Oxygenation of the Earth's atmosphere–ocean system: A review of physical and chemical sedimentologic responses

Marine and Petroleum Geology 32 (2012) 1e20

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Review article

Oxygenation of the Earth’s atmosphereeocean system: A review of physical and chemical sedimentologic responses P.K. Pufahl a, *, E.E. Hiatt b,1 a b

Department of Earth & Environmental Science, Acadia University, Wolfville, Nova Scotia B4P 2R6, Canada Department of Geology, University of Wisconsin Oshkosh, Oshkosh, WI 54901, USA

a r t i c l e i n f o

a b s t r a c t

Article history: Received 22 August 2011 Received in revised form 30 November 2011 Accepted 5 December 2011 Available online 14 December 2011

The Great Oxidation Event (GOE) is one of the most significant changes in seawater and atmospheric chemistry in Earth history. This rise in oxygen occurred between ca. 2.4 and 2.3 Ga and set the stage for oxidative chemical weathering, wholesale changes in ocean chemistry, and the evolution of multicelluar life. Most of what is known about this important event and the subsequent oxygenation history of the Precambrian Earth is based on either geochemistry or “data mining” published literature to understand the temporal abundance of bioelemental sediments. Bioelemental sediments include iron formation, chert, and phosphorite, which are precipitates of the nutrient elements Fe, Si, and P, respectively. Because biological processes leading to their accumulation often produce organic-rich sediment, black shale can also be included in the bioelemental spectrum. Thus, chemistry of bioelemental sediments potentially holds clues to the oxygenation of the Earth because they are not simply recorders of geologic processes, but intimately involved in Earth system evolution. Chemical proxies such as redox-sensitive trace elements (Cu, Cr, V, Cd, Mo, U, Y, Zn, and REE’s) and the ratio of stable isotopes (d56Fe, d53Cr, d97/95Mo, d98/95Mo, d34S, D33S) in bioelemental sediments are now routinely used to infer the oxygenation history of paleo-seawater. The most robust of these is the massindependent fractionation of sulfur isotopes (MIF), which is thought to have persisted under essentially anoxic conditions until the onset of the GOE at ca. 2.4 Ga. Since most of these proxies are derived from authigenic minerals reflecting pore water composition, extrapolating the chemistry of seawater from synsedimentary precipitates must be done cautiously. Paleoenvironmental context is critical to understanding whether geochemical trends during Earth’s oxygenation represent truly global, or merely local environmental conditions. To make this determination it is important to appreciate chemical data are primarily from authigenic minerals that are diagenetically altered and often metamorphosed. Because relatively few studies consider alteration in detail, our ability to measure geochemical anomalies through the GOE now surpasses our capacity to adequately understand them. In this review we highlight the need for careful consideration of the role sedimentology, stratigraphy, alteration, and basin geology play in controlling the geochemistry of bioelemental sediments. Such an approach will fine-tune what is known about the GOE because it permits the systematic evaluation of basin type and oceanography on geochemistry. This technique also provides information on how basin hydrology and post-depositional fluid movement alters bioelemental sediments. Thus, a primary aim of any investigation focused on prominent intervals of Earth history should be the integration of geochemistry with sedimentology and basin evolution to provide a more robust explanation of geochemical proxies and ocean-atmosphere evolution. Ó 2011 Elsevier Ltd. All rights reserved.

Keywords: Great oxidation event Earth oxygenation Ocean-atmosphere evolution Bioelemental Chemistry Alteration Sedimentology Diagenesis

1. Introduction

* Corresponding author. Tel.: þ1 902 585 1858; fax: þ1 902 585 1816. E-mail address: [email protected] (P.K. Pufahl). 1 Tel.: þ1 920 424 7001; fax: þ1 920 424 0240. 0264-8172/$ e see front matter Ó 2011 Elsevier Ltd. All rights reserved. doi:10.1016/j.marpetgeo.2011.12.002

One of the most intensely debated topics in the Earth sciences is the oxygenation of the Earth’s atmosphere and oceans, primarily because of their co-evolution with early life (e.g. Kasting, 1993; Catling et al., 2001; Canfield, 2005; Fedonkin, 2009). Spirited discussion began in 1964 with the publication of the “The Origin

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and Evolution of Atmospheres and Oceans” (Brancazio and Cameron, 1964). In 1973 dialogue shifted away from the notion that purely abiotic processes produced the early atmosphere when Cloud suggested that the deposition of large Paleoproterozoic iron formations was linked to a rise in photosynthetic oxygen (Cloud, 1973). More recently, Kasting and Siefert (2002) summarized the contemporary understanding of the influence of early life on the composition of the atmosphere. “microorganisms have probably determined the basic composition of the Earth’s atmosphere since the origin of life.” Holland (2002) hypothesized that the emergence of an aerobic biosphere did not represent a simple change in the volume of volcanic outgassing, but instead was related to a change from reducing to oxidizing volcanic gases. Zahnle et al. (2006) and Konhauser et al. (2009) proposed a decrease in atmospheric methane was the catalyst. Although it occurred over an extended interval of time (Wille et al., 2007; Voegelin et al., 2010), this rise in oxygen has become known as the Great Oxidation Event (GOE; Holland, 2002, 2006) and occurred between ca. 2.4 and 2.3 Ga (Fig. 1; Bekker et al., 2004; Holland, 2004, 2006; Frei et al., 2009; Guo et al., 2009). It marks the beginning of one the most significant changes the Earth has experienced, setting the stage for oxidative chemical weathering, wholesale

changes in ocean chemistry, and the evolution of multicellular life (Fig. 1). The first evidence for the oxygenation of the atmosphere was based on mineralogical changes with reduced detrital mineral phases such as pyrite and uraninite in sedimentary rocks giving way to hematite and other oxide phases (e.g. Cloud, 1968; Roscoe, 1969; Fleet, 1998; Rasmussen and Buick, 1999; Hazen et al., 2008). Most new data regarding the GOE, however, is geochemical in nature. Proxies such as trace element compositions (Cu, Cr, V, Cd, Mo, U, Y, Zn, and REE’s) and the ratio of stable isotopes (d56Fe, d53Cr, d97/95Mo, d98/95Mo, d34S, D33S) in iron formation, phosphorite, and black shale are now routinely used to indirectly deduce the redox conditions of paleo-seawater (Fig. 2A, B, C, D; Table 1). Iron formation, phosphorite and black shale are bioelemental sedimentary rocks that form from the nutrient elements Fe, P, and C, which are required for myriad life processes (Pufahl, 2010). Since the precipitation of these elements is so closely linked to biology, bioelemental sediments are not simply recorders of geologic processes, but are intimately involved in the evolution of the ocean-atmosphere system (e.g. Föllmi et al., 1993; Glenn et al., 2000; Simonson, 2003; Huston and Logan, 2004; Maliva et al., 2005; Holland, 2006; Bekker et al., 2010; Pufahl, 2010; Konhauser et al., 2011). Thus, their chemistry holds potentially

Figure 1. Seawater chemistry and Earth events as related to the three stages of ocean-atmosphere oxygenation (1, 2, 3). The degree of oxygenation immediately after the GOE is still largely unknown, but recent d53Cr data suggests that at ca. 1.9 Ga oxygen levels may have dipped to pre-GOE concentrations (Frei et al., 2009). See Figure 2 and Table 1 for a more complete summary of the geochemical data for Earth’s oxygenation. PAL ¼ present atmospheric levels; MIF ¼ mass-independent fractionation. Based on data from Farquhar et al. (2000), Condie et al. (2001), Canfield (2005), Fedonkin (2009), Johnston et al. (2009), Lyons and Reinhard (2009), Konhauser et al. (2011), and Nelson et al. (2010).

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Figure 2. Geochemical proxies used to understand Earth’s oxygenation. A) d34S data from sedimentary sulfides showing an increase in fractionation after ca. 2.4 Ga (Canfield, 2005). 33 The double dashed line is the estimated range in d34S values for SO2 4 . Lower dashed line is the maximum fractionation with sulfide. B) D S data from sedimentary sulfides (Farquhar et al., 2000; Farquhar and Wing, 2003). Mass independent S fractionations of 32S, 33S, and 34S indicate low atmospheric oxygen levels from ca. 3.8e3.0 Ga, an increase from ca. 2.7 to 2.4 Ga, and a permanent rise after ca. 2.4 Ga. The yellow horizontal line represents the range of values for mass dependent fractionation of S isotopes. C) d56Fe data from bulk shale samples, iron formations, and pyrite (Johnson et al., 2008). The yellow horizontal line marks the range in d56Fe values for Archean to modern, low-C and low-S clastic sedimentary rocks. Increased fractionation between ca. 2.7 and 2.5 Ga is the likely consequence of rising photosynthetic oxygen. D) d53Cr data from iron formations (Frei et al., 2009). The yellow horizontal line shows the range of values of magmatic Cr3þ-rich ores and minerals formed under high temperatures. Increased fractionation between ca. 2.8 and 2.6 Ga suggests a “whiff” or transient oxygen levels prior to the GOE. Decreased fractionation at ca. 1.9 Ga may record pre-GOE oxygen levels. E) Ni/Fe mole ratios for iron formations (Konhauser et al., 2009). Decline in Ni at ca. 2.7 Ga may have limited methanogens and contributed to the GOE.

important clues regarding the development of the early oceans and by extension, the atmosphere making them a logical target for application of new geochemical techniques. This attribute, together with the unprecedented development of technology, has spurred the recent surge in the geochemical investigation of Precambrian bioelemental sedimentary and meta-sedimentary rocks.

Although these technological advancements are resulting in publication of numerous datasets, it is problematic that our ability to measure chemical anomalies now surpasses our capacity to adequately understand them (Watson, 2008). This problem is exacerbated because data are often interpreted with little regard to sedimentology, stratigraphy, alteration, and basin evolution. Such

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Table 1 Geochemical proxies used to understand the Great Oxidation Event. Proxy

Environmental parameter

Host mineral(s)

Deposit type

Effects of alteration

d56Fe

Seawater Fe-(oxyhydr)oxide levels and effects on bacterial DIR. Seawater sulfate levels and effects on BSR.

Hematite Magnetite Siderite Pyrite Pyrite

Iron formation Unknown Black shale

Comments

Marked increase in fractionation between ca. 2.7 and 2.5 Ga reflects extensive radiation in DIR resulting from increased production of FeO. Photosynthetic oxygen likely caused the oxidation of Fe2þ to create FeO. Produced lower d56Fe values. d34S Black shale Unknown Marked increase in fractionation at ca. 2.4 Ga is coincident with GOE. Interpreted to record the transition from sulfate limited to sulfate unlimited bacterial sulfate reduction. Increased fractionation led to more variability in d34S values. 33 D S Absence of ozone and Pyrite Black shale Unknown MIF of S isotopes interpreted to record absence of free oxygen UV shield and effects on prior to the GOE. End of MIF interpreted to reflect development MIF of S isotopes. of ozone layer and a UV shield associated with the GOE. d53Cr Seawater and Iron formation Unknown, but Increase in the fractionation of Cr isotopes between ca. 2.8 and Cr3þ-oxides atmospheric oxygen associated inferred immobile. 2.6 Ga is interpreted to record a “whiff” of oxygen prior to the levels and generation of with FeO GOE. Decrease in the fractionation at ca. 1.9 Ga likely records 6þ Cr through oxic a dip in oxygen levels to pre-GOE values. Cr6þ is delivered to the chemical weathering. oceans during oxic chemical weathering and becomes immobile when reduced by Fe2þ to precipitate Cr3þ oxides associated with FeO. d97/95Mo, d98/95Mo Mo oxide levels in Mo sulfide Black shale Unknown Mo isotopic values suggest euxinic conditions prevailed after seawater and effects the GOE between ca. 1.4 and 1.7 Ga. Mo is removed from on the fractionation seawater by oxic adsorption processes. The isotopic of Mo isotopes. composition of these oxides is thought to be transferred to authigenic Mo sulfides precipitated under reducing conditions beneath the seafloor. Iron formation Interpreted to Provide information on whether the water column was oxygen REEs (negative Seawater oxygen Ce3þ on MneFeO. Phosphorite preserve a primary stratified. Because scavenging of Ce3þ-oxides by FeO is Ce anomaly) concentrations negligible; Ce4þ is scavenged on the surfaces of MneFeO signature. and Ce behaviour. producing the negative Ce anomaly. In this way the resulting low Ce concentration in seawater is transferred to the sediment. Trace elements Seawater and pore U oxide Iron formation Unknown A negative U anomaly and elevated Cr records accumulation water redox recorded Cr hydroxide Black shale under suboxic and oxic conditions. Elevated U, V, Cu, Cd, Zn, Mo, by differences in the V oxide Phosphorite and Ni reflects deposition under anoxic conditions. Such concentrations of Cr, U, Cu, Cd, Zn, Mo, differences in the trace element concentrations of shallow- and V, Cu, Cd, Zn, Mo, and and Ni sulfides deep-water lithofacies can indicate whether the water column Ni in sediment. was oxygen stratified. Mo enrichment in Mo sulfide Black shale Unknown Mo enrichment in black shale suggests a “whiff” of oxygen 50 seawater million years prior to the GOE. Increased delivery of Mo to the oceans via oxic chemical weathering is thought to have led to Mo enrichment in black shales that accumulated within anoxic environments. Ni decline in seawater Ni adsorbed to FeO Iron formation Unknown A decline in the Ni concentration of iron formation at ca. 2.7 Ga is interpreted to have contributed to the GOE by limiting methanogens. Ni is a bioessential nutrient for methanogens and without it their development was apparently limited allowing oxygen to accumulate in the atmosphere. Notes: Also see Figure 2. DIR ¼ dissimilatory iron reduction; FeO ¼ Fe-(oxyhydr)oxides; BSR ¼ bacterial sulfate reduction; MIF ¼ mass-independent fractionation; UV ¼ ultraviolet light; MneFeO ¼ Mn-Fe-(oxyhydr)oxides; Corg ¼ organic matter. Although many of these proxies are inferred to be directly related to seawater composition, because their host minerals are authigenic they in fact reflect processes that operated beneath the seafloor. Data are from Jarvis et al. (1994), Farquhar et al. (2000); Canfield (2001), Arnold et al. (2004), Klein (2005), Rouxel et al. (2005), Anbar et al. (2007), Johnson et al. (2008), Frei et al. (2009), Bekker et al. (2004, 2010), Konhauser et al. (2009), Lyons et al. (2009), Planavsky et al. (2009), Severmann and Anbar (2009).

context is critical to understanding whether anomalies represent paleoenvironmental conditions, are truly global in character, the result of local environmental factors, or the consequence of alteration of what largely are metamorphic rocks. The sedimentary record of the GOE spans ca. 100 million years and provides an excellent opportunity to examine the effect of this global geochemical revolution (cf. Watson, 2008) on interpreting major Earth events. The picture that has emerged of the Earth’s oxygenation is based almost exclusively on geochemistry. This approach has provided the broad brush-strokes required to understand this interval, but the fine lines necessary to refine this picture are only attainable by integrating geochemical data in a sedimentologic framework that permits the interpretation of depositional environments, oceanography, and subsequent alteration. The purpose of this review is to summarize what is known about the GOE from the bioelemental sedimentary record, and to re-examine the connection between sedimentology, basin history, and the geochemical proxies used to elucidate changes in ocean-atmosphere oxygenation.

2. The Great Oxidation Event and history of Earth’s oxygenation Although a great deal of controversy still exists about the oxygenation of the Earth (compare Holland, 2004 and Hoashi et al., 2009), there is a consistent interpretation of low Archean and Early Paleoproterozoic atmospheric oxygen levels (<1e100 ppm O2 in the atmosphere), which are followed by higher concentrations during the GOE that, after nearly a billion years, gave way to fully oxygenated conditions in the latest Neoproterozoic (Fig. 1; Holland, 2004, 2006; Canfield, 2005; Canfield et al., 2007; Narbonne, 2010). These stages are complex and multi-causal, and defined by times of significant change in the redox state of the ocean-atmosphere system (Huston and Logan, 2004; Canfield, 2005; Holland, 2006; Reddy and Evans, 2009). Support for the very low levels of oxygen prior to the GOE comes from the presence of detrital grains composed of reduced minerals, such as pyrite and uraninite, in sedimentary successions (e.g. Cloud,

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1968; Roscoe, 1969; Fleet, 1998; Rasmussen and Buick, 1999; Hazen et al., 2008), and more recently, non-mass-dependent sulfur isotope fractionation, which provides a proxy of oxygen in the atmosphere (Farquhar et al., 2000; Holland, 2006; Reddy and Evans, 2009; Lyons and Gill, 2010). Analysis of Archean and early Paleoproterozoic sedimentary sulfide and sulfate minerals has yielded anomalous variations in the abundance of the four stable isotopes of sulfur (32S, 33S, 34S, 36S). These anomalies are interpreted to result from mass-independent fraction (MIF; D33S, D36S; Figs. 1 and 2B; Table 1) involving gaseous sulfur species in the Precambrian atmosphere with coeval mixing into seawater that was marked by low sulfate concentrations (Farquhar et al., 2000; Canfield et al., 2000). MIF is driven by photochemical reactions involving high UV light flux. A prerequisite for these photochemical reactions is the absence of an effective UV shield such as ozone (Farquhar et al., 2000). Thus, the expression of MIF in sulfur isotopes is interpreted to reflect the near absence of free oxygen in the Archean and Early Paleoproterozoic (Farquhar et al., 2000, 2007). Although “whiffs” of oxygen are suggested in the Archean (Ohmoto et al., 2006; Anbar et al., 2007; Wille et al., 2007; Hoashi et al., 2009; Kato et al., 2009; Reinhard et al., 2009), the MIF of sulfur suggests that O2 concentrations in the Archean atmosphere were generally <105 of PAL (Kasting et al., 2001). Evidence from Mo isotopes and PGE concentrations, however, suggest that oxygen levels may have begun to rise between 2.7 and 2.5 Ga suggesting that the increase of atmospheric oxygen that led to the demise of MIF was not a simple linear trend (e.g. Wille et al., 2007). The end of MIF at ca. 2.4 Ga (Figs. 1 and 2) is interpreted to record the onset of the GOE (Bekker et al., 2004). During the following 100 million years oxygen levels are interpreted to have risen to >102 PAL

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(>0.2% or 2000 ppm; Pavlov and Kasting, 2002; Lyons and Reinhard, 2009). What is not known is whether oxygen levels through this protracted interval rose slowly or quickly, or whether the increase was constant, marked by punctuated increases, or some combination of these (compare Bekker et al., 2004; Ohmoto et al., 2006; Holland, 2006; Wille et al., 2007; Anbar et al., 2007; Lyons and Reinhard, 2009). Prior to the advent of oxygenic photosynthesis low oxygen levels were probably maintained in the Archean atmosphere and surface ocean by photo-dissociation of H2O molecules (Kasting et al., 1979). Photochemical breakdown of H2O releases H2O2, which in turn dissociates creating O2 (Kasting et al., 1985). Kasting and Walker (1981) determined that Archean oxygen concentrations would have been between 1012 and 1014 PAL in the presence of volcanic outgassed H2 and CO, but up to 4  108 PAL in the absence of such gases. Although low, these concentrations would have produced enough O2 to cause precipitation of hematite on the continents (Kasting and Walker, 1981). This suggests that red beds (Fig. 3A, B, C, D) should have formed long before the GOE (Kasting and Walker, 1981), yet the appearance of red beds in the stratigraphic record is often cited as evidence for the GOE (Fig. 1; Holland, 2002). The answer to this paradox lies in how red beds form. Walker (1976) showed that red beds are preserved during burial diagenesis in the presence of oxygenated groundwater when Fe-(oxyhydr)oxides that coat grains (Fig. 3C) recrystallize to form hematite (Fig. 3A, B, C). However, if groundwater was anoxic, iron (oxyhydr)oxides (Fig. 3A) formed at the Earth’s surface would have dissolved during early burial before quartz cement overgrowths could precipitate and protect these coatings, leaving no record of oxidation (e.g. Surdam and Crossey,

Figure 3. Development of red beds. A) Outcrop photo of the Eocene White River Formation, eastern Wyoming, USA. This sandstone is stained with Fe-(oxyhydr)oxides (limonite and goethite) that form “dust rims” on detrital grains; these coatings are concentrated on slightly more permeable laminae and highlight cross bed foresets. The metastable Fe-(oxyhydr) oxides will eventually recrystallize to form hematite making the rock red. B) Bright red hematite-stained quartz arenite and siltstone red beds from lacustrine facies of the 1.9 Ga Roraima Group, Guyana, South America. C) Photomicrograph in plane-polarized light from the 1.7 Ga Thelon Formation, Nunavut, Canada. Detrital quartz grains (Dq) in this eolian facies are coated with hematite dust rims (Hr) that underlie pore-filling quartz cement giving this quartz arenite a red color. D) Outcrop showing an upturned, ripple-marked bedding plane of the Paleoproterozoic (ca. 2.3 Ga) Lorrain Formation (fluvial facies), Huronian, Blind River, Ontario. This red bed succession was one of the examples originally used as evidence for the GOE. Photo courtesy of Steve Beyer.

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1987; Lovley et al., 1991). Thus, the appearance of red beds coincides with the GOE because groundwater became sufficiently oxygenated to retain Fe-(oxyhydr)oxides in the shallow burial realm and potentially preserve them in the sedimentary record. Most researchers agree that the evolution of oxygenic photosynthesis within cyanobacteria was the source of oxygen that caused the GOE (cf. Cloud, 1973). The timing of cyanobacterial evolution, however, is problematic since biomarkers indicate they may have evolved as early as ca. 2.9 Ga (Nisbet et al., 2007) and were abundant by ca. 2.7 Ga (Brocks et al., 2003, 2005; Canfield, 2005; Buick, 2008), at least 400 million years before the onset of the GOE. This long lag likely represents a period of inertia where oxygen-consuming chemical reactions prevented the rise of photosynthetic oxygen by consuming oxygen in inorganic reactions with reduced mineral phases and organic matter in the oceans (François and Gérard, 1986; Goldblatt et al., 2006; Saito, 2009). Recent molecular clock analyses of cyanobacteria lineages by Blank and Sánchez-Baracaldo (2010) further suggest the earliest cyanobacteria were restricted to freshwater environments until ca. 2.4 Ga when they diversified and exploited marine ecosystems. This diversification, extraordinary increase in habitat, and the resulting extensive organic carbon flux to the deep oceans could have caused a rapid increase in oxygen during the GOE. Based on this same molecular clock model, mat-forming cyanobacteria with filamentous forms, large sizes and that fixed nitrogen appeared at ca. 2.3 Ga (Blank and Sánchez-Baracaldo, 2010). Although photosynthetically produced oxygen was the primary driver of oxygenation during the GOE, a number of processes are postulated to have played a role in changing the redox state of thePaleoproterozoic ocean-atmosphere system. These include: (1) increased burial of organic matter (Des Marais et al., 1992; Melezhik et al., 2005); (2) loss of hydrogen to space from a methane-rich atmosphere (Kasting et al., 1979; Catling et al., 2001); (3) collapse of atmospheric methane (Zahnle et al., 2006; Konhauser et al., 2009); (4) changes in the redox potential of volcanic gases (Kump et al., 2001; Holland, 2002); (5) nutrient loading and increased production of cyanobacterial oxygen (Papineau et al., 2009); and (6) a period of major continental growth at the Archean-Proterozoic boundary (Godderis and Veizer, 2000). The collapse of a methane-rich atmosphere is also thought to have been an important contributor to the onset of Paleoproterozoic ice ages (Reddy and Evans, 2009). Little is known regarding oxygen levels immediately following the GOE (2.0e1.8 Ga; Fig. 1). Frei et al. (2009) interpret a decrease in oxygen, possibly dropping to pre-GOE levels, based on a change in chromium isotopic values from a small dataset (Fig. 2D; Table 1). Because a major interval of black shale deposition corresponds to this interval (Fig. 1), however, it is likely that any decrease in photosynthetic oxygen production would be at least partially compensated for by removal of organic carbon from the oceanatmosphere system. Thus, this interval should be explored further to elucidate whether there was in fact a major dip in oxygen concentration following the GOE (Fig. 1). Oxygen levels during the Earth’s middle age (ca. 1.85e0.85 Ga) apparently stabilized somewhere between 1 and 10% PAL (Fig. 1; Lyons and Reinhard, 2009). Such oxygen levels are hypothesized to have led to oxic chemical weathering of the continents, which oxidized sulfide minerals to produce sulfate that was delivered by rivers to the ocean. In this model, dissolved sulfate delivered to the oceans by rivers was transformed through bacterial sulfate reduction into sulfide causing euxinic conditions that developed at the end of the Paleoproterozoic (Canfield, 1998, 2005; Poulton and Canfield, 2011; Kendall et al., 2011). By ca. 1.85 the flux of sulfate was great enough to cause sulfidic intermediate and bottom waters (Fig. 1; Poulton et al., 2004; see also Pufahl et al., 2010). Widespread euxinia may have been perpetuated by thriving anoxygenic

photoautotrophs that tempered oxygen production by using sulfide as an electron donor (Johnston et al., 2009). These conditions are hypothesized to have prevailed for nearly a billion years and also perturbed the cycling of bioessential elements, possibly causing a long stasis in the evolution of eukaryotes (Anbar and Knoll, 2002). This period is often referred to as the “Boring Billion” because biological evolution is thought to have stagnated during this protracted interval (Anbar and Knoll, 2002; Holland, 2006). Oxygen concentrations increased to >10% PAL (>0.2% or 2000 ppm) during the Neoproterozoic ‘snowball glaciations’ (Fig. 1; Canfield, 2005; Holland, 2006). Ice cover that shrouded the Earth between ca. 740 and 630 Ma is thought to have slowed chemical weathering and delivery of sulfate to the oceans, causing the demise of widespread euxinia. This set the stage for the Earth’s transition from its prokaryote-dominated middle age by removing sulfide, a physiological barrier to eukaryote diversification (Johnston et al., 2010). For the first time in Earth history the complete dominance of oxygenic photosynthesis led to the ventilation of the deep ocean. By ca. 580 Ma bottom waters were oxygenated enough to stimulate the evolution of multicellular benthic animals (Canfield et al., 2007; Narbonne, 2010). With continued input of photosynthetic oxygen, Phanerozoic oxygen levels were achieved by ca. 540 Ma (Holland, 2006). 3. Bioelemental sediments and the record of Earth’s oxygenation The sedimentary and geochemical record of the GOE is preserved primarily in bioelemental sediments, a relatively new classification of sedimentary rocks that encompasses iron formation, chert, and phosphorite (Pufahl, 2010). Because bioelemental sediments are precipitated directly or indirectly by biological processes they are often associated with organic-rich deposits such as black shale, which can be included in the bioelemental spectrum since it contains biologically fixed C. The occurrence of bioelemental sediments through time reflects changes in ocean chemistry linked to climate change, biologic evolution, and tectonic processes (Fig. 4). These factors have influenced the biogeochemical cycling of Fe, Si, P and C (e.g. Logan et al., 1995) and the types of bioelemental sediments produced before, during, and after the GOE. Thus, the temporal distribution of bioelemental sediments provides a framework for understanding the nature of the GOE (Fig. 4) and associated long-term changes to ocean-atmosphere chemistry. Also important are changes in the stacking patterns of bioelemental lithofacies because the redox-sensitive minerals and chemical proxies they contain provide the most detailed information about shifts in water column oxygenation. The best records of seawater oxygenation come from pristine lithofacies. Pristine sedimentary facies are generally fine-grained and accumulate in calm environments. They are characterized by undisturbed water column precipitates and/or in situ authigenic minerals. In a very general sense, the occurrence of bioelemental sediments increased after the onset of the GOE and coincides with a conspicuous rise in the diversity of biologically precipitated minerals; this era of biomediated precipitation produced >2000 new oxide/ hydroxide species (Hazen et al., 2008; Sverjensky and Lee, 2010). As chert occurs in such close affinity with iron formation, phosphorite, and black shale it is discussed in relation to these sediments. 3.1. Iron formation Iron formation is a predominantly Precambrian, Fe-rich, marine chemical sedimentary rock (Figs. 1, 4e6; e.g. Gross (1983); Clout and Simonson, 2005; Klein, 2005; Bekker et al., 2010; Pufahl, 2010). The

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Figure 4. Temporal distribution of iron formation (red), ironstone (purple), phosphorite (yellow) and black shale (black). Based on deposit age, resource estimates and timing of Earth events in Glenn et al. (1994), Kholodov and Butuzova (2004), Condie et al. (2001), Klein (2005), Reddy and Evans (2009), and Bekker et al. (2010). Events: OP ¼ appearance of oxygenic photosynthesis; GOE ¼ Great Oxidation Event; BB ¼ Boring billion; CE ¼ Cambrian Explosion. Glaciations: 1 ¼ Mesoarchean; 2 ¼ Huronian; 3 ¼ Paleoproterozoic; 4 ¼ Neoproterozoic ‘Snow Ball’; 5 ¼ Ordovician; 6 ¼ Permian; 7 ¼ Neogene. Modified from Pufahl (2010).

original definition included a requirement of at least 15 wt. % Fe (James, 1954), but later workers have found this lower limit too restrictive (e.g. Klein, 2005). In weakly metamorphosed iron formation common minerals include the Fe-oxides hematite and magnetite as well as the silicates chert, greenalite, and stilpnomelane (Klein, 2005). Oxygenation of the ocean during the GOE, with either direct or indirect involvement of Fe-oxidizing bacteria, is believed to be responsible for deposition of all large Paleoproterozoic iron formations (Figs. 4e6; e.g. Cloud, 1973; Konhauser et al., 2002). In addition to the importance of iron formation as a recorder of oxygen levels on the early Earth, it is economically significant because it contains most of the world’s iron ore. 3.1.1. Temporal distribution The Archean is characterized by pyrite and magnetite-rich deepwater exhalative iron formation deposited in tectonically active areas around spreading centers associated with volcanic arcs. The dramatic rise in iron formation at ca. 2.8 Ga may correspond to the evolution of oxygenic photosynthesis (Nisbet et al., 2007) and resulting precipitation of ferrous Fe from the Archean ocean (Fig. 4). Although some evidence suggests that iron formation prior to this time was also linked to photosynthetic oxygen (e.g. Hoashi et al., 2009), most data indicate deposition through a combination of anoxygenic photosynthesis, dissimilatory iron reduction, oxygen produced via nonphototrophic sources, and episodic increases in the input of hydrothermal Fe and Si during mantle plume events

(Isley and Abbott, 1999; Konhauser et al., 2002; Pufahl, 2010; Bekker et al., 2010). The iron formation peak at ca. 2.5 Ga is interpreted to signal a shift from deep-water deposition to upwelling-driven, neritic accumulation on the expansive, unrimmed platforms that developed at the end of the Archean (Fig. 7; Cloud, 1973; Klein, 2005; Pufahl, 2010; Bekker et al., 2010). Such aerially extensive Paleoproterozoic iron formation formed in the full spectrum of shelf environments from an oxygen-stratified ocean born during the GOE (Pufahl, 2010). Precipitation occurred when ferrous Fe in upwelled, anoxic waters was either mixed with photosynthetically oxygenated seawater or oxidized during anoxygenic, bacterial photosynthesis (Fig. 7; Cloud, 1973; Klein, 2005; Konhauser et al., 2002; Bekker et al., 2010; Pufahl, 2010). Chert formed abiogenically primarily in subtidal environments where evaporitic concentration (Maliva et al., 2005) and Fe-redox pumping could saturate bottom- and pore water with silica (Fischer and Knoll, 2009; Pufahl, 2010). A suboxic seafloor was a prerequisite for Feredox pumping to saturate sediment with silica. Such conditions are interpreted to have occurred in coastal environments where photosynthetic oxygen oases impinged on the seafloor (Nelson et al., 2010; Pufahl, 2010). Silica was concentrated in pore water during burial when Fe-(oxyhydr)oxides dissolve below the suboxicanoxic redox interface (Fischer and Knoll, 2009; Pufahl, 2010), liberating adsorbed orthosilicic acid (Konhauser et al., 2007). A decline in iron formation through the GOE (Fig. 4) may reflect the increased precipitation of oxidized Fe from seawater as well as

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Figure 5. Iron formation lithofacies. When alteration is considered the mineralogic composition can be used to infer the redox conditions of paleo-seawater/pore water. Hematite ¼ suboxic; magnetite ¼ anoxic (Klein, 2005; Pufahl, 2010). A) Stromatolitic Paleoproterozoic Kona Dolomite, Northern Michigan, U.S.A. Dashed line highlights large stromatolite form. The production of oxygen by such cyanobacteria was responsible for the GOE. B) Laminated magnetite. Neoarchean Eagle Island Group, northwestern Ontario, Canada. C) Metamorphosed, laminated hematite and magnetite. Paleoproterozoic Negaunee Iron Formation, Northern Michigan, USA. D) Granular hematite-chert grainstone. Paleoproterozoic Sokoman Formation, Labrador, Canada. E) Granular iron formation with pebble sized rip-ups of magnetite and hematite mudstone. Paleoproterozoic Sokoman Formation, Labrador, Canada. F) Laminated magnetite and Fe-carbonate with rare magnetite mudstone intraclasts. Paleoproterozoic Sokoman Formation, Labrador, Canada.

a reduction in the delivery of Fe and Si to the ocean. As in the Archean, peaks in iron formation abundance through the Proterozoic have also been correlated to mantle plume activity (Isley and Abbott, 1999; Abbott and Isley, 2001). Deposition of iron formation on continental margins ceased at ca. 1.8 Ga, possibly due to the development of widespread euxinia (Fig. 4; Canfield, 1998; Poulton et al., 2004; Kendall et al., 2011). In a sulfidic water column dissolved sulfide is interpreted to have combined with ferrous Fe to form pyrite, titrating the ocean of dissolved Fe. An important change in the Precambrian Si cycle also occurred at this time and is marked by the end of subtidal chert deposition (Maliva et al., 2005). This change is thought to reflect waning hydrothermal input of Si and a decrease in Si derived from chemical weathering. Sulfidic ocean conditions are interpreted to have continued for nearly a billion years (Anbar and Knoll, 2002). Bioessential trace elements were largely removed from the oceans as sulfides associated with organic matter-rich sediments, which is thought to have contributed to the apparent lull in eukaryote evolution (Anbar and Knoll, 2002). The period that followed these changes is termed the ‘Boring Billion’ because there appears to have been little change in the atmosphereeocean and biological systems over this protracted interval of Earth history (Fig. 4). Iron formation finally reappears coincident with the Neoproterozic ‘snowball’ glaciations

between 740 and 630 Ma (Fig. 4; Klein, 2005; Reddy and Evans, 2009; Bekker et al., 2010). 3.1.2. Deposition and chemistry Unfortunately, there are only a few integrated studies that couple sedimentology, mineralogy and geochemistry of bioelemental deposits bracketing the GOE. Most of those that do focus on the disposition and chemistry of suboxic and anoxic lithofacies forming the large continental margin iron formations of the Paleoproterozoic (e.g. Beukes and Klein, 1990; Klein and Ladeira, 2000; Pickard et al., 2004; Pufahl and Fralick, 2004; Klein, 2005; Fralick and Pufahl, 2006; Fischer and Knoll, 2009; Pecoits et al., 2009). This is because the presence of a prominent oxygen chemocline is the primary control on facies mineralogy (Fig. 7; Pufahl, 2010). Deposition of suboxic lithofacies occurred along segments of the coastline where photosynthetic cyanobacteria produced oxygen (Figs. 5A and 7). These deposits are characterized by hematite (Fe2O3) and chert (SiO2; Fig. 8; Klein, 2005). Anoxic lithofacies are distinguished by the presence of magnetite (Fe3O4), greenalite ((Fe2þ, Fe3þ)2-3Si2O5(OH)4), or stilpnomelane (K(Fe2þ,Mg,Fe3þ)8 (Si,Al)12(O,OH)27$nH2O; Fig. 8; Klein, 2005). All of these minerals contain some reduced Fe, and reflect precipitation under extremely low oxygen concentrations (ca. 1020 pO2-water and were likely as low as 1070 pO2-water; Mel’nik, 1982).

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Figure 6. Iron formation lithofacies from the Neoarchean-Paleoproterozoic Hamersley Basin, Western Australia. As in Figure 5, mineralogy can reflect the redox conditions of paleoseawater/pore water. Arrows denote younging direction. A) Laminated magnetite and chert. Late Neoarchean Marra Mamba Iron Formation, Western Australia. B) Laminated magnetite and chert. Early Paleoproterozoic Joffre Iron Formation, Western Australia. C) Interlaminated magnetite, chert and fine-grained, hematitic grainstone laminae. Early Paleoproterozoic Joffre Iron Formation, Western Australia. D) Laminated magnetite and chert intercalated with thin beds of hematitic grainstone. Early Paleoproterozoic Joffre Iron Formation, Western Australia. Grainstones are interpreted as event deposits that carried granular sediment downslope from higher energy environments that were above the oxygen chemocline.

Indirect chemical proxies such as the Fe isotopic (d56Fe) and REE composition of iron formation have also been used to infer oxygenation history (Fig. 2C and 7; Table 1; e.g. Beukes and Klein, 1990; Klein, 2005; Johnson et al., 2008). The REE systematics of redox sensitive facies, however, seems more robust and easier to interpret, primarily because it is a direct measure of seawater Eh (Elderfield and Greaves, 1982) without the issues of potentially strong and not yet understood biologic fractionations (Johnson et al., 2008). In general, iron formation and chert facies formed in oxygenated marine environments have negative Ce anomalies and are enriched in heavy REE’s (HoeLu) when compared to shales (Klein, 2005). This is because although the oxidation of Ce3þ greatly reduces Ce solubility, oxidative scavenging on the surface of freshly precipitated Fe-(oxyhydr)oxides removes Ce from seawater (Ohta and Kawabe, 2001). Since the overall concentration of Ce is low, seawater is left depleted in Ce producing a negative Ce anomaly (e.g. Elderfield and Greaves, 1982; Piper et al., 1988). The enrichment of heavy REE’s (Byrne and Sholkovitz, 1996) is also interpreted to be the result of preferential oxidative removal of the other light REE’s (LaeDy) from seawater. These processes will only produce a “seawater pattern” if deposition occurs away from a terrigenous clastic source since siliciclastic material has no Ce anomaly or heavy REE enrichment (Watkins et al., 1995). Recently, the Ni concentration in iron formation has been used to infer both the timing and cause of the GOE (Fig. 2E; Table 1; Konhauser et al., 2009). A significant decrease in the Ni/Fe ratio at ca. 2.7 Ga is interpreted to correlate to a major drop in the concentration of Ni in seawater. Because of their insatiable appetite for Ni, this change likely limited methanogens in the Neoarchean and led to a concomitant reduction in the generation of atmospheric methane. With decreasing methane and the other environmental changes that occurred at the end of the Archean (Des Marais et al., 1992; Godderis and Veizer, 2000; Catling et al., 2001; Kump et al., 2001; Holland, 2002; Papineau et al., 2009) the stage was apparently set for the accumulation of cyanobacterial oxygen and the GOE.

3.2. Phosphorite Phosphorite is a bioelemental sedimentary rock rich in P, is often associated with coastal upwelling, and occurs almost exclusively in the Phanerozoic (Figs. 1, 4 and 9). It is defined as a rock with greater than 18 wt. % P2O5, but P2O5 can be as great as 40 wt. %, making these rocks an important fertilizer ore (Pufahl, 2010). Most published accounts of Proterozoic and Neoproterozoic phosphorites do not describe true phosphorite, but phosphatic deposits that contain much less than 18 wt. % P2O5. This distinction is important because uncritical reporting of phosphatic occurrences has resulted in an over estimation of Precambrian phosphorite, which has led to errors in assessing temporal abundance and understanding the Precambrian P cycle (e.g. Papineau, 2010). Phosphorite forms through phosphogenesis, the authigenic precipitation of francolite within sediment just beneath the seafloor (Glenn et al., 1994). Francolite is a highly substituted carbonate fluorapatite (Ca10-a-bNaaMgb(PO4)6-x(CO3)x-y-z(CO3$F)x-y-z(SO4)zF2). Its precipitation is microbially mediated and also controlled by the redox potential of bottom- and pore water (Jahnke et al., 1983; Glenn et al., 1994). Authigenic, biological, and hydrodynamic processes work together to form phosphatic laminae, in situ nodules or reworked granular beds (Föllmi et al., 1991; Föllmi, 1996). Phosphorite is the most important long-term sink in the global phosphorus cycle. In the Phanerozoic the majority of P in the oceans is sequestered in marine sediment on continental margins and beneath regions of active coastal upwelling (Filippelli, 2008; Fig. 10). Phosphorus is removed from nutrient-rich surface waters by phytoplankton and authigenically converted to francolite in accumulating organic-rich sediment through a series of microbially mediated redox reactions (Jahnke et al., 1983; Glenn et al., 1994). Bacterial sulfate reduction is the most efficient of these reactions at liberating organically bound P to pore water (Arning et al., 2009). Precipitation of francolite occurs when pore water becomes supersaturated with respect to calcium-phosphate (Glenn et al., 1994). Such phosphorite co-occurs with biogenic chert and black

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Figure 7. Continental margin iron formation. Lithofacies formed a sedimentary wedge that fines and thickens basinward. Coastal upwelling provided a sustained supply of anoxic bottom water rich in dissolved Fe and Si. Precipitation occurred in an oxygen stratified water column that was suboxic down to fair-weather wave base. Nearshore lithofacies consist of cross-stratified grainstones that are commonly stromatolitic. Laminated pristine lithofacies accumulated in low energy environments such as shallow lagoons and below fair-weather wave base on the middle and distal shelf. REE spidergrams show the behaviour of Ce across the shelf. A negative Ce anomaly is most pronounced along segments of the paleoshoreline that were oxygenated by photosynthesis. It disappears offshore where bottom and intermediate waters were anoxic. The positive Eu anomaly reflects the hydrothermal source of Fe (Klein, 2005 and references therein). SWB ¼ storm wave base; FWB ¼ fair-weather wave base. Modified from Pufahl (2010).

shale forming an upwelling triad of sediments. In areas not associated with prominent upwelling the concentration of phosphate in sediment is regulated by Fe-redox pumping (Fig. 11; Heggie et al., 1990). Preferential adsorption and release of phosphate on Fe-(oxyhydr)oxide is kinetically favoured in Phanerozoic seawater because it is severely silica-undersaturated (Konhauser et al., 2007).

3.2.1. Temporal distribution Phosphorite did not form in the Archean (Fig. 4), likely reflecting weathering of phosphate-poor, mafic crust under an anoxic atmosphere (Pufahl, 2010). The appearance of phosphorite in the Paleoproterozoic coincides with the GOE and the onset of oxic chemical weathering of the continents (Papineau, 2010; Pufahl, 2010). This relatively minor phosphatic episode was not associated with upwelling and unlike Phanerozoic phosphorites, restricted to shallow-water environments (Nelson et al., 2010). It occurred between 2.2 and 1.8 Ga, beginning just after the Huronian Glaciation and in the middle of the GOE (Papineau, 2010; Pufahl, 2010). This episode probably records an abrupt increase in the delivery of phosphate to the oceans. Increased phosphate likely fueled a corresponding increase in primary production that enhanced photosynthesis and the contribution of oxygen to the GOE (Papineau, 2010). This pulse may be the consequence of a switch to post-glacial continental chemical weathering under an oxygenated atmosphere from a long period dominated by mechanical weathering during the Huronian Glaciation (Papineau, 2010; Pufahl, 2010). Thus, the appearance of Paleoproterozoic phosphorite is directly linked to the GOE and the oxygenation of the oceans (Nelson et al., 2010; Pufahl, 2010). Phosphogenesis during this initial episode was restricted to segments of the shoreline that were silica undersaturated and oxygenated through microbial photosynthesis. These conditions permitted a combination of bacterial sulfate reduction and Fe-redox pumping to concentrate P in coastal sediment (Fig. 11; Nelson et al., 2010). Such shallow-water phosphorite is in stark contrast to upwelling-related, Phanerozoic phosphorites that accumulated in a range of shelf environments. This difference likely reflects the dissimilarity in the oxygenation state of the seafloor (Nelson et al., 2010). The anoxia that typified Precambrian intermediate and bottom water prevented Fe-redox pumping from operating in deeper settings (Fig. 11). During the onset of sulfidic ocean conditions the Fe and P cycles became decoupled, which led to the disappearance of phosphorite

Figure 8. Paragenesis typical of pristine iron formation in suboxic and anoxic paleoenvironments. Modified from Klein (2005) and Pufahl (2010).

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Figure 9. Precambrian phosphorite lithofacies. Most Precambrian phosphorites are unlike Phanerozoic phosphatic deposits because they do not form aerially extensive deposits. They generally consist of thin pristine phosphorite in peritidal environments and granular phosphatic lags in shallow-water lithofacies. A) Laminated pristine phosphorite. Subhedral crystals are pyrite, black blebs are organic matter and the honey-brown mineral between organic-rich laminae is francolite. B) Francolite peloids (brown) with greenalite cement (acicular crystals) surrounded by dolomite. Opaque square-shape is pyrite. Paleoproterozoic Ruth Formation, Labrador, Canada. Authigenic glauconite is commonly associated with such francolite grains indicating phosphogenesis along a suboxic seafloor (Pufahl, 2010). C) Phosphatic peloids on bedding surfaces (arrows) in cross-laminated chert. Paleoproterozoic Bijiki Iron Formation, northern Michigan, U.S.A. D) Francolite peloid (brown) cemented with ankerite. Paleoproterozoic Bijiki Iron Formation, northern Michigan, U.S.A.

at ca. 1.8 Ga (Fig. 4; Pufahl, 2010). The precipitation of pyrite in a euxinic water column decreased the potential for Fe-redox pumping, even in nearshore oxygen oases (Nelson et al., 2010). Widespread sulfidic conditions likely made bacterial sulfate reduction ineffective as a driver of phosphogenesis because phosphate would have been released to the water column where it could be efficiently recycled and not fixed as francolite in the sediment (Nelson et al., 2010). Phosphorite, like iron formation, was not deposited again until the Neoproterozoic (Fig. 4). 3.2.2. Deposition and chemistry Although rare, Proterozoic phosphorites hold great promise for refining what is known about changes in ocean redox structure (Melezhik et al., 2005; Pufahl, 2010), especially when coupled with the sedimentology and chemistry of co-occurring bioelemental sediments. Francolite readily incorporates a variety of redox sensitive trace elements into its crystal structure and thus, often preserves a record of pore water and bottom water Eh during deposition (Fig. 10; Jarvis et al., 1994). Trace elements generally replace Ca2þ in francolite, but can also be transferred to the sediment by absorption onto crystal surfaces, scavenging by organic matter, or substitution in sulfides (Jarvis et al., 1994). Enriched elements include Ag, Cu, Cr, V, Cd, Mo, Se, U, Y, and Zn, and the REEs La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, and Lu (e.g McArthur and Walsh 1984; Altschuler, 1980; Hiatt and Budd, 2003; Fig. 10). The most commonly used to infer redox conditions are Cu, Cr, V, Cd, Mo, U, and Zn (Fig. 10). All but U are mobilized and incorporated under reducing conditions. As in iron formation, the REE content of francolite records seawater values although it continues to absorb REE from pore water below the sedimentewater interface (Altschuler, 1980; Piper et al., 1988). As in iron formation, the presence of a prominent negative Ce anomaly indicates precipitation in oxygenated environments (Piper et al., 1988; Jarvis et al., 1994). In addition to trace elements, the stable isotopic composition of francolite can be used to understand the microbial processes

releasing phosphate to pore water (d13CCO3 ) and to determine precipitation temperature (d18OCO3 ; d18OPO4 ; Piper and Kolodny, 1987; Shemesh et al., 1988; Hiatt and Budd, 2001). Temperature calculations are sometimes coupled with trace element analysis to infer the redox conditions and paleooceanography of ancient seas (e.g. Hiatt and Budd, 2003).

Figure 10. Continental margin phosphorite and black shale. Phosphorite accumulates within organic-rich sediment beneath the sites of coastal upwelling. A pronounced oxygen minimum zone (OMZ) develops as benthic bacteria exhaust oxygen to degrade organic matter. Black shale is also associated with upwelling, but can form in calm, nutrient-rich coastal environments such as lagoons. The plots show redox-related changes in trace element concentrations across the shelf. In the nearshore a negative U anomaly and elevated Cr records accumulation under oxic and suboxic conditions. Elevated U, V, Cu, Cd, Zn, Mo, and Ni reflects deposition in deeper anoxic portions of the shelf. SWB ¼ storm wave base; FWB ¼ fair-weather wave base. Modified from Pufahl (2010).

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Figure 11. Extent of phosphogenesis resulting from Fe-redox pumping on Precambrian and Phanerozoic shelves. As Fe-(oxyhydr)oxides are buried beneath the Fe-redox boundary  they dissolve, liberating sorbed HPO24 to pore water. Francolite precipitation is limited in the sediment by the availability of seawater-derived F . Although important in stimulating phosphogenesis in the Phanerozoic, bacterial sulfate reduction was likely much less efficient at promoting the precipitation of francolite in the Precambrian because of the very low seawater sulfate levels. Thus, the difference in the size of phosphogenic regions in the Precambrian and Phanerozoic is interpreted to the consequence of the disparity in the oxygenation state of the seafloor. In the Precambrian, photosynthetically oxygenated nearshore environments possessed suboxic seafloors that facilitated Fe-redox pumping and phosphogenesis. Phosphogenesis could not occur in the middle and distal shelf because these regions were below the oxygen chemocline. Phosphogenesis in the Phanerozoic occurs across the entire shelf because the seafloor is generally well oxygenated. SWB ¼ storm wave base; FWB ¼ fair-weather wave base. Modified from Nelson et al. (2010).

3.3. Black shale Black shale is a dark, thinly laminated, carbonaceous finegrained clastic sedimentary rock (Fig. 12) that is rich in organic matter (>2 wt. %), sulfides (especially pyrite), and redox sensitive trace elements (U, V, Cu, and Ni; Arthur and Sageman, 1994; Piper and Calvert, 2009). It can form in a wide range of paleoenvironments from peritidal to deep basin settings, is often associated with phosphorite, and can be a hydrocarbon source rock (Fig. 10). Black shales are commonly interpreted as recording deposition beneath a highly productive surface ocean or within anoxic, sulfidic bottom waters, or a combination of both (Piper and Calvert, 2009). Recent work, however, suggests that high planktic productivity is the most important control on organic matter enrichment in marine sediment (e.g. Piper and Calvert, 2009 and references

therein). Organic matter accumulates because the rate of production and settling is greater than the rate of degradation of organic carbon on the seafloor (Pedersen and Calvert, 1990). Since processes of black shale deposition can occur across the spectrum of shelf environments, their occurrence is not always an indication of accumulation in a deep, open ocean basin. Processes leading to the formation of black shale are important because they link the various pools of carbon in the oceanatmosphere system (Arthur and Sageman, 1994). These processes govern carbon burial, which regulates climate, and oxygen levels by controlling the rate reduced C is sequestered in the geologic record (Holland, 2002; Canfield, 2005). Since P is the primary control on productivity over geologic time scales the phosphorus cycle ultimately determines the rate of organic matter burial and removal of carbon dioxide from the atmosphere.

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Figure 12. Black shale. A) Pristine phosphorite associated with black shale of the Permian Meade Peak Member (M), which is overlain by the Rex Chert Member. Permian Phosphoria Formation, Wyoming, U.S.A. From Pufahl (2010). B and C) Black shale from the Marra Mamba Iron Formation, Western Australia. Drill core WRL-1. Yellowish staining in (B) is from weathered sedimentary sulfides. Organic-rich laminae in (C) are commonly scoured by very fine-grained, thinly bedded sandstone layers. Dashed line highlights a scour surface. Arrows denote younging direction. D) Black shale from Joffre Iron Formation. Drill core SPD-50. Minute light-coloured specks within certain laminae are pyrite crystals.

3.3.1. Temporal distribution The temporal distribution of black shale is even less well constrained than that of phosphorite (Figs. 1 and 4), primarily because it forms components of other depositional systems. The record of Precambrian black shale is also severely biased given the rarity of preserved deep-sea sediments, and because they are easily eroded. In general, however, the timing of black shale deposition reflects periods when oxygen concentrations could increase in the atmosphereeocean system (Fig. 4; e.g. Berner, 2004). Secular changes in black shale deposition result from changes in carbon cycling in “active” surface ocean pools, in the atmosphere, on land, and in marine sediment, and carbon pools that cycle on much longer timescales (Burdige, 2006). Such changes are partly linked to the GOE (Des Marais et al., 1992), which apparently follows an episode of enhanced carbon burial in the late Archean (Fig. 4). This is the first of three noticeable peaks in black shale deposition during the Precambrian (Fig. 4; Condie et al., 2001). It is the least prominent and occurs in the Neoarchean between ca. 2.7 and 2.5 Ga. This initial pulse of black shale accumulation is thought to correspond to either a mantle plume event, which through climate warming, increased chemical weathering and nutrient delivery to the oceans (Condie, 2004), or a change in ocean currents (Condie et al., 2001) that resulted in initiation of upwelling along favorably positioned cratons. The sequestration of reducing organic matter during this episode is interpreted to have contributed to the GOE (Des Marais et al., 1992). The second pulse is a prominent event occurring between ca. 2.0 and 1.7 Ga (Fig. 4; Condie et al., 2001), just after the Huronian Glaciation. As with iron formation of this age, the accumulation of black shale is also correlated to mantle volcanism (Condie et al., 2001). Intense chemical weathering of post-glacial landscapes (Papineau, 2010; Pufahl, 2010) is interpreted to have increased P fluxes to the ocean that not only stimulated primary production and phosphogenesis (Nelson et al., 2010), but also black shale deposition as well. Black shale again becomes conspicuous in the Cryogenian between ca. 800e600 Ma (Fig. 4). Organic-rich mudstones, some of

which are phosphatic, accumulated during retreat of the “snowball” glaciations (Condie et al., 2001; Le Heron et al., 2009). Elevated surface ocean productivities were likely sustained by delivery of nutrients through glacial runoff and invigorated coastal upwelling (Papineau, 2010). The pronounced equator-to-pole temperature gradient that develops during glaciations leads to more energetic atmospheric circulation and thus, coastal upwelling, resulting in the widespread accumulation of organic-rich sediment (Vincent and Berger, 1985). Correspondence between peaks of black shale and those of iron formation deposition in the Precambrian (Fig. 4) highlights the importance that photosynthetic oxygen played in the accumulation of iron formation. This relationship also emphasizes the connection between ocean circulation and upwelling to deliver reduced iron and P to the photic zone. Like phosphorite, pulses of black shale deposition in the Phanerozoic are linked to ocean-climate feedback (Bluth and Kump, 1991; Arthur and Sageman, 1994). Prominent peaks are also the consequence of enhanced P burial from invigorated coastal upwelling or increased chemical weathering and delivery of phosphate to the oceans (Fig. 4; Glenn et al., 1994; Föllmi, 1996). 3.3.2. Deposition and chemistry Much information about fluctuations in seawater Eh is derived from paragenetic studies of black shale-hosted, authigenic minerals (Glenn and Arthur, 1988; Arthur and Sageman, 1994; Pufahl and Grimm, 2003; Raiswell et al., 2011). Textural relationships between glauconite, pyrite, francolite, and carbonate provide a high fidelity record of the physical, chemical, and biologic processes causing subtle shifts in redox potential (Glenn and Arthur, 1988; Pufahl and Grimm, 2003). These minerals precipitate through a series of microbially mediated redox reactions (Froelich et al., 1979; Glenn et al., 1994). In order of decreasing energy yield these reactions include oxic respiration, denitrification, transition metal oxide reduction, sulfate reduction, and methanogenesis. Geochemical evidence suggests that all but oxic respiration evolved by the late Archean (Garvin et al., 2009; Lyons and Gill, 2010), and aerobic

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heterotrophs evolved by ca. 2.1 Ga in response to increasing oxygen levels (Papineau et al., 2005). Because of widespread ocean anoxia bacteriaedriven reactions that produce and consume organic matter were not confined to below the seafloor, but also occurred within the water column. Pyrite precipitates below the sulfate redox interface through the conversion of monosulfides formed during bacterial sulfate reduction (Schieber, 2002). Sedimentary pyrite is often framboidal and finely disseminated (Raiswell, 1982; Wilkin and Arthur, 2001; Schieber, 2002), but discrete layers have been interpreted as recording precipitation and suspension settling through a euxinic water column (Poulton et al., 2004). Francolite precipitates in association with the microbial reduction of nitrate, Mn-oxides, Fe-oxides, and sulfate (Pufahl, 2010). Unlike the formation of pyrite, however, phosphogenesis is not a redox-controlled reaction, but is regulated only by the concentration of phosphate in pore water (Glenn et al., 1994). In addition to pyrite and francolite, other authigenic minerals that may precipitate in black shales include glauconite ((K,Na,Ca)1.2-2.0(Fe3þ,Al,Fe2þ,Mg)4.0[Si7-7.6Al1.0-0.4O20](OH)4$n(H2O)), calcite (CaCO3), dolomite (CaMg(CO3)2, and siderite (FeCO3), all of which can be used to further constrain pore water Eh. Glauconite forms first, within suboxic pore water at the Fe-redox interface, followed by pyrite, and then carbonate at progressively deeper levels in the sediment. Glauconite occurs as authigenic peloids, coatings, or cement. Carbonate minerals precipitate within the alkalinity maximum that develops during intense microbial respiration. The type of carbonate mineral produced depends on the availability of Ca2þ, Fe2þ and Mg2þ. Calcite forms from pore water with little Fe2þ and Mg2þ, whereas siderite precipitates in anoxic pore water with abundant Fe2þ (François and Gérard, 1986; Klein, 2005). Dolomite is created in pore water enriched in Ca2þ and Mg2þ when the bacterial reduction of SO2 4 , a kinetic inhibitor to dolomite precipitation, is converted to H2S (Baker and Kastner, 1981; Kastner, 1984; Wright and Wacey, 2005). These carbonate minerals are generally a microcrystalline cement that binds detrital grains and earlier authigenic phases together, but can also form displacive concretionary horizons (Kholodov and Butuzova, 2004). The bulk trace element content of black shales and the sulfur isotopic composition of co-occurring pyrite provide evidence of changing seawater Eh over longer timescales. An increase in the concentration of redox-sensitive trace elements and an increase in the fractionation of d34S roughly correspond to the onset of the GOE (Fig. 2A; Table 1). The restricted range of d34S values in pyrite prior to ca. 2.5 Ga is interpreted to reflect low seawater SO2 4 concentrations of the Archean; the consequence of negligible SO2 4 delivery from levels anoxic chemical weathering (Canfield, 2001). Low SO2 4 restrict bacterial sulfate reduction and produce little variation in d34S values (Canfield, 2001). After ca. 2.5 Ga fractionations increase dramatically to values expected for bacterial sulfate reduction, which is not limited by low sulfate concentrations. Higher sulfate levels were produced by weathering of pyrite under an oxygenated atmosphere (Canfield, 2001). MIF of sulfur isotopes (D33S; Figs. 1 and 2B; Table 1) in pyrite provides the best evidence of Precambrian atmospheric oxygen levels and the timing of the GOE (Farquhar et al., 2000; Farquhar and Wing, 2003). The nature of its onset is preserved in the record of multiple sulfur isotope distributions (d34S-D33S), which suggests that oxygen levels began to fluctuate ca. 150 million years prior to the permanent rise at ca. 2.4 Ga (Partridge et al., 2008; cf. Wille et al., 2007). The decreased variability and appearance of positive pyrite d56Fe values after ca. 2.3 Ga corroborate D33S data (Fig, 2; Rouxel et al., 2005), but it is unclear whether these changes reflect seawater composition or diagenesis (Johnson et al., 2008). Molybdenum isotopes from FeeMoeS precipitates in black shale

provide further clues (Lyons et al., 2009; Severmann and Anbar, 2009; Voegelin et al., 2010). d98/95Mo values corroborate the rise in oxygen levels ca. 150 million years before the accepted onset of the GOE (Wille et al., 2007; Voegelin et al., 2010). Molybdenum isotope data also suggest that although euxinic conditions may have eventually developed in the late Paleoproterozoic, the ocean was probably not a “global Black Sea” (Lyons et al., 2009). Another approach that is increasingly being used is the analysis of carbonate-associated-sulfur (CAS; e.g. Guo et al., 2009). Because CAS can acquire the isotopic composition of pore water and seawater (Burdett et al., 1989) it is particularly attractive as a paleoredox proxy in Precambrian limestones. Sulfur isotope data from associated pyrite also allows potential calculation of the offset between SO2 4 and H2S during bacterial sulfate reduction (Lyons and Gill, 2010), further constraining the nature of redox sensitive microbial processes in the sediment and water column. 4. Reading the record of Earth’s oxygenation: diagenetic and metamorphic effects Diagenesis and metamorphism can significantly alter sediment chemistry, especially in deposits as old as the Precambrian (e.g. Hayes et al., 1983; Ayalon and Longstaffe, 1988; Crusius and Thomson, 2000; Shields and Stille, 2001; Petsch et al., 2005; Gonzalez Alvarez and Kerrich, 2010; Hiatt et al., 2010). Thus, it is difficult to reconcile why so few studies of Precambrian depositional systems attempt to understand alteration of what are primarily metamorphic rocks and minerals before interpreting geochemical data. This is especially true since techniques exist to assess whether observed anomalies reflect conditions at the time of deposition, alteration, or a combination of both (e.g. Kendall et al., 2009). Unfortunately, technological breakthroughs that have allowed the rapid analysis of samples (Watson, 2008) also lead to the brisk publication of data without full assessment of potential alteration. 4.1. Sedimentology, basin evolution, and alteration The current level of understanding bioelemental sediment alteration is at about the same stage as understanding limestone diagenesis was 30 years ago. A basic tenet that has important implications for understanding geochemical proxies is that the most desirable deposits for geochemical analysis are pristine lithofacies. Like lime mudstones, fine-grained bioelemental facies usually represent seawater and authigenic precipitates with low hydraulic conductivities that tend to fix pore water chemistry close to the time of deposition (e.g. Kyser et al., 1998). This fact is commonly overlooked when extrapolating the chemistry of authigenic minerals to the overlying water column. Coarser facies have higher fluid/rock ratios and experience greater fluid fluxes during burial that often results in chemical compositions that are significantly different from original ones (e.g. Reeckman, 1981). Sedimentologic and basin evolution context are both paramount when interpreting geochemical data. A properly constrained depositional and post-depositional framework provides information on how oceanography, depositional environment, seawater and pore water chemistry, microbial biology, and alteration influence the chemical composition of bioelemental sediments. Without this perspective it is a challenge to interpret whether geochemical anomalies through the GOE are the consequence of local environmental factors or global in character (Lyons et al., 2009; Pufahl et al., 2010). In most cases ascertaining the nature of anomalies in Precambrian sedimentary rocks is especially difficult given the rarity of preserved deep-sea sediments (Pufahl et al., 2010). Unfortunately, this issue is often overlooked resulting in conclusions that are not fully supported by sedimentologic data. For example, key

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stratigraphic units in the Paleoproterozoic Pretoria Group, where much of the geochemical data is derived (e.g. Bekker et al., 2004; Bau and Alexander, 2006), are interpreted as epeiric sea deposits (Eriksson and Reczko, 1998; Eriksson et al., 2009). Sedimentologic evidence suggests that epeiric seas are typified by restricted circulation patterns that produce water masses with compositions that differ substantially from the open ocean (e.g. Hiatt and Budd, 2001; Piper, 2001; Algeo and Heckel, 2008). In all sedimentary basins diagenetic hydrostratigraphy is controlled primarily by lithofacies and resulting diagenetic reactions, as well as later cross-formational faulting (Hiatt and Budd, 2003; Hiatt et al., 2007; Holk et al., 2003; Hiatt and Kyser, 2007). Depositional environments determine internal hydrologic properties on a basin-scale because they control the composition, fabric and grain size of lithofacies (e.g. Hiatt and Budd, 2003). Widespread lateral flow of diagenetic fluids can occur over distances of hundreds of kilometers and at temperatures >200  C (Hiatt and Kyser, 2007; Kyser, 2007; Alexandre et al., 2009; Hiatt et al., 2010). Diagenetic fluids flow most intensively within coarse-grained facies that have not experienced intense cementation and are situated above unconformities or along internal discordances such as parasequence boundaries (Hiatt et al., 2003; Hiatt et al., 2007; Hiatt and Kyser, 2007). Stratigraphic boundaries often allow preferential diagenetic and metamorphic fluid flow that can reset relatively robust geochemical proxies, such as d34S (Pufahl et al., 2010). This long-term diagenetic hydrostratigraphy can involve multiple diagenetic events and can persist into deep burial settings (>5 km) over long periods of basin evolution, which in the case of Proterozoic basins can extend over 500 Ma (Holk et al., 2003; Hiatt et al., 2007, 2010; Kyser, 2007; Alexandre et al., 2009; Hiatt et al., 2010). Such heterogeneous alteration does not occur gradually through time, but during discrete episodes of pronounced diagenesis and metamorphism. The paragenesis of diagenetic and metamorphic mineral assemblages in Proterozoic sedimentary basin-hosted uranium and PbeZn deposits demonstrate that periods of elevated fluid flow and concomitant alteration are driven by tectonic events that changed basin hydrology (Kotzer et al., 1992; Polito et al., 2004, 2011; Alexandre and Kyser, 2005; Kyser, 2007; Alexandre et al., 2009; Hiatt et al., 2010; Polito et al., 2011). In these systems the punctuated recrystallization of iron oxides (Kotzer et al., 1992) and uraninite (Polito et al., 2004, 2011; Alexandre and Kyser, 2005; Polito et al., 2011) as well as the precipitation of diagenetic illite (Polito et al., 2004; Alexandre et al., 2009; Hiatt et al., 2010; and Polito et al., 2011) indicate fluid/rock ratios increased during regionalscale tectonic events, which created the hydraulic gradients necessary for fluid flow. All successions used to interpret the nature of the GOE have been subjected to these conditions. Thus, careful assessment of alteration using petrographic techniques should complement geochemical analyses to fully evaluate whether sedimentary successions contain chemical proxies that reflect paleoenvironment. 4.2. Mineralogy, chemistry, and alteration Diagenetic and metamorphic changes to bioelemental sediments are not only critical to fully understanding and interpreting the sedimentary record of the GOE, but also other important events in ocean-atmosphere evolution. Metamorphic mineral transformations in iron formation, primarily because of the detailed thermodynamic and paragenetic work of Klein (e.g. Klein, 2005; Figs. 5, 6 and 12), are the best understood aspect of bioelemental sediment alteration. During burial, authigenic greenalite and stilpnomelane change to minnesotaite ((Fe2þ,Mg)3Si4O10(OH)2; Fig. 8; Klein, 2005), a common alteration mineral. With increasing metamorphic grade amphiboles, pyroxenes, and fayalite are high-temperature reaction products. These relationships can be used to infer the original mineralogies of

15

iron formation, thus providing information regarding the paleoredox structure of the Precambrian ocean (Pufahl, 2010). Trace element concentration data can then be interpreted in terms of mineralogical changes, but little is known about the potential fractionation of isotopes within systems that are increasingly employed to interpret oceanographic conditions associated with the GOE, such as d56Fe d53Cr, d97/95Mo, and d98/95Mo. There are potentially significant fractionations of these isotopes between phases such as greenalite, stilpnomelane, and minnesotaite. Because REE ratios in iron formation are usually not significantly modified by alteration (Derry and Jacobsen, 1990) concentration patterns of REE’s are potentially useful trace element proxies. Coupling the stratigraphic correlation of metamorphosed Fe- and Si-rich facies to their REE composition further constrains ocean redox conditions (e.g. Derry and Jacobsen, 1990), and REE patterns could provide a potential method to evaluate other geochemical proxies. The trace element composition of phosphorite is more susceptible to diagenesis and metamorphism because elements within the “open” crystal structure of francolite can be mobilized and can fractionate (Bonnot-Courtois and Flicoteaux, 1989). Oxygen and carbon isotopes must also be used with caution. Isotopic exchange can affect d18O values from both the carbonate (d18OCO3 ) and phosphate (d18OPO4 ) sites, although the d18OCO3 francolite is more vulnerable to exchange with surrounding pore waters (e.g. McArthur and Herczeg, 1990). The result of post-burial exchange of carbon isotopes is best observed when d18OCO3 and d13CCO3 values are plotted against each other. As in altered limestones, such a plot is constrained at one end by seawater and the other by diagenetic francolite compositions (Jarvis et al., 1994). Because of the relative ease with which isotopic exchange occurs in francolite, caution should be exercised, especially when interpreting the stable isotopic composition of Precambrian phosphorite. The effects of diagenesis and metamorphism on the d56Fe, 98/95 d Mo, d34S, D33S composition of authigenic phases in bioelemental sediments are generally unknown. This is especially true for CAS, where sulfur does not sit within structural sites of carbonate minerals (Morse and Mackenzie, 1990; Marenco et al., 2008). Further work is also required to understand the full range of processes controlling isotopic fractionations prior to burial. Thus, much research is required before the true nature of geochemical anomalies (both spatial and stratigraphic) through the GOE can be fully assessed. 5. Integrated approach and future research Although the number of studies that combine sedimentology and geochemistry to understand the GOE and Earth’s subsequent oxygenation has increased in recent years (e.g. Beukes and Klein, 1990; Klein and Ladeira, 2000; Pickard et al., 2004; Klein, 2005; Fralick and Pufahl, 2006; Schröder and Grotzinger, 2007; Schröder et al., 2008; Fischer and Knoll, 2009; Pecoits et al., 2009; Poulton et al., 2010; Pufahl et al., 2010), most are geochemical investigations (e.g. Beaumont and Robert, 1999; Farquhar et al., 2000; Canfield et al., 2000, 2008; Catling et al., 2001; Shen et al., 2002; Yang et al., 2002; Bekker et al., 2003; Huston and Logan, 2004; Aharon, 2005; Brocks et al., 2005; Rouxel et al., 2005; Siebert et al., 2005; Johnston et al., 2006, 2009; Bottrell and Newton, 2006; Bau and Alexander, 2006; Frei et al., 2009; Guo et al., 2009; Johnson et al., 2008; Kendall et al., 2009; Konhauser et al., 2009; Lyons and Reinhard, 2009; Lyons et al., 2009; Planavsky et al., 2009, 2010; Severmann and Anbar, 2009; Lyons and Gill, 2010; Papineau, 2010; Voegelin et al., 2010; Basta et al., 2011). None use a fully integrated approach incorporating sedimentology, stratigraphy, alteration, and basin analysis to constrain the depositional and post-depositional context of bioelemental sediments

16

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(Fig. 13). Such a method mitigates the potential for incorrectly interpreting geochemical data because it not only provides information on how oceanography and depositional environment influenced sediment chemistry, but also the effects of seawater and pore water, microbial biology, and alteration. Central to this approach is the description of outcrop and drill core to understand lithofacies associations and stratal architecture. This allows the construction of a sequence stratigraphic framework to understand the evolution of paleoenvironments and ocean current systems through time (Catuneanu et al., 2009). Although documenting the sequence record in the Precambrian is difficult because of poor preservation, especially of deeper water lithofacies, and structural deformation (Miall, 2005), it can be done (Nelson et al., 2010). As in the Phanerozoic, attention must be given to the identification of lithofacies stacking patterns and breaks in sedimentation since each genetic unit, or systems tract, is defined by specific correlation of vertical and lateral facies trends and bounding surfaces (Catuneanu et al., 2009). Sampling for petrography and geochemistry should be lithofacies specific and interpreted in a sequence stratigraphic framework. Doing so permits the interpretation of petrographic and geochemical data in paleoenvironmental context and provides the backdrop for understanding the effects of post-depositional fluid flow and alteration on sediment composition (e.g. Kyser, 2007; Hiatt et al., 2010). Any geochemical analyses should be superseded

by petrographic investigations aimed at understanding mineral paragenesis (Fig. 13). Clarification of primary and secondary textures dictate what samples should be analyzed for their chemical composition. Once these relationships are understood, anomalies in high-resolution geochemical data across individual lithofacies can be properly assessed, elucidating any connection to alteration and if primary, whether they are of local or regional extent. 6. Conclusions The GOE marks the beginning of the most significant change in Earth history, setting the stage for wholesale changes in ocean chemistry and the evolution of multicellular life. It is the utmost expression of co-evolution between the geosphere and biosphere. The geosphere provided the chemical building blocks and ecological niches for early life, and the biosphere provided oxygen, which changed the nature of weathering, nutrient cycling, mobility of redox sensitive elements such as iron and uranium, and in turn provided environmental stresses that pushed life along new evolutionary pathways. Early understanding of the GOE was based on temporal trends in bioelemental sediments, changes in mineralogy such as iron mineral abundances (hematite and magnetite in iron formation), the disappearance of detrital phases (uraninite and pyrite), and the appearance of red beds in the continental sedimentary rock record. Knowledge of the GOE has been enhanced and refined using geochemical proxies derived from bioelemental sediments that span this 100 million year interval. These proxies paint a picture using broad brush-strokes that show the oxygenation of the atmosphereeocean system was more gradual than previously surmised and not a simple linear rise. We demonstrate in this review that the fine lines necessary to further focus this picture can only be attained by interpreting geochemical data in a sedimentologic and oceanographic framework that incorporates an understanding of diagenetic reactions. Although basin diagenetic hydrostratigraphy is rarely, if ever, considered when interpreting the geochemistry of sedimentary and metamorphic rocks, it is a prominent control on diagenesis. What becomes obvious is that geochemical trends often shift along lithofacies changes and sequence stratigraphic bounding surfaces because they have contrasting hydrologic properties. The holistic method advocated in this review mitigates the potential for incorrectly interpreting geochemical trends because it not only considers paleoenvironment and oceanography, but also assesses the effects of fluid flow on alteration. Such an approach should help determine whether trends are local, regional, or truly related to the GOE. Surprisingly, there are currently no studies that interpret highresolution geochemical data in a sequence stratigraphic framework to understand the subtle nuances of Earth’s oxygenation. The need to make such connections and understand the data in their full geologic context is imperative as technological advances continue to increase the rate at which geochemical data are generated. Caution should be exercised so that our ability to measure chemical anomalies keeps pace with our capacity to understand them. Developing a detailed appreciation of how chemical proxies respond to alteration should be a central focus of future work. Only once the alteration of bioelemental sedimentary rocks is better understood can the GOE and Earth’s subsequent oxygenation history be fully interpreted. Acknowledgements

Figure 13. Conceptual framework for integrating sedimentologic and geochemical studies of bioelemental sedimentary systems. TL ¼ transmitted light microscopy; RL ¼ reflected light microscopy; CL ¼ cathodoluminescent microscopy.

This paper was improved through critical review by P.G. Eriksson and four anonymous reviewers. We are grateful to N.P. James, T.K. Kyser, F. Pirajno, T. Clarke, G. Broadbent, D. Rossell, and P.W. Fralick

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for thoughtful discussions that led to this synthesis. Research was supported by a Natural Sciences and Engineering Research Council of Canada Discovery Grant and PetroCanada Young Innovator Award to PKP, and a University of Wisconsin-Oshkosh Research Professorship Grant and a Faculty Development Research Grant (FDR375) to EEH.

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