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Towards a magmatic ‘barcode’ for the south-easternmost terrane of the

Kaapvaal Craton, South Africa

by

ASHLEY PAUL GUMSLEY

DISSERTATION

Submitted in fulfilment of the requirements for the degree

of

MAGISTER SCIENTAE

in

GEOLOGY

at the

FACULTY OF SCIENCE

of the

UNIVERSITY OF JOHANNESBURG, SOUTH AFRICA

SUPERVISOR: M.W. KNOPER

CO-SUPERVISOR: M.O. DE KOCK

May 2013

i

DECLARATION

I hereby declare that this dissertation submitted for the Magister Scientae degree to the

Faculty of Science at the University of Johannesburg, apart from the help recognised, is my

own original work and has not been formally submitted in the past, or is being submitted,

for a degree or examination at any other university.

A.P. Gumsley

ii

iii

ACKNOWLEDGEMENTS

“Any system is the sum of its moving parts, and in this work, there is no difference, no matter

how small the part may be; as each bigger part is ultimately composed of a number of

equally critical smaller parts”

First, I wish to thank my two supervisors, Michael Knoper and Michiel de Kock. My

supervisors not only provided me with guidance and patience in my study, but also allowed

me the freedom to pursue my thoughts and feelings with regard to my work and what we

wished to accomplish. I would also like to thank them for their friendship, thoughts and

passion for geology, as well as their constructive criticisms and belief in me while working

on my thesis.

I am indebted to the Department of Geology at the University of Johannesburg, and more

specifically the Palaeoproterozoic Mineralisation Group which provided me with a

scholarship. In addition I wish to thank Richard Ernst, Wouter Bleeker and Ulf Söderlund

who not only provided finances for doing my U-Pb baddeleyite age dating through the

Supercontinent Project (www.supercontinent.org), but also helped with my baddeleyite

separation and TIMS age dating, particulary Ulf. The Jim and Gladys Taylor Trust must also

be thanked for providing me with living expenses while travelling and staying in Sweden

during my analytical work.

I would also like to acknowledge Rajesh Harirajan, Johan Olsson, Herman van Niekerk,

Bertus Smith, Lauren Blignaut, Nic Beukes, George Belyanin, Jan Kramers, Barbara Cavalazzi,

Andrea Agangi, Bryony Richards, Craig McClung, Christian Reinke, Fanie Kruger, Lisborn

Mangwane, Baldwin Tshivhiahuvhi, Diana Khoza, Eve Kroukamp, Herwe Wabo and Hennie

Jonker who all played a role in assisting me throughout my studies, whether it was through

friendship, advice, criticisms or help during my study, I cannot state this enough.

Last but not least I wish to thank my parents, without whose love and support I have

received over these many years, none of this would have been possible.

iv

v

ABSTRACT

The south-easternmost Kaapvaal Craton is composed of scattered inliers of Archaean

basement granitoid-greenstone terrane exposed through Phanerozoic cover successions. In

addition, erosional remnants of the supracrustal Mesoarchaean Pongola Supergroup

unconformably overlay this granitoid-greenstone terrane in the same inliers. Into this crust a

variety of Precambrian intrusions occur. These are comprised of SE-, ENE- and NE-trending

dolerite dykes. Also, the Hlagothi Complex intrudes into Pongola strata in the Nkandla

region, particularly the quartzites of the basal Mantonga Formation. The whole area,

including Phanerozoic strata, has in turn been intruded by Jurassic sills and dykes related to

the Karoo Large Igneous Province. All the rocks of the Archaean inliers, with the exception of

the Jurassic sills and dykes have been subjected to greenschist facies metamorphism and

deformation, with petrographic, Ar-Ar geochronologic and palaeomagnetic studies attesting

to this. This metamorphism and deformation is associated with the Mesoproterozoic

orogeny from the nearby Namaqua-Natal Mobile Belt located to the south. This orogeny has

a decreasing influence with distance from the cratonic margin, and is highly variable from

locality to locality. However, it is generally upper greenschist facies up to a metamorphic

isograd 50 km from the craton margin. Overprints directions seen within the palaeomagnetic

data confirm directions associated with the post-Pongola granitoids across the region and

the Namaqua-Natal Mobile Belt.

The dolerite dykes consist of several trends and generations. Up to five different

generations within the three Precambrian trends have potentially been recognised. SE-

trending dykes represent the oldest dyke swarm in the area, being cross-cut by all the other

dyke trends. These dykes consist of two possible generations with similar basaltic to basaltic

andesite geochemistry. They provide evidence of a geochemically enriched or contaminated

magma having been emplaced into the craton. This is similar to SE-trending dolerite dyke

swarms across the Barberton-Badplaas region to the north from literature. In northern

KwaZulu-Natal the SE-trending dolerite dyke swarms have been geochronologically,

geochemically and paleomagnetically linked to either ca. 2.95 or ca. 2.87 Ga magmatic

events across the Kaapvaal Craton.

vi

The 2866 ± 2 Ma Hlagothi Complex is composed of a series of layered sills

intruding into Nkandla sub-basin quartzites of the Pongola Supergroup. The sills consist of

meta-peridotite, pyroxenite and gabbro. At least two distinct pulses of magmatism have

been recognised in the sills from their geochemistry. The distinct high-MgO units are

compositionally different from the older Dominion Group and Nsuze Group volcanic rocks, as

well as younger Ventersdorp volcanic rocks. This resurgence of high-MgO magmatism is

similar to komatiitic lithologies seen in the Barberton Greenstone Belt. It is indicative of a

more primitive magma source, such as one derived from a mantle plume. A mantle plume

would also account for the Hlagothi Complex and the widespread distribution of magmatic

events of possible temporal and spatial similarity across the craton. Examples include the

layered Thole Complex, gabbroic phases of the ca. 2990 to 2870 Ma Usushwana Complex,

and the 2874 ± 2 Ma SE-trending dykes of northern KwaZulu-Natal already described above

and dated herein. A generation of NE-trending dolerite dykes in northern KwaZulu-Natal can

also be palaeomagnetically linked to this event with either a primary or overprint direction.

Flood basalts seen within the upper Witwatersrand and Pongola Supergroups (i.e., Crown,

Bird, Tobolsk and Gabela lavas) may also be related. This large, voluminous extent of

magmatism allows us to provide evidence for a new Large Igneous Province on the Kaapvaal

Craton during the Mesoarchaean. This new Large Igneous Province would encompass all of

the above mentioned geological units. It is possible that it could be generated by a short-

lived transient mantle plume(s), in several distinct pulses. This plume would also explain the

development of unconformities within the Mozaan Group. This is reasoned through thermal

uplift from the plume leading to erosion of the underlying strata, culminating in the eruption

of flood basalts coeval to the Hlagothi Complex. Marine incursion and sediment deposition

would occur during thermal subsidence from the plume into the Witwatersrand-Mozaan

basin. This magmatic event also assists in resolving the apparent polar wander path for the

Kaapvaal Craton during the Meso- to Neoarchaean. Between existing poles established for

the older ca. 2.95 Ga Nsuze event, to poles established for the younger ca. 2.65 Ga

Ventersdorp event, a new magnetic component for this ca. 2.87 Ga magmatic event can be

shown. This new component has a virtual geographic pole of 23.4° N, 53.4° E and a dp and

dm of 8.2° and 11.8° for the Hlagothi Complex, with a similar magnetic direction seen in one

generation of NE-trending dolerite dykes in the region. This new ca. 2870 Ma addition to the

magmatic barcode of the Kaapvaal Craton allows for comparisons to be made to other

vii

coeval magmatic units on cratons from around the world. Specific examples include the

Millindinna Complex and the Zebra Hills dykes on the Pilbara Craton. Precise age dating and

palaeomagnetism on these magmatic units is needed to confirm a temporal and spatial link

between all the events. If substantiated, this link would assist in further validating the

existence of the Vaalbara supercraton during the Mesoarchaean.

After the Hlagothi Complex event, different pulses of magma can be seen

associated with the Neoarchaean Ventersdorp event. A generation of NE-trending dolerite

dykes in the region was dated herein at 2652 ± 11 Ma. In addition, a primary Ventersdorp

virtual geographic pole established in Lubnina et al. (2010) from ENE-trending dolerite dykes

was confirmed in this study. This ENE-trending dolerite dyke has a virtual geographic pole of

31.7° S, 13.6° E and a dp and dm of 7.0° and 7.2°. This date and virtual geographic poles

from NE- and ENE-trending dolerite dyke swarms in northern KwaZulu-Natal match up with

NE- and E-trending palaeostress fields seen in the Neoarchaean Ventersdorp and proto-

Transvaal volcanics by Olsson et al. (2010). Both generations of dolerite dykes also

demonstrate variable geochemistry. The NE-trending dolerite dyke swarm is tholeiitic, and

the ENE dolerite dyke swarm is calc-alkaline. In addition, some of the tholeiitic NE-trending

dolerite dykes have a similar magnetic component to NE-trending dolerite dykes much

further to the north in the Black Hills area according to Lubnina et al. (2010). This magnetic

component is also similar to the Mazowe dolerite dyke swarm on the Zimbabwe Craton. The

NE-trending dolerite dykes in the Black Hills area differ geochemically from those in northern

KwaZulu-Natal though, but are also of ca. 1.90 Ga age. The Mazowe dolerite dyke swarm

was linked to the dyke swarm of the Black Hills dyke swarm through palaeomagnetic studies.

The Mazowe dolerite dyke swarm however is geochemically similar to the NE-trending

dolerite dykes of northern KwaZulu-Natal, creating greater complexity in the relationship

between the three dyke swarms. It is clear from the complex array of dolerite dyke swarms

and other intrusions into these Archaean inliers of northern KwaZulu-Natal, that much more

work on the dykes within the south-easternmost Kaapvaal Craton needs to be done. This will

resolve these complex patterns and outstanding issues with regard to their palaeo-tectonic

framework.

viii

ix

Table of Contents

DECLARATION i

ACKNOWLEDGEMENTS iii

ABSTRACT v

1. CHAPTER: 01 – INTRODUCTION 1 1.1. Statement Of The Problem 1 1.2. LIPs and reconstructions using barcoding and palaeomagnetism 6 1.3. Locality 13 1.4. Methodology 15 2. CHAPER: 02 – GEOLOGICAL SETTING 17 2.1. Regional Geology 17 2.1.1. The amalgamation of the Kaapvaal Craton 17 2.1.2. Meso- to Neoarchaean supracrustal successions and intrusions 20 2.1.3. Palaeoproterozoic supracrustal successions and intrusions 25 2.1.4. The Mesoproterozoic to the Mesozoic 28 2.2. Local Geology 29 2.2.1. The Archaean Basement 32 2.2.2. The Pongola Supergroup 32 2.2.3. The Hlagothi Complex 34 2.2.4. Dyke and sill swarms and provinces 35 3. CHAPTER: 03 – GEOLOGY 41 3.1. Introduction 41 3.2. The Hlagothi Complex 42 3.3. Dolerite Dykes 46 3.3.1. SE-trending dolerite dykes 49 3.3.2. ENE-trending dolerite dykes 52 3.3.3. NE-trending dolerite dykes 53 3.3.4. Dolerite dykes of other ages 56 4. CHAPTER: 04 – PETROGRAPHY 57 4.1. Introduction 57 4.2. The Hlagothi Complex 58 4.3. Dolerite Dykes 61 4.3.1. SE-trending dolerite dykes 62 4.3.2. ENE-trending dolerite dykes 64 4.3.3. NE-trending dolerite dykes 65 5. CHAPER: 05 – GEOCHEMISTRY 69 5.1. Methodology 69 5.2. The Hlagothi Complex 70 5.2.1. Rock Alteration/Classification 70 5.2.2. Magmatic Variation/Affinity 76 5.2.3. Further Characterisation/Tectonic Setting 76 5.3. Dolerite Dykes 80 5.3.1. Rock Alteration/Classification 82

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5.3.2. Magmatic Variation/Affinity 88 5.3.3. Further Characterisation/Tectonic Setting 92 6. CHAPTER: 06 – GEOCHRONOLOGY 97 6.1. Introduction 97 6.2. Ar-Ar Methodology 97 6.3. Ar-Ar Result(s) 98 6.4. U-Pb Methodology 101 6.5. U-Pb Result(s) 102 6.5.1. The Hlagothi Complex 102 6.5.2. Hlagothi Dyke Swarm 104 6.5.3. ‘Rykoppies’ Dyke Swarm 105 7. CHAPTER:0 7 – PALAEOMAGNETISM 107 7.1. Introduction 107 7.2. Methodology 108 7.3. The Hlagothi Complex 108 7.4. SE-trending dolerite dykes 116 7.5. ENE-trending dolerite dykes 119 7.6. NE-trending dolerite dykes 123 8. CHAPTER: 08 – DISCUSSION 127 8.1. Intrusion and metamorphism 127 8.2. Geochemistry and petrogenesis 131 8.3. Correlation to strata-bound igneous units 134 8.3.1. Correlation with the ca. 2.95 Ga Nsuze Group dykes and lavas 135 8.3.2. Correlation with the ca. 2.87 Usushwana and Thole layered complexes, as well as

Mozaan and Witwatersrand lavas 139 8.3.3. Correlation with the ca. 2.65 Ga Ventersdorp dykes and lavas 143 8.3.4. Correlation with the ca. 1.90 Ga Soutpansberg dykes and lavas 148 8.3.5. Dyke swarms of potentially other ages 151 8.4. Tectonic model and a new large igneous province 152 8.4.1. The Nsuze igneous event 153 8.4.2. The Hlagothi igneous event 155 8.4.3. The Ventersdorp igneous event 158 8.4.4. The Soutpansberg-Mashonaland igneous event 160 8.5. Palaeomagnetism 161 8.6. Correlations with the Pilbara Craton 165 9. CHAPTER: 09 – CONCLUSION 167 10. CHAPTER: 10 – REFERENCE(S) 171

APPENDIX: A – SAMPLE LOCALITIES 201

APPENDIX: B – PETROGRAPHY 203

APPENDIX: C – WHOLE-ROCK GEOCHEMISTRY 213

Chapter: 1 – Introduction ___________________________________________________________________________

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Chapter: 1

Introduction

1.1. Statement Of The Problem

Toward the goal of obtaining a magmatic ‘barcode’ and an apparent polar wander path

(APWP) for the Kaapvaal Craton during the Mesoarchaean to Palaeoproterozoic, a

geological, geochemical, geochronological and palaeomagnetic study was initiated on the

mafic intrusive units of the south-easternmost region of the Kaapvaal Craton. This particular

region has received only limited or no such study. These mafic intrusive units include: SE-,

NE- and ENE-trending dolerite dyke swarms, as well as the layered intrusion known as the

Hlagothi Complex. This portion of the Kaapvaal Craton is south of the Swaziland and

Barberton-Badplaas regions.

The Archaean Kaapvaal Craton in southern Africa is found within the eastern half of

South Africa and encompasses most of Lesotho and Swaziland, with its western extension

into Botswana. It is bounded on all sides by younger orogens (see Fig. 1). It is one of the few

cratons in the world that has retained an almost complete Mesoarchaean to

Palaeoproterozoic stratigraphy. This geologic record is also relatively unmetamorphosed

and undeformed, with several of these stratigraphic successions having been dated, such as

the Nsuze and Ventersdorp Supergroups (e.g., Eglington and Armstrong, 2004; Poujol et al.,

2003). The ca. 3.60 to 3.10 Ga Archaean granitoid-greenstone basement exposed in the

eastern part of the Kaapvaal Craton is also intruded by numerous mafic dyke swarms with

different trends, as well as a variety of sill provinces and layered complexes. Until the

reconnaissance study of Hunter and Reid (1987), there had been a complete lack of any

geological studies on which to base interpretations between mafic dyke and sill

emplacement and evolution of the crust on the Kaapvaal Craton. This work has since been

expanded upon by Havenga (1995), Hunter and Halls (1992), McCarthy et al. (1990), Meier

et al. (2009) and Uken and Watkeys (1997).


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Figure: 1 – The Kaapvaal Craton within the context of southern Africa, including all the surrounding belts.

These are in decreasing age: the Limpopo, Kheis / Magondi, Namaqua-Natal and Damara / Sinclair / Ghanzi-

Chobe / Gariep / Cape / Mozambique orogens; along with the Kasai, Rehoboth, Richtersveld and Zimbabwe

cratons and crustal blocks (modified after de Kock, 2007). Numbers 1 to 5 correspond to the Swaziland,

Witwatersrand, Pietersburg, Kimberley and Okwa terranes of the Kaapvaal Craton respectively. The focus area

of this study in northern KwaZulu-Natal, South Africa is denoted in red

Using geochemistry for example, it has been suggested that NE-trending dykes in the

Johannesburg Dome area fed the juxtaposed Klipriviersberg lavas (McCarthy et al., 1990).

SE-trending dykes in the Barberton-Badplaas area fed Nsuze lavas (Hunter and Halls, 1992).

Recently, these studies have been further expanded upon by the work of Klausen et al.

(2010), Lubnina et al. (2010), Olsson (2012) and Olsson et al. (2010; 2011), who focused on

dykes emplaced into the eastern, north-eastern and south-eastern basement of the

Kaapvaal Craton using geochronological, geochemical and palaeomagnetic studies. These

authors stated that there are at least three dominant regional-scale Precambrian dyke

swarms across these areas of the Kaapvaal Craton’s Archaean basement (see Fig. 4). These


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dyke swarms appear to collectively radiate from a centre within the eastern lobe of the ca.

2.06 Ga Bushveld Complex (Olsson et al., 2010; 2011; Uken and Watkeys, 1997). These

dolerite dyke swarms include:

A southerly ca. 2.95 Ga SE-trending dolerite dyke swarm that cuts through the ca. 3.6

to 3.1 Ga granitoid-greenstone terrane (Brandl et al., 2006).

A widespread ca. 2.65 Ga dolerite dyke swarm in the eastern Archaean basement.

Dykes of this swarm are NE-, E- and SE-trending.

A dense and extensive ca. 1.90 Ga N- to NE-trending dolerite swarm that can be

observed to cut across the ca. 2.66 to 2.06 Ga Transvaal Supergroup and Bushveld

Complex (Cawthorn et al., 2006; Eriksson et al., 2006;).

These authors have further stated that SE-trending dykes in the Barberton-Badplaas area

fed ca. 2.95 Ga Nsuze lavas, as was first proposed by Hunter and Halls (1992). In addition,

the radiating NE-, E- and SE-trending dykes in the Black Hills, Rykoppies and Barberton-

Badplaas areas are coeval with ca. 2.65 Ga Allanridge lavas within the Ventersdorp

Supergroup based on geochronological, geochemical and palaeomagnetic constraints.

Olsson et al. (2010) however, also speculated that lava successions within the proto-basinal

fills to the Transvaal Supergroup, such as the Wolkberg and Godwan Groups may be coeval

to these dolerite dykes. Klausen et al. (2010) studied the geochemistry of the different

dolerite dyke swarms and assigned them different geochemical attributes. This study was

elaborated on in Maré and Fourie (2012) within the Barberton-Badplaas area and was

critical on geochemical correlations made by Klausen et al. (2010). Maré and Fourie (2012)

illustrated a much greater geochemical variation with the different trending dyke swarms,

with different trends overlapping considerably geochemically. Klausen et al. (2010), Lubnina

et al. (2010), Olsson (2012) and Söderlund et al. (2010) have all argued for some NE-

trending dykes within the Black Hills area acting as feeders for ca. 1.90 Ga Soutpansberg

lavas. Thus most of these dykes can now be geochronologically, compositionally and

palaeomagnetically matched as potential feeder dykes to volcanic successions which are

known Large Igneous Provinces (LIPs) based on the original definition by Coffin and Eldholm

(1994). These studies now provide better data in order to re-evaluate the tectonic

association between mafic dyke emplacement in the Kaapvaal Craton and the various

Mesoarchaean to Palaeoproterozoic volcanic packages within the stratigraphy of the


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Kaapvaal Craton. This can also be further used to separate out the various dolerite dyke

swarms from each other.

Extensive work was also done on the northern Kaapvaal Craton by Jourdan et al.

(2004; 2006). Jourdan (2004; 2006) specifically addressed Proterozoic ages seen amongst

the presumed Jurassic ESE-trending Okavango dolerite dyke swarm and NNE-trending

Olifants River dyke swarm. These studies showed that the apparent triple junction formed

by radiating dyke swarms of Karoo age is not a Jurassic structure. It rather reflects

weakened lithospheric pathways that can no longer be considered a mantle plume marker

as previously proposed (Jourdan et al., 2004; 2006). It does not preclude the possible

existence of a mantle plume during the Precambrian along the same structure however.

Hanson et al. (2004a) targeted dolerite sills within the Waterberg and Soutpansberg strata

on the Kaapvaal Craton, and related them to the same ca. 1.90 Ga Soutpansberg Large

Igneous Province (LIP) already described above. Olsson (2012) has added more ages to this

same LIP on the north-eastern Kaapvaal Craton from NE-trending dolerite dykes. Hanson et

al. (2011) and Söderlund et al. (2010) have further described the Mashonaland sills on the

Zimbabwe Craton as being related to the greater ca. 1.90 Ga LIP. This LIP appears to be

extensive across both the Zimbabwe and Kaapvaal cratons with ages mostly varying from ca.

1.93 to 1.87 Ga.

The dolerite dyke swarms of the eastern and northern Kaapvaal Craton, and many

greenstone belts terminate at high angles to the current craton margins. This indicates that

the Kaapvaal Craton has been truncated by plate tectonic processes and/or mantle plumes

along rifted margins (Bleeker, 2003; Bleeker and Ernst, 2006). These dolerite dykes provide

potential piercing points with which to match spatially removed cratonic blocks. The

truncation of dyke swarms and greenstone belts can indicate that the Kaapvaal Craton may

have belonged to a larger ‘supercraton’. Bleeker (2003) defined a supercraton as a large

ancestral landmass of Archaean age with a stabilised core that on break-up spawned several

independently drifting cratons. The supercraton which the Kaapvaal was part of must have

been fragmented at different stages during the Mesoarchaean to Palaeoproterozoic. This

was prior to the amalgamation of the Limpopo, Kheis, Namaqua-Natal and Mozambique

mobile belts (e.g., Jacobs et al., 2008). Various possible Precambrian continental

arrangements have been proposed for the Kaapvaal Craton, particulary in the Neoarchaean


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(e.g., Aspler and Chiarenzelli, 1998; Rogers, 1996). The only consensus regarding nearest

neighbours for the Kaapvaal Craton during this time is the linkage with the Pilbara Craton,

called ‘Vaalbara’ (e.g., Cheney, 1996; de Kock et al., 2009; Eriksson et al., 2009; Nelson et al.,

1999; Wingate, 1998). Other possible nearest neighbour links with the Kaapvaal Craton

remain elusive during the Archaean-Palaeoproterozoic. The exception is the Grunehogna

crustal fragment of western Dronning Maud Land in eastern Antarctica (e.g., Basson et al.,

2004; Groenewald et al., 1991). The identification of other blocks formerly adjacent to other

sides of the Kaapvaal in the Mesoarchaean to Palaeoproterozoic is relatively unknown. The

Kaapvaal-Zimbabwe craton connection is known from at least the Palaeoproterozoic (e.g.,

Söderlund et al., 2010), although Hanson et al. (2011) proposed a greater than 2000 km

displacement.

The significance of such correlations and docking histories lies in accurate

palaeogeographic reconstructions. This is essential to understanding the full tectonic

framework for a particular craton or crustal fragment (e.g., Bleeker, 2003; Bleeker and Ernst,

2006; Ernst et al., 2013). This tectonic framework can be used in conjunction with all the

other remaining cratons and crustal fragments. This can help to further validate plate

tectonic models, processes and reconstructions back into the Archaean (e.g., Ernst et al.,

2013). In addition, there is the potential economic significance by tracing metallogenic belts

between crustal fragments. Examples include the Witwatersrand Basin and Bushveld

Complex. The ca. 2.06 Ma Bushveld LIP, is known for its economic platinum group element

(PGE), chromium, and nickel deposits. It follows that satellite intrusions of the Bushveld LIP

should be present on the former nearest neighbours to the Kaapvaal Craton within the

Palaeoproterozoic geologic record. The robust identification of former nearest neighbours

to the Kaapvaal Craton in the Neoarchaean to Palaeoproterozoic is of significance. It will

allow the correlation of geological units, prominent structures and sedimentary basins

between the Kaapvaal Craton and its former neighbours, such as the Pilbara Craton in

western Australia and the Grunehogna Craton of eastern Antarctica. Cratonic magmatic

barcoding in combination with palaeomagnetic studies and geochemistry on mafic

intrusions such as mafic dykes and sills offers a robust approach for producing well-

constrained reconstructions of the past, as discussed in the next section (Ernst et al., 2013).


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1.2. LIPs and reconstructions using barcoding and palaeomagnetism

LIPs were first defined in detail by Coffin and Eldholm (1994) as anomalously large volume

emplacements of predominantly mafic extrusive rock as well as related intrusives into the

crust. Many LIPs contain felsic components too, and sometimes carbonatites, kimberlites

and lamprophyres, which are often neglected (Bryan et al., 2002; Ernst et al., 2013). These

emplacements of magmatic material manifest in supracrustal volcanic successions, forming

flood basalts. The ‘plumbing system’ of such volcanic rocks is typically dyke swarms, sill

provinces and layered complexes, which represent the feeders to the volcanic succession

(e.g., Bryan and Ernst, 2008; Ernst and Buchan, 2001). LIPs have further been defined as

having a spatial extent of at least 0.1 MKm2 (Coffin and Eldholm, 2001). However, most LIPs

are greater than 1.0 MKm2, and are usually up to 10 km thick prior to erosion of the volcanic

pile (Courtillot and Renne, 2003; Eldholm and Coffin, 2000; Ernst et al., 2005). Where little

or no LIP supracrustal volcanic succession remains, minimum areal extent is usually

determined using the coverage of the feeder intrusive plumbing system related to that

particular event, despite the possible uncertainty (e.g., Yale and Carpenter, 1998). This is

particulary true in the Palaeozoic and Precambrian rock record. The bulk of LIPs are

generally transient, being deposited in less than 10 Ma, with most of the volcanism

occurring in less than 1 Ma (e.g., Coffin and Eldholm, 1994; 2005). In some cases, persistent

LIPs may last tens of millions of years and produce hotspot or seamount chains. A lot

depends on melt production rate and nature of the mantle processes involved (e.g., Ernst et

al., 2005 and references therein).

LIPs can be found globally throughout Earth’s history, mostly as continental and

oceanic flood basalts. The LIPs in the Mesozoic and Cenozoic being the best preserved.

Many Mesozoic and Cenozoic LIPs are associated with continental rifting and break-up, and

are commonly seen on volcanic-rifted margins (e.g., Coffin and Eldholm, 1994; 2001; Cox,

1980; Menzies et al., 2002; Storey, 1995; White and McKenzie, 1989). Volcanic-rifted

margins are truncated, thickened crust with LIPs, whereas nonvolcanic-rifted margins have a

LIP-free transition from continental to oceanic crust of normal thickness (Wilson et al.,

2001). Palaeozoic and Proterozoic LIPs are typically deeply eroded and fragmented,


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metamorphosed and deformed, with just their plumbing system usually remaining. LIP

fragments may also be found in the Archaean to a much more limited extent in greenstone

belts. In greenstone belts they occur as highly deformed and thick metamorphosed basalt

packages, with minor komatiite flows (Arndt, 1999; 2003; Bleeker, 2002; Campbell et al.,

1989; Kerr et al., 2000; Nelson, 1998; Tomlinson and Condie, 2001).

Coffin and Eldholm (1994) described LIPs as forming through processes unlike

‘normal’ seafloor spreading and subduction, being the manifestation of mantle-driven

processes, such as mantle plumes. Variations in depth into the Earth and fertility within the

mantle will lead to different amounts of melt, and may be responsible for the variety in size

and duration of LIPs (e.g., Ernst et al., 2005 and references therein). Apart from mantle

plumes, alternative models proposed for LIP development include lithospheric delamination

(e.g., Şengör, 2001), back-arc processes (e.g., Taylor, 1995), a continent over-riding a

spreading centre (Gower and Krogh, 2002), enhanced mantle convection at the edge of a

craton (shallow mantle ‘edge’ convection, e.g., King and Anderson, 1998), lithospheric

fracturing (e.g., Sheth, 1999), melting of fertile mantle (e.g. Anderson, 2005) and bolide

impact (e.g., Boslough et al., 1996). The study of LIPs and their relationship (or lack thereof)

with mantle plumes has intensified since the early 1990s as new information and techniques

have developed. These techniques include integrated mapping and remote sensing, seismic

tomography, ICP-MS trace element geochemistry, petrogenetic and geodynamic modelling,

as well as more precise palaeomagnetic measurements. The significant development in age

dating techniques has also assisted these new methodologies, particularly with regard to Ar-

Ar and U-Pb isotopic measurements.

LIPs can be associated with the breakup of continents (e.g., Bleeker and Ernst, 2006;

Ernst et al., 2013). When breakup does occur, the result is LIP remnants on conjugate

margins of continental crust (e.g. Courtillot et al., 1999; Storey, 1995). Studies of LIPs and

their associated mafic dyke swarms and sill provinces have also allowed further

understanding of the Earth’s palaeogeography through time, as well as the supercontinent

cycle (Bleeker and Ernst, 2006; Ernst et al., 2013). This can assist in the tracing of

metallogenic belts between continents, and also understanding of the evolving Earth system

as a whole (e.g., Bleeker and Ernst, 2006). Little reliable palaeogeographic information is

available for the Earth prior to 250 million years ago, which represents the time of final


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consolidation of the supercontinent Pangaea (Jacoby, 1981). The absence of preserved

oceanic crust that formed before the Pangaean supercontinent, the lack of an abundant a

fossil record before the Cambrian, and palaeomagnetic overprinting complicate plate

reconstructions in the Precambrian according to Bleeker and Ernst (2006) and Ernst et al.

(2013).

Continental crust usually consists of an Archaean cratonic core. Many cratonic cores

are composite, composed of a variety of terranes, which are in turn surrounded by

progressively younger terranes and orogenic belts (e.g., Bleeker, 2003; Ernst et al., 2013).

This crust was embedded in various supercontinental frameworks through time (Bleeker,

2003). However, these supercontinental frameworks become uncertain as determined from

continental geology using ages of granitic intrusions, variable orogenic belts, metamorphism

and deformation of sedimentary basins and piercing points such as structural trends and

dyke swarms (Bleeker and Ernst, 2006; Ernst and Bleeker, 2010; Ernst et al., 2013),

especially against non-distinct margins. This is the case when viewing the uncertainty

surrounding the ca. 1000 to 700 Ma supercontinent of Rodinia (Li et al., 2008).

Reconstructions of Columbia even further back into the Proterozoic (ca. 1800 to 1300 Ma),

and Kenorland from the Neoarchaean to Palaeoproterozoic respectively are, at best,

speculative and require further highly precise geochronology and palaeomagnetism (e.g.,

Ernst et al., 2013; Meert, 2012). Bleeker (2003) stated that several supercratons may have

existed during the Archaean, instead of one single supercontinent. Smaller scale (i.e., non-

global) reconstructions are possible with selected better preserved crustal fragments by

looking at the geology and palaeomagnetic record from one craton to the next. A good

example is the Kaapvaal and Pilbara (or Vaalbara) cratonic connection (Byerly et al., 2002;

Cheney, 1996; Cheney et al. 1988; de Kock et al., 2009; Nelson et al., 1999; Strik et al., 2001;

Trendall et al., 1990; Wingate, 1998; Zegers et al., 1998). Also, the growth of cratons

through looking at the docking histories of individual terranes within a single composite

craton can be established qualitatively, such as with the Superior Craton (Bleeker and Ernst,

2006) and the Kaapvaal Craton (Schmitz et al., 2004).

With modern analytical techniques, reconstructions as far back as ca. 2.7 Ga and

beyond may be possible, which is the age of stabilisation of a significant portion of the

Archaean cratons (e.g., Bleeker, 2003; Bleeker and Ernst, 2006; Ernst et al., 2013). Mafic


- 9 -

dyke swarms are integral parts of LIPs, being emplaced in a small time interval (e.g. Ernst et

al. 2013). These dyke swarms are usually spatially extensive, extending far into the cratonic

hinterland away from its margins. This allows for more distal portions to be less susceptible

to deformation and metamorphism, providing better geochronologic and palaeomagnetic

constraints for ‘key poles’, (e.g., Buchan, 2000; Buchan and Halls, 1990; Ernst et al., 2013;

Halls, 1982). Dykes also make good piercing points, as they are geophysically and

geochemically distinguishable, as well as being sub-vertical, which makes their preservation

insensitive to uplift and erosion, and avoids structural complication. Data gathered from

them thus allow continental fragments to be placed according to Bleeker and Ernst (2006):

At a specific latitude.

At a specific time.

With a known orientation such that piercing points provided by dyke swarms are

satisfied and in a position that optimises general geological continuity prior to break-

up and dispersal, provided they are not heavily altered along old sutures in orogenic

belts.

Baddeleyite (ZrO2) is an excellent geochronometer for mafic rocks, providing both accurate

and precise emplacement ages (Heaman and LeCheminant, 1993). Baddeleyite is also more

susceptible to being pseudomorphed by zircon during metamorphism instead of developing

zones, assuring that the ages reflect emplacement rather than metamorphic overprints

(Heaman and LeCheminant, 1993). The recent improvements in the separation and

recovering of baddeleyite grains from mafic dykes and sills, as well as layered intrusions

have significantly enhanced the success rate in dating both fine-grained and coarse-grained

silica-undersaturated rocks (Söderlund and Johansson, 2002). U-Pb age dating of

baddeleyite is important in providing emplacement ages, and allows for LIP remnants, such

as complex dyke swarm and sill provinces to be interpreted.


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Figure: 2 - Piercing points and craton reconstruction. (a) A hypothetical (super)craton with various geological

elements just prior to break-up. A LIP, with flood basalts and associated intrusions, is emplaced along the

incipient rift. (b) Break-up of the supercraton has spawned two cratons (A and B). As long as both cratons are

not too modified (e.g. South American and African conjugate margins), they are easily fitted together again us:

P-R, the fitting of promontories and re-entries along the rifted margins; PM, general correlation and fitting of

the conjugate passive margins; P1, piercing points and reconstruction of the LIP; P2, piercing points provided

by older sedimentary basins; P3, piercing points provided by an ancient orogenic front or fold-thrust belt; and

P4, the non-precise piercing points provide by orogenic internides. (c) The more general case where further

break-up up has occurred (craton C) and craton margins have been abraded, modified, and differentially

uplifted. Craton B was strongly uplifted and its sedimentary cover has been eroded. Piercing point P3, if still

recognizable as such, has strongly shifted, and an exhumed granitoid belt is unmatched in Craton A. Craton C

was also uplifted, virtually erasing piercing point P2. Dykes related to the LIP, however, remain on all three

cratons and precise age dating (x Ma) yields a critical clue that they might be part of a single event. Primary

palaeomagnetic data may yield additional geometrical clues (North arrows), if not palaeo-latitudes. (d)

Reconstruction of the original supercraton, based only on the precise piercing points and other information

derived from the dyke swarms (after Bleeker and Ernst, 2006)


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Figure: 3 – Magmatic barcode record of the Kaapvaal and Zimbabwe cratons. The width of individual bars

corresponds to the 2σ error in radiometric ages. Arrow depicts possible age match of mafic extrusions and

intrusions (dyke swarms, sill provinces and volcanic successions) between cratons. The record reveals three

post 2.0 Ga magmatic events common to both the Zimbabwe and Kaapvaal cratons, while no matches occur in

pre 2.0 Ga times, in favour of formation of Kalahari after ca. 2.0 Ga (modified after Söderlund et al., 2010)

An appropriate example would be the information than can be gained on deeply eroded

cratons dominated by Archaean rocks with indistinct and/or inter-mixed dyke swarms of

radiating or linear trends (e.g., the Superior Craton: Bleeker and Ernst, 2006; or the

Zimbabwe Craton: Söderlund et al., 2010; see Fig. 2 and 3). Combined with good geological,

geochemical and palaeomagnetic studies, age determinations of mafic intrusions yield

invaluable clues to ancient tectonic settings. They also potentially provide correlations to

supracrustal units such as volcanic successions in the stratigraphic record both within, and

between cratons and crustal blocks. This is critical for obtaining reliable global plate

reconstructions, especially for supercontinents and cratons of the past (e.g., Bleeker and

Ernst, 2006; Buchan et al., 1996; Ernst et al., 2013; Halls, 1982;). Ernst et al. (2013, and


- 12 -

references therein) state that using precise and accurate age matches among short-lived

mafic magmatic events is a first and highly efficient filter to identify whether two cratons,

which may now be distant, may have possibly been adjacent pieces of crust (see Fig. 3).

Such possible LIP events are usually associated with continental break-up, thus creating

conjugate margins (e.g., Bleeker and Ernst, 2006). These LIP events lead to the creation of a

unique mafic magmatic ‘barcode’ for each individual craton or terrane, which can be

compared and correlated to that of other cratons or terranes (Bleeker and Ernst, 2006).

“Several precisely dated magmatic events provide a ‘fingerprint’ or ‘barcode’ for a craton”

stated Bleeker (2003), Ernst and Bleeker (2010) and Ernst et al. (2013). Originally continuous

crustal fragments will share essential parts of their barcodes, and matches beyond a single

event will indicate a ‘sharing of events’ (e.g., Bleeker and Ernst, 2006). Obtaining magmatic

barcodes for all of the Archaean cratons worldwide would then greatly assist in continental

reconstructions, in addition to other key attributes such as a geological setting, distribution,

palaeomagnetism and geochemistry (Bleeker, 2003; Bleeker and Ernst, 2006; Ernst et al.,

2013).

In addition, there have been several previous palaeomagnetic studies (as listed in the

global palaeomagnetic database of Pisarevsky, 2005), particulary relevant for Neoarchaean

and Palaeoproterozoic units on the Kaapvaal Craton (e.g., de Kock et al. 2006; 2009; Evans

et al., 2002; Strik et al., 2007). However, the general problem with obtaining precise

Precambrian palaeomagnetic poles, not only in the Kaapvaal Craton but globally, has been

the absence of precise dates on the units studied palaeomagnetically. This problem was

noted by Buchan et al. (2000) who surveyed the available ‘key palaeomagnetic’ poles for

Laurentia and Baltica. These are palaeopoles for which there was a robust palaeomagnetic

direction that:

Averaged secular variation through comprehensive sampling of a single site, in

addition to providing data from multiple sites within an acceptable statistical error.

Demonstrated that the direction was primary on the basis of a field test, such as a

baked contact or fold test.

Had a precise age date of better than ± 20 Ma obtained through U-Pb or Ar-Ar

geochronology.


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Of the hundreds of poles published for Precambrian units, only a handful are considered

‘key poles’, and of these, the majority are on the Canadian shield. A reliable Mesoarchaean

to Palaeoproterozoic Apparent Polar Wander Path (APWP) for the Kaapvaal Craton is poorly

constrained. It must be said that the late Palaeoproterozoic APWP segment of the Kaapvaal

Craton is known to some extent (e.g., de Kock et al, 2006; Evans et al., 2002; Hanson et al.,

2004a; 2011)

1.3. Locality

The inliers of Archaean crust of northern KwaZulu-Natal in South Africa form part of the

south-easternmost terrane of the Kaapvaal Craton, and are the focus of this study (see Fig.

4). The largest of these inliers is the White Mfolozi inlier. These inliers consist of granitoid-

greenstone basement including the Nondweni and Ilangwe greenstone belts and fragments,

as well as the White Mfolozi and Nkandla portions of the supracrustal Pongola Supergroup.

Mafic- to ultramafic intrusions into the south-easternmost inliers of the Archaean crust

includes the Hlagothi Complex, which is the most prominent and well-studied. The Hlagothi

Complex is composed of a series of layered sills which intrude the Mesoarchaean Nsuze

Group of the Pongola Supergroup (du Toit, 1931; Groenewald, 1984, 1988, 2006). These sills

correlate well in terms of age and composition with the Thole Complex (Groenewald, 2006),

as well as gabbroic portions of the Usushwana Complex located further to the north. These

units potentially represent feeders to volcanic units within the upper Witwatersrand and

Pongola Supergroups. In addition, three large-scale dolerite dyke swarms of SE-, ENE- and

NE-trends may be found. However, sub-parallel dyke swarms within these dolerite dyke

trends may reveal the presence of additional dyke swarms in the area. These dyke swarms

may also be exploiting former lines of weakness in precursor dykes such as was noted in

Jourdan et al. (2006) for Jurassic aged dykes in the Okavango dolerite dyke swarm. The

dolerite dykes of northern KwaZulu-Natal cross-cut the granitic basement and associated

greenstone belts and fragments, as well as the Pongola Supergroup in some cases. Limited

work on the dykes was carried out by Klausen et al. (2010) and Lubnina et al. (2010), who

correlated them with dykes seen further to the north as was noted above. These authors


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carried out limited geochemical and palaeomagnetic studies within this region of the craton

on selected dolerite dykes.

Figure: 4 - Simplified geological map of the Kaapvaal Craton illustrating the exposed Archaean basement and

supracrustal successions relevant to this study, modified after Robb et al. (2006). The bottom right box is the

outline of the south-easternmost window of the craton, and forms the focus area for this study. Intrusions into

the eastern Kaapvaal Craton are demarcated in red. From north to south: Black Hills 2.65 and 1.90 Ga NE-

trending dykes, Rykoppies 2.65 E-trending dykes, Barberton-Badplaas 2.95 and 2.65 Ga SE-trending dykes and

1.90 NE-trending dykes, the 2.99 Ga Usushwana Complex and the 2.87 Thole Complex. Lastly the Hlagothi

Complex, SE-, NE- and ENE-trending dykes in the bottom right box in northern KwaZulu-Natal


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Lubnina et al. (2010) assigned ca. 2.95 Ga, 2.65 Ga and 1.90 Ga ages to the SE-, ENE- and NE-

trends respectively using palaeomagnetic studies. Klausen et al. (2010) however, assigned

ages of ca. 2.95 Ga and 1.90 Ga to the ENE- and NE-trending mafic dykes based on

geochemistry. The inliers of Archaean crust are overlain by the Natal Group and Karoo

Supergroup sedimentary rocks in which the above mentioned intrusions are not present.

The whole area, however, is in turn intruded by Jurassic SSE-trending dolerite dykes and sills

related to the breakup of Gondwana.

1.4. Methodology

This study focuses on the dyke and sill swarms and provinces on the south-easternmost

portion of the Kaapvaal Craton in northern KwaZulu-Natal, South Africa south of the

Swaziland and Barberton-Badplaas region. This south-easternmost window into the

Kaapvaal Craton has received little attention due to a variety of reasons including remote,

rugged terrane, complex geology associated with metamorphism and deformation due to

the proximity of the cratonic margin to the south with the Namaqua-Natal Mobile Belt. The

aim of this thesis is to obtain new palaeomagnetic and geochemical results on the intrusions

precisely dated during this study, following the same methodologies employed above by

Klausen et al. (2010), Lubnina et al. (2010) and Olsson et al. (2010).

Targets for dyke, sill and layered complex samples were identified using geological

and geophysical maps, as well as remote sensing imagery from Google Earth©. In addition,

this data was digitised on maps. This was used to calculate cumulative segment lengths and

strikes of dykes, which were plotted on histograms at 5° strike intervals in order to resolve

the dominant strike trends, following Klausen et al. (2010).

Dyke and sill samples were gathered from across the south-easternmost Kaapvaal

Craton in northern KwaZulu-Natal south of Vryheid and north of Eshowe. Geological field

relationships were studied, as was petrography, geochemistry and geochronology within

each intrusive unit. In general, several samples were gathered from each dyke and sill

outcrop, usually large or composite samples up to 1 kg. The coarser-grained central parts of


- 16 -

dykes and sills were preferentially sampled, particulary for geochronology and

geochemistry. Usually between 6 and 8 samples from these sampling sites were also taken

for palaeomagnetic studies, with the dyke centre to contact sampled. The contact zone with

the country rock, as well as the host rock itself was also sampled to produce a baked contact

test. Samples were taken from localities where the effects of alteration and weathering

were minimal, or were able to be mostly removed in the field. All sample processing and

preparation was done at the University of Johannesburg’s Department of Geology. The

exception was U-Pb baddeleyite separation, which was done at Lund University in Sweden.

Chapter: 2 – Geological Setting ___________________________________________________________________________

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Chapter: 2

Geological Setting

2.1. Regional Geology

The ca. 3.60 to 3.10 Ga Kaapvaal Craton is one of the world’s oldest and best preserved

granite-greenstone cratonic terranes. Most of the craton’s formation took place during the

Palaeo- to Mesoarchaean (see Fig. 5). It is located within South Africa, and occupies most of

Lesotho and Swaziland, with its north-western extension into Botswana. It is bounded on all

sides by younger orogens. It has a surface area of approximately 1.2 x 106 km2, with a

relatively cool, reduced lithosphere that extends down to a depth between 250 and 300 km

(de Wit et al., 1992). The Kaapvaal Craton has a near complete geologic record of

sedimentation and volcanism preserved within its Mesoarchaean to Palaeoproterozoic

stratigraphy (Hunter et al., 2006). Supracrustal successions cover more than 85% of the

basement. However, along the eastern and southeastern margins, large exposures of the

Archaean granitoid-greenstone basement can be seen. This basement hosts numerous mafic

to ultramafic intrusions, and in particular dyke swarms which are useful for magmatic

barcoding and palaeomagnetic studies.

2.1.1. The amalgamation of the Kaapvaal Craton

The Kaapvaal Craton formed during a series of Palaeo- to Mesoarchaean orogenic stages

according to de Wit et al. (1992), Eglington and Armstrong (2004) and Poujol et al. (2003).

Several granitoid sub-domains amalgamated, along with a number of greenstone belts

around a > 3600 Ma nucleus, which is the Ngwane gneiss of the Ancient Gneiss Complex in

north-western Swaziland. The process is akin to modern day plate tectonics (e.g., de Wit et

al., 1992; Eglington and Armstrong, 2004; Poujol et al., 2003).


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Figure: 5 - Simplified geological map of the Kaapvaal Craton illustrating the exposed Archaean basement and

supracrustal successions relevant to this study, modified after Robb et al. (2006). The bottom right rectangle is

the outline of the south-easternmost window of the craton, and forms the focus area for this study

Mantle xenoliths in kimberlites and their associated diamonds suggest the lithospheric keel

of the craton developed at approximately 3.5 Ga (Irvine et al., 2012). Its evolution can be

linked to four tectonic and geochronologically distinct domains or blocks. These include: the

ca. 3.6 to 3.1 Ga Swaziland block (of which the Ngwane gneiss is a part), the ca. 3.3 to 3.0 Ga

Witwatersrand block, the ca. 3.2 to 2.7 Pietersburg block and the ca. 3.0 to 2.7 Ga Kimberley

block (Eglington and Armstrong, 2004; Poujol et al., 2003). The most reliable age constraints


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for the Ngwane gneiss of the Ancient Gneiss Complex are zircon ages of 3644 ± 4 Ma on

tonalitic gneiss (Kröner and Compston, 1988), and 3663 ± 1 Ma on a nearby outcrop

(Schoene et al., 2008). Zircon xenocrysts within the Ancient Gneiss Complex have yielded

older ages of 3702 ± 1 Ma and 3683 ± 10 Ma (Kröner and Tegtmeyer, 1994; Kröner et al.,

1996), and may indicate an older crustal component. A variety of greenstone belts and a

series of granitoid intrusions amalgamated upon the Ancient Gneiss Complex up until ca.

3.10 Ga when the craton began to stabilise, forming a small proto-continental block (de Wit

et al., 1992). Among the numerous greenstone belts of the Kaapvaal Craton, the ca. 3.5-3.2

Ga Barberton Greenstone Belt is the largest, oldest and most rigorously investigated (e.g.,

Armstrong et al., 1990; Kamo and Davis, 1994; Kröner and Compston, 1988; Kröner et al.,

1996). Various tectonic regimes have been proposed for the Barberton Greenstone Belt,

with many proponents favouring the idea that the belt formed through subduction

processes similar to modern plate tectonic regimes (e.g., de Wit et al., 1992; Diener et al.,

2005; Dziggel et al., 2002; Lowe, 1994). However, other studies argue vertical tectonics

played a role, stating it is a structural complex of many discrete terranes, which were fused

together in tectonic stacking (e.g., Kröner et al., 1996), or formed through foundering (e.g.,

van Thienen et al., 2004), plumes and gravitational collapse (e.g., Anhaeusser, 2001) and

extensional orogenic collapse resulting in fault-bounded basins (e.g., Diener et al., 2005;

Kisters and Anhaeusser, 1995;). The Barberton Greenstone Belt, along with the numerous

other greenstone belts, and early granitoids were then stitched together by more potassic

granitoid batholiths of varying ages, predominantly at ca. 3230 and ca. 3090 Ma.

The craton is sub-divided into four major terranes as stated above (see Fig. 4), where

the presumed Archaean micro-continents amalgamated on to the oldest south-eastern

Swaziland block from the west and north (Eglington and Armstrong, 2004). The terrane

collision and suturing between the Swaziland block and the Witwatersrand block (forming a

greater Witwatersrand block) played an important role in the development of the

Mesoarchaean supracrustal basins (de Wit et al., 1992; Schmitz et al., 2004). In addition, the

collision between this greater Witwatersrand block and Kimberley block along the Colesberg

lineament to the west assisted. As the Dominion lavas erupted on the central half of the

craton and the Nsuze lavas on the eastern side, the craton continued to grow along the

northern and western margins. These juvenile arcs incorporating the Amalia, Kraaipan and


- 20 -

Madibe greenstone belts accreted during the protracted terrane collision between the

Witwatersrand and Kimberley blocks along the north-south Colesberg Lineament in the west

between ca. 2.93 and 2.88 Ga (Schmitz et al., 2004). The Kaapvaal Craton was already joined

with the Pietersberg terrane along the east-west Thabazimbi-Murchison lineament in the

north during this time according to de Wit et al. (1992).

2.1.2. Meso- to Neoarchaean supracrustal successions and intrusions

The Meso- to Neoarchaean stratigraphy and geochronology for the Kaapvaal Craton is

summarised in Table 1 (see Fig. 6). Fragments of a relatively complete Mesoarchaean to

Palaeoproterozoic stratigraphic sequence are remarkably well-preserved on the Kaapvaal

Craton. The stabilisation of the craton allowed a fundamental transition in style of crustal

evolution as supracrustal successions began to form. The oldest preserved of which is the

rift dominated 3074 ± 6 Ma volcanic–sedimentary Dominion Group on the centre of the

craton (Armstrong et al., 1991; Marsh, 2006). This succession represents the oldest cratonic

sedimentary basin on Earth, and although its present outline represents only a structural

remnant of the original depository, it may have occupied an area in excess of 320 000 km2

(Lowe and Tice, 2007). The lower Nsuze Group of the Pongola Supergroup on the eastern

side of the craton can possibly be correlated with the Dominion Group (Cole, 1994). Both

are rift-fill successions of volcanics and sediments. Both also bear testimony to the uplift

and erosion of the Kaapvaal Craton prior to the onset of sedimentation. The Kaapvaal

Craton then underwent further uplift and erosion, followed by thermal subsidence due to

post-rift/plume cooling. This led to the erosion of the majority of the Dominion Group

before the deposition of the proximal and shallow marine West Rand Group of the

Witwatersrand Supergroup, and the distal deep marine lower Mozaan Group (McCarthy,

2006). Both supergroups can be almost completely correlated bed for bed (Beukes and

Cairncross, 1991). Specific to this study, the 2914 ± 8 Ma Crown lava (Armstrong et al.,

1991), as well as the Bird lavas of the Witwatersrand Supergroup and the Tobolsk and

Gabela lavas from the Mozaan Group represent the best evidence for a Witwatersrand-

Mozaan correlation (see Fig. 7). Both sub-basins reflect upward-coarsening sequences.


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Table: 1 - Summary of reported ages for the Archaean supracrustal volcanic successions and related intrusions

Extrusives Intrusives Age Error System Mineral/Whole Rock Reference

Dominion lavas

3074 6 U-Pb zircon Armstrong et al. (1991)

Agatha lavas

3090 90 U-Pb zircon Burger and Coertze (1973)

3083 150 Rb-Sr whole rock Burger and Coertze (1973)

2985 1 U-Pb zircon Hegner et al. (1994)


2883 69 Rb-Sr whole rock Hegner et al. (1984)

2980 10 U-Pb zircon Mukasa et al. (2013)

2977 5 U-Pb zircon Nhleko (2003)



2934 114 Sm-Nd whole rock Hegner et al. (1984)

Barberton-Badplaas Dyke Swarm

2980 1 U-Pb baddelyeite Olsson (2012)

2967 1 U-Pb baddelyeite Olsson et al. (2010)


Usushwana Complex



2875 40 Rb-Sr, Sm-Nd whole rock Layer et al. (1988)

2871 30 Sm-Nd whole rock Hegner et al. (1984)

2870 38 Rb-Sr whole rock Davies et al. (1970)

2386 58 Ar-Ar pyroxene Layer et al. (1988)

2377 58 Ar-Ar pyroxene Layer et al. (1988)

2094 54 Ar-Ar amphibole Layer et al. (1988)

Crown lavas


Tobolsk lavas


Klipriviersberg/Derdepoort/Khanye lavas

2781 5 U-Pb zircon Wingate (1998)

2788 2 U-Pb zircon Moore et al. (1993)

2785 2 U-Pb zircon Moore et al. (1993)

2784 1 U-Pb zircon Grobler and Walraven (1993)

2782 2 U-Pb zircon Walraven et al. (1996)

2769 2 U-Pb zircon Walraven et al. (1994)


Goedgenoeg/Makwassie lavas

2733 4 U-Pb zircon de Kock et al. (2012)


2693 +60/-59 Rb-Sr whole rock Walraven et al. (1987)

2643 80 Rb-Sr whole rock van Niekerk and Burger (1978)

Rietgat lavas

2724 6 U-Pb zircon de Kock et al. (2012)

Rykoppies Dyke Swarm













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Figure: 6 – Cumulative Archaean stratigraphy for the Kaapvaal Craton, compiled by combining maximum

thickness estimates from various sources after Klausen et al. (2010) and references therein

The Witwatersrand sub-basin shows shallow marine deposits progressively overlain by more

fluvial braided river deposits (McCarthy, 2006). The lower stratigraphy of the Witwatersrand

Supergroup comprises the shale, banded iron formation and sandstone-dominant West

Rand Group. The upper stratigraphy consists of the sandstone and conglomerate-dominant

Central Rand Group. Two flood basalt successions occur in the Witwatersrand sub-basin,

with the 2914 ± 8 Ma Crown lava at the top of the West Rand Group, and the Bird lava in the

middle of the Central Rand Group (Armstrong et al., 1991). The origins of the lavas are

somewhat enigmatic, with changes in tectonic regime or plume magmatism having been

suggested according to Frimmel et al. (2005) and Nhleko (2003). Numerous tectonic settings

have been proposed for the Witwatersrand basin, with a foreland basin the preferred. The

Central Rand Group was deposited in a much smaller and restricted basin dominated by

uplift to the west and north (e.g., Myers et al., 1990).

The age of the Pongola Supergroup is poorly constrained, with an upper age given by

the 3107 +4/−2 Ma granitoid basement beneath the Pongola Supergroup (Kamo and Davis,

1994). A lower age limit is given by a cross-cutting granitoid intrusion and quartz-feldspar

porphyry, dated at 2863 ± 8 Ma and 2837 ± 5 Ma respectively (Gutzmer et al., 1999;

Reimold et al., 1993). There are also a variety of ages between 2985 ± 1 Ma and 2934 ± 114

Ma for the Nsuze (Agatha) lavas (Hegner et al., 1984, 1994; Mukasa et al., 2013 and

references therein; Nhleko, 2003).


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Figure: 7 - Stratigraphy of the Dominion, West Rand and Central Rand, Nsuze and Mozaan Groups (modified

after Duncan and Marsh, 2006; Gold, 2006; McCarthy, 2006). Correlation between Witwatersrand and Pongola

strata based on the work by Beukes and Cairncross (1991) and Cole (1994)


- 24 -

There is also the cross-cutting ca. 2990 Ma to 2860 Ma Usushwana Complex (Hunter and

Reid, 1987; Olsson, 2012). The Mozaan Group consists of several formations that lie

unconformably on top of the volcanic and rift-fill sedimentary rocks of the Nsuze Group. It is

composed of alternating sandstones and mudrocks, with iron formations nearer to the

stratigraphic base. The Klipwal diamictite marks the end of deeper marine deposition and a

transition into a shallower marine or continental setting (Gold, 2006).

The Tobolsk, Gabela and Ntanyana lavas occur stratigraphically within the upper

Mozaan basin, and are thought to be flood basalts produced from fissure eruptions in a

continental setting (Hammerbeck, 1982; Nhleko, 2003). The Tobolsk lavas unconformably

overlie the Delfkom (Nconga) Formation with an erosive base (Nhleko, 2003). The Tobolsk

lavas vary in thickness between 50 and 574 m and have been correlated with the Crown

Formation lavas. The Gabela lavas are approximately 150 m thick and have been

stratigraphically matched with the Bird Formation lavas (Beukes and Cairncross, 1991; Gold,

2006). The Ntanyana lavas occur at the top of the Mozaan Group in the Nongoma graben,

and have only be seen in drill core (Nhleko, 2003). They are approximately 64 m thick. Both

the Gabela and Ntanyana lavas consist of various flows intercalated with sediments and

tuffs. These lavas may also be associated with the Bird lavas of the eastern part of the

Witwatersrand basin (Beukes and Cairncross, 1991).

The Pongola Supergroup was in turn intruded by the Usushwana Complex and the

Thole Complex between 2990 Ma and ca. 2860 Ma (Hammerbeck, 1982; Hunter and Reid,

1987; Olsson, 2012). The SE-trending Barberton Badplaas dyke swarm has also recently

been dated between 2966 ± 1 Ma and 2967 ± 1 Ma in the south-eastern region of the

craton. This dyke swarm could possibly represent feeders to the Nsuze volcanic pile (Olsson

et al., 2010).

The Witwatersrand Basin is unconformably overlain by the volcanic-sedimentary ca.

2.70 Ga Ventersdorp Supergroup, in a setting once again dominated by rifting (Eriksson et

al., 2002; Marsh et al., 1992). The lower, middle and upper Klipriversberg, Platberg and Pniel

volcanic sequences respectively appear to have erupted relatively rapidly within

approximately 35 million years. This is shown in zircon ages of 2714 ± 8 Ma and 2709 ± 5 Ma

for the basal basaltic lavas of the Klipriviersberg Group and middle felsic lavas of the


- 25 -

Makwassie Group respectively (Armstrong et al., 1991). Author’s such as de Kock (2007), de

Kock et al. (2012) and Wingate (1998) argue that the significantly older 2782 ± 5 Ma

Derdepoort basalts represent the most basal rocks of the Ventersdorp Supergroup, and that

the 2714 ± 8 Ma age should be regarded as a minimum age. In addition, ages of 2733 ± 4

and 2724 ± 6 Ma on Platberg Group volcanic equivalents (Hartswater Group) by de Kock et

al. (2012). This further casts the 2714 ± 8 Ma and 2709 ± 5 Ma ages into doubt. This study

implies a volcanic history of less than 100 million years instead, from less than 2782 Ma to

greater than 2662 Ma (Cheney, 1998; Eriksson et al., 2002; Wingate, 1998). Possible 2701 ±

11 Ma to 2659 ± 3 Ma feeder dykes were also emplaced along radiating NE-, E- to SE-

trending directions known as the Rykoppies dyke swarm (Olsson, 2012; Olsson et al., 2010).

These were once inferred to be emplaced by the Bushveld Complex (Uken and Watkeys,

1997). A period of sedimentation then occurred after the Allanridge lavas terminated the

infilling of the Ventersdorp basin (van der Westhuizen et al., 1991). It is postulated that the

Ventersdorp basin was composed of NE-trending ‘Basin-and-Range’ type rift structures of

horst and graben blocks. The structures were found across the central and western parts of

the Kaapvaal Craton (Stanistreet and McCarthy, 1991; van der Westhuizen et al., 2006). It is

debated whether the Ventersdorp flood basalts (mainly the Klipriviersberg Group) are

related to the collision between the Kaapvaal and Zimbabwe cratons (Burke et al., 1985;

Light, 1982; Stanistreet and McCarthy, 1991) or to a mantle plume event (Eriksson et al.,

2002; Hatton, 1995).

The whole southeastern region of the Kaapvaal Craton was also intruded by

numerous potassic post-Pongola granites such as the 2863 ± 8 Ma Godlwayo and the 2671 ±

3 Ma Kwetta granites (Reimold et al., 1993).

2.1.3. Palaeoproterozoic supracrustal successions and intrusions

Several groups of mainly clastic sedimentary rocks unconformably overlie the Ventersdorp

Supergroup, with some volcanic rocks near the base (Eriksson and Reczko, 1995). These

groups are exposed as definite sedimentary basin-fills along the northern and eastern

exposed base of the Transvaal Supergroup. They are tentatively interpreted to have been

deposited in a rift environment (Eriksson et al., 2001). Lava in the Buffelsfontein Group of


- 26 -

one of these basin-fills has yielded Rb-Sr ages of ca. 2657-2659 Ma (SACS, 1993) and U-Pb

ages of 2664 ± 1 Ma (Barton et al., 1995), providing a depositional age for these successions.

The Kaapvaal Craton is disconformably overlain by two well-preserved supergroups:

The ca. 2.65 to 2.05 Ga Transvaal Supergroup towards the north, of which the

distribution is wholly confined to the craton.

The ca. 0.32 to 0.18 Ga Karoo Supergroup across the southern extent of the Kaapvaal

Craton, as well as overlying the nearby Namaqua-Natal Mobile Belt.

Both supergroups are comprised of relatively thick and varied sedimentary sequences

capped in each case by rapidly emplaced and extraordinarily voluminous igneous deposits,

i.e., the Bushveld LIP (Cawthorn et al., 2006) and the Karoo LIP (Duncan and Marsh, 2006).

The predominantly clastic and carbonate Transvaal Supergroup overlies the ‘proto-

basinal’ successions within its E-trending Transvaal Basin. It also unconformably overlies the

Ventersdorp Supergroup within the NE-trending Griqualand West basin and lesser Kanye

basin in Botswana (Eriksson et al., 2001). Minor lavas occur in the Abel Erasmus, Ongeluk-

Hekpoort and Machadodorp sequences (Crow and Condie, 1990). The 2222 ± 13 Ma

Ongeluk-Hekpoort lavas are the youngest dated unit in the Transvaal strata (Cornell et al.,

1996), but these lavas are conformably overlain by another few thousand meters of

sedimentary rock. The best age estimate for the base of the Transvaal Supergroup is derived

from detrital zircons in the 2642 ± 2 Ma Vryburg Formation from the Griqualand West sub-

basin. This formation is considered to be correlative with the Black Reef Formation in the

Transvaal sub-basin (SACS, 1980; Walraven and Martini, 1995). The Transvaal Supergroup is

capped and discordantly overlain and intruded by the massive Bushveld Complex (Cawthorn

et al., 2006).

This ca. 2.05 Ga old complex is made up of:

The 2061 ± 2 Ma bimodal Rooiberg Group lavas of basalt and rhyolite (Walraven,

1997).

The world’s largest layered mafic intrusion – the 2058 ± 2 to 2054 ± 1 Ma Rustenburg

layered suite at ca. 65 000 km2, which also hosts the world’s largest deposits of PGE’s,


- 27 -

Cr and V (Lee, 1996; Olsson et al., 2010; Scoates and Friedman, 2008), as well as Cu

and Ni in its satellite intrusions.

The late 2053 ± 12 Ma Rashoop granophyre and 2054 ± 2 Ma Lebowa granite suite

intrusions (Coertze et al., 1978; Walraven and Hattingh, 1993).

The large sill/lopolith that gave rise to the Rustenburg layered suite was likely emplaced

through centralised conduits (e.g., Clarke et al., 2009), as suggested by an absence of

recognised dyke-shaped satellite intrusions. Its emplacement may have occurred roughly at

the time that subduction was occurring along the western margin of the Kaapvaal Craton

prior to a collision with the Congo Craton along the Kheis-Magondi Belt, or even the

Zimbabwe Craton along the Limpopo Belt (e.g. Jacobs et al., 2008). It may also be a plume-

induced LIP, produced through delamination of the underlying lithosphere (e.g., Olsson et

al., 2011). Connected to the Bushveld Complex are smaller sill-, dyke- or plug-like satellite

intrusions in the central parts of the craton, such as the Moshaneng, Uitkomst and Marble

Hall complexes (Anhaeusser, 2006; Cawthorn et al., 2006).

The Bushveld event was superseded by predominantly clastic sedimentation within

an assumed foreland basin. This formed the red-bed sequences of the Waterberg Group.

Subsequent ca. 1.93 to 1.87 Ga igneous activity is recorded on the Kaapvaal Craton as

intrusive dykes and sills into the Waterberg and Soutpansberg groups (Hanson et al., 2004a;

2011). Also the Sibasa and Ngwanedzi lavas within the Soutspansberg Group were produced

in the same event (Barker et al., 2006; Barton, 1979; Crow and Condie, 1990). These

intrusives and lavas may be coeval with the 1928 ± 4 Ma Hartley lavas from the Olifantshoek

Supergroup on the western margin of the craton, which were caught up in the Kheis

orogeny (Cornell et al., 1998; Moen, 2006; van Niekerk, 2006). This ca. 1.90 Ga igneous

event is otherwise recognised as mafic dykes and sills across the northern and eastern

Kaapvaal Craton, as well as covering large parts of the by then attached Zimbabwe Craton

(e.g., Klausen et al., 2010; Olsson, 2012; Söderlund et al., 2010; Stubbs et al., 1999). The

dolerite dyke swarm associated with this event has been termed the ca. 1.90 Ga NE-

trending Black Hills dyke swarm by Olsson et al. (2010) and the Olifants River dyke swarm by

Uken and Watkeys (1997). The timing of collision of the Zimbabwe and Kaapvaal Cratons

along the Limpopo Belt is constrained by this magmatic record (Söderlund et al., 2010). This

is shown by the apparent absence of the ca. 2.05 Ga Bushveld magmatism on the Zimbabwe


- 28 -

Craton, and the corresponding absence of magmatism of the same age as the ca. 2.58 Ga

Great Dyke of Zimbabwe on the Kaapvaal Craton. This suggests that the Zimbabwe and

Kaapvaal Cratons were still separate at these times (Bleeker, 2003). Furthermore, since ca.

1.90 Ga magmatism is present in both Zimbabwe and Kaapvaal Cratons, then collision and

welding of these two cratons must have occurred between 1.90 Ga and 2.06 Ga (Söderlund

et al., 2010). However, palaeomagnetic evidence suggests an almost 2000 km displacement

between the Kaapvaal and Zimbabwe cratons during this time, adding to the complexity of

this LIP (Hanson et al., 2011).

2.1.4. The Mesoproterozoic to the Mesozoic

After the deposition of the Waterberg, Soutpansberg and Olifantshoek successions, the

deposition of a magnificent record of Mesoarchaean to Palaeoproterozoic volcanic/clastic

strata across the Kaapvaal Craton terminated. Successive orogenies occurred around the

Kaapvaal Craton as it was surrounded by active continental margins that led to the

expansion of the craton through crustal accretion. This formed a larger combined Kaapvaal-

Zimbabwe Craton, called the ‘Kalahari Craton’. Crustal growth occurred along the south and

eastern margins during the Mesoproterozoic Namaqua-Natal orogeny (Cornell et al., 2006;

Jacobs et al., 2008). The only record from the interior of the craton during this time comes

from intra-continental alkaline volcanism (Verwoerd, 2006).

Towards the end of the Namaquan Epoch, the interior of the Kalahari Craton

recorded the emplacement of the ca. 1.10 Ga Umkondo LIP across parts of what had now

grown into the Rodinian supercontinent (Hanson et al., 2004b; 2006). Rodinia then began to

break apart between ca. 1.00 and 0.75 Ga. The growth of the Gondwana and Pangaea

supercontinents was then recorded between ca. 0.65 to 0.50 Ga. This is evident in the

Mozambique, Damaran, Cape and Gariep orogenic belts around the Kalahari Craton, as well

as in the Antarctic, then occurred (Jacobs et al., 2008). Finally, the Kaapvaal Craton was

covered by Phanerozoic Karoo Supergroup sedimentary rocks, deposited up to the time of

the Jurassic emplacement of the Karoo LIP (Duncan and Marsh, 2006). The ca. 0.18 Ga Karoo

LIP is related to Gondwana break-up. This was when a vast area that stretches across South

America, southern Africa and Antarctica was affected by voluminous mafic magmatism.


- 29 -

Thus, the eastern Kaapvaal Craton potentially hosts Jurassic (Karoo) dykes that may be

difficult to distinguish from Precambrian dykes. While Karoo dykes and sills have been

distinguished on geological maps as less altered, this may be an unreliable criterion. For

instance, Jourdan et al. (2006) conducted a reconnaissance 40Ar/39Ar study on the NNE-

trending Olifants River dyke swarm in the north-eastern part of the Kaapvaal Craton. That

study revealed, contrary to map classifications, only mafic dykes of Precambrian age. In

general, however, it is suspected that the following intrusions are Jurassic:

N-trending dykes, because these parallel the Lebombo Group monocline (e.g.,

Klausen, 2009).

ESE-trending dykes across the north-eastern part of the Kaapvaal Craton, because

these parallel the Okavango dyke swarm (Jourdan et al., 2006).

All sills and potential feeders to these, near or within the Karoo Supergroup cover

across the southern parts of the eastern Kaapvaal Craton.

This tremendous record of sedimentary and volcanic strata and associated intrusions allows

for a formulation of a magmatic barcode for the Kaapvaal Craton. This enables comparison

against other cratons (see Fig. 8).

2.2. Local Geology

The south-easternmost terrane of the Kaapvaal Craton in northern KwaZulu-Natal hosts

numerous inliers of Precambrian basement (see Fig. 9). This terrane is truncated to the

south and east by the Natal Thrust Front. This marks the boundary between the Kaapvaal

Craton and the Mesoproterozoic Namaqua-Natal Mobile Belt to the south, and the

Neoproterozoic to early Palaeozoic Mozambique Belt to the east, which is covered by

Quaternary sands. The margin of the craton in the area thus is deformed and

metamorphosed. The intensity of this metamorphism and deformation decreases with

distance away from the cratonic boundary (Elworthy et al., 2000). The Precambrian inliers

are overlain by Phanerozoic rocks of the Natal and Karoo sedimentary successions, as well

as intruded by dykes and sills related to the Jurassic Karoo LIP. The inliers preserve the ca.


- 30 -

3.29 Ga to 2.61 Ga granitoid–greenstone basement, which is unconformably overlain by the

Pongola Supergroup (Elworthy et al., 2000; Matthews et al., 1989).

Figure: 8 – Magmatic barcode for the eastern and western sides of the Kaapvaal Craton with ages from a

variety of sources discussed in the text


- 31 -


- 32 -

2.2.1. The Archaean Basement

The basement of the Kaapvaal Craton in the area of study is composed predominantly of the

granitic Anhalt Suite in the White Mfolozi inlier, occupying the whole inlier east of the

Nondweni Greenstone Belt.

Locally, the Anhalt Suite is referred to as the Mvunyana granodiorite (du Toit, 1931;

Hunter, 1990; 1991; Hunter and Wilson, 1988; Hunter et al., 1992). It is intrusive into the ca.

3410 Ma Nondweni Greenstone Belt, having an age of ca. 3290 Ma (Matthews et al., 1989).

Grey tonalitic gneiss has also been documented in the area which predates the Mvunyana

granodiorite. Further south, the granitoids between Babanango and Nkandla remain

undifferentiated. They are however, similar to the Mvunyana granodiorite (Robb et al.,

2006). In addition, a wide variety of granitoids occur adjacent to the Ilangwe Greenstone

Belt near the margin with the Natal Thrust Front. These consist of early to late post-

Nondweni granitoids, gneisses and migmatites according to Mathe (1997). They become

increasingly metamorphosed and deformed closer to the Natal Thrust Front. Numerous

greenstone belt fragments have also been identified in the Buffalo River Gorge, the

Empagneni, as well as near the Nzimane areas and in the Mfule Gorge. Post-Pongola

granitoids were identified in northern KwaZulu-Natal, with the isolated Nzimane granite

having been dated at 2733 ± 3 Ma (Thomas et al., 1997). Further inliers of undifferentiated

granitoids that have as yet not been dated occur in the Buffalo River Gorge inlier, as well as

associated with the Empangeni greenstone belt fragments. All units are unconformably

overlain by the Pongola Supergroup.

2.2.2. The Pongola Supergroup

The Pongola Supergroup is exposed as inliers that are subdivided into two broad sub-basins

in the region: the bigger Pongola sub-basin and the smaller Nkandla sub-basin. The Nkandla

sub-basin is separated from the main Pongola sub-basin by a palaeo-high according to Cole

(1994). It preserves rocks of the Nsuze Group only. It is structurally more complex, with a

stratigraphy that is different from the main Pongola sub-basin according to Groenewald

(1984). Cole (1994) thought that the stratigraphy was essentially the same, although several


- 33 -

successions were structurally thickened or removed from the strata in the sub-basin. The

White Mfolozi inlier in northern KwaZulu-Natal falls into the main Pongola sub-basin area.

The smaller Archaean inliers in the Nkandla area preserve the Pongola successions falling

into the Nkandla sub-basin (see Fig. 10). The area being studied consists of pre-Pongola

granitoid basement rocks unconformably overlain by north-east dipping Pongola

Supergroup rocks in the main Pongola sub-basin. In the region of the Nkandla sub-basin, the

granitic basement is overlain by Pongola Supergroup rocks of the Nsuze Group only, dipping

gently to the south. Deformation and metamorphism is associated with the nearby

Namaqua-Natal Mobile Belt (Groenewald, 1984).

Figure: 10 – Geological map of the Nkandla sub-basin of the Pongola Supergroup showing the intrusions of the

Hlagothi Complex modified after Groenewald (2006). Sample sites for this study are indicated

The Nsuze Group comprises six formations in the White Mfolozi and Nkandla inliers,

with an average thickness of ~3.5 km: the Mantonga, Nhlebela/Pypklipberg, White Mfolozi,

Agatha, Langfontein, Mkuzane and Ekombe Formations from oldest to youngest according

to Cole (1994). The Mantonga Formation represents initial sedimentation of sandstones,

and minor shales and tuffs on top of the basement granite, and is related to basin formation

and subsidence through rifting. This was followed by continental rift volcanism of the

basaltic and andesitic Nhlebela/Pypklipberg lavas. Volcanism then ceased due to the


- 34 -

formation of a trailing plate margin, with subsidence and transgression depositing shallow

marine sandstones, minor shales and the Chobeni carbonates of the White Mfolozi

Formation (Cole, 1994). The dolomites have been documented as containing stromatolites

and algal mats (von Brunn and Mason, 1977). Uplift and erosion then occurred, after which

the volcanism of the basaltic to rhyolitic Agatha lavas were extruded. A small transgression

led to the deposition of Ntambo shales near the end of volcanism. The end of volcanism is

marked by Roodewal agglomerates and pyroclastics, which are part of the Langfontein

Formation. The Langfontein Formation then continues with sandstones deposited during a

marine transgression from subsidence. The basin then underwent further deepening,

depositing the Mkuzane Formation shales in a deeper marine setting. A further

unconformity seen only in the Nkandla sub-basin marks localised volcanism of the Ekombe

lavas ending Nsuze Group sedimentation and volcanism (Cole, 1994).

The Mozaan Group in the White Mfolozi inlier comprises a sedimentary sequence

consisting essentially of a lower arenaceous and an upper argillaceous sequence. According

to Linström (1987), it oversteps approximately 1 200 m of Nsuze Group strata in a south-

eastwards direction. The characteristics of the Mozaan Group succession in the White

Mfolozi area indicate that it was deposited as fluvial strata. This was followed by subsidence

and deposition of deep marine sedimentation during an overall period of sea-level

transgression (Linström, 1987). The upper sedimentary and volcanic packages of the

Mozaan Group in this region are absent.

2.2.3. The Hlagothi Complex

The Hlagothi Complex intrudes into this region of the craton (see Fig. 4 and 10). It is

composed of a series of layered sills of meta-peridotite, pyroxenite and gabbro. It intruded

into the basal quartzites of the Nsuze Group in the Nkandla sub-basin of the Pongola

Supergroup near the craton margin. These sills were named the ‘Hlagothi Complex’ by du

Toit (1931), who first mapped and described them, as the ‘differentiating products of a

single reservoir’. Groenewald (1984) carried out extensive mapping and petrography

coupled with some geochemistry. Groenewald (1984) noted at least five sills with E-trends.

The sills dip concordantly with the Nsuze beds gently to the south. The complex is scattered


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through small inliers of Archaean crust over 18 km in length and 8 km in width between the

towns of Babanango and Nkandla. All the rocks of these inliers have been deformed through

folding and faulting associated with the Natal Thrust Front. The lithologies appear massive,

except at the contacts with the quartzites where they become schistose. They are

predominantly overlain by undeformed Permian-Carboniferous Dwyka Group diamictites.

Prior to this study, the Hlagothi Complex was believed to be near contemporaneous with

the ca. 2.98 Ga Nsuze Group, with a poorly constrained whole rock Pb-Pb age of between

2980 and 3050 Ma (Hegner et al., 1981).

2.2.4. Dyke and sill swarms and provinces

The eastern areas of the Kaapvaal Craton host the largest exposures of the Archaean

basement of the craton. It hosts numerous intrusions of mafic dyke swarms (see Fig. 4)

which have received the most study until the present. At least three major dyke swarms

were identified from the work of Klausen et al. (2010), Lubnina et al. (2010), Olsson (2012)

and Olsson et al. (2010):

A ca. 2.95 Ga SE-trending swarm in the Barberton-Badplaas area.

A ca. 2.65 Ga NE-, E- and SE-trending swarm which is intermixed with the older 2.95

Ga dykes in the Barberton-Badplaas area, forming a radial pattern.

A ca. 1.90 Ga NE-trending swarm in the Black Hills area, also known as the Olifants

River swarm of Uken and Watkeys (1997).

The Archean basement of the south-easternmost Kaapvaal Craton is intruded by a

number of prominent mafic dyke swarms with several ages and trends. These can be

compared to dyke swarms located further north on the eastern side of the craton (e.g.,

Hunter and Reid, 1987; Olsson et al., 2010; Uken and Watkeys, 1997). There were no

accurate and precise, absolute ages on the eastern Kaapvaal Craton dykes prior to the U-Pb

dating by Olsson et al. (2010; 2011), Olsson (2012) and Söderlund et al. (2010). In the south-

easternmost region, this study concentrates on the dyke swarms limited to the Archaean

inliers of similar trends on which age dating, geochemistry and palaeomagnetism was done

further to the north on the eastern Kaapvaal Craton. This includes:


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NW- to SE-trending dolerite dykes, which may be compared to the same trending

Barberton-Badplaas dyke swarm of Olsson et al. (2010).

SW- to NE-trending dykes, which may be compared to the same trending Black Ridge

dyke swarm of Olsson (2012) and Söderlund et al. (2010) on the eastern Kaapvaal

Craton. This comparison was made by both Klausen et al. (2010) and Lubnina et al.

(2010), despite the almost 1000 km separation.

WSW- to ENE-trending dyke swarm which is comparable to the E-W trending

Rykoppies dyke swarm, for which new ages are available (Olsson, 2012; Olsson et al.,

2010). This comparison was made by Lubnina et al. (2010), but was contradicted by

Klausen et al. (2010), who inferred a relationship to the Barberton-Badplaas dyke

swarm of Olsson et al. (2010).

Prior to the present study, no age dating has been carried out on these dykes

directly, but in one locality a SE-trending dyke is seen to be cross-cut by an ENE-trending

dyke, which is in turn cross-cut by a SSE-trending dyke. The SSE-trending dyke also intrudes

into Phanerozoic Karoo strata, while all the other dyke trends are absent from it. Lubnina et

al. (2010) assigned ages of ca. 2.95, 2.65, 1.90 and 0.18 Ga for the SE-, ENE-, NE- and SSE-

trending dykes of the region respectively based on the cross-cutting relationships and

palaeomagnetic studies. Klausen et al. (2010) contradicted this with a possible age of ca.

2.95 Ga for the ENE-trending dykes using petrography and geochemistry. No other prior

study has been presented for the dykes in the south-easternmost area of the Kaapvaal

Craton. Therefore the following literature review below focuses on the dykes located

further north with which they may be compared (see Fig. 4).

The Barberton-Badplaas area of the Kaapvaal Craton has been more intensively

investigated because it includes the world-famous Barberton Greenstone Belt. The

remarkable high-standing SE-trending dolerite dyke ridges extend across many contrasting

pale granitoids, and these are believed to have been feeding Nsuze lavas within the Pongola

Supergroup (e.g., Hunter and Halls, 1992; Klausen et al., 2010; Lubnina et al., 2010; Olsson

et al., 2010). The assignment of the ca. 2.95 Ga age to the oldest SE-trending dyke swarm is

in agreement with the work of Hunter and Halls (1992) and Uken and Watkeys (1997). These

authors tentatively correlated these dykes with volcanism within the Nsuze Group of the

Pongola Supergroup. Neoarchaean post-Pongola Supergroup granitoids have been observed


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to cut many of the dykes that are orientated sub-parallel to the postulated Pongola Rift of

Weilers (1990). The ca. 2.65 Ga age for the E-trending dykes in the Rykoppies area is in

disagreement with the studies of Anhaeusser (2006) and Uken and Watkeys (1997), who

stated that these dykes represented feeder dykes to the E-W elongated Bushveld Complex.

However, more than one age of dyke events is present in the area for the following reasons:

A SE- to ESE-trending dyke in the Barberton-Badplaas area is of comparable age to the

E-trending Rykoppies dyke swarm located further north. This defines a fanning swarm

of this age in conjunction with coeval NE-trending dolerite dykes observed also in the

Black Hills area (Olsson et al., 2010; 2011)

SE-trending dykes may be subdivided into two geochemically different groups which

may be of different ages (e.g., Hunter and Halls, 1992; Klausen et al., 2010)

In detail, SE-trending dykes have two distinct trends between 120° and 150° which

probably represent different swarms.

In addition, there are sporadic N-trending dykes of presumed Jurassic age, and also more

subdued SW- to NE-trending dykes of likely ca. 1.90 Ga (Söderlund et al., 2010).

The Barberton-Badplaas dyke swarm is ~80 km wide, and can be followed from

Swaziland in the south-east for at least 100 km before disappearing under the cover

sequences of the Transvaal Supergroup in the north-west (Hunter and Halls, 1992; Hunter

and Reid, 1987; Uken and Watkeys, 1997). It has intruded into the Ancient Gneiss Complex,

but also into younger Archean granitoid bodies such as the 3227 ± 4 Ma Kaap Valley pluton

(Kamo and Davis, 1994), the 3212 ± 2 Ma Nelshoogte pluton (York et al., 1989) and the 3105

± 3 Ma Mpuluzi batholith (Kamo and Davis, 1994). Since the dolerite dykes of this swarm are

absent in the 2691 ± 2 Ma Mbabane pluton (Layer et al., 1989), Hunter and Halls (1992)

suggested that the majority of these dykes may have been emplaced between ca. 3.00 Ga

and 2.70 Ga. This is because they are generally cut by post-Pongola (Neoarchaean)

granitoids, which was confirmed by Olsson et al. (2010). This dyke swarm also parallels the

mafic ca. 2.99 to 2.87 Ga Ushushwana Complex (Hunter and Reid, 1987; Olsson, 2012).

Burke et al. (1985) stated that rocks of the Pongola Supergroup have been deposited in a

sub-parallel, NW-trending rift system, with dykes more commonly bifurcating towards the

north-west. This implies a predominant magma movement from the south-east (Hunter and


- 38 -

Halls, 1992). Also, the dykes have roughly similar compositions to Nsuze lavas (Klausen et

al., 2010). The dolerite dykes can be divided into high-Ti and low-Ti types (Hegner et al.,

1993). The high-Ti dykes have geochemical characteristics that resemble the 2984 ± 3 Ma

lavas from the Nsuze Group. It was argued that the low-Ti dykes have a geochemical

signature that is typical of early Proterozoic dykes (Hunter and Halls, 1992). Geochemical

variations (including that of Ti) can also be explained by variable degrees of crustal

contamination and low-pressure pyroxene and plagioclase fractionation of a common

primary basaltic melt (Hunter and Halls, 1992; Olsson et al., 2010). The latter explanation is

more consistent with the inability of Klausen et al. (2010) to make a similar low-Sr/V and

high-Sr/V discrimination between the Hunter and Halls (1992) high-Ti and low-Ti dykes,

respectively. The geochemical study of Maré and Fourie (2012) further contradicted the

inferences made by Klausen et al. (2010), showing even greater geochemical variability.

E-trending dykes appear to concentrate in the Rykoppies area of the eastern

Kaapvaal Craton, in addition to those in the south-easternmost Kaapvaal Craton in northern

KwaZulu-Natal (see Fig. 4). The E- to W-trending Rykoppies Dyke Swarm is more than 50 km

wide and 100 km long. This dyke swarm follows the long axis of the E-trending elongated

Transvaal Supergroup (including the Rustenburg Layered Suite). Relatively few NE-, S- and

SE-trending dykes appear to cross-cut this well-constrained succession. The age of the

sericitised and silicified shear zones in the Rykoppies area is unconstrained (Walraven and

Hartzer, 1986). However, they are oriented conspicuously sub-parallel to the ca. 2.95 Ga SE-

trending Barberton-Badplaas dyke swarm (Klausen et al., 2010; Lubnina et al., 2010; Olsson

et al., 2010). Such a tentative correlation may at first be discounted according to Lubnina et

al. (2010) by the prevalent right-lateral offsets of all ca. 2.65 Ga E-trending dykes along this

shear zone. The question is whether these systematic offsets were caused by a post-

intrusive tectonic shear, or by propagating dykes consistently following a pre-existing shear

zone in right-lateral fashion according to Lubnina et al. (2010). The following observations

convinced Lubnina et al. (2010) that the right-lateral offsets along these ca. 2.65 Ga dolerite

dykes are of a primary intrusive origin:

Three cases of conspicuous dykelets – located consistently to the left of a

corresponding major right-lateral offset – which resemble typical ‘bayonets’ extending

from magmatic dyke offsets.


- 39 -

Dykes do not appear nearly as sheared as the host rock.

The NL-23 sample of Lubnina et al. (2010) is offset by a greater amount than NL-22

across the same fault zone.

Thus, while being aware of the possibility of any secondary deformation, Lubnina et al.

(2010) were confident that samples from the sites within the Rykoppies area are unaffected

by this shear zone. The most prominent dolerite dyke within this swarm is the Rykoppies

dyke itself, which marks the central axis of the dyke swarm and predates the cross-cutting

Timbavati Gabbro, of ca. 1.10 Ga Umkondo age (Burger and Walraven, 1979, 1980;

Hargraves et al., 1994; Hanson et al., 2004b). The dyke swarm appears to be exposed only in

the Archaean basement and is discordantly overlain by the ca. 2.64 to 2.06 Ga Transvaal

Supergroup (Walraven and Martini, 1995). Thus Hunter and Reid (1987) proposed an

Archaean age for the Rykoppies Dyke Swarm. Uken and Watkeys (1997) suggested that the

swarm could be a feeder system to the ca. 2.06 Ga Bushveld Complex. This was not only

because of its E-trend parallel to the long axis of the Bushveld Complex, but also because of

a possible link with dykes and sills within the Transvaal Supergroup (e.g., Sharpe, 1981). This

view was also upheld in a recent compilation of South Africa’s mafic- to ultramafic intrusions

(Anhaeusser, 2006). However, consistent ca. 2.65 Ga ages negate previous inferences

suggesting that E-trending Rykoppies dykes were feeding the neighbouring E-W elongated

Bushveld Complex (Anhaeusser, 2006; Uken and Watkeys, 1997). Instead, this swarm

appears to be radiating (Olsson et al., 2010; 2011), and probably acted as feeders to the

upper (Allanridge) lava formation within the Ventersdorp Supergroup (Klausen et al., 2010;

Olsson et al., 2010).

Satellite imagery and aeromagnetic compilations clearly reveal the predominance of

NE-trending dykes in the Black Hills area, which appear to become denser towards the

Jurassic Lebombo-Mwenetzi-Okavango triple-junction (e.g., Jourdan et al., 2006; Klausen,

2009; Reeves, 2000). However, only Precambrian and no Phanerozoic dykes were recorded

among these in the Jourdan et al. (2006) 40Ar-39Ar reconnaissance study. An apparent kink

from a more E-trending swarm across the Kaapvaal Craton basement to a more northerly

trend across the Transvaal Supergroup is possibly an artefact of the juxtaposition of two

different swarms (Uken and Watkeys, 1997). There being a more E-trending ca. 2.65 Ga

dyke swarm that does not cut the Transvaal Supergroup and a more N-trending and


- 40 -

presumed ca. 1.90 Ga swarm that does (Klausen et al., 2010). This conjugate pattern

between approximately NE-trending dykes is evident throughout the north-eastern part of

the Kaapvaal Craton. The older, more E-trending set of dykes is more metamorphosed and

therefore, less suited for geochronology (Klausen et al., 2010). A ca. 1.90 Ga age for the NE-

trending Black Hills dyke swarm argues for a widespread magmatic event. This magmatic

event is correlatable to sills in the Waterberg Group (Hanson et al., 2004a; 2011), lavas

within the Soutpansberg Group (Barker et al., 2006), and even dykes and sills in

Mashonaland within the juxtaposed Zimbabwe Craton (Hanson, 2004a; 2011; Klausen et al.,

2010; Söderlund et al., 2010).

Chapter: 3 – Geology ___________________________________________________________________________

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Chapter: 3

Geology

3.1. Introduction

Northern KwaZulu-Natal south of the town of Vryheid in South Africa consists mainly of

Permian to Jurassic sedimentary rocks of the Natal and Karoo successions. Inliers of

Precambrian crust, have been exposed through the weathering and erosion of the

Phanerozoic strata as a result of uplift of the southern Africa region since the Cretaceous

(Partridge and Maud, 1987). The area is cross-cut by deeply incised braided and meandering

rivers, which have exposed these Precambrian rocks. To the east, along the coastal plain the

Phanerozoic sedimentary rocks have been in turn been covered by Quaternary sands.

The Precambrian rocks in these inliers consist of the Mesoarchaean to Neoarchaean

granite-greenstone terrane north of the town of Eshowe, with remnants also of the Pongola

Supergroup, which have been truncated to the south by the Natal Thrust Front running east-

west between Eshowe and Melmoth. This front is easily distinguished by an escarpment in

the area forming the boundary between the Archaean Kaapvaal Craton and the

Mesoproterozoic Namaqua-Natal Mobile Belt south of it. The Archaean inliers consist of

granitoids and the Nondweni and Ilangwe greenstone belts and fragments. Overlying this

are scattered remnants of the Pongola Supergroup, with the Nkandla sub-basin in the

vicinity of Nkandla separated from the Pongola sub-basin near Ulundi by a palaeo-high

(Cole, 1994; Groenewald, 1984). The Nkandla sub-basin preserves rocks only of the Nsuze

Group, whereas in the Pongola sub-basin, the complete Pongola sequence can be seen

(Cole, 1994). It is this Archaean crust that is the focus of this dissertation. It is poorly

studied, probably due to limited infrastructure and rugged terrane. The climate also renders

it difficult to work in this area over summer, with its heavy rainfall and high temperatures.

This has also led to a deep weathering profile and a scarcity of outcrop, making mapping

Chapter: 3 – Geology ___________________________________________________________________________

- 42 -

and sampling difficult, except along river section pavements, which are often flooded in

summer however.

Into these Archaean inliers numerous mafic- to ultramafic intrusions can be found in

the form of dyke swarms and layered complexes. These intrusions are useful for magmatic

barcoding and palaeomagnetic studies. These intrusions in this area have received little

study, and are also not mapped in great detail or differentiated. Intrusions into the crust

closer to the Natal Thrust Front in the proximity of the Ilangwe Greenstone Belt have not

been mapped or studied at all, nor have dykes present in the more scattered inliers of

Archaean crust in the Nkandla area. Sampling GPS co-ordinates from this study can be found

in Appendix A.

3.2. The Hlagothi Complex

The isolated and scattered Archaean inliers of the Nkandla area can be reached by turning

off the R68 between Babanango and Melmoth toward the town of Nkandla. The terrain

consists of rolling hills, with incised meandering rivers. These rivers expose the Archaean

basement beneath the mainly Phanerozoic Karoo strata consisting mostly of the Dwyka and

Ecca Groups of diamictite and shale. The inliers are dominated by the structurally complex

Nsuze Group remnants of the Nkandla sub-basin (Pongola Supergroup). Metamorphism and

deformation is also higher in this region, being approximately 50 km away from the Natal

thrust front. Into these inliers of Nsuze Group, the Hlagothi Complex intrudes (see Fig. 11).

This is the only intrusion which has been mapped or sampled in any detail in the area by du

Toit (1931) and Groenewald (1984; 1988; 2006). Outcrops of the granite-greenstone terrane

itself occur in the eastern regions adjacent to the Ilangwe inlier. This inlier is composed of

the Ilangwe Greenstone Belt and associated granitoids. An intrusive syenite has also been

identified and mapped in the vicinity of the Mhlatuze River. No dykes have been mapped or

reported, apart from feeder dykes to the Hlagothi Complex by Goenewald (1984).

Chapter: 3 – Geology ___________________________________________________________________________

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Figure: 11 – The Nsongeni and Hlagothi sheets of the Hlagothi Complex within the Archaean inliers of the

Nkandla area, with sample localities shown (modified after Groenewald, 1984). Sample sites for this study are

indicated, as well as section lines depicted in Figure 12

The Hlagothi Complex consists of alternating layered sills of peridotite, pyroxenite and

gabbro. It has been extensively altered through deformation and metamorphism associated

with the Namaqua-Natal Mobile Belt located approximately 50 km to the south. The

complex is scattered through small inliers of Archaean crust over 18 km in length and 8 km

in width between the towns of Babanango and Nkandla. The work of Groenewald (1984)

noted at least five sills with E-trends that vary in thickness from 50 to 250 m, with a

combined thickness of over 500 m measured along the Nsongeni and Nsuze rivers. The

thickest single sheet is about 200 m thick. The sills dip concordantly with the Nsuze

sedimentary strata gently to the south, varying from 10° in the northern Nsongeni sheet

Chapter: 3 – Geology ___________________________________________________________________________

- 44 -

exposures to 30° in the south within the outcrops of the Hlagothi sheets. However, the sills

do transgress from this accepted dip locally, and across observed bedding. The northern sills

are termed the Nsongeni sheets, and the southern sills the Hlagothi sheets, both of which

are located west of Nkandla. Scattered outcrops also occur east of Nkandla, and have been

termed the Wonderdraai sheets. Further to the east, poorer outcrop exposures are found,

most of which is predominantly along rivers from which the areal extent of the Hlagothi

Complex has been extrapolated. These sills are termed the Mhlatuze sheets. The whole

complex is predominantly overlain by undeformed Permian–Carboniferous Dwyka Group

diamictites (see Fig. 12).

Figure: 12 – Simplified E-W and N-S cross-sections (not to scale) through the various sheets of the Hlagothi

Complex. E-W section represents an 18 km long section from Wonderdraai to the headwaters of the Nsuze

River. The N-S section shows the 8 km long outcrop area along the upper Nsuze River valley (after Groenewald,

1984). HC-04 shows a representative hand specimen of the peridotites, whereas HC-08 shows a representative

hand specimen of the gabbros

Chapter: 3 – Geology ___________________________________________________________________________

- 45 -

Samples HC-01, HC-02, HC-03, HC-04, HC-05, HC-06, HC-07 and AG-I_core were

taken from the ultramafic sills within the Nsongeni sheets (see Fig. 9, 10 and 11). These sills

are dark grey-black, massive and fine- to medium-grained (see Fig. 13a). Along the contact

between the complex and the quartzite of the Nsuze Group, a zone of green, foliated and

fine-grained rock was observed. It was described in the field as talc-chlorite schist. It has

possibly developed along the contact between the complex and the quartzites due to stress

partitioning along a zone of competency contrast. One sample was taken from here, AG-

I_contact (see Fig. 13b and c). The pyroxenite sills are usually less than 10 m thick, and

consist of grey-green, massive and medium- to coarse-grained rocks. They are only observed

directly at the base of the Hlagothi sheets. No pyroxenite sill portions were judged adequate

for sampling due to weathering and/or alteration. The gabbros make up the bulk of the

complex and range from 10 to 70 m in thickness. The contact between the gabbros and the

underlying lithologies was not directly seen, but it was inferred within approximately 10 m.

The gabbros are coarse-grained and massive, being also grey in colour, with samples HC-08

and HC-09 taken from them in the Hlagothi sheets located to the south of the Nsongeni

sheet. In addition, one sample was taken from the Wonderdraai inlier. The Wonderdraai

inliers form part of a series of small, isolated inliers of Nsuze Group, and the Hlagothi

Complex rocks which intrude into them east of the main inliers containing the Hlagothi

Complex. Sample HC-10 was taken here, where a black coarse-grained and massive rock-

type, possibly of the Hlagothi Complex (although its exact relationship is unknown), was

seen to intrude into the Nsuze Group quartzite. It must be noted the contact itself wasn’t

directly visible. Skeletal plagioclase feldspar was clearly visible in hand specimen, as was

scattered grains of sulphides, such as pyrite. This rock has been termed a diorite (see Fig.

13d). Finally, samples AG-D, AG-E, AG-F, AG-G and AG-H were also collected from the

Mhlatuze sheets. These inliers host ultramafic to mafic intrusions which have been

correlated with the Hlagothi Complex located further west. They are also inter-sliced with

greenstone belt fragments and possibly volcanic samples of the Nsuze Group (Groenewald,

1984). Due to poor outcrop exposure which is limited to along river sections, as well as

structural and metamorphic complexities, the exact relationship of the Mhlatuze sheets

with the Hlagothi Complex are unknown. The river section sampled is composed of

alternating fine- to medium-grained gabbros and pyroxenites which range in colour from

dark grey to green, and are massive. In addition, the greater Hlagothi Complex is intruded by

Chapter: 3 – Geology ___________________________________________________________________________

- 46 -

porphyritic dykes which pre-date the Karoo Supergroup sedimentation in the region, and

which may be related to the Namaqua-Natal Mobile Belt to the south.

Figure: 13 – Selected geological features observed in the various Archaean inliers hosting intrusions of the

Hlagothi Complex in the Nkandla area. (a) The peridotite intrusions of the Nsongeni sheets of the complex

seen dipping gently toward the south. (b) The contact between the peridotites of the Nsongeni sheets (right)

and the basal Nsuze Group quartzites (left). (c) The schistose contact of the Hlagothi Complex with the

quartzites of the Nsuze Group. (d) The diorite seen in the Wonderdraai inlier.

3.3. Dolerite Dykes

The White Mfolozi inliers of northern KwaZulu-Natal are the largest inliers of Archaean crust

of the south-easternmost Kaapvaal Craton. The nearest towns in the region are Nondweni,

Chapter: 3 – Geology ___________________________________________________________________________

- 47 -

Melmoth and Vryheid. Nondweni is reached by turning off the R68 from Nqutu, whereas

the R34 from Melmoth to Vryheid bisects the more eastern side of the area. The area is

drained by the White Mfolozi River which has exposed a large area of rolling hills composed

of the Mvunyana granodiorite, as well as the Nondweni Greenstone Belt in the western

areas of the inliers as well. The Pongola Supergroup outcrops to the east and northeast,

which contain both Nsuze and Mozaan rocks. Good exposures of the Pongola Supergroup

can be seen along river sections, and are remarkably well preserved, illustrating a near

complete Pongola stratigraphy according to Cole (1994), with the exception of the

uppermost Mozaan Group. Metamorphism and deformation is far less than in the inliers

located to the south in the Ilangwe Greenstone Belt area. This is due to the White Mfolozi

inliers being at a greater distance from the Natal Thrust Front. The granodiorite to the south

becomes increasingly gneissic, and is undifferentiated. There is also a large amount of young

granitoids present. Pongola Supergroup rocks also occur west of the Ilangwe inliers, as well

as in scattered remnants from inliers south of the Nondweni Greenstone Belt.

These windows into the south-easternmost Kaapvaal Craton in northern KwaZulu-

Natal are more heavily cross-cut by a complex array of dolerite dykes and sills than seen in

areas immediately to the north. Some of these intrusions extend into the overlying Karoo

Supergroup and therefore are almost certainly Jurassic in age, and related to the Karoo LIP.

Dykes in the region have been mapped by Linström (1987), and predominantly show NE-

and ENE-trends. Lesser ~135° and 165° SE-trending dolerite dykes have been noted. The

135° dolerite dyke trend is seen in the north-western areas, and 165° trend seen in the

south-east (herein referred to as SE135 and SE165 dolerite dykes respectively). The ENE-

trending dolerite dykes were only noted in the south-western portions of the inliers, and

consist of a dominant ~075° trend only. The NE-trending dolerite dykes were noted to be

widespread across the whole area, but are predominantly in the north-western portions of

the inliers, and consist of a dominant ~030° dolerite dyke trend (NE030). A minor trend

across the whole region can be found at between 045° and 055° (herein referred to as

NE050). Samples of the dolerite dykes hosted in the oldest (Palaeo- to Neoarchaean),

Swaziland terrane basement granitoids in the area were collected. In general, tonalite-

trondhjemite-granodiorite (TTG) gneisses predominate within the Palaeo- to Mesoarchaean.

A granodiorite-monzonite-syenite (GMS) suite of intrusions makes up most of the late

Chapter: 3 – Geology ___________________________________________________________________________

- 48 -

Meso- to Neoarchaean (Robb et al., 2006). Neoarchaean granitoids are not directly relevant

to the study, except for the fact that for example, most SE-trending mafic dykes do not cut

the post-Pongola granitoids (GMS suites) seen further to the north in the Barberton and

Swaziland region. This is in agreement with this swarm’s syn-Pongola age of ca. 2.95 Ga

(Olsson et al., 2010). The south-easternmost Kaapvaal Craton includes the Nondweni and

Ilangwe greenstone belts. The lava sequences within these are generally too dark to provide

any good contrast to the same cross-cutting mafic dyke swarms that are more readily

mapped in equally old granitoids. The same is true of the mafic lavas from the Pongola. The

width and length of the dykes in the inliers are estimated from outcrop, as well as from

Google Earth© imagery and 1:250 000 geological maps from the Council of Geoscience for

the region (1:250 000 Dundee and Vryheid maps). These estimates should be regarded as

approximate (see Fig. 14). Cross-cutting relationships are described below (see Fig. 15).

Figure: 14 - Cumulative dyke lengths at 5° intervals for the White Mfolozi Archaean inliers on the south-

easternmost Kaapvaal Craton. Histogram (A) was generated for the north-western areas, and (B) for south-

eastern areas.

Chapter: 3 – Geology ___________________________________________________________________________

- 49 -

Figure: 15 – Field relationships seen within the Archaean inliers of northern KwaZulu-Natal, with the yellow

Karoo succession being weathered and eroded to expose the pink Archaean inliers of granitoids and lesser

greenstone belts. The supracrustal Pongola Supergroup is also shown in blue. The green SE-trending dolerite

dykes are cut by green ENE-trending dolerites which in turn are cross-cut by green NE-trending dolerite dykes.

These dykes are metamorphosed at greenschist facies and limited only to the Archaean inliers. The whole area

is then in turn cross-cut by unaltered grey S-trending dolerite dykes

3.3.1. SE-trending dolerite dykes

SE-trending dolerite dykes can be found throughout the White Mfolozi inliers, and are

absent from the overlying Karoo strata, indicating that they are Precambrian age. These

dykes are less visible in satellite imagery than the NE- and ENE-trending dykes within the

area. They generally show no or negative relief, with two possible generations of dolerite

dykes present in the inliers, i.e., two SE-trending sets of dolerite dykes offset by up to 30°

from one another in the northwest and southeast (SE135 and SE165). In one locality, already

noted by Lubnina et al. (2010) and Klausen et al. (2010) within the White Mfolozi River itself,

these dykes have been seen to be cross-cut by both SSE-trending dolerite dykes and ENE-

trending dykes. This makes them potentially the oldest in the inliers (see Fig. 15, and Fig.

17). They have also been noted in some instances to cross-cut the Pongola Supergroup

lithologies (Lubnina et al., 2010). The two different trends of SE-trending dolerite dykes

were studied. The first generation (SE165), represented by samples DY-01, AG-Bc and AG-K

Chapter: 3 – Geology ___________________________________________________________________________

- 50 -

(see Fig. 9), was only noted in the south-eastern portions of the White Mfolozi inliers. Dykes

of this generation are medium- to coarse-grained, green-grey and orientated at 165° with

no obvious fabric. A possible slight plunge of 10° to the vertical to the south-west was noted

in these dykes, and these dolerite dykes were observed to be approximately 10 m wide (see

Fig. 16a). In addition, another dyke generation (SE135) which was braided in nature, with

fine and coarse-grained phases was observed, along with devitrified glassy veinlets of dyke

material seen splaying off from the main dyke. Sample AG-J was taken from one of them.

Generally, they are aphanitic, however, and green, with no observable fabric. These dykes

are orientated at roughly 135° with variable widths of between 10 and 35 m. Xenoliths have

been observed within these dykes, accounting for their braided nature. Xenoliths are more

numerous along contacts between dykes and the country rock (see Fig. 16b). These SE-

trending dykes were preliminarily linked to the ca. 2.95 Ga dyke swarm that is dated near

Badplaas by Olsson et al. (2010).

Figure: 16 – Sample localities for the various SE-trending dolerite dykes. (a) a SE-trending dolerite dyke from

which sample AG-Bc was gathered (b) another SE-trending dyke from which sample AG-K was gathered.

Sample localities are also shown from various SE-trending dolerite dykes with hand sample pictures

Chapter: 3 – Geology ___________________________________________________________________________

- 51 -

Figure: 17 – The intersection of three dolerite dykes in the White Mfolozi inlier within the White Mfolozi River.

A SE-trending dolerite dyke was observed to be cross-cut by an ENE-trending dolerite dyke, which is in turn

cross-cut and off-set by a N-trending Jurassic aged dolerite dyke

Chapter: 3 – Geology ___________________________________________________________________________

- 52 -

3.3.2. ENE-trending dolerite dykes

The ENE-trending dolerite dykes are restricted to the south-eastern areas of the White

Mfolozi inliers (see Fig. 9). Like SE-trending dykes in the region, they are absent from the

overlying Karoo strata. In addition, these dykes are observable in satellite imagery from the

area, as vegetation is attracted to their outcrops; particulary aloes (see Fig. 18). They do not

vary as much in their width and direction as the SE-trending dolerite dykes. As was already

established by Klausen et al. (2010) and Lubnina et al. (2010) within the White Mfolozi River,

these dykes have been cross-cut by SSE-trending dolerite dykes, whereas they have in turn

cross cut the SE-trending braided dolerite dykes (SE135). This makes them possibly the

second oldest observed dykes in the inliers (see Fig. 15, 17, 18 and 19a). Klausen et al.

(2010) and Lubnina et al. (2010) have also observed them to be cross-cut by NE-trending

dolerite dykes. These dykes have been noted by Klausen et al. (2010) to be heavily altered

during metamorphism. They are largely indistinguishable from those within the E-trending

Rykoppies dyke swarm, and may likewise contain variable amounts of basement xenoliths.

These dykes were noted to be grey, massive and fine-grained (see Fig. 19b), with one

locality containing a large number of xenoliths, which are usually less than a centimetre to

greater than a meter long (Klausen et al., 2010).

Figure: 18 – Google Earth image with the strikes of various ENE-trending dolerite dykes indicated along with

sample localities. Sample sites for this study are indicated

Chapter: 3 – Geology ___________________________________________________________________________

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They have also been noted to be preferentially orientated in certain localities, according to

Klausen et al. (2010), at 30° to the vertical. In addition, the dykes were noted to be small,

massive and approximately 5 m in width, as well as orientated at approximately 075°. Sharp

contacts with the Mvunyana granodiorite are common. Sampling was done on these dykes,

with samples AG-Ba, AG-Bb and AG-Cb representing three dykes from the White Mfolozi

inlier (see Fig. 19).

Figure: 19 – Sample localities for the various ENE-trending dolerite dykes. (a) an ENE-trending dolerite dyke

near where a SE- trending dolerite dyke sample AG-J was gathered, (b) another ENE-trending dyke from which

samples AG-Ca and b were gathered. Sample localities with various ENE-trending dolerite dykes also shown

with hand sample pictures

3.3.3. NE-trending dolerite dykes

NE-trending dykes are found mostly in the north-western portions of the White Mfolozi

inlier, and intrude into the Precambrian basement, as well as the overlying Pongola

Supergroup strata. They are absent from the Phanerozoic cover of the Dwyka and Ecca

Chapter: 3 – Geology ___________________________________________________________________________

- 54 -

groups of the Karoo Supergroup. The NE-trending dolerite dykes with 030° (NE030) trends

are remarkably high-standing and wide, whereas the dykes with more subdued 045° to 055°

(NE050) trends are much thinner. These dykes were observed to cross-cut the SE-trending

dykes, making at least some of them younger than this event. ENE- trending dykes have also

been observed in turn to be cross-cut by some of these NE-trending dykes, in particular the

NE030 dolerite dyke trend (Klausen et al., 2010; Lubnina et al., 2010). No other cross-cutting

relationship of these dykes has been observed (see Fig. 15). The NE030 dykes are very

distinct on Google Earth imagery, being greater than 30 m wide, and also being marked by

linear features of outcrop and indigenous trees and shrubs in an otherwise undulating

landscape of grassland (see Fig. 20).

Figure: 20 - Google Earth map with strikes of NE-trending dolerite dykes indicated along with sample localities

Pink to white phenocrysts were observed in the NE030 dolerite dykes. In addition,

these dykes have been described as porphyritic by Klausen et al. (2010), with zoned

plagioclase feldspar phenocrysts ranging from 1 to 8 cm in length, with anorthite cores and

albite rims. The phenocrysts coarsen toward the centre of the dykes. These phenocrysts are

more sub-angular and anhedral near the margins of the dykes and more sub-rounded and

euhedral near the centres of the dykes. In addition, the dykes are not massive, showing

evidence of flow banding with the plagioclase feldspar phenocrysts forming bands varying

between 6 and 60 cm in width. The dykes are coarse-grained and dark grey-green in colour.

Chapter: 3 – Geology ___________________________________________________________________________

- 55 -

Sampling was done on these dykes in the north-western area, where sample AG-A was

taken (see Fig. 21b). These dykes have been interpreted to stretch as far as Swaziland and

Barberton by Klausen et al. (2010). However, a second generation of NE-trending dolerite

dykes was observed at 050° orientation (NE050). These dykes are less than 10 m in width

and are massive, dark-grey and fine-grained, with numerous dykelets splaying off them. No

phenocrysts were observed within these dykes, in contrast to the NE030 dolerite dykes

described above. These dykes were found to be widespread across the whole inlier(s), with

samples DY-02_m and DY-02_s and AG-Ca collected from them (see Fig. 21a).

Figure: 21 – Sample localities for the various NE-trending dolerite dykes. (a) a NE-trending dolerite dyke which

is < 5 m wide and from which sample DY-02 was gathered, (b) another NE-trending dyke from which sample

AG-A was gathered, note the phenocrysts. Both localities are in the north-western portions of the White

Mfolozi inliers. Sample localities with various NE-trending dolerite dykes are also shown with hand sample

pictures

Chapter: 3 – Geology ___________________________________________________________________________

- 56 -

3.3.4. Dolerite dykes of other ages

More pristine S- to SSE-trending dykes can be assumed to be Jurassic in age because they

run parallel to the Rooi Rand dyke swarm along the southern Lebombo monocline (Klausen,

2009). They may also be erratic step-feeders to juxtaposed sills, as was noted by Klausen et

al. (2010). These dykes can be easily distinguished from the altered NE-, ENE- and SE-

trending dolerite dykes, as they have experienced little metamorphism and deformation,

unlike the Precambrian examples. Samples AG-L and AG-M were taken from them near the

Natal Thrust Front in order to compare against the older generations of dolerite dykes,

however no further work was done on them during this study.

Chapter: 4 – Petrography ___________________________________________________________________________

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Chapter: 4

Petrography

4.1. Introduction

Petrographical and mineralogical studies were conducted on the 31 samples taken from a

variety of dolerite dykes and sills across the south-easternmost terrane of the Kaapvaal

Craton, including the Hlagothi Complex. Samples were selected from the two possible

generations of SE- and NE-trending dykes, as well as ENE-trending dykes. The various phases

of the Hlagothi Complex were also studied, including the meta-peridotites and gabbros

taken from the Nsongeni and Hlagothi sheets. The study was done principally by optical

microscopy in both transmitted and reflected light, as well as X-Ray diffractrometry (XRD). In

addition, selected features were studied by scanning electron microscopy (SEM), using

traditional imagery as well as electron dispersive spectrometry (EDS); see Appendix B for

further details.

Based on macroscopic observations in hand sample as well as microscopic

observations from thin and polished sections, the original mineralogy in the dolerite dykes

as well as the Hlagothi Complex was noted to be poorly preserved. In most cases this poor

preservation was due to alteration through metamorphism (or in some cases, weathering),

with a few exceptions. In addition, many of the dolerite dyke samples were too fine-grained

for good optical petrography study. Therefore, XRD was particulary useful in conjunction

with CIPW norms calculated following Cox et al. (1979), from whole-rock geochemistry

presented in the next chapter. Samples from the Mhlatuze and Wonderdraai sheets of the

Hlagothi Complex were omitted due to either being highly mineralised (HC-10), highly

altered (AG-E, AG-F and AG-H) or extremely fine-grained (AG-D and AG-G). The relationship

between these samples and the Hlagothi Complex was cast into doubt due to uncertain field

relationships. This was in addition to anomalous geochemical and palaeomagnetic data


- 58 -

discussed in later chapters. The fresh Karoo dolerite dykes are also not discussed (AG-L and

AG-M).


The various phases of the Hlagothi Complex were studied, including the meta-peridotites

and gabbros taken from the Nsongeni and Hlagothi sheets. Groenewald (1984; 1988; 2006)

proposed a four-fold subdivision of the complex based on petrographic studies, updating

the work of du Toit (1931). Groenewald (1984; 1988; 2006) noted that the Hlagothi Complex

generally consists of three types of lithologies, as well as a chilled margin phase against

upper contacts with the Nsuze Group in the gabbroic phases of the complex. Prior to

alteration and metamorphism, the lower portion of each sill consisted of cumulate

peridotites (feldspathic wehrlite, olivine websterite or lherzolite) and pyroxenites with clino-

and orthopyroxene (websterites). The upper portions consist of gabbros and gabbro-norites

according to Groenewald (1984). Primary minerals were probably olivine, ortho- and

clinopyroxene as well as plagioclase feldspar. Other minerals include magnetite, chromite

and biotite. Groenewald (1984; 1988; 2006) also noted chilled marginal phases along the

upper contacts of the gabbroic phase; but it was not observed in this study. Metamorphism

varies from lower to upper greenschist facies.

The peridotite mineralogy consists of fine- to medium-grained amphibole, chlorite,

talc, magnetite and serpentine; with talc or serpentine and magnetite replacing olivine;

chlorite and amphibole replacing orthopyroxene, and more minor clinopyroxene (see Fig.

22). There are instances where the original subhedral to anhedral clino- and orthopyroxene

are still preserved. Usually only relict textures of the cleavage directions remain, with the

pyroxenes almost always enclosing the olivine grains. Mostly the pyroxene has been

pseudomorphed by amphibole and chlorite. Grain size of the pyroxenes and olivine vary

between 0.2 and 7.0 mm. The amphibole and chlorite are usually intergrown, with the

amphibole forming small acicular and needle-like crystals. This amphibole-chlorite phase

can constitute 50-60 modal % of the peridotite.


- 59 -

Figure: 22 – Petrography of the Hlagothi Complex with (a) and (c) showing the meta-peridotites in plane-

polarised light and cross-polarised light, (b) and (d) showing the meta-gabbros in plane-polarised light and

cross-polarised light. Abbreviations are: ta – talc, cl – chlorite, cpx – clinopyroxene, amp – amphibole, mg –

magnetite and plg – plagioclase feldspar

The layers vary from olivine-rich to olivine-poor (10 to 40 modal % of the rock-type), with 1

to 5 mm subhedral to euhedral grains of olivine having broken down completely to talc or

serpentine, and with magnetite rimming the original cumulus olivine on its boundaries and

in cracks. Talc and serpentine can also be intergrown with each other. The original

plagioclase feldspar is mostly preserved, making up less than 20 modal % of the sample,

although it is highly sericitised and/or saussiritised. It is anhedral and occurs interstitially

within the groundmass. Magnetite and chromite form the minor phases. Two different

generations of magnetite and chromite are recognised: a minor generation of fine-grained

magnetite in the core of former olivine cumulus crystals, and larger blebby grains of

magnetite and chromite rimming the cumulus crystals. Magnetite can make up to 2 modal %


- 60 -

of the lithology, with chromite making up to 1 modal %. Orthoclase feldspar can sometimes

be seen occurring in the groundmass and makes up less than 1 modal % of the lithology.

XRD analysis recognised the presence of actinolite, clinochlore and lizardite/antigorite as

the dominant amphibole, chlorite and serpentine respectively. SEM also revealed the

presence of baddeleyite, ilmenite, pentlandite and biotite as accessory minerals in equal

amounts. All samples have been chloritised and serpentinised. At the contact with the

Nsuze Group quartzites, these peridotite bodies become completely altered into schists.

These schists are composed of talc, chlorite, amphibole and serpentinite, in which all the

relict textures have been completely obliterated. The finer talc, amphibole and chlorite are

typically intergrown, and may reflect the original groundmass of ortho- and clinopyroxene,

with irregular amounts of magnetite and chromite also present. The olivine is typically

replaced by talc, serpentine and chlorite. In weathered outcrop, clay minerals are common.

Between the peridotite and gabbro, small layers and lenses of rather homogeneous

pyroxenite occur. Although most samples were highly altered through weathering, relict

textures and minerals remain in fresher samples. These pyroxenites consist of ortho- and

minor clinopyroxene (which make up almost 70 modal % of the lithology) with plagioclase

feldspar consisting of up to 10% of the samples. The pyroxenites appear considerably less

altered than the peridotite through metamorphism however, with well-preserved crystals

and relict cleavage planes, particulary where amphibole has replaced pyroxene. However,

larger euhedral orthopyroxenes have been completely replaced by amphibole. Finer-grained

lath-like clinopyroxenes as well as equant and subhedral orthopyroxene are chloritised or

uralitised. These amphiboles are commonly intergrown in the groundmass between the

larger amphibole crystals. The orthopyroxene is in far greater abundance (60 modal %), and

typically encloses the clinopyroxene (10 modal %), creating a poikilitic texture. Plagioclase

feldspar grains show evidence of alteration, specifically sericitisation. Magnetite and

ilmenite occur in small amounts, making up less than 2 modal % of the total lithology. XRD

shows that the common amphiboles are actinolite and tremolite, with the presence of some

primary pigeonite, hornblende and augite. Clinochlore is the common chlorite. SEM

revealed chromite, orthoclase feldspar, apatite and baddeleyite.

The gabbros makeup the bulk of the Hlagothi Complex and have experienced the

most extensive alteration of the original pyroxenes and plagioclase feldspar. These minerals


- 61 -

have become, with rare exceptions, a quartz, plagioclase feldspar, amphibole, chlorite and

epidote mineral assemblage (see Fig. 22). The dominant plagioclase feldspar is albite.

Equant to elongate subhedral grains of amphibole are the primary alteration product of the

original orthopyroxenes, making up between 40 and 50 modal % of the gabbroic phase and

ranging between 1 and 15 mm in length. The amphibole is intergrown with chlorite, which

can also occur as an alteration product of plagioclase feldspar as well as pyroxene. Anhedral

plagioclase feldspar appears heavily sericitised, making up around 50 modal % of the

samples, and usually occurring in the groundmass; zoned, 0.2-2.0 mm, equant epidote has

also replaced it, along with about 10 modal % quartz. Intergrowths of plagioclase feldspar

also occur in the quartz. Biotite was also identified in association with the quartz. XRD

showed the main amphibole to be actinolite and hornblende. The hornblende appears

brown and euhedral, with grain sizes up to 2 mm. The plagioclase feldspar, when present, is

labrodorite. Approximately 4 modal % microcline feldspar is present too. Opaque minerals

are less common, and only about 1 modal % magnetite and ilmenite was noted. SEM did

reveal that what little magnetite and chromite is present is titanium-rich. Chromite, apatite

and baddeleyite occur in trace amounts. Groenewald (1984, 2006) noted very fine-grained

rocks within the uppermost gabbro sheets, generally within 4 m of the upper contact. This

fine-grained rock type is a chilled margin with the Nsuze Group quartzites; however, these

were not observed in this study.

4.3. Dolerite Dykes

All the dolerite dyke samples restricted to the Precambrian basement granitoid-greenstone

terrane of the south-easternmost inliers of the Kaapvaal Craton generally show high

concentrations of amphibole, chlorite and plagioclase feldspar in optical microscopy, which

was confirmed by XRD analyses. This type of alteration is indicative of low to intermediate

grades of metamorphism. It could also be possible syn-magmatic alteration through the

interaction of dyke material along the contact with meteoric or magmatic fluids. Almost no

talc was present in the dykes, which could suggest the presence of metamorphic fluids with

less than 10% carbon dioxide. However, this type of alteration is very typical of olivine,


- 62 -

which is not present within any of the dykes. Most samples show sericitised and epidotised

plagioclase feldspar intergrown with orthopyroxene that is replaced by amphibole (usually

actinolite). Fine-grained marginal rocks are typically highly weathered, whereas the dyke

centres are coarser, and much fresher. However, coarse-grained phenocrysts of plagioclase

feldspar can sometimes be seen along the margins of the dykes. Xenoliths and xenocrysts

from the country rock also tend to occur along margins of the dykes. Both phenocrysts and

xenocrysts can provide anomalous geochemical results. Granophyric textures along these

margins can also be common, with high amounts of interstitial quartz derived through

partial melting of the granodiorite host rock.

4.3.1. SE-trending dolerite dykes

Both possible generations of the SE-trending dolerite dykes are medium- to coarse-grained

and variably metamorphosed, with one dolerite dyke, AG-Bc, relatively fresh and unaltered

to a large extent (see Fig. 23). Both SE-trending dyke generations generally have lower

greenschist facies mineral assemblages, making identification of the primary mineralogy

more difficult. Most grains are greater than 0.5 mm in size. These dykes consist of

plagioclase feldspar (40 modal %), amphibole (20-40 modal %) and chlorite (10-20 modal %).

Other common minerals include a small amount of microcline feldspar and variety of

opaques, such as magnetite, which constitutes up to 5 modal % of the samples. Anhedral

plagioclase feldspar is variably sericitised and cloudy, with grain sizes ranging from 0.3-0.8

mm. Amphibole can be fibrous, and is a pseudomorph after orthopyroxene, which has been

uralitised and has grain sizes between 0.5 and 1.0 cm. The amphibole can be euhedral to

subhedral in some cases. Amphibole is often enclosed by large poikilitic plagioclase feldspar,

in which case primary orthopyroxene can be preserved. Chlorite is also common, and can

form as a pseudomorph after clinopyroxene and plagioclase feldspar, forming small granular

aggregates of flakes in the groundmass, which are rarely greater than 0.2 mm. It is usually

intergrown with actinolite or plagioclase feldspar. Various amounts of interstitial quartz can

occur, usually only up to 10 modal %, which may be granophyric, or may occur as

intergrowths with plagioclase feldspar. Under reflected light and SEM, magnetite and

ilmenite are the common opaque minerals.


- 63 -

Figure: 23 – Petrography of the SE-trending dolerite dykes with (a) and (c) showing dyke AG-Bc in plane-

polarised light and cross-polarised light, (b) and (d) showing dyke AG-J in plane-polarised light and cross-

polarised light. Abbreviations are: cpx – clinopyroxene, amp – amphibole, mg – magnetite, and plg –

plagioclase feldspar

Opaques account for less than 10 modal % of the samples, and are usually less than 0.2 mm.

They appear embayed with chlorite, with only skeletal crystals left. Small amounts of biotite,

apatite, baddeleyite, serpentine and epidote have been noted in two of the dykes within the

south-western areas. Indistinguishable material ranges from 10 to 30 modal % depending on

how fine-grained the rocks are. XRD analysis confirmed actinolite and tremolite as the

amphiboles present, clinochlore as the chlorite, and albite as the primary plagioclase

feldspar.


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4.3.2. ENE-trending dolerite dykes

All the dolerite dyke samples restricted to Precambrian basement show high concentrations

of amphibole and chlorite in optical microscopy after pyroxene. This type of alteration is

indicative of low grades of metamorphism. Almost all of the samples show sericitised

plagioclase feldspar intergrown with pyroxene replaced by amphibole (usually actinolite).

Transmitted and reflected light microscopy showed that the ENE-trending dykes are fine-

grained, with the mineralogy difficult to assess in thin section, with most grains usually

being less than 0.5 mm in size. The minerals that were present showed a high degree of

alteration to greenschist-facies mineral assemblages (see Fig. 24).

Figure: 24 – Petrography of the ENE-trending dolerite dykes with (a) and (c) showing dyke AG-Ba in plane-

polarised light and cross-polarised light, (b) and (d) showing dyke AG-Cb in plane-polarised light and cross-

polarised light. Abbreviation are: qz – quartz, and ch – chlorite


- 65 -

These dolerite dykes consist primarily of plagioclase feldspar (10-40 modal %), amphibole

(20-40 modal %) and chlorite (10-20 modal %), as determined with the assistance of CIPW

norms. Other minerals include quartz, microcline, calcite, apatite and various opaques. The

plagioclase feldspar mostly has elongate- to lath-like crystal shapes with a cloudy

appearance indicative of sericitisation, where visible. SEM imagery and EDS analysis showed

that the plagioclase feldspar grains range in size from 0.1-0.6 mm, and can be partially

enclosed by amphibole. The amphibole present is mostly fibrous or acicular, and is almost

certainly a pseudomorph after clinopyroxene, with grain sizes of less than 0.2 mm. Chlorite

can also form as a pseudomorph after clinopyroxene and plagioclase feldspar. Chlorite

forms small flakes (very seldom greater than 0.1 mm) within the plagioclase feldspar, or it

forms large aggregates. Quartz occurs in quantities ranging from 10 to 30 modal % of quartz,

and grain sizes from 0.2-0.5 mm. From 5-10 modal % opaques also occur, which can be up

to 0.2 mm. Indistinguishable material ranges from 10 to 50 modal %. Rarely there is calcite,

sphene and apatite present. XRD analysis confirmed the presence of actinolite as the

amphibole, with a smaller amount of hornblende, which could be a relict of the primary

product, as well as clinochlore as the chlorite and albite as the plagioclase feldspar. SEM

analysis found that the primary opaques include hematite or magnetite, ilmenite and rutile.

4.3.3. NE-trending dolerite dykes

The first described generation (NE030 in chapter 3) of dolerite dykes in chapter 3 consists of

the coarser-grained 030° NE-trending dykes (see Fig. 25a). Like the dolerite dykes

documented above, these dykes consist primarily of plagioclase feldspar, amphibole and

chlorite under optical microscopy, which have been altered during greenschist facies

metamorphism. Grain sizes are coarse, with most grains being greater than 0.5 mm.

Plagioclase feldspar phenocrysts, however, sometimes are greater than 1.0 cm. Plagioclase

feldspar is often sericitised, and comprises between 20-40 modal % of the dykes, and up to

60 modal % of the sample if phenocrysts are included outside of the groundmass. The

phenocrysts are zoned, and show no obvious signs of preferred orientation. Amphibole is

present in abundance. It is fibrous, comprising between 30-40 modal % of the samples, with

grain sizes usually less than 0.8 mm. It is usually a pseudomorph after clinopyroxene, which


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was uralitised. Small amounts of chlorite are also evident, usually up to 10 modal % of the

samples, with grain sizes typically at around 0.2 mm. It is a pseudomorph after clino-

pyroxene too. Quartz is rare; usually less than 5% of the sample, with grains up to 0.7 mm.

Opaques can be up to 10-15 modal % of the dykes, and have grain sizes of 0.4 mm. Some

sulphide mineralisation has been noted, usually pyrite. Indistinguishable material is usually

only about 5-10 modal %. XRD analysis confirmed the presence of actinolite as the

amphibole present, as well as clinochlore as the chlorite and albite as the plagioclase

feldspar. SEM analysis found that the primary opaques are ilmenite and titanium-rich

magnetite and chromite.

Figure: 25 – Petrography of the NE-trending dolerite dykes with (a) and (c) showing dyke AG-A (NE030) in

plane-polarised light and cross-polarised light, (b) and (d) showing dyke DY-02 (NE050) in plane-polarised light

and cross -polarised light. Abbreviations are: plg – plagioclase feldspar, amp – amphibole and mg – magnetite


- 67 -

The second generation of NE-trending dolerite dykes (NE050) is fine-grained and

highly altered at greenschist facies, with most grains usually less than 0.2 mm in size. This

makes it difficult to distinguish the original mineralogy (see Fig. 25b). Assisted by CIPW

norms from whole rock geochemistry, XRD and SEM analytical work, these dykes consist

primarily of plagioclase feldspar (10-40 modal %), amphibole (30-40 modal %), chlorite (5-20

modal %) and quartz (20-30 modal %). Other minerals include various opaques. The

plagioclase feldspar mostly has a cloudy appearance due to sericitisation. It grain size ranges

from 0.1 to 0.4 mm. Identifying the outline of the plagioclase from the surrounding

groundmass proved difficult. One amphibole present is brown hornblende and appears to

be primary, although it may be recrystallised. Other amphiboles are also present, but they

tend to be green-brown and are probably pseudomorphs after pyroxenes. Grain sizes for

amphibole are typically in the range of 0.2 mm. Chlorite forms as a pseudomorph after

pyroxene and plagioclase feldspar. It forms small aggregates in the groundmass and is

extremely fine-grained, being less than 0.1 mm in size. Large amounts of quartz have been

distinguished in quantities ranging from 20 to 30 modal % of the sample. It is usually

xenocrystic and intergrown with plagioclase feldspar. The quartz also usually shows

undulatory extinction and appears to be recrystallised. Grain sizes vary from 0.2- 0.5 mm in

size. The opaques are hard to distinguish, but are probably less than 10 modal % of the

sample. Opaque grains are typically less than 0.1 mm. Indistinguishable material ranges

from 20 to 50 modal %. XRD analysis confirmed the presence of actinolite and hornblende

as the amphiboles present, pigeonite as one of the pyroxenes, clinochlore as the chlorite,

and albite as the plagioclase feldspar. SEM analysis found that the opaques include

magnetite, ilmenite and rutile.


- 68 -

Chapter: 5 – Geochemistry ___________________________________________________________________________

- 69 -

Chapter: 5

Geochemistry

5.1. Methodology

A total of 17 samples from across the Hlagothi Complex in the inliers from the Nkandla sub-

basin of the Pongola Supergroup, were selected to provide a representative spectrum across

the various rock types for geochemistry. However, 6 of these samples provided anomalous

results. In addition, 11 samples were taken from selected SE-, NE- and ENE-trending dolerite

dykes. Sampling was done in such a way as to achieve homogeneity, and to avoid alteration.

For the analyses, samples were cleaned and any weathered/altered areas not removed in

the field were cut off using a diamond rock saw, paying particular attention to using the saw

as sparingly as possible. The cut pieces were then cleaned thoroughly in an ultrasonic bath

and oven-dried overnight. The samples were then crushed by hand, using a large mortar and

pestle, which was cleaned and dried with both water and acetone respectively to limit

possible contamination between each sample. The material was then sieved to collect only

rock fragments ≥ 1 cm3 in size, with no obvious evidence of being cut by the rock saw. The

remainder of the material was discarded because it was either too small or bearing cut

grooves or smears from the rock saw. This was done to limit contamination. The rock

fragments were then cleaned and dried in the ultrasonic bath and oven overnight once

again. This crushed material was then split and powdered for approximately 1 minute within

a chrome-steel ring mill. Between samples, all equipment was thoroughly brushed,

vacuumed, washed, cleaned and dried with water and then acetone. Between all samples,

the ring mill processed pure quartz sand. Following this, an approximate amount of 10 g of

sample was dispatched to ACME Laboratories in Vancouver, Canada (acmelab.com) for

major and trace element analysis. Major and trace element contents for all representative

whole-rock samples were determined using X-Ray Fluorescence Spectrometry (XRF) and

Inductively Coupled Plasma Mass Spectrometry (ICPMS), respectively. XRF was done for

major and minor elements on glass beads prepared from powdered whole-rock samples

with a sample-to-flux ratio of 1:10. Volatiles were determined by loss on ignition. Precision


- 70 -

for the different elements was better than ± 1 % of the reported values. Trace and rare

earth element (REE) compositions were determined using ICPMS. The samples were

dissolved in a 1.5 g lithium meta/tetraborate flux fusion, and the resultant molten bead was

then rapidly digested in a weak dilute nitric acid solution. Precision and accuracy based on

replicate analysis of international rock standards were between 2 and 5% (2σ) for most

elements and ± 10 % for U, Sr, Nd and Ni.

The whole rock geochemical analyses for the SE-, NE-, and ENE-trending dolerite

dykes and the Hlagothi Complex are presented in Table 2 and Table2 3; and additional

geochemical data can be found in Appendix C.

5.2. Hlagothi Complex

A total of eleven samples from across the Hlagothi Complex are presented here to

characterise and provide a broad, representative spectrum of the various rock types of the

complex for geochemistry. Similar fractionation histories for the different layered sills are

likely, and therefore the combined data are considered to provide broad geochemical

characteristics for the whole complex. The sampling is not considered adequate for

petrogenetic modelling, and no sampling of the pyroxenites was done. Samples from the

Wonderdraai and Mhlatuze sheets were omitted due to their uncertain relationships with

the complex and/or anomalous behaviour through alteration from weathering and/or

metamorphism. The whole rock geochemical analyses for the Hlagothi Complex are

presented in Table 2. Available geochemical data presented by Groenewald (1984) and

Hammerbeck (1982) for the Hlagothi Complex and Thole Complex is also shown in the

diagrams.

5.2.1. Rock Alteration/Classification

The alteration box plot of Large et al. (2001) demonstrates geochemical changes through

alteration for the major igneous rock types see in the previous chapter (see Fig. 26). The


- 71 -

gabbros of the Hlagothi Complex fall within the least altered field of basalt and andesite.

The pyroxenite and peridotite groups, however, show significant chloritisation. However, it

must be noted that this alteration diagram of Large et al. (2001) was developed specifically

for rhyolite, dacite, andesite and basalt rock types, and not ultramafic lithologies. This shows

the limitation of this diagram for these lithologies. In addition, while some alkalis may have

been leached from the Hlagothi Complex, all samples have maintained their original rock-

type classifications in Winchester and Floyd (1977) classification diagrams using SiO2, Nb/Y

and Zr/TiO2 (see Fig. 27). SiO2 and Nb/Y appear to be the best discriminators between the

rock types of the complex, with Zr/TiO2 being the least useful. SiO2 is the best for dividing

the peridotitic from the pyroxenitic and gabbroic phases, with Nb/Y good for distinguishing

between the pyroxenitic and gabbroic phases. This suggests that these rocks have not

suffered secondary large ionic lithophile (LIL) element mobilisation due to a clear correct

classification.

Table: 2 – Whole rock geochemical analyses on 11 samples from the Hlagothi Complex

HC-08 HC-09 AG-I_core AG-I_contact HC-01 HC-02 HC-03 HC-04 HC-05 HC-06 HC-07

SiO2 57.15 57.13 43.48 46.45 42.47 42.90 45.13 47.22 44.36 45.92 45.28

Al2O3 13.34 13.26 6.14 6.99 5.52 5.29 8.03 6.23 6.27 5.48 5.88

MnO 0.18 0.18 0.17 0.18 0.17 0.15 0.23 0.19 0.19 0.16 0.19

CaO 8.28 8.36 4.72 6.53 3.38 2.86 6.78 6.35 4.47 4.09 1.10

Na2O 2.55 2.40 0.05 0.05 0.01 0.01 0.12 0.04 0.02 0.03 0.01

K2O 0.87 0.86 0.21 0.02 0.09 0.06 0.02 0.01 0.13 0.12 0.03

Fe2O3T 11.69 11.61 11.50 9.70 12.37 11.89 11.73 10.77 11.83 12.19 13.21

MgO 4.16 4.05 24.86 22.06 26.21 26.65 20.67 22.02 24.13 24.00 24.47

TiO2 0.54 0.55 0.24 0.31 0.22 0.21 0.41 0.21 0.25 0.25 0.24

P2O5 0.07 0.09 0.05 0.05 0.05 0.04 0.07 0.04 0.06 0.05 0.05

Cr2O3 0.02 0.03 0.54 0.53 0.62 0.65 0.41 0.44 0.60 0.50 0.58

LOI 1.00 1.30 7.40 6.60 8.20 8.60 5.80 5.90 7.10 6.60 8.30

Total 99.85 99.82 99.36 99.47 99.31 99.31 99.40 99.42 99.41 99.39 99.34

Cs 0.40 0.30 4.00 0.05 1.70 1.30 0.05 0.05 3.80 2.00 0.50

Rb 22.10 22.50 15.10 1.10 9.00 7.10 1.30 1.00 13.60 11.70 2.60

Ba 214.00 206.00 9.00 2.00 4.00 13.00 5.00 3.00 3.00 3.00 3.00

Th 2.80 3.10 0.70 0.50 0.80 0.50 1.00 0.50 0.90 0.70 0.80

U 0.70 0.70 0.20 0.10 0.10 0.10 0.20 0.10 0.20 0.20 0.20

Nb 3.60 3.80 1.10 1.50 1.10 1.00 1.70 1.30 1.50 1.20 1.50

Ta 0.30 0.30 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10

La 13.50 13.60 4.1 5.9 4.4 3.1 4.9 3.5 3.3 4.6 3.3

Ce 27.60 27.50 8.6 11.8 8.7 6.9 10.2 7.0 7.5 9.6 6.7

Pb 0.90 1.20 1.2 1.3 1.4 0.8 1.4 1.3 1.1 0.8 0.8

Pr 3.15 3.26 1.05 1.54 1.00 0.64 1.27 0.82 0.89 1.16 0.82

Sr 118.20 118.90 31.1 21.9 12.9 10.3 9.7 17.9 19.5 22.6 22.5

Nd 12.20 12.30 3.90 5.90 4.40 3.10 5.80 3.40 3.30 4.80 3.80

Zr 79.70 82.90 27.90 36.60 24.80 24.30 45.90 24.60 31.10 29.30 27.90

Hf 2.10 2.40 0.80 1.00 0.50 0.80 1.20 0.70 0.70 0.80 0.80

Sm 2.44 2.59 1.01 1.38 0.86 0.82 1.42 0.86 0.88 1.08 0.86

Eu 0.68 0.71 0.34 0.35 0.21 0.21 0.36 0.18 0.23 0.26 0.17

Gd 2.65 2.80 1.10 1.54 0.98 0.96 1.74 0.91 1.09 1.12 1.02

Tb 0.50 0.53 0.18 0.26 0.15 0.17 0.30 0.15 0.18 0.17 0.16

Dy 3.12 3.29 1.23 1.70 1.11 1.14 2.03 1.06 1.28 1.18 1.10

Ho 0.72 0.77 0.26 0.35 0.22 0.25 0.39 0.19 0.24 0.24 0.24

Er 2.18 2.28 0.82 1.02 0.69 0.68 1.22 0.58 0.80 0.79 0.76

Tm 0.35 0.35 0.10 0.14 0.09 0.10 0.16 0.08 0.10 0.09 0.09

Yb 2.15 2.20 0.73 0.90 0.69 0.67 1.18 0.60 0.74 0.67 0.62

Y 19.70 20.00 7.00 9.80 6.00 6.50 11.00 5.20 6.80 6.40 6.00

Lu 0.35 0.37 0.10 0.14 0.09 0.10 0.18 0.08 0.11 0.10 0.10

V 262.00 269.00 108.00 121.00 98.00 96.00 155.00 119.00 111.00 119.00 109.00

Trac

e e

lem

en

ts (

pp

m)

Sample

Maj

or

ele

me

nts

(w

t. %

)


- 72 -

The three geochemical groups of the Hlagothi Complex can be identified, in agreement with

the major element classification.

Figure: 26 – Alteration box plot of Large et al. (2001) for the Hlagothi Complex

Complying with the recommendations of the IUGS sub commission of Le Bas et al.

(1986), the majority of samples from the Hlagothi Complex broadly fall within two

subdivisions of either basalts to picro-basalts and basaltic andesites on the total alkali-silica

(TAS) classification diagram (see Fig. 28a), with the basaltic andesites further separated into

two different groupings. Peridotites of the complex fall in the basalt to picro-basalt

grouping. The basalt to picro-basaltic compositions from the more ultramafic phases of the

complex contain <2 wt% Na2O+K2O and 43-51 wt% SiO2. The basaltic andesite group is

represented by the more mafic phases of the complex containing approximately 1 wt%

Na2O+K2O with 54-56 wt% SiO2 for the pyroxenites, 2-4 wt% Na2O+K2O and 53-57 SiO2 wt%

for the gabbros. Geochemically, the peridotites and pyroxenites have well defined ranges

whereas the gabbros are more variable. However, the limited number of samples from the

literature for the pyroxenites must be noted.


- 73 -

Figure: 27 – Trace element classification diagrams of Winchester and Floyd (1977) for the Hlagothi Complex


- 74 -

Figure: 28 - Total alkali-silica classification diagram of Le Bas et al. (1986) for the Hlagothi and Thole complexes

(a), and classification diagram showing the variation in K2O content with respect to SiO2 (b). Total alkali-silica

classification diagram for high-Mg rocks (c) of Le Bas et al. (2000)


- 75 -

The gabbroic portions of the Hlagothi Complex have a medium-potassium content of 0.5-1

wt% K2O (see Fig. 28b). The peridotitic and pyroxenitic portions show low-potassium

content of <0.5 wt% K2O. Another TAS classification diagram of Le Bas (2000) is useful for

high-Mg rocks, and this diagram classifies the peridotites as komatiites due to the fact that

the peridotites have more than 18 wt% MgO and less than 1 wt% TiO2 (see Fig. 28c).

The three geochemical groupings of the Hlagothi Complex are sub-alkaline and

broadly tholeiitic (see Fig. 29a), with the gabbroic phase straddling the tholeiitic to calc-

alkaline boundary. The pyroxenitic and peridotitic phases, however, are more tholeiitic on

the classification diagram of Irvine and Baragar (1971). The same trend is seen on the Jensen

(1976) classification diagram, with the pyroxenitic and peridotitic phases both defined as

komatiitic and the gabbroic phase as tholeiitic (see Fig. 29b).

Figure: 29 – AFM (a) and Jensen (b) classifications diagram of Irvine and Baragar (1971) and Jensen (1976) for

the Hlagothi Complex


- 76 -

5.2.2. Magmatic Variation/Affinity

Major and trace variation diagrams illustrate two distinct populations using Mg number

(Mg#), with no obvious trend line existing between the two groupings for the Hlagothi

Complex. The mafic- to ultramafic lithologies (peridotites and pyroxenites) can be defined as

SiO2-poor and MgO-rich (high Mg#), containing 20-30 wt% MgO, with a Mg# between 60-80.

The mafic to intermediate lithologies (gabbros) are more SiO2-rich and MgO-poor (low

Mg#). This population contains between 5-10 wt% MgO (low Mg#), with the Mg# from 30-

60. SiO2 content is 45-50 wt% for the high Mg# group and between 55-60 for the low Mg#

group. Total Fe as Fe2O3 is constant throughout the complex at between 5-10 wt%, as is

MnO at between 0.15-0.25 wt%. TiO2 content is <0.5 wt% and P2O5 <0.1 wt%. The low Mg#

end members of the Hlagothi Complex are enriched in Al2O3, CaO, Na2O and K2O at 10-15, 8-

12, 1.5-3.5 and 0.5-1.5 wt respectively. The high Mg# end members are depleted at 5-10, 2-

8, 0-1 and 0.0-0.4 wt% respectively for these oxides (see Fig. 30).

Ni and Cr are more enriched in the high Mg# portions of the Hlagothi Complex,

with between 500-2000 ppm for Ni and 2000-5000 ppm for Cr, as they are both compatible

elements. It must be noted however that Cr contamination is likely from the steel ring mill

during sample preparation. The low Mg# portions display between 0-400 ppm for Ni and 0-

1000 ppm for Cr. The low Mg# phases typically contain 150-250 ppm V, 1500-2500 ppm Y,

50-100 ppm Zr, 3-6 ppm Nb, 10-30 ppm Rb and 120-360 ppm Sr. In contrast, the high Mg#

phases have 100-200 ppm V, 500-1500 Y, 25-50 ppm Zr, 1-3 ppm Nb, 0-30 ppm Rb and 0-60

ppm for Sr. In addition, the low Mg# phases display high levels of variability, which the high

Mg# phases do not (see Fig. 31). SiO2, CaO and Ni can be used to distinguish between the

peridotites and pyroxenites.

5.2.3. Further Characterisation/Tectonic Setting

Using the multi-element trace geochemistry for the samples, spider diagrams were

constructed and normalised using data from McDonough and Sun (1995) to either primitive

mantle for the trace elements or C1 chondrite for the rare earth elements (see Fig. 32 and

33).


- 77 -

Figure: 30 – Major element variation diagrams for the Hlagothi Complex


- 78 -

Figure: 31 – Trace element variation diagrams for the Hlagothi Complex


- 79 -

All samples for the Hlagothi Complex exhibit an overall enriched multi-element pattern

when normalised to primitive mantle, with more or less consistently negative Nb-Ta

anomalies, a common feature of most mafic rocks on the Kaapvaal Craton (e.g., Duncan,

1987). For the Hlagothi Complex, all samples display enrichment in some of the large ion

lithophile (LIL) elements relative to primitive mantle, and are depleted in the high field

strength (HFS) elements relative to LIL elements. For the complex, two distinct populations

can be seen for high Mg# and low Mg# samples. The low Mg# population shows more

enrichment with respect to primitive mantle and C1 chondrite compared to the high Mg#

samples, which are less enriched. Both populations show negative Nb-Ta anomalies. The low

Mg# samples also display slight negative Pb, Sr, P, Eu and Ti anomalies, and the high Mg#

samples have deep negative Ba and Sr anomalies and slightly negative Eu and Ti anomalies.

They also have a slight positive Pb anomaly. Using the prior classification and discrimination

of the Hlagothi Complex, two distinct geochemical groupings of the Hlagothi Complex can

thus be further validated.

Figure: 32 – Trace element primitive mantle normalised spider diagram for the Hlagothi Complex (McDonough

and Sun, 1995)

The rare earth element (REE) patterns reflect the trends already seen above (see Fig.

33). For all the lithologies of the Hlagothi Complex there is a very slight enrichment of light

REE relative to heavy REE. There is also a small negative Eu anomaly in both the low and


- 80 -

high Mg# samples. The REE pattern also shows a general enrichment of the REEs in the low

Mg# samples relative to the high Mg# samples, with a relatively flat trend. This is illustrated

with (La/Yb)N, with both high Mg# and low Mg# phases having values of between 2.8 and

4.3.

Figure: 33 – Rare earth element C1 chondrite normalised diagram for the Hlagothi Complex (McDonough and

Sun, 1995)

Tectonic setting discrimination diagrams using trace elements were also plotted for the

Hlagothi Complex (see Fig. 34a, b and c). Both phases of the complex display an arc, plate

margin or continental flood basalt-type signature using a variety of tectonic discrimination

diagrams, suggestive of an enriched mantle signature beneath the Kaapvaal Craton or

crustal contamination.

5.3. Dolerite Dykes

In total, 4 samples of the SE-trending dolerite dykes were taken from the north-western as

well as south-eastern portion of the White Mfolozi Archaean basement inliers. Another 3

samples of the ENE- and 4 from NE-trending dolerite dykes were also taken across the

White Mfolozi Archaean basement inliers. Data from Klausen et al. (2010) and Maré and


- 81 -

Fourie (2012) was used for comparison from various SE-, NE- and ENE-trending dolerite

dykes in the region.

Figure: 34 – Tectonic discrimination diagrams for the Hlagothi Complex. (a) Discrimination diagram of Zr/Y

versus Ti/Y after Pearce and Gale (1977). (b) Mantle source discrimination diagram (Condie, 1997); DM –

depleted mantle; PM – primitive mantle; HIMU – high U/Pb mantle source; LC – lower continental crust; UC –

upper continental crust. (c) Ternary discrimination diagram of Ti, Zr and Y after Pearce and Cann (1973). 1 –

“within-plate” basalts, 2 – low-potassium tholeiite; 3 – ocean floor basalt and 4 – calc-alkaline basalt


- 82 -

The whole rock geochemical analyses of these dolerite dykes are presented in Table 3.

Table: 3 – Whole rock geochemical analyses on various trends of dolerite dykes in northern KwaZulu-Natal

5.3.1. Rock Alteration/Classification

The alteration box plot of Large et al. (2010) enables distinction of various types of

alteration, and is useful for rhyolites, dacites, andesites and basalts, as well as their intrusive

equivalents, such as dolerite dykes. The dolerite dykes generally fall along the boundary

between least altered and altered, with minimal alteration into the epidote and chlorite end

members (see Fig. 35). Alkalis may have leached from the dolerite dykes, but the samples

have maintained their original rock-type classifications in Winchester and Floyd (1977)

AG-J AG-Bc AG-K DY-01 AG-Ba AG-Bb AG-Cb AG-A DY-02m DY-02s AG-Ca

SiO2 49.05 49.15 53.38 53.87 55.52 50.88 55.61 48.37 49.34 49.44 51.63

Al2O3 13.77 14.36 13.80 10.92 14.92 14.96 14.90 14.92 14.76 14.27 12.41

MnO 0.15 0.17 0.18 0.16 0.14 0.16 0.14 0.19 0.2 0.2 0.13

CaO 9.06 11.08 9.76 6.4 6.58 9.97 6.84 9.52 8.32 9.66 7.44

Na2O 2.12 2.00 1.96 2.66 4.41 2.10 4.24 2.29 2.45 2.38 2.64

K2O 0.82 0.60 0.86 1.02 0.96 0.78 1.45 0.65 1.08 0.96 1.41

Fe2O3T 10.43 9.39 12.03 11.73 9.11 10.73 9.17 13.11 10.89 10.56 9.33

MgO 11.27 8.98 6.25 7.63 5.83 8.16 5.48 6.95 7.68 7.06 7.96

TiO2 0.78 0.43 1.16 1.00 0.51 0.94 0.52 1.27 0.86 0.86 1.23

P2O5 0.11 0.07 0.12 0.18 0.10 0.12 0.11 0.13 0.66 0.70 1.02

Cr2O3 0.02 0.03 0.04 0.05 0.02 0.07 0.02 0.04 0.03 0.04 0.08

LOI 2.00 2.30 1.20 2.80 1.70 0.90 1.30 2.30 3.10 3.20 3.20

Total 99.58 98.56 100.74 98.42 99.80 99.77 99.78 99.74 99.37 99.33 98.48

Cs 0.40 0.30 1.30 6.20 1 1.9 1.9 1.1 1.1 1.2 11.4

Rb 32.8 19.7 39.6 72.1 42.5 26.6 91.2 87 52.3 50.3 149.7

Ba 159 133 203 379 272 212 282 119 2385 2565 1807

Th 1.4 1.6 2.9 3.8 2.2 2.6 2 0.3 6.5 6 11.2

U 0.2 0.3 0.8 0.9 0.3 0.3 0.3 0.1 1 1.1 1.8

Nb 3.7 2.6 10.8 11.9 4.2 3.9 4.4 3.5 8.3 8.4 19.5

Ta 0.3 0.2 0.7 0.8 0.2 0.2 0.2 0.2 0.4 0.4 0.9

La 9.5 8.9 18.7 29.5 15.8 10.4 14.7 5.2 92.1 97.6 133.4

Ce 19.7 18.1 39.8 63.2 30.9 23.4 29.8 13.1 185.1 191.1 264.5

Pb 1.4 1.3 1.0 3.1 12.1 11.5 10 41.3 8.2 3.9 4.7

Pr 2.52 2.20 5.02 8.04 3.62 3.15 3.62 2.02 22.13 23.11 30.81

Sr 233.4 195.0 193.0 698.7 401.3 268.7 398.4 174.2 773.6 828.5 1141

Nd 11.3 9.0 22.1 31.3 14 13.5 14.1 9.8 80.5 84.9 112.8

Zr 72.7 58.3 142.5 161.0 95.1 98.2 96.5 78.1 150 144 358.8

Hf 2.1 2.4 3.4 3.8 2.3 2.8 2.3 2.1 3.3 3.0 7.8

Sm 2.53 2.12 4.76 5.98 2.78 3.48 2.8 2.93 12.08 12.71 17.04

Eu 0.89 0.74 1.41 1.65 1.06 1.01 0.94 1.13 3.23 3.35 4.29

Gd 2.91 2.35 5.19 4.89 2.84 3.93 2.88 3.66 8.01 8.42 11.25

Tb 0.47 0.43 0.80 0.74 0.47 0.70 0.48 0.67 1.02 1.07 1.43

Dy 3.10 2.63 4.90 3.83 2.73 4.19 2.77 4.22 4.9 5.06 6.71

Ho 0.65 0.56 1.00 0.72 0.58 0.92 0.59 0.96 0.88 0.87 1.20

Er 1.90 1.63 2.95 1.81 1.57 2.62 1.62 2.74 2.25 2.19 3.21

Yb 1.72 1.44 2.67 2.01 1.44 2.36 1.47 2.51 1.9 1.92 2.59

Y 17.1 14.8 27.1 18.7 16 24 15.9 25.1 25 24.4 34.7

Lu 0.26 0.22 0.42 0.28 0.23 0.37 0.23 0.4 0.29 0.29 0.4

Tm 0.27 0.23 0.41 0.27 0.24 0.38 0.24 0.41 0.32 0.32 0.45

V 172 209 238 135 187 172 185 303 189 192 141

NE-trendENE-trend

Maj

or

ele

me

nts

(w

t. %

)Tr

ace

ele

me

nts

(p

pm

)

Trend

Sample

SE-trend


- 83 -

classification diagrams using SiO2, Nb/Y and Zr/TiO2. This suggests that the effects of

weathering and metamorphism on the geochemistry of these dykes are minimal. Large

variability in SiO2 contents is in agreement with the major element classification for SE-,

ENE- and NE-trending dolerite dykes from basalt to dacite. Nb/Y distinguishes the dykes as

alkali or sub-alkali basalts. Zr/TiO2 distinguishes the dykes as andesitic basalt to dacite, as

well as alkali, demonstrating the weakness in using Zr/TiO2 in classification of the dolerite

dykes (see Fig. 36).

Figure: 35 – Alteration box plot of Large et al. (2001) for the various dolerite dyke trends

From the IUGS recommendations (Le Bas et al., 1986), the NE-, SE- and ENE-trending

dolerite dykes show increasing levels of variability with respect to SiO2 content. NE-trending

dolerite dykes show low variability, and are generally composed of basalts with between 47-

52 wt% SiO2, and 2-4 wt% Na2O+K2O. ENE-trending dolerite dykes are highly variable, with

basalt to andesite compositions of 49-63 wt% SiO2 and 2-6 wt% Na2O+K2O. The SE-trending

dolerites are intermediate in variability between NE- and ENE-trending dolerite dykes.


- 84 -

Figure: 36 – Trace element classification diagrams of Winchester and Floyd (1977) for the dolerite dykes of

northern KwaZulu-Natal compared to published data


- 85 -

Figure: 37 - Total alkali-silica classification diagram of Le Bas et al. (1986) for the dolerite dykes (a), and

classification diagram showing the variation in K2O content with respect to SiO2 after Le Bas et al. (2000), (b)


- 86 -

However, they are generally composed of basalts to basaltic andesites, and contain

between 2-4 wt% Na2O+K2O, as well as 49-56 wt% SiO2 respectively on the TAS classification

plot (see Fig. 37a). In addition, the SE-trending dolerite dykes can be shown to have

medium-potassium content (see Fig. 37b), having between 0.5-1.5 wt% K2O. The NE- and

ENE-trending dolerite dykes also have medium-potassium content between 0.5-1.5 as well

as between 0.5-2.5 wt% K2O for the NE- and ENE-trending dolerite dykes respectively.

However, two generations can be seen in NE-trending dolerite dykes in northern KwaZulu-

Natal: one of lower and one of higher potassium content, corresponding to the NE030 and

NE050 dolerite dykes respectively. The generation of low-K conforms to the variably

feldspar-phyric dykes of Klausen et al. (2010). In addition, although most dolerite dykes are

not high-Mg (greater than 18 wt% MgO), there is a proportion of SE- and ENE-trending

dolerite dykes across the south-eastern and south-easternmost areas that are boninitic

(greater than 8 wt% MgO and less than 0.5 wt% TiO2).

Klausen et al. (2010) illustrated by way of an AFM diagram that most of the andesitic

dykes on the eastern basement of the Kaapvaal Craton were correspondingly calc-alkaline,

whereas basaltic dykes are tholeiitic. Results here are similar (see Fig. 38a). All dykes are

also sub-alkaline in northern KwaZulu-Natal. There appears to be a large grouping of SE-

trending dykes with a trend that straddles the tholeiitic to calc-alkaline borderline, although

the dykes in this study in northern KwaZulu-Natal are all tholeiitic. The ENE-trending dykes

are clearly borderline tholeiitic to calc-alkaline, and can be compared with the Rykoppies

dyke swarm and the Barberton-Badplaas dyke swarm seen further to the north, both of

which are Archaean. NE-trending dolerite dykes can further be seen in two categories, a

clear tholeiitic trend like the other NE-trending dykes on the eastern Kaapvaal Craton, and a

more tholeiitic to calc-alkaline trend only seen in northern KwaZulu-Natal, comparable to

SE- and ENE-trending dolerite dykes. The first group can be compared to the ca. 1.90 Ga NE-

trending Palaeoproterozoic dykes observed further to the north in the Black Hills dyke

swarm according to Klausen et al. (2010). Using the Jensen plot, these trends are re-

affirmed. The SE-trending dykes showing tholeiitic affinity, different from tholeiitic to calc-

alkaline dykes seen further to the north, with some komatiitic affinity too in both regions.

ENE-trending dykes as above fall between the tholeiitic and calc-alkaline field, with the NE-

trending dykes showing tholeiitic affinity only. This illustrates that there is clearly a greater


- 87 -

complexity to the mafic dyke swarms seen in the northern KwaZulu-Natal window of the

south-easternmost Kaapvaal Craton (see Fig. 38b).

Figure: 38 – AFM (a) and Jensen (b) classifications diagram of Irvine and Baragar (1971) and Jensen (1976) for

the dolerite dykes of northern KwaZulu-Natal compared to dolerite dykes on the eastern Kaapvaal Craton

Using only the major and trace element classification diagrams, the SE- and ENE-

trending dolerite dykes do not deviate from the consistent separation from the ca. 2.95 Ga

SE-trending and radiating ca. 2.65 Ga NE-, E- to SE-trending dolerite dykes of Olsson et al.

(2010) and Klausen et al. (2010) seen further to the north. This is true also within northern

KwaZulu-Natal. Herein however, two groups of the NE-trending dykes can be identified

conclusively, in agreement with the major element classification, and geology seen within

chapter 3. One of these geochemical groups is not represented further to the north in

comparison with the data set of Klausen et al. (2010) and Maré and Fourie (2012), however.

5.3.2. Magmatic Variation/Affinity


- 88 -

Major and trace variation diagrams were constructed for the different trending dolerite

dykes using Mg# and compared to data obtained by Klausen et al. (2010) and Maré and

Fourie (2012) for the same trending dolerite dyke swarms (see Fig. 39 and 40). Almost all

the dykes vary between a Mg# of 15-65. The dolerite dykes seen within northern KwaZulu-

Natal in this study and within Klausen et al. (2010) fall between 30-50, making them far less

variable than further to the north. Mg# is lowest in the NE- and ENE-trending dolerite dykes,

with values from 15-45 and 20-50 in these dykes respectively. Mg# in SE-trending dolerite

dykes varies from 25-65. Fe is more enriched in the NE-trending dolerite dykes and depleted

in the SE- and ENE-trending dolerite dykes, with 15, 12 and 9 wt% FeO+Fe2O3 respectively.

The opposite is true of SE- and ENE-trending dolerite dykes, and particulary for SE-trending

dolerite dykes compared to NE-trending dolerite dykes, with values of MgO wt% of 10, 9

and 7 respectively. SiO2 content is 48-56 wt% for all the dolerite dykes in the region, with

ENE-trending dolerite dykes having up to 56 wt%, and NE-trending dolerite dykes the

lowest, at 48-51 wt%. MnO varies between 0.12-0.20 wt%. One NE-trending dolerite dyke

generation is clearly separated from the other in MnO, however. TiO2 content is in the range

from 0.5-1.5 wt%, with NE-trends being more enriched, and ENE-trends more depleted in

TiO2. P2O5 contents of the dykes are usually <0.2 wt%, except the one generation (NE030)

with P2O5 contents up to 0.65-0.95. Al2O3 varies between 11-15 wt%. Alkali wt% varies

between 2-3 for Na2O, and 0.5-1.5 for K2O. Na2O wt% in ENE-trending dykes can be as high

as 4 wt%. CaO varies between 6-12 wt% (see Fig. 39).

Ni and Cr contents are <400 and <1000 ppm respectively for all trends within

northern KwaZulu-Natal. V varies from 15-300 ppm, Y from 1500-3500 ppm and Zr contents

vary from between 50-250 ppm. Nb, Rb and Sr vary widely, with two generations of NE-

trending dolerite dykes being evident. Nb concentrations are generally 3 ppm for all dykes,

with some NE- and SE-trends varying up to 9-20 ppm. The same geochemical grouping is

seen in Sr, with most SE-trends and NE-trending dolerite dykes at 200 ppm and ENE-trends

at 380 ppm. The one geochemical grouping of NE dolerite dykes (NE050) can be up to

between 720-840 ppm Sr. Rb varies widely from 20-100 ppm (see Fig. 40). ENE-trending

dolerite dykes can be distinguished from the other dykes by their Na2O content, and the one

generation of NE-trending dykes (NE050) by their P2O5 and Sr contents. SE-trending dolerite

dykes appear to vary the most, and NE-trending dykes the least. However most dykes in


- 89 -

northern KwaZulu-Natal do not show the same level of variability as seen in dolerite dykes

across the wider eastern basement of the Kaapvaal Craton.


- 90 -

Figure: 39 – Major element variation diagrams for the dolerite dykes


- 91 -

Figure: 40 – Trace element variation diagrams for the Hlagothi Complex


- 92 -

5.3.3. Further Characterisation/Tectonic Setting

Using the multi-element trace geochemistry for the dolerite dyke samples, variation

diagrams were constructed and normalised to either primitive mantle for the trace

elements or C1 chondrite for the rare earth elements (McDonough and Sun (1995). All

samples from the SE- and ENE-trending dolerite dykes exhibit overall enriched multi-

elemental patterns with more or less consistently negative Nb-Ta anomalies (see Fig. 41),

which as noted before is a common feature of most mafic rocks on the Kaapvaal Craton

according to Duncan (1987). Using the prior classification and discrimination of the SE- and

ENE-trending dykes, the SE- and ENE-trending dykes are more LIL element enriched to

primitive mantle, with HFS elements less so, enabling comparison to the low Mg# samples

of the Hlagothi Complex. They also have distinct negative P and Ti anomalies, as well as

having generally higher positive Pb anomalies, and some slight positive K anomalies. This

makes the ENE-dolerite dykes essentially the same geochemically as SE-trending dolerite

dykes within the south-easternmost inliers of northern KwaZulu-Natal, as well as with SE-

and E-trending dykes seen within the Barberton-Badplaas and Rykoppies areas by Klausen et

al. (2010). However, ENE-trending dolerite dykes do have a slight negative Ba anomaly and a

more pronounced positive Pb anomaly. U and Th contents are also slightly higher. Not all

NE-trending dolerite dykes have consistently negative Nb-Ta anomalies however. The two

generations of NE-trending dykes stated already in prior chapters using the classification

and affinity diagrams can be further validated. The two generations of NE-trending dykes

exhibit unusual patterns, with two very distinct chemical groupings. One grouping (NE050) is

much more LIL and HFS element enriched than the other, and its pattern is comparable with

both SE- and ENE-trending dykes. It, however, has a pronounced positive Ba anomaly and

negative Pb and Ti anomalies. The second grouping (NE030) is characterised by very low Th

and U concentrations and no apparent negative Nb-Ta anomalies. These dykes are also

marked by positive Pb and K anomalies that can correlate with their feldspar contents. Both

these trends are different from observed NE-trending dolerite dykes to the north in the

Black Hills area by Klausen et al. (2010).

The REE patterns reflect the trends already seen above (see Fig. 42). For the SE- and

ENE-trending dolerite dykes there is a slight enrichment of light REE (LREE) relative to heavy

REE (HREE).


- 93 -

Figure: 41 – Trace element primitive mantle normalised spider diagrams for the dolerite dykes of northern

KwaZulu-Natal (McDonough and Sun, 1995). Grey areas denote data from Klausen et al. (2010) for the area.

REE patterns are also relatively smooth, with no obvious anomalies. For the NE-trending

dolerite dykes, both groupings are enriched relative to C1 chondrite, however, one

population is much more enriched than the other. The one generation of NE-trending dykes

(NE030) is also distinct with a flat trend, and no enrichment of LREE compared to HREE as is


- 94 -

observed in the other trend (NE050), which is highly enriched in LREE compared to HREE.

This gives dolerite dykes with the NE030 trend a very primitive signature.

Figure: 42 – Rare earth element C1 chondrite normalised diagrams (McDonough and Sun, 1995) for the various

dolerite dykes. Grey areas denote data from Klausen et al. (2010) for northern KwaZulu-Natal

It may also be possible to relate some of this depletion to the common abundance of large

and highly segregated feldspar phenocrysts, according to Klausen et al. (2010). Thus,

feldspar-rich samples exhibit the lowest element concentrations around some of the most

positive anomalies for elements that preferentially partition into feldspars (i.e., Rb, K, Pb

and Sr).


- 95 -

Figure: 43 – Tectonic discrimination diagrams for the dolerite dyke swarms. (a) Discrimination diagram of Zr/Y

versus Ti/Y after Pearce and Gale (1977). (b) Mantle source discrimination diagram (Condie, 1997); DM –

depleted mantle; PM – primitive mantle; HIMU – high U/Pb mantle source; LC – lower continental crust; UC –

upper continental crust. (c) Ternary discrimination diagram of Ti, Zr and Y after Pearce and Cann (1973). 1 –

“within-plate” basalts, 2 – low-potassium tholeiite; 3 – ocean floor basalt and 4 – calc-alkaline basalt

However, Klausen et al. (2010) noted on the other hand, that one feldspar-barren sample

exhibited the highest element concentrations around a negative Sr anomaly. These element


- 96 -

patterns correlate roughly with moderate phenocryst contents are consistent with flow-

segregated bands of feldspars in the NE030 generation of NE-trending dykes. The (La/Yb)N

ratio shows that SE-trending dolerite dykes are extremely variable, with values of between

3.8 and 10.0, whereas ENE-trending dykes vary between 3.0 and 7.7. The NE-trending

dolerite dykes show the greatest variability, with the one generation (NE050) between

32.90 and 35.0, whereas the NE030 generation is at 1.4.

Tectonic discrimination diagrams using trace elements were plotted for the SE-, ENE-

and NE-trending dyke swarms (see Fig. 43). The SE- and ENE-trending dykes display arc-,

plate margin or continental flood basalt-type signature, suggestive of an enriched mantle

signature beneath the Kaapvaal Craton or crustal contamination. The flatter REE-patterns of

the one generation of NE-trending dykes (NE030) may have been inherited from a more

primitive mantle source and show some affinities to oceanic plateau basalt. The other

generation of NE-trending dykes (NE050), however, shows an arc signature, as was seen to

be common in the other dykes too.

Chapter: 6 – Geochronology ___________________________________________________________________________

- 97 -

Chapter: 6

Geochronology

6.1. Introduction

In this chapter, preliminary Ar-Ar amphibole ages for NE- and SE-trending dykes are

presented. In addition, precise U-Pb baddeleyite emplacement ages for the Hlagothi

Complex as well as NE- and SE-trending dolerite dykes were obtained respectively. The SE-

trending swarm is herein referred to as the ‘Hlagothi’ Dyke Swarm. The NE-trending swarm

is tentatively compared to the Neoarchaean ‘Rykoppies’ Dyke Swarm seen further to the

north by Olsson et al. (2010). These results require a re-evaluation of previously proposed

linkages to volcanic and tectonic events within the relatively well-constrained

Mesoarchaean to Palaeoproterozoic South African stratigraphy (e.g., Johnson et al., 2006).

6.2. Ar-Ar Methodology

A total of three samples were selected for 40Ar/39Ar geochronology. They were crushed and

milled to fine sand, with finer sediment remaining in solution removed through water-based

separation. Hand-picked fresh amphibole separates, usually composite grains (5–10 grains

for step-heating experiments, defined below), were carefully selected under a binocular

microscope. The samples were irradiated for 20 hours in the NTP’s Safari1 nuclear reactor at

NECSA’s Pelindaba facility in South Africa. It was run at 20 MW, and this procedure typically

produced a J-factor of 0.09. It was run in position B2W along with standards hornblende

Hb3gr as well as McCure Mountains hornblende (MMHb). The total neutron flux density

during irradiation was 9.01018 neutrons/cm2. The estimated error bar on the corresponding

40Ar/39ArK ratio is ± 0.2 % (1σ) in the volume where the samples were included. After 3

weeks of "cooling", samples were placed in pits in an aluminium disc and evacuated in an

UHV sample port with a quartz window. Small clusters of amphibole (for step-heating


- 98 -

experiments) using a defocused beam from a continuous Nd-YAG laser beam (1064 nm),

and isotopic measurements of 40Ar, 39Ar, 38Ar, 37Ar and 36Ar were measured in sequence,

usually in 7 cycles, on a single collector MAP 250-15 mass spectrometer with electron

multiplier used in analogue mode. Step-heating experiments on amphibole bulk samples

were performed with a double-vacuum high-frequency furnace. The mass spectrometer is

composed of a 120° M.A.S.S.E. tube, a Baur–Signer GS 98 source and a Balzers electron

multiplier. The three samples were heated only with a few steps, with the aim of

discriminating between Jurassic and Precambrian dykes. The usual criteria according to

Jourdan et al. (2004) used to define a plateau age are:

At least 70 % of the 39Ar released.

A minimum of three successive steps in the plateau.

The integrated age of the plateau should agree with each apparent age of the

plateau within a two sigma confidence interval (2σ).

Plateau and integrated ages are given at the 2σ level, but individual apparent ages are given

at 1σ level. The uncertainties on the 40Ar/39Ar ratios of the monitors are included in the

calculation of the integrated and plateau age uncertainties, but the error on the age of the

monitor is not included in the calculation. Data acquisition and reduction were done using

in-house software. Blank correction was routinely done and signals were extrapolated to the

time of gas admission from the mass spectrometer. Errors were propagated by a Monte

Carlo procedure.

6.3. Ar-Ar Result(s)

Analytical Ar-Ar data are shown in Table 4, and the results are presented graphically (see

Fig. 44), with ages reflecting the uncertainties in Ar decay constants shown.


- 99 -

Table: 4 – Ar-Ar amphibole data for SE- and NE-trending dolerite dykes

In the step-heating experiments, a first step was performed to degas atmospheric and

alteration phase argon. The second and following steps (including fusion) represent 80% to

90% of the radiogenic 40Ar released. The three samples yielded poorly defined second- and

third-step ages of between 335-615 Ma for DY-02 and 1757-1486 Ma for AG-A, respectively.

However, the second- and third-step ages of DY-01 produced ages of 1440 Ma and 1538 Ma,

corresponding to the rather flat Gaussian curve in an age probability density distribution

diagram with a peak age at 1623 Ma. The biggest plateau age step for DY-02 and AG-A was

907 Ma and 1319 Ma, which released 44 and 45% 39Ar gas respectively, which corresponds

to the fourth step-heating age. The 1623 Ma age plateau for sample DY-01 produced 47% of

39Ar gas released in the sample. The plateau ages were: 1596 ± 16 Ma for DY-01, 831 ± 30

for DY-02 and 1388 ± 18 Ma for AG-A.

Step Name % 39Ar Age Ma ± 2σ % err Ca/K ± 2σ % err

DY02

10.0 5.96 335.24 112.08 2.08 0.20

10.4 30.05 614.65 16.24 0.78 0.08

10.8 43.78 907.49 25.44 5.58 0.54

11.2 13.43 852.48 38.10 12.98 1.25

11.6 3.08 747.70 239.61 26.38 2.55

12.0 3.70 1002.30 77.90 16.73 1.61

Pseudo-Plateau age 57.21 830.70 30.07 7.32 0.70

DY01

10.0 5.30 1439.88 20.82 0.65 0.11

10.5 20.90 1538.28 17.78 0.93 0.11

11.0 46.81 1622.53 18.05 9.30 1.08

11.5 11.98 1581.78 15.20 8.33 0.97

12.0 8.48 2026.77 19.66 11.63 1.35

12.5 0.88 1585.48 67.87 28.41 3.31


AG-A

10.0 1.24 1757.45 86.34 3.28 0.22

10.4 20.40 1486.45 22.82 10.69 0.58

10.8 45.33 1318.94 26.72 22.84 1.24

11.0 7.33 1337.91 20.07 24.89 1.36

11.4 15.08 1476.44 18.42 28.98 1.57

12.0 10.59 907.23 21.31 16.67 0.91



- 100 -

Figure: 44 - 40

Ar/39

Ar ratio spectra from dolerite dyke amphibole separates.


- 101 -

The Ca/K ratio spectra associated with the plateau ages display relatively irregular and

disturbed patterns, with values mainly ranging from 8-12 for DY-01, 6-26 for DY-02 and 23-

29 for AG-A, highlighting the role of alteration phases in these ages. The results obtained on

the samples by step-heating experiments display Proterozoic ages. Results are difficult to

interpret because it appears that the very variable apparent ages may be the result of

alteration (clearly visible on the analysed samples that were amphibole grains probably

derived from pyroxenes during greenschist facies metamorphism, accounting for the

variable Ca/K ratios and probably some excess argon). It is therefore not possible to

attribute precise emplacement ages to these Proterozoic samples, with the exception that

the ages may reflect overprinting.

6.4. U-Pb Methodology

Baddeleyite was extracted from rock samples weighing between 1-2 kg. The extraction was

performed at Lund University in Sweden, using the ‘water-based technique’ of Söderlund

and Johansson (2002). Approximately 10-20 baddeleyite grains were extracted from each

sample and between 1-6 grains were combined in each fraction analysed. U-Pb chemistry

and mass spectrometry were performed at the Laboratory of Isotope Geology at the

Swedish Museum of Natural History in Stockholm. Upon separation, the baddeleyite grains

were transferred to Teflon© capsules and washed in 2-3 M HNO3 on a hot plate and

repeatedly rinsed with distilled and deionised H2O. A mixture of HF and HNO3 was then

added to the capsules. The baddeleyite grains were then completely dissolved after 24

hours under high pressure and temperature (~210°C). After dissolution, samples were dried.

A small amount of a 205Pb-233-236U spike was added together with a 3.1 M HCl solution and

re-dissolved before being loaded on 50 μL columns filled with a pre-washed anion exchange

resin (Bio-Rad 200-400 Mesh chloride). The Zr-Hf-REE cut was washed out with 3.1 M HCl

solution and disposed of. U and Pb were washed out with H2O and collected in the same

Teflon© capsules as used for the dissolution. A small amount of H3PO4 was added to each

capsule and the samples were put on a hot plate to evaporate overnight. U and Pb were


- 102 -

loaded on the same single Re filament together with a small amount of silica gel produced

from the recipe of Gerstenberger and Haase (1997).

Baddeleyite fractions for samples AG-A, AG-B and AG-I were analysed by mass

spectrometry on a Thermo Finnigan Triton thermal ionisation multi-collector mass

spectrometer. The intensities of 208Pb, 207Pb, 206Pb and 205Pb were measured using Faraday

collectors whereas the intensity of 204Pb was measured simultaneously in an ETP SEM

equipped with an RPQ filter. SEM to Faraday gain was controlled by measuring a ~5-10 mV

signal between runs. The Pb isotopic measurements were performed at filament

temperatures in the 1200-1320°C range. Isotopes of U were measured subsequently in

dynamic mode on the SEM at filament temperatures exceeding 1350°C. Data reduction was

performed using an in-house program written in Microsoft Excel (Per-Olof Persson, Museum

of Natural History, Stockholm) with calculations from Ludwig (1991). Analytical results and

regressions have been calculated and plotted using the Excel Macro Isoplot (Ludwig, 2003);

decay constants for 238U and 235U follow those of Jaffey et al. (1971). All errors in age and

isotopic ratios are quoted at the 95% confidence level. Procedural blanks were 2 pg Pb and

0.2 pg U, and the initial Pb isotope compositions were corrected using the model

compositions of Stacey and Kramers (1975) at 2960 Ma, 2700Ma and 2060 Ma.

6.5. U-Pb Result(s)

Analytical U-Pb data are shown in Table 5, and the results are presented graphically (see Fig.

45), with ages reflecting the uncertainties in U decay constants shown.

6.5.1. Hlagothi Complex

Sample AG-I was collected from an approximately 200m thick layered sill along the

Nsongeni River. The sill has been identified as one of the type examples of the Hlagothi

Complex, and is a massive, medium-grained, green-grey meta-peridotite (Groenewald,

2006).


- 103 -

Figure: 45 – U-Pb Concordia diagrams showing results from the peridotites of the Hlagothi Complex (AG-I) and

dolerite dykes of the Hlagothi Dyke Swarm (AG-B) and ‘Rykoppies’ Dyke Swarm (AG-A).


- 104 -

Table: 5 – U-Pb baddeleyite data for the Hlagothi Complex, as well as SE- and NE-trending dolerite dyke

Sample AG-I yielded only 15 yellow-brown and tabular baddeleyite grains with lengths of

approximately 30 to 60 μm. This population of baddeleyite had definite traces of secondary

alteration, seen as frosty surfaces. The most transparent grains were preferentially selected

for the analyses. Regression comprising 3 fractions yielded an age of 2866 ± 2 Ma (MSWD =

0.040) which is interpreted as the crystallisation age. The lower intercept is 122 ± 69 Ma.

6.5.2. Hlagothi Dyke Swarm

AG-B was sampled in a river bed of exposed Mvunyana granodiorite 100m to the west of the

R34 road between Melmoth and Vryheid. It was collected from a massive, dark grey, coarse-

grained and sub-vertical 20m wide SE-trending dolerite dyke This SE-trending dyke has a

well-exposed strike length of ~1 km along the river, and is up to 20m wide. The country rocks

are the basement granitoids of the White Mfolozi Inlier. Mineral separation yielded 50

brown and tabular-shaped baddeleyite crystals, some with frosty surfaces (presumably due

to partial replacement of baddeleyite with zircon). Typical lengths of the crystals varied from

50-100 μm. Regression through seven analyses of AG-B yields an upper intercept age of

2874 ± 2 Ma (MSWD = 3.2), which is interpreted as the crystallization age of this sample,

and a lower intercept of 0 ± 100 Ma, which was a forced regression. Two single-grain

Analysis no. U/Th Pbc/206

Pb/207

Pb/ ± 2σ206

Pb/ ± 2σ207

Pb/206

Pb/207

Pb/ ± 2σ Concordance

(number of grains) Pbtot1 204Pb 235U % err 238U % err 235U 238U 206Pb % err

raw2 [corrected]3 [age, Ma]

Sample AG-A, 28.21037° S, 30.98243° E

Bd-a (3 grains) -3.4 0.308 172.2 10.3700 1.49 0.42026 1.49 2468.4 2261.7 2643.3 7.5 0.856

Bd-b (3 grains) 2.2 0.496 91.3 6.2388 4.22 0.25761 4.25 2009.9 1477.6 2612.2 21.2 0.566

Sample AG-B, 31.26239° E, 28.34448° S

Bd-a (3 grains) 17.7 0.007 7747.9 15.7843 0.23 0.55618 0.21 2863.8 2850.8 2873.0 1.7 0.992

Bd-b (1 grain) 4.8 0.042 1175.2 15.6440 0.28 0.55052 0.24 2855.3 2827.4 2875.1 2.0 0.983

Bd-c (2 grains) 7.5 0.026 2017.6 15.8214 0.31 0.55783 0.29 2866.1 2857.7 2872.0 1.8 0.995

Bd-d (5 grains) 23.4 0.009 5604.7 15.6433 0.28 0.55083 0.27 2855.3 2828.6 2874.1 1.4 0.984

Bd-e (6 grains) 16.9 0.011 4457.5 15.6318 0.28 0.55023 0.27 2854.6 2826.1 2874.7 1.3 0.983

Bd-f (2 grains) 23.5 0.032 1779.1 15.6616 0.43 0.55256 0.43 2856.4 2835.8 2870.9 2.0 0.988

Bd-g (1 grain) 31.5 0.018 3056.0 15.7423 0.29 0.55371 0.25 2861.3 2840.6 2875.9 2.5 0.988

Sample AG-I, 30.93429° E, 28.45418° S

Bd-a (3 grains) 8.5 0.095 505.9 13.7470 0.30 0.48854 0.26 2732.4 2564.4 2859.1 2.3 0.897

Bd-b (5 grains) 3.8 0.032 1527.2 15.0517 0.26 0.53347 0.24 2818.5 2756.1 2863.5 1.5 0.962

Bd-c (3 grains) 11.0 0.085 602.4 15.2452 0.50 0.54024 0.48 2830.7 2784.5 2863.8 2.6 0.972

3) isotopic ratios corrected for fractionation (0.1% per amu for Pb), spike contribution, blank (0.5 pg Pb and 0.05 pg U), and initial common

Pb. Initial common Pb corrected with isotopic compositions from the model of Stacey and Kramers (1975) at 2875 Ma

1) Pbc = common Pb; Pbtot = total Pb (radiogenic + blank + initial).2)

measured ratio, corrected for fractionation and spike.


- 105 -

fractions were analysed without U-Pb column chemistry, in order to obtain more

concordant data and to minimise the blank contamination.

6.5.3. ‘Rykoppies’ Dyke Swarm

Sample AG-A was collected from the NE-trending dyke, which has a well-exposed strike

length of approximately 11 km and a width of 50-100m. Baddeleyite extraction resulted in

approximately 100 dark-brown and wafer-shaped grains with an overall length of 50-100

μm. Regression yields an upper intercept age of 2652 ± 11 Ma and a lower intercept of 104

± 43 Ma. The sample had very low counts of U and Pb, and therefore only two fractions

were analysed. It was decided at the time not to continue with further analysis, and the

relatively high analytical uncertainty. This cannot be regarded at this stage as a rigorous

emplacement age, and is only used here as a preliminary result.


- 106 -

Chapter: 7 – Palaeomagnetism ___________________________________________________________________________

- 107 -

Chapter: 7

Palaeomagnetism

7.1. Introduction

A palaeomagnetic study has been conducted on the various inliers of the Hlagothi Complex,

as well as on a variety of SE-, ENE- and NE-trending dolerite dyke swarms across the

southeastern Kaapvaal Craton, thought to represent swarms with ages of ca. 2.95, 2.65 and

1.90 Ga (Lubnina et al., 2010). However, from the previous chapter, it is seen that there are ca.

2.87 Ga SE-trending dykes in the area, and possibly ca. 2.65 NE-trending dolerite dykes as well.

In the Nkandla inliers, between six and eight orientated samples were collected from

the Hlagothi Complex, in addition to one to three orientated samples from the basal ca. 2.95

Ga Nsuze Group quartzite host rock (Pongola Supergroup). Intrusive contacts were sampled

for a baked contact test. Between six and fourteen oriented samples were collected from

each dolerite sill or dyke site. In addition to this, four to eight samples were collected from a

host rock intruded by these dykes or sills, as well as the contact region. Most often the host

rock was an undifferentiated granitoid of either Palaeoarchaean or Mesoarchaean age.

Tectonic disruptions during Gondwana break-up are not thought to have significantly

rotated any of the sample sites because rift-related deformation was concentrated along

the African continent’s passive margins according to Klausen (2009). Also, Karoo Supergroup

sedimentary rocks display minimal dips, and dip very gently away from these margins and

across most of the southern part of the eastern Kaapvaal Craton.


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7.2. Methodology

The samples were collected using a water-cooled, portable, petrol drill. Most samples were

oriented with both a magnetic and sun compass. There were no significant differences

between sun and magnetic compass orientation measurements. The oriented drill core

samples were cut into standard cylinders of ~2 cm3 in size. Measurements were completed

at the palaeomagnetic laboratory of the University of Johannesburg, South Africa.

Remanence measurements were performed using a 2G Enterprises 755-4K Superconducting

Cryogenic Rock Magnetometer. Samples were progressively demagnetised using alternating

field (AF) and thermal demagnetisation methods. Demagnetisation generally consisted of

between 15 to 30 steps. Four AF steps were done in 2.5 mT steps from 2.5-10 mT, followed

by 22 systematically decreasing incremental temperature steps from 100-580°C. For

selected dolerite dyke specimens, 20 AF steps were done between 2-85 mT. AF

demagnetisation was done using a tumbling demagnetiser, while thermal demagnetisation

was achieved with stepwise heating within an ASC model TD48 shielded oven. NRM

components were visually identified by using stereographic and orthogonal projections

(Zijderveld, 1967). The directions of components were quantified via principal components

through least-squares analyses (Kirschvink, 1980), which is based in all cases on at least

three or more vector endpoints. Only components with a maximum angular deviation

(MAD) less than 10° were accepted for further interpretation. Mean directions were

calculated according to Fisher (1953). These calculations and a graphic representation of the

results were carried out using software by Jones (2002) and Cogné (2003).


A palaeomagnetic study has been applied to various inliers and lithologies of the Hlagothi

Complex. A total of 90 oriented samples were collected from 12 sampling sites from the

Hlagothi Complex as well as from possible eastward extensions of it along the Mhlatuze


- 109 -

River. Samples AG-I, HC-01, HC-03, HC-05 and HC-07 were gathered from the Nsongeni

sheets, and sample HC-08 from the Hlagothi sheets. A sample from the Wonderdraai sheets

was also collected, HC-10. In addition, samples (AG-D, AG-E, AG-F, AG-G and AG-H) were

taken from the possible eastward extension of the complex along the Mhlatuze River

(Mhlatuze sheets). Samples from the Mhlatuze sheets of the complex were omitted due to

erratic behaviour during demagnetisation, possibly through alteration or metamorphism

present in the inlier, as well uncertain structural relationships. A regional tilt correction was

applied to samples from the Nsongeni sheets, where a strike and dip was apparent,

particulary at the contact with the Nsuze Group quartzites. This was measured from outcrop

as 256°/13°. No tilt correction was applied to the Hlagothi or Wonderdraai sheets.

Table: 6 – Palaeomagnetic data for the Hlagothi Complex

Component Site LithologyLat

(in °N)

Long

(in °E)n/N L/P

Decl.

(in °)

Incl.

(in °)k

α95

(in °)

Decl.

(in °)

Incl

(in °)k

α95

(in °)

PLF AG-I

peridotite,

sheared

peridotite,

baked

quartzite

-28.5 30.9 14/14 14/0 330.9 -66.9 12.35 12.21 339.8 -37.5 12.35 12.21

HC-01 peridotite -28.5 30.9 8/8 8/0 76.8 -54.9 79.29 5.85 70.2 -68.4 78.74 5.87

HC-03 pyroxenite -28.5 30.9 6/6 6/0 83.7 -70.6 49.84 8.81 58.3 -83.8 49.14 8.80

HC-05 peridotite -28.5 30.9 5/5 5/0 86.3 -57.7 33.18 10.75 84.7 -71.6 33.21 10.74

HC-07 peridotite -28.5 30.9 3/11 3/0 76.0 -56.2 11.17 31.11 76.0 -56.2 11.17 31.11

HC-08 gabbro -28.5 30.9 3/6 3/0 50.2 -48.5 31.12 18.25 50.2 -48.5 31.12 18.25

HC-10 diorite -28.5 31.0 5/6 5/0 100.3 -70.4 16.36 17.33 100.3 -70.4 16.36 17.33

81.9 -61.1 86.0 13.4

B HC-03 pyroxenite -28.5 30.9 6/6 1/5 252.0 59.2 126.86 9.29 236.3 71.4 125.74 9.33

HC-01 peridotite -28.5 30.9 8/8 8/0 97.6 18.0 48.48 7.48 97.2 4.2 48.72 7.49

HC-05 peridotite -28.5 30.9 6/6 6/0 109.7 43.1 221.07 4.12 106.2 29.8 218.32 4.15

HC-07

sheared

peridotite,

baked

quartzite

-28.5 30.9 6/11 6/0 107.2 53.1 8.22 22.44 136.6 44.8 4.55 30.52

HC-08 gabbro -28.5 30.9 2/6 0/2 - - - - - - - -

AG-I

sheared

peridotite,

baked

quartzite

-28.5 30.9 5/14 5/0 83.1 59.9 25.90 16.65 121.5 46.8 25.98 16.62

100.3 44.1 16.7 23.1

AG-I peridotite -28.5 30.9 8/11 8/0 12.0 27.0 38.13 8.42 25.0 53.3 38.15 8.48

HC-07 peridotite -28.5 30.9 2/11 2/0 12.8 55.1 - - 12.8 55.1 - -n/N = number of samples included/number analysed, L/P = line verses plane least squares analyses, Decl. = mean declination in degrees, Incl. = mean inclination in

in degrees, k = Fisher’s precision parameter, except when in modified form for where both line and plane data were combined and α95 = radius of 95% confidence cone

about the mean

PLF is a present local field component. Components 'A' and 'B' are overprint directions associated with post-Pongola granitic intrusions between 2850 and 2650 Ma.

Component 'C' is related to an overprint from the Meso- to Neoproterozoic Namaqua-Natal orogen, whereas component 'D' is taken as the primary 2866 Ma direction.

Underlined values represent data used to calculate component means, which are given in bold letters.

D

Site location Present coordinates Tilt-corrected coordinates

Component A mean =

Component C mean =

A

C


- 110 -

At each site, between six and eight oriented samples were collected from the Hlagothi

Complex. In addition, one to three orientated samples from the basal ca. 2.95 Ga Nsuze

Group quartzite host rock (Pongola Supergroup) into which the Hlagothi Complex intrudes

was sampled for a baked contact test. The schistose contact, where present, was also

sampled. The palaeomagnetic results from the Hlagothi Complex are summarised in Table 6.

Examples of the demagnetisation behaviour and summaries of the remanent magnetisation

are also shown (Fig. 46, 47 and 48).

The meta-peridotites returned good intra-site reproducibility, whereas samples from

the meta-gabbro sites behaved more erratically during demagnetisation. Apart from

spurious low coercivity magnetic components, five stable and geologically significant

magnetic components (i.e., PLF, A, B, C and D) were identified. Each sample recorded a

maximum of two of these components. During AF demagnetisation and thermal

demagnetisation steps up to 530°C, peridotite samples from site AG-I revealed steep

northerly magnetic components with negative inclination, which is parallel to the present

local geomagnetic field at the sampling site. This component was labelled PLF (i.e., present

local field). The most widespread component (i.e., identified within peridotite, gabbro and

diorite samples) is a steep easterly magnetic component with negative inclination. This

component, labelled A, was variably demagnetised during AF and thermal demagnetisation

steps between 200°C and 515°C at six sampling sites. Following the removal of either PLF or

A components, three characteristic remanent magnetisations (i.e., B, C or D) were identified

in the various lithologies via thermal demagnetisation up to 580°C, but the signal of most

samples dropped below the noise level of the sample holder (1 x 10-9 A.m2) before this

temperature step was reached, and the samples started to show erratic behaviour.

Component B was only identified in meta-peridotite samples from site HC-03. It generally

demagnetised as great-circle arc trajectories away from component A towards an antipodal,

westerly and downward direction. One sample reached a stable end-point of

demagnetisation at 515°C, which allowed for definition of component B when combined

with great-circle arcs of the five other samples (McFadden and McElhinny, 1988). By far the

most widespread of the characteristic remanent magnetisations is component C. It was

identified within meta-peridotite, the sheared contacts, baked host quartzite of the Nsuze

Group, and meta-gabbro.


- 111 -

Figure: 46 – Representative sample demagnetisation behaviour using Zjiderveld projections for samples taken

from the Hlagothi Complex. Various components shown with coloured straight lines


- 112 -

Figure: 47 – Representative sample demagnetisation behaviour using equal area projections for samples taken

from the Hlagothi Complex


- 113 -

Figure: 48 – Equal-area plots of component means for components ‘sft’, ‘PLF’, ’C’, ‘B’, ’A’ and ’D’ fitted by the

least-squares line analysis method of Kirschvink (1980)


- 114 -

These units represent some of the most extensively altered lithologies sampled during this

study. After removal of A and PLF components, demagnetisation trajectories generally

followed linear tracts towards the origin. Least-squares analyses reveal moderately

negatively inclined eastward directed characteristic components. In contrast to this, the

demagnetisation of samples from the least altered lithologies (i.e., meta-peridotite), which

were generally collected far away from intrusive contacts, revealed a north-north-westerly

characteristic remanent component with moderate positive inclination.

Figure: 49 – Palaeomagnetic sample sites AG-I and HC-07 showing sample positions at the sites in order to produce a baked contact test. Also shown is the baked contact test using for site AG-I


- 115 -

We named this component D. This component was only identified from two sampling sites

(AG-I and HC-07), from eight and two samples, respectively (see Fig. 46, 47 and 48).

The baked contact test shows the existence of the same component C that was seen

in the Hlagothi Complex, as well as in the sheared contact between the complex and the

quartzites. The same component was seen in the quartzites of the Nsuze Group up to 30m

away from the contact, providing evidence toward a failed contact test. However, the

component D in two sites from the least altered sills of the Hlagothi Complex was revealed

only after the removal of component C and A within these sills.

Figure: 50 – VGPs plotted for each component observed within each site of the Hlagothi Complex (components

A, B, C and D), with the Kaapvaal Craton in its present orientation and position shown in orange


- 116 -

This indicates an origin before the overprinting by component C in all lithologies. Although

this is indicative of an inconclusive baked contact test, the complex, sheared contact and

host rock has been overprinted by component C, some of the least altered lithologies in the

Hlagothi Complex have retained an earlier high magnetic component, D (see Fig. 49).

The virtual geomagnetic poles (VGPs) are plotted below for each site for each

component, labelled poles A, B, C and D in relationship to the Kaapvaal Craton in its present

position (see Fig. 50), with A and B plotting at similar low latitudes, and C and D in

intermediate latitudes offset by up to 50°.

7.4. SE-trending dolerite dykes

Palaeomagnetic studies were conducted on two of the SE-trending dolerite dykes, from

which samples AG-K and AG-J were collected. In total, 31 oriented samples were collected

from the two sampling sites. Sample behaviour from the 10m wide SE-trending dolerite

dyke AG-J was erratic and chaotic, possibly due to alteration or metamorphism, or due to

the dyke’s anastomosing to braided nature, as well as the influence posed by the ENE- and

SSE-trending dolerite dykes 10m and 50m away respectively. For site AG-J, 10 samples were

taken of both the coarse-grained and fine-grained phases of the dyke, in addition to one

sample from a small glassy-textured dyke splaying off from the main dyke. Further, three

samples were obtained from within 0.3m of the contact with the surrounding Mvunyana

granodiorite. The granodiorite itself was sampled 10m away from the dyke in order to

produce a baked contact test. At the site AG-K, six oriented samples were collected from the

second 10m wide SE-trending dolerite dyke. In addition, five orientated samples were taken

from the contact zone and a xenolith of the country rock within the margin of the dyke. This

sampling zone was less than 1m on either side of the contact, with the Mvunyana

granodiorite. Three orientated samples 20m away in the unbaked granodiorite were also

collected. This was done in order to conduct a baked contact test. The palaeomagnetic

results from the SE-trending dolerite dykes are summarised in Table 7. Examples of the


- 117 -

demagnetisation behaviour and summaries of the remanent magnetisation are shown (see

Fig. 51).

Table: 7 – Palaeomagnetic data for SE-trending dolerite dykes

The SE-trending dolerite dykes returned good site reproducibility at AG-K, whereas

samples from the site AG-J were more erratic during demagnetisation, probably due to

alteration and/or metamorphism. Also, several phases are present. The effect of the two

later dyke events in the vicinity may also have affected the result. Two spurious low

coercivity magnetic components were noted in AG-K, whereas no low coercivity magnetic

components were identified in AG-J. One stable and geologically significant magnetic

component, C was identified from both sites however. During AF demagnetisation and

thermal demagnetisation steps up to 500°C, followed by alternating field demagnetisation

up to 8.5 mT, the dolerite dyke samples from site AG-K revealed a reproducible and

characteristic magnetic component. This component is in a moderate easterly direction,

with a positive inclination, which is similar to the Meso- to Neoproterozoic overprint

component, C, already identified for the Hlagothi Complex (see Fig. 46, 47 and 48). This

component was labelled C also, as both dolerite dykes are in the same geologic vicinity as

the Hlagothi Complex. This component is observed within the dolerite dykes, as well as the

contact zones and unbaked host rock of the Mvunyana granodiorite, suggesting this

component may represent a younger overprint. The means for these components are

plotted on stereographic projections (see Fig. 52), and are also plotted geographically as

VGPs (see Fig. 53), in a moderate northern latitude already seen in the component C from

the Hlagothi Complex (see Fig. 50).


(in °N)

Long

(in °E)n/N L/P

Decl.

(in °)

Incl.

(in °)k

α95

(in °)

Decl.

(in °)

Incl

(in °)k

α95

(in °)

AG-J dolerite -28.3 31.3 8/17 17/0 65.1 73.7 7.46 18.88 65.1 73.7 7.46 18.88

AG-K dolerite -28.4 31.2 13/14 13/0 92.1 46 37.37 7.21 92.1 46 37.37 7.21

n/N = number of samples included/number analysed, L/P = line verses plane least squares analyses, Decl. = mean declination in degrees, Incl. = mean inclination in


about the mean

Component 'C' is related to an overprint thought to be associated with the Meso- to Neoproterozoic Namaqua-Natal orogen


C


- 118 -

Figure: 51 – Representative sample demagnetisation behaviour using Zjiderveld and equal area projections for

all samples taken from the SE-trending dolerite dykes. Various components are highlighted with coloured

straight lines


- 119 -

Figure: 52 – Equal-area plots of component means for components ‘sft’ and ’C’, fitted by the least-squares line

analysis method of Kirschvink (1980)

Figure: 53 – VGPs plotted for each component observed within each site of the SE-trending dolerite dykes

(component C) with the Kaapvaal Craton in its present orientation and position

7.5. ENE-trending dolerite dykes

Palaeomagnetic studies were done on two of the ENE-trending dolerite dykes. Sample sites

included AG-B and AG-C. In total, 23 oriented samples were collected from these sites.

Samples from dolerite dyke AG-B showed chaotic behaviour, whereas sample behaviour

from the ENE-trending dolerite dyke AG-C was reproducible and consistent. For AG-C, seven


- 120 -

orientated samples were drilled into the dyke, and five samples were also drilled into the

contact zone with the surround Mvunyana granodiorite and the granodiorite itself in order

to produce a baked contact test. The palaeomagnetic results from this ENE-trending dolerite

dyke are summarised in Table 8. Examples of the demagnetisation behaviour and

summaries of the remnant magnetisation are also shown (Fig. 54 and 55).

Table: 8 – Palaeomagnetic data for the ENE-trending dolerite dyke

The ENE-trending dolerite dyke at site AG-C returned very good site reproducibility,

probably due to its homogeneity and fine grain size.


all samples taken from the ENE-trending dolerite dyke AG-C. The various components are illustrated with

straight coloured lines


(in °N)

Long

(in °E)n/N L/P

Decl.

(in °)

Incl.

(in °)k

α95

(in °)

Decl.

(in °)

Incl

(in °)k

α95

(in °)

E AG-C dolerite -28.3 31.3 9/13 9/0 312.2 47.4 33.65 8.48 312.2 -47.4 33.65 8.48

F AG-C dolerite -28.3 31.3 7/13 7/0 253.2 82 195.47 3.71 253.2 82 195.47 3.71

n/N = number of samples included/number analysed, L/P = line verses plane least squares analyses, Decl. = mean declination in degrees, Incl. = mean inclination in


about the mean

Components 'E' is an overprint direction associated with either the 180 Ma Karoo LIP or present local field

Component 'F' is taken as a primary 2650 Ma direction.



- 121 -

Figure: 55 – Equal area plots of component means for components ‘sft’, ’E’ and ‘F’ fitted by least squares line

analysis method of Kirschvink (1980)

The dyke produced two stable components at intermediate and high temperature steps

during demagnetisation. After a spurious low coercivity magnetic component was removed

by initial AF demagnetisation steps up to 0.5 mT within the dolerite dyke, component E was

then demagnetised from 0.5 mT up to temperature steps of 350°C. This component has an

intermediate negative north-west component, and is unique to the intrusion. The high

temperature component was isolated during thermal demagnetisation above 350°C up to

approximately 530°C, after which sample behaviour became erratic. This component has a

west-south-west orientation with a steep positive inclination, which due to a positive baked

contact test described below, is taken to be the primary direction for this ENE-trending

dolerite dyke.

Figure: 56 – VGPs plotted for each component observed within the ENE-trending dolerite dykes (component E

and F) compared with the Kaapvaal Craton in its present orientation and position


- 122 -

The means for these components are plotted on equal area projections (see Fig. 55), and

plotted geographically as VGPs (see Fig. 56), with component E in moderate northern

latitudes and F in moderate southern latitudes.

A baked contact test was attempted at locality AG-C. Samples were collected within

the baked contact zone up to 1m either side of the contact, as well as within the host rock

granitoids, which were collected up to about 20 m away from the 5m wide ENE-trending

dolerite dyke. The samples of baked granitoids in the contact zone exhibit two components

during demagnetisation.

Figure: 57 – Palaeomagnetic site AG-C showing sample positions at the site in order to produce a baked contact test. Also shown is the baked contact test using a equal-area projection


- 123 -

The low-coercivity or ‘soft’ component demonstrates a present-day geomagnetic field

component. An intermediate positive south-easterly component then unblocks up to

temperatures of about 500°C. A different component, however, was seen in the host rock

granodiorite collected away from the chilled margin and dyke, with two components seen

unblocking up to a 500°C. A component was seen toward the south-south-east with a

moderate negative inclination up to 300°C, as well as a component toward the north-west

with a moderate negative inclination from 300-500°C, showing a positive baked contact test.

(see Fig. 57).

7.6. NE-trending dolerite dykes

Palaeomagnetic studies were applied on both generations of the NE-trending dolerite dykes,

from which samples AG-A (NE030) and DY-02 (NE050) were collected. In total, 29 oriented

samples were collected from 2 sampling sites. Sample behaviour from AG-A was chaotic,

with only samples from DY-02 returning reproducible results. From this 3m wide dyke, six

samples were drilled into the dolerite dyke, and three into a small dolerite dyke splaying off

from the main dyke. In addition, three samples were drilled into the surrounding Mvunyana

granodiorite in order to produce a baked contact test. From the baked contact, four were

taken at the contact between the dyke and the host rock. The palaeomagnetic results from

the NE-trending dolerite dyke DY-02 are summarised in Table 9. Examples of the

demagnetisation behaviour and summaries of the remanent magnetisation are shown (Fig.

58 and 59).

Table: 9 – Palaeomagnetic data from a NE-trending dolerite dyke


(in °N)

Long

(in °E)n/N L/P

Decl.

(in °)

Incl.

(in °)k

α95

(in °)

Decl.

(in °)

Incl

(in °)k

α95

(in °)

D DY-02 dolerite -28.2 31.0 7/9 7/0 17.8 27.5 64.98 7.03 17.8 27.5 64.98 7.03n/N = number of samples included/number analysed, L/P = line verses plane least squares analyses, Decl. = mean declination in degrees, Incl. = mean inclination in


about the mean

Component 'D' is taken as a 2866 Ma direction associated with the Hlagothi Complex.



- 124 -


all samples taken from the NE-trending dolerite dyke DY-02. The one component is shown with a straight

coloured line

Figure: 59 – Equal area plots of component means for component ‘D’ fitted by the least-squares line analysis

method of Kirschvink (1980)

The NE-trending dolerite dyke returned very good site reproducibility at DY-02, probably

due to its homogeneity and fine grain size. The dyke produced one stable component at

intermediate and high temperature steps during demagnetisation. After a variety of

spurious low coercivity magnetic components were removed in initial AF demagnetisation

steps up to 10 mT within the dolerite dyke, component D was then demagnetised from 10


- 125 -

mT up to temperature steps of 540°C, and alternating field steps up to 85 mT, upon which

sample behaviour became erratic. This component has a positive and shallow north-north-

east direction, and is unique to the intrusion. The mean for this component is illustrated

below (see Fig. 59), with the VGP plotted in intermediate to high northern latitudes (see Fig.

60).

Figure: 60 – VGP plotted for the component observed within the NE-trending dolerite dyke DY-02 (component

D) with the Kaapvaal Craton in its present orientation and position

A baked-contact test was attempted at locality at DY-02. Sampling was carried out within

the baked-contact zone, as well as the host rock granitoids up to about 20 m away from the

NE-trending (NE050) dolerite dyke contact. The samples of baked granitoids produced

erratic and chaotic directions during demagnetisation. The baked-contact test is therefore

inconclusive, but suggesting that the dolerite dyke is different from the host rock in

magnetisation.


- 126 -

Chapter: 8 – Discussion ___________________________________________________________________________

- 127 -

Chapter: 8

Discussion

8.1. Intrusion and metamorphism

The inliers of the Palaeo- to Neoarchaean granitoid-greenstone terrane of the south-

easternmost Kaapvaal Craton are intruded by numerous dolerite dykes of SE-, ENE- and NE-

trends. These dykes are absent from the overlying Phanerozoic cover of the Karoo

Supergroup, indicating that they are Precambrian in age. The Mesoarchaean Pongola

Supergroup overlies this Archaean basement terrane. The Hlagothi Complex is intrusive into

the basal quartzite strata of the Mantonga Formation of the Nsuze Group, the Nkandla sub-

basin. It was once understood to be coeval with the ca. 2.95 Ga Nsuze Group (Hegner et al.,

1981). The Hlagothi Complex now has a newly determined U-Pb baddeleyite emplacement

age of 2866 ± 2 Ma, dated herein. It consists of a layered series of sills of meta-peridotite,

pyroxenite and gabbro. Several other intrusives into the Pongola strata have been noted

and mapped in the region by Groenewald (1984) and Gold (1993), with very little further

study. These deformed and metamorphosed inliers of Archaean crust are in turn overlain by

relatively fresh strata of the Karoo Supergroup diamictites and shales of the Dwyka and Ecca

groups. This strata is also intruded across the region by Jurassic SSE-trending dolerite dykes

and sills.

In addition, the Kaapvaal Craton on this south-easternmost side is in close proximity

to the craton margin, and the Namaqua-Natal Mobile Belt to the south and the

Mozambique Mobile Belt east of it. Units of the Namaqua-Natal orogen in places have been

thrust up and on to the craton. Elworthy et al. (2000), and Jacobs and Thomas (2001) noted

that the Kaapvaal Craton more than 50 km away from the Natal Thrust Front was

experienced lesser burial and exhumation from the orogeny, whereas localities further to

the south experienced greater metamorphism and deformation. Rb-Sr whole rock ages on

the granitoids to the south of this 50 km line returned a mean age of 967 ± 24 Ma, whereas


- 128 -

granitoids to the north of it produced ages of 2614 ± 74 Ma (Elworthy et al. 2000). The

Hlagothi Complex straddles this zone, being emplaced approximately 50 km from the

cratonic margin. The various Precambrian dolerite dyke generations of the White Mfolozi

inlier were emplaced between 50 and 80 km from the cratonic margin. Other dolerite dykes

are indeed present to the south of this 50 km line, but have not been mapped or described

in any detail. One generation each of SE- and NE-trending dykes have been dated in this

study by the U-Pb baddeleyite method at 2874 ± 2 Ma and 2652 ± 11 Ma. The first result is

thought to represent an emplacement age. The age of ca. 2652 Ma on the NE-trending

dolerite must, however, be regarded as preliminary. Ar-Ar amphibole dates were obtained

from SE- and NE-trending dolerite dykes in the White Mfolozi inlier as well. These dates are

1596 ± 16 for a SE-trending dolerite dyke and 831 ± 30 Ma and 1388 ± 18 Ma from NE-

trending dykes, respectively. These pseudo plateau ages, as well as their Ca/K ratios

however, were disturbed, and may reflect alteration and metamorphism related to the

cooling and exhumation of the Namaqua-Natal orogeny to the south, as no known plutonic

or volcanic event can be associated with these ages.

The rock types in this south-easternmost terrane of the Kaapvaal Craton have seen

strong alteration of the original lithologies at greenschist facies. The presence of uralitic

amphiboles, chlorite, talc, serpentine, epidote, and sericite attest to this in both the

Hlagothi Complex and variously trending dolerite dykes, as these secondary minerals are the

products of greenschist facies metamorphism and alteration of the original ortho- and

clinopyroxenes, olivine and plagioclase feldspar. Relict textures and relict primary pyroxene

in the dolerite dykes attest to the heterogeneous nature of the metamorphic alteration, as

do the 40Ar/39Ar ages obtained. This means that although the 50 km metamorphic isograd of

Elworthy et al. (2000) is correct, the area has sustained a much more protracted and

heterogeneous set of metamorphic conditions. This is also reflected in the palaeomagnetic

data (discussion to follow). From the data presented, and from the literature, it would

appear that there were metamorphic episodes or overprinting at ca. 2870, 2800 to 2600

Ma, 1600 to 800 Ma and 180 Ma. These events can be linked with the Hlagothi Complex LIP

event in the Mesoarchaean, post-Pongola intrusions and Ventersdorp volcanism in the

Neoarchaean, the protracted Namaqua-Natal orogeny in the Meso- to Neoproterozoic for

the region, and lastly the Jurassic Karoo LIP.


- 129 -

For the Hlagothi Complex, Groenewald (1984) argued that shallow intrusion of the

complex into water-rich rocks and subsequent syn-metasomatism placed the timing of

intrusion shortly after the deposition of the Nsuze Group, thus accounting for their

alteration. He did note post-intrusion metamorphic and deformation events, which could in

part be associated with the Meso- to Neoproterozoic Namaqua-Natal orogeny. The

palaeomagnetic and 40Ar/39Ar geochronological data provide additional information about

the possible timing of this post-intrusive alteration. A remanence component A identified in

this study was seen across all lithologies in the Hlagothi Complex, while component B was

revealed only within gabbroic samples from one site (see Fig. 61). These two magnetic

directions were isolated in the more altered lithologies, i.e. the peridotites, the sheared

contacts with the quartzites of the Nsuze Group, and the gabbros. Relict textures are

common, and in some cases there is unaltered minerals too, particulary pyroxenes.

Component B is antipodal to component A, and they are interpreted to represent a dual

polarity overprint. The VGPs of these two components are in a similar position as poles

obtained for the Agatha basalts of the Nsuze Group by Strik et al. (2007), and poles for the

2782 ± 5 Ma Derdepoort basalts by Wingate (1998) for example. It is interesting to note that

between 2850 and 2650 Ma, the whole southeastern Kaapvaal Craton experienced the

intrusion of voluminous potassic granitic plutons, collectively known as the post-Pongola

granites. The ages of these granites appear to be coeval with the Rb-Sr mean ages of 2614 ±

74 Ma obtained by Elworthy et al. (2000). We suggest that the Agatha basalt pole of Strik et

al. (2007) may very well also represent a magnetic overprint caused by the intrusion of

these post-Pongola granites into the rocks of the Pongola Supergroup in the area studied. In

addition, this pole shows a significant discrepancy to the similar age Nsuze dolerite dykes

and basalt pole of Lubnina et al. (2010), which is further enhanced by baked contact tests. A

primary component, F in the ENE-trending dolerite dykes across the region is also seen in

this study, and by Lubnina et al. (2010), which are in a similar region to ca. 2.65 Ga poles. It

is also similar to the Agatha basalt pole of Strik et al. (2007), further arguing that the Agatha

basalt pole is not primary.

Another component seen in most lithologies and interpreted as a younger magnetic

overprint is component C. It was seen in the peridotites and gabbros, as well as the

quartzites into which the Hlagothi Complex was intruded. In addition, some SE-trending


- 130 -

dykes appear to show the same component. Baked contact tests on the quartzites and

granitoids into which the Hlagothi Complex and SE-trending dolerite dykes intrude also

display the same component, and further attest to both the dykes and granitoid host rocks

being affected by the same regional metamorphic event. It cannot reflect a primary

direction. When compared to other known palaeopoles from the Kaapvaal Craton, the VGPs

for this component resembles the pole for the ca. 1050 Ma Ntimbankhulu pluton in the

Margate terrane of the Natal sector within the Namaqua-Natal Mobile Belt obtained by

Maré and Thomas (1998). The age of the Ntimbankhulu granite-charnockite pluton carries

significant uncertainty, but it is known to have been produced during the Meso- to

Neoproterozoic. It is believed component C represents an overprint related to this event,

given the SE-trending dolerite dykes’ and the Hlagothi Complex's proximity to the Natal

Thrust Front (see Fig. 9).

Other components include the possibility of a magnetic component D for the

Hlagothi Complex and component F seen within NE-trending dolerite dykes respectively.

Components D and F will be presented further in this chapter, as they may be primary.

Figure: 61 – Palaeopoles and VGPs for various units within the Mesoarchaean to Mesoproterozoic stratigraphy

of the Kaapvaal Craton (component codes with references, see Table: 10), as well as overprints associated with

the Namaqua-Natal Mobile Belt, with component D1 and D2 associated with the Hlagothi Complex and a NE-

trending dolerite dyke (NE050); A and B associated with overprinting from the post-Pongola granitoids in the

Hlagothi Complex; F with the Ventersdorp volcanics in ENE-trending dolerite dykes, and C1, C2 and C3 with

overprinting from the Namaqua-Natal orogenic belt in the Hlagothi Complex and the two SE-trending dolerite

dykes respectively


- 131 -

8.2. Geochemistry and petrogenesis

There is a geochemical separation between predominantly NE-trending depleted tholeiitic

basalt dykes (NE030), and more enriched, borderline tholeiitic to calc-alkaline dykes of other

trends. This includes ca. 2.65 Ga ENE-trending dykes, ca. 2.95 and 2.87 Ga SE-trending dykes

and NE-trending dolerite dykes (NE050) which possibly have an age of ca. 1.90 Ga. This is

consistent with more aeromagnetically distinct NE-trending dykes having higher modal iron-

titanium oxide contents, compared to more quartz-bearing ENE-, NE- and SE-trending

dolerite dykes (Klausen et al., 2010). However, these dykes are more resistant to weathering

and likely to form geomorphological ridges; providing a means to roughly discriminate

between Archean and Proterozoic dykes in the field according to Klausen et al. (2010).

Multi-element statistical discrimination and spider diagrams elaborate on the calc-alkaline

affinity of Archean ENE-, NE- and SE-trending dykes, which typically have higher LIL element

concentrations and steeper REE patterns, and where Sr/V and La/Yb ratios most clearly

distinguish these from the Proterozoic NE-trending tholeiites (Klausen et al., 2010).

Finally, a geochemical comparison with published data on coeval volcanic rocks

indicates according to Klausen et al. (2010) that:

The ca. 2.95 Ga and ca. 2.87 Ga SE-trending (and NE050) dolerite dykes, are feeders to

Nsuze and Mozaan Group lavas.

The ca. 2.65 Ga ENE-trending dykes resemble Allanridge Formation lavas more than

other Ventersdorp Supergroup lavas, and some NE-trending dykes may also be related

to Ventersdorp age events (NE030), although they are much more primitive in

composition. Geochemically, these NE030 dykes resemble the Mazowe dolerites on

the Zimbabwe Craton, which have been associated with the ca. 1.90 Ga event in the

past based on palaeomagnetic studies.

Some ca. 1.90 Ga NE-trending dykes might belong to ca. 1.90 Ga Soutpansberg-

Mashonaland LIP that extends across both the Kaapvaal and Zimbabwe cratons and

includes coeval lava remnants within both the Soutpansberg and Olifantshoek Group

based on geochemistry and palaeomagnetism.


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These constraints and inferences on the complex array of dykes across the south-

easternmost Kaapvaal Craton both complement and partly contradict existing tectonic

models for the major igneous events during the Mesoarchaean to Palaeoproterozoic.

Compositional variability among SE-, NE- and ENE-trending dolerite dykes as well as

the Hlagothi Complex is indicative of increasing concentration of incompatible elements

(e.g., Sr, V) with decreasing MgO content, which is consistent with earlier ilmenite

fractionation amongst the more borderline tholeiitic and calc-alkaline suites as was noted by

Klausen et al. (2010). However, there can be considerable geochemical range within a single

dyke, as was noted by Hunter and Halls (1992), with geochemical samples collected from

the same dolerite dyke outcrop, as well as by Maré and Fourie (2012) who noted significant

contamination between SE- and NE-trending dolerite dykes in the Badplaas-Barberton

region. In particular, a highly xenolithic or phenocrystic dyke can clearly demonstrate the

dyke’s geochemical heterogeneity, without any apparent differentiation trend, probably due

to variable contamination from xenoliths from the local host rock, or by granophyric quartz

along the dyke-host rock contact zone. Uncritical sampling by Hunter and Halls (1992) could

be explained through the sampling of serpentine pseudomorphs after olivine and

orthopyroxene crystals from ultramafic xenocrysts from the nearby Barberton Greenstone

Belt further to the north for example, as was noted by Klausen et al. (2010).

In addition, apart from obvious geochemical separation amongst the Hlagothi

Complex and different dyke trends, La/Yb was noted to be a slightly better discriminator

between basaltic andesite ENE-trending dykes and the basalt to basaltic andesite SE-

trending dykes in the KwaZulu-Natal area for example, and those located further north. In

addition, NE-trending dolerite dykes (NE030) from northern KwaZulu-Natal have slightly

lower La/Yb than NE-trending dykes from the Black Hills area. However, many dykes and

volcanic sequences on the craton show large amounts of possible crustal contamination, or

having a very enriched geochemical reservoir beneath the Kaapvaal Craton. The ca. 2.65 Ga

NE030 generation of NE-trending dolerite dykes in northern KwaZulu-Natal do not,

however.

The separation between compatible and incompatible elements among the Hlagothi

Complex and dolerite dykes is consistent with magma differentiation due to the


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fractionation of chromium spinel, olivine and possibly pyroxene. The tholeiitic peridotitic

phases of the Hlagothi Complex show higher concentrations of compatible elements,

whereas borderline tholeiitic to calc-alkaline SE-, ENE- and NE-trending dolerite dykes, and

the gabbroic phases of the Hlagothi Complex have higher concentrations of LIL elements.

This discrimination is consistent with inferences based on rock type characteristics and

petrography, where more calc-alkaline suites typically are LIL enriched (e.g., McCulloch and

Gamble, 1991) and experienced earlier iron-titanium oxide fractionation (e.g., Kuno, 1968)

than tholeiitic suites. In addition, geochemically, the Hlagothi Complex shows two or

possibly three types of basaltic and basaltic andesitic compositions, which corresponds to

the meta-peridotites and meta-pyroxenites/gabbros respectively, with very little

fractionation trend seen. Using immobile element discrimination plots, the same trend (or

lack thereof) is seen, which also provides evidence for little loss of alkalis during alteration,

and which is further affirmed using the alteration box plot of Large et al. (2001). This

alteration box is only applicable to mafic lithologies however. Petrographically, there is good

evidence for high levels of serpentinisation in the peridotites and chloritisation in the

peridotites and pyroxenites. However, such alteration could be associated with syn-

magmatic alteration or at most upper greenschist facies metamorphism, with no new crystal

growth of primary pyroxenes, amphiboles or garnets indicative of amphibolite facies

metamorphism. The composition types for the Hlagothi Complex bear no trend or similarity

to each other, although du Toit (1931) postulated that the Hlagothi Complex consisted of

the differentiating products of a single reservoir. It is quite possible that two separate pulses

of magma following the same magmatic pathways at slightly different times led to the

emplacement of the complex, which may be supported by the age difference of 4 to 8 Ma

accounting for 95% confidence interval errors between the Hlagothi Complex and SE-

trending dolerite dykes in the greater region. However, dating of the ‘younger’ pyroxenites,

gabbros and ‘feeder’ dolerite dykes would be needed to confirm this. The high MgO (Mg#)

pulse of the complex is geochemically similar to that of komatiites, as was noted by

Groenewald (1984), who did some basic petrogenetic modelling of the complex. This

provides evidence for a more primitive, tholeiitic and komatiitic mantle signature, which a

mantle plume could provide for example (Ernst and Buchan, 2003). It is worth noting that

such a geochemically primitive magma is absent from the Nsuze and Dominion groups. This

would imply a different source of magma during these events. Cole (1994), Groenewald


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(1984) and Gold (1994), however, did note the presence of ultramafic dykes and sills within

the Pongola Supergroup stratigraphy in the whole area, as well as within the basement

granitoid-greenstone terrane. These intrusions have not been studied in any great detail in

order to affect a comparison, or to compare with other such intrusions further to the north,

such as the Thole Complex in south-eastern Mpumalanga and Swaziland, which in itself has

not been studied in any significant detail. The low MgO (Mg#) end members of the Hlagothi

Complex do resemble the Pongola and Ventersdorp dykes, however, as well as portions of

the Usushwana Complex. In addition, there are high MgO lava flows at the base of the

Allanridge, as well as within the Klipriviersberg lavas, which are classed as komatiitic.

8.3. Correlation to strata-bound igneous units

Field evidence, petrographical descriptions and geochemical discriminations in tandem with

U-Pb baddeleyite ages are for the most part consistent with the presence of four or five

major Precambrian mafic dyke swarms on the south-easternmost Kaapvaal Craton in

addition to the Hlagothi Complex. Some of these are coeval with lava formations within

equally significant volcanic successions in the Mesoarchaean to Palaeoproterozoic South

African stratigraphy. In addition, as a result of palaeomagnetic investigations on the same

dyke swarms in the Kaapvaal Craton, at least four or five different components have also

been recognised apart from present local field directions, in addition or complimentary to

components already established by Lubnina et al. (2010). Their mean directions are

summarised below, according to Klausen et al. (2010), Lubnina et al. (2010), Olsson et al.

(2010) and Söderlund et al. (2010):

The ca. 2.95 Ga dykes within a SE-trending swarm in the Barberton-Badplaas area are

potential feeders to lavas within the Nzuse Group.

The ca. 2.65 Ga dykes within a radiating NE-, E to SE-trending swarm in the Black Hills,

Rykoppies and Barberton-Badplaas areas are potential feeders to lavas within the

Ventersdorp Supergroup or proto-basinal fills.

The ca. 1.90 Ga dykes within a NE-trending swarm are potential feeders to lavas

within the Soutpansberg Group in the Black Hills area.


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Klausen et al. (2010) noted that the Archean dykes appear to have been emplaced towards

the end of each, roughly coeval, volcanic event, whereas poorer age constraints on

Soutpansberg Group lavas preclude more accurate correlations with ca. 1.90 Ga feeders. A

decision needs to be made on whether to include any of northern KwaZulu-Natal’s SE-, NE-

and ENE-trending dykes in any of the above listed major igneous events, as well as which

dykes belong to other igneous events.

8.3.1. Correlation with the ca. 2.95 Ga Nsuze Group dykes and lavas

Uken and Watkeys (1997) considered the SE-trending dyke swarm, later dubbed the

Barberton-Badplaas dyke swarm by Olsson et al. (2010), to be associated with rifting at ca.

2.95 Ga. They linked these dykes to the volcanic units of the Nsuze Group, as well as

possibly the Dominion Group further to the west. However, their interpretation was based

on the observation that the Barberton-Badplaas dyke swarm structurally predates the

protobasinal Godwan sequence at the base of the Transvaal Supergroup. Furthermore,

Hunter and Halls (1992) noted that many SE-trending dykes cut ca. 3.00 Ga old tabular

granitoids such as the Mpuluzi and Heerenveen batholiths, but are not abundant in the ca.

2.7 Ga Mbabane Granite, for example, thus providing a maximum and minimum age for a

majority of SE-trending dykes. This, however, is not true for all of the SE-trending dolerite

dykes in the region. The ages of the two Barberton-Badplaas dykes determined by Olsson et

al. (2010) of 2966 ± 1 Ma and 2967 ± 1 Ma, lie within these estimated age-constraints, and

are more precisely synchronous with a 2968 ± 6 Ma porphyritic rhyolite, near the top of the

Agatha lavas (Mukasa et al., 2013). In the Pongola Supergroup, the ca. 4.6 km thick Nsuze

Group is made up of six formations, dominated by volcanic rocks and subordinate

sedimentary strata, specifically the Nhlebela/Pypklipberg, Agatha and Ekombe volcanic

formations. These rocks were deposited/extruded within a period of approximately 40

million years, bracketed between the U-Pb age of the 2985 ± 1 Ma and 2984 ± 3 Ma Agatha

lava unit (Hegner et al., 1993; 1994) and a 2934 ± 114 Ma and 2940 ± 22 Ma ages for the

same unit (Hegner et al., 1984). No ages have been obtained on the Nhlebela/Pypklipberg

and Ekombe units, and it could significantly enhance or invalidate stratigraphic correlations


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both within the Nsuze basin, as well as with the Dominion basin to the west if better age

constraints could be obtained.

The dykes were characterised geochemically by Hunter and Halls (1992) as belonging

to a group of low-Ti dykes, which contrast to high-Ti dykes, which were inferred to be

Palaeoproterozoic in age. Olsson et al. (2010) dated both generations, and the coeval age of

ca. 2950 Ma on both suggests that both low-Ti and high-Ti dolerite dykes are age

equivalents to the Nsuze lavas – that is, both groups can be linked to the same Pongola

rifting event. This interpretation could only tentatively be extended to encompass all the SE-

trending dykes in this part of the craton, which furthermore appear to converge onto the

present exposure of Nsuze lavas in the south-easternmost part of the craton, an

interpretation also supported by geochemical and palaeomagnetic correlations of Klausen et

al. (2010) and Lubnina et al. (2010). Klausen et al. (2010) assigned ages of ca. 2.95 Ga to

ENE-trending dolerite dykes in the south-easternmost region on the basis of geochemistry.

This argument is invalidated by the palaeomagnetic directions obtained on the ENE-trending

dykes in this region by both Lubnina et al. (2010) and the component F obtained in this

study. In addition, field, petrographic and geochemical arguments appear to favour the

possibility that the area is composed of two slightly different SE-trending dolerite dyke

orientations. At least one generation is cross-cut by ENE-trending dykes, which are

tentatively thought to be either ca. 2.95 by Klausen et al. (2010) or ca. 2.65 Ga by Lubnina et

al. (2010). The age of 2874 ± 2 on one SE-trending dolerite dyke in this study on this south-

easternmost region provides a link to the Hlagothi Complex however. This complicates the

tentative age assignment done by means of palaeomagnetic directions and geochemistry

correlation.

Lava samples from the Nzuse Group (Armstrong et al., 1986; Wilson and Grant,

2006) define a continuous sub-alkaline series that ranges from basaltic andesites to

rhyolites. Despite their resemblance to compositionally continuous volcanic suites within

present-day subduction-zone settings (McCulloch and Gamble, 1991), the Nsuze lavas are

not as depleted in iron and thereby straddle the calc-alkaline to tholeiitic boundary. Basaltic

andesite samples bear the closest resemblance to ca. 2.95 Ga SE-trending dykes in the

Barberton-Badplaas area, as well the SE-trending dykes within northern KwaZulu-Natal.

Some of the basalt to basaltic andesite SE-trending dolerite dykes in northern KwaZulu-Natal


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also define a more MgO-rich group that lines up with a more tholeiitic lava trend, whereas

more andesitic SE-trending dykes in the Barberton-Badplaas and northern KwaZulu-Natal

can possibly overlap the more calc-alkaline trend defined by the Allanridge Formation in the

Ventersdorp Supergroup.

Figure: 62 – SE-trending dolerite dykes in northern KwaZulu-Natal compared to ca. 2.95 Ga SE-trending

dolerite dykes and volcanics in the Barberton-Badplaas region with data from Klausen et al. (2010), Armstrong

et al. (1986) and Wilson and Grant (2006)

A correspondence analysis by Klausen et al. (2010) from the Nsuze Group lavas and

the ca. 2.95 Ga SE-trending dykes identified lava samples that were compositionally most

similar to the dykes. From these, a few samples with the highest La/Yb display multi-

element patterns that partially overlap the ca. 2.95 Ga dykes in both northern KwaZulu-

Natal and in the Barberton-Badplaas area. The ca. 2.95 Ga dykes in the Barberton-Badplaas

area and in northern KwaZulu-Natal generally have lower heavy rare earth element

compositions than the Nsuze Group lavas (see Fig. 62). Instead, the same Nsuze Group lavas

are slightly better matched with ESE-trending dykes in the Barberton-Badplaas area

according to Klausen et al. (2010), which also have the most similar La/Yb ratios; this opens

up the possibility that this swarm is ca. 2.95 Ga in age according to Klausen et al. (2010),

contradicting arguments by Lubnina et al. (2010), and the ca. 2.87 Ga obtained herein.


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Figure: 63 – Palaeomagnetic results for the Pongola Supergroup volcanics and related intrusives

Palaeomagnetic results of Lubnina et al. (2010) identified a component named P (with

corresponding pole position named BAD) in some of the SE-trending dolerite dykes already

dated by Olsson et al. (2010) in the Barberton-Badplaas area (see Fig. 63). These SE-trending

dolerite dykes were studied palaeomagnetically in the south-eastern and south-easternmost

areas of the Kaapvaal Craton. Therefore Lubnina et al. (2010) concluded that the ca. 2.95 Ga

age of these dykes applies also to the age of their magnetisation. A palaeomagnetic pole,

recalculated from this component (BAD), is located in high southern latitudes. This direction

is similar to that obtained for the ca. 2.95 Ga Nsuze basalts (NB), also obtained by Lubnina et

al. (2010) in the south-easternmost Kaapvaal Craton. These ca. 2.95 Ga poles are about 65°

to the west from the ca. 2.70 Ga pole obtained on the Allanridge Formation by Strik et al.

(2007) and de Kock et al. (2009), as well significantly different from the pole obtained from

the Agatha basalts by Strik et al. (2007). No Nsuze aged magnetic directions were obtained

in this study, despite drilling the same SE-trending dolerite dyke (NL-13) of Lubnina et al.

(2010). However, the dyke in question in this study (AG-J) was drilled closer to three

intersecting dykes, thus possibly explaining the differences in results. The drill holes from

the palaeomagnetic sampling of Lubnina et al. (2010) were not observed, and must have

been done further afield, which could also explain the better results away from the possible

baking and re-magnetisation of the other dolerite dykes seen in the site drilled in this study.

It is clear that there are probably two generations of SE-trending dolerite dykes in northern


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KwaZulu-Natal, conforming to both ca. 2.95 and 2.87 Ga ages. This hypothesis, however,

needs further study.

8.3.2. Correlation with the ca. 2.87 Usushwana and Thole layered complexes, as well as

Mozaan and Witwatersrand lavas

The Thole Complex, located further north in the Swaziland region of the main Pongola Basin,

consists of sills emplaced at several stratigraphic levels into the Pongola Supergroup,

Usushwana Complex and basement granitoids (Hammerbeck, 1982). These sills are also

layered, consisting of harzburgites at the base grading up into pyroxenites, gabbros and

norites. The complex bears the same two distinct compositional groupings as the 2866 ± 2

Ma Hlagothi Complex (Chapter 5), and has been recognised to be a potential correlative of

the Hlagothi Complex (Groenewald, 2006). These sills have also seen extensive alteration,

despite being far removed from the cratonic margin and its related metamorphism and

deformation. This would suggest alteration similar to that experienced by the potentially

synchronous Hlagothi Complex as postulated by Groenewald (1984), provided that the

alteration was not caused by known post-Pongola 2850 to 2650 Ma granites in the region.

Flood basalts of the Mozaan and Witwatersrand basins represent another prominent

magmatic event (or events) of a possibly similar age. Hammerbeck (1982) and Nhleko (2003)

noted that flood basalts present in the Mozaan Group of the Pongola basin (i.e., the

Tobolsk, Gabela and Ntanyana lavas) were produced from fissure eruptions in a continental

setting. Lavas of similar stratigraphic age are represented in the adjoining upper West Rand

Group and middle Central Rand Group of the Witwatersrand basin by the Crown and Bird

lavas. A 2914 ± 14 Ma age was reported by Armstrong et al. (1991) for the Crown lava in the

Central Rand Group, but this age should be treated as a maximum age. These lavas may

have seen crustal contamination, as they are geochemically distinct from the Mozaan lavas,

and are of dacitic composition. In addition, zircon grains were few, suffered lead loss and

were noted to consist of a variety of morphologies suggesting they may be xenocrysts. In

addition, detrital zircons obtained in stratigraphically lower successions in the greater basin

produced zircons of younger ages, further validating this argument that the zircons are

xenocrystic (see Fig. 64). Beukes and Cairncross (1991) argued that the Mozaan and


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Witwatersrand groups represent distal and proximal correlatives. A likely setting for these

continental flood basalts may be a short-lived transient mantle plume(s). These lavas are

located in the upper Mozaan Group of the Pongola Supergroup and the eastern half of the

upper Witwatersrand Supergroup closest to the Pongola Supergroup, especially in the case

of the less laterally extensive Bird lava. The Crown lava is much more widespread, however,

thinning toward the west and south. This suggests a plume centre in the Pongola basin

region in the east to south-east. One observable effect of this plume would be regional

uplift associated with doming above the plume (e.g., Sengör, 2001).

Figure: 64 – Detrital zircon ages ranges with peaks shown in black in the Witwatersrand and Mozaan basins

after Nhleko (2003). Data for the Crown and Tobolsk lavas are not detrital, however, and reflect possible

emplacement ages. The sedimentary strata (yellow) are seen to be interrupted by two volcanic formations

(grey). Red crosses are indicative of no data being available for the lava flows or sedimentary strata

The development of small local unconformities in the region’s sedimentary successions was

followed by the eruption of flood basalts, providing possible proof of uplift from a mantle

plume before the volcanism.


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The 2866 ± 2 Ma Hlagothi Complex gabbros and the 2874 ± 2 Ma generation of SE-

trending dolerite dykes in the south-easternmost Kaapvaal Craton can be shown to be

geochemically similar to the Mozaan lavas. In terms of major elements they show similar

amounts of SiO2, but the lavas are richer in Na2O and K2O, which possibly relates to

alteration, fractionation or contamination. Using trace element geochemistry, the Mozaan

lavas show an almost identical signature to the gabbroic phases of the Hlagothi Complex,

although they are more enriched generally, which is common with greater fractionation. In

addition, both magmatic events exhibit negative Nb-Ta anomalies, as well as a negative Sr

and anomalies, with a minor negative Ti anomaly. One exception is Th and U. Th is depleted

in the Mozaan lavas compared to Hlagothi Complex gabbros. U is enriched in the Mozaan

lavas compared to the Hlagothi Complex gabbros (see Fig. 65). These signatures, however,

are similar to many such volcanic successions and their plutonic equivalents across the

Kaapvaal Craton from the Mesoarchaean to the Jurassic, and show a possible contaminated

crustal source or derivation from an enriched sub-arc-like mantle wedge.

Figure: 65 – Geochemistry of the Hlagothi Complex in northern KwaZulu-Natal compared to Mozaan Group

volcanics with data from Nhleko (2003)

This signature is found even in rocks of known rift-related tectonic regimes such as the

Karoo LIP (Duncan, 1987; Klausen et al., 2010). Tectonic setting diagrams also demonstrate

emplacement within an arc-like regime.


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In addition, a magnetic component D was determined for the Hlagothi Complex once

the overprint component C was removed during demagnetisation (see Fig. 66). This

component shows no similarity to similar-aged magmatic events from the Archaean across

the Kaapvaal Craton. An exception is the ca. 2.05 Bushveld Complex, which is of different

age. In addition, the same magnetic component D was also obtained from a NE-trending

dolerite dyke located further to the north of the Hlagothi Complex. However, in the absence

of a conclusive palaeomagnetic field test, whether this component represents a primary

direction and thus a correlative to the Hlagothi Complex or whether it is an overprint

direction related to the intrusion cannot be determined. However, the influence of the

Hlagothi Complex across the south-eastern region of the craton can therefore not be

understated, with the existence of numerous possible similarly aged intrusions, volcanic

successions and thermal overprints in potentially older strata or intrusions.

Figure: 66 – Palaeomagnetic results for the Hlagothi Complex (D1) and NE-trending dyke (D2)

The complexity of the intrusions of northern KwaZulu-Natal is exemplified by the NE-

trending dykes (NE050 and NE030), whose ages indicate connections to ca. 2.87, 2.65 and

1.90 Ga magmatic events across the craton. This adds greater complexity into this area in

particular.


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8.3.3. Correlation with the ca. 2.65 Ga Ventersdorp dykes and lavas

The scarcity of geochronological data from the upper Ventersdorp and lower Transvaal

Supergroups makes the correlations of dolerite dyke ages of Olsson et al. (2010) difficult.

This is especially true in the absence of ages on the Allanridge lavas as the stratigraphically

youngest unit within the Ventersdorp Supergroup, as well as poor ages existing for lavas

within the Wolkberg and Godwan proto-basinal fills, for example. Indeed, speculation exists

on whether ages of the basal 2714 ± 8 Ma Klipriversberg lavas and Platberg lavas at 2708 ± 5

Ma of Armstrong et al. (1991) are correct, because they are at odds with the potentially

correlatable Derdepoort lavas at 2782 ± 5 Ma (Wingate, 1998). The Hartswater volcanics can

also be correlated with the Platberg volcanics, with ages of 2733 ± 3 Ma and 2724 ± 5 Ma

(de Kock et al., 2012), further casting doubt on the ages of 2714 ± 8 Ma and 2708 ± 5 Ma of

Armstrong et al. (1991). The correlation of the ca. 2701–2659 Ma Rykoppies dykes of Olsson

et al. (2010) and Olsson (2012) to the basaltic andesite lavas of the Allanridge Formation

may be possible, since this uppermost part of the Ventersdorp Supergroup is younger than

its 2709 ± 5 Ma Makwassie Formation from the middle part (Armstrong et al., 1991). This

inference agrees with a similarity in geochemical characteristics between the Rykoppies

dykes and lavas from the Allanridge Formation according to Klausen et al. (2010), although

the use of geochemistry in correlation for the volcanic sequences has to be cast in doubt,

especially given the geochemical similarities between the varieties of magmatic events,

particulary in the Archaean. Palaeomagnetic studies are in agreement however. The age of

2662 ± 3 Ma dates the emplacement of a younger sub-swarm within the Rykoppies dyke

swarm. All these dykes are interpreted at present to be of the same generation, despite the

significant scatter (Olsson, 2012). The later intrusive stage of the Rykoppies swarm coincides

with Rb-Sr whole rock ages known for the ‘protobasinal’ sequences of 2657 to 2659 Ma and

2664 ± 1 Ma for felsic lavas in the Buffelsfontein Group (SACS, 1993; Barton et al., 1995), at

the base of the Transvaal Supergroup.

The results of Olsson et al. (2010) and Olsson (2012) suggest that at 2.66 to 2.68 Ga,

the eastern Kaapvaal Craton was subjected to extension in an approximate north to south

direction, which allowed mantle-derived magmas to be emplaced as E- to W-trending dykes.

The timing of the Rykoppies swarm appears to coincide with a shift from mainly volcanic

rocks of the Ventersdorp Supergroup to predominately sedimentary rocks of the Transvaal


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Supergroup. According to Burke et al. (1985) and van der Westhuizen et al. (1991), the

Allanridge Formation lavas were formed during renewed rifting of the Kaapvaal Craton. The

rift structures of the Ventersdorp Supergroup have been assigned a NE-trend (Stanistreet

and McCarthy, 1991), but seismic investigations also reveal the presence of E- to W-trending

faults within rock units of the Ventersdorp Supergroup (Tinker et al., 2002). Eriksson et al.

(2001) proposed that the ‘protobasinal’ successions at the base of the Transvaal Supergroup

were deposited within a rift-type extensional tectonic setting and inferred a tectonic model

where the ‘protobasinal’ successions accumulated in ENE- to WSW-trending rift systems.

The age correlations of the younger sub-swarm of dykes with the ‘protobasinal’ sequences

and the proposed correlation between the older sub-swarm and the Allanridge Formation

support the idea that the 2.70 to 2.65 Ga Rykoppies Dyke Swarm marks the onset of rifting

of the Transvaal Basin, at a time when the rift system also changed direction from NE- to E-

trending. This interpretation is supported in the south-easternmost region of the Kaapvaal

Craton, with one generation of NE-trending dolerite dykes having been dated by U-Pb in

baddeleyite at 2652 ± 11 Ma. In addition, the magnetic component ‘F’ bears a similar ca.

2650 Ma direction among the ENE-trending dolerite dykes in the region, being supported by

a baked contact test not only in this study, but also across similar ENE-trending dolerite

dykes on the south-easternmost Kaapvaal Craton, as was determined by Lubnina et al.

(2010); see discussion below. Although it is difficult to verify ages, it is certain that both E-W

and SW-NE directed tectonic stresses were applied during the termination of Ventersdorp

volcanism and the onset of filling of the Transvaal basin, and that these events were seen

across the larger Kaapvaal Craton, as well as being present in this south-easternmost area

too.

Hatton (1995) and Eriksson et al. (2002) regard the Ventersdorp Supergroup as the

product of a mantle plume, possibly related to a global mantle plume event at ca. 2.70 Ga

(Condie, 1998; 2001). Hatton (1995) stated that the Ventersdorp magmas were extracted

from the mantle at shallow depths (ca. 18 to 40 km), and subsequently underwent crustal

contamination, accounting for the calc-alkaline geochemistry. A fast lava extrusion rate up

through the lower half of the Ventersdorp Supergroup from 2714 to 2709 Ma is consistent

with typical LIPs containing flood basalts (Armstrong et al., 1991). However, these ages are

almost certainly incorrect (Wingate, 1998; de Kock et al., 2012). Therefore the proposition


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that the Rykoppies dyke swarm might be a product of prolonged persistent plume activity

may not be correct. A mantle plume origin would agree with the radiating pattern of the

Rykoppies swarm from Olsson et al. (2010; 2011). However, the presence of a primitive

mantle signature of the 2652 ± 11 Ma NE-trending dolerite dyke swarm in northern

KwaZulu-Natal would argue against a radiating pattern, even if it is more indicative of a

plume based on geochemistry. In contrast to the mantle plume model, Burke et al. (1985)

and Stanistreet and McCarthy (1991) proposed a rifting model related to the collision

between the Kaapvaal and Zimbabwe cratons at ca. 2.70 Ga, where the Rykoppies dyke

swarm was emplaced within a hinterland horst and graben setting. The 2671 ± 2 Ma U-Pb

age of the Matok Granite in the Southern Marginal Zone of the Limpopo Belt (Barton and

van Reenen, 1992) can be temporally linked to the Rykoppies swarm, inferring coeval

emplacement of magmas in the eastern and northern Kaapvaal Craton and granite

intrusions along the northern margin of the craton. Also, peak metamorphism of the

Southern Marginal Zone has been constrained to 2691 ± 7 Ma according Kreissig et al.

(2001). However, the timing of the Kaapvaal-Zimbabwe collision is debatable, because

studies show that metamorphism and deformation within the Central Zone of the Limpopo

Belt occurred much later, at approximately 2.00 Ga (e.g., Kamber et al., 1995; Buick et al.,

2006; van Reenen et al., 2008; Kroner et al., 1998; Jaeckel et al., 1997). There are,

furthermore, no major dyke swarms and LIPs older than 2.00 Ga that mutually exist in both

the Kaapvaal and the Zimbabwe cratons, supporting an amalgamation collision forming the

Kalahari Craton after ca. 2.00 Ga (e.g., Söderlund et al., 2010). The proposed late formation

of Kalahari Craton does not rule out a continental back-arc setting as a plausible tectonic

model for the Rykoppies Dyke Swarm at ca. 2.70 to 2.65 Ga according to Olsson et al.

(2010). Olsson et al. (2010) state that the Rykoppies swarm could have developed in a back-

arc rift system at the same time as the northern margin of the Kaapvaal Craton underwent

compressional tectonics during south-directed subduction of oceanic lithosphere. Further,

Olsson et al. (2010) stated that support for this model lies in the interpretation that the

‘protobasinal’ sequences at the base of the Transvaal Supergroup represent intra-cratonic

depositories, whose provenance can be linked to erosion of Limpopo Belt sedimentary

source rocks from work by Eriksson et al. (2002). This would suggest a greater complexity to

the tectonic setting of the craton during the closure of the Ventersdorp basin and the

opening of the Transvaal basin.


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Ventersdorp Supergroup lava analyses show that the oldest ca. 2782 Ma

Derderpoort lavas are correlatable with the Klipriviersberg lavas (Wingate, 1998).

Geochemically distinct lava formations within the overall more basaltic Klipriviersberg

Group match NE-trending dykes near the Johannesburg dome according to McCarthy et al.

(1990). More andesitic Allanridge Formation lavas are, on the other hand, closer in age to

the ca. 2.65 Ga radiating NE-, E- to SE-trending dolerite dyke swarm on the eastern Kaapvaal

Craton. Furthermore, the Allanridge lavas show the best compositional overlap with both

Rykoppies E-trending dykes and more evolved SE-trending dykes from the Badplaas-

Barberton area (see Fig. 67), by Klausen et al. (2010).

Figure: 67 – ENE-trending dolerite dykes in northern KwaZulu-Natal (black) compared to ca. 2.65 Ga E-trending

dolerite dykes in the Rykoppies region with data from Klausen et al. (2010). Both generations of north-east

trending dolerite dykes in northern KwaZulu-Natal are also illustrated (red and blue), and compared to NE-

trending dolerite dyke data by Klausen et al. (2010) in northern KwaZulu-Natal. Geochemistry of the Allanridge

lavas of the Ventersdorp Supergroup is also shown with data from Crow and Condie (1988), Marsh et al.

(1992), Nelson et al. (1992) and Keyser (1998)

A similarity exists between more calc-alkaline Allanridge Formation lavas, most ca. 2.70 to

2.65 Ga E-trending dykes in the Rykoppies area, and MgO-poor SE-trending dykes from the

Badplaas-Barberton area (Klausen et al., 2010). The latter observation is consistent with

more evolved (andesitic) dolerites dykes being part of a radiating ca. 2.65 Ga pattern.

However, NE-trending ca. 2.65 Ga dolerite dykes have been dated on the south-easternmost


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region of the Kaapvaal Craton in this study, casting this proposed radiating pattern for ca.

2.70-2.65 Ga aged dykes and a mantle plume origin into doubt.

According to Klausen et al. (2010), the Allanridge Formation lavas plot relatively

close to variably contaminated samples from Rykoppies dyke swarm, as well as ca. 2.65 Ga

SE-trending dykes from the Badplaas-Barberton area. Apart from slightly more elevated Nb

and Sr, the basaltic andesites from the Allanridge Formation overlap the Rykoppies ridge

samples well. The ca. 2.65 Ga SE-trending dykes from the Badplaas-Barberton area have

overall lower but sufficiently similar trace element patterns compared to these lava

samples, suggesting that the entire radiating ca. 2.65 Ga swarm acted as a feeder system to

the Allanridge Formation according to Klausen et al. (2010). Klausen et al. (2010) made no

match based on geochemistry to dolerite dykes on the south-easternmost Kaapvaal Craton,

although the one generation of NE-trending dolerite dykes (NE030), as well as ENE-trending

dykes have a similar age and palaeomagnetism respectively. The ENE-trending dolerite

dykes in northern KwaZulu-Natal display a similar, albeit variable pattern with data from the

ENE-trending dolerite dykes in the Rykoppies area. However, a much more primitive ca. 2.65

Ga set of NE-trending dolerite exists (NE030), with the only dykes exhibiting a similar

geochemical trend coming from the Mazowe dykes on the Zimbabwe Craton – of speculated

ca. 1.90 Ga age, creating a great deal of controversy between the two sets of dykes trends,

ages and geochemistry. Work will need to be done in order to resolve this issue.

The dual polarity H component of Lubnina et al. (2010) is characteristic of E-trending

dolerite dykes of the Rykoppies swarm of the southeastern and eastern areas of the

Kaapvaal Craton. The primary origin of this component is supported by a positive contact

test. The age of these dykes is ca. 2.65 Ga, with this also bring the age of the magnetisation

(see Fig. 68). The RYK palaeopole, calculated from the H component, is close to the ca. 2.7

Ga pole for the Allanridge Formation (Strik et al., 2007; de Kock et al., 2009). However, the

Rykoppies direction is removed by about 20° from the Mbabane pluton direction dated at

2690 Ma (Layer et al., 1988; 1989). This pole, however, is not supported by field stability

tests, and can be left out from further discussion. The 2782 ± 5 Ma Derdepoort volcanics

have a similar direction, however (Wingate, 1998). The ENE-trending dolerite dykes on the

south-easternmost Kaapvaal Craton also show this same palaeopole direction, as shown in

this study (Component F), and in work done by Lubnina et al. (2010), indicative that


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Ventersdorp volcanic activity was present in the vicinity, with NE- and ENE-trending dolerite

dykes both showing evidence of it, even in the absence of Ventersdorp strata. The ca. 2.65

Ga NE-trending dolerite dykes, which show a primitive, uncontaminated mantle source

indicative of a plume, and perhaps a close proximity to one of the piercing points from

which the NE-trending swarm can be traced to the edge of the south-easternmost Kaapvaal

Craton. They have a very different geochemistry from the ca. 2.65 Ga ENE-trending dolerite

dykes shown above.

Figure: 68 – Palaeomagnetic results for the Ventersdorp volcanics and related intrusives, including the ENE-

trending dolerite dyke (F) and re-magnetised Hlagothi Complex (A and B)

8.3.4. Correlation with the ca. 1.90 Ga Soutpansberg dykes and lavas

The ca. 1.90 Ga dykes of Klausen et al. (2010) and Olsson (2012) are compared with lava

samples from the Soutpansberg Group (Crow and Condie, 1990; Bumby et al., 2001), as well

as sill samples from the underlying Waterberg Group (Hanson et al., 2004a), in addition to

Mashonaland sill and dyke samples, and Mazowe dyke samples from the Zimbabwe craton

(Stubbs et al., 1999; Stubbs, 2000). In contrast to the large compositional range among

Archean lavas (i.e., from basaltic andesites to rhyolites), almost all of the ca. 1.90 Ga lavas,

sills and dykes previously studied are exclusively basaltic tholeiites, which match most NE-


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trending dykes from the Black Hills area, but not the dated ca. 2.65 Ga NE-trending dykes in

northern KwaZulu-Natal.

Figure: 69 – ENE-trending dolerite dykes in northern KwaZulu-Natal (black) compared to ca. 1.90 Ga

Waterberg and Mashonaland sills and dykes, including the ca. 1.90 Ga NE-trending dolerite dykes of the Black

Hills area. Interestingly NE-trending dykes in northern KwaZulu-Natal do not compare well to any dykes and

sills located further to the north, with the exception of the Mazowe dykes on the Zimbabwe Craton

Despite what is stated in Klausen et al. (2010), a poor geochemical comparison invalidates

the proposed link between all of these different units, and the NE-trending dykes from

northern KwaZulu-Natal. The exception is the Mazowe dolerite dykes, which have

essentially flat REE patterns unlike the slightly LREE enriched patterns seen in the other

dykes and sills of supposedly coeval intrusions (see Fig. 69). There is REE data available for

the Soutpansberg Group lavas, Waterberg sills (Hanson et al., 2004), Mashonaland sills

(Stubbs et al., 1999), and Mazowe dykes, however (Stubbs, 2000). Klausen et al. (2010) used

only intrusions in multi-elemental comparisons, because especially poor LIL element

matches suggest that lava samples may have been altered. Klausen et al. (2010)

furthermore stated that:

The Waterberg sills have slightly higher Sr/V.

Some Mashonaland sills tend to have lower Sr/V.


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Mazowe dykes have (La/Yb)N values that are as low as those of NE-trending dykes in

northern KwaZulu-Natal.

Thus ca. 1.90 Ga NE-trending dykes of the Black Hills area define a restricted sub-

group of roughly parallel patterns which overlaps a tight cluster of the Mashonaland sills, as

stated by Klausen et al. (2010), and resemble the ca. 1.93 to 1.87 Ga post-Waterberg

dolerites of Hanson et al. (2004a). The Mazowe dykes from the northern Zimbabwe Craton

have much more depleted patterns, which resemble the variably feldspar-phyric NE-

trending dykes from northern KwaZulu-Natal, such as NE030, in the south-easternmost part

of the Kaapvaal Craton. Age dating was done on the NE-trending dykes of northern

KwaZulu-Natal at ca. 2.65 Ga in this study, but these dykes may be coeval with the Mazowe

dykes based on geochemistry too, casting further doubt on the suggestion that are part of a

ca. 1.90 Ga age LIP (Wilson et al., 1987; Klausen et al., 2010).

Figure: 70 – Palaeomagnetic results for the Bushveld Complex and post-Bushveld volcanics and related

intrusives, including also the Vredefort impact and Black Hills dykes (BH), in comparison to the pole obtained

by Lubnina et al. (2010) in northern KwaZulu-Natal from the NE-trending dolerite dykes there. See Table 10 for

sources of other poles

The component (M) of Lubnina et al. (2010) was found in NE-trending Black Hills

dolerite dykes from the north-eastern, south-eastern and south-easternmost areas (see Fig.

70). A palaeomagnetic pole, recalculated from this component (BHD), is similar to the 1870


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Ma Post-Waterberg dolerites (Hanson et al., 2004a), and also the 2990 to 2875 Ma

Ushuswana Complex (Layer et al., 1988; Olsson, 2012). The age of the dykes according to

Olsson (2012) and Söderlund et al. (2010) is ca. 1.90 Ga in the Black Hills area of the north-

eastern Kaapvaal Craton. However, over 500 km away, a NE-trending dolerite swarm on the

south-easternmost Kaapvaal Craton displayed a similar magnetic component (M), and the

primary origin of magnetisation in these dykes was supported by positive contact,

conglomerate and reversal tests. Lubnina et al. (2010) concluded that the age of

magnetisation is ca. 1.90 Ga in this region as well, in stark contrast to the ca. 2.65 Ga age,

complicating the already complex array of dolerite dyke swarms across this region of the

craton. NE-trending dolerite dykes studied for palaeomagnetism in this work, however,

showed chaotic distribution and the data thus could not be used, but both the

geochronology (done herein) and palaeomagnetism by Lubnina et al. (2010) have to be

accepted as correct. In the absence of further more detailed work, NE-trending dolerite

dykes may be composed of up to two, and possibly three dolerite dyke events at possibly ca.

2.87, 2.65 and 1.90 Ga. Further study is needed in order to resolve this issue too.

8.3.5. Dyke swarms of potentially other ages

The variably feldspar-phyric NE-trending dykes from northern KwaZulu-Natal have been

shown to be as equally depleted as the Mazowe dykes (Stubbs, 2000). However,

uncertainties regarding the presumed ca. 1.90 Ga age of the Mazowe dykes (Wilson et al.,

1987), the ca. 2.65 Ga age presented for the NE-trending dykes in northern KwaZulu-Natal,

and a greater than 1000 km separation between the Mazowe and NE-trending dykes

renders such a correlation difficult, although the palaeomagnetism and geochemistry

appears similar. Thus, unless the ca. 1.90 Ga Mazowe dyke age is incorrect and/or the ca.

2.65 Ga LIP was much more extensive, and was fringed by depleted mantle source areas,

more data is needed to resolve the correlation and tectonic setting of the Mazowe and NE-

trending dykes in northern KwaZulu-Natal, which may belong to a newly recognised

separate swarm.

In addition to the complexities observed above, Lubnina et al. (2010) identified

another component across the eastern and south-eastern Kaapvaal Craton. A palaeopole


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(NSA) was calculated from dykes of the south-eastern area of the craton, that is close to the

pole of Ongeluk Formation lavas with an age of ca. 2.22 Ga (Evans et al., 1997) and the

Gamagara Formation lavas with an age of ca. 2.13 Ga (Evans et al., 2002). Both these

formations occur within the greater Transvaal Supergroup, suggesting that magmatism of

this age may be extended across the larger craton.

The SE-trending dolerite dykes in northern KwaZulu-Natal are connected with two

events: one at ca. 2.95 Ga, and the other at ca. 2.87 Ga. It is clear that the relationship, if

any, between these dyke swarms needs to be resolved from the conflicting geochronology

and palaeomagnetism. This is also the case for the ages and palaeopoles amongst the ca.

2.65 Ga and ca. 1.90 Ga NE- and ENE-trending dolerite dykes discussed above and below.

More data is needed in order to resolve the age of ENE-trending dykes in the

northern KwaZulu-Natal and Rykoppies areas. Also, the dykes in the Black Hills area too,

because of their more pristine petrography, may have an age younger than the area’s ca.

1.90 Ga NE-trending dykes (Lubnina et al., 2010). One possibility is that these more pristine

dykes are coeval with the ca. 1.10 Ga Umkondo large igneous province (Hanson et al., 2006),

which is found extensively across the whole larger Kalahari Craton, and which may be

present in northern KwaZulu-Natal, too.

Another component is characteristic of SSE-trending dykes in the south-eastern

Kaapvaal Craton, as well as dykes and sills in the south-easternmost area, and which was

observed as a low- to medium temperature overprint within ENE-trending dykes within this

study. A palaeomagnetic pole (E), was recalculated from this component, and is close to the

present local field, as well as the ca. 0.18 Ga pole for a Karoo Dolerite (Hargraves et al.,

1997), with dykes and sills related to this event extensive across the south-easternmost

Kaapvaal Craton.

8.4. Tectonic model and a new large igneous province


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In summary, the petrological and palaeomagnetic discriminations between borderline

tholeiitic to calc-alkaline SE- and ENE-trending dolerite dykes, as well as the two generations

of tholeiitic and calc-alkaline NE-trending tholeiitic dykes are consistent with precise U-Pb

baddeleyite ages and palaeomagnetic results obtained herein, and by Olsson et al. (2010)

and Lubnina et al. (2010). This indicates the presence of a ca. 2.95 and ca. 2.87 SE-trending

dolerite dyke swarms on the south-easternmost Kaapvaal Craton, of which the ca. 2.95 Ga

dolerite dyke swarm is also seen further to the north in the Barberton-Badplaas area. In

addition, two pulses of ca. 2.65 NE- to ENE-trending dolerite dykes are seen across the

south-easternmost part of the craton which have different geochemistry related to NE- and

ENE-trending stresses related to the termination of the Ventersdorp and opening of

Transvaal basins across the greater craton.

8.4.1. The Nsuze igneous event

SE-trending dolerite dykes primarily outcrop across the south-eastern Kaapvaal Craton in

the Barberton-Badplaas area, as well as the south-easternmost area in the vicinity of

Vryheid-Melmoth, with this ~250 km wide swarm most likely coinciding with the

reconstructed extent of the Pongola Supergroup cover (Weilers, 1990). Sub-parallel trending

dykes and rift basin structures (Hunter and Halls, 1992; Uken and Watkeys, 1997), similar

dyke and lava ages (Hegner et al., 1994; Olsson et al., 2010), and roughly comparable

compositions and palaeomagnetic poles (Klausen et al., 2010; Lubnina et al., 2010)

strengthen previous suggestions of a co-genetic lava-feeder system within one of the

world’s oldest volcanic intra-continental rifts (Burke et al., 1985). An apparent

predominance of borderline tholeiitic to calc-alkaline basaltic and basaltic andesite

compositions in SE-trending dykes furthermore suggests that more evolved Nsuze Group

lavas generally differentiated in more elevated shallow crustal magma chambers, like the

elongated and sub-parallel SE-trending ca. 2990 Usushwana Complex that appears to have

been emplaced at the base of the Nsuze Group lava pile. This would also account for the

apparent lack of more primitive magma compositions within the volcanic succession and its

plutonic equivalents (Armstrong et al., 1986).


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Coeval granitoid ages across most of the Kaapvaal Craton’s north-eastern segment

(Olsson et al., 2010; Eglington and Armstrong, 2004; Zeh et al., 2009), are consistent with

active subduction along the craton’s northern margin, in conjunction with crustal

amalgamation onto the craton’s western margin along the Colesberg lineament (Poujol et

al., 2008; Schmitz et al., 2004). Mountain building along these margins could also provide

gold-rich sediments for south- to south-eastward directed fluvial systems into the combined

Witwatersrand-Mozaan basin (e.g., McCarthy, 2006). Such a tectonic setting has been used

to explain the calc-alkaline and andesitic character of Dominion Group lavas, and could

likewise explain the ca. 2.95 Ga Nsuze Group’s sub-alkaline series of basaltic andesites to

rhyolites, without any significant silica gap (e.g., Klausen et al., 2010; McCulloch and

Gamble, 1991). These roughly coeval lavas and associated SE-trending feeder dolerite dykes

could have been emplaced along the landward side of an active continental margin or within

a more inland continental back-arc setting (e.g., Olsson et al., 2010; Burke et al., 1985;

Winter and de La, 1987; Crow and Condie, 1988). Other geochemically based models do not

require such a subduction zone setting, favouring either crustally contaminated komatiitic

primary mantle melts according to Crow and Condie (1990), or basaltic primary melts from a

metasomatised lithospheric mantle (e.g., Klausen et al., 2010; Duncan, 1987; Marsh et al.,

1992). Both models could account for generation of LIL element enriched, borderline

tholeiitic to calc-alkaline and basaltic to andesitic magmas (with negative Nb, Ta anomalies

for example) within Archaean continental rifts across the Kaapvaal Craton. This includes the

Dominion, Nsuze and Ventersdorp lavas and their plutonic equivalents, which can also be

seen on the south-easternmost Kaapvaal Craton.

Structural evidence, including the SE-trending feeder dolerite dyke swarm, favours a

volcanic rift setting for both the Dominion (e.g., Lana et al., 2006) and Nsuze Group lavas

(Bickle and Eriksson, 1982). The north-west to south-east orientated Pongola volcanic rift is

also located very far from, and trends at a high angle to any subduction-zone along either

the northern or western margin of the Kaapvaal Craton (Burke et al., 1985). Furthermore,

the SE-trending swarm with north-westerly bifurcating dykes indicates emplacement from

the south-east (Hunter and Halls, 1992), and may have originated from an igneous plume

centre, near a hypothetical passive margin along the Kaapvaal Craton from which northern


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KwaZulu-Natal’s geochemically and palaeomagnetically similar SE-trending dykes also might

have been injected as part of a greater swarm, as suggested by Lubnina et al. (2010).

8.4.2. The Hlagothi igneous event

The 2866 ± 2 Ma sills and 2874 ± 2 Ma SE-trending dykes of the Hlagothi Complex, as well as

the Thole Complex, acted as the plumbing system that fed the Mozaan Group flood basalts.

The Usushwana Complex can also be shown to be composed of at least two magmatic

pulses between 2990 Ma and 2860 Ma (Olsson, 2012; Hammerbeck, 1982). The variation of

ages in the Usushwana Complex could be explained if some portions of the complex in fact

represent Hlagothi-Thole Complex age intrusions, and not the older 2990 Ma date obtained

by Olsson (2012). Shortly after the intrusion of the Hlagothi Complex, the whole region was

intruded by potassium-rich post-Pongola granites. These granites may have been the result

of partial melting induced by the culmination of the same thermal source as that which

generated the Hlagothi Complex.

The dated SE-trending dykes in the White Mfolozi inlier and the sills of the Hlagothi

Complex in this study, coupled with the Thole Complex and also possibly some gabbroic

phases within the Usushwana Complex, form a part of a newly recognised large igneous

province in the south-easternmost part of the craton (see Fig. 71). The extrusion of the

Mozaan lavas and potential correlatives in the central and eastern portions of the West

Rand Group and Central Rand Group of the Witwatersrand Supergroup may also belong to

this large igneous province. Geochemistry of the Hlagothi Complex and the SE-trending

dolerite dykes suggests two different pulses of magmatism, both with evidence of crustal

contamination or melting of a sub-arc-like mantle wedge. The geochemical evidence for an

arc-like tectonic setting and subduction is strong, but it is common for most volcanic rocks

of the Kaapvaal Craton, even those of known rifting settings, such as in the Nsuze and

Dominion groups (Klausen et al., 2010). There are also younger LIPs with arc-like signatures

acquired from enriched lithospheric mantle, such as the Karoo LIP, which was emplaced in a

known extensional tectonic regime (Duncan, 1987). Such a signature is most likely derived

from crustal contamination or contamination from metasomatised sub-continental

lithosphere, rather than reflecting an in-situ produced arc-like signature, and may also


- 156 -

explain the retention of LIL and depletion of HFS elements. It is also important to note that

the depositional environment of both the Nsuze and the Mozaan groups reflects an intra-

cratonic basin, and that erosional unconformities prior to volcanic deposition in the Mozaan

are suggestive of domal uplift before volcanism (see Fig. 72).

Figure: 71 – New magmatic barcodes for the eastern and western sides of the Kaapvaal Craton with ages from

a variety of sources discussed in the text, highlighting a new possible large igneous province


- 157 -

Figure: 72 – Tectonic setting for the proximal to distal Witwatersrand and Mozaan basin respectively, with a

large igneous province induced by a mantle plume leading to the upward coarsening sedimentary sequences in

the basin, culminating in flood volcanism being fed by a series of SE-trending dolerite dykes and layered

complexes


- 158 -

The basalt sequences in the upper Witwatersrand-Mozaan strata are thick, monotonous

sequences of lava, with little interflow of sediment suggestive of flood basalts (with the

exception of the Ntanyana lava unit), and were produced through fissure eruptions

(Hammerbeck, 1982; Nhleko, 2003). All these associated volcanic sequences, dykes, sills and

layered complexes occur over the eastern Witwatersrand block of the Kaapvaal Craton – an

area of approximately 100000 km2. This evidence is supportive of a LIP, following the

definition of Bryan and Ernst (2008).

8.4.3. The Ventersdorp igneous event

Roughly NE-, E- and SE-trending ca. 2.65 Ga dykes appear to be restricted within a nearly

150 km wide radiating dolerite dyke swarm that converges beneath a younger Transvaal

Supergroup basin (Olsson et al., 2011). More northerly-located ENE-trending dykes are

probably younger (less than ca. 1.90 Ga and likely ca. 1.10 Ga) in the north-eastern Kaapvaal

Craton, whereas tholeiitic ca. 2.65 Ga NE-trending dolerite dykes and calc-alkaline ENE-

trending dykes in the northern KwaZulu-Natal window may have fed the same roughly

coeval event(s), which is documented by geochronology and palaeomagnetism.

The Ventersdorp Supergroup lavas could have erupted in a foreland basin behind a

collision zone between the Zimbabwe and Kaapvaal cratons (McCarthy, 2006; Zeh et al.,

2009). Such a setting is supported by:

Roughly coeval metamorphism along the Limpopo Belt’s northern and southern zones

(Kramers et al., 2006)

A conjugate set of major transfer faults across the Kaapvaal craton, which is

compatible with approximately north-south directed compression (Stanistreet and

McCarthy, 1991).

The borderline tholeiitic to calc-alkaline character of Ventersdorp Supergroup lavas

and feeder dykes (Klausen et al., 2010).

However, this model is also questioned by:


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The ca. 2.0 Ga ages along the Limpopo Belt’s Central Zone (Kamber et al., 1995;

Mourie et al., 2008).

General uncertainties regarding age-determinations on fault displacements.

The above mentioned possibility that the calc-alkaline affinity of most intra-

continental basaltic rocks is inherited either from crustal contamination (Crow and

Condie, 1990) or the partial melting of a metasomatised continental lithosphere

(Duncan, 1987; Marsh et al., 1992).

Despite the plate tectonic setting, the rapid emplacement of the more basaltic

Klipriviersberg Group, including its komatiitic base flows, has also been attributed to the

presence of a mantle plume (Hatton, 1995; Eriksson et al., 2002). However, the older ages

on the Derderpoort basalts, as well as Hartswater rhyolites and tuffs cast doubt on this

(Wingate, 1998; de Kock et al., 2012)

A sequence of sub-swarms with differently trending feeder dykes and geochemistry

provides additional constraints on the tectonic evolution during the emplacement of the

Ventersdorp Supergroup. Firstly, NE-trending dykes that are geochemically matched to

individual lava formations within the Klipriviersberg Group around the Johannesburg dome

(McCarthy et al., 1990) reflect a south-west to north-east directed maximum compressive

palaeo-stress field that probably generated sub-parallel rift-structures (e.g., domino-block

faulting; van der Westhuizen et al., 2006) during their emplacement at 2782 Ma according

to Wingate (1998). The NE-trending dolerite dykes at ca. 2.65 Ga on the south-easternmost

Kaapvaal Craton could be coeval with these dykes, however they have a much more

primitive geochemistry and could instead be linked to the komatiitic flows at the base of the

Ventersdorp (Klipriviersberg lavas), or to the Mazowe dykes on the Zimbabwe Craton.

McCarthy et al. (1990) matched NE-trending dolerite dykes in the Johannesburg

Dome to the uppermost Lorain Formation of the Klipriviersberg Group, indicating a shift

towards a more east-west directed maximum compressive palaeo-stress field near the end

of Klipriviersberg Group lava eruptions, and not during Transvaal Supergroup times. Further

geochemical matching of ca. 2.80 to 2.65 Ga radiating NE-, E and SE-trending dolerite dykes

across eastern Kaapvaal with slightly more andesitic lavas from the younger Allanridge

Formation from less than 2708 Ma to greater than 2687 Ma according to Cheney (1996),


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conforms to McCarthy et al.’s (1990) latter palaeo-stress field however. Thereby, this

possibly identifies a regional shift towards an east-west orientated maximum compressive

stress field. This was maintained during the emplacement of the Platberg Group, Allanridge

Formation, Transvaal Supergroup (Eriksson et al., 2006), and the 2.05 Ga Bushveld Complex

(Cawthorn et al., 2006); i.e., for another ca. 600 million years. This would account for the ca.

2.65 Ga Ventersdorp palaeomagnetic component identified by Lubnina et al. (2010), on the

south-easternmost Kaapvaal Craton. This was seen in the ENE-trending dolerite dykes,

different from the ca. 2.65 Ga dolerite dykes in the same region. However, a supposed ca.

2.65 Ga radiating north-east, east to south-east pattern also converges toward an igneous

centre below the current eastern lobe of the Rustenburg layered suite, which may coincide

with the proposed mantle plume source for the Klipriviersberg Group lavas (Hatton, 1995;

Eriksson et al., 2002). This is not, however, supported by the ca. 2.65 Ga age of the NE-

trending dolerite dykes of south-easternmost Kaapvaal Craton.

8.4.4. The Soutpansberg-Mashonaland igneous event

Similar tholeiitic basalt compositions are consistent with a unified ca. 1.90 Ga swarm, which

may include some NE-trending dolerite dykes in the south-easternmost Kaapvaal Craton,

that have been matched palaeomagnetically by Lubnina et al. (2010) to ca. 1.90 Ga dykes

further north on the craton. The enriched NE-trending dykes in northern KwaZulu-Natal may

be part of a greater than 200 km wide swarm that extends across most of the eastern

Kaapvaal Craton already identified, partly hidden by coeval Soutpansberg Group deposits in

the north and truncated by the ca. 0.65 to 0.50 Ga East African (Mozambique)-Antarctic

Orogenic belt that bounds the eastern margin of the Kalahari Craton (e.g., Jacobs et al.,

2008). The locally ‘kinked’ pattern of more NNE-trending dykes across the eastern Transvaal

Supergroup and the underlying Kaapvaal segment (Uken and Watkeys, 1997) is more likely

to be primary (e.g., reflecting a regional stress that was distorted beneath the load of the

Transvaal basin, including the Bushveld Complex) rather than caused by any post-intrusive

left-lateral shear (Klausen et al., 2010).

No coeval NE-trending dykes appear to intersect the Zimbabwe Craton, because

Plumtree dykes are older (Söderlund et al., 2010) and 40Ar/39Ar ages (Jourdan et al., 2006)


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indicate that the ENE-trending Save-Limpopo dyke swarm is Mesoproterozoic, and therefore

most likely part of the Umkondo LIP (Hanson et al., 2006). Thus, only the Mashonaland sills

(Stubbs et al., 1999), and possibly some E-trending Mazowe dykes (Wilson et al., 1987;

Stubbs, 2000), are coeval with some of the NE-trending dolerite dykes on the south-

easternmost Kaapvaal Craton. It is possible that these intrusions concentrate near an

igneous palaeo-centre at the north-eastern edge of the Zimbabwe Craton, from which a

giant NE-trending swarm may have propagated laterally towards the south and west.

However, even if Mazowe dykes also are tholeiitic basalts, they are too depleted to have

been derived from an enriched mantle source, and are only geochemically similar to

northern KwaZulu-Natal’s NE-trending dykes in the opposite, south-easternmost part of the

Kaapvaal Craton, separated by more than 1000 km. In addition, Hanson et al. (2011) noted

paleomagnetic reconstructions consistent with a greater than 2000 km lateral displacement

being accommodated in the Limpopo orogenic belt that separates the Kaapvaal and

Zimbabwe cratons, further complicating the issue.

The more tholeiitic basalt character of most ca. 1.90 Ga and 2.65 Ga NE-trending

dykes (including lower LIL element and shallower, flat REE patterns) is distinctly different

from ca. 2.95 and ca. 2.65 Ga ENE- and SE-trending dykes, and is possibly related to a typical

continental rift setting, or perhaps even radiating from the suggested igneous centre near

the north-eastern edge of the Zimbabwe Craton. On the other hand, this extensive igneous

event also occurred during a complex tectonic period, including emplacement of the

Bushveld Complex and a ca. 2.0 Ga collision of the Zimbabwe and Kaapvaal cratons.

8.5. Palaeomagnetism

Our plot (see Fig. 73) of the Kaapvaal Craton palaeopoles in Table 10 illustrates the

considerable latitudinal drift undergone by the Kaapvaal Craton through the Mesoarchaean

and Neoarchaean. The palaeolatitude of the craton is shown to vary significantly from low

latitudes at ca. 2.95 Ga, to much higher latitudes at ca. 2.78 Ga and 2.65 Ga (see Fig. 73).


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Table: 10 - Summary of new and published palaeomagnetic data for the Kaapvaal Craton and adjacent orogens

used in this study

Pole Abbreviation Age (Ma) Age referenceLongitute

in °N

Latitude

in °Edp in °

dm in °

or

(A95 in °)

Paleomag reference

Nsuze basalts NB 2984 ± 3 Hegner et al ., 1993 67 105.6 5.3 9.2 Lubnina et al ., 2010

Badplaas

dyke swarmBAD 2967-2966 ± 1 Olsson et al ., 2010 63.6 105.4 2.3 4 Lubnina et al ., 2010

Hlagothi

comp. DD 2866 ± 2 this study 23.4 53.4 8.2 11.8 this study

NE-trending

dykesD ca. 2866 n/a n/a this study 43.9 55.2 4.2 7.7 this study

Agatha basalts AB 2977 ± 5 Nhleko, 2003 -9.4 333.0 (8.9) Strik, 2007

Usushwana

ComplexUC 2990-2860 n/a n/a Olsson et al ., 2011; Hunter and Reid, 1987 9.2 347 (7.6) Layer et al ., 1988; 1989

Hlagothi

comp. BB 2850-2750 n/a n/a this study -31.0 332.7 10.4 13.9 this study

Hlagothi

comp. AA 2850-2750 n/a n/a this study -24.4 338.1 (17.1) this study

Modipe

GabbroMG 2784 ± 3 Feinberg et al ., 2009 -32.8 30.9 (10.5) Evans and McElhinny, 1966

Gabarone

GraniteGG 2783 ± 2 Moore et al ., 1993 -35.0 284.0 Evans, 1967

Derdepoort

basaltsDB 2782 ± 5 Wingate, 1998 -39.6 4.7 (17.5) Wingate, 1998

ENE-trending

dykesF 2650 n/a n/a this study -76.2 16.7 8.59 24.2 this study

Allanridge

Formation

lavas

AFL 2709-2664 n/a n/a de Kock et al ., 2009 -67.6 355.8 (6.1) de Kock et al ., 2009

Rykoppies

dyke swarmRYD 2680-2668 n/a n/a Olsson et al ., 2010 -62.1 336 3.5 4.2 Lubnina et al ., 2010

Mbabane

PlutonMP 2687 ± 6 Layer et al., 1989 19.7 105.7 (9.7) Layer et al., 1989

Ongeluk lavas ONG 2222 ± 13 Evans et al ., 1997 -0.5 100.7 (5.3) Evans et al ., 1997

SE trending

dykes, south

of Badplaas

NSA ca . 2150 n/a n/a Lubnina et al ., 2010 5.9 93.4 12.0 20.4 Lubnina et al ., 2010

Gamagara/

Mapedi

Formation

BGM 2130 ± 92 Evans et al ., 2002 2.2 81.9 7.2 11.5 Evans et al ., 2002

Palaborwa

ComplexPB1 2060 ± 1 Heaman and LeCheminant, 1993 44.8 35.9 6.9 10.5 Morgan and Briden, 1981

Bushveld

ComplexBVC 2058 ± 2 Olsson et al ., 2010 19.2 30.8 (5.0) Letts et al ., 2009

Waterberg

Unconformity

bounded

sequence 1

WUBS1 2054 ± 4 Dorland et al ., 2006 51.3 36.5 (10.9) de Kock, 2007

Vredefort

impact

structure

VRED 2023 ± 4 Kamo et al ., 1996 40.7 22.3 11.6 15.7 Salminen et al. , 2009

Waterberg

Unconformity

bounded

sequence 2

WUBS2 ca . 1990 n/a n/a de Kock, 2007 -10.5 330.4 (9.8) de Kock, 2007

Hartley lavas HAR 1928 ± 4 Evans et al ., 2002 12.5 322.8 (16.0) Evans et al ., 2002

Post-

Waterberg

dyke swarm

PWD

1930-1870

n/a n/a Hanson et al ., 2004a 8.6 15.4 (17.3) Hanson et al ., 2004a

Black Hills

dyke swarmBH ca . 1900 n/a n/a Lubnina et al ., 2010 9.4 352 4.3 5.8 Lubnina et al ., 2010

Ntimbankulu

GraniteNG ca . 1050 n/a n/a Maré and Thomas, 1998 27.0 327.0 (18.0) Maré and Thomas, 1998

Namaqua-

NatalNN ca. 1050 n/a n/a Onstott et al., 1986 8.0 328.0 20 15

Hlagothi

comp. CC ca . 1050 n/a n/a this study 20.8 280.0 (19.6) this study

SE-trending

dolerite

dykes

C ca . 1050 n/a n/a this study 12.8 239.3 18.88 this study

Age uncertainty

(Ma)


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The significant change in the position of palaeopoles between 2.95, 2.87, 2.78 and 2.65 Ga

would argue against a single relatively stationary persistent mantle plume being responsible

for all the lavas, dykes and sills between the Nsuze, Hlagothi and Ventersdorp events. This

rather points toward short-lived transient plumes being operative in the Archaean. The

postulated ca. 2870 Ma Hlagothi plume would have strongly influenced the development of

the Pongola and Witwatersrand basins. It would have assisted in causing uplift and

subsidence forming several unconformities in the strata. These unconformities would have

been a result of uplift and erosion causing marine transgressions and regressions, with

cratonic flooding seen during maximum thermal subsidence of a plume.

Figure: 73 – Pole comparison between VGPs of components A, B, C (1, 2 and 3), D (1 and 2) and F is this study,

and VGPs and poles for the Archaean in (a), and the Proterozoic in (b) and (c). The various abbreviations in the

diagrams are shown in Table 10

The palaeolatitude of the Kaapvaal Craton changed significantly, moving from

northern high latitudes near the poles at ca. 2.95 Ga, to approximately 60° south at ca. 2.78

Ga and again to about 40° south at ca. 2.65 Ga, accounting for the newly established VGPs

using component D, A, B and finally F for a position during the time of the Ventersdorp

igneous event(s). Between ca. 2.65 Ga and ca. 2.22 Ga the Kaapvaal Craton changed

modestly in palaeolatitude (moving slightly northward) but rotated dramatically by about

140° clockwise, which may account for the lack of magmatic events seen during this time.

Between ca. 2.22 and ca. 2.15 Ga the position and orientation remained constant giving us


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more confidence that the data for both ca. 2.22 and 2.15 Ga events are reliable. By the time

of the ca. 2.05 Ga Bushveld Complex, the latitude of the craton remained constant and

there is a slight further clockwise rotation of the Kaapvaal Craton by about 40°. The

remaining key pole occurs at ca. 1.90 Ga by which time the palaeolatitude increased slightly

to the south, reaching about 45° south and there was still a further 80° clockwise rotation

(de Kock et al., 2006; Hanson et al., 2011). Of course there is a polarity ambiguity for each

value and therefore, it cannot be excluded that during some of these intervals the Kaapvaal

Craton was in the antipodal position (i.e., in the northern hemisphere and rotated by 180°).

As the APWP is developed for other cratonic blocks, this pattern for the Kaapvaal Craton can

be usefully compared in order to constrain Neoarchaean to Palaeoproterozoic

reconstructions involving the Kaapvaal Craton.

In addition, the ability to illustrate the presence of a magnetic remanence

component interpreted to be primary (i.e., component D) in the least altered lithologies of

the Hlagothi Complex is significant. This component is interpreted as primary, despite it

resembling Palaeoproterozoic palaeopoles from the Bushveld Complex, Waterberg Group

and Vredefort impact structure. The Hlagothi Complex is of Mesoarchaean age, and in a

location far removed from the thermal effects of the Bushveld Complex, the Vredefort

meteorite impact and LIPs within the sedimentary basins associated with the Waterberg and

Soutpansberg groups. Our VGP for this component is located in a position between the

Nsuze Group basalts and dykes obtained by Lubnina et al. (2010) and the post-Pongola

overprint palaeopoles, as well as the 2782 Ma Derdepoort basalt pole of Wingate (1998)

with which the ENE-trending dolerite dykes in the same area bear a similarity. The pole for

the Usushwana Complex is now considered discrepant, being cast into considerable doubt

given that it is unclear to which age group (i.e., 2990 Ma versus 2860 Ma) the pole belongs.

Also, the Usushwana Complex pole does bear remarkable similarity with the ca. 1.90 Ga

intrusions of the Palaeoproterozoic that are present in parts of the craton. Work will be

needed to be done to resolve this issue, which will assist in confirming the primary nature of

either the Nsuze or Hlagothi palaeopoles. The Hlagothi Complex VGP is further significantly

different from the poles for the ca. 2.65 NE-trending Rykoppies dyke swarm and other SE-

trending dykes (Lubnina et al., 2010), in addition to lavas of the Allanridge Formation (de

Kock et al., 2009).


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8.6. Correlations with the Pilbara Craton

The addition of the new ca. 2870 Ma LIP provides another intraplate event for the magmatic

barcode of the Kaapvaal Craton during the Archaean. This improved magmatic event

‘barcode’ can be compared to other similar magmatic event barcodes around the world (see

Fig. 74).

Figure: 74 – Correlations between the Kaapvaal and Pilbara cratons during the Meso- to Neoarchaean

reconstruction of Vaalbara

It has long been postulated that the Kaapvaal and Pilbara cratons were joined in the late

Archaean into a supercraton named ‘Vaalbara’ (e.g., Bleeker, 2003; Cheney, 1996; Cheney et

al., 1988; de Kock et al., 2009; Nelson et al., 1992; Wingate, 1998; Zegers et al., 1998). The

Vaalbara reconstruction is based on the correlation between magmatic pulses of the

Fortescue (Pilbara Craton) and Ventersdorp (Kaapvaal Craton) flood basalts at between 2.8

and 2.7 Ga; paleomagnetism also provides constraints (e.g., de Kock et al., 2009). If this

nearest neighbour relationship is correct, the Vaalbara reconstruction may also apply at ca.

2870 Ma (see Fig. 74). A potential correlative on the Pilbara Craton, the Millindinna Complex

yielded a poorly constrained Sm-Nd age of ca. 2860 Ma (Korsch and Gulson, 1986; Schmidt

and Embelton, 1985). However, more recent U-Pb age dating has obtained ages of 2925 and

3015 Ma for different Millindinna intrusions (e.g., van Kranendonk et al., 2007), suggesting

that the intrusions represent at least two distinct magmatic events of different ages (and

possibly a third if the ca. 2860 Ma Sm-Nd age was confirmed by U-Pb dating). Alternatively,


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we might consider the Zebra Hill dykes as possibly ca. 2870 Ma. The E-W trending Zebra Hill

dykes cut one of the ‘Millindinna’ suites of intrusions, the 2925 Ma Munni Munni intrusion,

and are unconformably overlain by the Hardey Sandstone at 2765 Ma (Barnes and Hoatson,

1994; Hoatson and Sun, 2002). If the ca. 2870 Ma Hlagothi event of the Kaapvaal Craton

were also confirmed to be present on the Pilbara Craton, this would further validate the

existence of Vaalbara back into the Mesoarchaean.

Figure: 75 – Palaeo-positions of the Pilbara Craton in relationship to the Kaapvaal Craton during the time of

the supercraton of Vaalbara, with two different configurations proposed based on palaeomagnetism and

geology: (A) the reconstruction of de Kock et al. (2009) and (B) the reconstruction of Zegers et al. (1998).

Chapter: 9 – Conclusion ___________________________________________________________________________

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Chapter: 9

Conclusion

The south-easternmost Kaapvaal Craton comprises a variety of inliers of Archaean basement

granitoid-greenstone terrane which is overlain by the supracrustal Mesoarchaean Pongola

Supergroup. These Precambrian rocks are exposed through the Phanerozoic cover of the

Karoo Supergroup sedimentary successions of the Dwyka and Ecca Group. To the south and

east, the craton is truncated by the Meso- to Neoproterozoic Namaqua-Natal mobile belt

which has metamorphosed and deformed the Archaean basement variably. However, it

usually increases as the cratonic margin is approached. The 50 km metamorphic boundary

marks the boundary between lower and upper greenschist facies metamorphism (Elworthy

et al., 2000). The Archaean basement is characterised by several intrusions that are absent

from the overlying Karoo strata. These intrusions comprise SE-, ENE- and NE-trending

dolerite dykes along with the Hlagothi Complex which intrudes the basement granitoid-

greenstones and overlying Pongola Supergroup strata. The SE- and NE-trending dolerite

dykes appear to consist of two different trends each. Trends of 030° and 050° dominate the

NE-trending dykes, and 135° and 160° trends dominate the SE-trending dolerite dykes. The

whole area including Phanerozoic strata has in turn been intruded by Jurassic sills and dykes

related to the Karoo LIP. The Precambrian inliers have been subjected to greenschist facies

metamorphism. The petrography indicates that the various primary mafic minerals of

olivine, ortho- and clinopyroxene have been pseudomorphed by talc, serpentine, chlorite

and amphibole at variable grades of greenschist facies metamorphism, while plagioclase

feldspar was sericitised. Pyroxene is uralitised. 40Ar/39Ar amphibole ages for the region also

illustrate this alteration from metamorphism, with most ages on the dykes indicating a

Namaqua-Natal overprint, with variable Ca/K ratios for the analysed amphiboles.

The dolerite dykes consist of several generations, with up to five having being

possibly recognised. SE-trending dykes represent the oldest dykes in the area, being cross-

cut by all the other dyke trends. These dykes probably consist of two generations with

similar basaltic to basaltic andesite and borderline tholeiitic to calc-alkaline geochemistry


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showing evidence of a geochemically enriched or contaminated magma being emplaced

into the craton. This is similar to other intrusions across the wider craton. These dykes have

been geochronologically, geochemically and paleomagnetically linked to either the ca. 2.95

or ca. 2.87 Ga magmatic events across the Kaapvaal Craton by Lubnina et al. (2010), a

conclusion that is supported by results in the present study.

The Hlagothi Complex, dated at 2866 ± 2 Ma as part of this dissertation comprises

a series of layered sills in northern KwaZulu-Natal. The sills consist of meta-peridotites,

pyroxenites and gabbros with at least two distinct pulses of magmatism. The distinct high

MgO (high Mg#) units are compositionally different from the older Dominion Group and

Nsuze Group volcanic rocks, as well as younger Ventersdorp volcanic rocks. This resurgence

of high MgO magmatism is similar to komatiitic lithologies in the Barberton Greenstone Belt,

and is indicative of a more primitive magma source, such as one derived from a mantle

plume. A mantle plume would account for the Hlagothi Complex, and the widespread

distribution of magmatic events of similar age and geochemistry across the craton. Potential

correlatives include the Thole Complex and gabbroic phases of the Usushwana Complex, the

2874 ± 2 Ma SE-trending dykes of northern KwaZulu-Natal, and flood basalts seen within the

upper Witwatersrand and Pongola supergroups such as the Crown, Bird, Tobolsk and Gabela

lavas. This large extent of intraplate magmatism allows us to propose a new large igneous

province for the Kaapvaal Craton during the Mesoarchaean that encompasses all of the

above mentioned geological units, and was generated by a short-lived transient mantle

plume, in one or two distinct pulses. This plume would also explain the development of

unconformities within the Mozaan Group through associated uplift and erosion, with

eruption of flood basalts coeval with the Hlagothi Complex during a mantle plume event,

and marine incursion and sediment deposition during subsidence from a thermal cooling.

This event also possibly re-magnetised some existing NE-trending dolerite dykes in the

vicinity, if indeed it is an overprint direction seen in the palaeomagnetic studies, and not a

primary direction.

Kaapvaal Craton during the Meso- to Neoarchaean formed an APWP from poles

established for the Nsuze event, to poles established for the Ventersdorp, incorporating the

new component for the Hlagothi Complex, and overprints seen in NE-trending dolerite

dykes. This new ca. 2870 Ma addition to the barcode of the Kaapvaal Craton allows new


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comparisons to be made to other possibly coeval units on cratons around the world. New

possible linkages include the Millindinna Complex, and the Zebra Hills dykes on the Pilbara

Craton. Precise dating and palaeomagnetism on these units is needed to confirm a link,

which if substantiated would assist in further validating the existence of Vaalbara during the

Mesoarchaean.

Following on from the Hlagothi Complex event, different pulses of magma

associated with the ca. 2.65 Ga Ventersdorp event occurred, with a NE-trending dolerite

dyke in the region dated at ca. 2.65 Ga, and a primary palaeomagnetic pole established by

Lubnina et al. (2010) in E-trending dolerite dykes across the region. These two directions

match the NE-trending and E-trending palaeostress fields seen in the Ventersdorp and

proto-Transvaal volcanic rocks, respectively. Both generations also demonstrate variable

geochemistry, with the NE-trending dykes having primary primitive tholeiitic magma, with

the only other known occurrence of this geochemical type being the assumed ca. 1.90 Ga

Mazowe dykes of the Zimbabwe Craton. The other generation of dolerite dykes from this

time, however, is calc-alkaline, being the ENE-trending dolerite dykes of the south-

easternmost region. In addition, some of the tholeiitic NE-trending dolerite dykes studied by

Lubnina et al. (2010) bear a similarity of magnetic components with NE-trending dykes

much further to the north in the Black Hills area, as well as with the Mazowe dolerite dykes

on the Zimbabwe Craton. It is clear from the complex array of dolerite dykes and intrusions

on the south-easternmost Kaapvaal Craton that much more work on the dolerite dykes

needs to be done in order to resolve these complex patterns in terms of their age,

geochemistry and palaeomagnetic components, which in turn will help resolve outstanding

issues with regard to their palaeo-tectonic framework within the larger craton.


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Chapter: 10 – Reference(s) ___________________________________________________________________________

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Chapter: 10

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Appendix: A – Sample Localities ___________________________________________________________________________

- 201 -

Appendix: A

Sample Localities

Locality Sample GPS co-ordinates

AG-D 28.55147°S, 31.15777°E

AG-E 28.55114°S, 31.15928°E

AG-F 28.55095°S, 31.15975°E

AG-G 28.55082°S, 31.15996°E

AG-H 28.55059°S, 31.15999°E

AG-I_core 28.45418°S, 30.93429°E

AG-I_contact 28.45418°S, 30.93429°E

HC-01 28.45418°S, 30.93429°E

HC-02 28.45683°S, 30.92145°E

HC-03 28.45796°S, 30.92024°E

HC-04 28.45837°S, 30.91893°E

HC-05 28.46264°S, 30.91957°E

HC-06 28.46320°S, 30.92132°E

HC-07 28.45813°S, 30.92923°E

HC-08 28.45804°S, 30.93224°E

HC-09 28.49097°S, 30.89525°E

HC-10 28.49181°S, 31.04262°E

AG-J 28.33221°S, 31.30280°E

AG-K 28.39073°S, 31.23518°E

DY-01 28.21946°S, 30.96515°E

AG-Bc 28.34447°S, 31.26238°E

AG-Cb 28.34353°S, 31.26481°E

AG-Ba 28.34447°S, 31.26238°E

AG-Bb 28.34447°S, 31.26238°E

AG-A 28.21037°S, 30.98243°E

AG-Ca 28.34353°S, 31.26481°E

DY-02_m 28.21253°S, 30.97588°E

DY-02_s 28.21253°S, 30.97588°E

DY-03 28.33222°S, 31.30280°E

AG-L 28.60811°S, 31.53068°E

AG-M 28.69957°S, 31.49188°E

Hlagothi Complex

SE-trending dolerite dyke

Other dolerite dykes

NE-trending dolerite dyke

ENE-trending dolerite dykes

Appendix: A – Sample Localities ___________________________________________________________________________

- 202 -

Appendix: B – Petrography ___________________________________________________________________________

- 203 -

Appendix: B

Petrography

The petrography and mineral identification on all 31 samples collected from the field was

investigated using optical microscopy, scanning electron microscopy (SEM) and X-ray

diffactrometry (XRD) in order to carry out a detailed mineralogical analysis.

Refected and transmitted light microscopy were performed on conventional

polished thin sections prepared at the Department of Geology, University of Johannesburg.

A Leica DMLP polarising research microscope equipped with an adapted Leica DC 200 digital

camera was used for optical microscopy and the acquisition of photomicrographs.

In addition SEM studies were conducted on the same set of polish thin sections upon

the completion of optical microscopy. Polished thin sections were carbon-coated and

microscopy was done on a Jeol 5600 SEM equipped with a Noran EDS detector, with a ultra

thin beryllium window at Spectrau, the central analytical facility at the University of

Johannesburg. The samples were examined with a 15 kV, 15 mA electron beam by means of

secondary and back scattered electron imaging, with mineral identification using semi-

quantiative EDS spot analyses.

XRD analyses was performed on all samples to assist petrographic studies in

identifying all major and minor mineral phases. Measurements were carried out using a

Pananalytical X’Pert diffractometer with a X’Celerator detector at Spectrau. Samples were

prepared as back-loaded powder pellets in aluminium metal sample holders. The analyses of

XRD patterns was done on X’Pert HighScore Plus software, with sample treatment involving

background removal and the stripping of Kɑ2 orbitals.

CIPW norms were also used from the whole rock geochemical analysis at Acme

Laboratories in Vancouver, Canada in order to assist in estimating the proportions of the

original pristine minerals before alteration.


- 204 -

Hlagothi Complex:

AG-I_core

AG-I_contact

HC-01

HC-02

HC-03

Mineral: Score:

actinolite 65

clinochlore 24

magnetite 39

talc 23

lizardite 29

Mineral: Score:

actinolite 58

clinochlore 44

talc 30

Mineral: Score:

actinolite 65

talc 22

clinochlore 35

magnetite 27

baddelyeite 25

Mineral: Score:

actinolite 64

talc 23

magnetite 31

baddelyeite 33

clinochlore 20


- 205 -

HC-04

HC-05

HC-06

HC-07

Mineral: Score:

actinolite 56

clinochlore 35

Mineral: Score:

actinolite 55

talc 36

clinochlore 31

Mineral: Score:

actinolite 65

clinochlore 25

lizardite 26

magnetite 35

talc 20

Mineral: Score:

actinolite 61

clinochlore 38

lizardite 22

magnetite 44

talc 26

Mineral: Score:

talc 40

magnetite 32

actinolite 25

clinochlore 27


- 206 -

HC-08

HC-09

HC-10

AG-D

AG-E

Mineral: Score:

quartz 55

albite 47

clinochlore 38

actinolite 44

epidote 22

Mineral: Score:

quartz 46

clinochlore 33

albite 52

actinolite 23

Mineral: Score:

albite 56

clinochlore 32

actinolite 35

Mineral: Score:

quartz 51

albite 48

clinochlore 34

actinolite 43

Mineral: Score:

clinochlore 40

talc 35

actinolite 30

antigorite 27


- 207 -

AG-F

AG-G

AG-H

SE-trending dolerite dykes:

AG-Bc

Mineral: Score:

albite 51

clinochlore 42

clinozoisite 31

actinolite 52

Mineral: Score:

actinolite 56

clinochlore 41

Mineral: Score:

albite 58

actinolite 52

clinochlore 33

Mineral: Score:

quartz 49

diopside 47

albite 43

actinolite 26

pigeonite 30


- 208 -

AG-J

AG-K

DY-01

ENE-trending dolerite dykes:

AG-Ba

Mineral: Score:

actinolite 57

clinochlore 35

albite 37

quartz 32

Mineral: Score:

actinolite 53

clinochlore 50

quartz 48

albite 44

Mineral: Score:

actinolite 53

albite 51

quartz 31

clinochlore 30

Mineral: Score:

albite 60

actinolite 56

clinochlore 39

microcline 26

quartz 42


- 209 -

AG-Bb

AG-Cb

NE-trending dolerite dykes:

AG-A

AG-Ca

Mineral: Score:

albite 50

anorthite 36

quartz 40

chlinochlore 40

Mineral: Score:

actinolite 57

albite 55

clinochlore 40

microcline 29

quartz 20

Mineral: Score:

actinolite 53

albite 42

clinochlore 41

magnetite 25

epidote 39

Mineral: Score:

quartz 53

actinolite 47

albite 46

clinochlore 40


- 210 -

DY-02m

DY-02s

SSE-trending dolerite dykes:

AG-L

AG-K

Mineral: Score:

quartz 56

albite 48

clinochlore 40

actinolite 40

Mineral: Score:

quartz 56

albite 48

clinochlore 40

actinolite 40

Mineral: Score:

quartz 56

hornblende 44

albite 34

ilmenite 25

Mineral: Score:

quartz 56

hornblende 44

albite 34

ilmenite 25


- 211 -

DY-03

Mineral: Score:

quartz 56

diopside 31

albite 44

biotite 33


- 212 -

Appendix: C – Whole-Rock Geochemistry ___________________________________________________________________________

- 213 -

Appendix: C

Whole-Rock Geochemistry

Hlagothi Complex:

HC-08 HC-09 AG-I_core AG-I_contact HC-01 HC-02 HC-03 HC-04 HC-05 HC-06 HC-07

SiO2 57.15 57.13 43.48 46.45 42.47 42.90 45.13 47.22 44.36 45.92 45.28

Al2O3 13.34 13.26 6.14 6.99 5.52 5.29 8.03 6.23 6.27 5.48 5.88

MnO 0.18 0.18 0.17 0.18 0.17 0.15 0.23 0.19 0.19 0.16 0.19

CaO 8.28 8.36 4.72 6.53 3.38 2.86 6.78 6.35 4.47 4.09 1.10

Na2O 2.55 2.40 0.25 0.25 0.21 0.21 0.22 0.24 0.22 0.23 0.21

K2O 0.87 0.86 0.21 0.22 0.29 0.26 0.22 0.21 0.33 0.32 0.23

Fe2O3T 11.69 11.61 11.50 9.70 12.37 11.89 11.73 10.77 11.83 12.19 13.21

MgO 4.16 4.05 24.86 22.06 26.21 26.65 20.67 22.02 24.13 24.00 24.47

TiO2 0.54 0.55 0.24 0.31 0.22 0.21 0.41 0.21 0.25 0.25 0.24

P2O5 0.07 0.09 0.05 0.05 0.05 0.04 0.07 0.04 0.06 0.05 0.05

Cr2O3 0.02 0.03 0.54 0.53 0.62 0.65 0.41 0.44 0.60 0.50 0.58

LOI 1.00 1.30 7.40 6.60 8.20 8.60 5.80 5.90 7.10 6.60 8.30

Total 99.85 99.82 99.56 99.87 99.71 99.71 99.70 99.82 99.81 99.79 99.74

Cs 0.40 0.30 3.00 1.50 1.70 1.30 1.10 1.10 2.80 2.00 1.50

Rb 22.10 22.50 15.10 1.10 9.00 7.10 1.30 1.00 13.60 11.70 2.60

Ba 214.00 206.00 7.00 3.00 4.00 4.00 3.00 3.00 3.00 3.00 3.00

Th 2.80 3.10 0.70 0.50 0.80 0.50 1.00 0.50 0.90 0.70 0.80

U 0.70 0.70 0.20 0.10 0.10 0.10 0.20 0.10 0.20 0.20 0.20

Nb 3.60 3.80 1.10 1.50 1.10 1.00 1.70 1.30 1.50 1.20 1.50

Ta 0.30 0.30 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10 0.10

La 13.50 13.60 4.1 5.9 4.4 3.1 4.9 3.5 3.3 4.6 3.3

Ce 27.60 27.50 8.6 11.8 8.7 6.9 10.2 7.0 7.5 9.6 6.7

Pb 0.90 1.20 1.2 1.3 1.4 0.8 1.4 1.3 1.1 0.8 0.8

Pr 3.15 3.26 1.05 1.54 1.00 0.64 1.27 0.82 0.89 1.16 0.82

Sr 118.20 118.90 31.1 21.9 12.9 10.3 9.7 17.9 19.5 22.6 22.5

Nd 12.20 12.30 3.90 5.90 4.40 3.10 5.80 3.40 3.30 4.80 3.80

Zr 79.70 82.90 27.90 36.60 24.80 24.30 45.90 24.60 31.10 29.30 27.90

Hf 2.10 2.40 0.80 1.00 0.50 0.80 1.20 0.70 0.70 0.80 0.80

Sm 2.44 2.59 1.01 1.38 0.86 0.82 1.42 0.86 0.88 1.08 0.86

Eu 0.68 0.71 0.28 0.35 0.21 0.21 0.36 0.18 0.23 0.26 0.17

Gd 2.65 2.80 1.10 1.54 0.98 0.96 1.74 0.91 1.09 1.12 1.02

Tb 0.50 0.53 0.18 0.26 0.15 0.17 0.30 0.15 0.18 0.17 0.16

Dy 3.12 3.29 1.23 1.70 1.11 1.14 2.03 1.06 1.28 1.18 1.10

Ho 0.72 0.77 0.26 0.35 0.22 0.25 0.39 0.19 0.24 0.24 0.24

Er 2.18 2.28 0.82 1.02 0.69 0.68 1.22 0.58 0.80 0.79 0.76

Tm 0.35 0.35 0.10 0.14 0.09 0.10 0.16 0.08 0.10 0.09 0.09

Yb 2.15 2.20 0.73 0.90 0.69 0.67 1.18 0.60 0.74 0.67 0.62

Y 19.70 20.00 7.00 9.80 6.00 6.50 11.00 5.20 6.80 6.40 6.00

Lu 0.35 0.37 0.10 0.14 0.09 0.10 0.18 0.08 0.11 0.10 0.10

V 262.00 269.00 108.00 121.00 98.00 96.00 155.00 119.00 111.00 119.00 109.00

Ni 30.9 30.9 1034.0 639.8 1262.0 1187.0 535.4 468.6 990.5 658.8 556.3

Co 50.4 49.5 112.5 85.3 116.4 114.4 91.0 96.9 113.6 107.7 116.3

Cu 86.3 94.8 1.0 1.9 1.0 0.8 1.3 1.0 3.2 11.2 2.6

Mo 5.8 7.9 1.3 1.0 0.8 0.8 2.9 0.6 0.8 0.7 0.3

Zn 51 46 14 14 27 22 66 27 24 19 18

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Dolerite Dykes:

AG-J AG-Bc AG-K DY-01 AG-Ba AG-Bb AG-Cb AG-A DY-02m DY-02s AG-Ca

SiO2 49.05 49.15 53.38 53.87 55.52 50.88 55.61 48.37 49.34 49.44 51.63

Al2O3 13.77 14.36 13.80 10.92 14.92 14.96 14.90 14.92 14.76 14.27 12.41

MnO 0.15 0.17 0.18 0.16 0.14 0.16 0.14 0.19 0.2 0.2 0.13

CaO 9.06 11.08 9.76 6.4 6.58 9.97 6.84 9.52 8.32 9.66 7.44

Na2O 2.12 2.00 1.96 2.66 4.41 2.10 4.24 2.29 2.45 2.38 2.64

K2O 0.82 0.60 0.86 1.02 0.96 0.78 1.45 0.65 1.08 0.96 1.41

Fe2O3T 10.43 9.39 12.03 11.73 9.11 10.73 9.17 13.11 10.89 10.56 9.33

MgO 11.27 8.98 6.25 7.63 5.83 8.16 5.48 6.95 7.68 7.06 7.96

TiO2 0.78 0.43 1.16 1.00 0.51 0.94 0.52 1.27 0.86 0.86 1.23

P2O5 0.11 0.07 0.12 0.18 0.10 0.12 0.11 0.13 0.66 0.70 1.02

Cr2O3 0.02 0.03 0.04 0.05 0.02 0.07 0.02 0.04 0.03 0.04 0.08

LOI 2.00 2.30 1.20 2.80 1.70 0.90 1.30 2.30 3.10 3.20 3.20

Total 99.58 98.56 100.74 98.42 99.80 99.77 99.78 99.74 99.37 99.33 98.48

Cs 0.40 0.30 1.30 6.20 1 1.9 1.9 1.1 1.1 1.2 11.4

Rb 32.8 19.7 39.6 72.1 42.5 26.6 91.2 87 52.3 50.3 149.7

Ba 159 133 203 379 272 212 282 119 2385 2565 1807

Th 1.4 1.6 2.9 3.8 2.2 2.6 2 0.3 6.5 6 11.2

U 0.2 0.3 0.8 0.9 0.3 0.3 0.3 0.1 1 1.1 1.8

Nb 3.7 2.6 10.8 11.9 4.2 3.9 4.4 3.5 8.3 8.4 19.5

Ta 0.3 0.2 0.7 0.8 0.2 0.2 0.2 0.2 0.4 0.4 0.9

La 9.5 8.9 18.7 29.5 15.8 10.4 14.7 5.2 92.1 97.6 133.4

Ce 19.7 18.1 39.8 63.2 30.9 23.4 29.8 13.1 185.1 191.1 264.5

Pb 1.4 1.3 1.0 3.1 12.1 11.5 10 41.3 8.2 3.9 4.7

Pr 2.52 2.20 5.02 8.04 3.62 3.15 3.62 2.02 22.13 23.11 30.81

Sr 233.4 195.0 193.0 698.7 401.3 268.7 398.4 174.2 773.6 828.5 1141

Nd 11.3 9.0 22.1 31.3 14 13.5 14.1 9.8 80.5 84.9 112.8

Zr 72.7 58.3 142.5 161.0 95.1 98.2 96.5 78.1 150 144 358.8

Hf 2.1 2.4 3.4 3.8 2.3 2.8 2.3 2.1 3.3 3.0 7.8

Sm 2.53 2.12 4.76 5.98 2.78 3.48 2.8 2.93 12.08 12.71 17.04

Eu 0.89 0.74 1.41 1.65 1.06 1.01 0.94 1.13 3.23 3.35 4.29

Gd 2.91 2.35 5.19 4.89 2.84 3.93 2.88 3.66 8.01 8.42 11.25

Tb 0.47 0.43 0.80 0.74 0.47 0.70 0.48 0.67 1.02 1.07 1.43

Dy 3.10 2.63 4.90 3.83 2.73 4.19 2.77 4.22 4.9 5.06 6.71

Ho 0.65 0.56 1.00 0.72 0.58 0.92 0.59 0.96 0.88 0.87 1.20

Er 1.90 1.63 2.95 1.81 1.57 2.62 1.62 2.74 2.25 2.19 3.21

Yb 1.72 1.44 2.67 2.01 1.44 2.36 1.47 2.51 1.9 1.92 2.59

Y 17.1 14.8 27.1 18.7 16 24 15.9 25.1 25 24.4 34.7

Lu 0.26 0.22 0.42 0.28 0.23 0.37 0.23 0.4 0.29 0.29 0.4

Tm 0.27 0.23 0.41 0.27 0.24 0.38 0.24 0.41 0.32 0.32 0.45

V 172 209 238 135 187 172 185 303 189 192 141

Ni 181.9 32.8 39.0 108.0 14.5 40.8 12.9 65.9 117.9 107.1 158.1

Co 62.3 47.1 43.6 59.6 53.7 48.0 51.2 57.2 47.0 51.9 38.5

Cu 72.2 34.5 59.4 212.8 36.4 89.7 30.6 71.3 36.7 35.2 11.2

Mo 2.6 4.3 4.2 2.6 3.1 3.0 3.8 3.0 2.7 3.9 3.7

Zn 38 15 45 27 45 29 31 82 85 79 82

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Miscellaneous:

DY-03 AG-L AG-M AG-D AG-E AG-F AG-G AG-H HC-10

SiO2 51.46 50.26 49.77 53.39 49.43 50.87 49.94 55.21 50.98

Al2O3 12.71 13.92 13.42 14.27 1.88 15.04 5.37 13.76 13.97

MnO 0.21 0.18 0.24 0.20 0.12 0.16 0.15 0.15 0.21

CaO 8.29 9.39 9.60 8.39 0.49 8.95 9.38 8.08 6.44

Na2O 2.45 2.11 2.02 3.83 0.01 2.85 0.18 5.16 4.16

K2O 1.41 0.63 0.76 0.32 0.01 1.13 0.02 0.17 1.37

Fe2O3T 16.09 11.41 14.76 11.84 7.00 9.75 9.66 9.56 15.87

MgO 4.07 8.95 5.77 5.39 30.81 8.36 19.88 5.70 1.96

TiO2 2.34 0.98 1.36 1.12 0.10 0.49 0.20 1.13 2.13

P2O5 0.36 0.15 0.13 0.13 0.02 0.05 0.02 0.11 0.91

Cr2O3 0.02 0.10 0.02 0.02 0.70 0.04 0.47 0.01 0.01

LOI 0.30 1.60 1.90 0.90 8.80 2.10 4.30 0.80 1.70

Total 99.74 99.71 99.75 99.83 99.51 99.77 99.65 99.84 99.68

Cs 5.5 12.0 0.6 0.2 0.1 0.3 0.1 0.1 16.8

Rb 68.3 55.8 12.4 3.3 0.3 19.9 0.3 1.1 72.2

Ba 372 285 94 75 5 441 4 45 775

Th 7.3 0.6 2.2 1.4 0.1 0.5 0.3 1.6 10.0

U 2.3 0.1 0.5 0.4 0.1 0.1 0.1 0.1 2.2

Nb 13.8 2.7 6.7 5.0 0.4 1.6 0.9 6.8 25.4

Ta 1.0 0.2 0.4 0.4 0.1 0.1 0.1 0.4 1.7

La 30.4 7.3 10.9 7.9 0.7 3.3 1.4 6.1 61.3

Ce 65.2 16.8 24.0 18.9 1.2 8.4 4.0 15.7 134.4

Pb 4.0 1.7 2.7 0.4 0.6 0.3 0.2 0.3 6.9

Pr 8.44 2.37 3.17 2.65 0.26 1.17 0.56 2.37 17.71

Sr 150.0 264.9 173.7 137.1 1.2 138.8 4.6 139.6 593.1

Nd 35.6 11.7 14.5 12.2 1.1 5.5 2.6 12.0 74.0

Zr 233.9 76.5 95.0 90.5 6.2 31.8 18.4 93.8 481.6

Hf 6.2 1.9 2.7 2.7 0.2 1.0 0.5 2.6 10.9

Sm 8.03 3.14 3.88 3.33 0.27 1.48 0.70 3.73 15.84

Eu 2.02 1.07 1.24 1.08 0.03 0.59 0.23 0.92 6.13

Gd 8.67 3.47 4.39 4.25 0.39 2.01 0.88 4.89 14.44

Tb 1.54 0.57 0.74 0.79 0.07 0.39 0.17 0.92 2.22

Dy 9.20 3.85 4.79 4.98 0.47 2.39 1.02 5.86 11.94

Ho 2.07 0.75 1.05 1.11 0.10 0.53 0.24 1.26 2.21

Er 5.87 2.11 3.02 3.25 0.28 1.65 0.71 3.78 6.00

Yb 5.36 1.77 2.81 3.08 0.29 1.44 0.66 3.46 4.97

Y 52.6 18.8 27.0 29.3 2.8 14.2 6.4 35.7 58.1

Lu 0.86 0.27 0.44 0.48 0.05 0.23 0.11 0.54 0.74

Tm 0.87 0.28 0.44 0.5 0.04 0.23 0.11 0.56 0.83

V 399 268 339 369 58 247 117 311 66

Ni 15.7 91.1 28.6 21.0 583.7 29.9 298.2 8.4 5.2

Co 47.3 49.8 46.2 43.7 82.4 46.5 94.6 37.0 31.0

Cu 113.2 103.8 180.8 1.7 4.4 48.2 7.5 2.4 21.2

Mo 4.9 3.7 3.8 3.8 0.6 2.5 0.9 3.4 4.9

Zn 58 42 42 18 12 12 14 5 109

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m)

Documents

dissertation - UJ IR - University of Johannesburg