Archaean Hydrothermal Systems in the Barberton Greenstone Belt and Their Significance as a Habitat for Early Life

Chapter

Abstract

Hydrothermal systems have played an important role in shaping the 3.54–3.23 Ga volcano-sedimentary succession of the Barberton greenstone belt. Evidence for relatively low-temperature (≤150°C) seafloor hydrothermal activity is widely recorded in extensive silicification of volcanic and sedimentary rocks, leaching of elements commonly mobile during water-rock interaction, and extensive hydraulic fracturing. Evidence for the presence of high-temperature hydrothermal vents is scarce and restricted to a massive sulphide deposit near the top of the succession. Many of the zones affected by seafloor alteration are spatially associated with traces of early life, such as carbonaceous matter and bioalteration features. Diffuse venting of low-temperature hydrothermal fluids was a widespread phenomenon on the Palaeoarchaean seafloor, making it an ideal habitat for hyperthermophiles and the possible birthplace of life during earlier times.

Keywords

Barberton greenstone belt Archaean Hydrothermal activity Seafloor alteration Silicification 

1 Introduction

The Barberton greenstone belt contains a well preserved Palaeoarchaean volcano-sedimentary succession that provides abundant evidence for the interaction of submarine volcanic rocks and seafloor sediments with the hydrosphere and ­biosphere. Petrological and geochemical studies have provided evidence for the existence, as early as 3.5 Ga, of a rich microbial ecosystem. This is preserved as filamentous and spherical structures of carbonaceous matter and microbial mats in cherts (Walsh 1992; Westall et al. 2001, 2006; Tice and Lowe 2006; Glikson et al. 2008) as well as possible bioalteration features in pillow lavas and hyaloclastites (Furnes et al. 2004; Banerjee et al. 2006). Many of these traces of life are associated with rocks that show evidence for hydrothermal activity on the Archaean ocean floor. Submarine hydrothermal systems, driven by circulation of seawater through volcanic rocks, constitute a possible site where life emerged and evolved. This paper provides a summary of features preserved in the Barberton greenstone belt that may be directly related to the presence of hydrothermal systems. The importance of these systems as a habitat of ancient life is briefly evaluated.

2 Geological Setting

The Barberton greenstone belt (Fig. 1) is one of the key belts for greenstone studies and represents a type locality of Palaeoarchaean supracrustal sequences. The Barberton greenstone belt consists of a NE–SW striking succession of supracrustal rocks, termed the Swaziland Supergroup, which ranges in age from 3.54 to 3.22 Ga (Viljoen and Viljoen 1969; de Ronde and de Wit 1994; Lowe and Byerly 2007a). The belt has a strike length of ca. 130 km, width of 10–35 km, and an approximate depth of 4–5 km, and is surrounded by granitoid domes and intrusive sheets, ranging in age from 3.51 to 3.10 Ga. The Swaziland Supergroup is subdivided into three stratigraphic units, the Onverwacht, Fig Tree and Moodies Groups (Fig. 2). These units have been tightly folded into a number of synclines that are separated by anticlines or shear zones. The metamorphic grade is greenschist facies, but locally reaches amphibolite facies close to the contact with the surrounding granitoid domes that intruded during several ­episodes of magmatism at ca. 3.45, 3.22, and 3.1 Ga (Kamo and Davis 1994).
Fig. 1

Geological map of the southwestern part of the Barberton greenstone belt (Modified from Hofmann and Bolhar 2007). TTG, tonalite-trondhjemite-granodiorite

Fig. 2

Stratigraphic sections of the Swaziland Supergroup from the western limb of the Onverwacht Anticline (Fig. 1) and from areas north of the Inyoka Fault, as exemplified by the stratigraphy of the Stolzburg Syncline (Modified from Hofmann 2005)

2.1 Onverwacht Group

The Onverwacht Group formed between 3.54 and 3.30 Ga (Kröner et al. 1996; Byerly et al. 1996) and consists of komatiites, komatiitic basalts, basalts and minor dacites. Sedimentary rocks make up less than 10% of the succession and formed in a deep to shallow marine environment. South of the Inyoka Fault, the Onverwacht Group has been subdivided into six formations, the Sandspruit, Theespruit, Komati, Hooggenoeg, Kromberg and Mendon Formations (Viljoen and Viljoen 1969; Lowe and Byerly 2007a). The Sandspruit and Theespruit Formations are separated from the Komati Formation by a fault zone, but the younger units are stratigraphically conformable. The Onverwacht Group is best developed and least deformed in the southwestern part of the belt northeast of Tjakastad (Fig. 1). Metamorphic grade is mainly greenschist facies, but locally reaches amphibolite facies, in particular in the Sandspruit and Theespruit Formations. Onverwacht Group rocks north of the Inyoka Fault have been grouped together as the Weltevreden Formation (Lowe and Byerly 2007a).

2.2 Fig Tree Group

The 3.26–3.23 Ga Fig Tree Group consists of a several kilometres thick ­siliciclastic-volcaniclastic sedimentary sequence that is capped by felsic volcanic rocks (Condie et al. 1970; Heinrichs 1980; Hofmann 2005; Lowe and Byerly 2007a). South of the Inyoka Fault, a variety of lithofacies are present that formed in deep- to shallow-water to alluvial environments (southern facies). The Loenen Formation is a sequence of sandstones and shales that contain abundant felsic detritus­. It is overlain by ferruginous shales and jaspilitic banded iron formations of the Ngwenya Formation. The Mapepe Formation consists of shale, tuffaceous shale, laminated felsic tuff, chert-clast conglomerates and, locally, baryte beds. Zircon dates from the Mapepe Formation vary from 3.26 to 3.23 Ma (Kröner et al. 1991; Byerly et al. 1996). The ca. 3.25 Ga Auber Villiers Formation is approximately 1.5 km thick and consists of dacitic, plagioclase-phyric volcaniclastic rocks and terrigeneous chert-clast conglomerates and sandstones.

In the northern part of the belt, the Fig Tree Group is mainly characterized by ­turbiditic sandstones and shales of the Ulundi, Sheba and Belvue Road Formations that formed in a relatively deep-water environment (northern facies). Kohler and Anhaeusser (2002) defined the Bien Venue Formation, which consists of quartz-­muscovite schist derived from dacitic to rhyodacitic volcaniclastic protoliths dated at ca. 3.25 Ma. Subordinate rock types include banded chert, phyllite, and biotite-plagioclase and chlorite schists, derived, in turn, from dacitic and basaltic precursors. Massive sulphide and baryte deposits are also present. The Schoongezicht Formation overlies the Belvue Road Formation in the central and northwestern part of the greenstone belt. It consists of plagioclase-rich turbidites intercalated with shale at the base and cross-bedded ­volcaniclastic sandstones and dacite clast conglomerates at the top. Intercalated felsic volcanic rocks have been dated at 3,226 ± 1 Ma (Kamo and Davis 1994).

2.3 Moodies Group

The Moodies Group was deposited close to 3.227 Ga, but not later than 3.11 Ga (Kamo and Davis 1994). It consists of shallow-marine to fluvial sandstone and conglomerate with minor shale and banded iron-formation. The Moodies Group is approximately 3 km thick and consists of several fining-upward sequences ranging from conglomerate or pebbly quartzose sandstone at the base, to a thick sandstone unit, to capping siltstone and shale (Anhaeusser 1976).

3 Hydrothermal Systems in the Onverwacht Group

The Onverwacht Group is a thick pile of submarine lava flows that erupted intermittently over a time period of ca. 250 Ma. Breaks in volcanic activity resulted in the deposition of thin interflow sedimentary units characterized by predominantly ­tuffaceous volcaniclastic material. At the same time, hydrothermal processes were operating on, and immediately below, the seafloor, leaving behind extensive zones of alteration that are mainly characterized by the presence of silicified volcanic rocks, bedded cherts, and chert veins.

3.1 Silica Alteration Zones

Silicification of volcanic rocks immediately below sedimentary chert horizons is a common phenomenon of Palaeoarchaean volcano-sedimentary successions in ­general and the Barberton greenstone belt in particular (Kitajima et al. 2001; Terabayashi et al. 2003; Van Kranendonk and Pirajno 2004; Van Kranendonk 2006; Hofmann and Wilson 2007; Hofmann and Harris 2008). Because of the high silica content of the alteration zones, early workers in the Barberton greenstone belt regarded some of these zones to consist of felsic volcanic rocks, which were later shown to be silicified mafic and ultramafic rocks (Paris et al. 1985; Lowe and Byerly 1986; Duchac and Hanor 1987). The origin of the silicification has been the subject of debate, with interpretations ranging from weathering of volcanic flow tops (Lowe and Byerly 1986) to various processes of hydrothermal alteration (de Wit et al. 1982; Paris et al. 1985; Duchac and Hanor 1987; Hanor and Duchac 1990; de Wit and Hart 1993).

In order to test competing models for the origin of the alteration zones, Hofmann and Harris (2008) investigated several sections of silicified volcanic rocks underlying bedded carbonaceous cherts that included most stratigraphic units of the Onverwacht Group, ranging from 3.54 to 3.30 Ga in age. These authors observed that each and every chert horizon is underlain by rocks that are invariably silicified, and locally carbonatised, in zones 10–>50 m thick, irrespective of their composition. Silicification affected lithologies that range from ultramafic to felsic in composition and include volcanic, volcaniclastic, and epiclastic rocks. Carbonate alteration is common in mafic to ultramafic rocks, but also in coarse-grained sedimentary rocks (Rouchon et al. 2009). An example of such an alteration zone developed in pillow basalt and capped by a thin chert horizon from the top of the Hooggenoeg Formation is shown in Figs. 3 and 4. In all alteration zones, the intensity of silicification increases stratigraphically upward towards the contact with the chert, showing a gradual increase in SiO2 content from the original igneous value to ∼90%. Enrichment of SiO2 is associated with a depletion of most other major elements, as monitored by element/Al2O3 or ­element/Hf ratios, including Fe2O3, MnO, MgO, and especially CaO and Na2O. K2O, together with Rb, and Ba, are enriched in these zones. In terms of trace ­elements, depletion was observed for some metals (Ni, Co, Cu, Zn), Sr, the heavy rare earth elements (REE) and Y, while Eu became enriched. Several alteration zones were observed to show zoning, with zones of carbonate alteration present 20–30 m below the top.
Fig. 3

Variations of selected geochemical data of silicified basalts (open symbols) with depth below chert H5c of the Hooggenoeg Formation. Analytical data of a bedded chert and a chert vein (filled symbols) are also shown (From Hofmann and Harris 2008)

Fig. 4

Features of the upper Hooggenoeg Formation (H5v). (a) Pavement of unaltered ocelli-bearing pillow basalt ca. 120 m below the capping chert bed. (b) Silicified and carbonatised pillow basalt ca. 20 m below the chert bed. The interpillow space is filled by secondary carbonate. (c) Network of carbonaceous chert veins transecting silicified basalt 2 m below the chert bed. (d) Photomicrograph (plane polarised light; scale bar is 1 mm) of highly silicified basalt shown in c. Large irregular patches consist of microcrystalline quartz and minor carbonate. Former plagioclase laths have been completely replaced by an intergrowth of quartz and phyllosilicates

Hofmann and Harris (2008) observed that δ18O values of silicified volcanic rocks show a positive linear relationship with silica content and plot between the mantle value (ca. 5.7‰) and a δ18O value of ca. 18‰ extrapolated for a composition of 100% SiO2. A Si isotope study was reported by Abraham et al. (2007) from the same alteration zone shown in Fig. 3. Unsilicified basalts display negative δ30Si values (−0.15‰ to −0.40‰), similar to average mafic magmatic rocks, whereas δ30Si for silicified basalts shifts towards more positive values (up to 0.8‰) with increasing grade of silicification, with overlying chert displaying the highest Si isotope values (up to 1.1‰).

Hofmann and Harris (2008) concluded that the element depletion–enrichment patterns and oxygen isotope data indicate low-temperature (ca. 100–150°C) hydrothermal processes for the origin of the alteration zones. Diffuse venting over large tracts of ocean floor was envisaged to have given rise to silicification of the volcanic­ rocks and of seafloor sediments to form cherts. The Si isotope variation may be attributed to simple mixing of a mafic volcanic and seawater endmember composition, both of which had values similar to the modern day. Apparent zonations of alteration, with silicification at the top and carbonate alteration at depth may be a result of differing fluid chemistries with depth. In the deeper part of the alteration zones, hydrothermal fluids may have been more alkaline, becoming more acidic towards the contact with the likely acidic Archaean ocean due to mixing with seawater.

3.2 Bedded Cherts

Bedded chert horizons are typically 1–20 m thick and were deposited on the seafloor in between phases of extrusive submarine volcanic activity. With the exception of the Komati Formation, bedded cherts are common throughout the Onverwacht Group and the lower part of the Fig Tree Group, but are rare to absent in the upper Fig Tree and Moodies Groups. They typically consist of microcrystalline quartz with minor amounts of phyllosilicates, Fe-Mg-Ca-carbonates, Fe-Mn and Ti-oxide minerals, disseminated sulphides and carbonaceous matter. Cherts are highly siliceous rocks (SiO2 commonly >95 wt%) that contain clastic or tuffaceous material of ultramafic to felsic composition replaced by microquartz. Despite the fact that addition of silica has resulted in dilution of most other elements, they have relatively high alkali contents and are enriched in many metals, particularly in Co, Ni, and Cu. The δ18O values of cherts are somewhat higher than for underlying silicified volcanic rocks and are generally in the range of 15–20‰. The highest δ18O value yet obtained for cherts of the Barberton greenstone belt is 22.1‰ (Knauth and Lowe 2003).

Bedded chert horizons consist of a variety of silicified sediments (Lowe 1999; Lowe and Fisher Worrel 1999; Hofmann 2005; Hofmann et al. 2006; Tice and Lowe 2006; Hofmann and Bolhar 2007). Silicified volcaniclastic sediments are common and include silicified beds of komatiitic ash and accretionary lapilli that are typically represented by light green to greenish grey cherts, the colouration due to finely dispersed chlorite and Cr-rich sericitic mica (Fig. 5a, b). Laminated cherts of various shades of grey to black represent mixtures of volcaniclastic material and carbonaceous matter. Carbonaceous cherts include thinly bedded heterolithic rocks (black and white banded cherts), and laminated and massive black chert varieties. Black and white banded cherts are made up of layers of carbonaceous chert and white-weathering translucent chert (Fig. 5c, d). Silicified orthochemical deposits include rare silicified evaporites (Fig. 5c), possibly primary sea-floor silica deposits in the form of translucent cherts, and banded iron formation and associated ferruginous rocks. Deposition took place predominantly in a low-energy, sub-wave base setting with episodic, high-energy current events during which coarse volcaniclastic­ material and, in rare cases, meteorite impact ejecta were deposited.
Fig. 5

Features of bedded chert. (a) Planar and wavy bedded green chert, consisting predominantly of komatiitic ash (Kromberg Formation). (b) Photomicrograph (plane polarised light; scale bar is 1 mm) of silicified volcaniclastic sandstone with accretionary lapilli and glass shards (Middle Marker). (c) Thinly bedded black-and-white chert transected by “evaporite” pseudomorphs (Buck Reef Chert). (d) Photomicrograph (plane polarised light; scale bar is 1 mm) of black-and-white chert showing lamina rich in carbonaceous grains (Buck Reef Chert)

It is generally agreed that silicification took place shortly after deposition of the precursor sediment on the seafloor (Lowe 1999), but the mechanism is a matter of debate. Many researchers regard the silicification as a direct result of hydrothermal processes on the seafloor (de Wit et al. 1982; Paris et al. 1985; Duchac and Hanor 1987). Hofmann and Harris (2008) suggested silicification being the result of diffuse flow of hydrothermal fluids that were also responsible for the silicification of the underlying rocks. Others refute the evidence for hydrothermal activity, and ascribe the formation of these rocks to marine processes that acted in a hot Archean ocean (e.g. Knauth and Lowe 2003; Tice and Lowe 2006). Cherts of the Barberton greenstone belt have much lower δ18O values relative to Phanerozoic cherts (difference of ≥10‰ on average; Perry 1967; Knauth and Lowe 2003). These signatures have been interpreted in various ways, but mainly being the result of one or more of the following: (1) lower δ18O values of Archaean seawater (Perry 1967; Kasting et al. 2006; Jaffres et al. 2007), (2) higher temperatures of the Archaean ocean (ca. 70°C; Perry 1967; Knauth and Lowe 2003), (3) temperatures of hydrothermal fluid–seawater mixtures and not the temperature of the Archaean ocean (Hofmann 2005; Hofmann and Harris 2008).

If seawater had δ18O values as low as −13‰ at 3.4 Ga (Jaffres et al. 2007), this would imply that the alteration took place at temperatures of around 20°C. Such low temperatures are inconsistent with the geochemical and mineralogical changes in the alteration zones that need elevated temperatures to form. The model of a hot Archaean ocean, potentially resulting in downward infiltration of hot fluids from the seafloor and leading to top down alteration of volcanic rocks, is inconsistent with a number of observations that indicate the presence of upwelling hydrothermal fluids, such as depletion of metals (Co, Ni and Cu) in subsurface volcanic rocks and their relative enrichment in overlying chert (Hofmann 2005; Hofmann and Harris 2008), and the development of chert veins, regarded as hydraulic fractures beneath chert beds necessitating high fluid pressures at times (Hofmann and Bolhar 2007).

3.3 Chert Veins

Veins of carbonaceous chert occur immediately beneath sedimentary chert horizons, transecting underlying silicified volcanic rocks. They are 0.1–3 m wide and cross-cut the host rock both perpendicular and sub-parallel to stratification (Fig. 6). A detailed field study of chert veins revealed the importance of syndepositional hydrothermal activity for their origin (Hofmann and Bolhar 2007). The geometry of the chert-filled veins (e.g. the presence of stratiform veins and the branching of vertical veins into stratiform ones), multiple vein fillings, in situ brecciation of earlier generations of vein fillings, and brecciation of the host rock suggest that they represent hydraulic fractures that were initiated by the forceful intrusion of overpressured fluids.
Fig. 6

Features of chert veins. (a) Two generations of chert veins cutting across highly silicified komatiite of the Mendon Formation. The vein filled with botryoidal chert is sub-parallel to stratification in the host rock. (b) Anastomosing chert veins in highly silicified sandstone below the Buck Reef Chert. The veins are subparallel to bedding. (c) Chert vein (margins not shown) containing fragments of silicified sedimentary host rock in a black chert matrix. Note strong brecciation of the fragments due to hydraulic fracturing. (d) Chert vein (note margin to silicified sediment in lower left) consisting of several generations of chert forming layers subparallel to the vein margins

Vein and bedded cherts of a single stratigraphic section are petrographically, isotopically as well as geochemically very similar. They are all characterized by sand-sized carbonaceous grains of apparent detrital origin. The differences between several pairs of vein chert and directly overlying bedded chert is <2‰ for C isotopes and <1.1‰ for O isotopes. The same applies to REE patterns and trace element ratios (Hofmann and Bolhar 2007; Hofmann and Harris 2008). All these observations suggest a genetic link between the veins and the sedimentary horizons.

While some workers regard chert veins as feeder channels for silicified sedimentary deposits in the Pilbara craton (e.g. Brasier et al. 2002; Van Kranendonk 2006), in the Barberton belt they have been regarded to represent fractures that were filled with sediment from above (Lowe and Byerly 1986; Hofmann and Bolhar 2007). In the model presented by Hofmann and Bolhar (2007), diffuse upflow of hydrothermal fluids in a shallow, but laterally very extensive hydrothermal system that was ­initially open to the ocean is envisaged during deposition of the sedimentary chert precursor. This system became temporarily closed during progressive sedimentation and hydrothermal silicification of the seafloor sediments that resulted in the bedded chert horizons. The cherts acted as a poorly permeable barrier for ascending fluids, resulting in the build-up of fluid overpressure and hydraulic fracturing of the volcanic­ rocks and the cap rocks at times. The hydraulic fractures were filled with sedimentary material from as yet unsilicified sediments overlying the chert for some time before the hydraulic system was sealed again and the next cycle of pressure build-up and fracturing took place. This eventually resulted in multiple injections of fluidized sedimentary material and hydrothermal fluids in cross-cutting veins, partly brecciating the host rock and incorporating host rock fragments into the veins.

3.4 Ironstone Pods: Archaean Hydrothermal Systems or Products of Recent Weathering

De Wit et al. (1982) and de Ronde et al. (1994) described discontinuous occurrences of lens-shaped “pods” of massive, specular haematite and goethite, up to 200 m in length, from the contact between Onverwacht komatiites and overlying sedimentary rocks at the base of the Fig Tree Group. One of these pods, which mainly consists of haematite (Fig. 7), was interpreted as a Fe-oxide-rich mound and ­discharge vent of 90–150°C hot hydrothermal fluids of Archaean age (de Ronde et al. 1994). Subsequent studies of the pod included mass spectroscopy that revealed a large variety of organic compounds (de Ronde and Ebbesen 1996), ­suggesting abundant organic productivity associated with the processes that led to ironstone pod formation. Fluid inclusion leachates obtained from quartz samples associated with the pod (Channer et al. 1997) yielded Br/Cl and I/Cl ratios higher than seawater, which was interpreted as a result of a secular change in ocean halide ratios through time. Additional fluid inclusion data on seawater and hydrothermal fluid endmember composition were reported by de Ronde et al. (1997).
Fig. 7

Geological map of the ironstone pod locality on Mendon Farm (Modified from Lowe and Byerly 2007b). The dip of the strata is subvertical. See Fig. 1 for locality

Lowe and Byerly (2003) studied Fe-oxide-rich deposits at different localities in the Barberton greenstone belt, including the ironstone pod described by de Ronde et al. (1994). They concluded that the ironstones were deposited on and directly below the modern ground surface by active groundwater and spring systems during relatively recent (Cenozoic) times. This reinterpretation resulted in some debate in the literature (de Ronde et al. 2004; Lowe and Byerly 2004), which has since cooled down with additional work published by Lowe and co-workers (Hren et al. 2006; Lowe and Byerly 2007b).

Debate has centred on the pod originally described by de Ronde et al. (1994) to be of Archaean age (Fig. 7). De Ronde et al. (2004) cited the following main arguments for the ancient age of this pod: (1) The pod is regarded to be part of the Archaean stratigraphy; it is cut by a porphyry intrusion and clasts derived from ironstone pods are present in overlying sedimentary units. (2) It contains rounded chert boulders interpreted as talus derived from scarps in the vicinity of the mound. (3) Quartz-haematite-goethite stockwork veins are present in the pod. (4) Its haematitic composition is unlike low-T spring deposits. Lowe and Byerly (2004, 2007b) argued for a recent origin of the ironstone pod, through partial replacement of Archaean chert host rocks by Fe-oxide-rich material. These authors based their interpretation on the following main observations: (1) No evidence was found that the pods are part of the stratigraphy, as they are not cut by Archaean porphyry or provided detritus upon erosion during Archaean times as claimed by de Ronde et al. (1994). Instead, they locally observed layering in the pod that dips parallel to the present slope. (2) Internal chert blocks are not rotated and represent in situ remnants not affected by replacement of chert by haematite. (3) The pods do not show evidence for deformation that characterizes the surrounding rocks and contain common open cavities, which would not have been preserved up to the present. Lowe and Byerly (2004, 2007b) further argue that the quartz stockworks represent remnants of Archaean quartz veins that were only partially replaced by ironstone.

Our own observations of ironstone pods in the Barberton greenstone belt tend to agree with a relatively recent origin of these features. Nevertheless, a large part of the work by de Ronde and co-workers was done on fluid inclusions of quartz veins that are of undoubted Archaean age. These quartz veins are commonly associated with chert veins interpreted to have formed during seafloor hydrothermal activity (Hofmann and Bolhar 2007) and may thus still provide evidence for hydrothermal fluid composition and temperature of an Archaean hydrothermal system. This is in contrast to the claim by Lowe and Byerly (2007b) that these features are a result of late-stage hydrothermal alteration of the belt (as late as 2.7 Ga). If the haematite-rich ironstone pod is a result of relatively recent spring activity, two questions still remain. One issue is the predominance of haematite in the pod, which generally only precipitates from relatively high-temperature fluids. This led Lowe and Byerly (2007b) to suggest a hot spring origin for the ironstone pod. No hot spring has so far been reported from the greenstone succession of the Barberton belt. Another major issue is the source of the iron. Lowe and Byerly (2007b) suggest that the iron was likely derived from the weathering of siderite-rich sedimentary units at depth, such as banded iron-formation. This is far from proven, and will require further investigations.

3.5 Heat Source for Onverwacht Hydrothermal Activity

Hydrothermal systems consist of two essential components, a heat source and a fluid phase (Pirajno 1992). While the fluid phase is most readily attributed to seawater circulating through the volcanic pile of the Onverwacht Group, the heat source is less easy to define. Most hydrothermal systems operating on the seafloor today are driven by heat generated from shallow-level magma chambers, especially those developed close to the mid-ocean ridges (Alt 1995). Other heat sources may include shallow-level magma chambers, intrusive igneous bodies or volcanic complexes associated with volcanic islands, oceanic plateaus and island arcs. Crustal sections close to these heat sources may experience extensive alteration related to the interaction with actively circulating >350°C hot hydrothermal fluids, which may locally discharge on the seafloor through black smokers. Further away from these zones fluid temperatures will be much lower, and hydrothermal fluid discharge is much more diffuse (Alt 1995; German and von Damm 2003).

Although the Onverwacht Group was deposited in a submarine setting and formed mainly as a result of subaquatic deposition of mafic-ultramafic lavas, there is no evidence that the volcanic rocks formed close to an Archaean mid-oceanic ridge. The lavas accumulated during an extensive period of time far longer than the average life span of modern oceanic crust. The alteration zones are laterally ­extensive and tabular, bearing no resemblance to the heavily faulted spreading ­centres along modern mid-ocean ridges. Instead, the Onverwacht Group is characterized by laterally continuous submarine lava flows more akin to large submarine shield volcanoes or oceanic plateaus (Cloete 1999; Hofmann and Harris 2008). Focused hydrothermal upflow zones in the form of epidosites and black smoker-type massive sulphides are lacking throughout the Onverwacht Group, suggesting diffuse upflow of hydrothermal fluids through the volcanic rocks.

Although there are several sub-volcanic intrusions in the Barberton sequence, ranging from ultramafic to felsic in composition, that could have been the heat source for the hydrothermal activity, there are many more examples of stratiform alteration zones in the Barberton greenstone belt that cannot be directly linked to such intrusions. Each volcanic sequence capped by chert is altered at the top, indicating that the alteration took place during a time interval of relative volcanic ­quiescence, when there was enough time for the deposition of finely laminated sediments before the onset of the next volcanic episode. Hofmann and Harris (2008) attributed the hydrothermal activity to heat derived from cooling of volcanic rocks and a high regional heat flow in an oceanic plateau-like setting in the vicinity of a long-lived source of hot, upwelling mantle that resulted in the establishment of shallow subseafloor convection cells. Shallow-level igneous intrusions may have provided additional heat, but they were not the main factor for elevated heat flow. Extensive silicification was the result of diffuse venting of hydrothermal fluids over broad areas of the ancient ocean floor, likely aided by seawater chemistry close to silica saturation and thus more conducive for widespread silicification (Hofmann and Harris 2008).

3.6 Buck Reef Chert: Deposition During Hydrothermal or Normal Marine Conditions?

While most of the Onverwacht cherts can be readily associated with active hydrothermal activity during their formation, this may not necessarily be the case for the Buck Reef Chert (BRC) for which a hydrothermal influence has been most strongly debated. The BRC is an unusually thick (up to 350 m) sequence of predominantly black-and-white banded cherts at the contact between the Hooggenoeg and Kromberg Formations along the western limb of the Onverwacht Anticline. It overlies a shallow intrusive to extrusive sequence (member H6 of Lowe and Byerly 2007a) of dacitic volcanic rocks, up to 2 km thick, dated at 3.45 Ga (de Wit et al. 1987; Armstrong et al. 1990; De Vries et al. 2006) and an epiclastic sedimentary unit of dacite-derived conglomerates and sandstones. The BRC has been subdivided into three facies (Fig. 8; Lowe and Fisher Worrel 1999; Tice and Lowe 2004, 2006): a basal evaporitic facies, containing silicified sandstones, chert and pseudomorphs after nahcolite; a lower division of platform facies black-and-white banded chert; a basin facies of banded ferruginous chert; and an upper division of black-and-white banded chert. A variety of fossil-like microstructures have been reported from the BRC (Walsh 1992; Tice and Lowe 2006). A felsic tuff at the base of the BRC has been dated at 3,416 ± 5 Ma (Kröner et al. 1991). Along the east limb of the Onverwacht Anticline, H6 is represented by a fining-upward sedimentary sequence of dacite-clast conglomerate, turbiditic sandstone and shale (Rouchon et al. 2009). These rocks are overlain by a sequence of ultramafic lapillistone and pillow basalt that are intercalated with several chert horizons that have been correlated with the BRC (Lowe and Byerly 2007a).
Fig. 8

Simplified section of the Buck Reef Chert (Modified from Hofmann and Bolhar 2007)

Debate has centred on the interpretation of the palaeoenvironment. De Vries (2004) and de Vries et al. (2006) consider the BRC to be the uppermost silicified part of the Buck Ridge volcano-sedimentary complex (BR-vsc), a series of arc-like basaltic and felsic volcanic rocks, capped by pervasively silicified sedimentary rocks. This complex is believed to be characterized by syndepositional normal faulting­, which widely controlled felsic magma upwelling and sedimentation. De Vries et al. (2006) advocate hydrothermal venting as a common growth-fault related ­phenomenon directly related to the depositional history of the unit, giving rise to features such as chert veins, hydraulic breccia bodies and quartz-filled cavities.

In contrast, Lowe and Byerly (2007a) regard the BRC to have been deposited on a subsiding volcanic platform representing the eroded felsic volcanic edifice of H6 that was emplaced well before deposition of the BRC. Syn-depositional normal faulting was observed to have taken place during deposition of the lower half of the BRC (Lowe and Fisher Worrel 1999). Tice and Lowe (2004, 2006) suggested that the BRC consists of normal marine sediments that were silicified through interaction with hot marine waters supersaturated with amorphous silica and related abundant carbonaceous matter to microbial mats in a non-hydrothermal environment.

Hofmann and Bolhar (2007) and Hofmann and Harris (2008) suggested hydrothermal activity during deposition of at least the lower part of the BRC (the evaporitic facies). Dacite and dacite-derived clastic sedimentary rocks below and interbedded with the evaporitic cherts are highly silicified and show the same chemical variations with depth as observed in all the other alteration zone below chert horizons. Multiple generations of cross-cutting veins of chert are common in the alteration zone. These are both stratiform as well as at a high angle to bedding, the latter extending several 100 m below the BRC.

If the evaporitic facies was indeed affected by syn-depositional hydrothermal activity, then the chert clast breccias commonly observed in this zone may be attributed to subsurface dissolution of evaporites by hydrothermal fluids rather than to evaporite dissolution in the phreatic zone (Lowe and Fisher Worrel 1999). A study of fluid inclusions in quartz veins and quartz-filled cavities by de Vries and Touret (2007) from the BRC gave conflicting results, although it pointed towards relatively low pressure of ∼100 bars for quartz precipitation, in support of a shallow water environment and synsedimentary hydrothermal activity.

In addition, the evaporites may not be precipitates from purely marine waters. Elongate crystal pseudomorphs interpreted to represent pseudomorphs after nahcolite (NaHCO3; Lowe and Fisher Worrel 1999) are associated with finely laminated and rippled cherts and grew concurrently with deposition of the host sediment. The nahcolite crystals have been interpreted as primary evaporitic minerals that grew in a hypersaline marine coastal environment with high bicarbonate and sodium contents. If the pseudomorphs are indeed former nahcolite and if the BRC was deposited during hydrothermal activity, there is a possibility that elevated concentrations of bicarbonate and/or sodium are derived from hydrothermal fluids. Alteration in the underlying dacite and dacite-derived sediments has resulted in the almost complete removal of Na from at least the uppermost 100 m of rock succession (<< 0.1 wt% Na2O, Hofmann and Harris 2008), as the Na content of the unaltered magmatic rocks was likely similar in composition to the genetically related Theespruit Pluton (ca. 7 wt% Na2O, Kleinhanns et al. 2003). It is thus possible that Na was derived from syn-depositional hydrothermal fluids. It needs to be emphasized that nahcolite has only been observed at the base of the BRC where it overlies altered felsic igneous rocks.

Chert veins were not observed cutting across the basal evaporitic facies. This may indicate that hydrothermal activity ceased before deposition of the overlying cherts or that the evaporitic facies acted as a barrier for ascending hydrothermal fluids. Cessation of hydrothermal activity can be attributed to reduced convection of seawater as the basal chert beds would have prevented continuous recharge of the hydrothermal system. If this was the case, deposition of cherts higher up in the BRC may have been less affected by hydrothermal activity. This may explain the abundance of soft-sediment deformation features in black-and-white banded cherts above the evaporitic facies (Tice and Lowe 2004), suggesting silicification acted more slowly. δ18O values for the BRC are amongst the highest values obtained by Knauth and Lowe (2003) from the Barberton greenstone belt and range from 15.7‰ to 21.4‰, which is in line with subdued hydrothermal activity.

If BRC deposition was indeed influenced by hydrothermal activity, then several potential heat sources need to be considered. The sill-like dacite intrusion underlying the BRC could have been the heat source, but this is unlikely for several reasons. (1) The dacite is erosively overlain by dacite-derived sandstones that are silicified at the top and may have had sufficiently cooled down prior to subaerial exposure and erosion not to have been responsible for the hydrothermal alteration as seen at the top. (2) Available age data suggest that the dacite formed long before the BRC was deposited (ca. 3.45 versus 3.42 Ga, see above). The heat source could thus have been an elevated regional geotherm or renewed volcanic activity that gave rise to the overlying volcanic rocks. The BRC is overlain by undated mafic/ultramafic lapillistones and tuffs of the Kromberg Formation, it is cut by abundant ultramafic dykes and sills that may be genetically related to the overlying unit, it contains rare layers of ultramafic accretionary lapilli, and has been regarded as interleaved laterally with mafic and ultramafic rocks (Ransom et al. 1999; Lowe and Byerly 2007a). All these observations suggest that mafic/ultramafic magmatism was ongoing during deposition of the middle to upper part of the BRC. It is very well possible that extensional deformation of the sequence was related to this magmatic event.

With the evidence for syndepositional volcanism and potential extensional deformation, it is hard to evoke the absence of elevated heatflow and hydrothermal activity during BRC deposition. This leads to the evaluation of a possible hydrothermal control of the deposition of ferruginous cherts. Ferruginous cherts in the BRC are composed mainly of microquartz and goethite, although siderite is regarded to be the main iron-bearing mineral at depth (Tice and Lowe 2004, 2006). Deposition was regarded to have taken place in a low-energy environment, far below storm wave base, from a stratified Archaean ocean consisting of a shallow, CO2-dominated layer and a deep, iron-rich layer. No hydrothermal influence was invoked, partly based on the lateral continuity of the different sedimentary facies, including the ferruginous basin facies (Tice and Lowe 2004, 2006). Recent mapping has revealed that the ferruginous facies may not be as continuous as previously thought (Fig. 9). In the map area, the unit of ferruginous cherts is discontinuous and shows changes in thickness along strike, with the maximum thickness being present in a lenticular area where rocks are strongly impregnated by goethite. The goethite forms layers that locally cross-cut bedding in the cherts and strongly resemble ironstone pod-like features described by Lowe and Byerly (2007b) from the BRC. Layered goethite was interpreted by these authors to represent precipitates from upward-migrating groundwater, with the Fe derived from dissolution of siderite bands in the ferruginous chert at depth. While the goethite is clearly secondary and likely relatively young, this does not mean that it could not have been derived from weathering of a lenticular body of unusual ferruginous rock at depth, possibly including sulphide minerals and derived from seafloor hydrothermal activity. These speculations will require subsurface investigations by drilling.
Fig. 9

Geological map of a portion of the Buck Reef Chert. The strata are younging to the NE. See Fig. 1 for locality

4 Hydrothermal Systems of the Fig Tree Group

The 3.26–3.23 Ga Fig Tree Group is a volcano-sedimentary sequence that is dominated by relatively fine-grained siliciclastic sediments and felsic volcanic and volcaniclastic rocks. Carbonaceous cherts are uncommon, but many of the shale horizons are rich in carbonaceous matter (Reimer 1975), although reports of microfossil finds have been met with scepticism (Schopf and Walter 1983). Evidence for hydrothermal activity during Fig Tree deposition is mainly preserved in the presence of baryte horizons, a massive sulphide deposit and localized hydrothermal alteration of the sedimentary rocks.

4.1 Baryte

Layers of baryte are present at a particular stratigraphic level within the lower Mapepe Formation in the central part of the greenstone belt and are well developed in the Baryte Valley Syncline area (Fig. 10) where they can be traced for several kilometres. The baryte is considered to be sedimentary in origin and has been studied­ by Heinrichs and Reimer (1977) and Reimer (1980). Baryte occurs in layers 13 cm thick on average (Fig. 11a) that are intercalated with green chert, dolomitic chert, chert pebble conglomerate and highly silicified sandstone and grit, all of which were deposited under a range of shallow-marine conditions (Lowe and Nocita 1999). Baryte occurs in three different varieties (Heinrichs and Reimer 1977; Reimer 1980). (1) Detrital baryte beds, consisting of slightly rounded baryte grains admixed with minor amounts of volcanic quartz, muscovite, pyrite, zircon and chromite, and showing cross-lamination (Fig. 11b). (2) Authigenic baryte crystals that form cauliflower-like structures locally and that grew at the sediment-water interface. (3) Authigenic, coarse baryte blades that replaced detrital baryte and grew as single crystals or bundles after deposition. Bao et al. (2007) reported ­multiple S and O isotope data from the baryte. No 17O anomaly was observed, while the average δ18O value is 10.6‰, close to that of the modern seawater sulphate value. Values of δ34S range narrowly from +3.7 to +6.2‰ while Δ33S values are negative, ranging from −0.69‰ to −0.31‰. Similar values were reported by Farquhar et al. (2000) from ca. 3.5 Ga baryte of the Warrawoona Group.
Fig. 10

Geological map of the Baryte Valley (After Pearton 1986). BVS: Baryte Valley Syncline. See Fig. 1 for locality

Fig. 11

(a) Outcrop of bedded detrital barite containing abundant blades of secondary barite oriented perpendicular to bedding. (b) Photomicrograph (plane polarised light; scale bar is 1 mm) of detrital barite admixed with minor detrital quartz and chromite

Detrital baryte was interpreted by Heinrichs and Reimer (1977) and Reimer (1980) to represent reworked baryte deposits of hydrothermal-exhalative origin. Multiple sulphur isotope values, in particular negative Δ33S values, suggest that the sulphate was derived from seawater, and ultimately from photolysis reactions in the atmosphere (Bao et al. 2007). Hydrothermal activity may have been roughly contemporaneous with baryte reworking and deposition. On the western limb of the Baryte Valley Syncline (Fig. 10), baryte beds directly overlie a succession of highly silicified shales, now chert, that are transected by numerous chert veins. Baryte values are high (>1,000 ppm) in several samples of these rocks (Hofmann 2005; Hofmann and Bolhar 2007). At a disused baryte mine (Main Workings, Fig. 10), chert and baryte veins and baryte-cemented breccias have been observed in cherts underlying the baryte beds (Pearton 1986). This suggests a genetic link between chert veins and reworked baryte, i.e. syndepositional hydrothermal activity. Horizons consisting of interbedded chert and baryte also occur in the Pilbara craton, where they are associated with swarms of zoned veins of chert and baryte, indicative of a hydrothermal origin (Nijman et al. 1999; Van Kranendonk and Pirajno 2004; Van Kranendonk 2006). The timing of baryte deposition in the Fig Tree Group is unclear, as dating of the Mapepe Formation in different structural domains has yielded conflicting results (Kröner et al. 1991; Byerly et al. 1996).

4.2 Hydrothermally Altered Shales and Sandstones

Shales and greywackes of the southern facies of the Fig Tree Group in the central part of the Barberton greenstone belt have been affected by strong alteration, in contrast to greywackes of the northern facies (Hofmann 2005). This alteration led to similar compositional changes as observed in the silica alteration zones of the Onverwacht Group: K, Rb and Ba were added, whereas Mg, Ca, Na and Sr were lost (Fig. 12). Palaeoweathering was excluded as a possible source for the alteration due to the restricted occurrence of altered rocks in the central part of the belt and low Al2O3 contents of Fig Tree greywackes and shales, suggestive of a palaeoclimate dominated by mechanical, rather than chemical, weathering processes.
Fig. 12

Distribution of selected major and trace elements of average Fig Tree Group sedimentary rock (greywacke and shale) of the southern facies normalised to average greywacke of the northern facies. Note the addition of K2O, Rb, and Ba, and the depletion of MgO, CaO, Na2O, Ni and Sr in rocks of the southern facies, which is interpreted to be a result of hydrothermal alteration

Because sedimentary baryte deposits are restricted to the central part of the greenstone belt where Fig Tree Group sedimentary rocks are highly altered, it is possible that this area was a geothermally active zone during Fig Tree times. Ba-rich hydrothermal fluids may have ascended mainly during lower Mapepe times and resulted, firstly, in local deposits of hydrothermal baryte that were reworked to form detrital baryte beds and, secondly, in the metasomatic alteration of Fig Tree strata in hydrothermal upflow zones. Syndepositional alteration may be related to the vicinity of the “baryte belt” to a tectonically active zone underlying, and possibly controlling parts of, the Fig Tree sedimentary basin (e.g. Proto-Inyoka zone of Heinrichs and Reimer 1977).

4.3 Bien Venue Massive Sulphide Deposit

The Bien Venue volcanogenic massive sulphide deposit is situated in felsic volcanic rocks of the Bien Venue Formation of the northern Fig Tree Group (Fig. 13; Murphy 1990; Ward 1999; Kohler and Anhaeusser 2002). Host rocks include a variety of tightly folded rhyolitic to rhyodacitic pyroclastic rocks, commonly containing quartz phenocrysts. The host rocks have been affected by hydrothermal alteration (silicification, sericitisation and chloritisation), followed by greenschist facies metamorphism and complex deformation, resulting in the formation of quartz-sericite schists. The ore body consists of 50–70 m wide, stratabound lenses of massive and disseminated sulphides, mainly pyrite, sphalerite and chalcopyrite, with minor galena, ­tennantite and native silver. The footwall ore is dominated by pyrite and chalcopyrite­. Towards the hangingwall, sphalerite, galena and baryte become common. This zonation­ of the ore body is similar to that in the much less deformed VMS deposits from the Strelley belt of the Pilbara craton that formed at almost the same time (e.g. Vearncombe et al. 1995). Murphy (1990) suggested that the sulphide ­mineralisation formed from Kuroko-type sea-floor exhalative activity, based on characteristic ­features such as stratabound mineralisation, ore zonations and spatial association with felsic volcanic rocks. While filamentous microfossils have been reported from cherts associated with the VMS deposit of the Strelley belt (Rasmussen 2000), no geobiological studies have so far been reported from the Bien Venue deposit.
Fig. 13

Geological map of the Bien Venue massive sulphide deposit (After Murphy 1990). See Fig. 1 for locality

4.4 Heat Source for Fig Tree Hydrothermal Activity

Felsic volcanic rocks of the Bien Venue Formation have been dated at 3,256 ± 1 Ma (Kohler 2003), which is the likely age of the Bien Venue VMS deposit. This is the same age as reported for the felsic volcanic Auber Villiers Formation (3,256 ± 4 Ma, Kröner et al. 1991; 3,253 ± 2 Ma, Byerly et al. 1996), suggesting a belt-wide event of felsic volcanic activity. A similar age has also been reported from dacitic tuffs at the base of the Mapepe Fm at one locality (3,258 ± 3 Ma, Byerly et al. 1996), which could suggest that at least some of the Mapepe baryte deposits formed at that time. The presence of baryte associated with the Bien Venue deposit would fit the interpretation that the Fig Tree Group both north and south of the Inyoka Fault was affected by a major hydrothermal event at ca. 3.255 Ga associated with felsic volcanic activity. No TTG plutons of this age have so far been identified in the Barberton Greenstone belt, although many of the highly sheared felsic rocks in the Bien Venue Formation may have originally formed as sub-volcanic sills. Instead there are a number of plutons with an age of 3.23 Ga that formed at the same time as the felsic volcanic rocks of the Schoongezicht Formation that forms the uppermost unit of the Fig Tree Group. Most authors (e.g. de Ronde and Kamo 2000; Kohler and Anhaeusser 2002; Lowe and Byerly 2007a) have interpreted the felsic volcanic activity in the Fig Tree Group to be associated with subduction-related magmatism in a continental arc or back-arc setting, environments conducive for hydrothermal activity.

5 Implications of Hydrothermal Activity for Early Life in the Barberton Greenstone Belt

Evidence for hydrothermal activity is widespread in the Barberton greenstone belt. This is not surprising, as the Swaziland Supergroup, with the exception of the Moodies Group, consists largely of volcanic and pyroclastic rocks that were intruded by TTG plutons and ultramafic layered complexes at several stages in the history of the succession. What remains to be answered more fully is how hydrothermal activity did affect the life forms that were present at that time and to what extent hydrothermal systems and sub-seafloor environments were occupied by microbial life. Furthermore, it is important to determine how hydrothermal alteration affected the morphology and composition of remnants of life as preserved in Barberton rocks. Research to answer these questions is still in its infancy.

Carbonaceous matter of likely organic origin (Walsh 1992; Westall et al. 2001; Tice and Lowe 2006; Glikson et al. 2008) is present in bedded cherts and in chert veins of the Onverwacht Group. In which environment this material formed is unresolved. If the carbonaceous matter formed in shallow parts of the ocean and represents background suspension sediments that settled out of the water column, its origin would be unrelated to seafloor hydrothermal activity. It is also possible that organic matter was produced by photo-/chemosynthetic bacteria close to the seafloor. Mat-like accumulations of carbonaceous matter have been described from bedded cherts (Westall et al. 2001; Tice and Lowe 2006). If biogenic and of chemosynthetic origin, low-temperature hydrothermal activity may thus have been conducive to life on the Archaean seafloor.

Alternatively, carbonaceous matter may have formed by subsurface biota that inhabited the hydrothermal systems and were deposited on the seafloor during hydrothermal effusions. Supporting the idea of a subsurface habitat are tubular features described from the rims of pillow basalts from the 3.3 Ga Kromberg Formation and interpreted to represent corrosion features by chemolithoautotrophs that colonized the originally glassy pillow rims (Furnes et al. 2004). If these ­structures are truly biogenic in origin, the microbes must have thrived while hydrothermal fluids were circulating the basaltic subseafloor environment.

Although thermophilic organisms may well have occupied the low-temperature hydrothermal systems at that time (e.g. Ueno et al. 2004), most of the carbonaceous matter present in the chert veins likely originated from the downward flow of ­sediment-laden seawater from the seafloor and was not produced at depth. This is indicated by its ubiquitous presence in subvertical veins extending to the palaeoseafloor, but its absence in stratiform chert veins that are commonly filled with botryoidal chert precipitate. The latter is interpreted to have formed when the hydrothermal system was closed off from the seafloor due to the impermeable chert cap (Hofmann and Bolhar 2007). A possible abiological carbon source includes the reduction of CO2 during serpentinisation of ultramafic rocks at depth (e.g. Horita and Berndt 1999) and redistribution of carbonaceous matter during hydrothermal venting, although these processes have been discounted to have played a major role for the origin of carbonaceous matter in the Barberton greenstone belt (Hofmann and Bolhar 2007). However, remobilisation of carbonaceous matter did take place, as indicated by its presence as secondary material in pore space and in veins (van Zuilen et al. 2007), indicating that morphological interpretations of carbonaceous structures need to be conducted with care.

Shales of the Fig Tree Group are rich in carbonaceous material (0.16–1.34 wt% TOC; Reimer 1975), which occurs as compacted carbonaceous grains. This material has very similar isotopic signatures compared to Onverwacht carbonaceous cherts (δ13C = −29‰ on average; Hofmann and Bolhar 2007), suggesting that most of the carbon must represent material that either formed and was reworked on the seafloor or settled out of the water column. A biogenic origin appears most plausible in the absence of a tenable model for the abiogenic origin of carbonaceous matter in thick and spatially widespread black shale sequences with light carbon isotope signatures comparable to those in younger settings. Carbonaceous matter is both present in hydrothermally altered and unaltered shales and throughout the Fig Tree Group, suggesting no direct link between life and hydrothermal activity, at least during Fig Tree times.

Notes

Acknowledgments

Research in the Barberton greenstone belt was supported by Deutsche Forschungsgemeinschaft (Ho 2507/1–1/2), University of the Witwatersrand Research Committee and National Research Foundation of South Africa (FA2005040400027). Numerous people ­provided access, support, and hospitality, including Johan Eksteen and Property Mokoena (Mpumalanga Parks Board), Colin Wille (Taurus Estate), Jan Maarten and Wilma van Rensburg (Sappi Forests) and Roelf le Roux and Chris Rippon (Barberton Mines).

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Copyright information

© Springer Science+Business Media B.V. 2011

Authors and Affiliations

  1. 1.Palaeoproterozoic Mineralization Research Group, Department of GeologyUniversity of JohannesburgJohannesburgSouth Africa

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