We report new petrological data for granulites from the Central Zone of the Limpopo Complex, southern Africa, and construct a prograde P-T path that traverses from high-pressure granulite-facies metamorphism to peak ultrahigh-temperature (UHT) metamorphism by rapid decompression, which was followed by further decompression and cooling. Mg-rich ( X Mg ~0.58) staurolite enclosed within poikiloblastic garnet in an Mg-Al-rich rock from the Beit Bridge area is rarely mantled by a sapphirine + quartz corona, suggesting the progress of the prograde dehydration reaction: staurolite + garnet → sapphirine + quartz + H 2 O. The symplectic sapphirine + quartz developed around staurolite probably implies decompression from P >14 kbar toward the stability of sapphirine + quartz at T ~1000 °C along a clockwise P-T path. The orthopyroxene + sillimanite + quartz assemblage mantled by cordierite aggregates in a pelitic granulite from the same area also suggests extreme metamorphism and subsequent further decompression. Various corona textures such as kyanite + sapphirine, sapphirine + cordierite, and orthopyroxene + cordierite were probably formed as a result of decompression cooling events. The prograde high-pressure metamorphism and the following UHT event relate to the collisional tectonics of the Zimbabwe and Kaapvaal Cratons, which are associated with the amalgamation of microcontinents during the Neoarchean.

Published whole-rock Sm-Nd and zircon Lu-Hf data from the Limpopo Complex and adjoining areas of the Zimbabwe and Kaapvaal Cratons provide insight into the regional crustal evolution and tectonic processes that shaped the complex. The Northern Marginal Zone of the complex, and the Francistown area of the Zimbabwe craton, represent an accretionary margin (active at 2.6–2.7 Ga) at the southern edge of that craton, at deep and shallow crustal levels, respectively. The Southern Marginal Zone represents a deep crustal level of the northern Kaapvaal Craton and was not an accretionary margin at the time of high-grade metamorphism (2.72–2.65 Ga). The syntectonic Matok granite was produced by crustal anatexis. In the Central Zone, the presence of ca. 3.5–3.3 Ga crust is indicated throughout its E-W extent by T Nd,DM model ages of metapelites and by zircon xenocrysts and their T Hf,DM model ages. The ca. 2.65 Ga granitoids in the Central Zone (the Singelele-type quartzofeldspathic gneisses in the Musina area, granitoids in the Phikwe Complex, Botswana, the so-called gray gneisses, and the Bulai charnockite) were formed by anatexis of such old crust, whereas 2.6 Ga juvenile (arc-related?) magmatism produced the Bulai enderbite, and may be a component in the Zanzibar gneiss. The Mahalapye granitoid complex in Botswana was formed by crustal anatexis at 2.0 Ga, but mafic and hybrid rocks of this age have a mantle-derived component. The data do not prohibit a collisional model for the Neoarchean high-grade metamorphic event in the Central Zone and Southern Marginal Zone of the Limpopo Complex.

Integrated geological studies in the Central Zone of the Limpopo Complex formed the basis for the construction of a composite deformation (D)–pressure (P)– temperature (T)–time (t) (D-P-T-t) diagram that shows the following: First, in the Neoarchean the Central Zone probably underwent high-pressure (HP) (P >14 kbar, T ~950 °C) conditions followed by near isothermal decompression to ultrahigh- temperature conditions (UHT) (T ~1000 °C, P ~10 kbar), before ca. 2.68 Ga. Second, the post-peak exhumation history linked to two distinct decompression cooling stages commenced at ca. 2.68 Ga and ended before the emplacement of the Bulai Pluton at ca. 2.61 Ga. Stage 1 started at P ~9 kbar, T = 900 °C, and culminated with the emplacement of leucocratic anatectic granitoids at ca. 2.65 Ga. Stage 2, linked to the development of major SW-plunging sheath folds and related shear zones, started at P ~6 kbar, T ~700 °C and ended at P ~5 kbar, T ~550 °C, before ca. 2.61 Ga. The rocks resided at the mid-crustal level for more than 600 m.y. before they were again reworked at ca. 2.02 Ga by a Paleoproterozoic event. This event commenced with isobaric (P ~5 kbar) reheating (T ~150 °C) of the rocks related to the emplacement at ca. 2.05 Ga of magma linked to the Bushveld Igneous Complex. This was followed by final exhumation of the Central Zone. The Neoarchean high-grade event that affected the Limpopo Complex is linked to a Himalayan-type collision of the Kaapvaal and Zimbabwe Cratons that resulted in over-thickened unstable crust and the establishment of HP and UHT conditions. This unstable crust initially responded to the compressional event by thrust-driven uplift and spreading of the marginal zones onto the two adjacent granite-greenstone cratons. The post-peak exhumation history was probably driven by a doming-diapiric mechanism (gravitational redistribution).

The Northern Marginal Zone (NMZ) of the Limpopo Belt, southern Africa, is a high-grade gneiss belt dominated by magmatic granulites of the charnoenderbite suite, which intruded minor mafic-ultramafic and metasedimentary rocks between 2.74 and 2.57 Ga. The intrusive rocks have crustal and mantle components, and occur as elliptical bodies interpreted as diapirs. Peak metamorphism (P ≤800 MPa, T = 800–850 °C) occurred at ca. 2.59 Ga. The highly radiogenic nature of the rocks in the NMZ, supplemented by heat from mantle melts, led to heating and diapirism, culminating in the intrusion of distinctive porphyritic charnockites and granites. Horizontal shortening and steep extrusion of the NMZ, during which crustal thickening was limited by high geothermal gradients, contrast with overthickening and gravitational collapse observed particularly in more recent orogens. The granulites were exhumed by the end of the Archean. The pervasive late Archean shortening over the whole of the NMZ contrasts with limited deformation on the Zimbabwe Craton, possibly owing to the strengthening effect of early crust in the craton. In the southeast of the NMZ, strike-slip kinematic indicators occur within the Transition Zone and the Triangle Shear Zone, where dextral shearing reworked the Archean crust at ca. 1.97 Ga.

Published models for the Limpopo Complex as a whole include Neoarchean (ca. 2.65 Ga) continent-continent collision, Turkic-type terrane accretion, and plume-related gravitational redistribution within the crust. Hypotheses proposed for parts of the complex are Paleoproterozoic (ca. 2.0 Ga) dextral transpression for the Central Zone, westward emplacement of the Central Zone as a giant nappe, and gravitational redistribution scenarios. In this chapter these models and hypotheses are reviewed and tested against new data from geophysics (chiefly seismics and gravity), isotope geochemistry (mainly Sm-Nd and Lu-Hf data), geochronology, and petrology. Among the whole-complex models, the plume-related gravitational redistribution model and the Turkic-type terrane accretion model do not satisfy the constraints. The Neoarchean collision model remains as a viable working hypothesis, whereby (in contrast to published versions) the Zimbabwe Craton appears to be the overriding plate, with the Northern Marginal and Central Zones of the Limpopo Complex as its (possibly Andean-type) active margin and shelf, respectively. Of the partial models, gravitational redistribution in the context of crustal thickening is compatible with Neoarchean collision and can explain features at the Complex–Kaapvaal Craton boundary. Paleoproterozoic dextral transpression in the Central Zone can be superimposed on Neoarchean collision, provided that it does not itself entail a continent collision. The Paleoproterozoic metamorphism is characterized by near-isobaric prograde paths, which (along with combined teleseismic and gravity data) suggest magmatic underplating. This could be related to the Bushveld Complex, and may have weakened the crust, leading to the focusing of regional strain into transcurrent movement in the Central Zone.

The magmatic and tectonic processes of the pre–2.5 Ga hot, young Earth differed profoundly from those of the modern planet. The ancient rocks differ strikingly in individual and collective composition, occurrence, association, and structure from modern rocks. Widespread forcing of Archean geology into plate-tectonic frameworks reflects unwarranted faith in uniformitarianism and in inappropriate chemical discriminants, and disregard for the lack of features that characterize plate interactions. Archean crust records extreme and prolonged internal mobility and was far too weak and mobile to behave as rigid plates, required, by definition, for plate tectonics. None of the geologic indicators of subduction, arc magmatism, and continental sundering, separation, and convergence have been documented. No Archean oceanic crust or mantle has been recognized, and the only known basement to supra-crustal rocks, including the thick basalts, high-Mg basalts, and ultramafic lavas that typify greenstone successions, consists of tonalite-trondhjemite-granodiorite (TTG) migmatites and gneisses. A thick global melabasaltic protocrust likely formed by ca. 4.45 Ga, and from it TTG suites were extracted by partial melting over the next 2 b.y. Delamination of the increasingly dense restitic protocrust enabled rise of lighter and hotter depleted mantle and hence more melting. The oldest known crustal materials are zircons, which scatter in age back to 4.4 Ga and are recycled in migmatites whose final crystallization was after 3.8 Ga, and in ancient sediments. Earth may have had a dense greenhouse atmosphere, not a hydrosphere, before 3.6 Ga, for the oldest proved supracrustal rocks are of that age, and older felsic crust may have been too hot to permit rise of dense melts. Rigid plates of lithosphere did not stabilize until a billion years after that and then were mostly small and local. Dense lavas erupted atop mobile felsic crust after 3.6 Ga produced a density inversion that was partly righted by sinking of the volcanic rocks and rising of the subjacent TTG. In some places, the early dense rocks retained cohesion and sank as synclinal keels between rising domiform diapiric batholiths. In others, the early dense rocks sank deep into mobile TTG crust, and only later in Archean time was the felsic substrate strong enough to enable dome-and-keel style. The TTG substrate rose slowly, with variable amounts of partial melting to generate more-fractionated melts and with additions of new TTG from the underlying protocrust, for hundreds of millions of years. The mantle beneath preserved cratons generated ultramafic melts that required a temperature ∼300°C hotter than modern asthenosphere ca. 3.5 Ga. Severe and prolonged lateral deformation was superimposed on large parts of some cratons during the era of volcanism and diapirism, obscuring dome-and-keel geology over broad tracts. Lower crust was at high temperature for prolonged periods and flowed pervasively, coupled discontinuously to the upper crust to produce lateral deformation therein. Rifting, separation, rotation, and collision of internally more rigid lithosphere fragments began ca. 2.1 Ga, but may have been dominantly intracontinental deformation, quite distinct from modern plate tectonics. The products of this regime differ greatly from those of Phanerozoic plate tectonics, and reflect a transitional era of erratically stiffening lithosphere. An early-depleted upper mantle has been progressively re-enriched, by delamination and subduction of crustal materials, while new “juvenile” crust derived from it has become progressively more depleted, during Pro-terozoic and Phanerozoic time.

The Archean Zimbabwe Craton in northeastern Zimbabwe is bounded along its northern margin by the east-trending Zambezi belt, which has been subdivided from south to north into the Archean Migmatitic Gneiss Terrain, overlain by Proterozoic gneiss units of the Marginal Gneiss Terrain and the Allochthonous Gneiss Terrain. The transition from the Zimbabwe Craton into the Migmatitic Gneiss Terrain occurs over a 10–15-km-wide zone across which the metamorphic grade increases from 3 to 4 Kbar and 500 °C in the northern part of the craton, to 6–7 Kbar and 700 °C in the Migmatitic Gneiss Terrain. This P-T (Pressure-Temperature) gradient was established in the late Archean and reflects exhumation of the Migmatitic Gneiss Terrain as a result of D 1 sinistral transpression due to oblique collision between the Zimbabwe Craton and an unknown rock mass to the north. D 1 structures are absent within the Marginal and Allochthonous Gneiss Terrains. The craton margin and Migmatitic Gneiss Terrain were subsequently intruded and crosscut by the 2.58 Ga Great Dyke. Between 1.1 and 0.5 Ga renewed tectonism occurred along the craton margin, forming the Zambezi belt. Deformation fabrics linked to this are mainly preserved in the Marginal and Allochthonous Gneiss Terrains and include 0.85–1.05 Ga high-P granulite-facies fabrics (D 2 ) extensively overprinted by D 3 fabrics that formed ∼0.80 b.y. ago. D 3 fabrics preserve mylonitic characteristics across the Marginal Gneiss Terrain, which is part of a 3–5-km-thick crustal-scale, shallowly north- to northeast-dipping shear zone that can be traced along the full length of the Zambezi belt north of the Zimbabwe Craton. This shear zone accommodated extension and exhumation of rocks in the Marginal Gneiss Terrain, the Migmatitic Gneiss Terrain, and the craton margin. A renewed burial and exhumation cycle affected the rocks of the Marginal Gneiss Terrain, the Migmatitic Gneiss Terrain, and the craton margin ca. 0.53 Ga, as units of the Allochthonous Gneiss Terrain overthrust the margin of the Zimbabwe Craton.