The Stratigraphy and Structure of the Western Saldania Belt, South Africa and Geodynamic Implications

Abstract

The western Saldania belt in South Africa exposes a section through a fore-arc region that records deformation, sedimentation and plutonism in the late Neoproterozoic and Cambrian (>560 to <510 Ma) along the obliquely convergent continental margin of the Kalahari Craton. The belt comprises two main and structurally overlying units. Imbricated and pervasively transposed marine metasediments and relics of oceanic crust constitute the structurally lower Swartland complex. Kinematics and strains indicate formation of the accretionary complex during tectonic underplating and top-to-the west and northwest thrusting related to the southeast-directed subduction of the Adamastor ocean below the Kalahari Craton. The Swartland complex is unconformably overlain by low-grade metamorphic metasediments and minor metavolcanic rocks of the Malmesbury group that represent the late-Neoproterozoic to Cambrian fore-arc basin fill. Sedimentary facies suggest the presence of a volcanic arc in the east, succeeded in the west by metaturbidites of the inner fore-arc basin and deeper-water metapelitic successions in the northwest interpreted to form slope-apron deposits overlying the toe of the prim and facing the ocean basin. The regional deformation of the fore arc is characterized by partitioned sinistral transpression related to oblique convergence. The absence of collisional structures and very limited exhumation of the rocks suggest a soft collisional event, probably as a result of slab break-off. This break-off may also account for the voluminous, syn- to late-tectonic plutonism of the Cape Granite Suite in the fore-arc region, peaking around 540–530 Ma and accompanying the waning stages of regional tectonism.

1 Introduction

The Saldania belt is an arcuate fold belt along the southern tip of Africa (Fig. 14.1) and part of the larger system of Pan-African belts that led to the assembly of southwest Gondwana in the late Neoproterozoic and early Palaeozoic (e.g., Hartnady et al. 1985; Gaucher et al. 2009). In southern Africa, the chronology and geological evolution of these events are well documented from the better exposed coastal Kaoko, Gariep and Vanrhynsdorp belts (e.g., Von Veh 1983, 1993; Gresse 1994; Frimmel and Frank 1998; Goscombe et al. 2003; Basei et al. 2005; Konopasek et al. 2005; Goscombe and Gray 2008; Miller 2008). Lithological sequences, metamorphic P–T paths and sinistral transpressive tectonics record the oblique and diachronous closure of oceanic basins culminating in the collision of South American cratons and intervening arc terranes with the Kalahari and Congo Cratons in Africa (Fig. 14.1). A similarly well-constrained depositional and kinematic framework has not been established for the southernmost Saldania belt. Poor outcrop conditions, the apparent monotony of supracrustal sequences and uniformly low grades of metamorphism have resulted in a number of contrasting views about the tectonostratigraphic make-up and overall geodynamic setting of the belt (Hartnady et al. 1974; Von Veh 1983; Rozendaal et al. 1999; Belcher and Kisters 2003; Gresse et al. 2006; Frimmel 2009; Frimmel et al. 2011, 2013; Buggisch et al. 2010; Rowe et al. 2010). This has prevented regional correlations with similarly old belts to the north so that the Saldania belt has only featured sporadically and commonly controversially discussed in paleogeographic reconstructions of southwest Gondwana (e.g., Gaucher et al. 2009; Miller et al. 2009; Frimmel et al. 2011; Rapela et al. 2011).

Fig. 14.1

Location of Neoproterozoic and Cambrian orogenic belts (orange) in the broader context of SW Gondwana in southern Africa and eastern South America (simplified after Rapela et al. 2011; Frimmel et al. 2013; Oriolo et al. 2016). Major strike-slip and thrust fault zones are shown and thin, solid lines indicate generalized structural formlines and trends

The Saldania belt (sensu lato) is made up of an east–west trending southern branch and a northerly trending western branch (Fig. 14.1) (Gresse et al. 2006). This contribution focuses on the western branch of the Saldania belt stretching from Cape Town in the south to north of the coastal town of Saldanha from which the belt derives its name (Fig. 14.2). This part documents sedimentation, plutonism and deformation in what probably represented the eastern realm of the Adamastor ocean or related oceanic basins along the western margin of the Kalahari Craton in pre-Gondwana times. Our understanding of the geology of the western Saldania belt has been fundamentally influenced by the benchmark paper by Hartnady et al. (1974), which proposed the belt to be underlain by three allochthonous or para-autochthonous domains or terranes separated by prominent and supposedly terrane bounding strike-slip fault zones (Fig. 14.2). This view also forms the basis for currently accepted lithostratigraphic subdivisions that discuss the stratigraphy of the belt in terms of distinct, fault-bounded tectonostratigraphic packages (e.g., South African Committee for Stratigraphy (SACS) 1980; Gresse et al. 2006; Table 14.1). Despite this, the term ‘Malmesbury Group’ has been retained for the supracrustal succession across purported terrane boundaries and the belt in its entirety. This is obviously not without problems and also acknowledges lithological and structural similarities of rocks in large parts of the belt. Moreover, the uniformly low grades of metamorphism, contiguity of structures without any major breaks and similar age of rocks are not indicative of the presence of distinct terranes. As a result, not only the extent and delineation of terranes (Frimmel 2009; Frimmel et al. 2011, 2013; Buggisch et al. 2010) but also the presence of distinct tectonostratigraphic domains (Belcher 2003; Belcher and Kisters 2003) have been questioned. This highlights our only rudimentary understanding of the structural geology and lithostratigraphic relationships and, as a result, the rather diverse interpretations of original depositional environments and geodynamic setting of the belt.

Fig. 14.2

Schematic geological map of the western Saldania belt showing the currently accepted subdivision of the belt into three tectonostratigraphic domains or terranes (Tygerberg, Swartland and Boland), separated by strike-slip faults and intruded by plutons of the Cape Granite Suite (modified after Rozendaal et al. 1999; Gresse et al. 2006). Note that the extent of and contacts between geological formations are largely inferred given the very poor outcrop conditions. Blank areas on the map correspond to regions with less than 1% outcrop. Locations mentioned in the text are shown

Table 14.1 Synopsis of recent stratigraphic subdivisions suggested for the western Saldania belt following Gresse et al. (2006) and Frimmel et al. (2013)

The primary purpose of this chapter is to review the structural and lithological inventory of the belt with a view to integrating some of the conflicting interpretations and unresolved controversies that surround the geology of the belt. For this we intend to highlight similarities and/or differences of lithological packages and regional strains, particularly across purported terrane boundaries, and against a background of more recent geochronological, petrographic and structural data (e.g., Villaros et al. 2009; Rowe et al. 2010; Farina et al. 2012; Frimmel et al. 2013), and our own work. It is not our aim to provide a new stratigraphic subdivision but rather to suggest a geological framework that acknowledges lithological affinities and reconciles distinct structural domains within the belt. We hope this will allow for a better integration of the western Saldania belt into the broader framework of Pan-African belts in southwest Gondwana.

2 The Main Geological Characteristics of the Western Saldania Belt

The western Saldania belt is a low-grade metamorphic fold belt underlain by supracrustal rocks collectively referred to as the Malmesbury Group. Aerially extensive syn-, late- and post-tectonic granites of the 550–510 Ma Cape Granite Suite (CGS) are intrusive into the Malmesbury Group (Fig. 14.2). The group comprises a late Neoproterozoic, predominantly clastic marine sedimentary succession. Mafic to felsic metavolcanic rocks are volumetrically minor and are either intrusive or structurally interleaved with the metasediments (Hartnady et al. 1974; Theron et al. 1992; Slabber 1995; Gresse et al. 2006). Following Hartnady et al. (1974), the Saldania belt has been subdivided into three tectonostratigraphic domains or terranes: a southwestern Tygerberg, a central Swartland and a northeastern Boland terrane (Fig. 14.2). The former two are separated by the prominent Colenso fault zone, whereas the Swartland is interpreted to be separated from the Boland terrane by the Piketberg-Wellington fault (Fig. 14.2). Detrital zircon ages (Armstrong et al. 1998; Frimmel et al. 2013) and zircon crystallization ages from intercalated tuffs (Kisters et al. 2015) indicate a depositional age of c. 560–555 Ma for at least the upper parts of the Malmesbury Group. The base of the group is not exposed anywhere and the thickness and upper age of the sequence as well as the nature of the basement remain elusive.

The second and characteristic component of the belt are voluminous S- and minor I- and A-type granites of the CGS. The granites crosscut regional folds and fabrics, and are largely devoid of magmatic or solid-state fabrics. This has commonly been suggested to indicate a late- to post-tectonic emplacement of the CGS (e.g., Scheepers 1995; Scheepers and Schoch 2006). In contrast, granites that have intruded into or along prominent fault zones are invariably deformed and gneissic. The most prominent example of this synmagmatic deformation is the large Darling batholith that has intruded the central Colenso fault zone over a strike length of more than 40 km (Fig. 14.2). The deformation of the granites illustrates not only the syntectonic timing of at least parts of the CGS but also the pronounced regional-scale partitioning of strain in the belt (Kisters et al. 2002). U-Pb zircon ages indicate emplacement of the granites between c. 550 and 510 Ma with a peak of plutonic activity between c. 540 and 530 Ma (Schoch 1975; Scheepers 1995; Da Silva et al. 2000; Scheepers and Schoch 2006; Villaros et al. 2009; Chemale et al. 2011; Farina et al. 2012). These intrusion ages also emphasize the only short timespan between the deposition, deformation and metamorphism of the Malmesbury Group and the onset of CGS plutonism. On a regional scale, earlier S-type granites are confined to the southwest of the Colenso Fault and the Tygerberg domain. I-type granites are only found east of the Colenso Fault and mainly in the Swartland domain (Fig. 14.2). The distribution of geochemically distinct granite suites on either side of the Colenso fault probably indicates the transcrustal extent of the fault zone, but has also been used to reiterate the terrane model for the belt (Rozendaal et al. 1999; Kisters et al. 2002; Frimmel et al. 2013).

The third but volumetrically minor lithological component of the belt are rocks of the Klipheuwel Group. The mainly clastic sedimentary rocks are somewhat transitional between the Pan-African basement of the Malmesbury Group and CGS, and the Paleozoic Cape Supergoup. Rocks of the Klipheuwel Group unconformably overlie folded strata of the Malmesbury Group and plutons of the CGS but are located below the regional unconformity with the overlying Cape Supergroup. For the most part, Klipheuwel Group rocks are confined to isolated, fault-bounded basins (Fig. 14.2). The rocks show steep dips, but it is not clear whether these strains are late-Pan-African or Permo-Triassic in age and part of the later Gondwanide evolution that has affected the southwest parts of the Gondwana supercontinent.

Structurally, the western Saldania belt is, for the most part, developed as a fold belt comprising northerly- to northwesterly-trending, gently doubly plunging, open to isoclinal and more or less upright to southwest-verging folds (F2) (Rabie 1948; Theron et al. 1992; Gresse et al. 2006; Rowe et al. 2010). Regional-scale thrusts have not been mapped, although this may be a function of the subdued topography and very limited outcrop. Domains of more complex low-angle fabrics (D1 fabric domains) and intensely disrupted stratigraphy are preserved in structural windows in the central parts of the belt, in the Swartland domain (Figs. 14.2 and 14.3) (Rabie 1948; Newton 1966; Hartnady et al. 1974; Belcher 2003; Belcher and Kisters 2003). Owing to the very poor outcrop conditions, the significance of these high-strain domains, their extent and also their contact relationship with the structurally overlying, relatively simple fold belt have not been resolved. The most prominent structural features of the belt are the northwesterly-trending subvertical Colenso and the less prominent Piketberg-Wellington fault zones (Rabie 1948, 1974; Schoch 1975; Theron et al. 1992; Kisters et al. 2002). Sinistral strike-slip kinematics are widely documented from the Colenso fault (Kisters et al. 2002) and an oblique slip, east-side up component is suggested by the deeper structural levels exposed on the eastern side of the fault (e.g., Hartnady et al. 1974; Rozendaal et al. 1999; Gresse et al. 2006). The location of the Piketberg-Wellington is largely inferred over much of its strike extent (Rabie 1974; Theron et al. 1992; Slabber 1995). Demonstrable displacement along the fault strands can only be shown to have affected the Paleozoic Cape Supergroup with a downthrow to the east (e.g., Belcher 2003).

Fig. 14.3

Geological map of the western Saldania belt, as in Fig. 14.2 but highlighting the distribution of structural and lithological packages extending across purported terrane boundaries, emphasizing the presence of two structurally overlying packages, a lower Swartland complex and an upper Malmesbury group. Note that these terms are not officially approved by SACS. Details of this subdivision, and lithological and structural affinities within these two overlying domains, are discussed in the text

Recent research has mainly focused on the delineation of actual tectonostratigraphic domains or terranes, and the original terrane model for the belt has undergone several permutations. Frimmel (2009) and Frimmel et al. (2011) suggest only the southwestern Tygerberg to represent a truly allochthonous terrane, separated from the autochthonous Swartland and Boland terranes by the Colenso fault zone. Buggisch et al. (2010) place the terrane boundary to the northeast and along the Piketberg-Wellington fault. The study by Buggisch et al. (2010) postulates the amalgamation of the Tygerberg and Swartland terranes against the Boland terrane during westward subduction and closure of an ocean basin. Supposed relics of this ocean basin are preserved as the metamafic Bridgetown Formation (Fig. 14.2; Table 14.1) between the Boland and Swartland terranes. In this scenario, the Swartland and Tygerberg terranes are suggested to have been part of South American cratons, whereas the Boland terrane represented the passive margin of the Kalahari Craton. Frimmel et al. (2013) report results of a detrital zircon study sampling different formations throughout large parts of the belt. The results indicate deposition of the rocks in the late Neoproterozoic, up to c. 560 Ma. However, zircon age spectra in the Boland domain show a far more prominent population of Mesoproterozoic zircons compared with the Tygerberg and Swartland domains that are dominated by Ediacaran-age zircons. Based on this, Frimmel et al. (2013) suggest a largely contemporaneous sedimentation of the rocks in the latest Neoproterozoic, but in two distinct basins and the two southwestern domains to represent part of a contiguous block, termed the Malmesbury terrane. In this scenario, the Piketberg-Wellingon fault is advocated as the main terrane boundary, separating the Malmesbury terrane in the southwest from an autochthonous Boland domain in the northeast. The results of the detrital zircon study also underline the similar late-Neoproterozoic timing of sedimentation in the Saldania belt and parts of the Gariep belt to the north (Fölling et al. 2002). Given the back-arc setting inferred for the upper Gariep Supergoup (e.g., Basei et al. 2005), Frimmel et al. (2013) suggest a similar depositional setting for the Malmesbury Group having formed in the same back-arc realm to the east of the Dionisio Cuchilla-Pelotas magmatic arc (Fig. 14.1), now situated in the Dom Feliciano Belt of northeastern Brazil and Uruguay. In contrast to this, regional studies have long maintained similarities and rather gradual contacts of rocks of the Malmesbury Group and the contiguity of structures and similar metamorphic grades across the main fault zones (Rabie 1948; Von Veh 1983; Rozendaal et al. 1999; Belcher 2003; Belcher and Kisters 2003; Rowe et al. 2010). These studies rather advocate the autochthonous nature of the Saldania belt and its development along the western margin of the Kalahari Craton. The Colenso Fault is recognized as a regional-scale fault zone and significant structural break, but not as a terrane bounding fault.

This brief synopsis highlights some of the discrepancies in our understanding of the overall structural evolution, but also stratigraphic framework, of the belt. The present review focuses on the structural inventory and lithological similarities or differences of sequences across purported terrane boundaries. Previous regional works have hinted at the presence of two overlying domains in the belt with distinct lithological and structural inventories and strain histories (Rabie 1948; Newton 1966; Hartnady et al. 1974; Belcher 2003; Belcher and Kisters 2003). These domains extend across the large strike-slip faults in the belt (Fig. 14.3), a view supported by the recent detrital zircon study by Frimmel et al. (2013). The upper and structurally seemingly more simple domain covers most of the Saldania belt and encompasses the former Tygerberg terrane, parts of the central Swartland and large parts of the eastern Boland terrane (Fig. 14.3). The lower domain is structurally more complex and lithologically hetereogeneous, and while strains are clearly regional, neither the significance nor the structural relationship between the two domains are fully understood.

3 Lithostratigraphic and Structural Relationships in the Western Saldania Belt

3.1 Lithological Inventory of the Lower Domain: The Swartland Complex

Rocks of the lower domain are mainly exposed in the cores of the regional-scale antiforms (F2 folds) of the Swartland and Spitskop domes (Fig. 14.3) that provide windows into the lower structural levels of the belt (e.g., Rabie 1948, 1974; Newton 1966; Hartnady et al. 1974; Belcher 2003). Lower domain rocks broadly correspond to the Swartland Subgroup of previous subdivisions (SACS 1980; Gresse et al. 2006), including the Berg River and Klipplaat formations (Table 14.1). The two formations form supposedly more or less coherent stratigraphic packages distinguished by varying abundances of metapsammites versus metapelites. However, regional mapping cannot confirm the presence of laterally continuous lithological packages or marker units, and there seems little or no correlation of sequences between outcrops (Belcher 2003; Belcher and Kisters 2003). We also include the lower parts of the Morreesburg Formation, previously known as the Porseleinberg Formation (Hartnady et al. 1974), and metamafic rocks of the Bridgetown Formation into the lower domain. Importantly and in contrast to accepted subdivisions (e.g., SACS 1980), we consider the largest parts of the regionally widespread Morreesburg Formation not to be part of this lower domain or the Swartland Subgroup, rather forming part of the upper domain (see below). Given the degree of structural dismemberment, a subdivision of lower domain rocks into formations is a rather moot point and the rocks rather form part of a tectonostratigraphic complex, henceforth referred to as the Swartland complex (Fig. 14.3; Table 14.1).

Lithologically, the Swartland complex comprises a heterogeneous assemblage of quartz-sericite and chlorite-muscovite schists and phyllites, chlorite- and talc-carbonate schists, quartzites and muscovite quartzites, highly sheared chert horizons, graphitic schists and limestones, and grey, massive to well-bedded limestones (Fig. 14.4a) (Rabie 1948, 1974; Hartnady et al. 1974; SACS 1980; Theron et al. 1992; Belcher 2003; Belcher and Kisters 2003; Gresse et al. 2006). The main distinguishing characteristic of the Swartland complex in outcrop is the presence of a pervasively developed bedding-parallel phyllitic foliation (S1; see below) (Fig. 14.4b, c). Primary bedding is preserved in more competent units such as limestones or quartzites, but is largely obliterated by the pervasive S1 foliation and bedding transposition in metapelitic schists and phyllites (Fig. 14.4b). A further characteristic is the abundance of foliation (S1) parallel and mostly transposed quartz and quartz-carbonate veins (Fig. 14.4c). Quartz veins may constitute up to 20 vol.% of individual outcrops and testify to the pervasive fluid flow during formation of the bedding-parallel foliation and bedding transposition.

Fig. 14.4

a The lithological heterogeneity and complexity of rocks of the Swartland complex are best observed in quarry operations, illustrated here by the simplified geological map of the De Hoek quarry, outside Piketberg. Contacts between units are almost exclusively structural in nature, transposed into parallelism with the pervasive S1 foliation. The quarry is located on the steep northeast-dipping limb of a later (F2), large-scale fold, the Spitskop Dome (Fig. 14.3), which accounts for the steep dips of bedding (S0) and the foliation (S1). b Oblique cross-sectional view (to the southeast) of a metre-scale, isoclinal fold (F1) that refolds bedding (S0), enveloped by the pervasive S1 axial planar foliation (De Hoek Quarry, Piketberg). F1 folds are observed on a centimetre to metre scale and indicate the near-pervasive transposition of original bedding features in rocks of the Swartland complex into the subhorizontal S1 foliation. c Oblique cross-sectional view of a quartz-rich phyllite illustrating the progressive transposition of fabric elements into the composite S1 foliation in rocks of the Swartland complex (here taken from a roadcut at Bothmaskloof Pass, outside Riebeck Kasteel). Tightly folded bedding laminations (S0, left-hand side) and centimetre-scale, isoclinally folded (F1), rootless quartz veins (annotated by fine yellow line) are transposed into the pervasive S1 foliation. The dismembered, asymmetric quartz vein at the centre suggests deformation during top-to-the west shearing. Coin for scale is 2.5 cm in diameter

Despite the pervasive fabrics, the predominantly metapelitic sequence and interlayered metapsammite units preserve characteristics of an originally marine, probably metaturbiditic succession (Theron et al. 1992; Belcher 2003). A marine origin can also be inferred for the structurally interleaved graphitic schists and limestones. Chlorite and talc-carbonate schists are interpreted as altered mafic metavolcanic rocks (Rozendaal et al. 1994; Slabber 1995; Gresse et al. 2006). These chlorite schists are volumetrically minor but common throughout the Swartland complex. In quarries and in borehole sections the metamafic rocks form up to several metres thick, typically isolated lens-like bodies enveloped by the S1 foliation (Rozendaal et al. 1994; Belcher 2003). Scattered outcrops of chlorite-actinolite-epidote-carbonate schists and minor ultramafic talc-carbonate schist define a northwest-trending, c. 15 by 3 km sliver in the eastern parts of the Swartland domain (Rabie 1948, 1974; Hartnady et al. 1974; Theron et al. 1992; Slabber 1995). This largest, more or less coherent metamafic unit forms the Bridgetown Formation (Hartnady et al. 1974; SACS 1980) situated along the southern down plunge extent of the upright to southwest-verging F2 Spitskop dome (Fig. 14.2). The highly altered mafic metavolcanic rocks are intercalated with cherts and dolomite units along the western margins of the Bridgetown Formation against quartz-sericite phyllites (Theron et al. 1992; Belcher 2003). Geochemical signatures indicate that the Bridgetown Formation rocks represent original tholeiitic metabasalts (Slabber 1995; Rozendaal et al. 1999).

3.2 Lithological Inventory of the Upper Domain: The Malmesbury Group

Rocks outside D1 fabric domains and in large parts of the Saldania belt show significantly lower fabric intensities. The mainly metasedimentary and minor metavolcanic rocks lack the bedding-parallel foliation (S1), and primary bedding and even intricate depositional features are well preserved so that sedimentary environments are better constrained (Fig. 14.5a, b) (e.g., Von Veh 1983; Theron et al. 1992; Rowe et al. 2010; Frimmel et al. 2013). Wherever recorded, way-up criteria indicate the normal younging of the rocks (Fig. 14.5b). In the following, we highlight lithological similarities and the distribution of different facies across much of the belt that point to reasonable regional correlations between units, also across purported terrane boundaries. Following the current terminology, we henceforth refer to these rocks as the Malmesbury group. Importantly, and in contrast to currently accepted lithostratigraphic classifications, the structurally lower rocks of the Swartland complex are not included in the Malmesbury group described here (Table 14.1).

Fig. 14.5

a Rocks of the Malmesbury Group show well-preserved bedding, here developed as alternating shale and greywacke units exposed on the vertical limb of a F2 fold, looking northwest in rocks of the Tygerberg Formation along the Atlantic seaboard at Ganzekraal. b Intricate sedimentary features, here cross-bedding in the Tygerberg Formation, are well preserved throughout the Malmesbury Group. Despite the tight folding in many places, Malmesbury Group rocks show only very limited internal strain, which is in marked contrast to rocks of the underlying Swartland complex (Fig. 14.4b, c). c Synsedimentary deformation features are very common, particularly in the Tygerberg Formation (e.g., Von Veh 1983). The features shown here are interpreted as sand boils or sand volcanoes, forming ovoid sand protrusions through shale formed during the liquefaction of unconsolidated sands during seismic events. Tygerberg Formation, Silverstroomstrand (Fig. 14.2). d Oblique view of a greywacke channel fill erosive into thick black shale units that dominate the northern facies of the Tygerberg Formation, looking southeast, Ganzekraal

In current subdivisions (SACS 1988; Gresse et al. 2006), the westernmost parts of the Saldania belt are only underlain by one stratigraphic unit, the Tygerberg Formation. In coastal exposures and large quarries around Cape Town, this formation is developed as a greywacke-dominated succession with intercalated shales and siltstones (Hartnady et al. 1974; Von Veh 1983; Rowe et al. 2010; Frimmel et al. 2013). The predominance of massive greywacke units, presence of angular feldspar fragments and poor sorting of the rocks together with the occurrence of conglomerates and thin limestone units have commonly been inferred to indicate relatively proximal and shallow-marine conditions of deposition (Theron 1984; Theron et al. 1992). Soft-sediment deformation features are common and have been suggested to indicate a rapid deposition of the sediments, but also as records of seismic activity in the unconsolidated sediments (Fig. 14.5c) (e.g., Von Veh 1983). Detrital zircons from the southern parts of the Tygerberg Formation point to a late-Neoproterozoic timing of sedimentation and the youngest concordant zircons yield ages of c. 565–560 Ma (Armstrong et al. 1998; Frimmel et al. 2013). These ages correspond well with U-Pb ages of an intercalated felsic tuff from the metavolcanic Bloubergstrand Member in the Tygerberg Formation, north of Cape Town, indicating an age of volcanism of c. 554.5 ± 5 Ma (Kisters et al. 2015). The amygdaloidal textures of the intermediate lavas and eutaxitic textures of the tuff units are consistent with the shallow marine or, locally, even subaerial conditions of volcanism of the Bloubergstrand Member and deposition of the Tygerberg Formation in this part of the belt (Von Veh 1983). Coastal exposures of the Tygerberg Formation in the north, north of Silverstroomstrand, are dominated by massive to laminated black shale. Intercalations of greywacke are subordinate and metapsammites mostly form 10–50 m wide erosive, often laterally and vertically stacked channel fills encased by thick units of black shales (Fig. 14.5d). The rocks resemble slope channel fills acting as sand bypasses and delivering sediment into deeper-water environments. This deeper marine northern facies has not been distinguished in the undifferentiated Tygerberg Formation. However, lateral stratigraphic correlations and age relationships between these different facies have to remain tentative given the only sporadic outcrop.

East of the Colenso Fault, large parts of the Swartland domain are underlain by rocks of the Moorreessburg Formation. The interlayered greywacke and shale units resemble those of the Tygerberg Formation and regional studies maintain the similarity between the two formations across the Colenso Fault (e.g., Von Veh 1983; Belcher and Kisters 2003). The lithological similarities are corroborated by near identical detrital zircon populations that point to the deposition of the rocks in the late Neoproterozoic and up to c. 560 Ma (Frimmel et al. 2013). The Franschhoek Formation in the central parts of the western Saldania belt is distinct from the laterally more extensive Tygerberg and Moorreessburg Formation in a number of ways. The sequence is dominated by quartzites, feldspathic greywackes and conglomerates, with only minor intercalated shales, but also tuffs and amygdaloidal lavas and intrusive quartz-porphyry dykes (De Villiers et al. 1964; Theron et al. 1992; Gresse et al. 2006). Clasts in conglomerates are made up of original quartz vein material, chert, and quartzite, but also granite and phyllite. In contrast to the laterally extensive Tygerberg and Morressburg formations, rocks of the Franschoek Formation form northwest-trending, elongate depositories that also occur on either side of the Colenso Fault on regional maps (Fig. 14.6) (Theron et al. 1992). The rocks have clearly been affected by the regionally upright, northwest-trending F2 fold and, as a result, show steep dips and, in places, a well-developed foliation and northwest-trending stretching lineation. Taken in conjunction, lithological and structural characteristics of the Franschhoek Formation indicate an origin of the rocks during the localized uplift, erosion and reworking of underlying rocks of the Swartland complex and Morreessburg Formation and during CGS plutonism pointing to a later, probably syntectonic (D2 shortening, see below), timing of deposition. This also agrees with the unconformable contacts of the Franschhoek formation against underlying phyllitic units, but also parts of the CGS (Theron et al. 1992).

Fig. 14.6

Simplified geological map, modified after Theron et al. (1992), showing the relationships of the laterally more extensive Tygerberg and Morreessburg formations against the Franschhoek Formation and Klipheuwel Group. The latter are mainly confined to northeast-trending and, in the case of the Klipheuwel Group, fault-bounded basins. Granitic dykes are intrusive into the Franschhoek Formation but not into the Klipheuwel Group. Note also the seemingly unconformable contacts of the Franschhoek Formation against plutons of the CGS

Rocks in the eastern parts of the western Saldania belt are grouped under the Boland Subgroup that comprises the Piketberg, Porterville, Norree and Brandwacht formations. Contacts are not exposed and most studies suggest the lateral interfingering of, or gradual contacts between, formations, and also correlations with rocks of the Franschhoek Formation to the west (Hartnady et al. 1974; Theron et al. 1992). In general, the Boland Subgroup comprises coarser-grained and less well-sorted clastic rocks compared with the western Tygerberg or Morreessburg Formations. The sequence comprises sandstone and feldspathic sandstone interlayered with silt and shale, greywacke, gritty sandstone, conglomerate units and impure limestones. Most studies suggest a near-shore marine to deltaic and fluival and non-marine depositional environment for the mainly coarse-clastic rocks (Gresse and Theron 1992; Rozendaal et al. 1999). Conglomerates of the Piketberg and Porterville formations contain pebbles of vein quartz, phyllite and greywacke, suggesting a very proximal source for the sediments probably derived from the reworking of underlying rocks, similar to the Franschhoek Formation. Frimmel et al. (2013) document detrital zircon age spectra from the Porterville Formation that are similar to rocks from the western parts of the Saldania belt and possibly point to a slightly younger age of sedimentation of at least parts of the Boland Subgroup compared with the western Tygerberg and Morreessburg formations. Differences in the age spectra relate to the much more prominent input of late-Mesoproterozoic and early Neoproterozoic zircons. This likely reflects the proximity of the Namaqua-age (1200–950 Ma) metamorphic hinterland to the immediate east (e.g., Rozendaal et al. 1999). Parts of the Piketberg Formation in particular are strongly foliated and lineated, and fabric development resembles those found in rocks of the Swartland complex. The Brandwacht Formation in the far eastern parts of the western Saldania belt is unique in that it contains prominent metavolcanic units of broadly andesitic composition intercalated with conglomerates, greywackes and metapelites. Contacts of the Brandwacht Formation against older rocks are interpreted as either a basal conglomerate or a tectonic melange, possibly along a thrust plane (Hartnady 1969).

3.3 Klipheuwel Group

Clastic rocks of the Klipheuwel Group are preserved in a number of rhomb-shaped, narrow (<2–3 km), elongate (>10 km), northwesterly-trending basins (SACS 1980) that are spatially closely associated with the Colenso and, to a lesser extent, Piketberg-Wellington fault zones (Figs. 14.2 and 14.6). Correlations between individual basins are limited, suggesting that individual depositories were not interconnected. The rocks commonly show a coarse-clastic basis with conglomerates and poorly sorted sandstones, arkoses and intercalated shales that unconformably overlie shales and/or phyllites of the Swartland complex and/or Malmesbury group, but also granitic rocks of the CGS. Pebbles in conglomerates and gritty sandstones include vein quartz, shale and phyllite, clearly derived from the underlying metasediments, but also boulders of the CGS. The Klipheuwel Group is said to attain a maximum thickness of up to 2000 m (Gresse et al. 2006), and towards its upper parts the basin fills are commonly finer grained and dominated by shale and intercalated sandstones. This has led to the subdivision of the group into a lower Magrug and an upper Populierbos formation (Theron et al. 1992). Rocks of the Klipheuwel Group show mainly steep dips and northwesterly strikes. In places, a steeply dipping northwesterly-trending foliation is developed in shale units. Frimmel et al. (2013) suggested a maximum age of deposition of c. 550 Ma based on detrital zircon ages from the Magrug Formation in the lower parts of the Klipheuwel Group. Unconformable contacts between rocks of the Klipheuwel Group and the Darling batholith in the type locality of the Klipheuwel quarry provide a tighter age constraint of <535 Ma for sedimentation. An upper age of c. 480–500 Ma is provided by the age of the unconformably overlying Cape Supergroup that is considered to be lower Ordovician in age (Thamm and Johnson 2006).

Tankard et al. (2009) suggest that the Klipheuwel Group represents a rift-related precursor to the Cape Supergoup, but most other studies interpret the group as late-stage molasse-type basins more closely related to the Pan-African evolution of the Saldania belt (e.g., Gresse et al. 2006). Numerous workers emphasize similarities with rocks of the Franschhoek Formation (Belcher and Kisters 2003; Frimmel et al. 2013). Lower fabric intensities, the lack of intrusive granite dykes, generally better preservation and the typically fault-bounded nature of the rocks may indicate a slightly later deposition of the Klipheuwel Group compared with that of the Franschhoek Formation.

3.4 Granites of the Cape Granite Suite

Granites of the CGS hardly ever feature in regional reconstructions and tectonic models for the Saldania belt when they are a volumetrically important component and a clear distinguishing feature compared with adjoining Pan-African belts in southern Africa. A detailed discussion of the CGS is beyond the scope of this chapter, so the reader is referred to, inter alia, Scheepers (1995), Scheepers and Schoch (2006), Stevens et al. (2007), Villaros et al. (2009, 2011), Chemale et al. (2011) and Farina et al. (2012) for more comprehensive accounts on the origin of the granites. However, any geodynamic considerations have to take the emplacement, petrogenesis and timing of the CGS into account that illustrate widespread partial melting of mid- and lower crustal levels underlying the belt. Aspects pertaining to the geodynamic setting of the granites are briefly presented below.

Most studies focus on the petrogenetic evolution of the peraluminous S-type granites (e.g., Stevens et al. 2007; Villaros et al. 2009, 2011; Harris and Vogeli 2010; Farina et al. 2012, to name but a few). S-type granites form five major and composite plutons that are confined to the west of the Colenso Fault (Fig. 14.2). Available age data points to an emplacement of the plutons between c. 550 and 535 Ma and as late as c. 525 Ma for the latest phases in the central Darling batholith (Da Silva et al. 2000; Villaros et al. 2009). Subvolcanic textures and miarolitic cavities in some of the c. 540 Ma S-type granites (e.g., Rozendaal and Bruwer 1995) point to fairly shallow emplacement levels of the plutons. This may also account for the presence of granitic pebbles in rocks of the Franschhoek and Piketberg formations and the Klipheuwel Group, and it also suggests only limited uplift and erosion of the rocks after their emplacement.

Most petrological studies discuss a lithologically heterogeneous but broadly aluminous metasedimentary mid- and lower crustal source for the S-type granites and partial melting of the rocks through fluid-absent reactions (e.g., Stevens et al. 2007; Villaros et al. 2009; Farina et al. 2012). As such, the granites provide windows into the deeper parts underlying the Saldania belt. Harris and Vogeli (2010) document oxygen isotope compositions of garnet in the c. 540 Ma S-type Pensinsula Pluton around Cape Town that are consistent with the partial melting of a metapelitic source with an oxygen isotope composition similar to that of the Malmesbury Group exposed on surface. Also from the Peninsula Pluton, Villaros et al. (2011) document ages as young as 570 Ma from cores of inherited zircons in the granites, showing magmatic overgrowths with ages of between 555 and 525 Ma. In another study, Villaros et al. (2009) document high-grade metamorphic assemblages from metapelitic xenoliths in the central Darling batholith that are compositionally similar to rocks of the Swartland complex. Mineral assemblages in xenoliths record P–T conditions of c. 10 kbar and T at c. 800 °C. On a similar point, although outside the CGS, Kisters et al. (2015) recorded xenocrystic zircon populations in 555 Ma old tuffs intercalated with the Tygerberg Formation that are near identical to detrital zircon populations from the Malmesbury Group and Swartland Complex (Frimmel et al. 2013). While being circumstantial, these results consistently point to Pan-African age, Malmesbury-type rocks in the anatectic source region of the magmas. This would indicate an at least structural thickness of Pan-African rocks of in excess of 25 km, about ten times the thickness of the similarly aged sequences in the Gariep belt to the north.

There are only a few studies of I-type granites of the CGS (for a comprehensive summary, see Scheepers and Schoch 2006). The I-type granites form elongate, northwest-trending plutons east of the Colenso Fault, largely intruding rocks of the central Swartland domain. Existing ages suggest that the 540–520 Ma I-type granites are slightly younger than S-types. However, the ever-improving geochronological database shows a considerable overlap of ages between S- and I-type granites and the two granite types are essentially coeval.

4 Structural Geology

4.1 D1 Structures and Fabrics of the Swartland Complex

The structural complexity and lithological heterogeneity of the Swartland complex is probably best appreciated in quarry operations that provide excellent 3D windows into the mélange-like sequence (e.g., Fig. 14.3a) (Belcher 2003; Belcher and Kisters 2003). The pervasive bedding-parallel phyllitic foliation (S1) is axial planar to and envelopes centimetre- to metre-scale isoclinal, intrafolial folds (F1) that testify to the widespread transposition of original bedding and the high-strain nature of the S1 foliation (Fig. 14.3b). Quartz and quartz-carbonate veins are abundant throughout the lower domain rocks where they appear almost invariably transposed into the S1 foliation, having undergone boudinage or pinch-and-swell within the foliation, particularly in phyllitic units (Fig. 14.3c). A northwesterly-trending, gently doubly plunging mineral and mineral stretching lineation (L1) contained in S1 is prominent throughout the Swartland complex. L1 parallel rodding textures are particularly common in chert units (Belcher 2003). In phyllites, F1 intrafolial folds show shallow northwesterly and/or southeasterly plunges parallel to the L1 stretching lineations indicative of the rotation of fold hinges into the regional stretch during bedding transposition.

Contacts between thicker lithological packages are commonly sheared, which also accounts for the discontinuous extent of units (Hartnady et al. 1974; Belcher 2003; Gresse et al. 2006) (see below). Competent limestone units commonly preserve bedding, and S1 fabric intensities are lower than in the pervasively transposed phyllites and schists. Instead, strain is localized along discrete, bedding-parallel calc-mylonites (Fig. 14.7a). Shear sense indicators such as S–C′ fabrics and rotated quartz-calcite aggregates are common and consistently point to a top-to-the-northwest sense of shear, parallel to the L1 lineation. Similarly, metre-scale duplex structures point to the imbrication of the rocks during top-to-the-northwest thrusting (Fig. 14.7b). Fault-propagation folds associated with thrusts show northeasterly trends at high angles to the thrust kinematics. Contrary to the pervasively transposed quartz veins in phyllitic units, carbonate and quartz-carbonate veins in limestones form a pervasive network of intersecting bedding-normal extensional veins and high-angle conjugate, and less abundant bedding-parallel veins (Fig. 14.7c). As such, D1 fabric development documents a distinctly mixed continuous-discontinuous deformation. Bedding transposition and the pervasive S1 foliation in schists and phyllites record, for the most part, continuous fabric development during distributed creep and ductile flow. The extensive high-angle vein networks and hydraulic breccias in more competent units record discontinuous brittle deformation, probably in the presence of temporarily supralithostatic fluid pressures. In general, foliation (S1) development, associated folding and boudinage of units, and the orientation of extensional and conjugate shear veins record the vertical shortening of lower domain rocks (Fig. 14.7c). Bedding-parallel thrusts, thrust-related folds and stretching lineations document a concomitant component of simple shear associated with a northwest stretch (Fig. 14.7a, b). Kinematic indicators consistently point to top-to-the-northwest and west thrusting along the low-angle thrusts. This strain pattern of vertical shortening and associated top-to-the-west and northwest non-coaxial shearing characterizes the Swartland complex throughout the central parts of the belt (Belcher 2003).

Fig. 14.7

a Vertical section of a mylonitic shear zone (top) and associated footwall deformation documenting top-to-the-northwest low-angle shearing in graphitic limestone of the De Hoek Quarry. Yellow lines annotate folded bedding and the shear zone boundaries. While strain is more distributed in phyllitic units (e.g., Fig. 14.3b, c), discrete calc-mylonites are developed in limestone units. b Metre-scale duplex structure in limestone suggesting top-to-the northwest and west low-angle thrusting at De Hoek quarry, Piketberg. c Cross-sectional view of conjugate calcite veins developed in low-strain domains within dark, graphitic limestones at limestone quarry, Riebeck West. The orientation of the conjugate sets together with the orientation of subvertical extensional veins points to the vertical shortening of the sequence during deformation and fluid flow. d Metasedimentary xenoliths within a finer-grained and low-strain phase of the Darling batholith, west of the Colenso fault (Fig. 14.2). Xenoliths preserving isoclinally folded bedding are common in the granites (here encircled xenolith on the left). These folds resemble F1 intrafolial folds developed in the Swartland complex and may indicate the wider extent of the lower, transposed domain at depth and across the Colenso fault

Rocks of the Swartland complex are not limited to the cores of the regional-scale Swartland and Spitskop domes, and several smaller windows expose transposed phyllitic units and, indeed, higher-grade metamorphic rocks. For example, Belcher (2003) describes isoclinally folded biotite gneisses from the farm Kanonkop north of Malmesbury (Fig. 14.2). These isolated outcrops highlight the much wider extent of the rocks in the western Saldania belt. This distinction has not been made on regional maps (e.g., Theron et al. 1992) but is clearly evident where low-angle foliations (dips < 40–50°) are indicated, pointing to the presence of D1 fabric domains. The distribution of these outcrops also suggests shallow dips along which rocks of the overlying Malmesbury Group rest on the Swartland complex, either along an unconformity (Belcher and Kisters 2003) or along a major low-angle fault (Rabie 1948). In this context, the abundance of high-grade metasedimentary xenoliths containing intrafolial folds and transposed bedding (e.g., Schoch 1975) typical of D1 fabric domains in the Darling batholith, west of the Colenso fault is conspicuous (Fig. 14.7d). Although circumstantial, the large number of these xenoliths may point to the extent of the Swartland complex at depth and across the Colenso fault.

4.2 D2 Regional Folding and Associated Strains

The D1 fabrics and strains described above are confined to rocks of the Swartland complex. The most obvious and regionally widespread structures that have affected the western Saldania belt are northerly- to northwesterly-trending folds and associated fabrics, collectively summarized under the D2 deformation event. Where rocks of the Swartland complex are exposed, F2 folds clearly refold earlier D1 fabrics, resulting in the steepening and crenulation of earlier low-angle fabrics (Hartnady et al. 1974; Belcher 2003; Belcher and Kisters 2003). Most mappable F2 folds are second- or third-order folds with wavelengths between c. 150 and 500 m (Fig. 14.8a). First-order F2 folds have wavelengths exceeding 3–5 km. The Swartland dome is probably the largest fold in the central parts of the Saldania belt with a width of c. 10 km. F2 folds are typically doubly plunging towards the northwest and southeast. The folds are more or less upright or steep southwesterly verging, but domains with northeast-verging folds also exist. Fold interlimb angles are highly variable and range from very gentle (>140°) to tight and isoclinal, even between adjacent folds. An upright, broadly axial planar cleavage (S2) is prominent in shaly units but is only very weak or absent in psammitic rocks (Fig. 14.8b). During their detailed analysis of F2 fold structures in the Tygerberg Formation on Robben Island, Rowe et al. (2010) showed the S2 foliation to be a transecting cleavage with a consistent clockwise rotation of S2 with respect to the axial planes for F2 folds. This finding can be extrapolated to F2 folds throughout much of the Saldania Belt and the S2 foliation commonly describes a clockwise rotation by 5–15° with respect to the axial surfaces of F2 folds (e.g., McGibbon 2012). Rowe et al. (2010) interpreted the transecting cleavage to indicate F2 folding during sinistral transpression. Quartz veins are common but are mainly restricted to more competent greywacke units. Along coastal exposures, quartz veins and vein sets show remarkably consistent orientations, also with respect to the orientation of F2 folds. Veins are high-angle extensional veins normal to the northwesterly-trending F2 folds or conjugate vein sets, both pointing to a component of subhorizontal, hinge-parallel north(west)-south(east) stretch during F2 folding (Fig. 14.8c). The intersection between bedding (S0) and the upright S2 foliation results in a variably developed, northerly- to northeasterly-trending intersection lineation (L2). Stretching lineations comparable to those in rocks of the lower domain are conspicuous by their absence.

Fig. 14.8

a Gentle, near-horizontal, upright F2 synform in black shales of the Tygerberg Formation, north of Ganzekraal, looking northwest, parallel to the trend of the fold. b Bedding (S0)–cleavage (S2) relationship in shale units on the southwest limb of an F2 antiform at Ganzekraal, looking northwest. c Oblique view of a central greywacke horizon interlayered with dark shales on the subvertical limb of a tight F2 fold at Ganzekraal. Systematic, high-angle quartz vein sets (annotated by yellow lines) are largely confined to the competent greywacke beds. Conjugate and extensional quartz veins indicate a layer-parallel stretch and, as such, hinge-parallel extension during F2 folding. d Magmatic fabric defined by the preferred orientation of k-feldspar megacrysts in the porphyritic phase of the Darling granite. Magmatic fabrics trend northwest-southeast, parallel to the Colenso fault, and are preserved in low-strain domains in the Darling batholith. e Gneissic textures in the Darling batholith result from the pervasive dynamic recrystallization of minerals and the overprint of originally magmatic textures. Here, augen textures are defined by ovoid, marginally recrystallized k-feldspar megacrysts and an anastomosing foliation defined by quartz ribbons and the preferred orientation of biotite. Plan view of the subvertical, northwest-southeast-trending foliation, parallel to the Colenso fault. f Plan view of mylonitic and protomylonitic textures in original granites of the Darling batholith define northwest-trending, up to 250 m wide mylonite belts overprinted on original magmatic and later solid-state gneissic fabrics. Shear sense indicators are common, here expressed by S–C′ fabrics, the majority of which indicate a sinistral sense of shear

Fig. 14.9

Schematic sketch illustrating the different lithological and structural elements of the Malmesbury fore arc as discussed in this chapter, made up of an underlying accretionary complex, the Swartland complex, and sediments of the overlying fore-arc basin, represented by the Malmesbury group. Later deformation by F2 folds and plutons of the CGS are not shown for clarity. Tectonic underplating during subduction in the deeper parts of the prism and deposition of fore-arc sediments at higher structural levels are contemporaneous and can be recorded between >560 Ma and at least 520 Ma. See text for further detailed discussion

4.3 D2 Strike-Slip Faults

The Colenso and Piketberg-Wellington fault zones are the largest structural features of the belt (Hartnady et al. 1974; Theron et al. 1992; Belcher 2003; Frimmel et al. 2011, 2013). The Colenso Fault is far more prominent and better exposed, and it can be traced for some 180 km along its northwesterly strike (Fig. 14.2). Deformation along the Colenso fault is distributed and the fault is developed as an up to 8 km wide anastomosing fault zone (Kisters et al. 2002). The best and most extensive exposures of the fault zone are in granites of the CGS, particularly the large, central Darling batholith. The batholith is bounded along its northeastern margin by mylonites and cataclasites related to the Colenso Fault. Intrusive and fabric relationships document the synmagmatic deformation of successively emplaced granite phases. Northwesterly-trending, steeply dipping magmatic foliations in the granites are commonly overprinted by high-temperature solid-state fabrics (Fig. 14.8d–f). This results in regionally widespread although variably pervasive gneissic textures (Schoch 1956; Theron et al. 1992; Kisters et al. 2002). Northwest-trending, subvertical mylonite and protomylonite zones show gradational contacts with magmatic and weak solid-state fabrics developed in large parts of the pluton. Brittle-ductile, anastomosing cataclasite and ultracataclasite zones indicate that deformation has continued under lower-grade conditions and accompanying the cooling of the granites. Magmatic and solid-state lineations are defined by the stretching of magmatic enclaves or xenoliths, the preferred alignment of k-felspar megacrysts or stretched quartz-feldspar aggregates in mylonites (Schoch 1975; Kisters et al. 2002). Lineations show mainly shallow northwest and/or southeast plunges consistent with the mainly strike-slip kinematics. The vast majority of magmatic and solid-state shear sense indicators point to sinistral kinematics (Fig. 14.8f), but there are also domains of seemingly conflicting dextral and sinistral kinematics. Available U-Pb ziron ages from granites of the Darling batholith point to deformation along the fault zones between at least 545 and 525 Ma (Da Silva et al. 2000; Villaros et al. 2009). The c. 510 Ma undeformed Klipberg granite (Scheepers 1995) crosscuts earlier magmatic and solid-state fabrics in the Darling batholith and indicates the cessation of deformation by that time. Northwest of the Darling batholith, fault rocks related to the Colenso Fault zones are developed in the Cape Columbine granite and Vredenburg adamellite. Faulting is evidenced by 2–3 km wide networks of anastomosing, northwesterly-trending mylonites, ultracataclasites and cataclasite zones (Schoch 1956). Shear sense indicators in these younger, 540–520 Ma granites mainly point to dextral kinematics, suggesting a late-stage reversal of shear along the Colenso fault (Kisters et al. 2002).

The Piketberg-Wellington fault is only poorly exposed and its actual trace and location are controversial (Rabie 1948, 1974; Slabber 1995; Belcher 2003). Rabie (1974) postulated the presence of the fault to account for the juxtaposition of sheared and transposed rocks of the Swartland complex in the west against only weakly deformed rocks of the Boland domain in the east. Notably, this area corresponds to the eastern limb of the first-order F2 Swartland dome and smaller Spitskop dome that are cored by the Swartland complex. Around Piketberg, faulting is evidenced by several northwest-trending, anastomosing fault strands and associated quartz veining. Where faulting can be identified with certainty, rocks of the Cape Supergroup have been displaced, forming downfaulted outliers surrounded by Pan-African basement. This suggests a significant component of post-Cape Supergroup displacement along the fault, with a mainly normal sense of movement and downthrow to the east. Along its southern extent, the trace of the Piketberg Wellington Fault can only be inferred (e.g., Theron et al. 1992). Despite this, Frimmel et al. (2013) suggest that the fault represents the major terrane boundary in the western Saldania Belt, but it should be emphasized that neither the location, nor the extent or actual timing and kinematics of the fault zone have been established with any certainty.

5 Discussion

The recognition of two structurally overlying tectonostratigraphic domains in the western Saldania belt is not new (e.g., Rabie 1948, 1974; Newton 1966; Hartnady et al. 1974; Belcher 2003; Gresse et al. 2006), but the significance of these contrasting domains for the overall evolution of the belt has never been discussed or fully appreciated. The detrital zircon study by Frimmel et al. (2013) demonstrates the, within error, identical late-Neoproterozoic age of rocks of the Malmesbury group and structurally underlying Swartland complex. Hence the notion expressed by Belcher and Kisters (2003) of an older D1 domain overlain by a younger, less deformed metasedimentary sequence cannot be upheld. This implies that the Swartland complex and the Malmesbury group represent two different structural levels, and we suggest that the contrasting lithological inventories, kinematics and strains illustrate a section through a fore-arc region situated along the western margin of the Kalahari Craton as has similarly been suggested by Von Veh (1983), Belcher (2003) and Rowe et al. (2010). The Swartland complex exposes the upper parts of an accretionary prism whereas rocks of the overlying Malmesbury group represent the relics of a reasonably coherent, although deformed, fore-arc basin on top of the accretionary prism. These aspects are detailed below.

5.1 Swartland Complex

The Swartland complex represents a highly sheared and transposed tectonostratigraphic sequence that records the offscraping and imbrication of marine sediments and slivers of oceanic crust (e.g., Hartnady et al. 1974; Belcher 2003). Strains record the vertical shortening of the sequence combined with a component of subhorizontal shear and top-to-the-west and northwest-thrusting along low-angle mylonites and phyllite zones. The combined vertical coaxial shortening strain and non-coaxial subhorizontal slip are diagnostic for accretionary systems formed by the underthrusting and basal accretion of rocks to the base of the accretionary complex, resulting in the thickening of the wedge during tectonic underplating of marine strata and oceanic crust (e.g., Fisher and Byrne 1987; Raimbourg et al. 2009; Malavieille 2010). The pervasive quartz and quartz-carbonate veins and vein networks document the progressive dehydration of underthrusted sediments during burial and basal accretion. Competent limestone units preserve the original geometries of these vein networks. The brittle vein networks and associated hydraulic breccias testify to episodically supralithostatic fluid pressures and associated seismic slip events during underplating of the rocks. In the regional context of southwest Gondwana, the top-to-the-west and northwest kinematics agree with an imbrication of the rocks during east- or southeast-directed convergence and subduction of oceanic crust below the Kalahari Craton. This corresponds to the kinematics recorded from coastal belts to the north (e.g., Hartnady et al. 1985; Von Veh 1993; Gresse 1995; Goscombe et al. 2003; Konopasek et al. 2005). The age spectra of detrital zircons in rocks of the Swartland complex provide further evidence of the setting of the rocks along an active continental margin. The ages of detrital zircons (Frimmel et al. 2013) suggest that deformation, burial and metamorphism of rocks in the Swartland complex followed shortly, within <5–10 Ma, the deposition of the rocks, typical of convergent plate margins (e.g., Cawood et al. 2012).

5.2 Malmesbury Group

Lithological assemblages and strains in the overlying Malmesbury group mark a sharp break against the underlying mélange-like rocks of the Swartland complex, although the actual contacts between the two are not exposed. Despite the only patchy outcrop and later folding of Malmesbury group rocks, the areal distribution of facies associations allows for the reconstruction of a depositional model for the Malmesbury group across purported terrane boundaries. The coarse-clastic sedimentary and intercalated volcanic rocks of the Boland Subgroup in the east define a northerly- to northwesterly-trending belt of non-marine to near-shore and shallow-marine sediments proximal to a volcanic arc located to the east. The proximal facies shows a pronounced influence and sediment input from the Kalahari Craton as an eastern hinterland (Tankard et al. 1982; Gresse and Theron 1992, 2006), which is underpinned by the prominence of Meso- to Neoproterozoic zircons in the rocks (Frimmel et al. 2013). To the west, the mainly proximal facies of the Boland Subgroup is replaced by the greywacke-dominated turbidite deposits of the Tygerberg and Morreessburg formations. The metaturbidites are laterally the most extensive units, showing mainly tabular geometries and a lack of channels at their base consistent with a depositional environment in the more central parts of the fore-arc basin. This agrees with the views of Von Veh (1983), who suggests deposition of the southern Tygerberg Formation along a progradational submarine fan along a continental slope. He also evokes the proximal facies of the southern Tygerberg Formation to be derived from rocks to the immediate east. The abundant soft-sediment deformation features agree with the deposition of the rocks along an active continental margin (see also Von Veh 1983) and represent either the near-surface expression of seismic events as they are recorded by hydrothermal breccias and vein networks preserved in the Swartland complex or very high sedimentation rates and gravitational instabilities as expected in this environment.

A deeper-water environment is indicated for the shale-dominated and stacked channel fills of the northern Tygerberg Formation north of Silverstroomstrand (Fig. 14.3). Hartnady et al. (1974) and Tankard et al. (1982) envisage a hemipelagic, trench-like sedimentary environment for this northern facies of the Tygerberg Formation. The shallow marine and, in places, subaerial depositional conditions in parts of the southern Tygerberg Formation may correspond to the location of a trench-slope break. This outer arc high separates the inner fore arc basin in the east (southern Tygerberg and Morreessburg formations) from the slope apron deposits to the west and northwest (northern Tygerberg Formation). In general, the regional distribution of facies indicates the transverse sediment input from the arc side via river systems and deltas in the east into and across the deepening fore-arc basin towards the west, as is common for fore-arc basins (e.g., Heller and Ryberg 1983; Santra et al. 2013). Contacts between formations and facies are gradational and across purported terrane boundaries.

The reworking of underlying rocks, including phyllites and granites of the CGS in rocks of the Franschhoek and Piketberg formations, is indicative of the deposition of these formations during or very shortly after plutonism and metamorphism at depth. Folded, unconformable contacts of the Franschhoek Formation against underlying phyllites suggest the syntectonic formation of the coarse-clastic rocks, possibly related to episodes of intrabasinal uplift during regional shortening (D2) along either growth folds or steepened breakthrough thrusts (e.g., Noda 2016). The different basins may have formed at different times but they correspond to similar processes that signify the progressive shortening (D2), localized uplift, erosion and deposition within the larger Malmesbury fore-arc basin between at least 550 Ma to probably 520 Ma. The synsedimentary deformation will also have an effect on the basin topography and, hence, the geometry of sediment supply paths and resulting flow and transport directions recorded in the sediments (e.g., Buggisch et al. 2010; Rowe et al. 2010). Lower strain intensities, better preservation and the strictly fault-bounded nature of basins of the Klipheuwel Group point to a slightly later sedimentation compared with, for example, that of the Franschhoek Formation, probably related to the late-stage uplift history.

5.3 Deformation of the Fore Arc (D2)

The western Saldania belt shows a very characteristic domainal structural pattern in which regions of upright to southwest-verging folds (F2), up to several tens of kilometres wide, are bounded by the several kilometre-wide vertical sinistral strike-slip corridor of the Colenso fault zone as the central structure. This pattern of orogen-parallel transcurrent shear zones bordering or enveloping folded domains is characteristic of obliquely convergent margins and points to the partitioning of the overall transpressional deformation into a contractional and a transcurrent component (Fig. 14.10a, b) (e.g., Robin and Cruden 1994; Tikoff and Teyssier 1994; Fossen and Tikoff 1998; Tikoff and Peterson 1998). Folded (F2) domains record the largely coaxial shortening component whereas strike-slip faults, such as the Colenso Fault, accommodate the transcurrent component of the strain. The subparallel orientation of F2 folds and the northwest-trending Colenso fault corridor also suggest that parts of the strike-slip component during oblique convergence were accommodated by the strike-slip shear zones, rather than being distributed across the belt (Fig. 14.10a, b). This pronounced strain partitioning may be related to the very oblique subduction and convergence angles of <20° (Rowe et al. 2010) and/or point to the decoupling of the shortening and strike-slip components in the belt as a result of weak fault rheologies, particularly of the main Colenso fault. Notably, the Colenso fault zone is, for most of its strike extent, developed in granites of the CGS. While this may be an artefact of the outcrop conditions, the progressive fabric development from magmatic via high-temperature to low-temperature solid-state fabrics underlines the spatial and temporal relationships between granite emplacement and strain localization (Fig. 14.8d–f). The granitic magmas are rheologically weaker than the surrounding wall rocks and magma emplacement will most likely result in the weakening of the fault zones. This in turn promotes strain localization and partitioning of the strike-slip component into the faults. Given that repeated granite sheeting in the Darling batholith is recorded between c. 545 and 525 Ma, magma emplacement may have, at least intermittently, contributed to the decoupling and strain partitioning in the belt over a period of over 20 Ma. Despite this, the fold hinge-parallel stretch (Fig. 14.8c) and the widely recorded transecting S2 cleavage (e.g., Rowe et al. 2010) in F2 fold domains is evidence for more distributed D2 strains. Given that magmatic sheeting is most likely episodic, one could argue that these more distributed strains have developed during episodes of fault lock-up and fault zone strengthening related to intervals when the faults were not intruded by granitic sheets.

Fig. 14.10

Schematic sketch illustrating the effects of a homogeneous distribution of strain and b strike-slip partitioned strain in transpressional zones on the orientation of fold hinges (after Fossen and Tikoff 1998). In (a), folds form at high angles to the shortening direction and fold hinges rotate into parallelism with the regional stretch during progressive deformation. In (b), the strike-slip component is partitioned into discrete transcurrent shear zones whereas the contractional component is accommodated by domains of largely coaxial shortening (folding) and fold hinges are subparallel to transcurrent shear zones. This latter case seems to be realized in the Saldania belt and either indicates the very weak rheology of the Colenso fault as the main transcurrent shear zone and/or a low angle (<20°) of plate convergence

On a regional scale, the D2 strike slip faults will not have contributed significantly to the shortening of the belt, but the open to tight and isoclinal F2 folds suggest a shortening of the fore arc by at least 30–50%, although the highly variable fold interlimb angles and lack of outcrop render these estimates tentative at best. Rocks of the Franschhoek and Piketberg formations testify to the synsedimentary deformation of the fore arc. This is consistent with the timing of strike-slip shearing recorded from the Colenso fault between at least 545 Ma and probably 520 Ma (Kisters et al. 2002). Given that the oldest c. 550 Ma granites of the CGS crosscut F2 folds (e.g., Theron 1984; Theron et al. 1992), D2 strains related to convergence and subduction can be recorded over at least 30 Ma from >550 to c. 520 Ma.

Sinistral transpressive kinematics are similarly documented from coastal belts to the north over a distance of >2000 km. There is general consensus that these strains document the oblique closure of the Adamastor ocean or related basins, and the eventual collision of terranes (Von Veh 1993; Gresse 1994, 1995; Goscombe et al. 2003; Konopasek et al. 2005; Goscombe and Gray 2008). Despite the similar regional kinematics, the structural style of upright folds and vertical strike slip faults in the western Saldania belt is distinct from, for example, the Vanrhynsdorp and Gariep belts to the immediate north. These northern belts are developed as east- to southwest-verging fold-and-thrust belts, with only little evidence of major vertical transcurrent faults. Importantly, these belts lack syn- to late kinematic granites such as those of the CGS in the Saldania belt. This may underscore the close relationship between synkinematic granite intrusions and the effects of at least temporary rheological weakening and resulting strain partitioning in the belts that gives rise to markedly different structural patterns in adjacent belts. The intrusion of the granites in turn likely reflects different basement structures underlying the belts.

5.4 The Deeper Structure of the Western Saldania Belt

Neither the base nor the top of the Pan-African rocks are exposed, so thickness estimates must remain tentative. However, metamorphic, geochemical and geochronological data from the CGS (Villaros et al. 2009, 2011; Harris and Vogeli 2010) and tuffs in the Tygerberg Formation (Kisters et al. 2015) point to the partial melting of metapelite-dominated, Pan-African age rocks at depths of >20–25 km. If correct, these thickness estimates are an order of magnitude greater compared to the thicknesses considered for similarly old rocks of the Gariep Supergroup to the north (e.g., Frimmel and Frank 1998; Gresse et al. 2006). A structural thickness of >20 km agrees with the thickness for accretionary prisms but is unlikely to be realized in a back-arc setting, as has been suggested in regional correlations of rocks of the Saldania with those of the Gariep belt (Frimmel et al. 2011, 2013). More precisely, the substantial thicknesses would point to the inner parts of the prism that has been structurally thickened through the underplating of sediments and oceanic crust. This in turn would correspond to the position of the unconformably overlying fore-arc basin sediments of the Malmesbury group above the inner parts of the wedge. In this scenario, the toe of the accretionary prism would be located to the far west, and the eastern and top parts of the frontal wedge would correspond to the deeper-water facies of the northern Tygerberg Formation forming the slope apron facing the ocean basin to the west (Rowe et al. 2010).

The plutons of the CGS may provide additional clues about the deeper structure of the western Saldania belt. Compositional variations in granites or granite suites are thought to mainly reflect source heterogeneities (e.g., Clemens et al. 2009, 2010). S-type granites are derived from the partial melting of aluminous clastic sediments, whereas I-type granites indicate former arc material or orthogneissic basement rocks in the source (e.g., Chappell and White 1974; Wall et al. 1987; Clemens 2003). The Colenso fault forms a sharp divide between S- and I-type granites in the western Saldania belt (Scheepers 1995; Scheepers and Schoch 2006; Stevens et al. 2007; Villaros et al. 2009), suggesting that the fault separates two compositionally different source regions at depth. This may imply the location of the Colenso fault at depth along the leading edge of the fore-arc basement (Fig. 14.9). Partial melting west of the fault has only affected the metasediments of the thickened accretionary wedge yielding the voluminous S-type granites. I-type granites east of the fault, in contrast, indicate a more heterogenous source made up of metasediments and older basement gneisses.

This raises the question as to the heat source for partial melting. Fore-arc regions are commonly considered to be the coldest part of convergent margins refrigerated by the subducting oceanic slab. The syn- to late-stage timing and fore-arc position of the CGS has previously been explained in terms of a slab break-off (e.g., Belcher 2003). Late tectonic fore-arc magmatism related to slab break-off is invoked for a number of convergent orogens, such as the Acadian of Maine (Schoonmaker et al. 2005) or, more recently, the Pan-African Damara belt in Namibia (Meneghini et al. 2014; Clemens et al. 2017). The slab break-off may also explain the lack of a hard collision in the western Saldania belt (e.g., Rozendaal et al. 1999). Unlike adjoining coastal belts to the north, rocks of the Saldania belt have not experienced high-grade metamorphism and lack evidence of a post-collisional uplift prior to the onset of Cape Supergoup sedimentation at c. 480–500 Ma. This would agree with a slab break-off in which slab delamination has resulted in a soft collision between the southern tip of the Kalahari Craton and the Rio de la Plata Craton.

6 Conclusions

This review is an attempt to present a different view of the geology of the western Saldania belt that is less based on the premise of allochthonous or parautochthonous domains and rather highlights correlations of lithological packages, facies and structures across purported terrane boundaries. The picture that emerges then is that of a deformed but reasonably well-preserved section through a fore arc that formed in the latest Neoproterozoic and into the Cambrian (>560 to <510 Ma) along the western margin of the Kalahari Craton. The structurally lower parts of the belt, the Swartland complex, record the imbrication of marine sediments and oceanic crust. Tectonic underplating and thrust imbrication have resulted in a complex tectonostratigraphic sequence with a thickness of likely >20–25 km. This section corresponds to the inner part of an accretionary prism that formed during southeast-directed oblique subduction of the Adamastor ocean below the Kalahari Craton. The accretionary prism is overlain by low-grade metamorphic metasediments and minor metavolcanic rocks with markedly lower fabric intensities, the Malmesbury group. Rocks of this group represent the late-Neoproterozoic to Cambrian fore-arc basin fill (Belcher 2003; Rowe et al. 2010). Contacts between formations are gradational and across purported terrane boundaries. Sedimentary facies suggest the presence of a volcanic arc in the east, succeeded in the west by shallow-marine turbidites of the inner fore-arc basin. Metapelitic successions and sandy channel fills in the far west and northwest indicate deeper-water deposits overlying the toe of the prim and facing the ocean basin. Late-stage, unconformably overlying coarse-clastic sediments in large parts of the belt represent syntectonic (D2) deposits that record the reworking of underlying sequences during shortening of the fore arc. Overall, the Malmesbury fore arc has been shortened by at least 30–50%, but strain (D2) is heterogeneous and partitioned into domains of coaxial northeast-southwest shortening and sinistral strike-slip shear. Coaxial shortening is documented by ubiquitous northerly- to northwesterly-trending, upright to southwest-verging folds (F2). The non-coaxial component of the strain is partitioned into regional-scale northwest-trending strike-slip fault zones, in particular the main Colenso Fault Zone. The sinistral transpressive strains agree with the oblique closure of the Adamastor ocean during the late Neoperoterozoic and probably until 520 Ma. The lack of burial metamorphism and subsequent exhumation that has also preserved original depositional environments and shallow intrusive features suggests that the deformation in the western Saldania records a soft collisional event, probably as a result of slab break-off. This break-off may also account for the voluminous, syn- to late-tectonic (550–510 Ma) granite plutonism of the CGS in the fore-arc region.