-rafted dropstone near the top of the proglacial Gaub Formation. (b) Dr Paul Hoffman (Havard scientist & snowball earth expert) examining evidence for snowball earth in Namibia. (c) Marine tillite (diamictite) from the younger Cryogenian glaciation within the Gaub Formation.

Figure 4: Stratigraphic cross-section of the western Otavi Fold Belt. Representative carbon isotope (δ13C) values are shown for the Otavi Group platformal carbonates. References Hoffman, P. F., Kaufman, A.J., Halverson, G.P. and Schrag, D.P. (1998). "A Neoproterozoic Snowball Earth." Science 281: 1342-1346.

A B

C


Road Guide: Uis to Otjiwarongo to Tsumeb Compiled by Dominique Tanner

This section compiles different excursions from The Roadside Geology of Namibia by Gabi Schneider (2008). Uis As we leave Uis, the road cuts across the Autseib Fault. This fault marks the boundary between the Northern and Central Zones of the Damara Orogen. Sadly, it is not well exposed here. To the W of the road we will see a range of white hills, which consist of marbles from the Karibib Formation. The small, roundish outcrops in the foreground consist of Damaran granite. Erongo Ring Dyke The Erongo Complex can be seen to the S. Hopefully we will have time to get out and look at boulders from the dolerite ring dyke, near a river. They contain labradorite, augite, and olivine. The magnetic signature suggests that the ring dyke might instead be a cone sheet (Pirajno, 1990).

Map of the road from Uis to Otjihaenamaparero (red). Arrows point to the age and locality of alkaline intrusives. The Erongo ring dyke is shown in yellow. Compiled from various sources using Google Earth.


Omaruru Omaruru is a town situated aside the Omaruru River – one of the largest ephemeral rivers in Namibia. It’s paleochannels form a delta beneath the Namib Desert. A prominent granite inselberg (the Omaruru Koppie) can be seen SE of the town. As you drive 9 kms from Omaruru towards Otjiwarongo, you can see a sharp-crested olivine-dolerite dyke, which runs parallel to the road for a couple of kilometers and forms dark hills to the East. These dykes are related to the emplacement of the Erongo Complex, and follow the structure of the Omaruru Lineament (see the black dotted line in the map on the left). Railway Bridge A railway bridge occurs 27 km from Omaruru (towards Otjiwarongo). Looking NW from here, you can see the banded and coarse-grained marbles which make up the distant Epako Range. Looking to the SE from this bridge provides views of the Erongo dyke swarms and the Omaruru Koppie. Skarnification Outcrop Damara marble outcrops about 49 km from Omaruru to the left of the road, and the next few kilometers are characterized by a complex mix of Damara meta-sediments, granites and a few post-Karoo dykes. While the granites form distinctive koppies or domes, low hills near the road consist of skarns. Intermingled granites, skarns and marbles occur in several roadcuts, many of which contain large calcite crystals formed during contact metamorphism. However the best outcrop occurs to the SE of the road, approximately 60 km from Omaruru. In this outcrop, pink Salem granite with biotite contains dark xenoliths at various stages of assimilation. Skarnification occurs in a mm-thick band along the contact between grey coarse-grained marble and the cross-cutting granite. Both the marble and the granite have in turn been cut by a 2 metre thick dolerite dyke. Kalkfeld Stopping at the Outjo turnoff provides a great view of the anorogenic alkaline complexes that haunt the region. To the SW, the Kalkfeld Carbonatite Complex forms a prominent landmark, whilst the pyramid-shaped Etaneno complex lies a few kilometers to the west, and the Ondumakorume Complex lies to the E, next to the road. To the N in the distance, the jagged peak of Klein Paresis and the massive Groot Paresis Complex (1806 m asl) can be seen. The road from Otjiwarongo to Tsumeb After leaving Otjiwarongo for Tsumeb, we enter the Northern region of the Damara Orogen with amphibolite-facies metamorphism (see Temma’s report). A number of hills and mountains can be seen on both sides of the road, consisting of Swakop-Group marble. In the E, the Honigberg rise can be seen, while the Jagershof Mountain (part of a NE-striking marble range) reaches 1822m asl. Otavi As we approach Otavi, we will see the elongated, dark Elefantenberg to the E of the road. This feature is made from diamictites, iron-formation and dolomites of the Abenab Subgroup, the Lower Tsumeb Subgroup and Otavi Group. We reach the rocks of the Otavi group, 5 km outside of Otavi. Approaching Tsumeb As we approach Tsumeb, we must negotiate the “Ten-Mile Pass”, where hills formed by the Elandshoek formation dolomites can be seen on either side of the road (see Dean’s report). These rocks are exposed in outcrops along the pass, while the low-lying area below the pass exposes the Abenab carbonates.


The Tsumeb Ore Body Written by Alex Moody

Geological Setting Tsumeb mine is situated in the Otavi mountain ranges in north eastern Namibia. It is comprised of a 1.7km deep hydrothermal ore deposit, situated in Precambrian dolomites. The mine itself is dominantly a copper, lead and zinc deposit, with substantial amounts of Germanium, assaying up to 0.83% (Otavi mountain land). Otavi mountain land shows substantial evidence of multiple glaciation events, with substantial glacial diamictite deposits, and changes in δ13C in carbonates under and overlying the diamictite deposits. Similar deposits are seen globally and are thought to indicate two separate global glaciations events. The timing of these events is contentious, with Hoffman et al. (1998) arguing for a time frame of 760-700mya, while Kennedy et al. (1998) and Folling and Frimmel (2002) argue for a more recent age of 580ma. Rapid deglaciation was most likely caused by a rapid and substantial increase in volcanic activity. This rapid warming allowed for the deposition of cap carbonates directly on top of glacial diamictite deposits. Regional metamorphism as a result of the continent-continent collision of the Congo and Kalahari Cratons, occurred at approximately 545mya.

Geology of the Otavi Mountain Land (based on Söhnge (1957) and the 1:250,000 map of the Geological Survey of Namibia (1999)) with locations of major deposits and some prospects. (Image sourced from A.F. Kamona, A. Günzel 2006. ) The hydrothermal fluids responsible for the formation of the ore are theorised to have most likely been produced by granites of Cambrian age, or post-Triassic ring complexes, which have a well defined ENE belt, which the Tsumeb, and Tsumeb West deposits lie in. Sr/Sr dating has at this stage proved inconclusive in determining an exact age of the deposit, and thus, its origin.


Description of Ore Body The Tsumeb ore body is a gently southward dipping ore body, which at a depth of approximately 600 metres, turns towards the north and dips more steeply, the ore deposit has been exploited at depths up to 1700 metres. The ore body is generally oval shaped, and reaches a maximum width of 180-60 metres. The geometry of the deposit has been heavily influenced by carbonates which the pipe is situated in.


Ore Body Mineralogy The mineralization in the Tsumeb pipe is associated with some wall rock alteration, a pseudo-aplite of as yet undetermined origin (which some studies suggest is a sedimentary rock formed during karsting, which was later metamorphosed by hot fluids which precipitated the ores), and solution collapse brecciation. The mineralization in Tsumeb is highly complex, with over 240 minerals described at the site prior to 2000, with Tsumeb being the type locality for over 55 of these. There are three dominant reasons for the spectacular array of minerals present in Tsumeb, 1. Whilst the dominant sulfide minerals at Tsumeb are Cu, Pb, Zn, and As, there are a large number of rare elements present, notably germanium, which is present in the mine in such substantial amounts, that it is estimated that the ore body contains a total of 2160 tonnes of Germanium. 2. The primary sulfide ore has been oxidised to varying degrees in three different zones throughout the deposit. 3. The deposit is located in a Karst pipe of a substantial carbonate formation. Because of the unique factors affecting Tsumeb, the minerals present can be broadly characterised into the following groups, with 28% of the minerals present being phosphates of arsenates of Cu, Pb, Zn, or As, 19% sulfides, 15% silicates, 11% sulfates, vanadates and molybdates, with the remaining minerals dominated by carbonates, which account for 9%. The early phases of sulphite ore deposit in Tsumeb were composed of Cu rich sulphide minerals, the earliest sulphide phase observed in Tsumeb is pyrite, but it is limited in abundance, chalcopyrite is the earliest substantial sulphide remain, although it still comprises only a small percentage of the remaining ore, and has been mostly replaced by later stages more rich in Cu. The next sulphide stages to become stable were chalcocite and bornite; it is most likely that the initial deposit was of a chalcocite-bornite solid solution, due to observation of exsolution textures in both these minerals. (A geochemical model for sulphide paragenesis and zoning). The final hypogene Cu phases formed are tennantite and enargite, both these minerals occur in veins and pods which cross cut the earlier chalcocite and bornite. Galena and sphalerite post date all other major sulphide phases. Using fluid inclusions in quartz which is intimately associated with the bornite and chalcocite deposits, indicate two stages of fluid based deposition, which, when accounting for likely formation pressures experienced peaks of between 210-240, and 260-280 degrees Celsius. The sodium content of the ore depositing fluids was also determined using quartz fluid inclusions, which gave an approximate NaCl content of 12 wt%.


Pegmatites: Did you know…? Written by Dominique Tanner

Pegmatite: A crystalline, intrusive rock composed of interlocking crystals >2.5 cm. It is well known that granitic pegmatites are enriched in the most incompatible elements – which is why some host rare gemstones such as topaz (Al2SiO4(F,OH)2), emeralds (Be3Al2(SiO3)6 with minor Cr and V) and tourmaline (Na(Li,Al)3Al6(BO3)3Si6O18(OH)4). …but did you know: At how low a temperature pegmatites can crystallise? Experimental petrology has shown that some of the (geologically) rare incompatible elements found in pegmatites, such as Li, B, P and H2O have been shown to depress the solidii of melts. Materials scientists have been aware of this for a long time, as they use these elements to flux the melting point of glass, ceramics and glazes. Sirbescu and Nabelek (2003, Geology) have found evidence that a Li-rich pegmatite crystallised between 350-400°C. They obtained these temperatures from primary fluid inclusions coexisting with pegmatitic melt inclusions in magmatic quartz from the Sn-mountain pegmatites in South Dakota, USA. I think that this is an extremely important study, as it means that there is significant overlap in P-T space between fluids (fumaroles up to 900°C) and melts (now <400°C)! Why pegmatites contain such large minerals? Pegmatitic minerals can be so large, that it is reported that a quarry was opened up in a single feldspar crystal (Lindgren, 1933)! Other notable occurrences include an 18m beryl, a 14m spodumene crystal, and a sheet of muscovite 5m x 3m (Vernon, 2004). In first year, we teach students that the rate of cooling determines the grain size of an igneous rock (i.e. aphanitic volcanic rocks and phaneritic plutonic rocks). However, granitic pegmatites occur as narrow dikes at shallow levels in the Earth’s crust – so why aren’t they aphanitic? The answer (according to Vernon, 2004) is that pegmatitic magmas exhibit very low rates of crystal nucleation, compared to mineral growth rates – due to the abundance of water, B and or P in the melt. The water and/or trace elements suppress nucleation of new crystals in the melt, resulting in rapid overgrowth on existing crystals. Mafic pegmatites exist! Traditionally we think of pegmatites as being granitic, since Bowen (1928) showed that granites are the most evolved magmas. The literature is quite clear that extensive fractionation concentrates incompatible elements, which in turn suppress nucleation – causing the growth of large crystals. However…. some layered mafic intrusions (LMIs) contain pegmatitic horizons. This includes pegmatitic Merensky Reef, that we will hopefully see when we visit the Rustenburg mines. The origin of mafic pegmatites on the Bushveld particularly is contentious (or ignored), as many authors consider the Merensky Reef to be “orthomagmatic”. This means that they think the Merensky Reef precipitated directly from a fresh pulse of mafic magma. But how could this cause the ore to form pegmatitic crystals? Perhaps it is more logical if the large crystals formed from extensive fractionation of mafic magma (causing a build-up in incompatible elements). This means that one of the world’s largest PGE reserves may have formed during the last – possibly hydrous – stages of crystallization. Keep this in mind when we see the Bushveld.


The Okorusu Fluorspar Mine Written by SJ Collum

Introduction Okorusu is an open pit fluorite mine located in northern Namibia within the Pan-African carbonaceous rock of the Damara Sequence (Fig. 1). It is approximately 48km north of Otjiwarongo and is the largest deposit of fluorite in Namibia. It has been mined since shortly after it was discovered in the early German colonial era. It has outstanding blue and green fluorite crystals, however green and purple crystals predominate. Yellow, grey and colourless fluorite crystals are also present. Aside from the excellent fluorite crystals, acid-grade fluorite concentrate is produced at Okorusu, for the production of hydrofluoric acid. Geology of Okorusu Alkaline silicate rocks and carbonatitic lithologies make up the Okorusu complex. The main carbonatite body occurs at the southern edge of the intrusion (Fig. 1). This body is composed of medium-grained, white to grey calcite and accessory apatite, quartz and feldspar. In the immediate vicinity of the Okorusu complex, especially in the south where the carbonatite body is exposed, a widespread and intense brecciation and fenitization aureole is developed. This area is rich in fractured and altered pyroxenite at the contact between the country rock limestones and greywackes. The fluorite mineralization occurs in the areas shown in black in Fig. 1, where the Damaran sedimentary carbonate and the fenitization aureole meet. The Damara System deformed and metamorphosed metasedimentary rocks consist of greywackes, micaceous and calcareous quartzites, various schists, conglomerates and crystalline, quartz-bearing limestone. The vein deposits of fluorite display typical replacement features, such as comb texture and coarse banding. Barite, smoky quartz, fluorapatite and calcite all occur as minor constituents of the fluorite mineralization.

Figure 1: Geology of the Okorusu carbonatite complex and associated fluorite mineralization. From Buhn et al. (2002).


Carbonatite at Okorusu The fluorite ore bodies at Okorusu have been emplaced primarily by the preferential replacement of carbonatites. This pegmatitic carbonatite is commonly present as pods, dykes and veinlets that intrude sodic fenitized metasedimentary rocks. The fluorite is mined from three pits at Okorusu, however only two pits have been extensively studied, the A pit and the B pit. The A pit has pegmatitic carbonatite, fine-grained calciocarbonatite and small amounts of pyroxene carbonate, a

green rock which exhibits a wide range of grain sizes. Pyroxene carbonatite is the main rock exposed at the B pit. The B pit (Fig. 2) also has pegmatitic carbonatite, fine-grained calciocarbonatite and local amount of course-grained calciocarbonatite in the hanging wall of pit B. Another interesting feature of pit B is a single vein of medium-grained high-radioactivity carbonatite. The carbonatite is dated between 137 and 124 Ma, and thought to be related to the Tristan da Cunha hot spot and the opening of the South Atlantic.

Figure 2: The B pit at the summit of Okorusu Mountain. From Hagni & Shivdasan (2000) Fluorite Orebodies at Okorusu Fluorite mineralization is caused by a fluid (a crustal brine or from an igneous source) interacting with crustal rock and/or waters. Fluorite mineralization occurs in two situations at Okorusu, the first is as a replacement deposit in the feldspar-bearing limonite-calcite rocks and the second in massive fluorite veins intersecting quartz-bearing marble bands. The replacement fluorite ores tend to be fine-grained and purple while the vug-filling fluorite tends to be green. The explosion of fluids form the Cretaceous Okorusu carbonatite caused the epigenetic mineralizing event responsible for the formation of the fluorite. Okorusu has low-salinity primary and secondary fluid inclusions in the fluorite and the fluorites also contain an abundance of high-field-strength elements. A study by Buhn et al. found a population of highly saline H2O-CO2 fluid-inclusions in the directly adjacent Pan-African country rock quartz. They concluded that that the heterogeneously trapped inclusions represented a primary, carbonatite-derived, mineralizing fluid that precipitated fluorite. The surrounding country rocks are considered to be representative of an orthomagmatic fluid responsible for the fluorite mineralization. The fluid formed a varied inclusion population which consists of everything between an aqueous brine, a CH4-bearing CO2 phase and solid phases, due to its extreme heterogeneity. This inclusion population lead Buhn et al. to conclude that the fluid in the inclusions was a sample of the carbonatitic fluid with only minor crustal contamination, which reacted with crustal carbonate rocks to precipitate fluorite. Fluorite Replacement of Pegmatitic Carbonatite – Evidence from Textures In general all traces of calcite in the original pegmatitic carbonatite are replaced by fluorite, however in some instances remnants of the partly replaced to un-replaced carbonatite remain locally in pods of fluorite ore that indicate the character of the replaced rock. In the rocks that have been completely replaced by fluorite titaniferous magnetite rims and goethite pseudomorphs indicate that the original rocks were pegmatitic carbonates. The titantiferous magnetite is present in rims around the fluorite ore deposits, remnants of coarse euhedral crystals in the pegmatitic carbonatite.


Figure 3: Hydrothermal mineralization at the Okorusu fluorite deposit. (a) Continuous fluorite I profiles of sample OK1, extending on both sides of metasomatized wallrock. (b) Close-up of the outermost portion of fluorite I sample OK2, showing the gradation between light yellow fluorite and colourless fluorite with purple laminations. The fluorite is coated with opaque phases. (c) Massive fluorite II (sample OK3) consists of closely spaced colourless and purple fluorite laminae. Note the different texture in comparison to fluorite I, with varied orientations of laminated fluorite crystals. (d) Fluorite I sample OK2. (e) Massive fluorite II sample OK7. Colourless and purple fluorite is capped by euhedral, smoky quartz, which is again capped by a Mn-rich crust on earliest calcite, followed by pure calcite. (f) Massive fluorite II sample OK8. The sample consists of intergrown purple fluorite and patches of fluorapatite rosettes. Vugs are filled with late calcite.


The pegmatitic carbonatite also has titantiferous magnetite disseminated through it. These coarse crystals were oxidized to equant goethite pseudomorphs during replacement of the carbonatite. Another interesting feature at Okorusu is large, platy crystals shapes of goethite in the fluorspar ore. The platy goethite was formed through oxidation of platy pyrite, which can be seen in the pegmatitic carbonatite. The fluorite ores also displays banded or planar features and are abundant in small platy vugs. The ore has a general tendency to a locally present alignment texture but typically lacks fluorite alignment features throughout the massive replacement fluorite.

Figure 4: Pegmatitic carbonatite in fluorspar ore, there the carbonatite consists of calcite (white), apatite (medium grey) and platy pyrite-pyrrhotite (dark grey). From Hagni & Shivdasan (2000)

Figure 5: Replacement remnants of pegmatitic carbonatite (bottom half of photo) in fluorspar ore. From Hagni & Shivdasan (2000)

Fluorite Replacement of Marble

A small proportion of the pit A ore and a significant proportion of the pit B ore are formed by the replacement of marbles instead of carbonatite. Over the two pits, about 20% of the fluorite ore is formed by marble replacement. Fluorite ores in the new C pit have been dominantly formed by the replacement of marble. These ores are characterized by strong foliation textures relict from the foliation of the replaced metasediments. Some of the fluorite concentrates have elevated silica, due to some of the metasediments containing potash feldspar which locks the potash feldspar and fluorite particles.

Figure 6: Fluorite ore (left) showing vertical foliation inherited from the replacement of marble and biotite schist (right). From Hagni & Shivdasan (2000) Beneficiation The carbonate source of the fluorite ore poses problems for the process of beneficiation, where the ore is separated into the mineral and gangue. The A pit at Okorusu has problems with excessive phosphorous (Fig. 7) while the B pit has problems with excessive silica. Both of these particles lock with the fluorite during beneficiation. While the purple fluorite ores in the A pit contain greater amounts of apatite than phosphorous, it is the greater amounts of phosphorous in the green fluorite that causes the major beneficiation problems. In the fluorite ore that has replaced marble, silica causes beneficiation problems.

Figure 7: a. Abundant large apatite crystals (white) contained in massive fluorite (medium gray) ore mine face. From Hagni & Shivdasan (2000)


The Otjihaenamaperero Dinosaur Footprints Written by Dylan Singh

Introduction Dinosaur footprints from the Late Triassic and Early Jurassic are common in the Karoo Age formations in the Waterberg Plateau Park, Mt. Etjo and the Omuramb-Omambonde region of Namibia. Of all these sites, the fossilised dinosaur footprints that form the trackways at the Otjihaenamaperero 92 Farm are the best preserved and most abundant. The Otjihaenamaperero 92 Farm dinosaur trackway is located on a farmstead 16km southeast of Kalkfeld, to the northeast of Omaruru. The dinosaur trace fossils are preserved in the deflation surfaces of the Upper Etjo Sandstone, an Aeolian fluvial sandstone the time equivalent of the Clarens Formation in South Africa. The Etjo Sandstone is composed of three sandstone units. The Lower sandstone and gravel unit contains only trace fossils from burrowing organisms that lived in freshwater (Schneider G & Schneider M, 2006). In the Middle sandstone unit a partial skeleton imprint of a Massospondylus, an early 6 metre long herbivorous prosauropod dinosaur, has been discovered (Schneider G & Schneider M, 2006). The Upper sandstone unit of the Etjo Formation is the largest of the three, formed from windblown sand deposits and is well sorted and heavily fractured. The Etjo Sandstone is overlain by the Rundu basalts which is the time equivalent of the flood basalts from the Drakensberg Group (NAMEX, 2009). Radiometric dating of the younger basaltic rock from the Drakensberg basalts gives a lower limit of 186-183Ma for the Upper Etjo Sandstone (NAMEX, 2009), indicating an Early Jurassic age of deposition. This was a time before Earth’s single continent Pangaea had completely separated into Gondwana and Laurasia. This formation has also been correlated to the Piramboia Formation of the Parana Basin in Brazil (NAMEX, 2009), which contains major oil reservoirs in tar sands. Other dinosaur ichnofossils found in the Etjo Sandstone in the Waterberg region include tracks believed to be from Quemetrisauropus princeps and Prototrisauropus crassidigitus, both early herbivorous dinosaurs. Footprints from cynodonts, mammal-like reptiles, have also been discovered. The dinosaur footprints in Otjihaenamaperero were first reported in 1925 by Von Huene, in 1932 Heinze identified two different sets of tracks. The fossil trackway at Otjihaenamaperero was declared a National Monument of Namibia on the 1st of August 1951. Environment and Deposition The sediment that formed the Etjo sandstone accumulated as Aeolian sands in conditions similar to the Namib Desert of today. The three sandstone units of the Etjo Formation show a transition in climate from semi-arid to fully arid conditions (Schneider G & Schneider M, 2006). Numerous reptiles, including dinosaurs, and mammal-like reptiles lived in the inter-dune areas of this subtropical ‘wet desert’ environment during the Late Triassic. The increasingly arid conditions during the Early Jurassic forced these animals to congregate near the remaining water bodies, waterholes, small lakes and rivers fed only by occasional rains. The sands at Otjihaenamaperero held one such vital water source (Merkel B, 2004) and it was during this time that the footprints were left in the soft, wet sands surrounding a dwindling body of water by some of the animals that relied so desperately on them. As the Jurassic progressed, these water bodies became scarcer as conditions became even drier. It is assumed that sometime after the footprints were formed that, due to the increasing aridity of the landscape, dinosaurs completely disappeared from this area. After the imprints were made in the soft sandy sediment they were covered by further layers of windblown sand. The pressure of the overburdening sediment caused the sand to lithify over time, preserving the footprints as trace fossils.


Otjihaenamaperero Dinosaur Footprints The dinosaur trackway at Otjihaenamaperero contains three sets of tracks. All the footprints were imprinted in wet inter-dune sand by the hind feet of two species of early three-toed (tridactyl), bipedal saurischian (lizard-hipped) therapod (a group of carnivorous dinosaurs from which birds are descended) dinosaur. The two different tracks are situated on separate slabs of sandstone. One slab contains two intersecting older trackways (Merkel B, 2004), each containing over 30 footprints, which can be traced for a distance of approximately 30 m. The footprints are approximately 23cm long and 19cm wide, with a distance of between 67 and 90cm between them (Strobel A & Strobel R, 2011). One set of tracks heads in a north-westerly direction while the other goes south west. The other sandstone slab contains smaller footprints and a shorter trackway. These footprints are about 7cm long, have an average stride length of 35cm and can be followed for just over 10m in an easterly direction (Strobel A & Strobel R, 2011). Because most footprints are a poor record of foot structure (Lockley M, 1991), and because no fossils relating to the footprints have been discovered in the Upper Etjo sandstone formation, it is difficult to determine the morphology of the trackmaker and make a positive identification. Above left: large single imprint from the older tracks (Merkel B, 2004). Above right: smaller footprints from the younger trackway (Merkel B, 2004).


Dino 1 The track dimensions and stride lengths of the larger sets of tracks indicate that they belong to a dinosaur between three and six meters in length. These trace fossils are generally believed to have been left by a Ceratosaurus (Schneider G & Schneider M, 2006; Strobel A & Strobel R, 2011), meaning ‘horned lizard’. Named for a pair of short horns on the end of its snout, Ceratosaurus was discovered in 1884. It is very similar to the larger Allosaurus, being both closely related and contemporary. Ceratosaurus was a common, widespread and successful dinosaur during the Jurassic. Ceratosaurus nasicornis is known in Africa from fossils found in Namibia and Tanzania and Utah and Colorado in North America. Other species have also been discovered in America and Portugal. Ceratosaurus has been found to have lived in all manner of environments; from desert-like conditions to swampland. Its head and jaw were huge in comparison to its body, filled with large bladed teeth. It had short, powerful arms, and like many primitive dinosaurs, had four clawed fingers instead of three. It had a long, flexible tail similar to a crocodile’s and is believed to have been a strong swimmer. C. nasicornis was powerfully built and could weigh up to one tonne. Fossil evidence suggests that it may have hunted in groups or pairs. Primarily an active hunter, they would have preyed largely upon the giant herbivorous dinosaurs of the Jurassic, but would have supplemented this by scavenging. Above: C. nasicornis profile (Feenixx Publishing Inc., 2007). C. nasicornis (Wikipedia, 2007)


Dino 2 The smaller footprints are believed to have been produced by Megapnosaurus (Schneider G & Schneider M, 2006; Strobel A & Strobel R, 2011), an early therapod whose name means ‘big dead lizard’. There are two species of Megapnosaurus; M. rhodesiensis, the African species known from the Forest Jurassic Sandstone of Zimbabwe, as well as the Clarens and Elliot Formations in South Africa, and the American species M. kayentakatae. M. kayentakatae was double crested while M. rhodesiensis was uncrested. Discovered by M. Raath in 1969, this dinosaur was originally named Syntarsus, meaning ‘fused tarsus’ for its fused ankle, by its discoverer. It was later renamed, as Syntarsus had already been claimed by a genus of beetle, in the absence of Raath as he was presumed dead. Megapnosaurus had a slender, streamlined body with a lightweight skeleton; it was one of the first terrestrial animals with a form designed for speed. Despite measuring up to three metres long from head to tail, it would have weighed only about 32 kilograms. Like Ceratosaurus, it had a primitive four-fingered configuration to the hands at the ends of its long grasping arms. Megapnosaurus is a coelophysid, an early group of carnivorous dinosaurs that hunted in packs dominant during the Triassic; one of the few to survive into the Jurassic. Fossil evidence shows that Megapnosaurus was also a pack hunter, at one site in Zimbabwe over 30 individuals were found together. It was hypothesised that they were buried and preserved by a flash flood. They are thought to have lived alongside freshwater where they hunted small reptiles and fish. Above left: M. rhodesiensis profile (Wikipedia, 2012). Above: M. rhodesiensis (Natural History Museum, London, 2010)


The Windhoek Meteorites Written by Sean Jefferson

Windhoek is the capital and largest city within Namibia. It is located in the Khomas Highland plateau area 1,700 metres above sea level. What is most interesting about the city of Windhoek is the meteorite display which is located within Post Street Mall. In all there are 33 pieces on display weighing between 195 kg and 555 kg. The meteorites displayed are a selection of the Gibeon Meteorites, named after the closest town of Gibeon where they fell. The Meteorites are in fact the largest known shower of extra-terrestrial bodies ever to land on Earth; to date 120 pieces have been recorded however the exact number is unknown due to them being smuggled out of the country over many years.

The photograph above shows the Gibeon Meteorites on display at the Post Street Mall taken from http://www.namibiatourism.com.na/ Radiometric dating of the composition of the Gibeon Meteorites has aged them at 4 billion years old with the astronomical origin believed to have been from an exploding supernova. Although the date of collision with Earth is unknown it is estimated to be between 200-220 million years ago. Trajectory reconstruction of the flight of the original meteoroid body which created the Gibeon Meteorite shower show that it would have been about 4 x 4 x 1.5 m entering at a reasonable low trajectory (20º) from north-westerly direction. The evidence for this is shown in the wide spread of the Gibeon Meteorites themselves across an ellipse of 275 x 100 km. The Gibeon Meteorites composition is that of an octahedrite the most common type of iron meteorite and consists largely of iron and nickel in the form of alloys taenite and kamacite. Cobalt, phosphorus, chromium, copper and carbon are also found in small amounts as well as traces of germanium, zinc and gallium. It is however the inclusion of the Taenite and Kamacite alloys which has allowed scientists to reveal more about the history of Gibeon Meteorites and the original meteoroid.


A texture known as a Widmanstätten pattern is present throughout the Gibeon meteorites. The crystalline patterns become visible when the meteorites are cut, polished, and acid etched. This is formed by the super slow cooling. At temperatures below 900 to 600 ºC both alloys are stable, Kamacite with lower Ni-content (5 to 15% Ni) and Taenite with high Ni (up to 50%). Octahedrite meteorites have a nickel content which lies somewhere between the two, and so during super slow cooling events the precipitation and growth of Kamacite and Taenite occur. This occurs through the process of diffusion a very slow process, for crystals to have grown to the size witnessed within the meteorites a cooling rate of 100 to 10,000ºC/Myr is required. This proves that the meteoroid was in fact in circulation in space for a long period of time before collision with the earth.

Left is an example the Widmanstätten pattern seen within a slice of a Gibeon Meteorite. Taken by http://geology.com/meteorites/iron-meteorites.shtml

Right is a polished section example of light Kamacite bands which are bordered by darker Taenite ribbons taken from the New England Meteorical Services http://www.meteorlab.com/METEORLAB2001dev/images/gib100x2.gif


The Bushveld Complex Written by Heather Marman

This section outlines the how the Bushveld Complex was formed, and will cover an introduction to the complex, an overview of the two components of the complex (a layered mafic suite and a granitic suite) addressing their formation and mineralogy, and the economic importance of these suites. The majority of the information is from A First Introduction to the Geology of the Bushveld Complex and those aspects of South African geology that relate to it by Hugh Eales (2001). Introduction to the Bushveld Complex The Bushveld Complex is an intrusive igneous complex made up of a mafic layered sequence capped by a felsic granitic sequence. The outcrop distribution can be seen in the below image. The whole complex intruded around 2.06Ga into the Transvaal Supergroup, a thick sequence of sedimentary units and lava flows covering an area of 500,000 km2. The complex contains economic deposits of platinum group elements (PGEs), chromium, iron, titanium, vanadium, tin and fluorine.

The exposed area of the Bushveld complex showing the mafic layered rocks shaded black and the granites with crosses. Sourced from Eales (2001) pp8. The whole complex is divided into 5 limbs which under modern theory are not thought to be connected at depth rather they are likely different offshoots of a common magma source. The Layered Suite; its formation, melt evolution and mineralogy The Rustenburg Layered Suite (RLS), is a 7km layered sequence of dark mafic material. It’s derived from mantle melt from a mantle diapir which rose after a sinking basin reactivated a heat source. The associated upward welling pressure of the diapir produced fractures along which magma (from the diapir) was able to flow and be emplaced. This idea supports the theory that the Bushveld complex is a series of arms and not joined continuously underground. Further development of this concept by C.J. Hutton in 1995, suggested that the diapir rose to within 18km of the surface and caused the crustal melts which formed the Bushveld Granites.


The formation depth of the RLS must have been at least 5km deep, determined by the chemistry of the surrounding metamorphosed material. The massive chamber was filled over a period of time through a secession of magma pulses. The first evidence that there was more than one magma event is the RLS exists as a series of layers with altering mineralogy, and some layers themselves show a cyclic layered sequence of mineral deposition (discussed below). The real evidence for a multi-magma event was the study of strontium isotope ratios. Throughout the layered sequence the isotopes vary significantly. If this were truly one event the ratio should remain relatively constant. The physical appearance of the RLS changes up-sequence due to a distinct change in mineralogy. The most identifiable aspect is the change of the cumulates within the rock. The cumulate mineral sequence from base up is: chromite → olivine → orthopyroxene → plagioclase → clinopyroxene → magnetite → alkali-feldspar. This is not a clean transition the whole way up the suite, rather a general trend with some layers being repeated in a cyclic manner conceptually represented to the left. These mineralogical changes led to characterising the RLS into zones. Whilst this is relatively arbitrary it is common reference and is used throughout the literature. The alterations in mineralogy also give an indication of what was happening in the magma chamber at a given time. The following is a brief outline of the zones: Marginal Zone – Is the area or “shell” around the base and sides of the complex and is dominantly, harzburgite, orthopyroxenite and norite. Lower Zone – Is a cyclic unit of olivine and orthopyroxene rich units indicating several injections of magma occurred during its formation. Lower Critical Zone – At this level olivine levels decline. This zone contains some of the most extensive layers of chromitite. Upper Critical Zone – First appearance of feldspar and contains a cyclic unit of chromitite, harzburgite, pyroxenites, norite, and anorthosite. This zone contains the famous and all-important Merensky Reef and Upper Group 2 (UG2) chromitite layer. Main Zone – At its thickest this unit is 4000m. It represents the time of maturing of the residual liquid which led to concentrations of vanadium and iron in the upper zone. Upper Zone – Full of iron rich minerals, i.e. magnetite, but also iron rich phases of minerals such as olivine (fayalite) and pyroxene. Plagioclase also moves to being sodic and then potassic in the upper most layers. Finally hydrous minerals appear and the significant depletion of metals throughout the whole complex leaves an excess of silica and free quartz appears. Apatite is found in this zone due to phosphates incompatibility and desire to stay in the fluid. Despite across the board agreement that the entire RLS formed from multiple magma injections the formation of the specific mineral sequence and particularly what caused the occurrence of cyclic sequences are still debated. Several ideas of how this has occurred have been proposed and the most reasonable suggestions proposed for the distinct mineralogical change up-sequence are:

• Alteration of pressure during crystallisation • Changes in oxygen content • Repeated injection of chemically the same liquids (for the formation of cyclic units within

layers i.e. the olivine pyroxene repetition in the lower zone.) • Sinking of crystals and therefore accumulating them on the chamber base

The cyclic sequence repeatedly seen in the RLS. Sourced from Eales (2001) pp40.


The Economic Importance of the Layered Suite After the Bushveld sequence was emplaced there was a long time period where the Transvaal Supergroup area underwent significant amounts of weathering and erosion, removing the upper zones of the Bushveld Complex in some locations. It is only because of this that the PGE bearing layers, Merensky Reef, and the UG2 Chromitite layer, became exposed at the surface and occur at mineable depths. The UG2 Chromitite layer with an average thickness of 1m is the largest current source of PGMs. In this unit, PGEs are contained within chromite and sulfides. In contrast the Merensky Reef is a pyroxenite layer, anywhere between 25cm and 14m thick, with enrichment of PGEs being dominated by sulfide inclusions and sulfide solid solution. As a group in the Bushveld Complex the PG minerals present are a combination of sulfides (40%), iron alloys (40%), tellurides (11%), arsenides (6%), and 3% gold and silver alloys. Because of the high proportion of PGEs preferentially moving into sulfide minerals, both as microscopic inclusions and solid solution, there is a direct correlation between the amount of sulfides present and the concentration of PGMs. There are again many theories to why this enrichment has occurred for the UG2 and Merensky layers and the most likely explanations are:

• the sulfides and chromites scavenged the PGEs from the magma; • side reactions were occurring during deposition of the chromites and sulfides which caused

the PGEs to accompany these minerals in deposition; or • the sulfide layers had already formed and trapped PGEs from fluids moving through the

system. The cause of the PGE levels of enrichment is another scientific quandary as the volume of magma required to reach the concentrations found is massive. The favoured explanation at present is the mixing of magma layers through a process called finger mixing, which is the insertion of new hot magma which rises as a plume due to its lower density and then protrudes ‘fingers’ of magma into the layers beneath it as it cools resulting in a layer of new chemistry. The general idea of how this process can cause the precipitation of large amounts of chromite and PG mins is displayed in the triangle diagram below. For the chromitite layers to have formed as thickly as they have this is certainly the reasonable explanation as chromium has low solubility in basaltic magma and would not have formed in its thick concentrated layers as it has through normal crystallisation of a magma chamber . In normal instances to have the thickness of chromitite present in the Bushveld complex (25m) there would have needed to be a liquid thickness of 25,000m.

The mixing of liquid with composition ‘a’ with a liquid of composition ‘b’ to have precipitation of chromite at composition ‘c’. Sourced from Eales (2001) pp65.


The Merensky Reef is a unit of particular interest and during its formation geologists do know that something happened in the magma chamber involving a new liquid. Prior to the reef being laid down there had to be a major pulse of new magma as there is a stark strontium ratio change between the base and top of the Merensky unit. The Merensky Reef also contains interesting structures dubbed potholes which make the unit not entirely laterally continuous. The structures are the result of scouring by newly injected magma into the material below which has not yet entirely consolidated, which has led to the deposition of the reef to occur within these scours. The potholes are also intimately associated with areas of dip change suggesting there is also an element of slumping involved in causing their formation. Some potholes are large enough that they have coalesced large amounts of platinum ore and have been targets of exploration. Another target of platinum exploration particularly in the early days was platinum contained in iron alloys found in discordant bodies which cut across the RLS. The bodies are pipe-like and some contain platinum in their cores. Their origin is still unknown and something which makes them even more interesting is they are not standard intrusive bodies. The UG2 chromitite layer in particular can be seen tracking across the bodies at the same height it would normally appear (a bit like a dotted line of bedding rather than a solid line). This suggests there may be a chemical reaction component to these structures but what it is has yet to be determined. The two spinels in the layered suite (chromite and magnetite) are very separated due to the evolution of the melt during crystallisation. However both have formed enough volume and concentration to be of economic value. The chromitite layers have already been discussed previously, however the magnetite layers of the upper zone, which formed later and from a much more evolved iron-rich liquid, also contain economic levels of vanadium. This is from simple substitution of vanadium for iron. The vanadium is not evenly distributed between all magnetite layers but displays a preference to be in the earliest deposited layers as its distribution coefficient means it prefers to enter the solid phase. The Bushveld Granites The Bushveld Granites termed the Lebowa Granite Suite, are 3-4km thick and formed above and after the layered sequence (RLS inclusions are present in the granite). The formation occurred 3.05Ga from a volume of magma in excess of 200,000km3. As mentioned above it has been proposed that the heat source for these crustal melts was the Layered Suite insertion. Within the granites there are different types, which have been subdivided into Nebo, Bobbejaankop, Lease, and Makhutso. The Nebo granite is coarse grained and has gradational changes in minerals through the sequence. The implication on how the granite formed is the magma crystallised from the base up and underwent fractional crystallisation in the process. As this process occurred the liquid evolved and the minerals precipitating altered. Bobbejaankop shows much more water present which caused cavities in the rock to form and also caused some low grade alterations in feldspar. The combination of water rich fluids and fractional crystallisation resulted in concentrations of tin and fluorine, both of which were mined from this unit throughout the 20th century. Bobbejaankop was intruded by the Lease group which is a series of fine grained dykes and sheets. The last group is the Makhutso, a biotite rich unit which intrudes into the Nebo group and formed from the melting of Archean gneisses. Fitting, as a great portion of the basement rocks of the African continent are old highly metamorphosed rocks.


Ore Mineral Help Guide Written by Tarun (Tazz) Whan

Sulfides and sulfosalts

Formula System/ Hard/streak Identifying features

Aresnopyrite (R,N) FeAsS Monoclinic/ 5.5-6/Blk

Silver-white looking pyrite with D. cleavage {001}

Bismuthinite (N) Bi2S3 Orthorhombic/ 2-2.5/Pb grey

Looks like stibnite, lead grey. P cleavage on {010},

Bornite (R) Cu5FeS4 Orthorhombic/ 3/Grn-B

Irreg. fracture, peacock blue colour tarnishes Cu-red

Calvertite (TL T) Cu5Ge0.5S4 Isometric/ 4.5

Black granular anhedral-subhedral xtals in matrix of chalcocite veins

Chalcocite Cu2S Monoclinic/2.5-3/Grey-Blk

Blue-black to grey

Chalcopyrite (R,O,N)

CuFeS2 Tetragonal/3.5/green blk

Brass to honey yellow, soft, looks like gold, blue-red tarnish

Cinnabar HgS Trigonal/2-2.5/bright red

Red-brown, lead grey, V.soft, dense

Cooperite (TL B) PtS Tetragonal/4-5 Steel grey, no cleavage, V. dense, conchoidal fracture

Covellite CuS Hexagonal/1.5-2/grey-black

From light to dark to Indigo blue, Perfect {0001} cleavage

Galena (O) PbS Isometric/2.5/ Pb-grey

Perfect cleavage {001},{010},{100}, V.soft, dense, silver cubes

Germanite (TL T) Cu13Fe2Ge2S16 Isometric/4/ dark grey-blk

No cleavage, red-brown, hardness 3

Maldonite (N) Au2Bi Isometric/1.5-2 V.V.dense, copper-red to black to silver white, Au-qtz veins

Millerite NiS Trigonal/3-3.5/ Grn-Blk

Sml radiating acicular, brassy yellow-to black, V. dense, similar to pyrite.

Molybdenite (R,N) MoS2 Hexagonal/1-1.5/Blue-Blk

V. soft, perfect cleavage, flexible plates, lead grey

Pyrite (R,O,N) FeS2 Isometric/6-6.5/Grn-Blk

Pale brass yellow, cubic xtals, no cleavage.

Pyrrhotite (O,N) Fe1-xS (0<x<0.17) Monoclinic, hexagonal/3.5-4/grey-blk

Bronze to coppery yellow, magnetic when hexagonal.

Realgar As4S4 Monoclinic/1.5-2/orange-red

Red, short prismatic, resinous-greasy luster, smells like garlic when heated

Rustenburgite (TL B)

(Pt,Pd)3Sn2 Isometric/5 Yellow with white tint, granular xtals in matrix

Sphalerite (O,N) ZnS Isometric/3.5-4/brn-white

Yellow-brown, dodecahedral cleavage, resinous lustre

Sulfosalt (Metal,Semimetal)xSx varies Ask mav

Oxides and hydroxides

Formula System/ Hardness Identifying features

Anatase (B) TiO2 Tetragonal/ 5.5-6

Blue-black,red-brown, conchoidal uneven fracture, hard 5-6, dense

Brannerite (R) (U4+,REE,Th,Ca)(Ti,Fe3+,Nb)2(O,OH)6

Monoclinic/4.5-5.5

Black-brownish olive green, conchoidal fract, vit-resinous lustre

Brucite Mg(OH)2 Trigonal/2.5 white-grey, flakey, malible, tabular,. Looks mica-ish

Carnotite (R) K2(UO2)2(VO4)2.3H2O Monoclinic/2 Bright yellow, silky lustre, dense, perfect micaceous cleavage

Chromite (B) Fe(Cr,Al,Fe)2O4 Isometric/5.5 Black-brown-black, minor magnetic, no cleavage, greasy-metallic

Cuprite (B) Cu2O Isometric/3.5-4 Red-translucent, oxidised Cu-mineral, conch. Fracture, cubic xtals


Diaspore (B) AlO(OH) Orthorhombic/6.5-7

White-grey, vitreous, pearly, complete cleavage, alteration mineral

Goethite (B,O) α-Fe3+O(OH) Orthorhombic/5-5.5

Rust, dense

Hematite (B,R,O) Fe2O3 Trigonal/ 5-6

Red-brown, no cleavage, various habits, non-magnetic, hard 5-6

Ilmenite (B,R) Fe2+TiO3 Trigonal/ 5-6

Black-brown black, often twins, no cleavage, brittle. Conchl-uneven frac.

Magnetite (B,O) Fe2+Fe3+2O4 Isometric/5.5-

6.5 Highly magnetic, black, conch frac. No cleagave.

Pyrolusite (B) MnO2 Tetragonal/ 2-6.5

Grey-black, dull and metallic, found in Mn veins, xtals rare

Quartz Moganite

SiO2 Hexagonal/7/ Monoclinic/6

Qtz you know! Moganite is amorphous silica like chalcedony

Rutile (B) TiO2 Tetragonal/ 6-6.5

Brown-redish, metallic, Hard 6, cleavage, prismatic xtals, twins

Spinel (B,R) MgAl2O4 Isometric/ 8

often perfect octohedrons, dense, poor cleavage

Tyuyamunite (R) Ca(Y,REE)(CO3)2(OH).H2O

Orthorhombic/1.5-2

Canary yellow, pearly on cleavage, micaceous fract, waxy-dull massive

Uraninite UO2 Isometric/ 5-6

V.V dense. Radioactive, often metallic black, changes with Pb cont.

Silicates


Actinolite (B,N) {Ca2}{(Mg,Fe2+)5}(Si8O22)(OH)2

Monoclinic/ 5-6

Amphibole group, dark-light green, prisms, fibrous, 2 cleavages @ 600

Aegirine (B,O) NaFe3+Si2O6 Monoclinic/ 6-6.5

Pyroxene group, often in Si under-saturated, green-brown, vitreous

Albite (B) NaAlSi3O8 Triclinic/ 6-6.5

Plag, colourless to white, polysynthetic twins,

Allanite-(La) (B) {Ca,La}{Al2Fe2+}(Si2O7)(SiO4)O(OH)

Monoclinic/ 5.5-6

Dense, submetallic-resinous lustre. Brownish-pitch black, radioactive

Andradite (B,O) Ca3Fe3+2(SiO4)3 Isometric/

6.5-7.5 Garnet, dodecahedral xtals, no cleavage, adamantine lustre

Anothite (B,N) CaAl2Si2O8 Triclinic/ 6-6.5

Plag, colourless to white, parallel or criss crossing twins,

Augite (B,O) (Ca,Na)(Mg,Fe2+,Al,Fe3+,Ti)[(Si,Al)2O6]

Monoclinic/ 5-6

Pyroxene group, intermediate Hedenburgite-diopside, green

Boltwoodite (R) K,Na)(UO2)[HSiO4] · 0.5H2O

Monoclinic/ 3.5-4

Pale yellow-orange-yellow, sub vitreous-waxy, dense, translucent

Clinozoisite (N) {Ca2}{Al3}(Si2O7)(SiO4)O(OH)

Monoclinic/ 6.5

Epidote, dense, elongated prismatic, perfect cleavage, vitreous lustre

Diopside (B) CaMgSi2O6 Monoclinic/ 5-6

Pyroxene, dense, prismatic-granular, radiating agg. xtals, green-brown

Enstatite (B) (Mg,Fe2+)2[SiO3]2 Orthorhombic/ 5.5

Pyroxene group, yellowish-green, fiberous-platey masses. Good cleav.

Epidote (B) {Ca2}{Al2Fe3+}(Si2O7)(SiO4)O(OH)

Monoclinic/ 6-7

Prismatic, columnar xtals, finely striated shiny faces, green-black

Olivine (B) (Mg,Fe2+)2SiO4 Orthorhombic/ 6.5-7

You know! Green, conchoidal fracture.

Haiweeite (R) Ca(UO2)2[Si5O12(OH)2] · 4.5H2O

Monoclinic/ 3.5

Pale yellow, pearly lustre, small radiating fibrous aggregates

Hedenbergite (B,N) CaFe2+Si2O6 Monoclinic/ 5-6

Pyroxene group, oft radiating xtals, dense, green-black, perfect cleav.

Laumontite (B) CaAl2Si4O12 · 4H2O Monoclinic/ 3.5-4

Oft opaque, white, elongated prisms with vert. striations, fibrous-radiating.

Microcline (B) KAlSi3O8 Triclinic/ 6-6.5

K-spar polymorph, white-pink, blue-green, twinning common, 900 cleav.


Muscovite (N) KAl2(AlSi3O10)(OH)2 Monoclinic/ 2-2.5

Mica, flexable elastic flakes, sometime microcrystalline (sericite)

Natrolite (B) Na2Al2Si3O10 · 2H2O Orthorhombic/ 5-5.5

Zeolite, light, white-ish, globular aggs of fibrous radiating needles

Nepheline (B,O) (Na,K)AlSiO4 Hexagonal/ 5.5-6

Feldspathoid, compact granular aggs,white-grey-green, greasy lustre

Phlogopite (B,O) KMg3(AlSi3O10)(OH,F)2

Monoclinic/ 2.5-3

Mica, foliated aggs, light brown-yellow, elastic-flexible-transparent.

Prehnite (B) Ca2Al2Si3O12(OH) Orthorhombic/ 6-6.5

Greenish-white reinform, stalactitic or mammillary aggs translucent vitreous

Sodalite (B) Na8(Al6Si6O24)Cl2 Isometric/ 5.5-6

Compact masses, bright blue, white-grey with green tints. dodec xtals

Talc (N) Mg3(Si4O10)(OH)2 Monoclinic/ 1

Scaly, foliated aggs, greasy feel, perfect cleavage, greenish-white

Titanite (B,R,O,N) CaTi(SiO4)O Monoclinic/ 5-5.5

Sphene, dense, wedge shaped or prismatic, resinous, oft brown-yellow

Tremolite (N) Ca2}{Mg5}(Si8O22)(OH)2

Monoclinic/ 5-6

Amphibole, fibrous asbestiform, SS with actinolite, white-greyish-green

Uvarovite (B) Ca3Cr2(SiO4)3 Isometric/ 6.5-7.5

Garnet, dense, emerald green small xtals, no cleavage,

Zircon (B,R) ZrSiO4 Tetragonal/ 7.5

Stubby, prismatic, occas. dipyramidal xtals, yellow-brown-grey-green

Β-Uranophane (R) Ca(UO2)2(HSiO4)2 · 5H2O

Monoclinic/ 2.5-3

Yellow, vitreous, transparent, translucent, oft small radiating aggs

Sulfates, Molybdates and

Tungstates


Anhydrate CaSO4 Orthorhombic/ 3-3.5

Fibrous masses-breaks into sml cubes, translucent-transparent vitreous-pearly, alters to gypsm

Baryte (B,O) BaSO4 Orthorhombic/ 2.5-3.5

V. dense, tabular xtals, prismatic cleavage, massive, granular, rosettes

Gypsum (B) CaSO4 · 2H2O Monoclinic/ 2

Clear tabular xtals, perfect cleavage into slightly flexible/inelastic flakes

Jarosite (B) KFe3+ 3(SO4)2(OH)6 Trigonal/hexagonal/2.5-3.5

V. sml tabular or pseudo-cubic xtals, brown, assoc. with other rust mins.

Scheelite (N) Ca(WO)4 Tetragonal/ 4.5-5

Adamantine, vitreous, tan-yellow, V. dense, can be comp colour zoned

Wulfenite (B) Pb(MoO4) Tetragonal/ 2.5-3

Adamantine-resinous, yellow-black, V dense, distinct cleavage

Phosphates, Arsenates and

Vanadates


Fluroapatite (B,R,O)

Ca5(PO4)3F Hexagonal/ 5

Colour varies, vitreous lustre, Poor/indistinct cleavage

Monazite-(La) (B) (La,Ce,Nd)(PO4) Monoclinic/ 5-5.5

Brown-yellowish, adamantine resinous, twins common, conch. frac

Carbonates and Halides


Calcite (B,R,O,N) CaCO3 Trigonal/ 3

Any colour, 3 @1200 cleavages, rhoms, vitreous-pearly

Cerussite (B) PbCO3 Orthorhombic/ 3-3.5

V. dense, conch frac, V brittle, white-grey, vitreous-adamintine-resinous

Fluorite CaF2 Isometric/ 4

Purple-blue-pink-brown, cubic cleavage, vitreous/dull when massive

Siderite (B) FeCO3 Trigonal/ 3.5-4.5

Like calcite but red-brown, more dense, harder


Appendix of Useful Figures:

The Moose-o-gram (or Sr isotopes in the Bushveld)


IUGS Quartz-Alkali Feldspar-Plagioclase-Feldspathoid (foid) diagrams

QAPF for volcanic rocks

QAPF for intrusive rocks


IUGS Classification for Mafic Intrusive Rocks

Blevin’s Porphyry Classification

Arco’s Wet & Dry Crystallisation Sequence DRY MAGMAS

chromite

plagioclase clinopyroxene

magnetite

WET MAGMAS spinel

olivine clinopyroxene Ca-plagioclase

amphibole magnetite

The Albite-Water Phase Diagram


Hydrothermal Fluid Source Diagrams

From “Ore-forming processes”

Robb, 2004

Documents

Delia the Dalmatian on ‘Dalmatian Rock’€¦ · Delia the Dalmatian on ‘Dalmatian Rock’ (spotted anorthosite), Eastern Bushveld. Photo courtesy of Lew Ashwal. Acknowledgements