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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.

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