Oxygenation of the Earth’s atmosphere and ocean system: A review of physical and chemical...

at SciVerse ScienceDirect

Marine and Petroleum Geology 32 (2012) 1e20

Contents lists available

Marine and Petroleum Geology

journal homepage: www.elsevier .com/locate/marpetgeo

Review article

Oxygenation of the Earth’s atmosphereeocean system: A review of physicaland chemical sedimentologic responses

P.K. Pufahl a,*, E.E. Hiatt b,1

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

a r t i c l e i n f o

Article history:Received 22 August 2011Received in revised form30 November 2011Accepted 5 December 2011Available online 14 December 2011

Keywords:Great oxidation eventEarth oxygenationOcean-atmosphere evolutionBioelementalChemistryAlterationSedimentologyDiagenesis

* Corresponding author. Tel.: þ1 902 585 1858; faxE-mail address: [email protected] (P.K. Pufah

1 Tel.: þ1 920 424 7001; fax: þ1 920 424 0240.

0264-8172/$ e see front matter � 2011 Elsevier Ltd.doi:10.1016/j.marpetgeo.2011.12.002

a b s t r a c t

The Great Oxidation Event (GOE) is one of the most significant changes in seawater and atmosphericchemistry in Earth history. This rise in oxygen occurred between ca. 2.4 and 2.3 Ga and set the stage foroxidative chemical weathering, wholesale changes in ocean chemistry, and the evolution of multicelluarlife. Most of what is known about this important event and the subsequent oxygenation history of thePrecambrian Earth is based on either geochemistry or “data mining” published literature to understandthe 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. Becausebiological processes leading to their accumulation often produce organic-rich sediment, black shale canalso be included in the bioelemental spectrum. Thus, chemistry of bioelemental sediments potentiallyholds 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 theratio of stable isotopes (d56Fe, d53Cr, d97/95Mo, d98/95Mo, d34S, D33S) in bioelemental sediments are nowroutinely used to infer the oxygenation history of paleo-seawater. The most robust of these is the mass-independent fractionation of sulfur isotopes (MIF), which is thought to have persisted under essentiallyanoxic conditions until the onset of the GOE at ca. 2.4 Ga. Since most of these proxies are derived fromauthigenic minerals reflecting pore water composition, extrapolating the chemistry of seawater fromsynsedimentary precipitates must be done cautiously.

Paleoenvironmental context is critical to understanding whether geochemical trends during Earth’soxygenation represent truly global, or merely local environmental conditions. To make this determina-tion it is important to appreciate chemical data are primarily from authigenic minerals that are diage-netically 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 adequatelyunderstand 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 anapproach will fine-tune what is known about the GOE because it permits the systematic evaluation ofbasin type and oceanography on geochemistry. This technique also provides information on how basinhydrology and post-depositional fluid movement alters bioelemental sediments. Thus, a primary aim ofany investigation focused on prominent intervals of Earth history should be the integration of geochem-istry with sedimentology and basin evolution to provide amore robust explanation of geochemical proxiesand ocean-atmosphere evolution.

� 2011 Elsevier Ltd. All rights reserved.

: þ1 902 585 1816.l).

All rights reserved.

1. Introduction

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

mailto:[email protected]

www.sciencedirect.com/science/journal/02648172

http://www.elsevier.com/locate/marpetgeo

http://dx.doi.org/10.1016/j.marpetgeo.2011.12.002



P.K. Pufahl, E.E. Hiatt / Marine and Petroleum Geology 32 (2012) 1e202

and Evolution of Atmospheres and Oceans” (Brancazio andCameron, 1964). In 1973 dialogue shifted away from the notionthat purely abiotic processes produced the early atmosphere whenCloud suggested that the deposition of large Paleoproterozoic ironformations was linked to a rise in photosynthetic oxygen (Cloud,1973). More recently, Kasting and Siefert (2002) summarized thecontemporary understanding of the influence of early life on thecomposition of the atmosphere. “microorganisms have probablydetermined the basic composition of the Earth’s atmosphere sincethe origin of life.”

Holland (2002) hypothesized that the emergence of an aerobicbiosphere didnot represent a simple change in the volumeof volcanicoutgassing, but instead was related to a change from reducing tooxidizing volcanic gases. Zahnle et al. (2006) and Konhauser et al.(2009) proposed a decrease in atmospheric methane was the cata-lyst. Although it occurred over an extended interval of time (Willeet al., 2007; Voegelin et al., 2010), this rise in oxygen has becomeknown as the Great Oxidation Event (GOE; Holland, 2002, 2006) andoccurred 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 thebeginning of one the most significant changes the Earth has experi-enced, setting the stage for oxidative chemicalweathering,wholesale

Figure 1. Seawater chemistry and Earth events as related to the three stages of ocean-atmoslargely unknown, but recent d53Cr data suggests that at ca. 1.9 Ga oxygen levels may have dcomplete summary of the geochemical data for Earth’s oxygenation. PAL ¼ present atmosph(2000), Condie et al. (2001), Canfield (2005), Fedonkin (2009), Johnston et al. (2009), Lyon

changes in ocean chemistry, and the evolution of multicellular life(Fig. 1).

The first evidence for the oxygenation of the atmosphere wasbased on mineralogical changes with reduced detrital mineralphases such as pyrite and uraninite in sedimentary rocks giving wayto hematite and other oxide phases (e.g. Cloud, 1968; Roscoe, 1969;Fleet, 1998; Rasmussen and Buick, 1999; Hazen et al., 2008). Mostnew 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 shaleare now routinely used to indirectly deduce the redox conditions ofpaleo-seawater (Fig. 2A, B, C, D; Table 1). Iron formation, phosphoriteand black shale are bioelemental sedimentary rocks that form fromthe nutrient elements Fe, P, and C, which are required for myriad lifeprocesses (Pufahl, 2010). Since the precipitation of these elements isso closely linked to biology, bioelemental sediments are not simplyrecorders of geologic processes, but are intimately involved in theevolution of the ocean-atmosphere system (e.g. Föllmi et al., 1993;Glenn et al., 2000; Simonson, 2003; Huston and Logan, 2004;Malivaet al., 2005; Holland, 2006; Bekker et al., 2010; Pufahl, 2010;Konhauser et al., 2011). Thus, their chemistry holds potentially

phere oxygenation (1, 2, 3). The degree of oxygenation immediately after the GOE is stillipped to pre-GOE concentrations (Frei et al., 2009). See Figure 2 and Table 1 for a moreeric levels; MIF ¼ mass-independent fractionation. Based on data from Farquhar et al.s and Reinhard (2009), Konhauser et al. (2011), and Nelson et al. (2010).

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).The double dashed line is the estimated range in d34S values for SO2�

4 . Lower dashed line is the maximum fractionation with sulfide. B) D33S 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 increasefrom 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 datafrom 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-Sclastic 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 betweenca. 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 ironformations (Konhauser et al., 2009). Decline in Ni at ca. 2.7 Ga may have limited methanogens and contributed to the GOE.

P.K. Pufahl, E.E. Hiatt / Marine and Petroleum Geology 32 (2012) 1e20 3

important clues regarding the development of the early oceansand by extension, the atmosphere making them a logical target forapplication of new geochemical techniques. This attribute, togetherwith the unprecedented development of technology, has spurredthe recent surge in the geochemical investigation of Precambrianbioelemental sedimentary and meta-sedimentary rocks.

Although these technological advancements are resulting inpublication of numerous datasets, it is problematic that our abilityto measure chemical anomalies now surpasses our capacity toadequately understand them (Watson, 2008). This problem isexacerbated because data are often interpreted with little regard tosedimentology, stratigraphy, alteration, and basin evolution. Such

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

Proxy Environmentalparameter

Host mineral(s) Deposit type Effects ofalteration

Comments

d56Fe SeawaterFe-(oxyhydr)oxidelevels and effects onbacterial DIR.

HematiteMagnetiteSideritePyrite

Iron formationBlack shale

Unknown Marked increase in fractionation between ca. 2.7 and 2.5 Gareflects extensive radiation in DIR resulting from increasedproduction of FeO. Photosynthetic oxygen likely caused theoxidation of Fe2þ to create FeO. Produced lower d56Fe values.

d34S Seawater sulfate levelsand effects on BSR.

Pyrite Black shale Unknown Marked increase in fractionation at ca. 2.4 Ga is coincident withGOE. Interpreted to record the transition from sulfate limited tosulfate unlimited bacterial sulfate reduction. Increasedfractionation led to more variability in d34S values.

D33S Absence of ozone andUV shield and effects onMIF of S isotopes.

Pyrite Black shale Unknown MIF of S isotopes interpreted to record absence of free oxygenprior to the GOE. End of MIF interpreted to reflect developmentof ozone layer and a UV shield associated with the GOE.

d53Cr Seawater andatmospheric oxygenlevels and generation ofCr6þ through oxicchemical weathering.

Cr3þ-oxidesassociatedwith FeO

Iron formation Unknown, butinferred immobile.

Increase in the fractionation of Cr isotopes between ca. 2.8 and2.6 Ga is interpreted to record a “whiff” of oxygen prior to theGOE. Decrease in the fractionation at ca. 1.9 Ga likely recordsa dip in oxygen levels to pre-GOE values. Cr6þ is delivered to theoceans during oxic chemical weathering and becomes immobilewhen reduced by Fe2þ to precipitate Cr3þ oxides associatedwithFeO.

d97/95Mo, d98/95Mo Mo oxide levels inseawater and effectson the fractionationof Mo isotopes.

Mo sulfide Black shale Unknown Mo isotopic values suggest euxinic conditions prevailed afterthe GOE between ca. 1.4 and 1.7 Ga. Mo is removed fromseawater by oxic adsorption processes. The isotopiccomposition of these oxides is thought to be transferred toauthigenic Mo sulfides precipitated under reducing conditionsbeneath the seafloor.

REEs (negativeCe anomaly)

Seawater oxygenconcentrationsand Ce behaviour.

Ce3þ on MneFeO. Iron formationPhosphorite

Interpreted topreserve a primarysignature.

Provide information on whether the water column was oxygenstratified. Because scavenging of Ce3þ-oxides by FeO isnegligible; Ce4þ is scavenged on the surfaces of MneFeOproducing the negative Ce anomaly. In this way the resultinglow Ce concentration in seawater is transferred to the sediment.

Trace elements Seawater and porewater redox recordedby differences in theconcentrations of Cr, U,V, Cu, Cd, Zn, Mo, andNi in sediment.

U oxideCr hydroxideV oxideCu, Cd, Zn, Mo,and Ni sulfides

Iron formationBlack shalePhosphorite

Unknown A negative U anomaly and elevated Cr records accumulationunder suboxic and oxic conditions. Elevated U, V, Cu, Cd, Zn, Mo,and Ni reflects deposition under anoxic conditions. Suchdifferences in the trace element concentrations of shallow- anddeep-water lithofacies can indicate whether the water columnwas oxygen stratified.

Mo enrichment inseawater

Mo sulfide Black shale Unknown Mo enrichment in black shale suggests a “whiff” of oxygen 50million years prior to the GOE. Increased delivery of Mo to theoceans via oxic chemical weathering is thought to have led toMo enrichment in black shales that accumulated within anoxicenvironments.

Ni decline in seawater Ni adsorbed to FeO Iron formation Unknown A decline in the Ni concentration of iron formation at ca. 2.7 Gais interpreted to have contributed to the GOE by limitingmethanogens. Ni is a bioessential nutrient for methanogens andwithout it their development was apparently limited allowingoxygen 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 ¼Mn-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 representpaleoenvironmental conditions, are truly global in character, theresult of local environmental factors, or the consequence of alter-ation of what largely are metamorphic rocks.

The sedimentary record of theGOE spans ca.100millionyears andprovides an excellent opportunity to examine the effect of this globalgeochemical revolution (cf. Watson, 2008) on interpreting majorEarth events. The picture that has emergedof the Earth’s oxygenationis based almost exclusively on geochemistry. This approach hasprovided the broad brush-strokes required to understand thisinterval, but the fine lines necessary to refine this picture are onlyattainable by integrating geochemical data in a sedimentologicframework that permits the interpretation of depositional environ-ments, oceanography, and subsequent alteration. The purpose of thisreview is to summarize what is known about the GOE from thebioelemental sedimentary record, and to re-examine the connectionbetween sedimentology, basin history, and the geochemical proxiesused to elucidate changes in ocean-atmosphere oxygenation.

2. The Great Oxidation Event and history of Earth’soxygenation

Although a great deal of controversy still exists about theoxygenation of the Earth (compare Holland, 2004 and Hoashi et al.,2009), there is a consistent interpretation of low Archean andEarly Paleoproterozoic atmospheric oxygen levels (<1e100 ppm O2in the atmosphere), which are followed by higher concentrationsduring the GOE that, after nearly a billion years, gave way to fullyoxygenated 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 ofsignificant change in the redox state of the ocean-atmospheresystem (Huston and Logan, 2004; Canfield, 2005; Holland, 2006;Reddy and Evans, 2009).

Support for the very low levels of oxygen prior to the GOE comesfrom the presence of detrital grains composed of reduced minerals,such as pyrite and uraninite, in sedimentary successions (e.g. Cloud,


1968; Roscoe, 1969; Fleet, 1998; Rasmussen and Buick, 1999;Hazen et al., 2008), and more recently, non-mass-dependent sulfurisotope fractionation, which provides a proxy of oxygen in theatmosphere (Farquhar et al., 2000; Holland, 2006; Reddy andEvans, 2009; Lyons and Gill, 2010). Analysis of Archean and earlyPaleoproterozoic sedimentary sulfide and sulfate minerals hasyielded anomalous variations in the abundance of the four stableisotopes of sulfur (32S, 33S, 34S, 36S). These anomalies are interpretedto result from mass-independent fraction (MIF; D33S, D36S; Figs. 1and 2B; Table 1) involving gaseous sulfur species in the Precam-brian atmosphere with coeval mixing into seawater thatwas marked by low sulfate concentrations (Farquhar et al., 2000;Canfield et al., 2000). MIF is driven by photochemical reactionsinvolving high UV light flux. A prerequisite for these photochemicalreactions is the absence of an effective UV shield such as ozone(Farquhar et al., 2000). Thus, the expression of MIF in sulfurisotopes is interpreted to reflect the near absence of free oxygen inthe 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; Hoashiet al., 2009; Kato et al., 2009; Reinhard et al., 2009), the MIF ofsulfur suggests that O2 concentrations in the Archean atmospherewere generally <10�5 of PAL (Kasting et al., 2001). Evidence fromMo isotopes and PGE concentrations, however, suggest that oxygenlevels may have begun to rise between 2.7 and 2.5 Ga suggestingthat the increase of atmospheric oxygen that led to the demise ofMIF was not a simple linear trend (e.g. Wille et al., 2007). The end ofMIF at ca. 2.4 Ga (Figs. 1 and 2) is interpreted to record the onset ofthe GOE (Bekker et al., 2004). During the following 100 millionyears oxygen levels are interpreted to have risen to >10�2 PAL

Figure 3. Development of red beds. A) Outcrop photo of the Eocene White River Formation, egoethite) that form “dust rims” on detrital grains; these coatings are concentrated on slightlyoxides will eventually recrystallize to form hematite making the rock red. B) Bright red hemRoraima Group, Guyana, South America. C) Photomicrograph in plane-polarized light from tfacies are coated with hematite dust rims (Hr) that underlie pore-filling quartz cement givingplane of the Paleoproterozoic (ca. 2.3 Ga) Lorrain Formation (fluvial facies), Huronian, Blind Rifor the GOE. Photo courtesy of Steve Beyer.

(>0.2% or 2000 ppm; Pavlov and Kasting, 2002; Lyons andReinhard, 2009). What is not known is whether oxygen levelsthrough this protracted interval rose slowly or quickly, or whetherthe increase was constant, marked by punctuated increases, orsome combination of these (compare Bekker et al., 2004; Ohmotoet 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 levelswere probably maintained in the Archean atmosphere and surfaceocean by photo-dissociation of H2O molecules (Kasting et al., 1979).Photochemical breakdown of H2O releases H2O2, which in turndissociates creating O2 (Kasting et al., 1985). Kasting and Walker(1981) determined that Archean oxygen concentrations wouldhave been between 10�12 and 10�14 PAL in the presence of volcanicoutgassed H2 and CO, but up to 4 � 10�8 PAL in the absence of suchgases. Although low, these concentrations would have producedenough 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 formedlong before the GOE (Kasting andWalker, 1981), yet the appearanceof red beds in the stratigraphic record is often cited as evidence forthe GOE (Fig. 1; Holland, 2002). The answer to this paradox liesin how red beds form. Walker (1976) showed that red beds arepreserved during burial diagenesis in the presence of oxygenatedgroundwater when Fe-(oxyhydr)oxides that coat grains (Fig. 3C)recrystallize to form hematite (Fig. 3A, B, C). However, if ground-water was anoxic, iron (oxyhydr)oxides (Fig. 3A) formed at theEarth’s surface would have dissolved during early burial beforequartz cement overgrowths could precipitate and protect thesecoatings, leaving no record of oxidation (e.g. Surdam and Crossey,

asternWyoming, USA. This sandstone is stained with Fe-(oxyhydr)oxides (limonite andmore permeable laminae and highlight cross bed foresets. The metastable Fe-(oxyhydr)atite-stained quartz arenite and siltstone red beds from lacustrine facies of the 1.9 Gahe 1.7 Ga Thelon Formation, Nunavut, Canada. Detrital quartz grains (Dq) in this eolianthis quartz arenite a red color. D) Outcrop showing an upturned, ripple-marked beddingver, Ontario. This red bed successionwas one of the examples originally used as evidence


1987; Lovley et al., 1991). Thus, the appearance of red beds coincideswith the GOE because groundwater became sufficiently oxygenatedto retain Fe-(oxyhydr)oxides in the shallow burial realm andpotentially preserve them in the sedimentary record.

Most researchers agree that the evolution of oxygenic photo-synthesis within cyanobacteria was the source of oxygen thatcaused the GOE (cf. Cloud, 1973). The timing of cyanobacterialevolution, however, is problematic since biomarkers indicate theymay have evolved as early as ca. 2.9 Ga (Nisbet et al., 2007) andwere abundant by ca. 2.7 Ga (Brocks et al., 2003, 2005; Canfield,2005; Buick, 2008), at least 400 million years before the onset ofthe GOE. This long lag likely represents a period of inertia whereoxygen-consuming chemical reactions prevented the rise ofphotosynthetic oxygen by consuming oxygen in inorganic reactionswith 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 byBlank and Sánchez-Baracaldo (2010) further suggest the earliestcyanobacteria were restricted to freshwater environments until ca.2.4 Gawhen they diversified and exploitedmarine ecosystems. Thisdiversification, extraordinary increase in habitat, and the resultingextensive organic carbon flux to the deep oceans could have causeda rapid increase in oxygen during the GOE. Based on this samemolecular clock model, mat-forming cyanobacteria with filamen-tous forms, large sizes and that fixed nitrogen appeared at ca. 2.3 Ga(Blank and Sánchez-Baracaldo, 2010).

Although photosynthetically produced oxygen was the primarydriver of oxygenation during the GOE, a number of processes arepostulated to have played a role in changing the redox state of the-Paleoproterozoic ocean-atmosphere system. These include: (1)increased burial of organic matter (DesMarais et al., 1992; Melezhiket al., 2005); (2) loss of hydrogen to space from a methane-richatmosphere (Kasting et al., 1979; Catling et al., 2001); (3) collapseof atmosphericmethane (Zahnle et al., 2006;Konhauseret al., 2009);(4) changes in the redox potential of volcanic gases (Kump et al.,2001; Holland, 2002); (5) nutrient loading and increased produc-tionof cyanobacterial oxygen (Papineauet al., 2009); and (6) aperiodof major continental growth at the Archean-Proterozoic boundary(Godderis and Veizer, 2000). The collapse of a methane-rich atmo-sphere is also thought to have been an important contributor tothe onset of Paleoproterozoic ice ages (Reddy and Evans, 2009).

Little is known regarding oxygen levels immediately followingthe GOE (2.0e1.8 Ga; Fig. 1). Frei et al. (2009) interpret a decreasein oxygen, possibly dropping to pre-GOE levels, based on a changein chromium isotopic values from a small dataset (Fig. 2D; Table 1).Because a major interval of black shale deposition correspondsto this interval (Fig. 1), however, it is likely that any decreasein photosynthetic oxygen production would be at least partiallycompensated for by removal of organic carbon from the ocean-atmosphere system. Thus, this interval should be explored furtherto elucidate whether there was in fact a major dip in oxygenconcentration 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 hypothesizedto have led to oxic chemical weathering of the continents, whichoxidized sulfide minerals to produce sulfate that was delivered byrivers to the ocean. In this model, dissolved sulfate delivered tothe oceans by rivers was transformed through bacterial sulfatereduction into sulfide causing euxinic conditions that developed atthe end of the Paleoproterozoic (Canfield, 1998, 2005; Poulton andCanfield, 2011; Kendall et al., 2011). By ca. 1.85 the flux of sulfatewas great enough to cause sulfidic intermediate and bottomwaters(Fig. 1; Poulton et al., 2004; see also Pufahl et al., 2010). Widespreadeuxinia may have been perpetuated by thriving anoxygenic

photoautotrophs that tempered oxygen production by using sulfideas an electron donor (Johnston et al., 2009). These conditions arehypothesized to have prevailed for nearly a billion years and alsoperturbed the cycling of bioessential elements, possibly causinga long stasis in the evolution of eukaryotes (Anbar and Knoll, 2002).This period is often referred to as the “Boring Billion” becausebiological evolution is thought to have stagnated during this pro-tracted interval (Anbar and Knoll, 2002; Holland, 2006).

Oxygen concentrations increased to >10% PAL (>0.2% or2000 ppm) during the Neoproterozoic ‘snowball glaciations’ (Fig. 1;Canfield, 2005; Holland, 2006). Ice cover that shrouded the Earthbetween ca. 740 and 630 Ma is thought to have slowed chemicalweathering and delivery of sulfate to the oceans, causing thedemise of widespread euxinia. This set the stage for the Earth’stransition from its prokaryote-dominated middle age by removingsulfide, a physiological barrier to eukaryote diversification(Johnston et al., 2010). For the first time in Earth history thecomplete dominance of oxygenic photosynthesis led to the venti-lation of the deep ocean. By ca. 580 Ma bottom waters wereoxygenated enough to stimulate the evolution of multicellularbenthic animals (Canfield et al., 2007; Narbonne, 2010). Withcontinued input of photosynthetic oxygen, Phanerozoic oxygenlevels were achieved by ca. 540 Ma (Holland, 2006).

3. Bioelemental sediments and the record of Earth’soxygenation

The sedimentary and geochemical record of the GOE ispreserved primarily in bioelemental sediments, a relatively newclassification of sedimentary rocks that encompasses iron forma-tion, chert, and phosphorite (Pufahl, 2010). Because bioelementalsediments are precipitated directly or indirectly by biologicalprocesses they are often associated with organic-rich deposits suchas black shale, which can be included in the bioelemental spectrumsince it contains biologically fixed C.

The occurrence of bioelemental sediments through time reflectschanges in ocean chemistry linked to climate change, biologicevolution, and tectonic processes (Fig. 4). These factors haveinfluenced the biogeochemical cycling of Fe, Si, P and C (e.g. Loganet al., 1995) and the types of bioelemental sediments producedbefore, during, and after the GOE. Thus, the temporal distribution ofbioelemental sediments provides a framework for understandingthe nature of the GOE (Fig. 4) and associated long-term changes toocean-atmosphere chemistry.

Also important are changes in the stacking patterns ofbioelemental lithofacies because the redox-sensitive minerals andchemical proxies they contain provide the most detailed informa-tion about shifts in water column oxygenation. The best recordsof seawater oxygenation come from pristine lithofacies. Pristinesedimentary facies are generally fine-grained and accumulate incalm environments. They are characterized by undisturbed watercolumn precipitates and/or in situ authigenic minerals.

In a very general sense, the occurrence of bioelemental sedimentsincreased after the onset of the GOE and coincides with a conspic-uous 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). Aschert 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, marinechemical sedimentary rock (Figs. 1, 4e6; e.g. Gross (1983); Clout andSimonson, 2005; Klein, 2005; Bekker et al., 2010; Pufahl, 2010). The

Figure 4. Temporal distribution of iron formation (red), ironstone (purple), phosphorite (yellow) and black shale (black). Based on deposit age, resource estimates and timing ofEarth 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 ofoxygenic 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 toorestrictive (e.g. Klein, 2005). In weakly metamorphosed iron forma-tion common minerals include the Fe-oxides hematite and magne-tite aswell as the silicates chert, greenalite, and stilpnomelane (Klein,2005).Oxygenationof the oceanduring theGOE,with eitherdirect orindirect involvement of Fe-oxidizing bacteria, is believed to beresponsible for deposition of all large Paleoproterozoic iron forma-tions (Figs. 4e6; e.g. Cloud,1973; Konhauser et al., 2002). In additionto the importance of iron formation as a recorder of oxygen levels onthe early Earth, it is economically significant because it containsmostof the world’s iron ore.

3.1.1. Temporal distributionThe Archean is characterized by pyrite andmagnetite-rich deep-

water exhalative iron formation deposited in tectonically activeareas around spreading centers associated with volcanic arcs.The dramatic rise in iron formation at ca. 2.8 Ga may correspond tothe evolution of oxygenic photosynthesis (Nisbet et al., 2007) andresulting precipitation of ferrous Fe from the Archean ocean (Fig. 4).Although some evidence suggests that iron formation prior to thistime was also linked to photosynthetic oxygen (e.g. Hoashi et al.,2009), most data indicate deposition through a combination ofanoxygenic photosynthesis, dissimilatory iron reduction, oxygenproduced via nonphototrophic sources, and episodic increases inthe 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 signala shift from deep-water deposition to upwelling-driven, neriticaccumulation on the expansive, unrimmed platforms that devel-oped at the end of the Archean (Fig. 7; Cloud, 1973; Klein, 2005;Pufahl, 2010; Bekker et al., 2010). Such aerially extensive Paleo-proterozoic iron formation formed in the full spectrum of shelfenvironments from an oxygen-stratified ocean born during theGOE (Pufahl, 2010). Precipitation occurred when ferrous Fe inupwelled, anoxic waters was either mixed with photosyntheticallyoxygenated seawater or oxidized during anoxygenic, bacterialphotosynthesis (Fig. 7; Cloud, 1973; Klein, 2005; Konhauser et al.,2002; Bekker et al., 2010; Pufahl, 2010). Chert formed abiogeni-cally primarily in subtidal environments where evaporiticconcentration (Maliva et al., 2005) and Fe-redox pumping couldsaturate bottom- and pore water with silica (Fischer and Knoll,2009; Pufahl, 2010). A suboxic seafloor was a prerequisite for Fe-redox pumping to saturate sediment with silica. Such conditionsare interpreted to have occurred in coastal environments wherephotosynthetic oxygen oases impinged on the seafloor (Nelsonet al., 2010; Pufahl, 2010). Silica was concentrated in pore waterduring burial when Fe-(oxyhydr)oxides dissolve below the suboxic-anoxic 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 reflectthe increased precipitation of oxidized Fe from seawater as well as

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 largestromatolite 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 SokomanFormation, 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 theArchean, peaks in iron formation abundance through the Protero-zoic have also been correlated to mantle plume activity (Isley andAbbott, 1999; Abbott and Isley, 2001).

Deposition of iron formation on continental margins ceased atca. 1.8 Ga, possibly due to the development of widespread euxinia(Fig. 4; Canfield, 1998; Poulton et al., 2004; Kendall et al., 2011). Ina sulfidic water column dissolved sulfide is interpreted to havecombined with ferrous Fe to form pyrite, titrating the ocean ofdissolved Fe. An important change in the Precambrian Si cycle alsooccurred at this time and is marked by the end of subtidal chertdeposition (Maliva et al., 2005). This change is thought to reflectwaning hydrothermal input of Si and a decrease in Si derived fromchemical weathering.

Sulfidic ocean conditions are interpreted to have continuedfor nearly a billion years (Anbar and Knoll, 2002). Bioessential traceelements were largely removed from the oceans as sulfides asso-ciated with organic matter-rich sediments, which is thought tohave contributed to the apparent lull in eukaryote evolution (Anbarand Knoll, 2002). The period that followed these changes is termedthe ‘Boring Billion’ because there appears to have been little changein the atmosphereeocean and biological systems over this pro-tracted interval of Earth history (Fig. 4). Iron formation finallyreappears 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 chemistryUnfortunately, there are only a few integrated studies that

couple sedimentology,mineralogy and geochemistryof bioelementaldeposits bracketing the GOE. Most of those that do focus on thedisposition and chemistry of suboxic and anoxic lithofacies formingthe 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 becausethe presence of a prominent oxygen chemocline is the primarycontrol on facies mineralogy (Fig. 7; Pufahl, 2010).

Deposition of suboxic lithofacies occurred along segments of thecoastline where photosynthetic cyanobacteria produced oxygen(Figs. 5A and 7). These deposits are characterized by hematite(Fe2O3) and chert (SiO2; Fig. 8; Klein, 2005). Anoxic lithofaciesare 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 mineralscontain some reduced Fe, and reflect precipitation under extremelylow oxygen concentrations (ca. 10�20 pO2-water and were likely aslow as 10�70 pO2-water; Mel’nik, 1982).

Figure 6. Iron formation lithofacies from the Neoarchean-Paleoproterozoic Hamersley Basin, Western Australia. As in Figure 5, mineralogy can reflect the redox conditions of paleo-seawater/pore water. Arrows denote younging direction. A) Laminated magnetite and chert. Late Neoarchean Marra Mamba Iron Formation, Western Australia. B) Laminated magne-tite and chert. Early Paleoproterozoic Joffre Iron Formation,Western Australia. C) Interlaminatedmagnetite, chert and fine-grained, hematitic grainstone laminae. Early PaleoproterozoicJoffre Iron Formation, Western Australia. D) Laminated magnetite and chert intercalated with thin beds of hematitic grainstone. Early Paleoproterozoic Joffre Iron Formation, WesternAustralia. 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 REEcomposition of iron formation have also been used to inferoxygenation history (Fig. 2C and 7; Table 1; e.g. Beukes and Klein,1990; Klein, 2005; Johnson et al., 2008). The REE systematics ofredox sensitive facies, however, seems more robust and easier tointerpret, primarily because it is a direct measure of seawater Eh(Elderfield and Greaves, 1982) without the issues of potentiallystrong and not yet understood biologic fractionations (Johnsonet al., 2008). In general, iron formation and chert facies formed inoxygenated marine environments have negative Ce anomalies andare enriched in heavy REE’s (HoeLu) when compared to shales(Klein, 2005). This is because although the oxidation of Ce3þ greatlyreduces Ce solubility, oxidative scavenging on the surface of freshlyprecipitated Fe-(oxyhydr)oxides removes Ce from seawater (Ohtaand 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 enrich-ment of heavy REE’s (Byrne and Sholkovitz, 1996) is also inter-preted to be the result of preferential oxidative removal of the otherlight REE’s (LaeDy) from seawater. These processes will onlyproduce a “seawater pattern” if deposition occurs away from aterrigenous clastic source since siliciclastic material has no Ceanomaly or heavy REE enrichment (Watkins et al., 1995).

Recently, the Ni concentration in iron formation has been usedto infer both the timing and cause of the GOE (Fig. 2E; Table 1;Konhauser et al., 2009). A significant decrease in the Ni/Fe ratioat ca. 2.7 Ga is interpreted to correlate to a major drop in theconcentration of Ni in seawater. Because of their insatiable appetitefor Ni, this change likely limited methanogens in the Neoarcheanand led to a concomitant reduction in the generation of atmo-spheric methane. With decreasing methane and the otherenvironmental 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 cyanobacterialoxygen and the GOE.

3.2. Phosphorite

Phosphorite is a bioelemental sedimentary rock rich in P, is oftenassociated with coastal upwelling, and occurs almost exclusively inthe Phanerozoic (Figs. 1, 4 and 9). It is defined as a rock with greaterthan 18 wt. % P2O5, but P2O5 can be as great as 40 wt. %, makingthese rocks an important fertilizer ore (Pufahl, 2010). Most pub-lished accounts of Proterozoic and Neoproterozoic phosphorites donot describe true phosphorite, but phosphatic deposits that containmuch less than 18 wt. % P2O5. This distinction is important becauseuncritical reporting of phosphatic occurrences has resulted inan over estimation of Precambrian phosphorite, which has led toerrors in assessing temporal abundance and understanding thePrecambrian P cycle (e.g. Papineau, 2010).

Phosphorite forms through phosphogenesis, the authigenicprecipitation of francolitewithin sediment just beneath the seafloor(Glenn et al., 1994). Francolite is a highly substituted carbonatefluorapatite (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 bythe redox potential of bottom- and pore water (Jahnke et al., 1983;Glenn et al., 1994). Authigenic, biological, and hydrodynamicprocesseswork together to formphosphatic laminae, in situ nodulesor reworked granular beds (Föllmi et al., 1991; Föllmi, 1996).

Phosphorite is the most important long-term sink in theglobal phosphorus cycle. In the Phanerozoic the majority of P in theoceans is sequestered in marine sediment on continental marginsand beneath regions of active coastal upwelling (Filippelli, 2008;Fig. 10). Phosphorus is removed from nutrient-rich surface watersby phytoplankton and authigenically converted to francolite inaccumulating organic-rich sediment through a series of microbiallymediated redox reactions (Jahnke et al., 1983; Glenn et al., 1994).Bacterial sulfate reduction is the most efficient of these reactionsat liberating organically bound P to pore water (Arning et al., 2009).Precipitation of francolite occurs when pore water becomessupersaturated with respect to calcium-phosphate (Glenn et al.,1994). Such phosphorite co-occurs with biogenic chert and black

Figure 7. Continental margin iron formation. Lithofacies formed a sedimentary wedgethat fines and thickens basinward. Coastal upwelling provided a sustained supply ofanoxic bottom water rich in dissolved Fe and Si. Precipitation occurred in an oxygenstratified water column that was suboxic down to fair-weather wave base. Nearshorelithofacies consist of cross-stratified grainstones that are commonly stromatolitic.Laminated pristine lithofacies accumulated in low energy environments such asshallow lagoons and below fair-weather wave base on the middle and distal shelf. REEspidergrams show the behaviour of Ce across the shelf. A negative Ce anomaly is mostpronounced along segments of the paleoshoreline that were oxygenated by photo-synthesis. It disappears offshore where bottom and intermediate waters were anoxic.The positive Eu anomaly reflects the hydrothermal source of Fe (Klein, 2005 andreferences therein). SWB ¼ stormwave base; FWB ¼ fair-weather wave base. Modifiedfrom Pufahl (2010).


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

Figure 8. Paragenesis typical of pristine iron formation in suboxic and ano

3.2.1. Temporal distributionPhosphorite did not form in the Archean (Fig. 4), likely reflecting

weathering of phosphate-poor, mafic crust under an anoxic atmo-sphere (Pufahl, 2010). The appearance of phosphorite inthe Paleoproterozoic coincides with the GOE and the onset of oxicchemical weathering of the continents (Papineau, 2010; Pufahl,2010). This relatively minor phosphatic episode was not associatedwith upwelling and unlike Phanerozoic phosphorites, restrictedto shallow-water environments (Nelson et al., 2010). It occurredbetween 2.2 and 1.8 Ga, beginning just after the Huronian Glaciationand in the middle of the GOE (Papineau, 2010; Pufahl, 2010).This episode probably records an abrupt increase in the delivery ofphosphate to the oceans. Increased phosphate likely fueled a corre-sponding increase in primary production that enhanced photosyn-thesis and the contribution of oxygen to the GOE (Papineau, 2010).This pulse may be the consequence of a switch to post-glacialcontinental chemical weathering under an oxygenated atmospherefrom a long period dominated by mechanical weathering duringthe Huronian Glaciation (Papineau, 2010; Pufahl, 2010). Thus, theappearance of Paleoproterozoic phosphorite is directly linked tothe GOE and the oxygenation of the oceans (Nelson et al., 2010;Pufahl, 2010).

Phosphogenesis during this initial episode was restricted tosegments of the shoreline that were silica undersaturated andoxygenated through microbial photosynthesis. These conditionspermitted a combination of bacterial sulfate reduction and Fe-redoxpumping to concentrate P in coastal sediment (Fig. 11; Nelson et al.,2010). Such shallow-water phosphorite is in stark contrast toupwelling-related, Phanerozoic phosphorites that accumulated ina range of shelf environments. This difference likely reflects thedissimilarity in the oxygenation state of the seafloor (Nelson et al.,2010). The anoxia that typified Precambrian intermediate andbottom water prevented Fe-redox pumping from operating indeeper settings (Fig. 11).

During the onset of sulfidic ocean conditions the Fe and P cyclesbecame decoupled, which led to the disappearance of phosphorite

xic paleoenvironments. Modified from Klein (2005) and Pufahl (2010).

Figure 9. Precambrian phosphorite lithofacies. Most Precambrian phosphorites are unlike Phanerozoic phosphatic deposits because they do not form aerially extensive deposits. Theygenerally consist of thinpristine phosphorite inperitidal environments and granular phosphatic lags in shallow-water lithofacies. A) Laminatedpristine phosphorite. Subhedral crystals arepyrite, 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 grainsindicating 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.

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


at ca. 1.8 Ga (Fig. 4; Pufahl, 2010). The precipitation of pyrite ina euxinic water column decreased the potential for Fe-redoxpumping, even in nearshore oxygen oases (Nelson et al., 2010).Widespread sulfidic conditions likely made bacterial sulfatereduction ineffective as a driver of phosphogenesis because phos-phatewould have been released to thewater columnwhere it couldbe efficiently recycled and not fixed as francolite in the sediment(Nelson et al., 2010). Phosphorite, like iron formation, was notdeposited again until the Neoproterozoic (Fig. 4).

3.2.2. Deposition and chemistryAlthough 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 withthe sedimentology and chemistry of co-occurring bioelementalsediments. Francolite readily incorporates a variety of redoxsensitive trace elements into its crystal structure and thus, oftenpreserves a record of pore water and bottom water Eh duringdeposition (Fig. 10; Jarvis et al., 1994). Trace elements generallyreplace Ca2þ in francolite, but can also be transferred to the sedi-ment by absorption onto crystal surfaces, scavenging by organicmatter, 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 conditionsare Cu, Cr, V, Cd, Mo, U, and Zn (Fig. 10). All but U are mobilized andincorporated under reducing conditions. As in iron formation, theREE content of francolite records seawater values although itcontinues to absorb REE from pore water below the sedimente-water interface (Altschuler, 1980; Piper et al., 1988). As in ironformation, the presence of a prominent negative Ce anomaly indi-cates precipitation in oxygenated environments (Piper et al., 1988;Jarvis et al., 1994).

In addition to trace elements, the stable isotopic compositionof 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). Temperaturecalculations are sometimes coupled with trace element analysis toinfer the redox conditions and paleooceanography of ancient seas(e.g. Hiatt and Budd, 2003).

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 boundarythey dissolve, liberating sorbed HPO4

2- to pore water. Francolite precipitation is limited in the sediment by the availability of seawater-derived F�. Although important in stimulatingphosphogenesis in the Phanerozoic, bacterial sulfate reduction was likely much less efficient at promoting the precipitation of francolite in the Precambrian because of the very lowseawater 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 theoxygenation state of the seafloor. In the Precambrian, photosynthetically oxygenated nearshore environments possessed suboxic seafloors that facilitated Fe-redox pumping andphosphogenesis. Phosphogenesis could not occur in the middle and distal shelf because these regions were below the oxygen chemocline. Phosphogenesis in the Phanerozoicoccurs 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 fine-grained clastic sedimentary rock (Fig. 12) that is rich in organicmatter (>2 wt. %), sulfides (especially pyrite), and redox sensitivetrace elements (U, V, Cu, and Ni; Arthur and Sageman, 1994; Piperand Calvert, 2009). It can form in a wide range of paleoenviron-ments from peritidal to deep basin settings, is often associated withphosphorite, and can be a hydrocarbon source rock (Fig. 10).

Black shales are commonly interpreted as recording depositionbeneath a highly productive surface ocean or within anoxic, sulfidicbottom waters, or a combination of both (Piper and Calvert, 2009).Recent work, however, suggests that high planktic productivityis the most important control on organic matter enrichmentin marine sediment (e.g. Piper and Calvert, 2009 and references

therein). Organic matter accumulates because the rate of produc-tion and settling is greater than the rate of degradation of organiccarbon on the seafloor (Pedersen and Calvert, 1990). Sinceprocesses of black shale deposition can occur across the spectrumof shelf environments, their occurrence is not always an indicationof accumulation in a deep, open ocean basin.

Processes leading to the formation of black shale are importantbecause they link the various pools of carbon in the ocean-atmosphere system (Arthur and Sageman, 1994). These processesgovern carbon burial, which regulates climate, and oxygen levels bycontrolling the rate reduced C is sequestered in the geologic record(Holland, 2002; Canfield, 2005). Since P is the primary controlon productivity over geologic time scales the phosphorus cycleultimately determines the rate of organic matter burial andremoval of carbon dioxide from the atmosphere.

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. PermianPhosphoria 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 fromweathered sedimentary sulfides. Organic-rich laminae in (C) are commonly scoured by very fine-grained, thinly bedded sandstone layers. Dashed line highlights a scoursurface. 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 distributionThe temporal distribution of black shale is even less well con-

strained than that of phosphorite (Figs. 1 and 4), primarily becauseit forms components of other depositional systems. The record ofPrecambrian black shale is also severely biased given the rarity ofpreserved deep-sea sediments, and because they are easily eroded.In general, however, the timing of black shale deposition reflectsperiods when oxygen concentrations could increase in the atmos-phereeocean system (Fig. 4; e.g. Berner, 2004). Secular changesin black shale deposition result from changes in carbon cycling in“active” surface ocean pools, in the atmosphere, on land, and inmarine sediment, and carbon pools that cycle on much longertimescales (Burdige, 2006).

Such changes are partly linked to the GOE (Des Marais et al.,1992), which apparently follows an episode of enhanced carbonburial in the late Archean (Fig. 4). This is the first of three noticeablepeaks in black shale deposition during the Precambrian (Fig. 4;Condie et al., 2001). It is the least prominent and occurs in theNeoarchean between ca. 2.7 and 2.5 Ga. This initial pulse of blackshale accumulation is thought to correspond to either a mantleplume event, which through climate warming, increased chemicalweathering and nutrient delivery to the oceans (Condie, 2004),or a change in ocean currents (Condie et al., 2001) that resultedin initiation of upwelling along favorably positioned cratons.The sequestration of reducing organic matter during this episode isinterpreted to have contributed to the GOE (Des Marais et al., 1992).

The second pulse is a prominent event occurring between ca. 2.0and 1.7 Ga (Fig. 4; Condie et al., 2001), just after the HuronianGlaciation. As with iron formation of this age, the accumulation ofblack 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 Pfluxes to the ocean that not only stimulated primary productionand phosphogenesis (Nelson et al., 2010), but also black shaledeposition as well.

Black shale again becomes conspicuous in the Cryogenianbetween ca. 800e600 Ma (Fig. 4). Organic-rich mudstones, some of

which are phosphatic, accumulated during retreat of the “snow-ball” glaciations (Condie et al., 2001; Le Heron et al., 2009). Elevatedsurface ocean productivities were likely sustained by delivery ofnutrients through glacial runoff and invigorated coastal upwelling(Papineau, 2010). The pronounced equator-to-pole temperaturegradient that develops during glaciations leads to more energeticatmospheric circulation and thus, coastal upwelling, resulting inthe widespread accumulation of organic-rich sediment (Vincentand Berger, 1985).

Correspondence between peaks of black shale and those of ironformation deposition in the Precambrian (Fig. 4) highlights theimportance that photosynthetic oxygen played in the accumulationof iron formation. This relationship also emphasizes the connectionbetween ocean circulation and upwelling to deliver reduced ironand P to the photic zone.

Like phosphorite, pulses of black shale deposition in the Phan-erozoic are linked to ocean-climate feedback (Bluth andKump,1991;Arthur and Sageman, 1994). Prominent peaks are also the conse-quence of enhanced P burial from invigorated coastal upwelling orincreased chemical weathering and delivery of phosphate to theoceans (Fig. 4; Glenn et al., 1994; Föllmi, 1996).

3.3.2. Deposition and chemistryMuch 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 andGrimm, 2003; Raiswell et al., 2011). Textural relationships betweenglauconite, pyrite, francolite, and carbonate provide a high fidelityrecord of the physical, chemical, and biologic processes causingsubtle shifts in redox potential (Glenn and Arthur, 1988; Pufahland Grimm, 2003). These minerals precipitate through a series ofmicrobially mediated redox reactions (Froelich et al., 1979; Glennet al., 1994). In order of decreasing energy yield these reactionsinclude oxic respiration, denitrification, transition metal oxidereduction, sulfate reduction, and methanogenesis. Geochemicalevidence suggests that all but oxic respiration evolved by the lateArchean (Garvin et al., 2009; Lyons and Gill, 2010), and aerobic


heterotrophs evolved by ca. 2.1 Ga in response to increasing oxygenlevels (Papineau et al., 2005). Because of widespread ocean anoxiabacteriaedriven reactions that produce and consume organicmatterwere not confined to below the seafloor, but also occurredwithin thewater column.

Pyrite precipitates below the sulfate redox interface through theconversion of monosulfides formed during bacterial sulfatereduction (Schieber, 2002). Sedimentary pyrite is often framboidaland finely disseminated (Raiswell, 1982; Wilkin and Arthur, 2001;Schieber, 2002), but discrete layers have been interpreted asrecording precipitation and suspension settling through a euxinicwater column (Poulton et al., 2004). Francolite precipitates inassociation with the microbial reduction of nitrate, Mn-oxides,Fe-oxides, and sulfate (Pufahl, 2010). Unlike the formation ofpyrite, however, phosphogenesis is not a redox-controlled reaction,but is regulated only by the concentration of phosphate in porewater (Glenn et al., 1994).

In addition to pyrite and francolite, other authigenicminerals 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 ofwhich can be used to further constrain pore water Eh. Glauconiteforms first, within suboxic pore water at the Fe-redox interface,followed by pyrite, and then carbonate at progressively deeperlevels in the sediment. Glauconite occurs as authigenic peloids,coatings, or cement. Carbonate minerals precipitate within thealkalinity maximum that develops during intense microbial respi-ration. The type of carbonate mineral produced depends on theavailability of Ca2þ, Fe2þ and Mg2þ. Calcite forms from pore waterwith little Fe2þ and Mg2þ, whereas siderite precipitates in anoxicpore water with abundant Fe2þ (François and Gérard, 1986; Klein,2005). Dolomite is created in pore water enriched in Ca2þ andMg2þ when the bacterial reduction of SO2�

4 , a kinetic inhibitor todolomite precipitation, is converted to H2S (Baker and Kastner,1981; Kastner, 1984; Wright and Wacey, 2005). These carbonateminerals are generally a microcrystalline cement that binds detritalgrains and earlier authigenic phases together, but can also formdisplacive concretionary horizons (Kholodov and Butuzova, 2004).

The bulk trace element content of black shales and the sulfurisotopic composition of co-occurring pyrite provide evidence ofchanging seawater Eh over longer timescales. An increase in theconcentration of redox-sensitive trace elements and an increase inthe fractionation of d34S roughly correspond to the onset of the GOE(Fig. 2A; Table 1). The restricted range of d34S values in pyrite prior toca. 2.5 Ga is interpreted to reflect low seawater SO2�

4 concentrationsof the Archean; the consequence of negligible SO2�

4 delivery fromanoxic chemical weathering (Canfield, 2001). Low SO2�

4 levelsrestrict bacterial sulfate reduction andproduce little variation in d34Svalues (Canfield, 2001). After ca. 2.5 Ga fractionations increasedramatically to values expected for bacterial sulfate reduction,which is not limited by low sulfate concentrations. Higher sulfatelevels were produced by weathering of pyrite under an oxygenatedatmosphere (Canfield, 2001).

MIF of sulfur isotopes (D33S; Figs. 1 and 2B; Table 1) in pyriteprovides the best evidence of Precambrian atmospheric oxygenlevels and the timing of the GOE (Farquhar et al., 2000; Farquharand Wing, 2003). The nature of its onset is preserved in therecord of multiple sulfur isotope distributions (d34S-D33S), whichsuggests that oxygen levels began to fluctuate ca. 150 million yearsprior to the permanent rise at ca. 2.4 Ga (Partridge et al., 2008; cf.Wille et al., 2007). The decreased variability and appearance ofpositive pyrite d56Fe values after ca. 2.3 Ga corroborate D33S data(Fig, 2; Rouxel et al., 2005), but it is unclear whether these changesreflect 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 risein oxygen levels ca. 150 million years before the accepted onset ofthe GOE (Wille et al., 2007; Voegelin et al., 2010). Molybdenumisotope data also suggest that although euxinic conditions mayhave eventually developed in the late Paleoproterozoic, the oceanwas probably not a “global Black Sea” (Lyons et al., 2009).

Another approach that is increasingly being used is the analysisof carbonate-associated-sulfur (CAS; e.g. Guo et al., 2009). BecauseCAS can acquire the isotopic composition of pore water andseawater (Burdett et al., 1989) it is particularly attractive as a pale-oredox proxy in Precambrian limestones. Sulfur isotope data fromassociated pyrite also allows potential calculation of the offsetbetween SO2�

4 and H2S during bacterial sulfate reduction (Lyonsand Gill, 2010), further constraining the nature of redox sensitivemicrobial 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 sedimentchemistry, especially in deposits as old as the Precambrian (e.g.Hayes et al., 1983; Ayalon and Longstaffe, 1988; Crusius andThomson, 2000; Shields and Stille, 2001; Petsch et al., 2005;Gonzalez Alvarez and Kerrich, 2010; Hiatt et al., 2010). Thus, it isdifficult to reconcile why so few studies of Precambrian depositionalsystems attempt to understand alteration of what are primarilymetamorphic rocks and minerals before interpreting geochemicaldata. This is especially true since techniques exist to assess whetherobserved 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 therapid analysis of samples (Watson, 2008) also lead to the briskpublication of data without full assessment of potential alteration.

4.1. Sedimentology, basin evolution, and alteration

The current level of understanding bioelemental sedimentalteration is at about the same stage as understanding limestonediagenesis was 30 years ago. A basic tenet that has importantimplications for understanding geochemical proxies is that themostdesirable deposits for geochemical analysis are pristine lithofacies.Like lime mudstones, fine-grained bioelemental facies usuallyrepresent seawater and authigenic precipitates with low hydraulicconductivities that tend to fix pore water chemistry close to thetime of deposition (e.g. Kyser et al., 1998). This fact is commonlyoverlookedwhen extrapolating the chemistry of authigenicmineralsto the overlying water column. Coarser facies have higher fluid/rockratios and experience greater fluid fluxes during burial that oftenresults in chemical compositions that are significantly different fromoriginal ones (e.g. Reeckman, 1981).

Sedimentologic and basin evolution context are both paramountwhen interpreting geochemical data. A properly constrained depo-sitional and post-depositional framework provides informationon how oceanography, depositional environment, seawater andpore water chemistry, microbial biology, and alteration influencethe chemical composition of bioelemental sediments. Withoutthis perspective it is a challenge to interpret whether geochemicalanomalies through the GOE are the consequence of local environ-mental factors or global in character (Lyons et al., 2009; Pufahlet al., 2010). In most cases ascertaining the nature of anomalies inPrecambrian sedimentary rocks is especially difficult given therarity of preserved deep-sea sediments (Pufahl et al., 2010). Unfor-tunately, this issue is often overlooked resulting in conclusions thatare not fully supported by sedimentologic data. For example, key


stratigraphic units in the Paleoproterozoic Pretoria Group, wheremuch 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). Sedimentologicevidence suggests that epeiric seas are typified by restricted circu-lation patterns that produce water masses with compositions thatdiffer substantially from the open ocean (e.g. Hiatt and Budd, 2001;Piper, 2001; Algeo and Heckel, 2008).

In all sedimentary basins diagenetic hydrostratigraphy iscontrolled primarily by lithofacies and resulting diagenetic reac-tions, 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 proper-ties on a basin-scale because they control the composition, fabric andgrain size of lithofacies (e.g. Hiatt and Budd, 2003). Widespreadlateral flow of diagenetic fluids can occur over distances of hundredsof kilometers and at temperatures >200 �C (Hiatt and Kyser, 2007;Kyser, 2007; Alexandre et al., 2009; Hiatt et al., 2010). Diageneticfluids flow most intensively within coarse-grained facies thathave not experienced intense cementation and are situated aboveunconformities or along internal discordances such as parasequenceboundaries (Hiatt et al., 2003; Hiatt et al., 2007; Hiatt and Kyser,2007). Stratigraphic boundaries often allow preferential diageneticand metamorphic fluid flow that can reset relatively robustgeochemical proxies, such as d34S (Pufahl et al., 2010). This long-termdiagenetic hydrostratigraphy can involve multiple diagenetic eventsand canpersist into deep burial settings (>5 km) over long periods ofbasin evolution, which in the case of Proterozoic basins can extendover 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 throughtime, but during discrete episodes of pronounced diagenesis andmetamorphism. The paragenesis of diagenetic and metamorphicmineral assemblages in Proterozoic sedimentary basin-hosteduranium and PbeZn deposits demonstrate that periods of elevatedfluid flow and concomitant alteration are driven by tectonic eventsthat 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 thepunctuated recrystallization of iron oxides (Kotzer et al., 1992) anduraninite (Polito et al., 2004, 2011; Alexandre andKyser, 2005; Politoet al., 2011) as well as the precipitation of diagenetic illite (Politoet al., 2004; Alexandre et al., 2009; Hiatt et al., 2010; and Politoet al., 2011) indicate fluid/rock ratios increased during regional-scale tectonic events, which created the hydraulic gradients neces-sary for fluid flow. All successions used to interpret the nature of theGOE have been subjected to these conditions. Thus, careful assess-ment of alteration using petrographic techniques should comple-ment geochemical analyses to fully evaluate whether sedimentarysuccessions contain chemical proxies that reflect paleoenvironment.

4.2. Mineralogy, chemistry, and alteration

Diagenetic and metamorphic changes to bioelemental sedimentsare not only critical to fully understanding and interpreting thesedimentary record of the GOE, but also other important events inocean-atmosphere evolution.Metamorphicmineral transformationsin iron formation, primarily because of the detailed thermodynamicand paragenetic work of Klein (e.g. Klein, 2005; Figs. 5, 6 and 12), arethe best understood aspect of bioelemental sediment alteration.During burial, authigenic greenalite and stilpnomelane change tominnesotaite ((Fe2þ,Mg)3Si4O10(OH)2; Fig. 8; Klein, 2005), a commonalteration 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

iron formation, thus providing information regarding the paleoredoxstructure of the Precambrian ocean (Pufahl, 2010). Trace elementconcentration data can then be interpreted in terms of mineralogicalchanges, but little is known about the potential fractionation ofisotopes within systems that are increasingly employed to interpretoceanographic conditions associated with the GOE, such asd56Fe d53Cr, d97/95Mo, and d98/95Mo. There are potentially significantfractionations of these isotopes between phases such as greenalite,stilpnomelane, and minnesotaite.

Because REE ratios in iron formation are usually not significantlymodified by alteration (Derry and Jacobsen, 1990) concentrationpatterns of REE’s are potentially useful trace element proxies. Couplingthe stratigraphic correlation of metamorphosed Fe- and Si-rich faciesto their REE composition further constrains ocean redox conditions(e.g. Derry and Jacobsen, 1990), and REE patterns could providea potential method to evaluate other geochemical proxies.

The trace element composition of phosphorite is more suscep-tible to diagenesis and metamorphism because elements withinthe “open” crystal structure of francolite can be mobilized andcan fractionate (Bonnot-Courtois and Flicoteaux, 1989). Oxygenand carbon isotopes must also be used with caution. Isotopicexchange can affect d18O values from both the carbonate (d18OCO3

)and phosphate (d18OPO4

) sites, although the d18OCO3�francolite ismore vulnerable to exchange with surrounding pore waters(e.g. McArthur and Herczeg, 1990). The result of post-burialexchange of carbon isotopes is best observed when d18OCO3

andd13CCO3

values are plotted against each other. As in altered lime-stones, such a plot is constrained at one end by seawater and theother by diagenetic francolite compositions (Jarvis et al., 1994).Because of the relative ease with which isotopic exchange occurs infrancolite, caution should be exercised, especially when interpret-ing the stable isotopic composition of Precambrian phosphorite.

The effects of diagenesis and metamorphism on the d56Fe,d98/95Mo, d34S, D33S composition of authigenic phases in bio-elemental sediments are generally unknown. This is especially truefor CAS, where sulfur does not sit within structural sites of carbonateminerals (Morse andMackenzie,1990;Marenco et al., 2008). Furtherwork is also required to understand the full range of processescontrolling isotopic fractionations prior to burial. Thus, muchresearch is required before the true nature of geochemical anomalies(both spatial and stratigraphic) through the GOE can be fullyassessed.

5. Integrated approach and future research

Although the number of studies that combine sedimentologyand geochemistry to understand the GOE and Earth’s subsequentoxygenation 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; Schröderet al., 2008; Fischer and Knoll, 2009; Pecoits et al., 2009; Poultonet al., 2010; Pufahl et al., 2010), most are geochemical investiga-tions (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; Siebertet 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; Planavskyet 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 depo-sitional and post-depositional context of bioelemental sediments


(Fig. 13). Such a method mitigates the potential for incorrectlyinterpreting geochemical data because it not only provides infor-mation on how oceanography and depositional environmentinfluenced sediment chemistry, but also the effects of seawater andpore water, microbial biology, and alteration.

Central to this approach is the description of outcrop and drillcore to understand lithofacies associations and stratal architecture.This allows the construction of a sequence stratigraphic frameworkto understand the evolution of paleoenvironments and oceancurrent systems through time (Catuneanu et al., 2009). Althoughdocumenting the sequence record in the Precambrian is difficultbecause of poor preservation, especially of deeper water lithofacies,and structural deformation (Miall, 2005), it can be done (Nelsonet al., 2010). As in the Phanerozoic, attention must be given tothe identification of lithofacies stacking patterns and breaks insedimentation since each genetic unit, or systems tract, is definedby specific correlation of vertical and lateral facies trends andbounding surfaces (Catuneanu et al., 2009).

Sampling for petrography and geochemistry should belithofacies specific and interpreted in a sequence stratigraphicframework. Doing so permits the interpretation of petrographic andgeochemical data in paleoenvironmental context and provides thebackdrop for understanding the effects of post-depositional fluidflowand alteration on sediment composition (e.g. Kyser, 2007; Hiattet al., 2010). Any geochemical analyses should be superseded

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

by petrographic investigations aimed at understanding mineralparagenesis (Fig. 13). Clarification of primary and secondarytextures dictatewhat samples should be analyzed for their chemicalcomposition. Once these relationships are understood, anomalies inhigh-resolution geochemical data across individual lithofacies canbe properly assessed, elucidating any connection to alteration and ifprimary, whether they are of local or regional extent.

6. Conclusions

The GOE marks the beginning of the most significant changein Earth history, setting the stage for wholesale changes in oceanchemistry and the evolution of multicellular life. It is the utmostexpression of co-evolution between the geosphere and biosphere.The geosphere provided the chemical building blocks and ecolog-ical niches for early life, and the biosphere provided oxygen,which changed the nature of weathering, nutrient cycling, mobilityof redox sensitive elements such as iron and uranium, and inturn provided environmental stresses that pushed life along newevolutionary pathways.

Early understanding of the GOE was based on temporal trendsin bioelemental sediments, changes in mineralogy such as ironmineral abundances (hematite and magnetite in iron formation),the disappearance of detrital phases (uraninite and pyrite), andthe appearance of red beds in the continental sedimentary rockrecord. Knowledge of the GOE has been enhanced and refinedusing geochemical proxies derived from bioelemental sedimentsthat span this 100 million year interval. These proxies painta picture using broad brush-strokes that show the oxygenation ofthe atmosphereeocean system was more gradual than previouslysurmised and not a simple linear rise.

We demonstrate in this review that the fine lines necessary tofurther focus this picture can only be attained by interpretinggeochemical data in a sedimentologic and oceanographic frameworkthat incorporates an understanding of diagenetic reactions. Althoughbasin diagenetic hydrostratigraphy is rarely, if ever, consideredwheninterpreting the geochemistry of sedimentary and metamorphicrocks, it is a prominent control on diagenesis.What becomes obviousis that geochemical trends often shift along lithofacies changesand sequence stratigraphic bounding surfaces because they havecontrasting hydrologic properties. The holistic method advocatedin this review mitigates the potential for incorrectly interpretinggeochemical trends because it not only considers paleoenvironmentand oceanography, but also assesses the effects of fluid flow onalteration. Such an approach should help determine whether trendsare local, regional, or truly related to the GOE.

Surprisingly, there are currently no studies that interpret high-resolution geochemical data in a sequence stratigraphic frame-work to understand the subtle nuances of Earth’s oxygenation. Theneed tomake such connections and understand the data in their fullgeologic context is imperative as technological advances continue toincrease the rate at which geochemical data are generated. Cautionshould be exercised so that our ability to measure chemical anom-alies keeps pace with our capacity to understand them. Developinga detailed appreciation of how chemical proxies respond toalteration should be a central focus of future work. Only once thealteration of bioelemental sedimentary rocks is better understoodcan the GOE and Earth’s subsequent oxygenation history be fullyinterpreted.

Acknowledgements

This paper was improved through critical review by P.G. Erikssonand 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


for thoughtful discussions that led to this synthesis. Research wassupported byaNatural Sciences andEngineeringResearch Council ofCanada Discovery Grant and PetroCanada Young Innovator Award toPKP, and a University of Wisconsin-Oshkosh Research ProfessorshipGrant and a Faculty Development Research Grant (FDR375) to EEH.

References

Abbott, D., Isley, A., 2001. Ocean Upwelling and Mantle-plume Activity: Paleo-magnetic Tests of Ideas on the Source of the Fe in Early Precambrian IronFormations, vol. 352. Geological Society of America, Special Paper, pp. 323e339.

Aharon, P., 2005. Redox stratification and anoxia of the early Precambrian oceans:implications for carbon isotope excursions and oxidation events. PrecambrianResearch 137, 207e222.

Alexandre, P., Kyser, T.K., 2005. Effects of cationic substitutios and alteration inuraninite, and implications for the dating of uranium deposits. CanadianMineralogist 43, 1005e1017.

Alexandre, P., Kyser, K., Thomas, D., Polito, P., Marlatt, J., 2009. Geochronology ofunconformity-related uranium deposits in the Athabasca Basin, Saskatchewan,Canada and their integration in the evolution of the basin. Mineralium Deposita44, 41e59.

Algeo, T.J., Heckel, P.H., 2008. The Late Pennsylvanian Midcontinent Sea of NorthAmerica: a review. Palaeogeography, Palaeoclimatology, Palaeoecology 268,205e221.

Altschuler, Z.S.,1980. The geochemistry of tracemetals inmarine phosphorites: part 1.Characteristic abundances and enrichment. In: Bentor, Y.K. (Ed.), 1980. MarinePhosphorites. SEPM Special Publication, vol. 29. SEPM, Oklahoma, USA, pp.19e30.

Anbar, A.D., Knoll, A.H., 2002. Proterozoic ocean chemistry and evolution: a bio-inorganic bridge? Science 297, 1137e1142.

Anbar, A.D., Duan, Y., Lyons, T.W., Arnold, G.L., Kendall, B., Creaser, R.A.,Kaufman, A.J., Gordon, G.W., Scott, C., Garvin, J., Buick, R., 2007. A whiff ofoxygen before the Great Oxidation Event? Science 317, 1903e1906.

Arnold, G.L., Anbar, A.D., Barling, J., Lyons, T.W., 2004. Molybdenum isotopeevidence for widespread anoxia in Mid-Proterozoic oceans. Science 304, 87e90.

Arthur, M.A., Sageman, B.B., 1994. Marine black shales: depositional mechanismsand environments of ancient deposits. Annual Reviews of Earth and PlanetaryScience 22, 499e551.

Arning, E.T., Birgel, D., Brunner, B., Peckmann, J., 2009. Bacterial formation ofphosphatic laminites off Peru. Geobiology 7, 295e307.

Ayalon, A., Longstaffe, F.J., 1988. Oxygen isotope studies of diagenesis and pore-water evolution in the western Canada sedimentary basin: evidence from theUpper Cretaceous basal Belly River sandstone. Journal of Sedimentary Petrology58, 489e505.

Bau, M., Alexander, B., 2006. Preservation of primary REE patterns without Ceanomaly during dolomitization of Mid-Paleoproterozoic limestone and thepotential re-establishment of marine anoxia immediately after the “GreatOxidation Event”. South African Journal of Geology 109, 81e86.

Baker, P.A., Kastner, M., 1981. Constraints on the formation of sedimentary dolomite.Science 213, 215e216.

Basta, F.F., Maurice, A.E., Fontboté, L., Favarger, P., 2011. Petrology and geochemistryof banded iron formation (BIF) of Wadi Karim and Um Anab, Eastern Desert,Egypt: implications for the origin of Neoproterozoic BIF. Precambrian Research187, 277e292.

Bekker, A., Sial, A.N., Karhu, J.A., Ferreira, V.P., Noce, C.M., Kaufman, A.J.,Romano, A.W., Pimentel, M.M., 2003. Chemostratigraphy of carbonates from theMinas Supergroup, Quadrilátero Ferrífero (Iron Quadrangle), Brazil: a strati-graphic record of Early Proterozoic atmospheric, biogeochemical and climaticchange. American Journal of Science 303, 865e904.

Bekker, A., Holland, H.D., Wang, P.-L., Rumble III, D., Stein, H.J., Hannah, J.L.,Coetzee, L.L., Beukes, N.J., 2004. Dating the rise of atmospheric oxygen. Nature427, 117e120.

Bekker, A., Slack, J.F., Planavsky, N., Krape�z, Hofmann, A., Konhauser, K.O.,Rouxel, O.J., 2010. Iron formation: the sedimentary product of a complexinterplay among mantle, tectonic, oceanic, and biospheric processes. EconomicGeology 105, 467e508.

Beaumont, V., Robert, F., 1999. Nitrogen isotope ratios of kerogens in Precambriancherts: a record of the evolution of atmosphere chemistry? PrecambrianResearch 96, 63e82.

Berner, R.A., 2004. The Phanerozoic Carbon Cycle: CO2 and O2. Oxford UniversityPress, New York, 158 p.

Beukes, N.J., Klein, C., 1990. Geochemistry and sedimentology of a facies transition efrom microbanded to granular iron-formation e in the early ProterozoicTransvaal Supergroup, South Africa. Precambrian Research 47, 99e139.

Blank, C.E., Sánchez-Baracaldo, P., 2010. Timing of morphological and ecologicalinnovations in the cyanobacteria e a key to understanding the rise inatmospheric oxygen. Geobiology 8, 1e23.

Bluth, G.J.S., Kump, L.R., 1991. Phanerozoic paleogeology. American Journal ofScience 291, 284e308.

Bonnot-Courtois, C., Flicoteaux, R., 1989. Distribution of rare-earth and some traceelements in Tertiary phosphorites from the Senegal Basin and their weatheringproducts. Chemical Geology 75, 311e328.

Bottrell, S.H., Newton, R.J., 2006. Reconstruction of changes in global sulfur cyclingfrom marine sulfate isotopes. Earth-Science Reviews 75, 59e83.

Brancazio, P.J., Cameron, A.G.W., 1964. The origin and evolution of atmospheres andoceans. In: Proceedings of a Conference, Held at the Goddard Institute for SpaceStudies, NASA, New York, April 8e9, 1963. Wiley, New York.

Brocks, J.J., Buick, R., Logan, G.A., Summons, R.E., 2003. A reconstruction of Archaeanbiological diversity based on molecular fossils from the 2.78 to 2.45 billion-year-old Mount Bruce Supergroup, Hamersley Basin, Western Australia.Geochimica et Cosmochimica Acta 67, 4321e4335.

Brocks, J.J., Love, G.D., Summons, R.E., Knoll, A.H., Logan, G.A., Bowden, S.A., 2005.Biomarker evidence for green and purple sulphur bacteria in a stratifiedPalaeoproterozoic sea. Nature 437, 866e870.

Buick, R., 2008. When did oxygenic photosynthesis evolve? Philosophical Trans-actions of the Royal Society B 363, 2731e2743.

Burdett, J.W., Arthur, M.A., Richardson, M., 1989. A Neogene seawater sulphurisotope age curve from calcareous pelagic microfossils. Earth and PlanetaryScience Letters 94, 189e198.

Burdige, D.J., 2006. Geochemistry of Marine Sediments. Princeton University Press,Princeton, New Jersey, 609 p.

Byrne, R.H., Sholkovitz, E.R., 1996. Marine chemistry and geochemistry of thelanthanides. In: Gschneidner Jr., K.A., Eyring, L. (Eds.), 1996. Handbook on thePhysics and Chemistry of the Rare Earths, vol. 23. Elsevier Science B, pp. 497e593.

Canfield, D.E., 1998. A new model for Proterozoic ocean chemistry. Nature 396,450e453.

Canfield, D.E., 2001. Biogeochemistry of sulphur isotopes. Reviews in Mineralogyand Geochemistry 43, 607e636.

Canfield, D.E., 2005. The early history of atmospheric oxygen: homage to Robert M.Garrels. Annual Review of Earth and Planetary Science 33, 1e36.

Canfield, D.E., Habicht, K.S., Thamdrup, B., 2000. The Archean sulfur cycle and theearly history of atmospheric oxygen. Science 288, 658e661.

Canfield, D.E., Poulton, S.W., Narbonne, G.M., 2007. Late-Neoproterozoic deep-oceanoxygenation and the rise of animal life. Science 315, 92e95.

Canfield, D.E., Poulton, S.W., Knoll, A.H., Narbonne, G.M., Ross, G., Goldberg, T.,Strauss, H., 2008. Ferruginous conditions dominated later Neoproterozoic deep-water chemistry. Science 321, 949e952.

Catling, D.C., Zahnle, K.J., McKay, C.P., 2001. Biogenic methane, hydrogen escape, andthe irreversible oxidation of early Earth. Science 293, 839e843.

Catuneanu, O., Abreu, V., Bhattacharya, J.P., Blum, M.D., Dalrymple, R.W.,Eriksson, P.G., Fielding, C.R., Fisher, W.L., Galloway, W.E., Gibling, M.R., Giles, K.A.,Holbrook, J.M., Jordan, R., Kendall, C.G.St.C., Macurdo, B., Martinsen, O.J.,Miall, A.D., Neal, J.E., Nummedal, D., Pomar, L., Posmentier, H.W., Pratt, B.R.,Sarg, J.F., Shanley, K.W., Steel, R.J., Strasser, A., Tucker, M.E., Winker, C., 2009.Towards the standardization of sequence stratigraphy. Earth-Science Reviews 92,1e33.

Cloud, P.E., 1968. Atmospheric and hydrospheric evolution on the primitive Earth,vol. 160. Science. 729e736.

Cloud, P.E., 1973. Paleoecological significance of the banded iron-formation.Economic Geology 68, 1135e1143.

Clout, J.M., Simonson, B.M., 2005. Precambrian iron formations and iron formation-hosted ore deposits. Economic Geology, 643e679.

Condie, K.C., 2004. Supercontinents and superplume events: distinguishing signalsin the geologic record. Physics of the Earth and Planetary Interiors 146,319e332.

Condie, K.C., Des Marais, D.J., Abbott, D., 2001. Precambrian superplumes andsupercontinents: a record in black shales, carbon isotopes, and paleoclimates?Precambrian Research 106, 239e260.

Crusius, J., Thomson, J., 2000. Comparative behaviour of authigenic Re, U, and Moduring reoxidation and subsequent long-term burial in marine sediments.Geochimica et Cosmochimica Acta 64, 2233e2242.

Derry, L.A., Jacobsen, S.B., 1990. The chemical evolution of Precambrian seawater:evidence from REEs in banded in iron formation. Geochimica et CosmochimicaActa 54, 2965e2977.

Des Marais, D.J., Strauss, H., Summons, R.E., Hayes, J.M., 1992. Carbon isotopeevidence for the stepwise oxidation of the Proterozoic environment. Nature359, 605e609.

Elderfield, H., Greaves, M.J., 1982. The rare earth elements in seawater. Nature 296,214e219.

Eriksson, P.G., Reczko, B.F.F., 1998. Contourites associated with pelagic mudrocksand distal delta-fed turbidites in the Lower Proterozoic Timeball Hill Formationepeiric basin (Transvaal Supergroup), South Africa. Sedimentary Geology 120,319e335.

Eriksson, P.G., Rautenbach, C.J.de.W., Wright, D.T., Bumby, A.J., Catuneanu, O.,Mostert, P., vanderNeut, M., 2009. Possible evidence for episodic epeiric marineand fluvial sedimentation (and implications for palaeoclimatic conditions), c.2.3e1.8 Ga, Kaapvaal craton, South Africa. Palaeogeography, Palaeoclimatology,Palaeoecology 273, 153e173.

Farquhar, J., Wing, B.A., 2003. Multiple sulphur isotopes and the evolution of theatmosphere. Earth and Planetary Science Letters 213, 1e13.

Farquhar, J., Bao, H., Thiemens, M., 2000. Atmospheric influence of Earth’s earliestsulphur cycle. Science 289, 756e758.

Farquhar, J., Peters, M., Johnston, D.T., Strauss, H., Masterson, A., Wiechert, U.,Kaufman, A.J., 2007. Isotopic evidence for Mesoarchean anoxia and changingatmospheric sulphur chemistry. Nature 449, 706e709.

Fedonkin, M.A., 2009. Eukaryotization of the early biosphere: a biogeochemicalaspect. Geochemistry International 47, 1265e1333.


Filippelli, G.M., 2008. The global phosphorus cycle: past, present, and future.Elements 4, 89e95.

Fischer, W.W., Knoll, A.H., 2009. An iron shuttle for deepwater silica in Late Archeanand early Paleoproterozoic iron formation. Geological Society of AmericaBulletin 121, 222e235.

Fleet, M., 1998. Detrital pyrite in Witwatersrand gold reefs: X-ray diffractionevidence and implications for atmospheric evolution. Terra Nova 10, 302e306.

Föllmi, K.B., 1996. The phosphorus cycle, phosphogenesis and marine phosphate-rich deposits. Earth-Science Reviews 40, 55e124.

Föllmi, K.B., Garrison, R.E., Grimm, K.A., 1991. Stratification in phosphatic sediments:illustrations from the Neogene of California. In: Einsele, G., Ricken, W.,Seilacher, A. (Eds.), Cycles and Events in Stratigraphy. Springer-Verlag, Berlin,pp. 492e507.

Föllmi, K.B., Weissert, H., Lini, A., 1993. Nonlinearities in phosphogenesis and P-carbon coupling and their implications for global change. In: Wollast, R.,Mackenzie, F.T., Chou, L. (Eds.), Interactions of C, N, P, and S BiogeochemicalCycles and Global Change. NATO ASI Series. Springer-Verlag, Berlin, pp. 447e474.

Fralick, P.W., Pufahl, P.K., 2006. Iron formation in Neoarchean deltaic successionsand the microbially mediated deposition of transgressive systems tracts. Jour-nal of Sedimentary Research 76, 1057e1066.

François, L.M., Gérard, J.-C., 1986. Reducing power of ferrous iron in the Archeanocean, 1. Contribution of photosynthetic oxygen. Paleoceanography 1, 355e368.

Frei, R., Gaucher, C., Poulton, S.W., Canfield, D.E., 2009. Fluctuations in Precambrianatmospheric oxygenation recorded by chromium isotopes. Nature 461,250e253.

Froelich, P.H., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A., Heath, G.R., Cullen, D.,Dauphin, P., Hammond, D., Hartman, B., Maynard, V., 1979. Early oxidation oforganic matter in pelagic sediments of the eastern equatorial Atlantic: suboxicdiagenesis. Geochimica et Cosmochimica Acta 43, 1075e1090.

Garvin, J., Buick, R., Anbar, A.D., Arnold, G.L., Kaufman, A.J., 2009. Isotopic evidencefor an aerobic nitrogen cycle in the latest Archean. Science 323, 1045e1048.

Glenn, C.R., Arthur, M.A., 1988. Petrology and major element geochemistry of Perumargin phosphorites and associated diagenetic minerals: authigenesis inmodern organic rich sediments. Marine Geology 80, 231e268.

Glenn, C.R., Föllmi, K.B., Riggs, S.R., Baturin, G.N., Grimm, K.A., Trappe, J., Abed, A.M.,Galli-Oliver, C., Garrison, R.E., Ilyin, A.V., Jehl, C., Rohrlich, V., Sadaqah, R.M.Y.,Schidlowski, M., Sheldon, R.E., Siegmund, H., 1994. P and phosphorites: sedi-mentology and environments of formation. Eclogae Geologicae Helveticae 87,747e788.

Glenn, C.R., Prevot-Lucas, L., Lucas, J., 2000. Marine Authigenesis: From Global toMicrobial, vol. 66. SEPM, Tulsa, Oklahoma, Special Publication, 536 p.

Godderis, Y., Veizer, J., 2000. Tectonic control of chemical and isotopic compositionof ancient oceans: the impact of continental growth. American Journal ofScience 300, 434e461.

Goldblatt, C., Lenton, T.M., Watson, A.J., 2006. Bistability of atmospheric oxygen andthe Great Oxidation. Nature 443, 683e686.

Gonzalez Alvarez, I., Kerrich, R., 2010. REE and HFSE mobility due to protracted flowof basinal brines in the Mesoproterozoic Belt-Purcell Supergroup, Laurentia.Precambrian Research 177, 291e307.

Gross, G.A., 1983. Tectonic systems and the deposition of iron-formation. Precam-brian Research 20, 171e187.

Guo, Q., Strauss, H., Kaufman, A.J., Schröder, S., Gutzmer, J., Wing, B., Baker, M.A.,Bekker, A., Jin, Q., Kim, S.-T., Farquhar, J., 2009. Reconstructing Earth’s surfaceoxidation across the Archean-Proterozoic transition. Geology 37, 399e402.

Hayes, J.M., Kaplan, I.R., Wedeking, K.W., 1983. Precambrian organic geochemistry,preservation of the record. In: Schopf, J.W. (Ed.), Earth’s Earliest Biosphere:Its Origin and Evolution. Princeton University Press, Princeton, New Jersey,pp. 93e134.

Hazen, R.M., Papineau, D., Bleeker, W., Downs, R.T., Ferry, J.M., McCoy, T.J.,Sverjensky, D.A., Yang, H., 2008. Mineral evolution. American Mineralogist 93,1693e1720.

Heggie, D.T., Skyring, G.W., O’Brien, G.W., Reimers, C., Herczeg, A., Moriarty, D.J.W.,Burnett, W.C., Milnes, A.R., 1990. Organic carbon cycling and modern phos-phorite formation on the East Australia continental margin: an overview. In:Notholt, A.J.G., Jarvis, I. (Eds.), Phosphorite Research and Development. TheGeological Society of London, Oxford, pp. 147e168.

Hiatt, E.E., Budd, D.A., 2001. Sedimentary phosphate formation in warm shallowwater: new insights into the palaeoceanography of the Permian Phosphoria Seafrom analysis of phosphate oxygen isotopes. Sedimentary Geology 145,119e133.

Hiatt, E.E., Budd, D.A., 2003. Extreme paleooceanographic conditions in a Paleozoicoceanic upwelling system: organic productivity and widespread phospho-genesis in the Permian Phosphoria Sea. In: Chan, M.A., Archer, A.W. (Eds.), 2003.Extreme Depositional Environments: Mega End Members in Geologic Time, vol.370. Geological Society of America, pp. 245e264. Special Paper.

Hiatt, E.E., Kyser, T.K., 2007. Sequence stratigraphy, hydrostratigraphy, andmineralizing fluid flow in the Proterozoic Manitou Falls Formation, easternAthabasca Basin, Saskatchewan. In: Jefferson, C.W., Delaney, G. (Eds.), EXTECHIV: Geology and Uranium Exploration TECHnology of the Proterozoic Atha-basca Basin, Saskatchewan and Alberta. Geological Survey of Canada,pp. 489e506. Bulletin 588. (also Saskatchewan Geological Society SpecialPublication 18: Geological Association of Canada, Mineral Deposits Division,Special Publication 4).

Hiatt, E.E., Kyser, T.K., Fayek, M., Polito, P., Holk, G.J., Piciputi, L.R., 2007. Early quartzcements and evolution of paleohydraulic properties of basal sandstones in

three Paleoproterozoic continental basins: evidence from in situ d18O analysis ofquartz cements. Chemical Geology 238, 19e37.

Hiatt, E.E., Palmer, S.E., Kyser, T.K., O’Connor, T.K., 2010. Basin evolution, diagenesisand uranium mineralization in the Paleoproterozoic Thelon Basin, Nunavut,Canada. Basin Research 22, 302e323.

Hoashi, M., Bevacqua, D.C., Otake, T., Watanabe, Y., Hickman, A.H., Utsunomiya, S.,Ohmoto, H., 2009. Primary haematite formation in an oxygenated sea 3.46billion years ago. Nature Geoscience 2, 301e306.

Holk, G.J., Kyser, T.K., Chipley, D., Hiatt, E.E., Marlatt, J., 2003. Mobile Pb-isotopes inProterozoic sedimentary basins as guides for the exploration of uraniumdeposits. The Journal of Geochemical Exploration 80, 297e320.

Holland, H.D., 2002. Volcanic gases, black smokers, and the Great Oxidation Event.Geochimica et Cosmochimica Acta 66, 3811e3826.

Holland, H.D., 2004. The geologic history of seawater. In: Elderfield, H. (Ed.), 2004.The Oceans and Marine Geochemistry. Treatise on Geochemistry, vol. 6.Elsevier-Pergamon, Oxford, pp. 583e625.

Holland, H.D., 2006. The oxygenation of the atmosphere and oceans. PhilosophicalTransactions of the Royal Society B 361, 903e915.

Huston, D.L., Logan, G.A., 2004. Barite, BIFs and bugs: evidence for the evolutionof the Earth’s early hydrosphere. Earth and Planetary Science Letters 220,41e55.

Isley, A.E., Abbott, D.H., 1999. Plume-related mafic volcanism and the deposition ofbanded iron formation. Journal of Geophysical Research 104, 15461e15478.

Jahnke, R.A., Emerson, S.R., Roe, K.K., Burnett, W.C., 1983. The present day formationof apatite in Mexican continental margin sediments. Geochimica et Cosmo-chimica Acta 47, 259e266.

James, H.L., 1954. Sedimentary facies of iron-formation. Economic Geology 49,235e293.

Jarvis, I., Burnett, W.C., Nathan, Y., Almbaydin, F.S.M., Attia, A.K.M., Castro, L.N.,Flicoteaux, R., Hilmy, M.E., Husain, V., Qutawnah, A.A., Serjani, A., Zanin, Y.N.,1994. Phosphorite geochemistry: state-of-the-art and environmental concerns.Eclogae Geologicae Helveticae 87, 643e700.

Johnson, C.M., Bear, B.L., Roden, E.E., 2008. The iron isotope fingerprints of redoxand biogeochemical cycling in modern and ancient Earth. Annual Review ofEarth and Planetary Sciences 36, 457e493.

Johnston, D.T., Poulton, S.W., Fralick, P.W., Wing, B.A., Canfield, D.E., Farquhar, J.,2006. Evolution of the oceanic sulphur cycle at the end of the Paleoproterozoic.Geochimica et Cosmochimica Acta 70, 5723e5739.

Johnston, D.T., Wolfe-Simon, F., Pearson, A., Knoll, A.H., 2009. Anoxygenic photo-synthesis modulated Proterozoic oxygen and sustained Earth’s middle age.PNAS 106, 16925e16929.

Johnston, D.T., Poulton, S.W., Dehler, C., Porter, S., Husson, J., Canfield, D.E.,Knoll, A.H., 2010. An emerging picture of Neoproterozoic ocean chemistry:insights from the Chuar Group, Grand Canyon, USA. Earth and Planetary ScienceLetters 290, 64e73.

Kasting, J.F., 1993. Earth’s early atmosphere. Science 259, 920e925.Kasting, J.F., Liu, S.C., Donahue, T.M., 1979. Oxygen levels in the prebiological

atmosphere. Journal of Geophysical Research 84, 3097e3107.Kasting, J.F., Walker, J.C.G., 1981. Limits on oxygen concentration in the prebiological

atmosphere and the rate of abiotic fixation of nitrogen. Journal of GeophysicalResearch 86, 1147e1158.

Kasting, J.F., Siefert, J.L., 2002. Life and the evolution of Earth's atmosphere. Science296, 1066e1068.

Kasting, J.F., Holland, H.D., Pinto, J.P., 1985. Oxidant abundances in rainwater and theevolution of atmospheric oxygen. Journal of Geophysical Research 90,10497e10510.

Kasting, J.F., Pavlov, A.A., Siefert, J.L., 2001. A coupled ecosystem-climate model forpredicting the methane concentration in the Archean atmosphere. Origins ofLife and Evolution of the Biosphere 31, 271e285.

Kastner, M., 1984. Control on dolomite formation. Nature 311, 410e411.Kato, Y., Suzuki, K., Nakamura, K., Hickman, A.H., Nedachi, M., Kusakabe, M.,

Bevacqua, D.C., Ohmoto, H., 2009. Hematite formation by oxygenated ground-water more than 2.76 billion years ago. Earth and Planetary Science Letters 278,40e49.

Kendall, B., Creaser, R.A., Selby, D., 2009. 187Re e187Os Geochronology of Precam-brian Organic-rich Sedimentary Rocks. Geological Society, London, SpecialPublications, vol. 326 85e107.

Kendall, B., Gordon, G.W., Poulton, S.W., Anbar, A.D., 2011. Molybdenum isotopeconstraints on the extent of late Paleoproterozoic ocean euxinia. Earth andPlanetary Science Letters 307, 450e460.

Kholodov, V.N., Butuzova, G.Y., 2004. Problems of siderite formation and iron oreepochs: communication1. Types of siderite-bearing iron ore deposits. Lithologyand Mineral Resources 39, 389e411.

Klein, C., 2005. Some Precambrian banded iron-formation (BIFs) from around theworld: their age, geologic setting, mineralogy, metamorphism, geochemistry,and origin. American Mineralogist 90, 1473e1499.

Klein, C., Ladeira, E.A., 2000. Geochemistry and petrology of some Proterozoicbanded iron-formations of the Quadrilátero Ferrífero, Minas Gerais, Brazil.Economic Geology 95, 405e427.

Konhauser, K.O., Kappler, A., Roden, E.E., 2011. Iron in microbial metabolisms.Elements 7, 89e93.

Konhauser, K.O., Hamade, T., Raiswell, R., Morris, R.C., Ferris, F.G., Southam, G.,Canfield, D.E., 2002. Could bacteria have formed the Precambrian ironformation? Geology 30, 1079e1082.

Konhauser, K.O., Lalonde, S.V., Amskold, L., Holland, H.D., 2007. Was there really anArchean phosphate crisis? Science 315, 1234.


Konhauser, K.O., Pecoits, E., Lalonde, S.V., Papineau, D., Nisbet, E.G., Barley, M.E.,Arndt, N.T., Zahnle, K., Kamber, B.S., 2009. Oceanic nickel depletion anda methanogen famine before the Great Oxidation Event. Nature 458,750e753.

Kotzer, T.G., Kyser, T.K., Irving, E., 1992. Paleomagnetism and evolution of fluids inthe Proterozoic Athabasca Basin, northern Saskatchewan, Canada. CanadianJournal of Earth Science 29, 1474e1491.

Kump, L.R., Kasting, J.F., Barley, M.E., 2001. Rise of atmospheric oxygen and the“upside-down” Archean mantle. Geochemistry, Geophysics, Geosystems 2,1025e1035.

Kyser, T.K., 2007. Fluids, basin analysis, and mineral deposits. Geofluids 7, 238e257.Kyser, T.K., James, N.P., Bone, Y., 1998. Alteration of Cenozoic cool-water carbonates

to low-Mg calcite in marine waters, Gambier Embayment, South Australia.Journal of Sedimentary Research 68, 947e955.

Le Heron, D.P., Craig, J., Etienne, J.L., 2009. Ancient glaciations and hydrocarbonaccumulations in North Africa and the Middle East. Earth-Science Reviews 93,47e76.

Logan, G.A., Hayes, J.M., Hieshima, G.B., Summons, R.E., 1995. Terminal Proterozoicreorganization of biogeochemical cycles. Nature 376, 53e56.

Lovley, D.R., Phillips, E.J.P., Lonergan, D.J., 1991. Enzymatic versus non enzymaticmechanisms for Fe(III) reduction in aquatic sediments. Environmental Scienceand Technology 25, 1062e1067.

Lyons, T.W., Gill, B.C., 2010. Ancient sulphur cycling and oxygenation of the earlybiosphere. Elements 6, 93e99.

Lyons, T.W., Reinhard, C.T., 2009. Oxygen for heavy-metal fans. Nature 461, 179e181.Lyons, T.W., Anbar, A.B., Severmann, S., Scott, C., Gill, B.C., 2009. Tracking euxinia in

the ancient ocean: a multiproxy perspective and Proterozoic case study. AnnualReview of Earth and Planetary Sciences 37, 507e534.

Maliva, R.G., Knoll, A.H., Simonson, B.M., 2005. Secular change in the Precambriansilica cycle: insights from chert petrology. GSA Bulletin 117, 835e845.

Marenco, P.J., Corsetti, F.A., Kaufman, A.J., Bottjer, D.J., 2008. Environmental anddiagenetic variations in carbonate associated sulfate: an investigation of CAS inthe Lower Triassic of the western USA. Geochimica et Cosmochimica Acta 72,1570e1582.

McArthur, J.M., Walsh, J.N., 1984. Rare-earth geochemistry of phosphorites. Chem-ical Geology 47, 191e220.

McArthur, J.M., Herczeg, A., 1990. Diagenetic Stability of the Isotopic Composition ofPhosphate-oxygen: Palaeoenvironmental Implications, vol. 52. GeologicalSociety, London, Special Publications, pp. 119e124.

Melezhik, V.A., Fallick, A.E., Hanski, E.J., Kump, L.R., Lepland, A., Prave, A.R.,Strauss, H., 2005. Emergence of the aerobic biosphere during the Archean-Proterozoic transition: challenges of future research. GSA Today 15, 4e11.

Mel’nik, Y.P., 1982. Precambrian Banded Iron-formations, Physiochemical Condi-tions of Formation. Elsevier, New York, 310 p.

Miall, A.D., 2005. Testing for eustatic sea-level control in the Precambriansedimentary record. Sedimentary Geology 176, 9e16.

Morse, J.W., Mackenzie, F.T., 1990. Geochemistry of Sedimentary Carbonates.Elsevier, Amsterdam, Netherlands, 707 p.

Narbonne, G.M., 2010. Ocean chemistry and early animals. Science 328, 53e54.Nelson, G.J., Pufahl, P.K., Hiatt, E.E., 2010. Paleoceanographic constraints on

Precambrian phosphorite accumulation, Baraga Group, Michigan, USA. Sedi-mentary Geology 226, 9e21.

Nisbet, E.G., Grassineau, N.V., Howe, C.J., Abell, P.I., Regelous, M., Nisbet, R.E.R., 2007.The age of Rubisco: the evolution of oxygenic photosynthesis. Geobiology 5,311e335.

Ohmoto, H., Watanabe, Y., Ikemi, H., Poulson, S.R., Taylor, B.E., 2006. Sulphur isotopeevidence for an oxic Archaean atmosphere. Nature 442, 908e911.

Ohta, A., Kawabe, I., 2001. REE (III) adsorption onto Mn dioxide (d-MnO2) and Feoxyhydroxide: Ce (III) oxidation by d-MnO2. Geochimica et Cosmochimica Acta65, 695e703.

Papineau, D., 2010. Global biogeochemical changes at both ends of the Proterozoic:insights from phosphorites. Astrobiology 10, 165e181.

Papineau, D., Mojzsis, S.J., Coath, C.D., Karhu, J.A., McKeegan, K.D., 2005. Multiplesulphur isotopes of sulfides from sediments in the aftermath of Paleoproter-ozoic glaciations. Geochimica et Cosmochimica Acta 69, 5033e5060.

Papineau, D., Purohit, R., Goldberg, T., Pi, D., Shields, G.A., Bhu, H., Steele, A.,Fogel, M.L., 2009. High primary productivity and nitrogen cycling after thePaleoproterozoic phosphogenic event in the Aravalli Supergroup, India.Precambrian Research 171, 37e56.

Partridge, M.A., Golding, S.D., Baublys, K.A., Young, E., 2008. Pyrite paragenesis andmultiple sulphur isotope distribution in late Archean and early PaleoproterozoicHamersley Basin sediments. Earth and Planetary Science Letters 272, 41e49.

Pavlov, A.A., Kasting, J.F., 2002. Mass-independent fractionation of sulphur isotopesin Archean sediments: strong evidence for an anoxic Archean atmosphere.Astrobiology 2, 27e41.

Pecoits, E., Gingras, M.K., Barley, M.E., Kappler, A., Posth, N.R., Konhauser, K.O., 2009.Petrography and geochemistry of the Dales Gorge banded iron formation:paragenetic sequence, source and implications for palaeo-ocean chemistry.Precambrian Research 172, 163e187.

Pedersen, T.F., Calvert, S.E., 1990. Anoxia vs productivity; what controls theformation of organic-carbon-rich sediments and sedimentary rocks? AAPGBulletin 74, 454e466.

Petsch, S.T., Edwards, K.J., Eglinton, T.I., 2005. Microbial transformations of organicmatter in black shales and implications for global biogeochemical cycles.Palaeogeography, Palaeoclimatology, Palaeoecology 219, 157e170.

Pickard, A.L., Barley, M.E., Krape�z, B., 2004. Deep-marine depositional setting ofbanded iron formation: sedimentological evidence from interbedded clasticsedimentary rocks in the early Palaeoproterozoic Dales Gorge Member ofWestern Australia. Sedimentary Geology 170, 37e62.

Piper, D.Z., 2001. Marine chemistry of the Permian Phosphoria formation and basin,southeast Idaho. Economic Geology 96, 599e620.

Piper, D.Z., Kolodny, Y., 1987. The stable isotopic composition of a phosphoritedeposit: d13C, d34S, and d18O. Deep-Sea Research 34, 897e911.

Piper, D.Z., Calvert, S.E., 2009. A marine biogeochemical perspective on black shaledeposition. Earth-Science Reviews 95, 63e96.

Piper, D.Z., Baedecker, P.A., Crock, J.G., Burnett, W.C., Loebner, B.J., 1988. Rare earthelements in the phosphatic-enriched sediment of the Peru Shelf. MarineGeology 80, 269e285.

Planavsky, N., Rouxel, O., Bekker, A., Shapiro, R., Fralick, P., Knudsen, A., 2009. Iron-oxidizing microbial ecosystems thrived in late Paleoproterozoic redox-stratifiedoceans. Earth and Planetary Science Letters 286, 230e242.

Planavsky, N.J., Rouxel, O.J., Bekker, A., Lalonde, S.V., Konhauser, K.O., Reinhard, C.T.,Lyons, T.W., 2010. The evolution of the marine phosphate reservoir. Nature 467,1088e1090.

Polito, P.A., Kyser, T.K., Marlatt, J., Alexandre, P., Bajwah, Z., Drever, G., 2004.Significance of alteration assemblages for the origin and evolution of theProterozoic Nabarlek unconformity-related uranium deposit, Northern Terri-tory, Australia. Economic Geology 99, 113e139.

Polito, P.A., Kyser, T.K., Alexandre, P., Hiatt, E.E., Stanley, C., 2011. Advances inunderstanding the Kombolgie Subgroup and unconformity-related uraniumdeposits in the Alligator rivers uranium field and how to explore for themusing lithogeochemical principles. Australian Journal of Earth Sciences 58,453e474.

Poulton, S.W., Fralick, P.W., Canfield, D.E., 2004. The transition to a sulphidic oceanw1.84 billion years ago. Nature 431, 173e177.

Poulton, S.W., Fralick, P.W., Canfield, D.E., 2010. Spatial variability in oceanic redoxstructure 1.8 billion years ago. Nature Geoscience 3, 486e490.

Poulton, S.W., Canfield, D.E., 2011. Ferruginous conditions: a dominant feature ofthe ocean through Earth’s history. Elements 7, 107e112.

Pufahl, P.K., 2010. Bioelemental sediments. In: James, N.P., Dalrymple, R.W. (Eds.),Facies Models, fourth ed. Geological Association of Canada, pp. 477e503.

Pufahl, P.K., Fralick, P.W., 2004. Depositional controls on Palaeoproterozoic ironformation accumulation, Gogebic Range, Lake Superior region, USA. Sedimen-tology 51, 791e808.

Pufahl, P.K., Grimm, K.A., 2003. Coated phosphate grains: proxy for physical,chemical, and ecological changes in seawater. Geology 31, 801e804.

Pufahl, P.K., Hiatt, E.E., Kyser, T.K., 2010. Does the Paleoproterozoic Animikie Basinrecord sulfidic ocean transition? Geology 38, 659e662.

Raiswell, R., 1982. Pyrite texture, isotopic composition and the availability of iron.American Journal of Science 282, 1244e1263.

Raiswell, R., Reinhard, C.T., Derkowski, A., Owens, J., Bottrell, S.H., Anbar, A.D.,Lyons, T.W., 2011. Formation of syngenetic and early diagenetic iron minerals inthe late Archean Mt. McRae Shale, Hamersley Basin, Australia: new insights onthe patterns, controls and paleoenvironmental implications of authigenicmineral formation. Geochimica et Cosmochimica Acta 75, 1072e1087.

Rasmussen, B., Buick, R., 1999. Redox state of the Archean atmosphere: evidencefrom detrital minerals in ca. 3250e2750 Ma sandstones from the Pilbara Craton,Australia. Geology 27, 115e118.

Reddy, S.M., Evans, D.A.D., 2009. Palaeoproterozoic supercontinents and globalevolution: correlations from core to atmosphere. In: Reddy, S.M.,Mazumder, R., Evans, D.A.D., Collins, A.S. (Eds.), Palaeoproterozoic Supercon-tinents and Global Evolution. Geological Society, London, Special Publications,vol. 323, pp. 1e26.

Reeckman, S.A., 1981. Relationships between lithology and diagenesis in a lime-stone-marl facies transition: an electron microprobe study. Sedimentology 28,643e652.

Reinhard, C.T., Raiswell, R., Scott, C., Anbar, A.D., Lyons, T.W., 2009. A late Archeansulfidic sea stimulated by early oxidative weathering of the continents. Science326, 713e716.

Roscoe, S.M., 1969. Huronian Rocks and Uraniferous Conglomerates in the CanadianShield,GeologicalSurveyofCanadaPaper68e40.Geological SurveyofCanada.205p.

Rouxel, O.J., Bekker, A., Edwards, K.J., 2005. Iron isotopes constraints on the Archeanand Paleoproterozoic ocean redox state. Science 307, 1088e1091.

Saito, M.A., 2009. Less nickel for more oxygen. Nature 458, 714e715.Schieber, J., 2002. Sedimentary pyrite: a window into the microbial past. Geology

30, 531e534.Schröder, S., Grotzinger, J.P., 2007. Evidence for anoxia at the Ediacaran-Cambrian

boundary: the record of redox-sensitive trace elements and rare earthelements in Oman. Journal of the Geological Society London 164, 175e187.

Schröder, S., Bekker, A., Beukes, N.J., Strauss, H., van Niekerk, H.S., 2008. Rise inseawater sulphate concentration associated with the Paleoproterzoic positivecarbon isotope excursion: evidence from sulphate evaporates in thew2.2e2.1 Gyr shallow-marine Lucknow Formation, South Africa. Terra Nova 20,108e117.

Severmann, S., Anbar, A.D., 2009. Reconstructing paleoredox conditions througha multitracer approach: the key to the past is the present. Elements 5,349e364.

Shemesh, A., Kolodny, Y., Luz, B., 1988. Isotopic geochemistry of oxygen and carbonin phosphate and carbonate of phosphorite francolite. Geochimica et Cosmo-chimica Acta 52, 2565e2572.


Shen, Y., Canfield, D.E., Knoll, A.H., 2002. Middle Proterozoic ocean chemistry:evidence from the McArthur Basin, Northern Australia. American Journal ofScience 302, 81e109.

Shields, G., Stille, P., 2001. Diagenetic constraints on the use of cerium anomaliesas palaeoseawater redox proxies: an isotopic and REE study of Cambrianphosphorites. Chemical Geology 175, 29e48.

Siebert, C., Kramers, J.D., Meisel, Th., Morel, Ph, Nägler, Th.F., 2005. PGE, Re-Os, andMo isotope systematics in Archean and early Proterozoic sedimentary systemsas proxies for redox conditions of the early Earth. Geochimica et CosmochimicaActa 69, 1787e1801.

Simonson, B., 2003. Origin and evolution of large Precambrian iron formations. In:Chan, M.A., Archer, A.W. (Eds.), 2003. Extreme Depositional Environments:Mega End Members in Geologic Time, vol. 370. Geological Society of America,pp. 231e244. Special Paper.

Surdam, R.C., Crossey, L.J., 1987. Integrated diagenetic modeling: a process-orientedapproach for clastic systems. Annual Reviews of Earth and Planetary Sciences15, 141e170.

Sverjensky, D.A., Lee, N., 2010. The Great Oxidation Event and mineral diversifica-tion. Elements 6, 31e36.

Vincent, E., Berger, E.H., 1985. Carbon dioxide and polar cooling in theMiocenee theMonterey hypothesis. In: Proceedings of the Chapman Conference on NaturalVariations in CarbonDioxide and the Carbon Cycle. American Geophysical Union,pp. 455e468.

Voegelin, A.R., Nägler, T.F., Beukes, N.J., Lacassie, J.P., 2010. Molybdenum isotopes inlate Archean carbonate rocks: implications for early Earth oxygenation.Precambrian Research 182, 70e82.

Walker, T.R., 1976. Diagenetic origin of continental red beds. In: Falke, H. (Ed.), TheContinental Permian in Central, West and South Europe. D. Reidel PublishingCo., Dordrecht, pp. 240e282.

Watkins, R.T., Nathan, Y., Bremner, J.M., 1995. Rare earth elements in phosphoriteand associated sediment from the Namibian and South African continentalshelves. Marine Geology 129, 111e128.

Watson, E.B., 2008. Quiet revolution in the geochemical sciences. Elements 4,219e220.

Wilkin, R.T., Arthur, M.A., 2001. Variations in pyrite texture, sulphur isotopecomposition, and iron systematic in the Black Sea: evidence for Late Pleistoceneto Holocene excursions of the O2 and H2S redox transition. Geochimica etCosmochimica Acta 65, 1399e1416.

Wille, M., Kramers, J.D., Nägler, T.F., Beukes, N.J., Schröder, S., Meisel, Th.,Lacassie, J.P., Voegelin, A.R., 2007. Evidence for a gradual rise of oxygen between2.6 and 2.5 Ga from Mo isotopes and Re-PGE signatures in shales. Geochemicaet Cosmochimica Acta 71, 2417e2435.

Wright, D.T., Wacey, D., 2005. Precipitation of dolomite using sulphate-reducingbacteria from the Coorong Region, South Australia: significance and implica-tions. Sedimentology 52, 987e1,008.

Yang, W., Holland, H.D., Rye, R., 2002. Evidence for low or no oxygen inthe late Archean armosphere from the w2.76 Ga Mt. Roe #2 paleosol,Western Australia: Part 3. Geochimica et Cosmochimica Acta 66,3707e3718.

Zahnle, K., Claire, M., Catlin, D., 2006. The loss of mass-independent fractionation insulfur due to a Palaeoproterozoic collapse of atmospheric methane. Geobiology4, 271e283.

Oxygenation of the Earth’s atmosphere and ocean system: A review of physical and chemical...

Documents

Transcript of Oxygenation of the Earth’s atmosphere and ocean system: A review of physical and chemical...