Abstract Although the gold mining industry in the Witwatersrand has passed its maturity, the Witwatersrand basin remains by far the most important known gold-producing province. Most of the gold occurs together with rounded pyrite and uraninite, and in places with bituminous carbon, along degradation surfaces in coarsegrained siliciclastic rocks of the 2.97 to 2.91 Ga West Rand Group, and more importantly, in similar rocks of the 2.89 to 2.71 Ga Central Rand Group. Significant correlation between gold and detrital zircon distribution, as well as sedimentary lithofacies, preservation of detrital micronuggets in places, and absolute age data provide evidence of a detrital origin of the gold, which was derived from the weathered products of 3.1 to 2.9 Ga greenstone belts to the north and west, and transported and concentrated by fluvial selective entrainment processes. Eolian deflation led to further upgrading of the placer deposits. Microstructural, mineral chemical, and isotopic evidence indicate a detrital origin also for rounded compact pyrite and uraninite particles, from which reducing conditions are inferred for the Archean meteoric environment.These conditions prevailed until at least 2.64 Ga, when the last Witwatersrand-type placer deposits were laid down at the base of the Transvaal basin. A series of tectonothermal overprints are recognized in the Witwatersrand strata. These began with a 2.83 Ga deformation event near the western margin of the Witwatersrand basin in response to accretion of an oceanic arc to the western margin of the Kaapvaal craton. Compression followed by continental rifting led to the deposition of the Ventersdorp Supergroup at 2.70 to 2.67 Ga as a consequence of collision between the Kalahari and Zimbabwe cratons and subsequent orogenic collapse of the Limpopo belt. Thermal subsidence and associated deposition of the 2.64 to 2.43 Ga Chuniespoort Group led to low-grade burial metamorphism of the Witwatersrand strata, to basin dewatering, oil migration, and partial gold mobilization within the Witwatersrand basin fill. Burial to similar depths was achieved again during the deposition of the 2.35(?) to 2.25 Ga Pretoria Group after a major erosive phase. Inferred magmatic underplating and intrusion of large amounts of mafic to ultramafic melts into the lower to middle crust during the emplacement of the 2.06 to 2.05 Ga Bushveld Igneous Complex led to a thermal metamorphic and possibly also metasomatic overprint of the lower parts of the Witwatersrand sequence, but it had little effect on the distribution of the gold. Further local gold mobilization was triggered again by the 2.02 Ga Vredefort meteorite impact that created a secondary permeability. Later overprints during orogenic events along the margins of the Kaapvaal craton, i.e. the 2.0 to 1.8 Kheis, the 1.2 to 1.0 Namaqua, and the 0.6 to 0.5 Ga Pan-African orogenies had little or no significant effect on the distribution of the Witwatersrand gold.

Abstract This paper provides a critical review of advances made in understanding of sedimentary environments, geochemical processes, and biological systems that contributed to the deposition and diagenetic evolution of the exceptionally well-preserved and large iron formations of the late Neoarchean to very early Paleoproterozoic Ghaap-Chuniespoort Group of the Transvaal Supergroup on the Kaapvaal craton (South Africa) and the time equivalent Hamersley Group on the Pilbara craton (Western Australia). These iron formations are commonly assumed to have formed coevally but in separate basins, and they are often used as proxies for global ocean chemistry and paleoenvironmental conditions at ~2.5 Ga. However, lithostratigraphic and paleogeographic reconstructions show that the iron formations formed in a single large partly enclosed oceanic basin along the margins of the ancient continent of Vaalbara. Furthermore, although large relative to other preserved iron formations, the combined Transvaal-Hamersley basin is miniscule compared to marginal basins of the modern ocean system so that the succession probably documents secular changes in depositional environments of that basin rather than of the global ocean at the time. The iron formations comprise a large variety of textural and mineralogical rock types that display complex lateral and vertical facies variations on basinal scale. Based on detailed analyses of these variations it is concluded that the iron formations were deposited in environments that ranged from very deep-water basinal settings far below storm-wave base and the photic zone to very shallow-platform settings above normal wave base. Precipitation of both iron and silica took place from hydrothermal plumes in a dynamically circulating ocean system that was not permanently stratified. Ferric oxyhydroxide was the primary iron precipitate in virtually all of the iron formation facies. This primary precipitate is now represented by early diagenetic hematite in some of the iron formations. However, in both deep- and shallow-water iron formations most of the original ferric oxyhydroxides have been transformed by dissimilatory iron reduction to early diagenetic siderite and/or magnetite in the presence of organic carbon. Precipitation of ferric oxyhydroxides in very deep water below the photic zone required a downward flux of photosynthetically-derived free oxygen from the shallow photic zone. In these deep-water environments, under microaerobic conditions, chemolithoautotrophic iron-oxidizing bacteria may have played an important role in precipitation of ferric oxyhydroxides and acted as a source of primary organic matter. With basin fill even shallow-shelf embayments were invaded by circulating hydrothermal plume water from which ferric oxyhydroxides could be precipitated in oxygenated environments with high primary organic carbon productivity and thus iron reduction to form hematite-poor siderite- and magnetite-rich clastic-textured iron formations. Depositional models derived from the study of the iron formations along the Neoarchean-Proterozoic boundary can be applied to iron formations of all ages in both the Archean and later Paleoproterozoic. The facies architecture of the iron formations determines to a large degree the textural attributes, composition, and stratigraphic setting of high-grade iron ores hosted by them. Detailed facies information thus would assist in improving genetic models for high-grade iron ore deposits. Future research should be guided in this direction, especially in some of the very large iron ore districts of Brazil and India where very little is known about the composition and facies variations of the primary iron formation hosts and possible controls on localization of high-grade ores.