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Chapter 5.8 Tectono-Metamorphic Controls on Archean Gold Mineralization in the Barberton Greenstone Belt, South Africa: An Example from the New Consort Gold Mine

Chapter 5.8 Tectono-Metamorphic Controls on Archean Gold Mineralization in the Barberton Greenstone Belt, South Africa: An Example from the New Consort Gold Mine

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Chapter 5.8: Tectono-Metamorphic Controls on Archean Gold Mineralization



Fig. 5.8-1. Geological map of the Barberton Greenstone Belt, showing the extension of the northern

and southern terranes, and the location of the Saddleback–Inyoka Fault System (modified after de

Ronde and de Wit (1994)). The occurrence of major gold deposits is restricted to the northern terrane.



The hypozonal New Consort gold mine in the Palaeo- to Mesoarchean Barberton Greenstone Belt, South Africa, forms one of four currently active mines that are clustered along

the northern margin of the greenstone belt. The mine is generally interpreted as having

formed during the late tectonic evolution of the belt in the Mesoarchean between ca. 3.2

and 3.1 Ga (de Ronde et al., 1991b). As such, it represents one of the oldest known orogenic

gold deposits on Earth (Groves et al., 1998; Goldfarb et al., 2001), although older deposits

have been documented for the Pilbara Craton (Zegers et al., 2002). The timing of mineralization relative to the metamorphic evolution of the greenstones is a significant problem

that relates to this and other Archean gold deposits, as this relationship has profound implications for the fluid source and tectonic setting in which mineralization occurred. In

this paper, we review the tectonic and metamorphic setting of this hypozonal gold deposit

within the overall tectonic framework of the greenstone belt, in an attempt to characterize the possible tectonic scenarios that were responsible for the formation of these gold

deposits.



5.8-2. Geological Setting



701



5.8-2. GEOLOGICAL SETTING

The Barberton Greenstone Belt, South Africa, forms part of the oldest nucleus of the

Kaapvaal Craton (e.g., de Wit et al., 1992). It contains a well-preserved ca. 3570–3220 Ma

volcano-sedimentary sequence, which is surrounded by various generations of tonalitetrondhjemite-granodiorite (TTG) gneiss domes and sheet-like potassic granites emplaced

between ca. 3500–3100 Ma (Fig. 5.8-1: e.g., Armstrong et al., 1990; Kamo and Davis,

1994; de Ronde and de Wit, 1994). The regional structural framework of the belt is dominated by large, upright synforms separated either by thrust faults, or narrow anticlines.

A characteristic feature of the belt is that the metamorphic grade is generally low, but increases towards the sheared contacts with surrounding TTG gneisses. This contact has been

interpreted as an extensional detachment that separates the greenschist facies greenstone

belt from mid-crustal basement gneisses (Kisters et al., 2003; Diener et al., 2005).

The greenstone belt sequence, assigned to the Swaziland Supergroup, has been subdivided into three stratigraphic units. From base to top, these include: (1) the Onverwacht

Group, dominated by ultramafic and mafic volcanic rocks; (2) the Fig Tree Group, a metaturbiditic succession made up of greywackes, shales, and cherts; and (3) the Moodies

Group, characterized by coarse-grained clastic sedimentary rocks, mainly including sandstones and conglomerates (Fig. 5.8-1: SACS, 1980). Major differences in age relationships,

depositional environments and sediment provenances between rocks to the north and south

of the Saddleback-Inyoka Fault System in the centre of the belt (Fig. 5.8-1), suggests that a

major suture zone separates the greenstone belt into a northern and southern terrane (e.g.,

de Ronde and de Wit, 1994; Kamo and Davis, 1994; Lowe and Byerly, 1999, this volume;

Lowe, 1994, 1999b). Consequently, the stratigraphy of the two terranes will be described

separately.

5.8-2.1. Southern Terrane

Onverwacht Group rocks to the south of the Saddleback–Inyoka Fault range between ca.

3550–3300 Ma. They mainly include mafic to ultramafic volcanic rocks, intercalated felsic

volcaniclastic sequences, and minor units of clastic and chemical sedimentary rocks (e.g.,

Viljoen and Viljoen, 1969a; de Wit et al., 1987b; Kamo and Davis, 1994; Byerly et al.,

1996; Dziggel et al., 2006a). At the base, they are intruded by an early generation of TTG

granitoid plutons, most of which have been dated at ca. 3445 Ma (e.g., Armstrong et al.,

1990). The overlying Southern Facies of the Fig Tree Group comprises shales, sandstones,

conglomerates, dacitic to rhyodacitic volcanic and volcaniclastic rocks, and minor chemical sedimentary rocks (Lowe and Byerly, 1999). U-Pb zircon dating of felsic volcaniclastic

rocks reveal ages of ca. 3260–3230 Ma (e.g., Kröner et al., 1991a; Byerly et al., 1996).

Deposition took place in alluvial, fan-delta, and shallow to moderately deep subaqueous

environments (e.g., Lowe and Nocita, 1999). The coarse-clastic sedimentary rocks of the

uppermost Moodies Group are dominated by quartz-rich sandstones that were probably deposited in several isolated basins (e.g., Heubeck and Lowe, 1999). The age of the Moodies

Group is poorly constrained, however, a minimum age of 3216 +2/−1 Ma (U-Pb zircon



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Chapter 5.8: Tectono-Metamorphic Controls on Archean Gold Mineralization



dating; Kamo and Davis, 1994) is suggested by the emplacement age of the post-kinematic

Dalmein Pluton that crosscuts the Kromberg syncline (Fig. 5.8-1).

5.8-2.2. Northern Terrane

Onverwacht Group rocks to the north of the Saddleback–Inyoka Fault have been assigned

to the Weltevreden Formation (e.g., Lowe and Byerly, 1999). In addition to the predominantly mafic and ultramafic volcanic rocks and minor chemical sedimentary rocks, this formation is characterized by abundant layered ultramafic intrusive complexes (Anhaeusser,

2001). A Nd isochron age of 3286 ± 29 Ma on komatiites suggests that the deposition

of this unit might correlate time wise with the uppermost Onverwacht Group rocks of the

Southern Facies (Lahaye et al., 1995). The base of the Onverwacht Group in this part of

the greenstone belt has been intruded by TTG granitoid plutons that yield U-Pb zircon

ages of ca. 3230 Ma (Fig. 5.8-1: e.g., Kamo and Davis, 1994). Compared to the ‘Southern

Facies’, the ‘Northern Facies’ of the Fig Tree Group has a more turbidite-like character,

and consists of carbonaceous shales, ferruginous cherts, greywackes, dacitic volcaniclastic

rocks, turbiditic sandstones, and minor conglomerates. U-Pb zircon data of 3226 ± 6 Ma

and 3225 ± 6 Ma indicate deposition largely coeval with the ‘Southern Facies’ (Kamo and

Davis, 1994; Lowe and Byerly, 1999). The age of the Moodies Group in the Northern Terrane remains uncertain, although Layer (1986) presented evidence that it was overprinted

during emplacement of the Kaap Valley Pluton at ca. 3214 ± 4 Ma.

5.8-2.3. Tectonic Evolution and Timing of Gold Mineralization

The greenstone belt and surrounding granitoid–gneiss terrain record a long, complex,

and polyphase tectono-magmatic history (e.g., de Ronde and de Wit, 1994; Lowe, 1994,

1999b). The episodic tectonic evolution involved at least three periods of deformation.

The first significant tectonic event (D1 ) took place at ca. 3445–3416 Ma, and was restricted to Onverwacht Group rocks in the southern part of the greenstone belt (e.g., de

Ronde and de Wit, 1994). The D1 deformation was closely associated with the intrusion of

an early generation of TTG granitoid plutons along the southern margin of the greenstone

belt (Fig. 5.8-1; e.g., de Wit et al., 1987a, 1987b; de Wit, 2004).

The D2 deformation occurred during the late Palaeoarchean (ca. 3227–3230 Ma) and

has been ascribed to a short-lived accretionary episode that affected the entire greenstone belt (de Ronde and de Wit, 1994; Kamo and Davis, 1994). The D2 deformation

was associated with the amalgamation of the southern and northern terranes along the

Saddleback–Inyoka Fault system and resulted in substantial crustal thickening (e.g., de

Ronde and de Wit, 1994; Kamo and Davis, 1994; de Ronde and Kamo, 2000). At this time,

early TTG granitoid gneisses and the associated basal sequences of the Onverwacht Group

along the southern margin of the greenstone belt, namely the Sandspruit and Theespruit

Formations, were buried to mid- to lower-crustal depth, where they experienced highpressure amphibolite facies metamorphism (Dziggel et al., 2002, 2005; Diener et al., 2005,

2006; Moyen et al., 2006). P-T conditions of up to 650 ◦ C and 15 kbar indicate very low



5.8-2. Geological Setting



703



apparent geothermal gradients of ca. 12 ◦ C/km, the lowest apparent geothermal gradient

ever been recorded in the Archean rock record (Moyen et al., 2006). Deposition of the Fig

Tree Group, as well as the syn-deformational molasse-type sedimentation of the Moodies

Group most probably correlate with this event (e.g., de Ronde and de Wit, 1994; Kamo

and Davis, 1994; de Ronde and Kamo, 2000). The D2 deformation is interpreted as a

result of crustal convergence and the accretion of the southern terrane during northwardsdirected subduction (Diener et al., 2006; Moyen et al., 2006). This interpretation is also

supported by the presence of syn (D2 )-tectonic TTG granitoid plutons along the northern

margin of the greenstone belt, such as the Nelshoogte and Kaap Valley Plutons (Fig. 5.81), which have been dated at 3236 ± 1 Ma and 3227 ± 1 Ma (U-Pb zircon; Kamo and

Davis, 1994; de Ronde and Kamo, 2000). Further, high-grade basement gneisses and associated supracrustal rocks along the northern margin of the greenstone belt record distinctly

higher geothermal gradients ( 30 ◦ C/km) than those in the south. This is consistent with

advective heating of upper plate rocks during TTG emplacement (Dziggel et al., 2006b).

D2 accretion was either synchronous with, or immediately followed by, a period of synorogenic extension and solid-state doming that eventually resulted in the steepening of

fabrics during the orogenic collapse of the belt (e.g., Kisters et al., 2003; Dziggel et al.,

2006b).

Currently available data on the timing of gold mineralization suggest that the gold deposits in the Barberton Greenstone Belt formed more than 100 million years after the main

accretionary D2 event (e.g., de Ronde et al., 1991b). Consequently, most workers regard

the late-tectonic evolution of the greenstone belt to be of particular importance for the

gold mineralization (e.g., de Ronde et al., 1991b, 1992; de Ronde and de Wit, 1994). During D3 , many of the steepened thrusts appear to have been reactivated as strike slip shear

zones, when most of the gold appears to have been precipitated (e.g., de Ronde et al., 1992;

Robertson et al., 1993). The D3 deformation was associated with the emplacement of voluminous sheet-like potassic granites to the north and south of the greenstone belt, such

as the ca. 3100 Ma Mpuluzi and Nelspruit batholiths (Fig. 5.8-1; e.g., Kamo and Davis,

1994; de Ronde and de Wit, 1994). Time constraints on the shear-zone-hosted gold mineralization are given by a ca. 3126 Ma porphyry dyke that predates shearing at the Fairview

Mine (Fig. 5.8-1), and a ca. 3084 Ma age for hydrothermal rutile associated with gold

mineralization (U-Pb dating; de Ronde et al., 1991b). This late-tectonic evolution, however, has remained speculative mainly because of the lack of robust geochronological data

constraining the age of strike-slip shearing within the greenstone belt, but also due to the

lack of knowledge about the influence of the D3 deformation on the final architecture of

the granite–greenstone terrain. Further complicating this picture is that the emplacement of

the late potassic granites seems to be related to a craton-wide period of intracratonic magmatism that was driven by the development of a crescent-shaped, juvenile arc along the

northern and western margins of the Kaapvaal Craton (e.g., Poujol et al., 2003). Thus, the

Mesoarchean magmatic and tectonic accretion at ca. 3.1 Ga appears to have been localized

some 200–300 km to the north and west of the Barberton Greenstone Belt, pointing to a

rather atypical intracratonic setting for the formation of the orogenic gold deposits.



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Chapter 5.8: Tectono-Metamorphic Controls on Archean Gold Mineralization



5.8-3. CHARACTERISTICS OF GREENSCHIST FACIES GOLD DEPOSITS

Although more than 350 gold deposits have been recorded in the Barberton Greenstone Belt, the bulk of the ca. 320 t of gold (more than 85%) has been produced from the

Sheba–Fairview, New Consort, and Agnes–Princeton mining complexes in the northern

part of the belt (Fig. 5.8-1; e.g., Anhaeusser, 1986; Ward, 1999). The majority of these and

associated smaller gold deposits are clustered in a crescent-shaped zone in greenschist facies supracrustal rocks along the eastern margin of the ca. 3230 Ma, tonalitic Kaap Valley

Pluton (Fig. 5.8-1). Gold mineralization in these deposits is structurally controlled by east–

west trending strike-slip shear zones that are situated in, or in close proximity to, major D2

structures that have been interpreted to have originated as thrust faults (e.g., de Ronde et

al., 1992). The gold deposits within these D3 -related, brittle–ductile shear zones show similar alteration characteristics with gold-bearing quartz–carbonate veins hosted in fuchsite-,

sericite-, and carbonate-rich alteration zones. Due to the close spatial association between

the gold deposits and the Kaap Valley Pluton, de Ronde et al. (1992) interpreted this tonalite

body as a major impermeable barrier for the mineralizing fluids. Fluid inclusion data show

that gold was deposited from a H2 O–CO2 –NaCl fluid due to phase separation at conditions

of ca. 300 ◦ C and 1 kbar, corresponding to a depth of ca. 3–4 km (de Ronde et al., 1992).

Based on the large apparent time gap between regional greenschist facies metamorphism

and gold mineralization (>100 million years), as well as the shallow crustal depth of the

greenstone belt (ca. 4–8 km; de Beer et al., 1988), the source of the mineralizing fluids was

interpreted to be external to the greenstone belt.

In the following sections, we review the characteristics of one of the better documented

greenschist facies gold deposits, the Sheba gold mine, in order to emphasize the structural

control and relative timing of gold mineralization in the greenschist facies gold deposits

of the Barberton Greenstone Belt. This is followed by a more detailed description of the

structural and metamorphic evolution of the New Consort gold mine.

5.8-3.1. Sheba Gold Mine

Gold mineralization at the Sheba gold mine is situated in the Sheba Hills, dominated by the

arcuate Eureka and Ulundi Synclines of post-Moodies Group age (Figs. 5.8-1 and 5.8-2;

e.g., Anhaeusser, 1986). These northerly-verging folds are separated by the Sheba Fault, a

regional-scale strike-slip shear zone. Arenaceous Moodies Group rocks dominate the Eureka Sycline to the north of the Sheba Fault, whereas the Ulundi Syncline to the south

of the Sheba Fault is made up of Onverwacht and Fig Tree Group rocks (Wagener and

Wiegand, 1986; Robertson et al., 1993). Host rocks to the mineralization in this part of

the greenstone belt are mainly Fig Tree and Moodies Group sedimentary rocks, including

quartzites, shales, sandstones, and greywackes. Within the Ulundi Syncline, Onverwacht

Group volcanic rocks (mainly talc–carbonate schists and cherts) outline km-scale, tight- to

isoclinal folds that are locally referred to as the Sheba Anticlines (Fig. 5.8-2). The Sheba

Fault has been interpreted as an originally low-angle thrust fault that was reactivated during



5.8-3. Characteristics of Greenschist Facies Gold Deposits

705



Fig. 5.8-2. Simplified geological map of the area around the Sheba and Fairview mines, including major structures (modified after Wagener

and Wiegand (1986)). Inset shows a rose diagram plot illustrating the trend of shear zones in the Main Reef Complex (Robertson et al., 1993),

as well as an interpretation of the data.



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Chapter 5.8: Tectono-Metamorphic Controls on Archean Gold Mineralization



post-Moodies deformation. During refolding and arcuation of the Eureka and Ulundi Synclines, the Sheba Fault was reactivated as a dextral strike-slip fault (e.g., Robertson et al.,

1993). Gold mineralization is predominantly situated in the Ulundi Syncline, where it is

spatially closely associated with the steeply southerly dipping contact between greywackes

and shales of the Fig Tree Group and the more competent cherts of the Onverwacht Group.

Most of the gold mineralization is hosted by shear zones that dip steeply to the SE

and SW (Robertson et al., 1993). Gold is mainly hosted in quartz–carbonate veins along

these fractures, or occurs in association with disseminated sulfides (mainly arsenopyrite

and pyrite) within the adjacent wall rocks. Based on detailed structural analyses of the Main

Reef Complex (MRC) in the westernmost part of the Sheba mine (Fig. 5.8-2), Robertson

et al. (1993) interpreted the geometry of fracture zones associated with the gold mineralization as second- and third-order structures formed during WNW–ESE crustal shortening.

They distinguished three sets of shear zones: (i) the NE-trending and SE-dipping main

shear zones, which are sub-parallel to bedding; (ii) the steep shear zones, which strike E–

W and dip steeply to the south, and (iii) the cross shear zones, a set of shallowly southerly

dipping shear zones that cross-cut the earlier structures. Zones of high-grade gold mineralization occur at the intersection of the main and cross shear zones, and plunge at low

to moderate angles to the NE. This fracture pattern is consistent with the D3 -related, dextral strike slip shearing along the Sheba Fault. In their model, Robertson et al. (1993)

interpreted the main shear zones to represent the principal displacement shears (Y-shears),

whereas the conjugate shear fractures would correspond to R and P shears (Fig. 5.8-2,

inset). The cross shear zones would represent high-angle antithetic (R ) shears. The geometry indicates a WNW–ESE directed maximum principal stress during shearing, consistent

with the proposed direction of regional crustal shortening during D3 (e.g., de Ronde and

de Wit, 1994; Belcher and Kisters, 2006a).



5.8-4. THE NEW CONSORT GOLD MINE

The New Consort gold mine is situated in the eastern part of the Jamestown schist belt

(Figs. 5.8-1 and 5.8-3). In contrast to the gold deposits described above, mineralization

at New Consort is hosted by distinctly higher-grade metamorphic rocks. The complexly

folded and imbricated volcano-sedimentary sequence is situated in the immediate hangingwall of the basal granitoid–greenstone contact, which separates the generally greenschist

facies greenstone belt from the mid-crustal gneisses of the Stentor Pluton (Dziggel et al.,

2006b). This contact has been interpreted as an extensional detachment along which basement gneisses have been exhumed, most likely in the course of the orogenic collapse of

the belt at ca. 3230 Ma. A distinctive feature of the gold mineralization at the New Consort

gold mine is the development of laterally extensive mineralized horizons in wall-rocks of

significantly different metamorphic grade. The gold mineralization is mainly structurally

controlled by highly silicified ductile shear zones at, or near, the structural contact between

rocks of the Onverwacht and Fig Tree Groups. This contact, locally referred to as the Consort Bar, reaches a thickness of up to several meters, and epigenetic ore shoots may occur in



5.8-4. The New Consort Gold Mine

707



Fig. 5.8-3. Simplified geological map of the New Consort gold mine area in the eastern part of the Jamestown schist belt (modified after

Anhaeusser and Viljoen (1965); Anhaeusser (1972)). The Western, Central and Eastern Zones mark the areas investigated by Otto et al. (2007),

and are projections from underground. See text for explanation. Section A-B shown in Fig. 5.8-4.



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Chapter 5.8: Tectono-Metamorphic Controls on Archean Gold Mineralization



different stratigraphic levels at, or near, this structural break. Stratigraphically, the exposed

rocks comprise Onverwacht Group mafic and ultramafic volcanic rocks and intercalated

aluminous felsic schists and cherts that are overlain by argillaceous and coarse clastic sedimentary rocks of the Fig Tree and Moodies Groups (Fig. 5.8-3). The volcano-sedimentary

sequence has been folded into two major synclines that plunge at low to moderate angles

to the SE and E, namely the Top Section and 3 Shaft Synclines (Fig. 5.8-3). The structural and metamorphic setting of the main ore bodies, as well as the characteristics of ore

and alteration assemblages associated with the gold mineralization, are outlined below. It

should be noted that the sequence of deformation events described in the following refers

to the local structural evolution based on overprinting structural fabrics developed in the

mine workings and surroundings. The nomenclature does not correspond to the regional

tectonic events outlined in the chapters on regional geology (e.g. Lowe and Byerly, this

volume), although we try to establish a correlation of local and regional fabrics and deformation events.

5.8-4.1. Structural Evolution

The earliest fabric recorded along the granite–greenstone contact is a moderately southerly

dipping S1 foliation. S1 is parallel to the granite–greenstone contact and to bedding, attributed to the transposition of lithologies (Figs. 5.8-3 and 5.8-4). This D1 shearing affected

both the greenstone sequence, as well as the structurally underlying gneisses of the Stentor

Pluton. The orientation of fabrics and kinematic indicators contained within them point

to the fact that D1 -related fabrics are associated with the syn- to post-collisional extension of this mid- to lower-crustal segment (Dziggel et al., 2006b). In the area north of the

New Consort gold mine, two main structural domains can be distinguished (Fig. 5.8-4):

the early S1A fabrics are restricted to the amphibolite facies gneisses of the Stentor Pluton

and overlying supracrustal rocks. Associated mineral stretching lineations plunge at low

to moderate angles to the W and SW, parallel to the fold axes of open to tight F1A folds.

The SW plunge of L1A stretching lineations is also parallel to the general NE–SW trend of

constrictional fabrics in the southern parts of the greenstone belt, and thus is in agreement

with the postulated direction of orogen-parallel extension (Kisters et al., 2003). In contrast,

S1B fabrics are mainly confined to a 0.5–1 km wide high strain belt, where they overprint

the earlier fabrics. These fabrics are recorded in both amphibolite and retrograde greenschist facies mylonites (Fig. 5.8-4), pointing to a considerable increase in strain intensity

during retrogression. The associated mineral stretching lineations are either downdip, or

plunge at low to moderate angles to the SE, parallel to the fold axes of isoclinal intrafolial folds. Kinematic indicators consistently point to a Stentor Pluton-up/greenstone-down

sense of movement with a sinistral strike slip component. The S1B fabrics have been linked

to the last stages of exhumation that were characterized by solid-state doming of footwall gneisses in response to the increasing buoyancy contrast during extensional shearing

(Dziggel et al., 2006b). The folding of the shallowly plunging Top Section and 3 Shaft

Synclines has been linked to the D1C deformation (Fig. 5.8-4). Their formation most probably occurred in direct response to the emplacement and solid-state doming of hot, ductile



5.8-4. The New Consort Gold Mine



Fig. 5.8-4. Geological cross-section through the granite–greenstone contact, illustrating the structural relations and distribution of metamorphic

domains (section A-B in Fig. 5.8-3; modified after Dziggel et al. (2006b)). All stereographic projections are equal area plots and to the lower

hemisphere.



709



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Chapter 5.8: Tectono-Metamorphic Controls on Archean Gold Mineralization



basement rocks into shallower crustal levels, resulting in the folding of earlier fabrics and

the overlying stratigraphic units. Because this folding also affected the Consort Bar, much

of the shearing (but not the mineralization) along this shear zone must predate the regional

folding (Fig. 5.8-3).

The gold mineralization has been related to the local D2 deformation, which has been interpreted to correspond to the regional D3 tectonic event (Tomkinson and Lombard, 1990;

Harris et al., 1995; Otto et al., 2007). In general, the fabrics related to this deformation

event remain steep, but are variable in orientation and partitioned into discrete high-strain

zones (Otto et al., 2007). In the central parts of the mine, the local D2 deformation resulted

in the formation of a prominent, steeply dipping shear zone system known as the Shires

shear zone (Fig. 5.8-3). The Shires shear zone is made up of an anastomozing network of

discrete, N–S to NW–SE trending shear zones, and can be up to 350 m wide. Individual

shear zones vary from a few cm to 5 m in thickness (Harris et al., 1995). Earlier studies on

the kinematics along this shear zone proposed a west-block-up and minor sinistral sense

of movement that appears to displace the Consort Bar by 1200 m on surface (Tomkinson and Lombard, 1990; Harris et al., 1995). This suggests NW–SE directed subhorizontal

shortening during the D2 deformation. However, the occurrence of gold in shear zones

parallel to, or within, the Consort Bar along the northern, roughly E–W trending, limbs

of the Top Section and 3 Shaft Synclines indicates that D2 shearing was not restricted to

the central parts of the mine (Fig. 5.8-3). In these shear zones, the mylonitic S2 fabrics

are essentially coplanar with the regional S1 foliation, and thus, not easily distinguished.

The D2 deformation is also marked by the intrusion of numerous syn-kinematic pegmatite

dykes, in an orientation strongly controlled by D2 shear zones (Harris et al., 1995). These

pegmatites locally crosscut the mineralization, but may also be folded and commonly display mylonitic fabrics along their margins, pointing to a continuous deformation during

their emplacement. Rb-Sr and Sm-Nd garnet isotope analyses yield an age of ca. 3040 Ma

(Harris et al., 1993). The last deformation event recorded (D3 ) is characterized by the development of brittle normal faults, collectively referred to as the Blue Jackets fault system,

which crosscut all earlier structures (Fig. 5.8-3; Tomkinson and Lombard, 1990; Harris et

al., 1995).

Detailed structural underground mapping by Otto et al. (2007) revealed that the geometry of the three ore bodies currently mined at New Consort varies according to the exact

structural position within the mine. The location of the three areas investigated, projected

from underground, is shown in Fig. 5.8-3. The Western Zone ore body is located on the

shallowly southerly dipping, northern limb of the 3 Shaft Syncline, about 150 m below

surface. The mineralization of the Central Zone is located in the Consort Bar within the

Shires Shear Zone system and ca. 1400 m below surface. The Eastern Zone ore body is

located at the northern limb of the Top Section Syncline, ca. 1400 m below surface and

ca. 60 m in the footwall of the Consort Bar, and represents an example of the so-called

Footwall Lens mineralization (see Section 5.8-4.3).

In accordance with the structural framework outlined above, Otto et al. (2007) identified

3 phases of deformation (Fig. 5.8-5). In all three localities, the variably dipping S1 foliation strikes approximately east–west, and is interpreted to correlate with the S1B foliation



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Chapter 5.8 Tectono-Metamorphic Controls on Archean Gold Mineralization in the Barberton Greenstone Belt, South Africa: An Example from the New Consort Gold Mine

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