Scholarly article on topic 'Stagnant lids and mantle overturns: Implications for Archaean tectonics, magmagenesis, crustal growth, mantle evolution, and the start of plate tectonics'

Stagnant lids and mantle overturns: Implications for Archaean tectonics, magmagenesis, crustal growth, mantle evolution, and the start of plate tectonics Academic research paper on "Earth and related environmental sciences"

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Abstract of research paper on Earth and related environmental sciences, author of scientific article — Jean H. Bédard

Abstract The lower plate is the dominant agent in modern convergent margins characterized by active subduction, as negatively buoyant oceanic lithosphere sinks into the asthenosphere under its own weight. This is a strong plate-driving force because the slab-pull force is transmitted through the stiff sub-oceanic lithospheric mantle. As geological and geochemical data seem inconsistent with the existence of modern-style ridges and arcs in the Archaean, a periodically-destabilized stagnant-lid crust system is proposed instead. Stagnant-lid intervals may correspond to periods of layered mantle convection where efficient cooling was restricted to the upper mantle, perturbing Earth's heat generation/loss balance, eventually triggering mantle overturns. Archaean basalts were derived from fertile mantle in overturn upwelling zones (OUZOs), which were larger and longer-lived than post-Archaean plumes. Early cratons/continents probably formed above OUZOs as large volumes of basalt and komatiite were delivered for protracted periods, allowing basal crustal cannibalism, garnetiferous crustal restite delamination, and coupled development of continental crust and sub-continental lithospheric mantle. Periodic mixing and rehomogenization during overturns retarded development of isotopically depleted MORB (mid-ocean ridge basalt) mantle. Only after the start of true subduction did sequestration of subducted slabs at the core-mantle boundary lead to the development of the depleted MORB mantle source. During Archaean mantle overturns, pre-existing continents located above OUZOs would be strongly reworked; whereas OUZO-distal continents would drift in response to mantle currents. The leading edge of drifting Archaean continents would be convergent margins characterized by terrane accretion, imbrication, subcretion and anatexis of unsubductable oceanic lithosphere. As Earth cooled and the background oceanic lithosphere became denser and stiffer, there would be an increasing probability that oceanic crustal segments could founder in an organized way, producing a gradual evolution of pre-subduction convergent margins into modern-style active subduction systems around 2.5 Ga. Plate tectonics today is constituted of: (1) a continental drift system that started in the Early Archaean, driven by deep mantle currents pressing against the Archaean-age sub-continental lithospheric mantle keels that underlie Archaean cratons; (2) a subduction-driven system that started near the end of the Archaean.

Academic research paper on topic "Stagnant lids and mantle overturns: Implications for Archaean tectonics, magmagenesis, crustal growth, mantle evolution, and the start of plate tectonics"

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Stagnant lids and mantle overturns: Implications for Archaean tectonics, magmagenesis, crustal growth, mantle evolution, and the start of plate tectonics

Jean H. Bédard

Geological Survey of Canada, GSC-Québec, 490 de la Couronne, Québec, Qc G1K 9A9, Canada

ARTICLE INFO

ABSTRACT

Article history: Received 30 May 2016 Received in revised form 25 November 2016 Accepted 11 January 2017 Available online xxx

Keywords:

Archaean

Mantle-overturn

Stagnant-lid

Continental crust

Oceanic crust

Subcretion

The lower plate is the dominant agent in modern convergent margins characterized by active subduction, as negatively buoyant oceanic lithosphere sinks into the asthenosphere under its own weight. This is a strong plate-driving force because the slab-pull force is transmitted through the stiff sub-oceanic lith-ospheric mantle. As geological and geochemical data seem inconsistent with the existence of modernstyle ridges and arcs in the Archaean, a periodically-destabilized stagnant-lid crust system is proposed instead. Stagnant-lid intervals may correspond to periods of layered mantle convection where efficient cooling was restricted to the upper mantle, perturbing Earth's heat generation/loss balance, eventually triggering mantle overturns. Archaean basalts were derived from fertile mantle in overturn upwelling zones (OUZOs), which were larger and longer-lived than post-Archaean plumes. Early cratons/continents probably formed above OUZOs as large volumes of basalt and komatiite were delivered for protracted periods, allowing basal crustal cannibalism, garnetiferous crustal restite delamination, and coupled development of continental crust and sub-continental lithospheric mantle. Periodic mixing and reho-mogenization during overturns retarded development of isotopically depleted MORB (mid-ocean ridge basalt) mantle. Only after the start of true subduction did sequestration of subducted slabs at the core-mantle boundary lead to the development of the depleted MORB mantle source. During Archaean mantle overturns, pre-existing continents located above OUZOs would be strongly reworked; whereas OUZO-distal continents would drift in response to mantle currents. The leading edge of drifting Archaean continents would be convergent margins characterized by terrane accretion, imbrication, subcretion and anatexis of unsubductable oceanic lithosphere. As Earth cooled and the background oceanic lithosphere became denser and stiffer, there would be an increasing probability that oceanic crustal segments could founder in an organized way, producing a gradual evolution of pre-subduction convergent margins into modern-style active subduction systems around 2.5 Ga. Plate tectonics today is constituted of: (1) a continental drift system that started in the Early Archaean, driven by deep mantle currents pressing against the Archaean-age sub-continental lithospheric mantle keels that underlie Archaean cratons; (2) a subduction-driven system that started near the end of the Archaean.

© 2017, China University of Geosciences (Beijing) and Peking University. Production and hosting by Elsevier B.V. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/

licenses/by-nc-nd/4.0/).

1. Introduction

Although plate tectonics is the fundamental unifying theory in the Earth sciences there is no consensus on when or how it began. Modern plate tectonics is often described as subduction tectonics because the pull of subducting slabs is the strongest plate driving

E-mail address: Jeanh.bedard@canada.ca.

Peer-review under responsibility of China University of Geosciences (Beijing).

force (Forsyth and Uyeda, 1975; Chapple and Tullis, 1977; Richardson, 1992; Bercovici et al., 2000; Anderson, 2001; Conrad and Lithgow-Bertelloni, 2002, 2004; Hamilton, 2007). The andes-itic bulk composition of continental crust inspired proposals that continents grew by assembly of subduction-generated oceanic arc terranes, or by maturation of subduction zones into continental arcs (Taylor, 1967; Langford and Morin, 1976; Kelemen, 1995; Davidson and Arculus, 2006; Percival et al., 2006; Lee et al., 2007; Shirey et al., 2008; Polat, 2012; Arndt, 2013; Jagoutz and Kelemen, 2015). For genesis of Archaean continental crust,

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J.H. Bédard / Geoscience Frontiers xxx (2017)

volumetrically dominated by the TTG suite (Tonalite, Trondhjemite and Granodiorite: Martin, 1987; Moyen and Martin, 2012), some see a major role for garnet-facies melting of subducted oceanic crust in a hotter ambient mantle (Martin, 1986,1993; Rapp et al., 1991 ; Arculus, 1999). It is also proposed that much of the juvenile input that feeds continental crustal growth through time was added through subduction-driven accretion of plume-generated oceanic plateau crust (Boher et al., 1991; Hill et al., 1992; Stein and Goldstein, 1996; Abbott et al., 1997; Albarède, 1998; Arculus, 1999; Benn and Kamber, 2009), with subsequent reworking and refining of the accreted juvenile plateau crust by ongoing or later subduction magmatism and terrane collisions. Given the paramount role proposed for subduction in these continental crustal growth scenarios it is vital to test these hypotheses by unambiguously identifying subduction products in the geological record. Correct identification of ancient arcs is particularly important, because other plate tectonic manifestations (ridges, transforms) have low-to-nil preservation potentials. Alternative non-plate tectonic subduction-less mechanisms for Archaean continental crust genesis involve remelting the base of magmatically or tectonically thickened basaltic crustal sections, or of delaminated lower crust (Campbell and Hill, 1988; Zegers and van Keken, 2001 ; van Thienen et al., 2004a,b; Smithies et al., 2005a; Bédard, 2006; Van Kranendonk et al., 2007b).

Today, it is possible to see robust volcanic arcs (arcuate chains of active volcanos) located above seismically detectable dipping oceanic slabs, and to directly measure convergence between upper and lower plates. Identification of subduction in the rock record, however, requires recognition in the field and laboratory of characteristic volcanic arc-related lithofacies and geochemical signatures, and that the rock facies that characterize fore-arc, arc, and back-arc environments should have the correct architecture (cf. Condie, 2015). This becomes increasingly difficult in the deep past due to degradation of the geological record (e.g. Goodwin, 1996; Myers, 2001; Cawood et al., 2013), leading to over-dependence on geochemical palaeotectonic classifications (e.g. Pearce and Cann, 1973; Pearce and Norry, 1979; Wood, 1980). Unfortunately, trace element signatures commonly used to identify arc magmas are mimicked by crustal contamination, making black-box classifications especially unreliable for ensialic basalts, such as the eponymous continental flood basalts and many Archaean basalts (e.g. Wang and Glover, 1992; Green et al., 2000; Pearce, 2008; Condie, 2015; Li et al., 2015).

A second major criterion used to identify ancient arcs is evidence for compressional tectonics (terrane thrusting, nappes, regional folding and shearing); on the assumption that horizontal tectonics is equivalent to plate tectonics. Mapping and structural studies have recognized compressional horizontal tectonics and terrane assembly in Archaean accretionary orogens (Van Kranendonk et al., 2002; Percival et al., 2004; Tomlinson et al., 2004; Smithies et al., 2005b; Nutman and Friend, 2007, 2009; Boily et al., 2009; Windley and Garde, 2009; Czarnota et al., 2010; Kisters et al., 2010; Chardon et al., 2011; Hickman and van Kranendonk, 2012; Leclerc et al., 2012; Polat et al., 2015). On the other hand, compelling evidence for vertical tectonics in Archaean granite greenstone terrains, typically with shear sense indicators indicating ascent of granitoid domes and coeval development of pinched supracrustal synclines, has led to increasing acceptance of sagduction/partial convective overturn as an important Archaean intra-crustal redistribution and maturation process (Mareschal and West, 1980; Hickman, 1984; Choukroune et al., 1995; Chardon et al., 1996,1998; Collins et al., 1998; Bailey, 1999; deBremond d'Ars et al., 1999; Van Kranendonk et al., 2002, 2004, 2007a, 2009; Bédard et al., 2003, 2013; Robin and Bailey, 2009; Lin et al., 2013; Thébaud and Rey, 2013; François et al., 2014; Kamber, 2015;

Wiemer et al., 2016). Although vertical tectonic models have long been viewed as the an-Archic antithesis of plate tectonics, this is a false dichotomy as the two tectonic styles are not mutually exclusive, and evidence for coeval vertical and horizontal tectonics is now recognized in the Archaean record (Lin, 2005; Van Kranendonk, 2010, 2011a; Bedard et al., 2013; Lin et al., 2013; Harris and Bedard, 2014a,b). So the key question is not whether there was vertical or plate tectonics in the Archaean, but whether horizontal tectonics necessarily requires modern-style plate tectonic driving forces?

Bedard et al. (2013) and Harris and Bedard (2014a,b) argued that the distinctive Archaean rock associations, structures and geochemical signatures are best explained if continental drift started very early in Earth's history (ca. 3.9 Ga), long before active subduction (ca. 2.5 Ga), allowing horizontal tectonics and terrane accretion to occur throughout the Archaean in the absence of subduction. They argued that Archaean continental drift (like modern continental drift) is primarily the result of a traction force exerted by mantle flow on the Archaean-age sub continental lith-ospheric mantle (SCLM) keel that underlies cratons (cf. Forsyth and Uyeda, 1975; Bokelmann, 2002a,b; Bokelmann and Silver, 2002; Eaton et al., 2004; Conrad and Lithgow-Bertelloni, 2006; Eaton and Frederiksen, 2007; Husson et al., 2012; Kaban et al., 2015). Harris and Bedard (2014a,b) corroborated this hypothesis by documenting Himalayan-style orogenesis and extrusion tectonics on Venus, a planet lacking subduction zones and spreading ridges, implying that horizontal tectonics and continental drift cannot simply be equated to uniformitarian plate tectonics. Instead, Harris and Bedard (2014b) proposed that plate tectonics today is constituted of two systems: (1) a bottom-up continental drift system driven by mantle currents that started in the Early Archaean; and (2) a top-down subduction-driven system that began near the Archaean/Proterozoic boundary. In this discussion paper, the Archaean subduction-less continental drift model is combined with existing (and some new) field, structural, metamorphic and geochemical data, and with concepts emerging from recent thermo-mechanical models, so as to develop an overarching nonplate tectonic hypothesis of Archaean geodynamics and magmagenesis.

The debate about how and when plate tectonics began is handicapped by terminological ambiguities and imprecision. I argue that many differences of opinion regarding Archaean tectonics and magmagenesis can be resolved if a formal distinction is made between active subduction and passive subcretion/imbrica-tion. I also suggest that although the upwelling zones of Archaean mantle overturns are plume-like in their physics, they differ from modern plumes in scale, longevity and possibly source chemistry. The definitions of reworking and recycling from Hawkesworth et al. (2010) are slightly modified. Crustal reworking: the remobilization of pre-existing crust (mafic or felsic) by partial melting or erosion at sites within the crust. Recycling: the introduction of crust (either oceanic or continental) into the mantle by whatever process (subduction, delamination or foundering), making it available for resampling by younger mantle melting events.

2. Active subduction

Active subduction of oceanic lithosphere characterizes many convergent margins and is a fundamental part of plate tectonics today. The purpose of the review that follows is to provide a point of comparison for putative Archaean analogues, so as to show how pre-2.5 Ga convergent margin environments may have differed from modern ones.

Active Subduction is a process by which oceanic lithosphere descends in an organized manner as a coherent slab into the mantle

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because of its negative buoyancy (Forsyth and Uyeda, 1975; Molnar and Atwater, 1978; Stern, 2002, 2010; Turcotte and Schubert, 2002; Billen, 2008; Fig. 1A). The density of the old, cold, sub-oceanic lithospheric mantle (SOLM) is the most important driving force, with lesser contributions from the negative thermal buoyancy of the oceanic crust and progressive conversion of basalt or gabbro to eclogite (Turcotte and Schubert, 2002; van Hunen and van den Berg, 2008). As oceanic lithosphere subducts, the basaltic and sedimentary carapace initially experience an increase in pressure but only limited heating, locally forming blueschist facies meta-morphic rocks when convergence rates are especially rapid (Peacock, 1992, 1996; Stern, 2002). Conversely, some subduction zones retreat too fast for the upper plate to keep pace, creating mantle upwellings, high heat flows and exerting a tensional force on the upper plate, potentially leading to: (1) fore-arc extension and seafloor spreading (e.g. Stern and Bloomer, 1992; Bedard, 1999; Bedard et al., 2000; Schroetter et al., 2003); (2) development of backarc spreading ridges (Tatsumi et al., 1990); (3) detachment of ribbon continents (Arakawa et al., 2000); and (4) creation of hot, episodically-extensional, accretionary orogens (Collins, 2002a,b; Kemp et al., 2009).

After the downgoing subducting slab clears the base of the upper plate it encounters hot, convecting, asthenospheric wedge mantle (Peacock, 1996; Syracuse et al., 2010). Mantle wedge convection is primarily driven by mechanical coupling with the downgoing slab (McKenzie, 1969; Peacock, 1996), with enhanced upflow of asthenosphere into the wedge when a slab retreats. Heat diffusion from the wedge mantle into the subducting slab (Stein and Stein, 1996) causes wet sediment and basalt to devolatilize and possibly melt (Wyllie, 1988; Rapp et al., 1991; Martin, 1999). The processes by which subduction-related melt/fluid are generated and transferred into the mantle wedge are diverse (Gill, 1981; Tatsumi, 1989; McCulloch and Gamble, 1991; Rapp et al., 1991; Hawkesworth et al., 1993; Tatsumi and Eggins, 1995; Davidson, 1996; Myers and Johnston, 1996; Tatsumi and Kogiso, 1997; Poli and Schmidt, 2002; Grove et al., 2006; Bebout et al., 2007; Kelley et al., 2010; Pirard and Hermann, 2015). Water plays a major role in transferring elements like K, Ba, Rb and U into the wedge mantle, leaving insoluble elements like Nb and Ta behind in the slab. In some arcs the strongly hydrated and altered basaltic carapace and capping sediment of the downgoing slab melt (Wyllie, 1988; Peacock et al., 1994; Elliott et al., 1997; Hawkesworth et al., 1997; Johnson and Plank, 1999; Skora et al., 2015), in part due to fric-tional heating (Turcotte and Schubert, 1973; Peacock, 1996). Experimental data suggest that high-pressure anatexis of metabasalt commonly produces residual rutile, a phase that retains Nb-Ta-Ti in the slab but allows Th, Zr, Hf, LREE (Light Rare Earth Elements) to be added to the wedge mantle (Green and Pearson, 1987; Ryerson and Watson, 1987; Patino-Douce and Johnston, 1991; Rapp et al., 1991; Klemme et al., 2002). Addition of a slab-derived, water-rich arc component to the wedge lowers the solidus temperature and increases melt productivity in comparison to dry mantle melting (Kostopoulos, 1991; Pearce and Peate, 1995; Grove et al., 2006; Page et al., 2008). The decreased density of metasomatized wedge peridotite may cause it to ascend diapirically and melt by decompression (Gerya and Yuen, 2003). Melts generated from fluid-metasomatized wedge mantle have source-metasomatic geochemical signatures (Davidson, 1996; Pearce, 2008), with individual arc suites showing trends that are sub-parallel to the MORB-OIB array on a Th/Yb vs. Nb/Yb diagram (MORB = mid ocean ridge basalt, OIB = ocean island basalt), but which are systematically offset to higher Th/Yb (Fig. 2A). This systematic offset results because variably depleted wedge mantle tends to receive roughly the same amount and types of metasomatic agents from any given slab segment. The gabbroic middle and lower crust of slabs

probably plays only a subordinate role in arc magmagenesis, because gabbroic cumulates are inherently more refractory and contain less water than seafloor-altered basalts, and because there is insufficient time for the cold core of the subducting slab to heat up. Simply put, for an oceanic slab to actively subduct it must be cold, and if it is cold, then only the uppermost crust can heat up enough to melt significantly.

As an arc matures an increased degree of fractionation and crustal recycling generates evolved calc alkaline magmas (e.g. Gill, 1981; Arculus, 1994; Müntener et al., 2001; Grove et al., 2003b; Blatter et al., 2013; Lee and Bachmann, 2014; Keller et al., 2015). Subduction beneath a continent provides abundant opportunities for crustal reworking (assimilation) and modification of isotopic signatures (Hildreth and Moorbath, 1988; Sisson et al., 1996), but crustal contamination can usually be distinguished from source-metasomatism as it creates oblique trends on the Th/Yb vs. Nb/Yb diagram (Fig. 2E, Pearce, 2008). Boninites are wet, high-SiÜ2-MgO melts with low incompatible trace element contents that form when strongly depleted mantle residues are fluxed by arc fluids (Crawford et al., 1989; Pearce et al., 1992; Parkinson and Pearce, 1998; Bédard, 1999; Pagé et al., 2008). The source of Phanerozoic boninites is more depleted in incompatible elements than the residues of MORB extraction, which probably requires extensive hydrous melting in a prior arc environment (Crawford et al., 1989; Pearce et al., 1992; Pearce and Peate, 1995; Bédard, 1999; Pagé et al., 2009). Late-stage potassic magmas (e.g. shoshonites) are small-degree melts of mantle domains previously metasomatized by hydrous fluids/melts derived from subducted continental material (Eklund et al., 1998; Callegari et al., 2004; Scarrow et al., 2009; Campbell et al., 2014; Cai et al., 2015; Lu et al., 2015).

Delamination (or foundering) of high-density lower crustal arc cumulates and restites is now widely recognized (Herzberg et al., 1983; Kay and Kay, 1991; Rudnick, 1995; Arculus, 1999; Tatsumi, 2000; Jull and Kelemen, 2001; Müntener et al., 2001; Saleeby et al., 2003; Behn and Kelemen, 2006; Greene et al., 2006; Lee et al., 2006; DeCelles et al., 2009; Ducea, 2011; Hacker et al., 2015; Jagoutz and Kelemen, 2015), and helps to generate the overall intermediate nature of arcs from primary melts of basaltic composition. Relamination occurs when buoyant lithologies subducted into denser upper mantle rise up to be emplaced at the base of the crust (Hacker et al., 2011; Jagoutz and Kelemen, 2015). This can occur by imbrication where buoyant material is thrust into or just beneath lower arc crust (Grove et al., 2003a; Ducea et al., 2009); ascent along a subduction channel (Gerya et al., 2002; Warren et al., 2008); or rise of diapirs through the mantle wedge (Gerya and Yuen, 2003; Chatterjee and Jagoutz, 2015). The Ti-Nb-enriched material that remains after devolatilization and melting of the downgoing oceanic slab, the SÜLM, and portions of the upper plate removed by subduction erosion and lower crustal delamina-tion, are thought to accumulate at the core-mantle boundary, and are partly recycled later as components of mantle plume OIBs (Hofmann and White, 1982; Hofmann, 1997; Tatsumi, 2000; White, 2015; Castillo, 2016).

3. Convergent margins without active subduction?

At some convergent margins the lower plate plays a more passive role and cannot drive its own motion. For example, there are several places where oceanic ridges must have been subducted (e.g. Abratis and Wörner, 2001; Burkett and Billen, 2009; Cole and Stewart, 2009; Morell, 2015), despite their inherent buoyancy and problems with the transmission of slab-pull tensional stresses through the ridge-axis discontinuity. Successful subduction of spreading ridges leads to formation of slab windows through which the asthenosphere wells up, producing localized regional uplift and

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eruption of alkali basalts (e.g. Gorring et al., 1997; D'Orazio et al., 2000; Haschke et al., 2006). Large oceanic plateaux (Coffin and Eldholm, 1994; Taylor, 2006) commonly are too buoyant to subduct (Cloos, 1993; Abbott et al., 1997); a classic example being the choking of the Vitiaz Trench due to arrival of the Ontong-Java Plateau, followed by nucleation of a new subduction zone behind the accreted plateau (Mann and Taira, 2004). That some oceanic plateau segments have successfully subducted (e.g. Hikurangi Plateau beneath North Island, New Zealand) suggests that there is a size/density cutoff limit for successful subduction of buoyant crustal segments (Davy et al., 2008). Arrival of buoyant material at a trench that does not halt subduction completely, typically causes large-scale erosion and recycling of upper plate crust into the mantle (von Huene and Scholl, 1991 ; Clift et al., 2005, 2009; Scholl and von Huene, 2007; DeCelles et al., 2009; Coldwell et al., 2011; Stern, 2011). Flat-slab segments occur where the underthrust oceanic lithosphere is anomalously buoyant and resists subduction, and these segments are commonly associated with orogenic pulses (Fig. 1B: Kay and Mpodozis, 2002; Ramos and Folguera, 2009; Flament et al., 2015; Margirier et al., 2015; Hu et al., 2016).

Because the continental crust is positively buoyant compared to the mantle it too mostly resists subduction, although high-pressure facies assemblages in continental crustal rocks attest to partial subduction followed by exhumation (Chopin, 2003). Partial subduction of passive continental margins beneath peri-continental arcs is commonly attributed to the traction exerted by an attached negatively buoyant old oceanic slab, which eventually breaks off after the buoyant continental margin enters the trench (Wortel and Spakman, 2000; van Hunen and Allen, 2011 ). There are many consequences to slab breakoff, including a magmatic pulse, cessation of the slab-pull force, tectonic underplating, and rebound of the partly-subducted continental margin (Davies and von Blackenburg, 1995; Hildebrand and Bowring, 1999; Chemenda et al., 2000; Duretz and Gerya, 2013; Hildebrand and Whalen, 2014). Breakoff-induced rebound of partly-subducted passive margins is one of the preferred mechanisms for ophiolite obduction (Schroetter et al., 2003, 2005,2006; Tremblay et al., 2011 ). A variant of this scenario is underthrusting of continental crust beneath continental crust, as in the Himalayas.

A common view about docking/accretion of buoyant terranes (arcs, microcontinents, oceanic plateaux) and successful ridge and oceanic plateau subduction, is that there must be extensive adjoining segments of active subduction involving coherent, steeply-dipping, negatively-buoyant oceanic slabs; so as to provide the force needed to drive convergence and entrain discontinuities and buoyant 'pips' to depth (cf. van Hunen et al., 2000, 2002; Taramon et al., 2015). In principle, the Indian continent could have been dragged down by an attached leading edge oceanic slab, but as this slab is no longer attached to India, this cannot explain ongoing convergence (Negredo et al., 2007; Capitanio and Replumaz, 2013; Duretz and Gerya, 2013). A possible active subduction-related explanation for continuing convergence between India and Asia is that the Indian Ocean plate is mechanically coupled with the Indian craton, and that both move north because of slab pull forces generated by subduction of normal oceanic

lithosphere beneath Asia. An alternative is to posit a more active role for continents (van Hunen et al., 2000; DeCelles et al., 2009; Alvarez, 2010; Bédard et al., 2013; Harris and Bédard, 2014b; Chen et al., 2015), where ongoing convergence between India and Asia is caused by pressure on India's SCLM keel induced by northward flow of mantle from the sub-African upwelling zone, with more localized pushes coming from individual plumes that detach from it (Alvarez, 2010; Burke, 2011; Becker and Faccenna, 2011; Cande and Stegman, 2011). Motion of the American continents may also be primarily driven by westwardly flowing mantle pushing against their lithospheric cratonic keels (Bokelmann and Silver, 2002; Bokelmann, 2002a,b; Eaton et al., 2004; Conrad and Lithgow-Bertelloni, 2006; Eaton and Frederiksen, 2007; Husson et al., 2012; Harris and Bédard, 2014b; Kaban et al., 2015), with accretionary orogens and crustal imbrication zones forming on the West coast when unsubductable terranes dock against the advancing continents (Coney et al., 1980; Dickinson, 2004; Nelson and Colpron, 2007; Greene et al., 2009; Ramos, 2009; Hildebrand and Whalen, 2014).

As terrane accretion by failed subduction beneath drifting continents is not primarily driven by active subduction of the lower plate, it seems illogical to refer to such types of convergent margins as subduction zones. It would clarify the debate about when subduction began if terms such as 'subduction zones' and 'subduction margins' were only applied to ancient convergent margins and accretionary orogens when evidence for subduction is clear.

4. Archaean and Hadean convergent margins: active subduction or passive subcretion?

Numerical thermo-mechanical and analogue modelling (van Thienen et al., 2004a,b; van Hunen and van den Berg, 2008; Sizova et al., 2010, 2015; van Hunen and Moyen, 2012) yields insight into driving forces and fluxes in a hotter early Earth, complementing field, structural and geochemical studies. The most important driver for active subduction today is the negative buoyancy of the cold, dense, sub-oceanic lithospheric mantle (SOLM). As a crust-dominated oceanic lithosphere provides only a weak driving force, and is characterized by frequent drips or breakoffs instead of coherent plate behaviour (van Hunen and van den Berg, 2008; van Hunen and Moyen, 2012); then existence of a SOLM layer would also seem a necessary premise for hypotheses of Archaean and Hadean oceanic lithospheric subduction to be permissible. It is the presence of a SOLM in the van Hunen and Moyen (2012) models that gave their oceanic lithosphere enough density and strength to create and transmit a slab-pull force, keeping the oceanic slab moving forward even as the tip dripped off periodically. Plate stiffness also affects the way deformation is partitioned. The 2-dimensional models of Sizova et al. (2010) explored fixed-rate (5 cm/yr) convergence between an oceanic and continental terrane at different potential temperatures. At their lowest model mantle temperatures they reproduced one-sided modern-style active subduction; whereas at higher temperatures (mantle T >250 °C above present day), imposed horizontal motion was accommodated by internal deformation of the weaker oceanic

Figure 1. Cartoon illustrating differences between: (A) active subduction, (B) subduction of buoyant terranes due to trenchward continental drift, and (C) Archaean subcretion tectonics. (A) Active subduction; the thermal slab pull force acting on a stiff oceanic lithosphere is largely responsible for lower plate motion and arc magmas are created when fluids and melts escape the downgoing slab to flux melting in the convecting mantle wedge. (B) Trenchward continental drift has successfully subducted a smaller buoyant pip, creating a flat-slab segment, an orogenic pulse, and causing a shift in the location of active volcanism. Arrival of larger buoyant terranes at a trench can terminate subduction. (C) Archaean subcretion tectonics occurs when the drifting continent pushes against unsubductable oceanic lithosphere. Imbricated basalt-dominated slabs heat up and melt to create syn-kinematic TTG suites (yellow). Localized tectonic thickening in the foreland allows formation of sedimentary belts (pale blue). This scenario differs fundamentally from active subduction (A) as the lower plate is passive, cannot cause plate motion, and there are no source-metasomatic magmatic arcs formed through focussed slab devolatilization. Localized steepening/eclogitization of imbricated segments in the orogenic foreland could trigger focussed mantle upwelling above a zone where seawater-altered basalts (and basaltic drips) lose their fluids, creating subordinate arc-like magmatic suites.

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Figure 2. Log Th/Yb vs. Log Nb/Yb diagram from Pearce (2008). (A) Shows the Mariana arc data and source-metasomatic and melting vectors from Pearce (2008). The OIB-MORB array is defined from modern seafloor and ocean island data. Extent of source depletion is shown. The dashed isopleths show effect of adding equal amounts of a component derived from the subducting slab to mantle sources with different degrees of prior depletion. Note that the Mariana arc rocks form an array parallel to the MORB—OIB array. (B and C) Show data from the Chibougamau segment of the Abitibi greenstone belt from Leclerc et al. (2011,2016); with the order of symbols in the legend corresponding to the stratigraphic order. The low-MgO (<3%) tholeiites are equivalent to the Type III Na-rhyolites from Leclerc et al. (2011) and Bedard et al. (2013); and are restricted to the Waconichi Formation. Calc-alkaline rocks (C) are not discriminated for MgO content. Note that the oblique data array defined by the coeval rocks of the Chibougamau area overlaps exactly with the distribution of the younger Blake River Group (from Bedard et al., 2013) and Shebandowan area (Lodge, 2016). Oblique data arrays are typical for Archaean volcanic suites and cannot be explained with the source metasomatic model illustrated in panel A. If these oblique arrays are attributed to source metasomatism, then fertile mantle must always be more contaminated than depleted mantle, which seems implausible. (D) Shows felsic plutonic rocks from the Chibougamau area as 'X' symbols (data from Leclerc et al., 2016). Tonalite-trondhjemite-granodiorite suite data from the Douglas Harbour domain of the NE Superior (unpublished data, see Bedard et al., 2003 for location) are shown as coloured symbols. Crustal contamination could incorporate any of these TTG rocks/magmas, such that unique AFC paths are not expected. (E) Mixing and AFC models are superimposed on the datafields. Bulk mixing lines are shown as bold lines. Only the 30% felsic component tick is labelled. The mafic pole 'B' is the average of 3 of the most primitive, least-contaminated

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lithosphere. This led Sizova et al. (2010) to conclude that weak, buoyant, Early Archaean oceanic lithosphere was probably unsubductable (cf. Oxburgh and Parmentier, 1977).

Sizova et al. (2015) examine in-situ maturation of Archaean crust at high mantle potential temperatures. In these 2-D models, systematic mantle upflow leads to the development of a coherent oceanic plate, where ridge push forces coupled with basal drag compress the adjoining continent, creating a convergent margin with short-lived subduction/drip events similar to those discussed by van Hunen and Moyen (2012). Although some features of the end-stage Sizova et al. (2015) models resemble modern-style active subduction, the formation of a moving oceanic plate with a spreading centre in these models may be the result of a strong, long-lived, focused mantle upflow zone. Linear ridge-like mantle upwellings may not be typical of the early Earth, however, which more likely had a less organized, unstable, 3-D upwelling pattern. The large ridge-push force (thermal dome) created by the strong, focused, mantle upwelling in the Sizova et al. (2015) model also contrasts with the very small ridge-push force that typifies modern ridges (Russo and Silver, 1996) where mantle upflow is mostly a passive response to plate divergence (e.g. Blackman and Forsyth, 1992).

Episodic and intermittent Archaean subduction has been posited by Moyen and van Hunen (2012), who suggested that Neoarchaean calc-alkaline events with a periodicity of 5—10 Ma in the Abitibi greenstone belt might reflect short-lived breakoff 'arcs' superimposed upon a background of plume-related tholeiitic magmatism. Although the models are instructive and basaltic drips from underthrust basaltic crust seem probable, several aspects of this interpretation are inconsistent with stratigraphic and geochemical data (e.g. Fig. 2B—E). I do not repeat all of the arguments previously advanced against the Archaean subduction hypothesis (absence of true ophiolites, rarity of andesites and accretionary mélanges.etc.) and refer the reader to Hamilton (1998), Bédard et al. (2003, 2013), Bédard (2006, 2013), Harris and Bédard (2014a,b), Thurston (2015) and Kamber (2015) for overviews. It was previously argued that calc-alkaline volcanics in the Abitibi belt are not arc-related, but formed through intermittent melting at the base of a thick (or thickened) basaltic crust (Bédard et al., 2013). Additional arguments against the multiple arc model for the Abitibi greenstone belt are developed next.

(1) Most Abitibi calc-alkaline magmas do not closely resemble Phanerozoic arc magmas in terms of either geochemistry or volcanic facies (Bédard et al., 2013; Thurston, 2015). When plotted on the Th/Yb vs. Nb/Yb diagram (Fig. 2C and d), Abitibi volcanic suites show oblique trends, requiring variable admixture into uncontaminated mantle melts of a high-Th/Yb component (cf.; Barley, 1986; Pearce, 2008; Barnes et al., 2012; Bédard et al., 2013). This is a feature common to most Archaean volcanic suites (Pearce, 2008; Bédard et al., 2013); whereas Phanerozoic arcs typically yield arrays that are parallel to the MORB-OIB array (e.g. Fig. 2A; Pearce, 2008). In the Abitibi case, Bédard et al. (2013) argued that the high-Th/Yb contaminant was a trondhjemitic anatectic melt derived from young, isotopically juvenile metabasalts located in the lower crust; leaving garnet, pyroxene and a titanate mineral in their

residues. They argued that contamination occurred in two ways (Fig. 2C—E): (1) low-degree anatectic melts bled into through-flowing mafic magma by a process that resembles AFC (Assimilation-Fractional Crystallization: DePaolo, 1981); (2) there was episodic bulk mixing between basalt and anatectic felsic melt in the complex conduit system (e.g. Lafleche et al., 1992; Bedard et al., 2009).

(2) The sediment-starved hiatuses that separate many basalt sequences in Archaean greenstones (e.g. Thurston et al., 2008, 2012) do not appear to contain any distal pyroclastic calc-alkaline airfall deposits. High-viscosity evolved continental arc magmas typically form subareal constructional strato-volcanos that shed abundant intermediate to felsic detritus to adjoining basins; and long-lived continental arcs are proposed to have constituted the terranes adjoining the Abitibi belt (e.g. Percival et al., 2006). If a long-lived Andean margin was really located nearby, then why are there no correlatable calc-alkaline proximal epiclastic to distal airfall deposits located within these long intervals of chemical sedimentation (e.g. Thurston et al., 2012)?

(3) As the intermittent Abitibi 'arcs' are inferred (e.g. Wyman, 1999) to have alternated with plumes (so as to explain the intervening komatiites and tholeiites), it would be necessary for new subduction zones to have nucleated repeatedly on short timescales, as a major plume head would probably have obliterated the previous subduction zone slab (Bedard, 2013). Moyen and van Hunen (2012) argued that the periodicity of subduction breakoff events in their simulations matches that of Abitibi belt calc-alkaline pulses (ca. 5—10 Ma). However, the interval between the major formation-rank Abitibi belt calc-alkaline pulses is commonly much shorter than this. For example, the youngest Tisdale calc-alkaline magma is dated at 2703 Ma (ages from Ayer et al., 2002); is overlain by the Kinojevis plume magmas, and then by the Kinojevis calc-alkaline magmas at 2701 Ma. Thus, only ca. 2 Ma separates the two putative arc successions. Very high and unrealistic convergence rates would be needed to get a coherent slab of oceanic lithosphere back to a depth of >100 km where it could obtain heat from a convecting mantle wedge so as to generate the younger 2703 Ma Kinojevis arc magmas. Conversely, calculations byJagoutz and Behn (2013) suggest that the periodicity of lower crustal delamination/dripping events ranges from <0.5 to ca. 5 Ma for dense unstable layers from <1 to ca. 15 km thick, a rhythm which could explain the 1st-order tholeiitic/ calc-alkaline cyclicity observed in the Abitibi stratigraphy without need of active subduction.

(4) Thin (10 s of m), laterally-discontinuous calc-alkaline lavas and tuffs embedded within the otherwise monotonous tholeiite-dominated lava plain sequences are compositionally identical to the thicker (<1 km), formation-rank calc-alkaline units (Fig. 2C). These thin calc-alkaline units could not plausibly represent additional, miniature, syn-plume arcs.

If Bedard and Harris (2014) are correct in interpreting the juvenile Neoarchaean belts of the Southern Superior as reflecting

coeval formation of small extensional ocean basins (Red Sea like?)

between detached ribbon-continents, then convergent margin

Bruneau basalts from Leclerc et al. (2016). Felsic poles are: 'A' the average of the most siliceous Waconichi Type I-II calc-alkaline rhyodacites (TTG-like with strong HREE-depletion); and 'L' the most siliceous Waconichi Type III Na-Rhyolites. The dark blue bold line marked C shows the effect of 90% closed system fractional crystallization from basalt B of a 3-phase olivine + plagioclase + clinopyroxene (14:57:29) fractionation assemblage. Other conditions, as previously modelled in Leclerc et al. (2011), are QFM+2, 0.7% H2O, and 0.25 GPa, with a liquidus temperature of 1230 °C. Partition coefficients typical for basaltic melts were used, and kept constant for simplicity. The red dotted line shows 90% solidification with an AFC (assimilation—fractional—crystallization) model, done with a stepwise approach involving assimilation into basalt B of the 'A' rhyodacite with an assimilation/crystallization ratio of 0.5.

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interpretations for most Abitibi calc-alkaline pulses would necessarily be invalid. On the other hand, the youngest and largest volumes of calc-alkaline magmatism in the Abitibi sub-province (e.g. Blake River Group: Ayer et al., 2002; Ross et al., 2011; Shebandowan Belt: Lodge, 2016) are roughly coeval with the start of convergence and terrane accretion, and it is plausible to hypothesize that these youngest calc-alkaline sequences might have formed at accre-tionary convergent margins where some oceanic crustal segments were stiff and cold enough to experience active subduction (as defined above) for a short length of time, as suggested by Moyen and van Hunen (2012). However, this hypothesis seems inconsistent with the fact that the <2700 Ma Blake River Group and Shebandowan mixed tholeiitic—calc-alkaline sequences do not form arrays similar to modern arcs, but are compositionally indistinguishable from the older calc-alkaline units in the Abitibi (Fig. 2C—E). In fact, only a few studies have identified truly arc-like lithofacies and source metasomatic signatures in Archaean volcanic suites (e.g. Szilas et al., 2012: Greenland; DeJoux et al., 2014: Yil-garn; Smithies et al., 2005b: Pilbara), suggesting that if modernstyle active subduction zones did occasionally form in the Archaean, they were probably short-lived and did not drive large plates. Alternative, non-subduction scenarios for genesis of isolated arc-like Archaean magma suites will be proposed later (cf. Bedard et al., 2013; Harris and Bedard, 2014a) in the context of continent-driven imbrication of Archaean oceanic crust (Fig. 1C).

5. Stagnant-lid models

Stagnant-lid convection (Richter, 1985; Solomatov and Moresi, 1997, 2000; Moresi and Solomatov, 1998; Sleep, 2000; van Thienen et al., 2005; Ernst, 2007; Piper, 2013; O'Neill and Debaille, 2014) is characterized by an immobile, conductively-cooled lid, where underlying mantle develops unsteady convection cells (Fig. 3A). When the rate of volcanic resurfacing is high the conductive end-member stagnant-lid scenario is perturbed and may be overwhelmed by melt advection to the surface and delamination of cold lower crust into the mantle (van Thienen et al., 2004a,b). In an unstable/transitional stagnant-lid regime like Venus, hot spot volcanism is considered the primary heat loss mechanism (Phillips et al., 1991; Noltimier and Sahagian, 1992; Ernst et al., 2001) and frequent volcanic resurfacing events bury, rework and recycle much of the pre-existing crust (cf.; van Thienen et al., 2004a,b; Hansen, 2007a, b). As calculated by van Thienen et al. (2005), advection of heat by voluminous basaltic eruptions could potentially have expelled enough heat to dissipate radiogenic heat and cool a stagnant-lid early Earth.

Induced thermal convection beneath stagnant-lid crust was modelled by Johnson et al. (2014); yielding behaviours that depend on mantle temperature. In their hotter scenario (Johnson et al., 2014; 1600 °C potential temperature), small, unsteady convection

cells developed in the aesthenospheric mantle beneath the conductive mechanical boundary layer (45 km oceanic crust and 55 km of lithospheric mantle, cf. van Thienen et al., 2005). In these models, basalt that erupted at the surface sagducted through the softened crust and mixed into the mantle in less than 5 Myr. Rapid delivery to the Moho of seafloor-metamorphosed mafic volcanics by sagduction (Vlaar, 1986; Zegers and van Keken, 2001; van Thienen et al., 2004a; Bédard, 2006; Lin et al., 2013; François et al., 2014; Fischer and Gerya, 2016) would provide a ready source of wet fusible metabasalt that could generate TTG-like anatectic magma (Palin et al., 2016). Another important feature of the Johnson et al. (2014) 1600 °C model is that unsteady aesthe-nospheric convection cells rapidly erode and destroy the SOLM that was initially present, making stagnant-lid Early Archaean oceanic lithosphere unsubductable. As the Earth cooled and mantle potential temperatures decreased with time it is inferred that: (1) the vigour of sagduction would decrease, contributing to the change of structural style between the Archaean (common dome-and-basin structure) and the Post-Archaean (long linear orogens with thrust and fold belts); (2) the efficiency of erosion of the SOLM by the unsteadily convecting aesthenosphere would tail off, stiffening and densifying the oceanic lithosphere and making segments of it progressively more similar to the modern oceanic lithosphere, and hence, progressively more subductable.

Most thermal and convection models applied to the Archaean and Hadean (Bickle, 1978, 1986; Sleep and Windley, 1982; Nisbet and Fowler, 1983; Hargraves, 1986; Hynes, 2014) have assumed that the Archaean oceanic lithosphere formed by seafloor spreading, producing thermal, rheological and density profiles similar to those of modern oceanic lithosphere. There is an element of circularity in this assumption, as it requires Archaean SOLM to thicken systematically with distance away from a linear spreading-ridge thermal anomaly (Parsons and Sclater, 1977, Fig. 1A), which would make old Archaean oceanic lithosphere both stiff and dense. Conversely, in a stagnant-lid context of unsteady aesthenospheric convection cells, crustal segments would never drift very far from a thermal anomaly, never form a significant SOLM thickness, and so remain unsubductable. The absence of a SOLM from stagnant-lid oceanic lithosphere could perhaps also explain why ophiolitic mantle sections have yet to be clearly documented in Archaean greenstone terrains; the reason being that only thick, solidly attached SOLM sections can be preserved during obduction. The assumption that oceanic crust formed by seafloor-spreading also has repercussions on thermal evolution models, since mobile lid tectonics is an efficient way to evacuate heat from Earth (Moresi and Solomatov, 1998; van Thienen et al., 2005), to the point that assuming a modern plate tectonic regime leads to a 'thermal catastrophe' when today's Urey ratio is extrapolated to Archaean time (Korenaga, 2006). Silver and Behn (2008) showed that in stagnant-

Figure 3. Scaled cartoon illustrating a generic Archaean stagnant-lid crust and mantle overturn system. (A) Between overturn events (sometime between 3.9 and 2.7 Ga), Earth's surface is a mosaic of basalt + komatiite shield volcanos (green). Oceanic crust (ca. 40 km thick) and continents (pink) with accompanying SCLM (sub continental lithospheric mantle) formed during a previous overturn event(s), yielding a complementary residual depleted upper mantle (DM). In the repose period between overturns, the conductive lid induces small-scale convection cells that cool the upper DM. Small thermal upwellings prevent formation of a stable SOLM (sub oceanic lithospheric mantle), leaving the oceanic lithosphere unsubductable. The lower mantle is posited to be a mixture of DM and lower mantle resulting from previous overturn events, and grades down to less depleted mantle compositions that resemble PM (primitive mantle). EER is an early enriched reservoir ponded at the core-mantle boundary. The lower mantle does not cool efficiently and heats up through addition of heat from the core (large red arrows) and internal radioactive heating, eventually triggering the next overturn (panel B). (B) During a mantle overturn event, thermally buoyant lower mantle ascends as large plume-like structures. Mantle in the overturn upwelling zones (OUZOs) melts by decompression, delivering large volumes of mafic-ultramafic melts with PM-like geochemical signatures. During ascent and outflow of the OUZO mantle there is extensive mixing with ambient DM, with delaminated residues of crustal melting, and with foundered crust; producing a hybrid/rehomogenized mantle (khaki colour). If the OUZO rises beneath stagnant-lid oceanic crust (center of image), large volumes of mafic/ultramafic crust affected by seawater alteration are recycled. Rafts of older mafic crust may remain embedded in younger igneous crust (cf. Kamber, 2015). On the right, an OUZO ascends beneath a pre-existing craton and reworks it. Much of the SCLM is resorbed (cf. Sobolev et al., 2011), exposing lower crust to mantle heat, triggering extensive remelting (yellow) of pre-existing TTGs, leading to formation of enderbite (pyroxene-tonalite) and granites. Basalt and komatiite lavas drip down through soft lower crust (sagduction). Outflow from the spreading OUZO stretches and fragments the older craton, rafting off ribbon-continents and creating narrow oceanic tracts like the Abitibi. In regions more distal to the OUZO, continents experience lateral mantle flow and drift, growing by lateral accretion/subcretion/reworking of stagnant-lid crust (Bedard et al., 2013).

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lid mode, heat loss is less than in active-lid mode, potentially resolving this problem.

Stagnant-lid lithosphere is both a thermal and mechanical boundary layer. Heat loss is partly due to conduction (stagnant-lid ss), but there is additional cooling through advection of magma to the surface (eruption) and advection of cold basalt and restites into the mantle as negative diapirs (delamination). Sizova et al. (2015) and Fischer and Gerya (2016) examine in-situ maturation of Archaean/Hadean crust at high mantle potential temperatures, and show that crustal reworking, partial convective crustal overturn and delamination-driven recycling of garnet-bearing lower crust into the mantle as a result of plume magmatism probably played a major role in the early stages of continent formation (cf. Bedard, 2006; Van Kranendonk et al., 2007b; Smithies et al., 2009; Johnson et al., 2014). In these non-subduction models, crustal underplating by hot decompressing mantle and its partial melts trigger extensive lower crustal anatexis (Palin et al., 2016), followed by vertical intra-crustal exchange (partial convective overturn) and hybridization between basalt and partial melts of the crust. Dense lower crustal cumulates, anatectic restites or foundered metabasalt delaminate into the mantle. Delamination exposes the lower crust to new pulses of heat as mantle wells up to fill the space.

What would be the consequences of a stagnant-lid induced-convection system on the early Earth mantle cooling system? Because the lid loses heat from its top it creates a thermal gradient in the mantle, which should induce convection (Fig. 3A). The 2-D models of Johnson et al. (2014) suggest that induced thermal convection cells are unsteady with wavelengths of ca. 300—400 km. It seems extremely unlikely that such cells could taper down all the way to the core-mantle boundary, as this would imply unrealistic aspect ratios between 10/1 and 20/1. This implies that the convection cells beneath stagnant-lid crust would terminate at a mid-mantle depth; which in turn, suggests that stagnant-lid phases were periods of layered mantle convection, where efficient convective/advective cooling of the uppermost mantle did not extend to the deep mantle. This would have prevented the deep mantle from cooling and have perturbed the balance between Earth's heat generation and heat loss. Heat generated through radioactive decay in the lower mantle and the large heat flux out of the core would have accumulated in the lower mantle, eventually triggering overturn/mixing/cooling events (Fig. 3B). This broad prediction is consistent with the concept of periodic mantle overturns (see below, O'Neill et al., 2007, 2013), and with variations of magmatic potential temperatures through time (e.g. Herzberg et al., 2007, 2010; Herzberg and Gazel, 2009).

6. Archaean oceanic crust: what was it like and is any of it preserved?

Given the absence of convincing Archaean ophiolites (with refractory lithospheric mantle, and cogenetic sheeted dykes, lavas and cumulates), hypotheses for the formation of Archaean oceanic lithosphere by seafloor-spreading lack geological corroboration. As the vast majority of Archaean basalts differ from MORB in having higher SiO2, lower TiO2, and more enriched incompatible element concentrations (Fig. 4; Condie, 2005, 2015; Bedard, 2006; Leclerc et al., 2011; Kamber, 2015; Condie et al., 2016), then there is no geochemical support for a uniformitarian seafloor-spreading model for Archaean MORB crust. Bedard et al. (2013) and Harris and Bedard (2014a) argued that the absence of Archaean arcs and ridges, with consequent absence of associated plate-boundary forces, implies a stagnant-lid regime should have dominated the Archaean Earth; and that the Abitibi greenstone belt is an example of Neoarchaean stagnant-lid oceanic crust (cf. Kamber, 2015; Thurston, 2015). Application of a stagnant-lid model suggests

most of Earth should have been covered by a mosaic of intermittent basaltic shield volcanos, with crustal thickness being limited by the basaltic solidus and depth-dependent conversion to dense garnetiferous assemblages. This is consistent with data from spherules indicating a paucity of felsic rocks in impact-excavated Archaean oceanic crust (Krull-Davatzes et al., 2014). This crust differs from Phanerozoic oceanic crust because it did not form by seafloor-spreading, so explaining the absence of Archaean sheeted dyke systems.

Most Archaean greenstone belts contain komatiites and primitive tholeiitic basalts that: lack marked LREE enrichment and negative Nb—Ta anomalies (the so-called non-arc basalts); yield high calculated Archaean mantle potential temperatures (e.g. Herzberg et al., 2010; Herzberg, 2011; Condie et al., 2016); and are widely attributed to mantle plumes (Campbell et al., 1989; Tomlinson and Condie, 2001; Arndt, 2003). Phanerozoic oceanic plateau magmas scatter between PM (Primitive Mantle) and DEP (Depleted Plume component) on the Zr/Nb vs. Nb/Th diagram (Fig. 4, Condie, 2005) and are interpreted to be derived from ascending plume heads, locally including a deep-seated depleted plume component (DEP; Kerr et al., 1995; Breddam, 2002; Condie, 2005). Archaean tholeiites have geochemical signatures very similar to those of Phanerozoic oceanic plateaux (Fig. 4A), and it is widely considered that oceanic plateaux are their closest analogues (Puchtel et al., 1998; Hollings et al., 1999; Kerrich et al., 1999 Kerr et al., 2000; Arndt et al., 2001; Polat and Kerrich, 2001; Arndt, 2003; Condie, 2003, 2005; Ernst and Buchan, 2003; Smithies et al., 2005a,b; Kerrich and Polat, 2006). This has led to proposals that many Archaean greenstones formed as plume-related oceanic plateaux, and that continents grew in part through docking of such oceanic plateaux.

Although there are resemblances between Archaean greenstones and Phanerozoic oceanic plateaux; there are also differences in facies, thickness and longevity. For example, the Ontong-Java Plateau crust is about 35 km thick and is dominated by tholeiitic basalts (Kerr, 2014). Assuming a 7-km-thick oceanic crust as a starting point implies that most of the Ontong-Java plateau formed as a single, short-lived (<5 Myr), ca. 28-km-thick eruptive/under-plating event. A thick section of dense rocks in the lower Ontong-Java crust is interpreted to represent the cumulate extract (Farnetani et al., 1996; Miura et al., 2004). Although aggregate sections of well-preserved Archaean greenstone belts are comparable in thickness to the Ontong-Java plateau, individual basaltic pulses in Archaean greenstone belts are typically only 2—7 km thick and are punctuated by intervals of chemical sedimentation and/or by calc-alkaline sequences (Ayer et al., 2002; Thurston et al., 2008, 2012; Leclerc et al., 2011; Bédard et al., 2013; Kamber, 2015; Thurston, 2015). In the Abitibi belt, basaltic ± komatiitic pulses are separated (on average) by 5 to 10 Ma intervals and span a total age range of >50 Ma (Ayer et al., 2002), whereas pulses of similar thickness in the East Pilbara are spaced over a longer timespan of ca. 350 Ma (Van Kranendonk et al., 2002; Hickman and Van Kranendonk, 2012), and magmatic cycles of the Barberton belt cover ca. 320 Ma (Lowe and Byerly, 2007; Van Kranendonk et al., 2009; Van Kranendonk, 2011b).

Although their primary mantle melts are probably more primitive (O'Hara and Herzberg, 2002; Herzberg and Rudnick, 2012), evolved tholeiitic basalts are the predominant constituent of most greenstone belts, with an average MgO of ca. 7 wt.% computed for the Abitibi (Kamber, 2015). As these basalts are not in equilibrium with peridotitic mantle, and follow low-pressure cotectic fractional crystallization paths (O'Hara and Herzberg, 2002; Leclerc et al., 2011), large volumes of complementary crustal cumulates are required. The most probable place to situate this differentiation is in sill-like magma chambers (Bédard et al., 2009); similar to those

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Figure 4. Log Zr/Nb vs. Nb/Th plot. The OPB (oceanic plateau basalt) and OIB (ocean island basalt) fields are from Condie (2005) whereas the N-MORB (or depleted basalt) and Arc (or hydrated basalt) fields are from Condie (2015). Canonical PM (primitive Mantle), DM (depleted mantle), uc (upper crust), DEP (depleted plume component), and the recycled OIB components EM1, EM2 and HiMu, are from Condie (2005). The grey field encloses all Archaean basalt data shown in Condie (2015). Roy Group Neoarchaean volcanic data (Leclerc et al., 2011, 2016) are from the Chibougamau area, Abitibi belt. (A) Shows tholeiitic Roy Group basaltic lavas. Note the relative rarity of N-MORB and OIB magma signatures, and the overwhelming predominance of PM-like basalts in the Archaean population. A few Roy Group basalts scatter towards a DEP component. Iceland field from Condie (2005). (B) Shows Roy Group calc-alkaline volcanics, and the Type III Rhyolite suite volcanics from the Lemoine member of the Waconichi Formation, Roy Group. These scatter into the 'arc' field and towards the REC OIB components.

thought to be responsible for fractionation of many oceanic plateau basalts (Kerr, 2014), continental flood basalts (Thompson, 1972; Cox, 1993; Bedard et al., 1984, 2012), and MORB-like basalts at spreading ridges (Dicket al., 1991; Bedard, 1993, 2015; Maclennan

et al., 2001 ). However, reflection seismic images and gravity models show that thick accumulations of lower crustal cumulates are absent from the base of dominantly felsic Archaean crust (Artemieva, 2009; Abbott et al., 2013; Yuan, 2015); whereas estimates of Phanerozoic lower crust are markedly more mafic (Rudnick and Fountain, 1995; Hacker et al., 2015), in part due to the continued presence of underplating mafic intrusions and cumulates (Quick et al., 1994; Farnetani et al., 1996; Miura et al., 2004; Sinigoi et al., 2011). Where then are the lower-crustal cumulates needed to complement the evolved compositions of Archaean basalts? Modelling results suggest that any large masses of dense cumulate (or restite) that formed probably foundered spontaneously into the low-density underlying mantle (Herzberg et al., 1983; Vlaar et al., 1994; Jull and Kelemen, 2001; van Thienen et al., 2004a,b; Behn and Kelemen, 2006; Herzberg, 2014; Sizova et al., 2015; Fischer and Gerya, 2016).

It is argued here that isotopically-juvenile greenstone belts like the Abitibi represent the best existing samples of Archaean oceanic crust (Bédard et al., 2013; Harris and Bédard, 2014a; Kamber, 2015; Thurston, 2015). As komatiites are also present there, the involvement of deep-rooted, hot, mantle upwellings seems required. It will be argued in a later section that batches of Archaean oceanic crust were generated during major, long-lived (ca. 100 Ma) mantle overturn/resurfacing events characterized by anomalously high mantle melt fluxes and temperatures. Mantle melt fluxes would have decreased markedly in the intervals between overturns, when newly-formed Abitibi-like oceanic crust would have lapsed into a more quiescent, stagnant-lid ss phase that resembles predictions of heat-pipe models (O'Reilly and Davies, 1981; Turcotte, 1989; Moore and Webb, 2013).

7. Where do Archaean calc-alkaline magmas come from?

Another important distinction between Archaean greenstones and Phanerozoic oceanic plateaux is that the latter typically do not have substantial intervals of calc-alkaline eruptives and swarms of diapiric TTG intrusions. Archaean calc-alkaline volcanics, LREE-enriched basalts with negative Nb—Ta anomalies, and rare boni-nitic lavas that are intercalated with the 'non-arc' tholeiites and komatiites are an important point of contention in the different interpretations of how Archaean greenstones formed. Many believe that the existence of such magmas requires a subduction environment (see Hamilton, 1998; and Bédard et al., 2013 for critical reviews), forcing the overall stratigraphy to have formed through complex plume/arc/backarc interactions (e.g. Wyman, 1999, 2013; Jenner et al., 2009; Wyman and Kerrich, 2010; Arndt, 2013). The contrasting explanation for intermittent pulses of calc-alkaline magma is that a high basalt/komatiite flux allows lower crustal rocks to be remelted episodically (Campbell and Hill, 1988; Bédard, 2006; Van Kranendonk et al., 2007b; Smithies et al., 2009; Barnes et al., 2012; Bédard et al., 2013; Barnes and Van Kranendonk, 2014; Thurston, 2015; Kamber, 2015); an inference consistent with thermal modelling results indicating that large-scale emplacement of crustal sills can trigger extensive lower crustal anatexis (Huppert and Sparks, 1988; Bergantz, 1989; Sandiford et al., 2004; Dufek and Bergantz, 2005; Annen et al., 2006; Annen, 2011; Solano et al., 2012).

As discussed above, the clear differences in volcanic facies and geochemical signatures between Archaean calc-alkaline sequences and modern arc suites (e.g. Fig. 2) appears to be inconsistent with the generation of Archaean calc-alkaline sequences in modernstyle active subduction zones. A common objection to non-subduction lower crustal anatexis models for Archaean calc-alkaline suites (Arndt, 2013; Wyman, 2013; among others) is a perceived problem with the amount of water needed to generate

large (continent-forming) volumes of felsic magma. Arndt (2013) believed that the middle and lower crust of Archaean terrains could not have yielded much felsic melt because, like the Ontong-Java Plateau, they were dominated by refractory anhydrous cumulate rocks. Wyman (2013) made a similar point and questioned how the crust can experience penetrative hydrothermal metamorphic overprints as: ' a large proportion of plateau thickening occurs via the emplacement of sills at a wide range of depths within the edifice ... which will not undergo the same type of alteration found in the upper parts of normal oceanic crust.'

These objections to lower crustal basalt anatexis models are invalidated by the following observations and arguments:

(1) In contrast to the static oceanic plateau crust model of Arndt (2013), Archaean crustal sections were probably extremely dynamic, and there is good evidence from the East Pilbara and Superior cratons (and modelling) for rapid rates of sagduction that would bring seawater-altered basalts down to the Moho (Collins and Van Kranendonk, 1999; Lin et al., 2013; François et al., 2014; Johnson et al., 2014; Sizova et al., 2015; Fischer and Gerya, 2016). Rapid descent of cold negative basaltic di-apirs explains the survival of water-rich greenschist facies metabasaltic cores within km- to decimetre- scale supracrustal slivers that have amphibolite-grade margins (Bédard, 2003). Simply put, the field evidence implies that intra-crustal negative diapirs descended too fast for complete thermal equilibration with their hotter surroundings. Thus, it is plausible to suggest that even relatively low-grade, water-rich meta-volcanics may have been delivered in large volumes to the base of the Archaean crust through partial crustal convective overturn, and would yield abundant TTG magma when next underplated by mafic/ultramafic sills.

(2) Although many Archaean basaltic formation-scale units formed as high-volume, rapid-effusion events; magmatism was intermittent, as demonstrated by the presence of thick intercalated BIF (banded iron formation) and chert sequences (Ayer et al., 2002; Thurston et al., 2008, 2012). During the long intervals between magmatic pulses, water would have penetrated deeply into the fissured, porous submarine volcanic pile. When new magma pulses eventually arrive, thousands or millions of years later, feeder intrusions would drive the intense hydrothermal circulation that formed Archaean vol-canogenic sulphide deposits (Gibson et al., 1999; Hannington et al., 2003; Drieberg et al., 2013). By comparison, earthquakes are recorded in the mantle immediately beneath modern slow-spreading ridge axes (Toomey et al., 1988), which requires that seawater be able to penetrate to mantle depths (>10 km) between magmatic pulses that are separated only by a few thousand years. Consequently, extensive rehydration of Archaean greenstone crust in the intervals between volcanic pulses separated by millions of years seems likely.

(3) Abitibi belt lower crustal (>30 km in depth) garnet granulites and amphibolites exposed in the Kapuskasing structure show direct evidence of hornblende-breakdown anatexis (Hartel and Pattison, 1996), with tonalitic leucosomes having ages that correspond to TTG emplacement at higher levels (Benn and Kamber, 2009). A similar lower crustal source of TTG is favoured in many other greenstones (e.g. Kroner et al., 2013). Mass balance and geochemical arguments require that the extraction of each gram of felsic melt from a basaltic source must generate 3 to 4 times as much restite (if anatectic) or cumulates (if they are fractionation residues) (Arndt and Goldstein, 1989; Rapp and Watson, 1995; Rudnick, 1995; Foley et al., 2000; Bédard, 2006, 2013; Zhang et al., 2013). Thus, even a small amount of TTG melt formation in the lower crust implies that abundant

restites and/or cumulates composed of dense pyroxenites and garnet pyroxenites must also have formed. Basalt and gabbro emplaced near the base of a thick crust would also transform into dense, gravitationally unstable, garnet-rich metamorphic assemblages (e.g. Lyons et al., 2007; Johnson et al., 2014). Thus, any substantial thickness of lower crustal cumulates, solidified intrusions or anatectic residues that did form should have quickly foundered (Herzberg et al., 1983; Vlaar et al., 1994; Ducea and Saleeby, 1998; Jull and Kelemen, 2001; van Thienen et al., 2004a,b; Bédard, 2006; Behn and Kelemen, 2006; Herzberg, 2014; Sizova et al., 2015; Fischer and Gerya, 2016), allowing hot mantle (and mantle melts) to well up and underplate fusible lower-crustal metabasaltic and felsic rocks. Thus, lower crustal reworking and cumulate/restite delamina-tion seem the simplest way to account both for frequent generation of subordinate Archaean calc-alkaline magmas, and for the absence of lower crustal cumulates and restites.

Wyman (2013) argued that remelting the base of a static basaltdominated plateau could not produce batholith-scale TTG provinces (i.e. continental crust), stating: 'If the mafic intrusions were emplaced into seawater-altered crust, then it is likely that a series of small volume melts might occur at a range of depths associated with successive intrusions rather than the batholith-forming processes required by the monodynamic model' In fact, Wyman's (2013) predicted outcome closely resembles what is observed in the Abitibi belt, which has small-volume felsic and calc-alkaline volcanics scattered throughout a basalt-dominated stratigraphic sequence. Abitibi felsic melt chemistry indeed reflects melt generation at a variety of depths, with type I-II dacites/rhyodacites originating from the lower crust where garnet is stable in the residue, whereas Type III sodic-rhyolites reflect shallower melting conditions (Thurston and Fryer, 1983; Lesher et al., 1986; Leclerc et al., 2011; Bédard et al., 2013). Although it is referred to as a greenstone belt, the Abibiti is in fact composed of >50% felsic intrusions, many of which are geochemically correlatable to the felsic extrusive rocks (Fig. 2D), and which could have been derived from anatexis of Abitibi-like basalts (Bédard et al., 2013).

The anomalously productive calc-alkaline Shebandowan segment of the Abitibi belt (Lodge, 2016) might reflect the presence of a pod of older felsic crust in its source, as some dacites there have unusually low eNd 2720 Ma values (+0.6 and -0.2). Isolated domains with older material are known to be present elsewhere in the Abitibi belt's lower crust (Benn and Kamber, 2009; Thurston, 2015). Having a mass of fusible older tonalite or meta-rhyolite in the lower crust would evidently increase the proportion of felsic magma generated during anatexis, creating magma mixing-homogenization zones and a crustal filter that could block ascent of coeval basalt (cf. Hildreth, 1981; Hildreth and Moorbath, 1988; Leclerc et al., 2011), potentially yielding bimodal tholeiite/calc-alkaline volcanic suites. Contemporary examples have been documented of continental rafts stranded in ocean basins during rifting (Weis and Frey, 2002; Funck et al., 2003; Begg et al., 2009; Dean et al., 2015), and I speculate that a felsic crustal raft was stranded under the Shebandowan area (Lodge, 2016) when the Abitibi Sea opened (Bédard and Harris, 2014).

I would argue that craton-forming (batholith-scale) Archaean felsic pulses characterize crustal domains where there was already a lot of felsic material available to be remelted (unlike the Abitibi); an inference confirmed by isotopic data and the ubiquitous inherited zircon cores in North-East Superior TTGs and granites (Bédard, 2003,2006; Bédard et al., 2003; Boilyetal., 2009; Maurice et al., 2009). It is inferred that continental growth and stabilization is inherently multi-cyclic. The first reheating events would generate small amounts of felsic melt from fusible lower crustal

Figure 5. Plan view cartoon showing indentation of Southern Superior accreted terranes by a southwardly-drifting composite continent. The figure is generic, and corresponds to ca. 1000 km x 500 km. Continental crust shown in orange, oceanic crust in pale green. Complex deformation patterns formed ahead of a complex indentor are shown schematically as rifts and fold traces. The Abitibi oceanic crust is interpreted to have been extruded towards the East, in accord with the structural interpretation of Harris and Bedard (2014a). Major compressional faults formed ahead of the indentor prevent ascent of gold-carrying metamorphic fluids. When the Abitibi was extruded, major faults would have shifted from compressional (circled c) to transtensional structures (circled e), allowing sudden fault-guided ascent of the metamorphic fluids, so explaining the association of late-orogenic lode gold deposits (stars) with major faults. Pressure shadows behind sialic microcontinents also represent favourable sites for gold deposition (cf. Leclerc et al., 2012).

metabasaltic domains that may or may not segregate into discrete intrusions, and may not erupt. Subsequent underplating magma pulses would generate new primary TTG melt from metabasalt, but also preferentially remelt existing TTG veins and bodies. With time, the proportion of felsic material in the crust would increase to the point where the crust becomes 'continental' (Bedard, 2006). A multi-cyclic model explains why entrained restitic garnet is so rarely observed in TTG magmas that have geochemical signatures requiring residual garnet. Although the volume of felsic material generated during an individual crustal reworking event may be roughly proportional to the amount of felsic material already present in the crustal source volume, a craton-scale TTG-forming pulse would still require a very large mafic/ultramafic magmatic under-plate to provide anatectic heat (Raia and Spera, 1997; Milidragovic et al., 2014). Evidently, the Earth is not now covered with cratons, so 'normal' Archaean oceanic crust would only rarely have evolved into continental crust. As argued by Bedard (2006, 2013), continental crust would have formed above long-lived mantle upwelling instabilities large enough to build a thick volcanic pile that could cannibalize its own base to yield voluminous gravitationally unstable pyroxenitic to eclogitic restites, delamination of which triggered new mantle melting events, leading to coupled autocatalytic differentiation of the continental crust and SCLM.

8. Subduction or subcretion?

Much of the controversy about when plate tectonics began can be mitigated if an explicit distinction is made between active subduction and subcreting/imbricating convergent margins. The lower plate is the active agent in modern convergent margins

characterized by active subduction, as negatively buoyant oceanic lithosphere sinks into the aesthenosphere under its own weight (Fig. 1A). In contrast, the lower plate is almost completely passive during subcretion and imbrication, as unsubductable buoyant oceanic or continental lithosphere is partly overridden by advancing continents (Fig. 1C). Bedard et al. (2013) proposed that most Archaean convergent margins were not characterized by active subduction, but by subcretion and imbrication of buoyant stagnant-lid oceanic lithosphere ahead of drifting continents. This subduction-less scenario provides a self-consistent explanation for the generation of Archaean compressional tectonics and terrane accretion, accounts for the distinctive Archaean magmatic, strati-graphic, metamorphic and structural patterns, and accommodates the presence of subordinate source-metasomatic arc-like magmatic suites (Fig. 1C). Subcretion as envisaged here is kinematically similar to some types of modern convergent margins where large amounts of supracrustal rocks relaminate the upper plate (Barr et al., 1999; Grove et al., 2003a; Ducea et al., 2009; Hacker et al., 2015; Taramon et al., 2015).

Buoyant Archaean oceanic crust lacking a SOLM should grade from a cold upper section to a hotter and softer lower crust, implying that hot-orogen models with wide tectonic forelands would better describe their deformation style (Choukroune et al., 1995; Benn, 2006; Cagnard et al., 2006; Rey and Coltice, 2008; Harris and Bedard, 2014a,b). Ongoing convergence would be accommodated largely by flow in the lower crust and mantle beneath docked terranes, with the brittle-ductile carapace riding on the ductile substrate (Chardon et al., 2009, 2011). Although this type of crust is probably too weak to maintain significant long-term relief (England and Bickle, 1984; Rey and Houseman, 2006), the

upper crust would be colder and stiffer as a result of hydrothermal cooling and could thicken dynamically as duplexes that would occasionally emerge and be eroded, providing an explanation for the development of major syn-orogenic metasedimentary belts (the sedimentary sub-provinces of Card and Ciesilski, 1986) that separate some accreted Archaean terranes (Bédard et al., 2013; Bédard and Harris, 2014). The imbricated/subcreted package predicted to form ahead of a drifting continent (Figs. 1C and 5) is very similar to the imbricated Archaean orogenic geometry previously proposed by many (Oxburgh, 1972; Helmstaedt and Schulze, 1989; deWit et al., 1992; Smithies et al., 2003; Canil, 2004; Griffin et al., 2004; Lee, 2006; Griffin and O'Reilly, 2007; Lee et al., 2008; Van Kranendonk, 2010), with large volumes of altered basalt introduced to depths >30—40 km allowing formation of correspondingly large pulses of anatectic syn-kinematic TTG. The imbricated structure of the southern Superior craton (Calvert et al., 1995; Calvert and Ludden, 1999; White et al., 2003; Benn, 2006) could have formed in this way (Grey and Pysklywec, 2010; Bédard and Harris, 2014), as might the multiple fault-bounded 'terranes' and structural panels of the North Atlantic craton (Windley and Bridgewater, 1971; Bridgwater et al., 1974; Myers, 1984, 2001; Nutman and Friend, 2007, 2009; Windley and Garde, 2009; Dziggel et al., 2012; Polat et al., 2015), the mafic accretionary system that formed north of the Kaapval craton (Eglington and Armstrong, 2004; Zeh et al., 2009), and the accreted West Pilbara terranes (Smithies et al., 2005b; Hickman and van Kranendonk, 2012). Imbrication and subcretion to a drifting continent (Bédard and Harris, 2014) can explain the systematic polarity and short collision interval of terrane docking events documented in the Southern Superior (Percival et al., 2006), and the shift to thicker, less felsic crust in younger parts of Australian cratons (Yuan, 2015). Because a moving continent is unlikely to have had a straight leading edge, promontory-reentrant geometries would create complex structural patterns in the tectonic foreland, possibly involving terrane extrusion (Harris and Bédard, 2014a,b; Fig. 5). The deep sections of imbricated foreland terranes would generate gold-carrying metamorphic fluids in abundance (Goldfarb and Groves, 2015), but the compressional regime ahead of indentors would seal faults and hinder ascent of metamorphic fluids. As foreland terranes escape from an indentor's compressional footprint (Harris and Bédard, 2014a), faults would shift from being compressional to extensional, allowing metamorphic fluids to drain out of their lower crustal sources en masse, potentially explaining the association of late-kinematic Abitibi belt lode gold deposits with major strike-slip faults (Fig. 5).

Blueschist facies rocks mostly form when cold oceanic crust subducts at high convergence rates (Ernst, 1988; Peacock, 1996; Stern, 2005; Brown, 2009). Because this facies forms at shallow levels (<40 km), the temperature of the mantle is largely irrelevant, such that secular changes in mantle temperature cannot explain the absence of blueschists in the deep past. Palin and White (2016) argue that the high MgO content of Archaean oceanic crust formed from a higher-temperature mantle precludes stabilization of blueschist facies metamorphic assemblages, and that consequently, the absence of blueschists from the ancient record (Stern, 2005) cannot be used as an argument against the existence of Archaean subduction. Although computed primary melt compositions of Archaean non-arc basalts do indeed have compositions between 15 and 25 wt.% MgO (Herzberg and Rudnick, 2012), the average greenstone basalt typically has only 7 wt.% MgO (Kamber, 2015), such that blueschist facies assemblages could form if these rocks were subjected to high-pressure/low-temperature conditions. Thermobarometric data on Archaean 'high-pressure' rocks (Collins and Van Kranendonk, 1999; Moyen et al., 2006; Brown, 2009; Van Kranendonk, 2011b; François et al., 2014) show that they are not in

the blueschist facies, but preserve P—T conditions that are consistent with either partial convective overturn or transient compres-sional crustal thickening. A potential explanation for the absence of ancient blueschists is that stagnant-lid oceanic crust would not subduct out of the way upon the approach of a drifting continent (in contrast to most oceanic lithosphere today) but would resist, decreasing convergence rates to the point where blueschists could not form.

Modern active subduction is a plate-driving force because the tensional stress generated by slab-pull can be transmitted through a stiff oceanic plate (Turcotte and Schubert, 2002; van Thienen et al., 2004a,b). An induced-convection stagnant-lid system would probably not develop or preserve a SOLM (see above, Johnson et al., 2014); and so would be too weak to transmit the tensional stresses needed to create a subduction/seafloor-spreading couple. Thus, if the Archaean oceanic lithosphere lacked a SOLM layer, there should have been no organized oceanic plate behaviour. If underthrust to depths where eclogitization of the crust becomes a significant force, weak Archaean oceanic crust would tend to drip off (van Hunen and Moyen, 2012). Perhaps enough of the force generated by eclogitization could be transmitted to keep the lower plate moving, leading to episodic drips from the leading edge of an actively advancing oceanic slab as suggested by van Hunen and Moyen (2012) and Moyen and van Hunen (2012)? Whether an Archaean oceanic slab could be drawn down this way probably depends on whether a stiff SOLM layer was present, as was assumed by van Hunen and Moyen (2012) and Hynes (2014). As Earth cooled and the background oceanic lithosphere stabilized a SOLM layer, there would be an increasing probability that oceanic lithosphere might start to actively subduct when subjected to compression.

Active subduction generates island arc and continental arc magmas with source-metasomatic geochemical signatures (Fig. 2A, Davidson, 1996; Pearce, 2008; Bédard et al., 2013; see above). Such source-metasomatic signatures are rare in Archaean packages, with most Archaean calc-alkaline suites having oblique trends indicative of crustal contamination (Fig. 2B—D, Pearce, 2008; Barnes et al., 2012; Bédard et al., 2013; Barnes and Van Kranendonk, 2014). During subcretion beneath the leading edge of a drifting continent (Fig. 1C; Bédard et al., 2013), the downthrust Archaean oceanic crust would not generally be overlain by a convecting mantle wedge, and so could not generate magmas with arc-like source-metasomatic geochemical signatures. Instead, oceanic slabs would mostly stack up ahead and beneath a refractory SCLM. After thermal relaxation, heated, subcreted metabasalt would form anatectic TTGs, soaking up most available volatiles. Volatiles and TTG released from subcreted basalt that impregnate the overlying SCLM may emerge later as sanukitoids (Laurent et al., 2011, 2014). Geochemical and isotopic data imply that some Archaean potassic shoshonites and sanukitoids formed through introduction of K2O-rich sediment and/or TTG-like melt into the mantle (Wyman and Kerrich, 1993; Laurent et al., 2011, 2014; Wyman, 2013); and it is commonly argued that this is the result of subduction. I agree that formation of sanukitoids in the late stages of craton stabilization reflects introduction of fusible water-bearing material into the mantle at convergent margins, but the bulk of the evidence suggests these were not sites of active subduction (as defined above). It could also be argued that peri-continental subcretion-imbrication (Bédard et al., 2013; Fig. 1C) is more effective at introducing the K-rich components needed for sanukitoid genesis into the shallow mantle source region than active subduction.

Laurent et al. (2014) reviewed the progressive evolution seen in many cratons from a long TTG stage to a short terminal pulse of extensive crustal reworking and the appearance of true granites (Bickle et al., 1993; Bleeker, 2002; Boily et al., 2009; Maurice et al.,

2009). The late granites are coeval with sanukitoid and shoshonitic magmatism in the youngest cratons. Laurent et al. (2014) argued that the first appearance of sanukitoids at ca. 3 Ga records the beginning of modern-style subduction—collision processes. I argue instead that the appearance of sanukitoids records emergence above sea-level and erosion of mature continental crust containing high-K2O granites, as only erosion of granite can create sediments with high K2O contents. Flament et al. (2013) showed that nascent cratons are commonly capped by extensive basalts and are too hot and weak (Bailey, 1999; Flament et al., 2011) to allow felsic plutons to be eroded in any volume. For a continent to emerge above sea level and be eroded, the lower crust must be rigid enough to not collapse spontaneously. One way to strengthen the lower crust is by extracting radioactive elements (U—Th—K) through repeated reworking. I suggest that the absence of pre-3 Ga sanukitoids reflects a higher radioactive heat productivity in the early Archaean that retarded crustal rigidification, continental emergence and erosion (cf. Vlaar, 2000).

The proposed existence of Archaean and Hadean boninites is also claimed to demonstrate early Earth subduction and plate tectonics (Boily and Dion, 2002; Frei et al.,2004; Hoffmann etal.,2010; Adam et al., 2012; Wyman, 2013; Turner et al., 2014). Some Archaean boninitic magmas may have formed by volatile fluxing of a depleted source (see discussion in Smithies (2002) and Smithies et al. (2004)), as did Phanerozoic boninites (Crawford et al., 1989; Bedard, 1999; Page et al., 2009). However, most Archaean 'bonin-ites' are significantly less depleted in incompatible elements than are their Phanerozoic cousins (Fig. 6); which implies that the mantle that generated Archaean boninitic magmas was generally less depleted than the sources of Phanerozoic boninites. The extreme depletion of Phanerozoic boninite mantle sources is attributed to a prior stage of extensive, volatile-fluxed melting in subduction zone environments (Pearce and Peate, 1995; Grove et al., 2006; Page et al., 2009). Consequently, the lesser depletion of Archaean 'boninites' would imply that the source mantle from which they originated did not form as residues of extensive wet melting in subduction zone environments. Instead, I suggest that eclogitization of oceanic slab leading edges in an imbricated accretionary stack creates a geometry that permits localized source-metasomatic mantle flux melting (Fig. 1C), so accounting for the intermittent appearance among Archaean suites of source-metasomatic calc-alkaline suites and occasional boninitic magmas (Harris and Bedard, 2014a). Specifically, as a downthrust basaltic package crosses the garnet-in reaction depth, density increases would steepen the downthrust angle and open up space allowing hot mantle to well up (Fig. 1C). Release of fluids from the steepening metabasalt slab into the upwelling mantle could generate magmas with arc-like source-metasomatic geochemical signatures. If the upwelling mantle had previously lost a melt fraction, it would show more depleted trace element contents and form Archaean boninitic magmas. Because these downthrust eclogitized oceanic slabs would lack a thick SOLM, they would be extremely weak and would tend to drip off (vanThienen et al., 2004a,b; van Hunen and Moyen, 2012), providing another vector for the transfer of fluids into hot upwelling mantle.

It is vitally important for the debate on when subduction started that the term 'subduction' not be used as a loosely-defined catch-all for any process that transfers surface rocks to depth. For example, it is often affirmed that Archaean-age mantle eclogite xenoliths from the SCLM must have formed through oceanic crust subduction because they have: (a) low-temperature stable isotopic signatures that imply interaction with seawater; and (b) major and trace element evidence which indicates a gabbroic cumulate protolith (Neal et al., 1990; Jacob et al., 1994; Schulze et al., 2000; Barth et al., 2001; Richardson et al., 2001; Farquhar et al., 2002; Zack et al.,

2003; Shirey and Richardson, 2011; Smit et al., 2014a,b; Burnham et al., 2015). This interpretation is non-unique, however, as delamination/foundering of metabasalt and associated gabbroic intracrustal feeders into the upper mantle (Bedard, 2006), and imbrication and subcretion of basalt-dominated stagnant-lid crust during continental drift/accretion events (Bedard et al., 2013), both provide effective, non-subduction mechanisms for the introduction of these geochemical signatures into the mantle. As the altered basalts in Archaean greenstone belts are commonly interlayered with chemical sediments, there is the potential to also transfer anomalous Li—Be—Pb—W-isotopic signatures and U—Pb concentrations into the upper mantle without subduction. Large-scale recycling into the mantle of hydrothermally altered oceanic crust and chemical sediments through delamination, subcretion and overturns (see below) may explain some of the Pb-isotopic paradoxes discussed by Shirey et al. (2008), Kamber (2015), and Castillo (2016); and harks back to proposals by Albarede and Michard (1986) and Albarede (1998) for recycling into the mantle of altered seafloor rocks.

9. Mantle overturns

O'Neill et al. (2007, 2013) proposed that the early Earth was a stagnant-lid planet punctuated by periodic overturn/resurfacing events (cf. Stein and Hofmann, 1994; Davies, 1995; Griffin et al., 2014), while also being pummelled by bolides (Simonson et al., 2009; Glikson and Vickers, 2010; Bottke et al., 2012; Marchi et al., 2014). This concept provides a rationale for many first order features of Hadean and Archaean geology. As extremely high decompression-melting fluxes are predicted for overturn upwell-ing zones (OUZOs, Fig. 3B), then these are inferred to be the most likely places to form and rework a thick mafic crust, in accord with in-situ crust maturation models for craton genesis (Campbell and Hill, 1988; Bedard et al., 2003; Smithies et al., 2005a; Bedard, 2006; Champion and Smithies, 2007; Van Kranendonk et al., 2007b; Sizova et al., 2015; Fischer and Gerya, 2016). This periodic overturn system would have continued until secular decay of radioactive isotopes cooled the mantle to the point where heat loss could keep pace with heat production. Since active-lid tectonics (subduction-seafloor spreading) is more efficient than stagnant-lid systems at extracting heat from the mantle (Moresi and Solomatov, 1998), the initiation of modern-style plate tectonics should have increased Earth's cooling efficiency and either aborted subsequent overturn events or lengthened the intervals between them (cf. Condie and Aster, 2010; Condie et al., 2016). Previously, large-volume mantle upwellings or superplumes were attributed to avalanches of subducted slab megaliths that had temporarily accumulated at the upper/lower mantle interface (Stein and Hofmann, 1994; Albarede, 1998; Condie, 1998; Nelson, 1998; Bleeker, 2003; Machetel and Humler, 2003; Pysklywec et al., 2003; Davies, 2008). However, revised estimates of the Mg-perovskite reaction Clapeyron slope (Katsura et al., 2003; Fei et al., 2004) no longer favour the accumulation of transition zone slab megaliths. Although the end-results of avalanche-driven upwellings and overturn-driven upwellings are somewhat similar, the mantle overturn scenario has a different driving force and involves up-wellings that are larger in scale and much longer-lived (O'Neill et al., 2007, 2013; Griffin et al., 2014).

Although overturn scenarios provide a rationale for the scale and longevity of mantle upwelling events in the Archaean, there are many uncertainties as the numerical models are sensitive to boundary conditions, thermo-physical constants, constituting equations, and terminology. O'Neill et al. (2007, 2013) suggested that the early Earth experienced alternations between stagnant-lid and subduction phases, a pattern they labelled 'episodic

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subduction'. However, in their models the crust is efficiently coupled to mantle flow and accumulates above downwelling zones to form large masses that go catastrophically unstable as a result of their low temperatures and eclogitization. The term 'subduction' is especially misleading as the downwellings do not resemble modern-style active subduction, and the entrainment and buildup of lid material reflects a strong mechanical coupling between crust and mantle that is built into the model. In whole-mantle scale models it is nearly impossible to incorporate the fine-scale complexities of boundary layers. Seismic work at active ridges (e.g. Garmany, 1989) and studies of ophiolites (e.g. Bedard, 1993, 2015), exposed continental lower crust (Quick et al., 1994; Sinigoi et al., 2011), as well as geochemical and petrological studies of highvolume basaltic events (e.g. Cox, 1993), all show that the Moho mechanical and density discontinuity is commonly exploited by abundant sill-like intrusions. If Archaean Moho was similarly lubricated as a result of high melt productivity during overturns, then the lithosphere would not be efficiently mechanically coupled to the subjacent flowing aesthenosphere, and so the crust would not be entrained very efficiently. This suggests that overturns would mainly have been caused by inefficiencies in how the Earth evacuated heat during the stagnant-lid intervals (Fig. 3). Although the non-dimensionalized models of O'Neill et al. (2007, 2013) do not exactly constrain timescales of overturn and repose, they do suggest that overturn events would be long-lived with correspondingly long repose times so as to allow enough heat to build up to trigger the next overturn. If the Neoarchaean Superior reworking event is the result of a mantle overturn (Bedard and Harris, 2014), then this overturn would have lasted ca. 120 Myr (between ca. 2800 Ma and ca. 2680 Ma). If it is correct to attribute most large-scale crustal recycling events to mantle overturns, then the age peaks in the zircon record provide an empirical constraint on overturn periodicity (Griffin et al., 2014), even though the size of preserved zircon age peaks would shift with time as a result of reworking.

The oldest Hadean zircons attest to a crust-forming event at ca. 4.34 Ga, followed by intermittent small-degree remelting (Kemp et al., 2010; Bell et al., 2014; Nebel et al., 2014). The evidence for intermittent small-scale magmatic-thermal pulses that did not destroy the overall coherency of the crust is consistent with evolution during a stagnant-lid interval. The first major overprinting event occurred at ca. 3.9—3.85 Ga, with the second at ca. 3.35—3.4 Ga (Bell et al., 2011, 2014; Griffin et al., 2014; Nebel et al., 2014). These events were formally attributed to mantle overturns by Griffin etal. (2014); who suggested that the period ca. 3.1—2.9 Ga coincided with the rapid buildup of the SCLM recorded by Os model ages, which peaked at 2.7 Ga (Pearson et al., 2007; Griffin et al., 2014). Griffin et al. (2014) interpreted this to signify that most of the SCLM formed after 3 Ga as a result of larger, more rapid mantle overturns. They also suggested that the rarity of crustal zircon age data in the 3.1—2.9 Ga interval is the result of large-scale reworking of crust caused by mantle overturn activity; with the 2.9—2.5 Ga period reflecting continued mantle overturns and the beginnings of plate tectonics.

While the evidence is consistent with alternating mantle overturn/stagnant-lid phases between 4.34 and 2.5 Ga, I do not believe that plate tectonics began much before 2.5 Ga, mainly

because that age is the time when major isotopic systems started to evolve differently (Valley et al., 2005; Hawkesworth et al., 2010), and corresponds to the major break in the rock record (Hamilton, 1998; Bédard et al., 2003, 2013; Keller and Schoene, 2012). In essence, greenstone + TTG packages formed at 2.73 Ga are almost indistinguishable from those generated at 3.6 Ga, and many other facets of Archaean geology seem inconsistent with the existence of widespread early active subduction and seafloor-spreading, as discussed previously (Bédard et al., 2013).

If the stagnant-lid/mantle overturn scenario outlined above is broadly correct, then early Earth stagnant-lid cooling intervals may also correspond to periods of layered mantle convection. This has important implications for geochemical Earth evolution models. Early mass balance geochemical box models for the generation and evolution of Earth's major geochemical reservoirs regarded the depleted MORB-source mantle (DMM) as the direct complement to continental crust extraction (Allègre et al., 1979,1983a,b; Jacobsen and Wasserburg, 1979; DePaolo, 1980; O'Nions et al., 1980; see review in White, 2015). Layered mantle convection was invoked to satisfy mass balance and allow survival beneath the DMM reservoir of a near-primitive lower mantle source of OIB (Stein and Hofmann, 1994; O'Nions and Tolstikhin, 1996). Identification of the modern OIB source as true Primitive Mantle (PM) has now been abandoned because OIB (and MORB) are now recognized as complex mixtures of depleted and enriched components (Hofmann, 2003; Stracke, 2012; Castillo, 2015, 2016; White, 2015). Survival into the present of a true primitive mantle source is no longer favoured because geochemical and isotopic data require a more complex early Earth history, involving early mantle depletion/crust formation events coupled with either sequestration of a hidden enriched reservoir (EER = early enriched reservoir) in the deep mantle (Boyet and Carlson, 2006; Tolstikhin et al., 2006; Shirey et al., 2008) or its loss during accretion impacts (Caro et al., 2008). Removal of the EER was followed by the Moon-forming giant impact at ca. 4.47 Ga (Bottke et al., 2015), formation and solidification of the second magma ocean, and then possibly by a mantle overturn event that yielded the oldest preserved fragments of Hadean oceanic crust (Shirey et al., 2008).

Modern seismic data show that most of the middle and upper mantle is traversed by subducting slabs and hot upwelling zones, precluding the existence of stable layered mantle convection today (Jordan et al., 1993; French and Romanowicz, 2015), and contributing to the abandonment of 2-reservoir layered mantle evolution models. However, the geophysical data, although conclusive, cannot preclude the existence of episodic layered mantle convection in the Archaean and Hadean, when Earth was hotter and probably convected differently. An episodic mantle overturn model allows a fresh look at how geochemical reservoirs might have evolved in the Hadean and Archaean. Overturn upwelling zones (OUZOs) are predicted to create high heat flows and yield a high flux of mafic-ultramafic magma. Older oceanic or continental crust situated above upwelling zones should therefore have experienced large-scale reworking and/or resurfacing. Felsic anatectic rocks generated by lower crustal melting during overturns would be difficult to recycle into the mantle and would contribute to secular growth of continental crust. As each overturn would rework basaltic crust and generate new felsic crust, this would produce

Figure 6. (A) Yb (ppm) and (B) TiO2 (wt.%) vs. MgO (wt.%) of Archaean magmas with claimed boninitic affinities. Colours are coded by continent. Background fields defined by Betts Cove ophiolite data, which has sheeted dykes and lower lavas of low-Ti boninite-ss affinity (green field), and a thick overlying unit of depleted arc tholeiites (BCDT, Mt.Misery Formation, yellow field: Bédard, 1999; Bédard et al., 2000). A unit with intermediate characteristics (BC I, blue field), mistakenly referred to as boninitic in the original publications, falls along tholeiitic trace-element melting trends (unpublished model results). Among the Archaean examples, only the most Ti-poor Whundo examples have TiO2 contents comparable to true boninites, whereas the vast majority of claimed Archaean boninites do not. A few Isua and Mallina samples overlap the upper part of the Betts Cove field for Yb, but these do not extend to typical Phanerozoic boninite abundances of ca. 0.1 wt.% TiO2 (Kerrich et al., 1998; Wyman et al., 1999; Polat et al., 2002; Polat and Hofmann, 2003; Srivastava et al., 2004; Manikyamba et al., 2005; Wyman and Kerrich, 2012).

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crustal 'growth spurts' (Dhuime et al., 2012; Naeraa et al., 2012). Conversely, delamination of dense anatectic restites and eclogitized basalt would intermittently have returned crust-like isotopic signatures and some trace-elements back into the mantle. It seems inevitable that large-scale mantle overturns would mix a substantial part of the depleted upper mantle and its overlying stagnant-lid oceanic crust with less-depleted upwelling mantle (Fig. 3B). Preservation of geochemical islands with extinct isotope signatures (Bennett et al., 2007; Rizo et al., 2012; Debaille et al., 2013) and of possible Hadean relics (O'Neil et al., 2011, 2012) show that re-homogenization was never complete, however. On the other hand, each overturn event would resurface large segments of the Earth, generating new crops of oceanic crust and complementary depleted upper mantle (Fig. 3B; Bennett, 2003; Valley et al., 2005; Kemp et al., 2010). It is inferred that the lowermost part of the mantle would be more PM-like, as it would have been less influenced by mixing with the upper DM. In this model, most Archaean basaltic and komatiitic magmas would have formed during overturns from upwelling PM-like lower mantle, with only subordinate volumes of basalt being generated between overturn events from more depleted upper mantle. Application of

this periodic rehomogenization framework to the evolution of mantle geochemistry produces results that are very similar to some older models (Fig. 7). For example, cycling between whole-mantle and layered mantle convection was postulated by Stein and Hofmann (1994) to explain crustal growth pulses and retarded development of depleted Nd isotopic mantle signatures. Similarly, I suggest that Hf—Nd isotopic stagnation during the Hadean—Archaean is a consequence of periodic overturn/mixing events that buffered the mantle near CHUR/PM (CHUR = Chondritic Uniform Reservoir), while also allowing progressive sequestration of the continental crust (cf. Allègre et al., 1979; DePaolo, 1980). Once plate tectonics started, the modern DMM source would have grown from the Neoarchaean mantle as a result of large-scale sequestration of subducted oceanic lithosphere at the core mantle boundary.

Dhuime et al. (2011) and Arndt (2013) noted that misfits between Archaean model and crystallization ages are lessened if the source mantle was less depleted than canonical depleted mantle (cf. Francis, 2003; Condie, 2011; Griffin et al., 2014); but attribute the enrichment to continuous re-introduction of a continental crust component into the mantle by subduction of continent-derived sediments. If a continental component was indeed being

Age (Ga)

Figure 7. Mantle isotopic differentiation model. Grey field and dots mark in zircon data from Hawkesworth et al. (2010) and Kemp et al. (2010). CHUR = Chondritic Uniform Reservoir, DMM = zero-age depleted MORB mantle. Two reference depleted mantle evolution curves are shown, the lower asymptotic dash-dot curve (NK) is the standard continental incremental extraction model of Nagler and Kramers (1998), whereas the upper dashed black line (G2) is the end-member MORB-DM model from Griffin et al. (2014). The yellow box at the right is an inferred CHUR-like primitive mantle. A 4.5 Ga crust-forming/upper mantle-depleting event creates a depleted upper mantle, basaltic crust, and small amounts of felsic crust. The coloured arrows illustrate trends created by incubation of melting-induced variations in parent-daughter ratios from Kemp et al. (2010), and corresponding to 176Lu/177Hf appropriate for UC (orange, upper felsic crust), BC (green, basaltic crust), and DM (purple, depleted mantle). Similar slopes are assumed for younger differentiation events, which are only shown schematically. It is inferred that each overturn (vertical red line) partly homogenized a PM-like lower mantle with an upper DM, and recycled abundant mafic crust and smaller amounts of felsic upper crust, so generating a hybrid mantle (pink boxes) that is closer to CHUR than commonly proposed for juvenile mantle melts. Although most of the depleted mantle formed by previous events is remixed into the PM-like upwelling mantle, domains of depleted upper mantle may survive, with trajectories indicated by the dashed purple arrow (Dms = depleted mantle survivor), and can account for generation of occasional 'very' depleted signatures in magmas. Reho-mogenenized mantle 3 is shown schematically as having ingested large volumes of an ultra-depleted mantle domain that survived previous homogenization events. In contrast, rehomogenenized mantle 4 is shown as having ingested very little of the older depleted mantle layer, but is contaminated by abundant recycled (delaminated) basaltic crust, and so lies closer to the CHUR line. If overturns mostly ceased at 2.5 Ga as a result of the beginning of plate tectonics, then the mantle would no longer be periodically rehomogenized and would become increasingly depleted with time, generating modern DMM because of continuous loss of basalt to the core-mantle boundary by subduction, and periodic sequestration of felsic melts in the upper continental crust. A possible 'last' overturn is shown at 1.9 Ga.

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ploughed back into the mantle on the scale needed to retard Nd and Hf isotopic evolution paths for over a billion years, then it could be argued that a large proportion of mantle-derived Neoarchaean magmas (and modern MORB) should have 'primary' arc-like trace element signatures. However, the majority of Archaean magmas that occupy the 'arc' field on the Zr/Nb vs. Nb/Th diagram (Fig. 4B), show oblique arrays on the Th/Yb vs. Nb/Yb diagram (Fig. 2B and C; Pearce, 2008) that require an origin through crustal contamination. Contaminated magmas must therefore be screened out of this discussion, allowing the geochemical characteristics of uncontam-inated melts extracted directly from ambient Archaean mantle to be examined. Abitibi basalts of the Roy Group are typical Neo-archaean 'non-arc' magmas that show few signs of shallow contamination, mostly falling within the MORB-OIB mantle melting array of the Th/Yb vs. Nb/Yb diagram (Fig. 2B, see also Bedard et al., 2013; their Fig. 7). These Roy Group tholeiitic basalts scatter between canonical PM and enriched mantle on the Condie (2005) Zr/Nb vs. Nb/Th diagram (Fig. 4A), as does the dominant population of Archaean basalts from Condie (2015). Thus, it seems more likely that common Archaean tholeiites with nearly flat chondrite-normalized trace element profiles do not sample depleted mantle that was re-enriched by subduction-mediated recycling of continental crust back into the mantle, as suggested by Dhuime et al. (2011) and Arndt (2013), but were derived directly from a mantle source that resembles PM in terms of trace element abundances. Guitreau et al. (2012) proposed a long-lived PM-like lower mantle reservoir that has steadily shrunk though geologic time. The model proposed here (Fig. 3) is similar in many respects, but posits that the PM-like, CHUR-like mantle source of most Archaean basalts was maintained dynamically throughout the Archaean as a consequence of repeated, overturn-driven homoge-nization events.

Given the PM-like trace element signature of a large proportion of Archaean basalts, it is thus plausible to consider (Fig. 7) that their mantle source was close to the chondritic evolution line and may have resembled the CHUR-like evolutionary path proposed by Stein and Hofmann (1994). If true, this undermines many existing crustal growth rate calculations as it implies that much of what is interpreted as 'recycled' because it has £Nd and £Hf isotopic signatures below the DMM evolution line (Hawkesworth et al., 2010; Iizuka et al., 2010, 2013; Condie, 2011; Dhuime et al., 2011, 2012) is directly mantle-derived and 'juvenile'. In the scenario advocated here, Archaean magmas that fall well above CHUR (Fig. 7) are not the main evolutionary stem of the bulk mantle from which modern DMM emerged, but are samples of localized or ephemeral upper DM domains (Bennett et al., 1993).

Many have argued that lateral accretion of oceanic plateau crust contributed to growth of the continental crust. In existing views (Boher et al., 1991; Kimura and Ludden, 1995; Stein and Goldstein, 1996; Abbott et al., 1997; Albarede, 1998; Arculus, 1999) accretion of plume-generated oceanic plateaux was primarily subduction-driven, with subduction magmatism reworking and refining accreted oceanic plateau crust. Alternatively, in the Archaean continental drift model (Bedard et al., 2013; Harris and Bedard, 2014a,b), unsubductable buoyant Archaean stagnant-lid 'oceanic lithosphere' is accreted to the leading edge of drifting continents. Accretion of an Archaean oceanic crust similar to that of the Abitibi greenstone belt differs from accretion of modern spreading-ridge oceanic crust and oceanic plateaux, as the Archaean sections include interbedded chemical sediments, and calc-alkaline volca-nics and plutons (cf. Ayer et al., 2002; Leclerc et al., 2011; Thurston et al., 2012; Kamber, 2015; Thurston, 2015). The accreted Archaean juvenile oceanic bulk crust would also be more evolved than their parental magmas because of pre-accretion loss of dense ultramafic cumulates and restites by basal delamination, and because many of

the basalts constituting this crust may have assimilated lower crustal felsic anatectic melts. This is consistent with trace element modelling and phase equilibrium data indicating that most Archaean TTGs were derived from basalts that were more enriched in incompatible elements than MORB (Bédard, 2006; Moyen et al., 2007; Smithies et al., 2009; Zhang et al., 2013; Martin et al., 2014; Naeraa et al., 2014). An important point to remember is that large-scale melting (to form TTG batholiths) of an accreted Abitibi-like crust would inevitably involve the calc-alkaline units as well as the tholeiites. Thus, much of the accreted Archaean oceanic crust probably had fractionated trace element inventories with LILE-enrichment and negative HFSE signatures, and would yield larger proportions of felsic melt when reworked. A weighted average of the Abitibi belt supracrustal sequence with associated syn-volcanic plutons might represent a reasonable estimate of the material added to a growing continental nucleus by lateral bulk accretion of oceanic crust in the Archaean.

The discussion has focussed on whole-mantle overturn events, but partial overturns or lesser instabilities are also possible. Griffin et al. (2014) proposed a large 2.5 Ga peak because their data contains many analyses from Australia, whereas other compilations (Condie and Aster, 2010) yield a markedly smaller peak size for this age. Does the absence of a large 2.5 Ga peak in other cratons reflect a partial overturn event beneath Australia, or is it a function of where the individual cratons were located with respect to OUZOs (Fig. 2)? Just as 2-D sheetlike upwelling zones beneath slow spreading oceanic ridges tend to break up into individual 3-D di-apirs at the shallowest levels (e.g. Ceuleneer et al., 1988; Jousselin et al., 1998; Choblet and Parmentier, 2001), the apex of Archaean OUZOs may be equally complex (cf. Romanowicz and Gung, 2002). There is much yet to be learned about the style of the upwelling and downwelling instabilities during mantle overturn events; and the transition to whole-mantle convection that characterizes Earth today.

10. Are plumes and OUZOs different?

The Archaean literature is riddled with 'plumes'. This reflects the requirement for hot upwelling zones so as to generate komatiites (Nisbet et al., 1993; Herzberg et al., 2007; Arndt et al., 2008), as this seems the only way to explain the high degrees of peridotite melting (>50%) required for their formation and the inferred presence of majoritic garnet in the source of some komatiites (Herzberg and Ohtani, 1988; Barley et al., 2000; Kerrich and Xie, 2002; Sproule et al., 2002; Arndt, 2003; Robin-Popieul et al., 2012). In plume-dominated scenarios for greenstone belt genesis, the dominant 'non-arc' tholeiites that show many resemblances to modern oceanic plateau magmas represent the products of slightly lesser degrees of melting of 'undepleted' PM-like sources or of hybrid (PM + DM) plume heads (Campbell et al., 1989; Laflèche et al., 1992; Arndt et al., 1997; Herzberg et al., 2010; Rollinson, 2010; Guitreau et al., 2012; Herzberg and Rudnick, 2012); much as was proposed for OUZO-generated basalts (Fig. 3B).

Although the argument for extensive melting in hot Archaean upwelling mantle seems strong, there appear to be differences between the products of modern plumes and the products of Archaean OUZOs; as Archaean basalts and komatiites seem to lack some of the geochemical (Condie, 2005, 2015; Condie et al., 2016) and isotopic (Campbell and Griffiths, 1992, 1993) signatures of modern plumes. Specifically, the enriched recycled components (EM1, EM2, HIMU) that characterize modern OIB-plume magmas seem to be rare or absent from Archaean magmatic suites (Fig. 4, Condie, 2005, 2015; Condie et al., 2016); although Fig. 4B shows that a few Archaean basalts scatter towards a DEP component, thought to be derived from a depleted mantle facies of some type

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(Kerr et al., 1995; Chauvel and Hemond, 2000; Breddam, 2002; Fitton et al., 2003; Stracke et al., 2003; Castillo, 2015, 2016; Buchs et al., 2016). The recycled enriched components (EM1, EM2, HIMU) are widely believed to sample deeply subducted (or delaminated) oceanic and continental lithosphere that accumulated at the core-mantle boundary (Hofmann and White, 1982; Zindler and Hart, 1986; Chauvel et al., 1992; White, 1995; Hofmann, 2003; Li et al., 2014; Castillo, 2015; Garapic et al., 2015; White, 2015). The large low shear-wave velocity provinces (LLSVPs) identified beneath Africa and the Pacific (Dziewonski et al., 2010) are long-lived structures that are thought to focus the ascent of mantle plumes (Burke, 2011; Glisovic and Forte, 2014). In this interpretation, modern plumes are thermal and compositional instabilities that tap slab graveyards at the core-mantle boundary (Sleep, 1988; Montelli et al., 2006), with plume ascent being focused by the edges of LLSVPs. In contrast to Phanerozoic OIB-plume magmas, most 'non-arc' Archaean tholeiites are strongly clustered around PM (primitive mantle) compositions in the Zr/Nb vs. Nb/Th diagram (Fig. 4A), with only rare examples plotting in the characteristic OIB fields that represent subduction-recycled enriched components. As existing concepts and data imply that modern plumes contain components that were introduced to the deep mantle by subduction zones, then the lack of these specific geochemical signals in common Archaean tholeiites would seem to imply a different source, not yet affected by subduction.

Another difference is the apparent absence from the Archaean record (Blichert-Toft et al., 1996; Goodwin, 1996; Thurston, 2015) of silica-undersaturated sodic OIB magmas like alkali basalts, basan-ites, and nephelinites (e.g. S0rensen, 1974; Fitton and Upton, 1987; Hoernle and Schmincke, 1993; Sisson et al., 2009). Archaean 'alkaline' magmas include: rare carbonatites (Bedard and Chown, 1992; Cavell et al., 1992; Villeneuve and Relf, 1998), the uncommon ferropicrites which may or may not be 'alkaline' (Kitayama and Francis, 2014), and low-volume post-tectonic potassic sanuki-toids, syenites and shoshonites (Shirey and Hanson, 1984; Lafleche et al., 1991; Dostal and Mueller, 1992; Blichert-Toft et al., 1996; Bedard, 1996; Stevenson et al., 1999; Wyman et al., 2006, 2015; Laurent et al., 2011,2014; DeJoux et al., 2014). It has been suggested (N.T. Arndt, personal communication) that the absence of alkaline magmas (and the recycled enriched components) from Archaean magmas reflects ubiquitously higher mantle temperatures, which would preclude formation and survival of low-degree melts. However, as Archaean mantle temperatures were heterogeneous (±400 °C: Herzberg et al., 2010; Condie et al., 2016), then large scale overturn-driven upwellings would likely affect cooler mantle domains as well as hotter ones, and the colder domains should melt less than the hotter ones, potentially allowing formation, eruption, and preservation of low-degree alkaline magmas. In addition, the degree of decompression melting should decrease radially away from upwelling axes (O'Hara, 1995), such that cooler domains that upwell more slowly on the periphery of the main ascending mantle plume would melt less than along the axis and could generate low degree melts. Furthermore, carbonatites are considered to be low-degree mantle melts, and are found in Archaean terrains (Bedard and Chown, 1992; Cavell et al., 1992; Villeneuve and Relf, 1998).

11. Impact of mantle overturns on older continents

The Archaean history of individual continents can be rationalized in the context of proximity to OUZOs or to complementary downwelling zones. I suggest that pre-existing continents located directly above an OUZO would have been almost completely reworked (Fig. 3B), and that evidence of this is preserved in the Superior craton, where ca. 60% of the older crust was reworked in

the Neoarchaean (Bedard et al., 2003; Boily et al., 2009; Maurice et al., 2009; Bedard and Harris, 2014; Milidragovic and Francis, 2014). Bedard and Harris (2014) argued that the mantle outflow from a 2.8—2.7 Ga OUZO caused the southern Superior to break up into ribbon continents separated by nascent oceanic tracts like the Abitibi greenstone belt (Fig. 3B), whereas remelts of older crust kept pace with crustal extension in the North-East Superior (Bedard, 2003; cf.; Mole et al., 2012). Enderbites (pyroxene tonalites) constitute about half of the Neoarchaean pulse in the NorthEast Superior (Percival et al., 1992; Percival and Mortensen, 2002; Bedard, 2003; Milidragovic and Francis, 2014). Geochemically, these igneous enderbites are almost indistinguishable from typical hornblende-biotite TTGs (Figs. 2D and 8), and have Nd-isotopic signatures and zircon cores inherited from older felsic protoliths (Boily et al., 2009), leading Bedard and Harris (2014) to suggest that they represent bulk remelts of older TTG. Other constituents of this North-East Superior Neoarchaean pulse range from isotopically juvenile tonalites and trondhjemites formed by anatexis of basalts, to granodiorites and granites with strong inherited crustal signatures and abundant inherited zircon cores derived by remelting older tonalites and granodiorites (Boily et al., 2009; Maurice et al., 2009). Geothermometry indicates that the North-East Superior enderbites were emplaced at ca. 1050 °C (Bedard, 2003), which is close to or above the liquidus temperature of tonalite (Schmidt and Thompson, 1996), and which is probably too high for self anatexis of lower crust by radioactive heating (Bodorkos and Sandiford, 2006). A voluminous mafic/ultramafic magmatic underplate is needed to provide enough heat to rework 60% of the continental crust on such a scale (cf. Milidragovic and Francis, 2014). Given the association with juvenile basalts and komatiites and absence of compressional deformation in the North-East Superior in this time interval, a mantle overturn scenario seems the best explanation for the longevity of the Neoarchaean magmatic pulse, and the large volume and lateral extent of reworked felsic magmas found there.

It was inferred that parts of the Superior craton's pre-Neoarchaean SCLM would have been destroyed by mantle up-welling, exposing fusible older felsic crust to mantle heat (Bedard and Harris, 2014; Fig. 3B). Extensive recycling of SCLM during overturn events seems at odds with the preservation of ancient SCLM keels beneath most cratons. The major peak in SCLM Os model ages starts to grow at ca. 3.1 Ga (Pearson et al., 2007; Griffin et al., 2014), but there are older, smaller peaks at 3.4,3.8 and 4.3 Ga. The SCLM age pattern is similar to the pattern seen in the crustal zircon age data, with a very large Neoarchaean peak and many smaller older ones, a pattern that many believe is a preservation signal, not a growth signal (Armstrong, 1991; Belousova et al., 2010; Condie and Aster, 2010; Hawkesworth et al., 2010). Given the clear evidence for extensive reworking of older crustal rocks to form Neoarchaean felsic magmas (e.g. Bedard, 2006; Foley, 2008; Boily et al., 2009; Maurice et al., 2009; Belousova et al., 2010; Huang et al., 2013; Satkoski et al., 2013), one could ask why SCLM roots could not also have been repeatedly reworked (e.g. Sleep, 1994; Jurine et al., 2005; Sobolev et al., 2011; Bedard and Harris, 2014; Smit et al., 2014a,b; Yang and Leng, 2014)? This implies that many Archaean SCLM segments may have a polycyclic history, much like the crust. It is possible that older crustal domains in the Superior craton survived the Neoarchaean resurfacing event specifically because they were underlain by protective rumps of unresorbed older SCLM.

Older continents that were not astride an OUZO would still be affected but would not be reworked to the same extent and in the same way. Perhaps off-OUZO continents would receive a few basaltic/komatiitic pulses, but not enough to trigger complete and rapid crustal convective overturn? On the other hand, it seems plausible to suggest that the lateral flow of the upper mantle away

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Figure 8. N-MORB normalized spidergrams comparing North-East Superior TTGs and enderbites from the Douglas Harbour domain, NE Superior (Bedard et al., 2003). Individual lines show tonalites and trondhjemites and are colour-coded from most primitive (coldest colours) to most silicic (hottest colours). The two shaded fields show representative sets of mela-enderbites (= tonalites) and leuco-enderbites (= trondhjemites) from the same area. Normalization factors from Sun and McDonough (1989).

from the OUZO would press against the SCLM keels of distal, preexisting continents, causing them to drift and accrete oceanic lithosphere to their leading edge (Bedard et al., 2013; Figs. 1C and 3B). OUZO-triggered drift events would be episodic, short lived (ca. 50—100 Myr?), and may not have had the same drift vector from one overturn to the next. If there are OUZOs, there must also be overturn downwelling zones. However, these downwelling zones are probably not long-lived enough to allow aggregation of all existing cratons to form a supercontinent, unless the continental fragments were already clustered together in its vicinity.

12. Conclusions

Episodic mantle overturns may be a consequence of an early Earth stagnant-lid crust system. Geological data suggest that the Hadean to Archaean mantle may have overturned many times, that overturn phases were protracted (ca. 100 Myr?), and that they alternated with stagnant-lid intervals (ca. 300—500 Myr?). I propose that most Archaean oceanic crust formed in bursts over large parts of Earth during major mantle overturn/resurfacing events characterized by anomalously high mantle melt fluxes and temperatures. The stagnant-lid basaltic crust generated during an overturn event, and the subjacent upper mantle reservoir from

which it was extracted, would form complementary, ephemeral, enriched/depleted isotopic reservoirs. Melt fluxes would decrease markedly in the intervals between mantle overturns, and newly-formed oceanic crust would lapse into a quiescent, stagnant-lid phase that resembles end-member heat-pipe models.

The thermal gradient in the mantle beneath stagnant-lid crust would induce unsteady, shallow convection cells that would only cool the upper mantle, possibly resulting in phases of layered mantle convection. This would have perturbed Earth's heat generation/loss balance because heat generated through radioactive decay and the large heat flux out of the core would remain trapped in the lower mantle, eventually triggering the next overturn (Fig. 3). Periodic overturns would act to rehomogenenize the mantle and retard its radiogenic isotopic evolution (Fig. 7), bringing into question many existing crustal growth scenarios. Only when true subduction started at the end of the Archaean would mantle differentiation towards the depleted MORB mantle (DMM) have begun in earnest, driven principally by the sequestration of subducted slabs at the core-mantle boundary.

As upwelling zones related to Archaean mantle overturns were larger and longer-lived than post-Archaean mantle plumes, and appear to lack enriched plume components (EM1, EM2, HiMu) introduced into the deep mantle by subduction, they may be

J.H. Bédard / Geoscience Frontiers xxx (2017)

distinct from modern plumes and are called overturn upwelling zones (OUZOs) to facilitate discussion. It is inferred that continents began to form above long-lived OUZOs, as these are sites where large volumes of hot magma would have ascended over a protracted interval. This provides the necessary context for the extensive and protracted reworking of the lower crust required to generate craton-scale TTG-suites. Delamination of restites from TTG-forming events would refertilize the subjacent mantle, triggering further mantle melting events, and eventually creating an evolved crust and ultra-depleted SCLM pair (Bedard, 2006).

Pre-existing continents located above OUZOs would be strongly reworked and Bedard and Harris (2014) argued that a mantle up/ outflow from the 2.8—2.7 Ga mantle overturn extensively reworked and partly disaggregated the Superior craton. Older continents that were not astride an OUZO would not be reworked to the same extent, but lateral mantle flow away from an OUZO would press against the SCLM keels of distal continents, causing them to drift and grow by leading-edge terrane accretion of Abitibi-like crust (Bedard et al., 2013). Such a continental-drift accretion model can account for the systematic polarity and short collision interval of Archaean terrane docking events in the Southern Superior accre-tionary orogeny, and elsewhere. As a drifting continent advances, the cold upper parts of foreland 'terranes' would deform and thicken as duplexes that could emerge and be eroded, generating syn-orogenic metasedimentary belts. Imbrication/subcretion of metabasaltic packages would introduce large volumes of wet basalt to depths where they could yield anatectic syn-kinematic TTGs. Promontory-reentrant geometries would lead to complex structural patterns and can account for the location and timing of some lode gold camps (Fig. 5). The downthrust angle of colder oceanic segments embedded in such an accretionary system may be steepened by eclogitization (Fig. 1C), creating space for mantle to well up and interact with fluids and melts emanating from the dipping/dripping basaltic terrane; generating subordinate suites of syncollisional magmas with arc-like source-metasomatic geochemical signatures. Buoyant Archaean stagnant-lid oceanic crust would resist the advance of the drifting continents, decreasing convergence rates and precluding formation of blueschists.

Plate tectonics today is constituted of two systems: a bottom-up continental drift system and a top-down subduction-driven system. The lower plate is the active agent in modern convergent margins characterized by active subduction, as negatively buoyant oceanic lithosphere sinks and rolls back under its own weight. Modern active subduction is a plate-driving force because slab-pull forces can be transmitted through the thick, stiff, sub-oceanic lithospheric mantle (SOLM). In the Archaean stagnant-lid context of unsteady induced convection cells, a given crustal segment would never drift away from a focussed, linear thermal anomaly (the spreading ridge), never form a significant SOLM thickness, and so would remain unsubductable. In consequence, typical Archaean oceanic lithosphere would have been extremely weak and probably could not have stabilized a subduction/seafloor-spreading couple. The distinctive Archaean rock associations and structures can best be explained if continental drift started early in Earth's history (ca. 3.9 Ga), long before active subduction began (ca. 2.5 Ga). It is argued that most Archaean convergent margins were not characterized by active subduction but were places where stagnant-lid oceanic lithosphere was passively imbricated and subcreted by drifting continents (Bedard et al., 2013). As Earth cooled and the background oceanic lithosphere became colder and stiffer, there would be an increasing probability that older crustal segments in these mafic accretionary orogens would founder in a more organized way when subjected to compression, a pattern that would gradually evolve into modern-style active subduction around 2.5 Ga.

Acknowledgments

Clark Friend and Nick Arndt provided formal reviews that helped me improve the manuscript. I thank Lyal Harris, Jean-François Moyen, Kent Condie, Phil Thurston, Bill Davis and Claude Herzberg for comments on earlier versions. Any errors that remain are wholly mine, however. I also benefited from discussions with Craig O'Neill, Ellis Hoffmann, Hugh Smithies, Martin VanKranendonk, Shoufa Lin, and all who have critiqued my ideas in the past. François Leclerc generously provided the updated Chibougamau database. I especially thank my Division Head Andrée Bolduc for her support of this non-Program research. This is ESS/NRCan/GSC contribution # 20160233.

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