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1 The Institute for Geoscience Research, Department of Applied Geology, Curtin University of Technology, GPO Box U 1987, Perth, WA 6845, Australia
2 Department of Geology and Geophysics, Yale University, New Haven, CT 06520-8109, USA
* Corresponding author (e-mail: S.Reddy{at}curtin.edu.au)
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Abstract |
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 Top Abstract Evolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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The connections illustrated above are based on concepts or a few well-constrained examples. The actual record of supercontinents on Earth is not yet well enough known to verify or modify the models in deep time. Prior to Pangaea (approximately 0.25–0.15 Ga) and Gondwana-Land (0.52–0.18 Ga), the possible configurations and even existence of Neoproterozoic Rodinia (c. 1.0–0.8 Ga) are intensely debated (Meert & Torsvik 2003; Li et al. 2008; Evans 2009). Prior to Rodinia, an earlier supercontinental assemblage at 1.9–1.8 Ga has been suggested. Although only preliminary and palaeomagnetically untested models of this supercontinent have been published (Rogers & Santosh 2002; Zhao et al. 2002, 2004), its assembly appears to have followed tectonic processes that are remarkably similar to those of the present day (Hoffman 1988, 1989). This supercontinent is referred to by various names (e.g. Columbia, Nuna, Capricornia) but here, and below, we refer to it as Nuna (Hoffman 1997).
It is unclear whether Nuna's predecessor was a large supercontinent, or whether it was one of several large, but distinct coeval landmasses (Aspler & Chiarenzelli 1998; Bleeker 2003). Nonetheless, numerous large igneous provinces, with ages between 2.45 and 2.2 Ga, perforate the world's 35 or so Archaean cratons and could represent an episode of globally widespread continental rifting at that time (Heaman 1997; Buchan et al. 1998; Ernst & Buchan 2001).
Given that the Palaeoproterozoic era is defined chronometrically at 2.5–1.6 Ga (Plumb 1991), it thus encompasses one or more episodes, perhaps cycles, of global tectonics. As enumerated by the following examples, these tectonic events coincide with fundamental changes to the Earth as an integrated system of core, mantle, lithosphere, hydrosphere, atmosphere, and biosphere. Understanding these changes requires the integration of seemingly disparate geoscience disciplines. One pioneering review of this sort (Nance et al. 1986) has been followed by an incredible wealth of precise geochronological data constraining ages within the Palaeoproterozoic geological record. In this chapter we have compiled the currently available data from core to atmosphere, from late Archaean to late Mesoproterozoic time (3.0–1.2 Ga), to provide an overview of Earth-system evolution and an up-to-date temporal and spatial framework for hypotheses concerning the global transition through Earth's ‘middle ages’.
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Evolution of the core |
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TopAbstractEvolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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Recent estimates of the age of inner core nucleation, on a theoretical basis constrained by the geochemical data, are typically about 1 Ga (Labrosse et al. 2001; Nimmo et al. 2004; Butler et al. 2005; Gubbins et al. 2008). Much discussion on this topic is confounded with discussions of the intensity of Earth's ancient geomagnetic field. This is because two of the primary energy drivers of the geodynamo are thought to be thermal and compositional convection in the outer core due to inner core crystallization (Stevenson et al. 1983). Early attempts to determine the palaeointensity of Earth's magnetic field, which is among the most laborious and controversial measurements in geophysics, suggested an abrupt increase in moment near the Archaean–Proterozoic boundary (Hale 1987). Subsequent refinements to techniques in palaeointensity (e.g. the single-crystal technique applied by Smirnov et al. 2003) have generated mixed results, but several of the measurements indicate a strong field in the earliest Proterozoic (Fig. 1a). More traditional palaeointensity techniques complete the later Proterozoic time interval with results of generally low palaeointensity (Macouin et al. 2003). There is currently no systematic test among the various palaeointensity techniques, so absolute Precambrian palaeointensity values remain ambiguous. In addition, the large apparent increase in palaeointensity at c. 1.2 Ga (Fig. 1a) is underpinned by sparse data that lack some standard reliability checks (Macouin et al. 2004). However, we tentatively explore the possibility that all the available records provide reliable estimates of ancient geomagnetic field strength.
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Mantle evolution |
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TopAbstractEvolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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Numerical modelling is one approach that has received considerable attention in attempting to constrain Earth cooling models. Such modelling commonly utilizes the simple relationship between radiogenic heat production, secular cooling and surface heat flux. Critical to these models is the way in which surface heat flux is calculated and how this heat flow is scaled to mantle convection over time. ‘Conventional’ models commonly calculate surface heat flux using scaling laws that assume a strong temperature dependency on viscosity such that hotter mantle convects more vigorously, thereby increasing surface heat flux. As clearly enunciated by Korenaga (2006), applying conventional heat flux scaling from current conditions back through geological time predict unrealistically hot mantle temperatures before 1 Ga and lead to the so called ‘thermal catastrophe’ (Davies 1980). Several numerical models have therefore addressed different ways of modifying the scaling laws to avoid these unrealistic mantle temperatures. One way of doing this is by assuming a much higher convective Urey ratio (i.e. the measure of internal heat production to mantle heat flux) in the geological past (e.g. Schubert et al. 1980; Geoffrey 1993) than modern day estimates (Korenaga 2008c). Although this assumption overcomes problems of the thermal catastrophe in the Mesoproterozoic, the solution leads to high Archaean mantle temperatures that appear inconsistent with empirical petrological data for mantle temperatures (Abbott et al. 1994; Grove & Parman 2004; Komiya 2004; Berry et al. 2008) (Fig. 1b).
An alternative way of alleviating the problem of a Mesoproterozoic ‘thermal catastrophe’ is to assume layered mantle convection rather than whole-mantle convection (e.g. Richter 1985). A layered mantle has been inferred as a means of maintaining distinct geochemical reservoirs, particularly noble gas compositions measured between mid-ocean ridge basalts and plume-related ocean island basalts (e.g. Allegre et al. 1983; O'Nions & Oxburgh 1983) and large-ion lithophile elements budgets of the crust and mantle (e.g. Jacobsen & Wasserburg 1979; O'Nions et al. 1979). The nature of this layering differs between different models (c.f. Allegre et al. 1983; Kellogg et al. 1999; Gonnermann et al. 2002). However, critical to this argument is that the modelling of layered versus whole-mantle convection and its affect on present day topography supports the latter (Davies 1988), and seismic tomographic data provide evidence of subducting slabs that penetrate the lower mantle (e.g. van der Hilst et al. 1997). These observations are difficult to reconcile with the classic model of layered convection (see van Keken et al. 2002 for review).
With increasing geochemical data from a range of different sources, geochemical constraints on mantle evolution are becoming more refined, and the two-layer mantle reservoir models have necessarily become more complex (see reviews of Graham 2002; Porcelli & Ballentine 2002; Hilton & Porcelli 2003; Hofmann 2003; Harrison & Ballentine 2005). These data have led to the formulation of models in which some of the chemical heterogeneity is stored in the core or deep mantle (Porcelli & Elliott 2008), or is associated with a unmixed lower mantle ‘magma ocean’ (Labrosse et al. 2007), or is associated with lateral compositional variations (Trampert & van der Hilst 2005), or is explained by filtering of incompatible trace elements associated with water release and melting associated with magnesium silicate phase changes in the mantle transition zone (Bercovici & Karato 2003). However, a recent development in layered mantle geodynamics and the understanding of core–mantle interaction is the significant discovery of the post-perovskite phase transition (Murakami et al. 2004) and the realization that such a transition seems to account for many of the characteristics of the seismically recognized D'' of the lower mantle (see Hirose & Lay 2008 for a recent summary). The large positive Clapeyron slope of the perovskite–post-perovskite phase transition means that cold subducting slabs that reach the core–mantle boundary are likely to lead to enhanced post-perovskite formation and form lateral topography on the D'' layer (Hernlund et al. 2005). With the addition of slab material to the D'' layer, and its potential for enhanced melting (Hirose et al. 1999), it is likely that there will be significant geochemical implications for lower mantle enrichment and melting (see for example Kellogg et al. 1999; Labrosse et al. 2007), in addition to the potential for mantle plume generation (Hirose & Lay 2008), both of which are yet to be resolved. The post-perovskite stability field intersects only the lowermost depths of the present mantle geotherm (Hernlund et al. 2005) and the secular cooling rate of the lowermost mantle is not well enough known to estimate the age at which this post-perovskite phase first appeared in Earth history. However, this is likely to be of importance in the secular evolution of the deep Earth.
To summarize, current geodynamical and geochemical models for the modern mantle are therefore necessarily complex and it seems that the answer may reside in a mixture of whole-mantle and episodic layered convection that is closely linked to large-scale plate tectonic processes and affects chemical heterogeneities preserved at a range of scales (Tackley 2008).
Despite the advances in understanding the geodynamics of the present-day mantle, extrapolation to the Palaeoproterozoic mantle remains elusive. In the last few years, developments in the modelling of secular cooling of the mantle have utilized surface heat flux scaled to a ‘plate tectonic’ model involving mantle melting at mid-ocean ridges (Korenaga 2006) or intermittent plate tectonic models in which subduction flux, indicated by geochemical proxies (e.g. the Nb/Th ratios of Collerson & Kamber 1999 Fig. 1e), varies over time (Silver & Behn 2008a). These models overcome the problems of Mesoproterozoic thermal runaway, calculated back from present day conditions via conventional scaling laws, in the first case by taking account of depth-dependent mantle viscosity variations as a function of melting (Hirth & Kohlstedt 1996) and in the second by reducing the amount of heat loss at times of plate tectonic quiescence (Silver & Behn 2008a). Importantly, such models predict significantly reduced plate velocities during the Neoarchaean and Palaeoproterozoic eras than conventional models do (Fig. 1c) because of either the increased difficulty of subducting thicker dehydrated lithosphere (Korenaga 2008b) or the episodic nature of the supercontinent cycle (Silver & Behn 2008a). However, they remain controversial (Korenaga 2008a; Silver & Behn 2008b), particularly in light of recent evidence for a temperature control on lower crustal thermal conductivity (Whittington et al. 2009) and empirical estimates of Archaean mantle temperatures (Berry et al. 2008) that are higher than the model predictions (Fig. 1b). Even so, the development of models that are intimately linked to plate tectonic processes and that involve the formation and subduction of strong, plate-like lithosphere that controls mantle convection (e.g. Tackley 2000; Bercovici 2003), provide predictions from the core to the atmosphere that can be empirically constrained by the extant geological record, for example secular variations in crust and mantle geochemistry (Fig. 1d), the amalgamation and dispersal of supercontinents (Fig. 1i) and chemistry of the oceans and atmosphere (Fig. 1k, q).
Another feature of mantle evolution that has received considerable scientific attention is the formation and secular development of mantle plumes (e.g. Ernst & Buchan 2003 and references therein). One of the expressions of the impingement of mantle plumes on Earth's lithosphere is the development of large igneous provinces (LIPS) (Ernst & Buchan 2001, 2003) manifest as continental or oceanic flood basalts and oceanic plateaus. The geological record of LIPS (summarized by Ernst & Buchan 2001; Abbott & Isley 2002) recognizes a general decrease of size in younger LIP events, and an episodic distribution in time (Fig. 1d). Time-series analysis of LIPS recognizes a c. 330 Ma cycle in the period from 3.0–1.0 Ga upon which weaker, shorter duration cycles are superimposed (Prokoph et al. 2004). As pointed out by Prokoph et al. (2004), no simple correlation exists between the identified LIP cycles and possible forcing functions, though current global initiatives to date mafic volcanism precisely (e.g. Bleeker & Ernst 2006) should improve the likelihood of establishing such correlations. Currently, the correlations reported in Figure 1 show that LIP activity at c. 2.7–2.8 Ga coincides with global banded iron formation (BIF) and orogenic gold formation, while a second event at 2.45 Ga corresponds with a the major formation of Superior-type BIF. In addition, the major 2.7 Ga peak in LIP activity lies temporally close to the peak in juvenile crust formation (Fig. 1g), a correlation that has led to the suggestion that these features may be geodynamically related (Condie 2004) and which may explain observed gold enrichment (Brimhall 1987; Pirajno 2004).
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Crustal growth and emergence |
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TopAbstractEvolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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One widely used approach has been the analysis of fine-grained sedimentary rocks, thought to provide a statistical representation of the upper crust, to establish the increase in volume of the continental crust over time (e.g. Allegre & Rousseau 1984; Taylor & McLennan 1985, 1995). Other approaches include the direct analysis of zircon age populations from juvenile continental crust (Condie 1998), or indirect analysis by assessing the chemical evolution of the depleted upper mantle from which continental crust is believed to be ultimately sourced (e.g. Bennett 2003 and references therein) (Fig. 1f). More recent developments include the integration of in situ stable and radiogenic isotope analysis (Valley et al. 2005; Hawkesworth & Kemp 2006b).
Despite the difficulties in constraining the composition of continental crust (Rudnick et al. 2003) and the likely requirement of two-stage differentiation (e.g. Arndt & Goldstein 1989; Rudnick 1995; Kemp et al. 2007), several fundamentally different growth models dominate the literature (for recent summary see Rino et al. 2004): rapid differentiation of the crust early in Earth's history and subsequent recycling such that there has been little subsequent increase in volume over time (e.g. Armstrong 1981, 1991); growth mainly in the Proterozoic (Hurley & Rand 1969; Veizer & Jansen 1979); high growth rates in the Archaean followed by slower growth in the Proterozoic (Dewey & Windley 1981; Reymer & Schubert 1984; Taylor & McLennan 1995) and growth accommodated by discrete episodes of juvenile crust formation (McCulloch & Bennett 1994; Condie 1998, 2000).
The first of these models (Armstrong 1981, 1991) seems unlikely in the light of relatively constant 
18O from Archaean rocks that show little evidence of extensive crustal recycling prior to 2.5 Ga (Valley et al. 2005). These data also point towards significant crustal growth in the Archaean and so are also inconsistent with the models of dominantly Proterozoic growth (Hurley & Rand 1969; Veizer & Jansen 1979). In the case of progressive or episodic crustal growth, mantle depletion events should also mimic continental growth. Despite possible complexities associated with crustal recycling and questions regarding the nature of mantle convection, Re–Os data from peridotites and platinum group alloys indicate mantle depletion events that cluster at 1.2, 1.9 and 2.7 Ga (Pearson et al. 2007) (Fig. 1e). A temporally similar peak at 2.7–2.5 Ga is recorded in Nb/Th data (Collerson & Kamber 1999) following polynomial fitting of the data (Silver & Behn 2008a) and in the 4He/3He ocean island basalt data inferred at 2.7 Ga and 1.9 Ga (Parman 2007) (Fig. 1e). Although the age constraints on the mantle depletion events recorded by the He data are poor, the pattern of Os and Nb/Th data are similar to that documented by the temporal distribution of juvenile continental crust (Fig. 1g) and considered to reflect the formation of supercontinents (Fig. 1i) (Condie 1998; Campbell & Allen 2008), possibly linked to large-scale mantle overturn events (Condie 2000; Rino et al. 2004), and the cessation of, or decrease in, subduction flux (Silver & Behn 2008a). Recent models integrating chemical differentiation with mantle convection also predict the episodicity of juvenile crust formation (Walzer & Hendel 2008), as do large-scale mantle overturn events that are thought to take place on a timescale of several hundred million years (Davies 1995).
Despite the above correlations, the pattern of juvenile crust ages has been argued to be a consequence of the preservation potential within the supercontinent cycle, in particular the ability to preserve material inboard of arcs (Hawkesworth et al. 2009). Although this seems a reasonable interpretation based on the crustal record, the temporal link between mantle depletion events (Collerson & Kamber 1999; Parman 2007; Pearson et al. 2007) and juvenile crust (Condie 1998; Campbell & Allen 2008) is less easy to explain by preservational biases and remains a compelling observation for linking these processes. Our current preference is therefore for continental growth models that involve continental crust formation in the Archaean with subsequent reworking, recycling and the addition of juvenile material via episodic processes through the Proterozoic.
The growth of continental crust, its evolving volume and its thickness are intimately related to the evolution of the mantle (e.g. Hynes 2001). These characteristics also play a critical role in continental freeboard, the mean elevation of continental crust above sea level, and the emergence of the continents (e.g. Wise 1974; Eriksson et al. 2006). The emergence of the continental crust above sea level in turn influences the nature of sedimentation (Eriksson et al. 2005b) and enables weathering that affects ocean and atmospheric compositions. Several lines of evidence point to continental emergence around the Archaean–Proterozoic boundary, primarily in the form of distinctive geochemical trends that require low-temperature alteration and crustal recycling. One such line of evidence comes from a recent compilation of 18O data from the zircons of juvenile rocks, which show a clear trend of relatively constant Archaean values with increasing values at c. 2.5 Ga (Valley et al. 2005) (Fig. 1f). This trend mimics those recorded in Hf and Nd isotopic data from juvenile granitic rocks (Fig. 1f), which is thought to represent depletion of the mantle associated with extraction of the continental crust (Bennett 2003). The pattern of oxygen isotopic variation in zircon is explained by complex contributions of various processes (Valley et al. 2005). However, a requirement is for a component of low-temperature fractionation commonly interpreted to be associated with continental weathering, the recycling of supracrustal rocks and subsequent melting.
The most recent data to place temporal constraints on continental emergence comes from an analysis of submarine versus subaerial LIP (Kump & Barley 2007) (Fig 1f), which shows an abrupt increase in the secular variation of subaerial LIPs at c. 2.5 Ga. This is thought to have had significant global repercussions with respect to the increase in atmospheric oxygen levels (Kump & Barley 2007) and is discussed further below. Recent modelling of continental emergence that links continental freeboard with different models for cooling of the Earth indicates that the emergent continental crust was only 2–3% of the Earth's surface area during the Archaean, a stark contrast to present day values of c. 27% (Flament et al. 2008).
Sedimentary rocks document the physical, chemical, and biological interface between the Earth's crust and its changing Precambrian surface environments (see Eriksson et al. 2005b for an overview). Palaeoproterozoic sedimentary basins share many of the same features as their modern counterparts, whether in rift settings (Sengor & Natal'in 2001), passive margins (Bradley 2008), strike-slip basins (e.g. Ritts & Grotzinger 1994), or foreland basins (e.g. Grotzinger et al. 1988). Ironically, Precambrian sedimentary structures can be much easier to decipher than their Phanerozoic counterparts due to the lack of bioturbation in pre-metazoan depositional environments. However, there are some notable differences between Precambrian and modern clastic sedimentation systems. Most notably, sandstones of pre-Devonian river systems are commonly characterized by sheet-braided geometry with greater channel widths ascribed to the lack of vegetative slope stabilization (e.g. Long 2006). The deposits of specific palaeoenvironments such as aeolianites have a temporal distribution modulated by long-term preservational potential and possible relationships to phases of supercontinental cyclicity (Eriksson & Simpson 1998), as is also the case for glaciogenic deposits, which are summarized below. Also discussed below are factors determining the temporal variation in redox-sensitive mineral clasts such as detrital pyrite and uraninite (Fig. 1p), and the abundances and chemical compositions of banded iron-formations (Fig. 1k), evaporites (Fig. 1n), and carbonate chemistry and structure (Grotzinger 1989; Grotzinger & James 2000).
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Supercontinents |
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TopAbstractEvolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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There are additional names for putative Palaeoproterozoic continental assemblages. Capricornia (Krapez 1999) refers to a model for the early amalgamation of Australia and hypothesized adjacent cratons Laurentia, India, Antarctica, and Kalahari. Arctica is defined as the postulated assemblage of Siberia with northern Laurentia, and Atlantica comprises the proposed long-lived amalgamation of West (and parts of northern) Africa, Amazon, São Francisco–Congo, and Rio de la Plata (Rogers 1996). Of these, a Siberia–northern Laurentia connection, if not directly adjacent then slightly separated, is allowed by numerous independent palaeomagnetic comparisons from 1.5 to 1.0 Ga (Pisarevsky & Natapov 2003; Pisarevsky et al. 2008; Wingate et al. 2009), and lack of tectonic activity after 1860 Ma in southern Siberia would suggest that this position had been established by that time (Poller et al. 2005; Pisarevsky et al. 2008). A long-lived Atlantica continent is difficult to test palaeomagnetically, due to a dearth of reliable results from its constituent cratons (Meert 2002; Pesonen et al. 2003). Minor transcurrent motions between West Africa and Amazon, sometime between original craton assembly at 2.1 Ga (Ledru et al. 1994) and their Gondwanan amalgamation in the Cambrian (e.g. Trindade et al. 2006), are proposed based on limited palaeomagnetic data from those two blocks (Onstott & Hargraves 1981; Onstott et al. 1984; Nomade et al. 2003). Such minor amounts of relative displacement would preserve the proposed tectonic correlations among the Atlantica cratons, but further palaeomagnetic testing is needed.
Nuna, commonly under the guise of ‘Columbia’, is commonly reconstructed with many cratonic juxtapositions taken from inferred models of Rodinia (Rogers & Santosh 2002; Zhao et al. 2002, 2004; Hou et al. 2008a, b). These models lack robust palaeomagnetic constraints, and the reconstructions are commonly distorted by unscaled cut-outs from a Mercator projection of Pangaea. When palaeomagnetic data are incorporated into geometrically accurate reconstructions, the sparsity of reliable results has led authors to conflicting conclusions of either non-existence of a 1.8 Ga supercontinent altogether (Meert 2002), or one that accommodated substantial internal shears (Laurentia–Baltica motions illustrated in Pesonen et al. 2003), or the consideration of at most a few fragments with applicable data from discrete time intervals (Li 2000; Salminen & Pesonen 2007; Bispo-Santos et al. 2008).
The most robust long-lived juxtaposition of Palaeo–Mesoproterozoic cratons is that of Laurentia and Baltica throughout the interval 1.8–1.1 Ga. First proposed on geological grounds and named ‘NENA’ (northern Europe–North America, Gower et al. 1990), this juxtaposition finds palaeomagnetic support from numerous results throughout that interval, defining a common apparent polar wander path when poles are rotated according to the reconstruction (Evans & Pisarevsky 2008; Salminen et al.). The NENA reconstruction is distinct in detail from the commonly depicted juxtaposition of Hoffman (1988) that has been reproduced in Zhao et al. (2002, 2004), which is not supported palaeomagnetically. NENA is a specific reconstruction between two cratons and is distinct from Nuna, the supercontinent proposed to have assembled at 1.9–1.8 Ga without particular palaeogeographic specifications. By coincidence, the two names are similar, and NENA appears to be a robustly constrained component of Nuna. Payne et al. extend the Palaeo–Mesoproterozoic apparent polar wander comparisons to include Siberia, North and West Australia, and the Mawson Continent (Gawler craton with original extensions into Antarctica). The Australian proto-continent is restored in that analysis by the same sense of relative rotations as proposed by Betts & Giles (2006). As more data from other cratons accumulate through the 1.8–1.5 Ga interval, this approach should lead to a successful first-order solution of Nuna's palaeogeography.
Kenorland is the name given to the palaeogeographically unspecified supercontinent that might have formed in the Neoarchaean era (Williams et al. 1991). Its etymology derives from the Kenoran orogeny (Stockwell 1982) that represents cratonization of the Superior craton at about 2.72–2.68 Ga (Card & Poulsen 1998; Percival et al. 2006) and of that age or younger in the Neoarchaean on other cratons (Bleeker 2003). Breakup of Kenorland would be represented by numerous large igneous provinces starting with a global pulse at 2.45 Ga (Heaman 1997). Subsequently, the meaning of ‘Kenorland’ has varied. Aspler & Chiaranzelli (1998) referred to Kenorland as the specified palaeogeography of ancestral North America, with an interpretation of the Trans-Hudson and related orogens that accommodated at most accordion-like oceanic opening and reclosing, but not extensive reshuffling of cratons. Baltica and Siberia were included in unspecified palaeogeographic configurations. A second proposed supercontinent (‘Zimvaalbara’ comprising Zimbabwe, Kaapvaal, Pilbara and perhaps São Francisco and cratons in India) is proposed to have assembled and begun to break up somewhat earlier than Kenorland, at 2.9 and 2.65 Ga, respectively; yet its final fragmentation was conjectured at 2.45–2.1 Ga, simultaneous with that of Kenorland (Aspler & Chiarenzelli 1998).
Bleeker (2003) chose a different nomenclature that referred to ‘Kenorland’ as a possible solution to Neoarchaean–Palaeoproterozoic palaeogeography whereby all (or most) cratons were joined together; he also proposed an alternative solution of palaeogeographically independent ‘supercratons’ that would each include a cluster of presently preserved cratons. Three supercratons were exemplified and distinguished by age of cratonization: Vaalbara (2.9 Ga), Superia (2.7 Ga), and Sclavia (2.6 Ga). Vaalbara has the longest history of recognition and palaeomagnetic testing (Cheney 1996; Wingate 1998; Zegers et al. 1998) with the most recent tests allowing a direct juxtaposition (Strik et al. 2003; de Kock et al. 2009). Superia has become more completely specified by the geometric constraints of precisely dated, intersecting dyke swarms across Superior, Kola–Karelia, Hearne, and Wyoming cratons through the interval 2.5–2.1 Ga (Bleeker & Ernst 2006). Aside from Slave craton, the additional elements of Sclavia remain unknown. Lastly, a completely different view of Kenorland (Barley et al. 2005) entails inclusion of all (or most) of the world's cratons, following assembly as late as 2.45 Ga, and breakup at 2.25–2.1 Ga. This definition would be consistent with a global nadir in isotopic ages during the 2.45–2.25 Ga interval, proposed earlier as a time of supercontinent assembly (Condie 1995), or a more recently and radically, a cessation of plate tectonics altogether (Condie et al. 2009).
Palaeogeographic constraints on these proposed Neoarchaean–Palaeoproterozoic supercontinents and supercratons, other than the aforementioned Vaalbara, are limited by a small number of reliable palaeomagnetic data (Evans & Pisarevsky 2008). The best progress has been made in the Superior craton, where a series of precise U–Pb ages on mafic dyke swarms has been closely integrated with palaeomagnetic studies at the same localities, and with particular attention to baked-contact tests to demonstrate primary ages of magnetic remanence (e.g. Buchan et al. 2007; Halls et al. 2008). If a similar strategy is applied to other cratons then the solution of Kenorland or supercraton configurations will be much closer to realization.
Finally, Piper (2003) incorporated the extant Archaean–Palaeoproterozoic cratons into a long-lived supercontinent (duration 2.9–2.2 Ga) named ‘Protopangaea’. This assemblage mirrors Piper's proposed Meso-Neoproterozoic supercontinent ‘Palaeopangaea’ (Piper 2007). Both are based on broad-brush compilations of the entire database containing a complex mixture of primary and secondary magnetizations rather than on precise palaeomagnetic comparisons of the most reliable data. Meert & Torsvik (2004) point out some of the problems with this approach, and Li et al. (2009) demonstrate specific quantitative refutations of the reconstructions.
Linked to the supercontinent debate is the controversy regarding the timing of initiation of a modern-style of plate tectonics (cf. Stern 2005; Cawood et al. 2006). A range of geological, geodynamic and geochemical constraints (recently summarized by Condie & Kroner 2008; Shirey et al. 2008; van Hunen et al. 2008), suggest the strong likelihood of plate tectonic behaviour in the Palaeoproterozoic and possibly back into the Archaean; an inference also made by Brown (2007) based on the presence of characteristic high pressure–lower temperature and high temperature–lower pressure metamorphic mineral assemblages associated with subduction and arcs respectively (Fig. 1h). Recent numerical models are consistent with this interpretation (e.g. Labrosse & Jaupart 2007). For a recent summary of the different aspects of the ongoing debate the reader is referred to Condie & Pease (2008) and references therein.
In summary, although Rodinia's configuration remains highly uncertain, a working model of its predecessor Nuna is being assembled rapidly with the acquisition of new geochronological and palaeomagnetic data considered in the context of global tectonic constraints. It remains uncertain whether there was a supercontinent at all during the Archaean–Proterozoic transition, and if so, when it existed and how its internal configuration of cratons was arranged. Figure 1i depicts these uncertainties using the template illustration of Bleeker (2003). Creation of a global stratigraphic database (Eglington et al. 2009) not only illustrates the growing wealth of global age constraints on the Palaeoproterozoic rock record, but also shows which stratigraphic units are in greatest need of dating, and which units are most promising for successful preservation of primary palaeomagnetic remanence directions that will allow further progress in supercontinent reconstruction.
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Minerals and mineral deposits |
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TopAbstractEvolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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The evolving ocean and atmosphere |
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TopAbstractEvolution of the coreMantle evolutionCrustal growth and emergenceSupercontinentsMinerals and mineral depositsThe evolving ocean and...The evolving palaeoclimate and...Extra-terrestrial influences:...DiscussionAcknowledgmentsReferences |
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Anomalous variation among the four stable sulphur isotope ratios, deviating from purely mass-dependent effects, are common in Archaean and earliest Proterozoic sedimentary pyrites (Farquhar et al. 2000). The deviations are thought to arise from gas-phase sulphur reactions in the upper atmosphere with concomitant mixing into seawater, with further constraints on sulphur speciation provided by increasingly refined datasets (e.g. Ono et al. 2003, 2009a, b). The termination of this non-mass-dependent fractionation signal is now estimated at c. 2.4–2.32 Ga (Bekker et al. 2004). Atmospheric modelling suggests that it was caused by either the rise of O2 above 10–5 of present atmospheric levels (Pavlov & Kasting 2002) or loss of CH4 below a critical threshold of c. 10 ppmv due to a shrinking ecological role of methanogenic producers (Zahnle et al. 2006). The collapse of methane from formerly high levels in the Archaean atmosphere (reviewed by Kasting 2005) probably plays a large role, not only the oxygenation history, but also the occurrence of Palaeoproterozoic ice ages, which will be discussed below. The combined data point to the 2.4–2.3 Ga interval as host to a ‘Great Oxidation Event’ (GOE: Cloud 1968; Holland 2002) during which Earth's surface environment changed profoundly and irreversibly.
Possible changes in atmospheric oxidation state can be estimated by age variations in the non-mass-dependent signal (Fig. 1o). Many data now exist for the crucial interval of 3.0–2.4 Ga (Ohmoto et al. 2006; Farquhar et al. 2007; Kaufman et al. 2007; Ono et al. 2009a, b) and perhaps the most compelling evidence for a ‘whiff’ of oxygen prior to 2.4 Ga is furnished by Mo and Re concentration peaks in black shales at 2.5 Ga (Anbar et al. 2007; Wille et al. 2007). Much of the inferred global nature of these peaks depends on the correlation of signals between the Pilbara and Kaapvaal cratons, whereas a subsequent palaeomagnetic reconstruction of the blocks places them in direct juxtaposition (de Kock et al. 2009). More records of similar nature are needed from other cratons through this interval, to substantiate the global nature of any hypothesized oxygenation events. Localized, photosynthetically produced ‘oxygen oases’ could attain ppm-level O2 concentrations within plumes dissipating into the methane-rich Neoarchaean troposphere in a timescale of hours to days (Pavlov et al. 2001).
The causes of the GOE remain unclear. Traditionally attributed to development of photosystem II in cyanobacteria (Cloud 1968; Kirschvink & Kopp 2008) or enhanced burial rates of organic carbon produced by that process (Karhu & Holland 1996), other potential long-term sources of oxidizing agents involve dissociation of atmospheric methane coupled to H2 escape (Catling et al. 2001; Catling & Claire 2005), or the changing oxidative state of the upper mantle coupled to hydrothermal alteration of seafloor basalts (Kasting et al. 1993; Kump et al. 2001; Holland 2002). Finally, a qualitative empirical relationship between continental collisions and global oxygen increases over the past three billion years (Campbell & Allen 2008), linked to carbon and sulphur burial through enhanced physical weathering and sediment transport in mountain belts, is rendered particularly speculative for the GOE due to the poorly understood history of Palaeoproterozoic supercontinents or supercratons (see above). Widespread consideration of the non-biological alternatives has been motivated primarily by the apparent antiquity of molecular biomarkers for photosynthesizing organisms as old as 2.7 Ga, about 300–400 million years prior to the GOE (Brocks et al. 1999); however, the reliability of those records is currently debated (Eigenbrode et al. 2008; Rasmussen et al. 2008; Waldbauer et al. 2009).
Returning to the ocean, banded iron formations are a prima facie example of non-uniformitarianism in the Earth's palaeoenvironment. Deposited widely prior to 1.85 Ga (Fig. 1k), they are nearly absent in the subsequent geological record, apart from a few small occurrences or unusual associations with Neoproterozoic low-latitude glacial deposits of putative Snowball Earth events (Kirschvink 1992; Klein & Beukes 1992). The marked concentration of iron formation in the Palaeoproterozoic era has traditionally been construed as an indication of atmospheric oxidation at about that time (e.g. Cloud 1968) and that element is no doubt part of the story. However, recent research on iron formations using geochemical and isotopic tracers is painting a more refined picture of oceanic evolution.
Iron formations are broadly categorized into two classes, one with close association to volcanic successions (Algoma-type), and one without (Superior-type) (Gross 1965; Gross 1983). The Algoma-type iron formations are distributed throughout the Archaean sedimentary record, whereas Superior-type deposits only become significant at 2.6 Ga and later (Huston & Logan 2004) (Fig. 1k). By 2.5 Ga, Superior-type iron formations predominate, to the extent that ‘Siderian’ was informally proposed as the first period of the Palaeoproterozoic era (Plumb 1991). With refined global geochronology, the global peak in Palaeoproterozoic iron formation (e.g. Klein 2005 and references therein) appears to split into two modes, at c. 2.45 Ga and c. 1.9 Ga (Isley & Abbott 1999; Huston & Logan 2004), although some large deposits remain imprecisely dated (e.g. Krivoy Rog, Ukraine; Simandou, West Africa). The temporal distribution of iron formation is closely matched by the occurrence of thick, massive seafloor calcium carbonate cements (Grotzinger & Kasting 1993), commonly with crystalline microtextures linked to high Fe concentrations in seawater (Sumner & Grotzinger 1996).
Detailed petrological and facies analysis of iron formations has led to models with varying degrees of ocean stratification (e.g. Klein 2005; Beukes & Gutzmer 2008), but of concern to the present compilation are the broader trends in Earth's palaeoenvironmental evolution. Do the time-varying abundances of iron formations through the Archaean–Proterozoic transition represent fundamental changes in oceanography? At the older end of that age range, the answer is likely to be ‘no’. Trendall (2002) discussed the requirement of tectonic stability to allow accumulation of the giant iron formations, and with the exception of the c. 1.9 Ga deposits, many of the large iron formations appear at the first submergence of each craton following initial amalgamation. If so, the initial peak in iron formations between 2.7 and 2.4 Ga (Fig. 1k) has causative correlation with other records such as continental emergence (Fig. 1f). The apparently abrupt end of the first pulse of iron formations at c. 2.4 Ga remains to be tested by further geochronology (Krivoy Rog, Simandou), but its close temporal association with the rise of atmospheric oxygen strongly suggests a causal connection. Possible feedbacks involving phosphorus limitation on primary production (Bjerrum & Canfield 2002) have been disputed on quantitative grounds (Konhauser et al. 2007). However, fractionations of Fe isotopes in sedimentary pyrite appear to support the close coincidence between the timing of oxygenation in the atmosphere and the ocean (Rouxel et al. 2005).
Renewal of iron-formation deposition at 2.0–1.8 Ga could also indicate profound changes in the hydrosphere. The so-called ‘Canfield ocean’ model (Canfield 1998, 2005; Anbar & Knoll 2002) is the suggestion that after atmospheric oxidation, the increased weathering of continental sulphide minerals brought reactive sulphate ions into the marine realm, where sulphate-reducing bacteria responded with enhanced pyrite formation, thus stripping seawater of its hydrothermally generated Fe. In this model, the disappearance of iron formations at c. 1.8 Ga reflects the change to more reducing, sulphidic conditions rather than the deep ocean turning oxic. Meanwhile, the upper layer of seawater would remain oxygenated in contact with the post-GOE atmosphere; this first-order stratification is proposed to have persisted until the end of Precambrian time (Canfield et al. 2008). Despite some criticism (e.g. Holland 2006), the model has passed initial tests using depth-dependent iron speciation and sulphur stable isotopic variations in Mesoproterozoic marine sediments (Shen et al. 2002, 2003), patterns of trace metal concentrations (Anbar & Knoll 2002), isotopes (Arnold et al. 2004) and molecular biomarkers (Brocks et al. 2005).
Amid these novel geochemical proxies for hydrospheric and biospheric evolution, more traditionally studied isotopic systems show equally dramatic variations. One of the largest seawater carbon-isotopic anomalies in Earth history is known as the ‘Lomagundi’ or ‘Jatuli’ event respectively after the c. 2.1 Ga strata in Zimbabwe or Karelia where it was first documented (cf. Schidlowski et al. 1975) with typical values as high as +10
(Fig. 1m). An elegantly simple model attributed the peak to organic-carbon burial associated with evolution of oxygenic photosynthesis and, consequently, the rise of atmospheric oxygen (Karhu & Holland 1996). As discussed below, however, there is (contested) evidence that the advent of oxygenic photosynthesis could have preceded the isotopic peak by at least several hundred million years. Although the most enriched 13C values were probably generated in restricted, evaporative basins, such enrichments may well have been mere additions to a global peak, as indicated by the global extent of the signal (Melezhik et al. 1999, 2005a, b).
Rather than a single, long-lived Lomagundi-Jatuli event spanning the entire interval of 2.2–2.06 Ga (Karhu & Holland 1996), evidence has now accumulated for multiple positive excursions with intervening non-enriched values through that period (Buick et al. 1998; Bekker et al. 2001, 2006). Recent U–Pb dating on interstratified volcaniclastic rocks in Fennoscandia has constrained the end of the event to c. 2.06 Ga (Melezhik et al. 2007), but the earliest enrichments have been more difficult to date. The oldest strongly 13C-enriched carbonate units, in the Duitschland Formation of South Africa (Bekker et al. 2001), are older than 2316±7 Ma based on Re–Os of diagenetic pyrite from the stratigraphically overlying Timeball Hill shales (Bekker et al. 2004). Those same pyritic shales lack a non-mass-dependent sulphur-isotope fractionation signal (Bekker et al. 2004) and thus the onset of the Lomagundi-Jatuli event(s) coincides exactly, to the best knowledge of available ages, with atmospheric oxidation above 10–5 of present atmospheric levels. These events directly follow the only Palaeoproterozoic ice age with a 13C-depleted cap carbonate, which left a record on two palaeocontinents if the correlations of Bekker et al. (2006) are correct. The relationship between Palaeoproterozoic ice ages and Earth system evolution will be explored further, below.
The strontium isotopic composition of seawater, as recorded in carbonate rocks, involves more difficult analytical methods, and although well established for the Phanerozoic Eon, is relatively poorly constrained for Precambrian time (Shields & Veizer 2002). Data from the Palaeoproterozoic and surrounding intervals are sparse, and subject to uncertainties in age as well as the minimum-altered values. Nonetheless, a noticeable pattern of increasing 87Sr/86Sr, away from the inferred (slowly increasing) mantle value, characterizes the general Palaeoproterozoic trend (‘seawater’ curve in Fig. 1l). As Shields & Veizer (2002) noted, this probably indicates greater continental emergence and riverine runoff of radiogenic strontium through the Archaean–Palaeoproterozoic interval. Shields (2007) provided a more sophisticated model of these processes, including rough estimates of carbonate/evaporite dissolution as a contributor to the riverine runoff component, to produce a percentage estimate of the riverine contribution to seawater 87Sr/86Sr ratios through time (dashed curve in Fig. 1l). Even with this more detailed model, the first-order interpretation remains valid. However, the critical period for testing the connection with continental emergence (Fig. 1f), that is 2.7–1.9 Ga, is represented by a scant number of data (Shields & Veizer 2002).
Long-term secular evolution of oceanic chemistry can also be measured by compositional and sedimentological trends in carbonates and evaporites. Grotzinger (1989) and Grotzinger & James (2000) noted the abundance of carbonate platforms mirroring the growth of large sedimentary basins due to stable cratonization, much like the pattern observed for iron formations. The latter study also illustrated the successive peaks in ages of: aragonite crystal fans/herringbone calcite, tidal flat tufas, molar tooth structures, giant ooids, and various biogenic features of Archaean–Proterozoic sedimentary history. Not all of these records are well understood, but as noted above, disappearance of aragonite crystal fans and herringbone calcite is best correlated to the end of iron formation deposition and thus broadly to the rise of atmospheric oxygen.
Ocean palaeochemistry can also be inferred from the record of evaporite deposits, which in the Palaeoproterozoic era consist almost entirely of pseudomorphs after the original minerals. Evans (2006) compiled volume estimates for the largest evaporite basins through Earth history, summarized here in Figure 1n. A previous compilation (Grotzinger & Kasting 1993) described sulphate evaporites as old as c. 1.7 Ga and no older, but there is more recent recognition of common sulphate pseudomorphs in sedimentary successions at c. 2.2–2.1 Ga (reviewed by Pope & Grotzinger 2003; Evans 2006). Those gypsum- or anhydrite-bearing strata are usually associated with redbeds and 13C-enriched carbonates of the Lomagundi-Jatuli event (Melezhik et al. 2005b). Evaporite deposits of both younger (Pope & Grotzinger 2003) and older age (Buick 1992; Eriksson et al. 2005a) contain an evaporative sequence from carbonate directly to halite, excluding sulphate. Temporal correlation of halite-dominated evaporites with the peaks in iron-formation (Fig. 1k, n) conforms to the model of Fe-oxide seawater with low sulphate content (pre-2.4 Ga, plus 2.0–1.8 Ga), alternating with Fe-sulphidic deepwater driven by mildly oxidized surface water with higher sulphate content (Anbar & Knoll 2002). A rising oceanic sulphate reservoir through the Proterozoic Eon is also inferred by modelling rates of sulphur isotope excursions through sedimentary sections (Kah et al. 2004; Fig. 1n).
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The evolving palaeoclimate and biosphere |
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The eukaryotic fossil record, prior to 1.2 Ga, is equally controversial. Examples are described here in order of increasing age. At the younger end of the interval covered by this review (Fig. 1t), the record of bangiophyte red algae, in northern Canada, presents the oldest phylogenetically pinpointed eukaryotic body fossils (Butterfield 2000) and a robust starting point for considering older examples. The taxonomic affiliations of such older ‘eukaryotic’ fossils is inferred from their sizes and complexities (Knoll et al. 2006), including many simple and ornamented acritarchs, and various filamentous forms. Within the macroscale, the next older putative eukaryotic fossil is Horodyskia, found in c. 1.5–1.1 Ga strata in Western Australia and North America (reviewed by Fedonkin & Yochelson 2002; Grey et al. 2002; Martin 2004). Informally known as ‘strings of beads’, Horodyskia is difficult to place taxonomically; Knoll et al. (2006) consider it to be ‘a problematic macrofossil whose eukaryotic affinities are probable, but not beyond debate.’
The next two older macrofossils have been purported to preserve the trails of motile, multicellular organisms and are more contentious than the younger taxa outlined above. The first, in the Chorhat sandstone of the lowermost Vindhyan basin in India (Seilacher et al. 1998) attains Palaeoproterozoic antiquity on the merits of two concurrent and independent high-precision U–Pb studies (Rasmussen et al. 2002b; Ray et al. 2002). However, Seilacher (2007) has subsequently introduced a viable alternative hypothesis that the Chorhat traces could represent (biogenic?) gas structures trapped beneath a microbial mat. The second, consisting of discoidal and furrowed impressions in sandstone of the Stirling Ranges in Western Australia (Cruse & Harris 1994; Rasmussen et al. 2002a) has recently been dated to the interval 1960–1800 Ma (Rasmussen et al. 2004). An extensive discussion on these putative trace fossils retains the original interpretation of their being produced by ‘motile, mucus-producing, probably multicellular organisms’, which on the basis of size alone were probably eukaryotic (Bengtson et al. 2007). However, recent discovery of furrowed trails produced by extant Gromia amoebas may provide an adequate explanation for the Stirling biota (Matz et al. 2008); such an explanation needs further testing.
The oldest likely eukaryotic body fossil, Grypania spiralis, is found in the c. 1.88 Ga Negaunee iron-formation (Han & Runnegar 1992; Schneider et al. 2002), coeval with the spectacular palaeontological record of the nearby Gunflint Chert (Tyler & Barghoorn 1954; Fralick et al. 2002) and only slightly younger than the equally impressive Belcher Islands microflora (Hofmann 1976) at c. 2.0 Ga (Chandler & Parrish 1989). Classification of Grypania is based on its morphological similarity to Mesoproterozoic occurrences from North China and North America (Walter et al. 1990). A recent report describing a spinose acritarch in amphibolite-grade Archaean metasedimentary rocks of South Australia (Zang 2007) seems less convincing.
Apart from body fossils, evidence for eukaryotic life in the Palaeoproterozoic also includes the molecular biomarker record of steranes. Sterol biosynthesis is largely, although not entirely, limited to the eukaryotic realm (see discussion in Kirschvink & Kopp 2008; Waldbauer et al. 2009). Sensationally old steranes were identified in the 2.7 Ga Jeerinah Formation of Western Australia (Brocks et al. 1999, 2003a, b). However, Rasmussen et al. (2008) attributed these signals to a secondary fluid migration into the boreholes, at some unknown time after c. 2.16 Ga regional metamorphism. There are other Palaeoproterozoic sterane biomarker records. Dutkiewicz et al. (2006) and George et al. (2008) found them in sediments of the basal Huronian Supergroup (c. 2.4 Ga), with a signal that pre-dates c. 1.9 Ga Penokean metamorphism, and Dutkiewicz et al. (2007) discovered them in the c. 2.1 Ga Francevilian series of Gabon, with a signal that pre-dates supercritcality of the Oklo natural nuclear reactor at 1.95±0.04 Ga (Gauthier-Lafaye & Weber 2003).
Waldbauer et al. (2009) conducted a benchmark study in attempts to demonstrate the syngeneity of their observed sterane biomarker signal, obtained from c. 2.65–2.45 Ga strata in South Africa. The molecular fossils are described as pre-metamorphic, and vary according to sedimentary facies in correlative sections from adjacent drillcores. Nonetheless, the carbonate formations in those drillcores have been pervasively remagnetized at about 2.2–2.1 Ga (de Kock et al. 2009), indicating basinwide low-grade hydrothermal fluid infiltration-at the same age within error as, and possibly in direct palaeogeographic proximity to, the regional metamorphic event on the adjacent Pilbara craton as described by Rasmussen et al. (2008).
We return to the Palaeoproterozoic glacial deposits, which are classically used to infer the palaeoenvironmental conditions in which these biological innovations occurred. Following an almost entirely ice-free Archaean history, the Palaeoproterozoic world was exposed to at least three ice ages, which appear to have penetrated deep into the tropics (Evans et al. 1997, global constraints reviewed by Evans 2003). These ice ages, lesser known but seemingly of equal severity to their more widely publicized Neoproterozoic ‘snowball Earth’ counterparts (Hoffman & Schrag 2002), are generally rather poorly constrained in age to within the interval 2.45–2.22 Ga. As with the Neoproterozoic ice ages, estimating the number of Palaeoproterozoic glaciations is complicated by the fact that the diamictites themselves are commonly the principal items of correlation among cratons through this interval (e.g. Aspler & Chiarenzelli 1998; Bekker et al. 2006).
The end of the second among three ice ages, recorded in the Huronian succession and correlative strata in Wyoming, is marked by a 13C-depleted ‘cap carbonate’ unit that may be broadly comparable to those better developed after Neoproterozoic ice ages (Bekker et al. 2005). The South African sections contain two distinct sequences of diamictite and overlying 13C-depleted carbonate (Bekker et al. 2001), unconformably overlain by a third diamictite, the Makganyene Formation, in turn overlain by flood basalt and variably Mn-rich carbonate and ironstone units (Kirschvink et al. 2000) with near-zero 13C values (Bau et al. 1999). Palaeomagnetic data from the flood basalt indicates deep tropical palaeolatitudes, constituting the best evidence of its kind for a Palaeoproterozoic snowball Earth event (Evans et al. 1997; Kirschvink et al. 2000; Evans 2003). Within the limits of existing age constraints, the Makganyene ice age could be correlative with the uppermost Huronian glaciation at 2.23 Ga (Bekker et al. 2006), or, all three Huronian glacial levels could be distinctly older (Kopp et al. 2005). Given that the Lomagundi-Jatuli positive carbon-isotope excursion(s) began as early as 2.32 Ga, the near-zero 13C values in the post-Makganyene carbonate units are anomalously negative and warrant comparison with other Proterozoic postglacial cap carbonate sequences. Considering the high greenhouse forcing required to offset the Palaeoproterozoic ‘faint young Sun’ (Sagan & Mullen 1972), escape from any ‘snowball’ climate regime of that age would have required tens of millions of years of volcanic outgassing uncompensated by silicate weathering (Tajika 2003). If all of the Palaeoproterozoic glacial deposits represent so-called ‘hard’ snowball ice ages, then Earth's panglacial climate mode would have occupied a substantial fraction of time in the 2.45–2.22 Ga interval.
As discussed above, the disappearance of non-mass-dependent sulphur isotope fractionation and the onset of highly 13C-enriched carbonates of the Lomagundi isotopic event are both located stratigraphically within the broad age range of these glaciations. More precisely, if the Bekker et al. (2006) correlations between North America and South Africa are correct, then the oldest carbonate-capped glacial deposits (Bruce and Rooihoogte Formations) constitute the stratigraphical boundary between two fundamentally distinct states of Earth's palaeoenvironment. Closely below this level are the final vestiges of detrital pyrite/uraninite deposition (Roscoe 1973) and non-mass-dependent sulphur isotope fractionation (Papineau et al. 2007). Closely above the level are the entirely mass-dependent-fractionated pyrites of the Timeball Hill Formation, dated at 2.32 Ga and representing the rise of atmospheric oxygen (Bekker et al. 2004; Hannah et al. 2004) and the oldest carbonates with strongly enriched 13C values indicating the onset of the Lomagundi-Jatuli isotopic excursions (Bekker et al. 2001). Glacial deposits with cap carbonates thus appear to be closely related to the rise of atmospheric oxygen. Collapse of the methane-rich, pre-2.4-Ga greenhouse due to atmospheric oxygenation (see above and Kasting 2005) could well be a trigger for the low-latitude ice ages, perhaps in addition to the silicate weathering removal of carbon dioxide due to the widespread and largely subaerial large igneous provinces at 2.45 Ga (Melezhik 2006; Kump & Barley 2007). But the ice ages themselves could also have contributed further to rapid pulses of oxygen production: Liang and co-workers (Liang et al. 2006) postulated the mechanism of hydrogen peroxide trapping in ice throughout the duration of a ‘hard’ snowball stage, which would be released suddenly to the oceans upon deglaciation.
This glaciation-oxygenation scenario has been developed further, as reviewed by Kirschvink & Kopp (2008). In that model, the hydrogen peroxide plume into the oceans upon panglacial melting would constitute the evolutionary driver of intracellular oxygen-mediating enzymes, which are seen as a necessary precursor to oxygenic photosynthesis. The post-Makganyene sequence, extraordinarily rich in Mn, would represent the final and irreversible oxidation of the deep oceans (Kirschvink et al. 2000). Two outstanding problems with the timing of the model are as follows: (1) the Lomagundi-Jatuli positive isotopic excursion, apparently requiring burial of photosynthetically produced organic carbon, is found in pre-Makganyene strata (Bekker et al. 2001); and (2) increasingly more rigorous biomarker studies provide compelling evidence for photosynthetic organisms as old as 2.7 Ga (e.g. Eigenbrode et al. 2008; Waldbauer et al. 2009).
Following the well known 2.45–2.2 Ga ice ages, the next nearly 1.5 billion years has traditionally been noted as entirely ice-free (Evans 2003). Recent documentation of periglacial features at c. 1.8 Ga (Williams 2005) contests this conclusion. Nonetheless, the dominant climate state was nonglacial throughout most of Palaeoproterozoic–Mesoproterozoic time. Even in a mildly oxygenated, post-GOE atmosphere, methane is increasingly favoured as a minor but powerful greenhouse gas to combat the low luminosity of the Mesoproterozoic Sun (Pavlov et al. 2003; Kasting 2005; Kah & Riding 2007).
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Extra-terrestrial influences: bolides and Earth-Moon orbital dynamics |
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30 km diameter) craters in the 3.0–1.2 Ga time interval: Yarrabubba at an unknown age younger than its c. 2.65 Ga target rocks (Macdonald et al. 2003), Keurusselkä at an unknown age younger than its c. 1.88 Ga target rocks (Hietala et al. 2004) and Shoemaker with a maximum age of 1.63 Ga (Pirajno et al. 2003, 2009). As there are no older preserved impact craters than c. 2.4 Ga (Earth impact database 2009), all knowledge of prior impact history must be inferred from the lunar record, or determined from ejecta beds containing either spherules (Simonson & Glass 2004) or anomalous concentrations of siderophile elements (Glikson 2005). For the time interval investigated here, the most prominent impact record is found in the sedimentary cover of the Vaalbara supercraton. At least three distinct spherule beds can be recognized; some readily correlated between Australia and South Africa, within the interval 2.63–2.49 Ga (Simonson et al. 2009). It is unknown what effects these impacts, undoubtedly a small subset of the total Archaean–Palaeoproterozoic bolide flux to the Earth, had on the ancient surface environment. Completion of the IGCP509 global stratigraphic database (Eglington et al. 2009) will help identify suitable sedimentary basins for finding ejecta blankets from the large craters described above.
The orbital parameters of Earth and the Moon can be gleaned from the sedimentary record of tidal rhythmites, which are usually in fine-grained mudstones and siltstones (e.g. Williams 2000). They can also be found in sandstone crossbed foresets (Mazumder 2004; Mazumder & Arima 2005), including the oldest rhythmites in the geological record at c. 3.2 Ga (Eriksson & Simpson 2000). The most complete calculation of orbital parameters from tidal rhythmites can be found in Williams (2000), who listed two alternative calculations for the 2.45 Ga Weeli Wolli banded iron formation in the Hamersley Ranges of Western Australia: one assuming that the lamina couplets (microbands) represented annual increments, the other assuming that they represented fortnightly cycles. Trendall (2002) has discussed how the annual microband model is consistent with U–Pb age constraints through the Hamersley succession, and by implication, that the fortnightly alternative is not. The annual microband model predicts 17.1±1.1 hours in the solar (Earth) day, 514±33 solar days per year, and an Earth–Moon distance of 51.9±3.3 Earth radii (compared to 60.27 at present) for the earliest Palaeoproterozoic (Williams 2000).
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Discussion |
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Many of the physical models used to extrapolate modern planetary dynamics back into deep time suffer from the necessary simplifications of tractability, and many of the historical data proxies are incomplete or contentious. However, an emphasis on interdisciplinary Earth-system science has led to multiple working hypotheses for the interrelationships among the various proxy records, and we are particularly inspired by the following 10 recent developments.
Supercontinents may indeed be the centrepiece of the long-term Earth system, but their history is one of the least constrained elements in Figure 1. Nonetheless, there is hope for eventual understanding. New isotopic methods for precise geochronological calibration of deep time have made these comparisons possible. The stratigraphies of most of the world's Precambrian cratons are now increasingly constrained by acquisition of such precise rock ages, and there is no sign of slowing. Dedicated global working groups such as IGCP projects have fostered frequent, direct communication among researchers around the world. An emphasis on interdisciplinary science has led to multiple working hypotheses for the interrelationships among the various proxy records.
From core to surface, our next major advances will likely arise through determining accurate ways to measure ancient geomagnetic field strength; obtaining more complete records of mantle plume activity; developing novel methods for estimating crustal growth and continental emergence; solving the palaeogeography of supercontinents and supercratons, with consequent ‘ground-truthing’ of proposed tectonic processes and mineral deposit evolution; creating new chemical and isotopic proxies for the evolution of mantle, crust, and surface; dating these records precisely with new analytical techniques; and integrating these strands of data into robust geodynamic models. With these continuing advancements the Archaean–Proterozoic transition is at last coming into focus.
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Acknowledgments |
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33S data, summarized by