Oxygenation of the_earth_s_atmosphere_an

Review article
Oxygenation of the Earth’s atmosphereeocean system: A review of physical
and chemical sedimentologic responses
P.K. Pufahl a,*, E.E. Hiatt b,1
a
Department of Earth & Environmental Science, Acadia University, Wolfville, Nova Scotia B4P 2R6, Canada
b
Department 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 2011
Received in revised form
30 November 2011
Accepted 5 December 2011
Available online 14 December 2011
Keywords:
Great oxidation event
Earth oxygenation
Ocean-atmosphere evolution
Bioelemental
Chemistry
Alteration
Sedimentology
Diagenesis
a b s t r a c t
The Great Oxidation Event (GOE) is one of the most significant changes in seawater and atmospheric
chemistry in Earth history. This rise in oxygen occurred between ca. 2.4 and 2.3 Ga and set the stage for
oxidative chemical weathering, wholesale changes in ocean chemistry, and the evolution of multicelluar
life. Most of what is known about this important event and the subsequent oxygenation history of the
Precambrian Earth is based on either geochemistry or “data mining” published literature to understand
the temporal abundance of bioelemental sediments. Bioelemental sediments include iron formation,
chert, and phosphorite, which are precipitates of the nutrient elements Fe, Si, and P, respectively. Because
biological processes leading to their accumulation often produce organic-rich sediment, black shale can
also be included in the bioelemental spectrum. Thus, chemistry of bioelemental sediments potentially
holds clues to the oxygenation of the Earth because they are not simply recorders of geologic processes,
but intimately involved in Earth system evolution.
Chemical proxies such as redox-sensitive trace elements (Cu, Cr, V, Cd, Mo, U, Y, Zn, and REE’s) and the
ratio of stable isotopes (d56
Fe, d53
Cr, d97/95
Mo, d98/95
Mo, d34
S, D33
S) in bioelemental sediments are now
routinely 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 essentially
anoxic conditions until the onset of the GOE at ca. 2.4 Ga. Since most of these proxies are derived from
authigenic minerals reflecting pore water composition, extrapolating the chemistry of seawater from
synsedimentary precipitates must be done cautiously.
Paleoenvironmental context is critical to understanding whether geochemical trends during Earth’s
oxygenation represent truly global, or merely local environmental conditions. To make this 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 adequately
understand them.
In this review we highlight the need for careful consideration of the role sedimentology, stratigraphy,
alteration, and basin geology play in controlling the geochemistry of bioelemental sediments. Such an
approach will fine-tune what is known about the GOE because it permits the systematic evaluation of
basin type and oceanography on geochemistry. This technique also provides information on how basin
hydrology and post-depositional fluid movement alters bioelemental sediments. Thus, a primary aim of
any investigation focused on prominent intervals of Earth history should be the integration of geochem-
istry with sedimentology and basin evolution to provide a more robust explanation of geochemical proxies
and ocean-atmosphere evolution.
Ó 2011 Elsevier Ltd. All rights reserved.
1. Introduction
One of the most intensely debated topics in the Earth sciences is
the oxygenation of the Earth’s atmosphere and oceans, primarily
because of their co-evolution with early life (e.g. Kasting, 1993;
Catling et al., 2001; Canfield, 2005; Fedonkin, 2009). Spirited
discussion began in 1964 with the publication of the “The Origin
* Corresponding author. Tel.: þ1 902 585 1858; fax: þ1 902 585 1816.
E-mail address: peir.pufahl@acadiau.ca (P.K. Pufahl).
1
Tel.: þ1 920 424 7001; fax: þ1 920 424 0240.
Contents lists available at SciVerse ScienceDirect
Marine and Petroleum Geology
journal homepage: www.elsevier.com/locate/marpetgeo
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doi:10.1016/j.marpetgeo.2011.12.002
Marine and Petroleum Geology 32 (2012) 1e20

and Evolution of Atmospheres and Oceans” (Brancazio and
Cameron, 1964). In 1973 dialogue shifted away from the notion
that purely abiotic processes produced the early atmosphere when
Cloud suggested that the deposition of large Paleoproterozoic iron
formations was linked to a rise in photosynthetic oxygen (Cloud,
1973). More recently, Kasting and Siefert (2002) summarized the
contemporary understanding of the influence of early life on the
composition of the atmosphere. “microorganisms have probably
determined the basic composition of the Earth’s atmosphere since
the origin of life.”
Holland (2002) hypothesized that the emergence of an aerobic
biosphere did not represent a simple change in the volume of volcanic
outgassing, but instead was related to a change from reducing to
oxidizing volcanic gases. Zahnle et al. (2006) and Konhauser et al.
(2009) proposed a decrease in atmospheric methane was the cata-
lyst. Although it occurred over an extended interval of time (Wille
et al., 2007; Voegelin et al., 2010), this rise in oxygen has become
known as the Great Oxidation Event (GOE; Holland, 2002, 2006) and
occurred between ca. 2.4 and 2.3 Ga (Fig. 1; Bekker et al., 2004;
Holland, 2004, 2006; Frei et al., 2009; Guo et al., 2009). It marks the
beginning of one the most significant changes the Earth has experi-
enced, setting the stage for oxidative chemical weathering, wholesale
changes in ocean chemistry, and the evolution of multicellular life
(Fig. 1).
The first evidence for the oxygenation of the atmosphere was
based on mineralogical changes with reduced detrital mineral
phases such as pyrite and uraninite in sedimentary rocks giving way
to hematite and other oxide phases (e.g. Cloud, 1968; Roscoe, 1969;
Fleet, 1998; Rasmussen and Buick, 1999; Hazen et al., 2008). Most
new data regarding the GOE, however, is geochemical in nature.
Proxies such as trace element compositions (Cu, Cr, V, Cd, Mo, U, Y,
Zn, and REE’s) and the ratio of stable isotopes (d56
Fe, d53
Cr, d97/95
Mo,
d98/95
Mo, d34
S, D33
S) in iron formation, phosphorite, and black shale
are now routinely used to indirectly deduce the redox conditions of
paleo-seawater (Fig. 2A, B, C, D; Table 1). Iron formation, phosphorite
and black shale are bioelemental sedimentary rocks that form from
the nutrient elements Fe, P, and C, which are required for myriad life
processes (Pufahl, 2010). Since the precipitation of these elements is
so closely linked to biology, bioelemental sediments are not simply
recorders of geologic processes, but are intimately involved in the
evolution of the ocean-atmosphere system (e.g. Föllmi et al., 1993;
Glenn et al., 2000; Simonson, 2003; Huston and Logan, 2004; Maliva
et al., 2005; Holland, 2006; Bekker et al., 2010; Pufahl, 2010;
Konhauser et al., 2011). Thus, their chemistry holds potentially
Figure 1. Seawater chemistry and Earth events as related to the three stages of ocean-atmosphere oxygenation (1, 2, 3). The degree of oxygenation immediately after the GOE is still
largely unknown, but recent d53
Cr data suggests that at ca. 1.9 Ga oxygen levels may have dipped to pre-GOE concentrations (Frei et al., 2009). See Figure 2 and Table 1 for a more
complete summary of the geochemical data for Earth’s oxygenation. PAL ¼ present atmospheric levels; MIF ¼ mass-independent fractionation. Based on data from Farquhar et al.
(2000), Condie et al. (2001), Canfield (2005), Fedonkin (2009), Johnston et al. (2009), Lyons and Reinhard (2009), Konhauser et al. (2011), and Nelson et al. (2010).
P.K. Pufahl, E.E. Hiatt / Marine and Petroleum Geology 32 (2012) 1e202

important clues regarding the development of the early oceans
and by extension, the atmosphere making them a logical target for
application of new geochemical techniques. This attribute, together
with the unprecedented development of technology, has spurred
the recent surge in the geochemical investigation of Precambrian
bioelemental sedimentary and meta-sedimentary rocks.
Although these technological advancements are resulting in
publication of numerous datasets, it is problematic that our ability
to measure chemical anomalies now surpasses our capacity to
adequately understand them (Watson, 2008). This problem is
exacerbated because data are often interpreted with little regard to
sedimentology, stratigraphy, alteration, and basin evolution. Such
Figure 2. Geochemical proxies used to understand Earth’s oxygenation. A) d34
S 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 d34
S values for SO2À
4 . Lower dashed line is the maximum fractionation with sulfide. B) D33
S data from sedimentary sulfides
(Farquhar et al., 2000; Farquhar and Wing, 2003). Mass independent S fractionations of 32
S, 33
S, and 34
S indicate low atmospheric oxygen levels from ca. 3.8e3.0 Ga, an increase
from ca. 2.7 to 2.4 Ga, and a permanent rise after ca. 2.4 Ga. The yellow horizontal line represents the range of values for mass dependent fractionation of S isotopes. C) d56
Fe data
from bulk shale samples, iron formations, and pyrite (Johnson et al., 2008). The yellow horizontal line marks the range in d56
Fe values for Archean to modern, low-C and low-S
clastic sedimentary rocks. Increased fractionation between ca. 2.7 and 2.5 Ga is the likely consequence of rising photosynthetic oxygen. D) d53
Cr data from iron formations
(Frei et al., 2009). The yellow horizontal line shows the range of values of magmatic Cr3þ
-rich ores and minerals formed under high temperatures. Increased fractionation between
ca. 2.8 and 2.6 Ga suggests a “whiff” or transient oxygen levels prior to the GOE. Decreased fractionation at ca. 1.9 Ga may record pre-GOE oxygen levels. E) Ni/Fe mole ratios for iron
formations (Konhauser et al., 2009). Decline in Ni at ca. 2.7 Ga may have limited methanogens and contributed to the GOE.
P.K. Pufahl, E.E. Hiatt / Marine and Petroleum Geology 32 (2012) 1e20 3

context is critical to understanding whether anomalies represent
paleoenvironmental conditions, are truly global in character, the
result of local environmental factors, or the consequence of alter-
ation of what largely are metamorphic rocks.
The sedimentary record of the GOE spans ca.100 millionyears and
provides an excellent opportunity to examine the effect of this global
geochemical revolution (cf. Watson, 2008) on interpreting major
Earth events. The picture that has emerged of the Earth’s oxygenation
is based almost exclusively on geochemistry. This approach has
provided the broad brush-strokes required to understand this
interval, but the fine lines necessary to refine this picture are only
attainable by integrating geochemical data in a sedimentologic
framework that permits the interpretation of depositional environ-
ments, oceanography, and subsequent alteration. The purpose of this
review is to summarize what is known about the GOE from the
bioelemental sedimentary record, and to re-examine the connection
between sedimentology, basin history, and the geochemical proxies
used to elucidate changes in ocean-atmosphere oxygenation.
2. The Great Oxidation Event and history of Earth’s
oxygenation
Although a great deal of controversy still exists about the
oxygenation of the Earth (compare Holland, 2004 and Hoashi et al.,
2009), there is a consistent interpretation of low Archean and
Early Paleoproterozoic atmospheric oxygen levels (<1e100 ppm O2
in the atmosphere), which are followed by higher concentrations
during the GOE that, after nearly a billion years, gave way to fully
oxygenated conditions in the latest Neoproterozoic (Fig. 1; Holland,
2004, 2006; Canfield, 2005; Canfield et al., 2007; Narbonne, 2010).
These stages are complex and multi-causal, and defined by times of
significant change in the redox state of the ocean-atmosphere
system (Huston and Logan, 2004; Canfield, 2005; Holland, 2006;
Reddy and Evans, 2009).
Support for the very low levels of oxygen prior to the GOE comes
from the presence of detrital grains composed of reduced minerals,
such as pyrite and uraninite, in sedimentary successions (e.g. Cloud,
Table 1
Geochemical proxies used to understand the Great Oxidation Event.
Proxy Environmental
parameter
Host mineral(s) Deposit type Effects of
alteration
Comments
d56
Fe Seawater
Fe-(oxyhydr)oxide
levels and effects on
bacterial DIR.
Hematite
Magnetite
Siderite
Pyrite
Iron formation
Black shale
Unknown Marked increase in fractionation between ca. 2.7 and 2.5 Ga
reflects extensive radiation in DIR resulting from increased
production of FeO. Photosynthetic oxygen likely caused the
oxidation of Fe2þ
to create FeO. Produced lower d56
Fe values.
d34
S Seawater sulfate levels
and effects on BSR.
Pyrite Black shale Unknown Marked increase in fractionation at ca. 2.4 Ga is coincident with
GOE. Interpreted to record the transition from sulfate limited to
sulfate unlimited bacterial sulfate reduction. Increased
fractionation led to more variability in d34
S values.
D33
S Absence of ozone and
UV shield and effects on
MIF of S isotopes.
Pyrite Black shale Unknown MIF of S isotopes interpreted to record absence of free oxygen
prior to the GOE. End of MIF interpreted to reflect development
of ozone layer and a UV shield associated with the GOE.
d53
Cr Seawater and
atmospheric oxygen
levels and generation of
Cr6þ
through oxic
chemical weathering.
Cr3þ
-oxides
associated
with FeO
Iron formation Unknown, but
inferred immobile.
Increase in the fractionation of Cr isotopes between ca. 2.8 and
2.6 Ga is interpreted to record a “whiff” of oxygen prior to the
GOE. Decrease in the fractionation at ca. 1.9 Ga likely records
a dip in oxygen levels to pre-GOE values. Cr6þ
is delivered to the
oceans during oxic chemical weathering and becomes immobile
when reduced by Fe2þ
to precipitate Cr3þ
oxides associated with
FeO.
d97/95
Mo, d98/95
Mo Mo oxide levels in
seawater and effects
on the fractionation
of Mo isotopes.
Mo sulfide Black shale Unknown Mo isotopic values suggest euxinic conditions prevailed after
the GOE between ca. 1.4 and 1.7 Ga. Mo is removed from
seawater by oxic adsorption processes. The isotopic
composition of these oxides is thought to be transferred to
authigenic Mo sulfides precipitated under reducing conditions
beneath the seafloor.
REEs (negative
Ce anomaly)
Seawater oxygen
concentrations
and Ce behaviour.
Ce3þ
on MneFeO. Iron formation
Phosphorite
Interpreted to
preserve a primary
signature.
Provide information on whether the water column was oxygen
stratified. Because scavenging of Ce3þ
-oxides by FeO is
negligible; Ce4þ
is scavenged on the surfaces of MneFeO
producing the negative Ce anomaly. In this way the resulting
low Ce concentration in seawater is transferred to the sediment.
Trace elements Seawater and pore
water redox recorded
by differences in the
concentrations of Cr, U,
V, Cu, Cd, Zn, Mo, and
Ni in sediment.
U oxide
Cr hydroxide
V oxide
Cu, Cd, Zn, Mo,
and Ni sulfides
Iron formation
Black shale
Phosphorite
Unknown A negative U anomaly and elevated Cr records accumulation
under suboxic and oxic conditions. Elevated U, V, Cu, Cd, Zn, Mo,
and Ni reflects deposition under anoxic conditions. Such
differences in the trace element concentrations of shallow- and
deep-water lithofacies can indicate whether the water column
was oxygen stratified.
Mo enrichment in
seawater
Mo sulfide Black shale Unknown Mo enrichment in black shale suggests a “whiff” of oxygen 50
million years prior to the GOE. Increased delivery of Mo to the
oceans via oxic chemical weathering is thought to have led to
Mo enrichment in black shales that accumulated within anoxic
environments.
Ni decline in seawater Ni adsorbed to FeO Iron formation Unknown A decline in the Ni concentration of iron formation at ca. 2.7 Ga
is interpreted to have contributed to the GOE by limiting
methanogens. Ni is a bioessential nutrient for methanogens and
without it their development was apparently limited allowing
oxygen to accumulate in the atmosphere.
Notes: Also see Figure 2. DIR ¼ dissimilatory iron reduction; FeO ¼ Fe-(oxyhydr)oxides; BSR ¼ bacterial sulfate reduction; MIF ¼ mass-independent fractionation;
UV ¼ ultraviolet light; MneFeO ¼ Mn-Fe-(oxyhydr)oxides; Corg ¼ organic matter. Although many of these proxies are inferred to be directly related to seawater composition,
because their host minerals are authigenic they in fact reflect processes that operated beneath the seafloor. Data are from Jarvis et al. (1994), Farquhar et al. (2000); Canfield
(2001), Arnold et al. (2004), Klein (2005), Rouxel et al. (2005), Anbar et al. (2007), Johnson et al. (2008), Frei et al. (2009), Bekker et al. (2004, 2010), Konhauser et al. (2009),
Lyons et al. (2009), Planavsky et al. (2009), Severmann and Anbar (2009).

1968; Roscoe, 1969; Fleet, 1998; Rasmussen and Buick, 1999;
Hazen et al., 2008), and more recently, non-mass-dependent sulfur
isotope fractionation, which provides a proxy of oxygen in the
atmosphere (Farquhar et al., 2000; Holland, 2006; Reddy and
Evans, 2009; Lyons and Gill, 2010). Analysis of Archean and early
Paleoproterozoic sedimentary sulfide and sulfate minerals has
yielded anomalous variations in the abundance of the four stable
isotopes of sulfur (32
S, 33
S, 34
S, 36
S). These anomalies are interpreted
to result from mass-independent fraction (MIF; D33
S, D36
S; Figs. 1
and 2B; Table 1) involving gaseous sulfur species in the Precam-
brian atmosphere with coeval mixing into seawater that
was marked by low sulfate concentrations (Farquhar et al., 2000;
Canfield et al., 2000). MIF is driven by photochemical reactions
involving high UV light flux. A prerequisite for these photochemical
reactions is the absence of an effective UV shield such as ozone
(Farquhar et al., 2000). Thus, the expression of MIF in sulfur
isotopes is interpreted to reflect the near absence of free oxygen in
the Archean and Early Paleoproterozoic (Farquhar et al., 2000,
2007). Although “whiffs” of oxygen are suggested in the Archean
(Ohmoto et al., 2006; Anbar et al., 2007; Wille et al., 2007; Hoashi
et al., 2009; Kato et al., 2009; Reinhard et al., 2009), the MIF of
sulfur suggests that O2 concentrations in the Archean atmosphere
were generally <10À5
of PAL (Kasting et al., 2001). Evidence from
Mo isotopes and PGE concentrations, however, suggest that oxygen
levels may have begun to rise between 2.7 and 2.5 Ga suggesting
that the increase of atmospheric oxygen that led to the demise of
MIF was not a simple linear trend (e.g. Wille et al., 2007). The end of
MIF at ca. 2.4 Ga (Figs. 1 and 2) is interpreted to record the onset of
the GOE (Bekker et al., 2004). During the following 100 million
years oxygen levels are interpreted to have risen to >10À2
PAL
(>0.2% or 2000 ppm; Pavlov and Kasting, 2002; Lyons and
Reinhard, 2009). What is not known is whether oxygen levels
through this protracted interval rose slowly or quickly, or whether
the increase was constant, marked by punctuated increases, or
some combination of these (compare Bekker et al., 2004; Ohmoto
et al., 2006; Holland, 2006; Wille et al., 2007; Anbar et al., 2007;
Lyons and Reinhard, 2009).
Prior to the advent of oxygenic photosynthesis low oxygen levels
were probably maintained in the Archean atmosphere and surface
ocean by photo-dissociation of H2O molecules (Kasting et al., 1979).
Photochemical breakdown of H2O releases H2O2, which in turn
dissociates creating O2 (Kasting et al., 1985). Kasting and Walker
(1981) determined that Archean oxygen concentrations would
have been between 10À12
and 10À14
PAL in the presence of volcanic
outgassed H2 and CO, but up to 4 Â 10À8
PAL in the absence of such
gases. Although low, these concentrations would have produced
enough O2 to cause precipitation of hematite on the continents
(Kasting and Walker, 1981).
This suggests that red beds (Fig. 3A, B, C, D) should have formed
long before the GOE (Kasting and Walker, 1981), yet the appearance
of red beds in the stratigraphic record is often cited as evidence for
the GOE (Fig. 1; Holland, 2002). The answer to this paradox lies
in how red beds form. Walker (1976) showed that red beds are
preserved during burial diagenesis in the presence of oxygenated
groundwater when Fe-(oxyhydr)oxides that coat grains (Fig. 3C)
recrystallize to form hematite (Fig. 3A, B, C). However, if ground-
water was anoxic, iron (oxyhydr)oxides (Fig. 3A) formed at the
Earth’s surface would have dissolved during early burial before
quartz cement overgrowths could precipitate and protect these
coatings, leaving no record of oxidation (e.g. Surdam and Crossey,
Figure 3. Development of red beds. A) Outcrop photo of the Eocene White River Formation, eastern Wyoming, USA. This sandstone is stained with Fe-(oxyhydr)oxides (limonite and
goethite) that form “dust rims” on detrital grains; these coatings are concentrated on slightly more permeable laminae and highlight cross bed foresets. The metastable Fe-(oxyhydr)
oxides will eventually recrystallize to form hematite making the rock red. B) Bright red hematite-stained quartz arenite and siltstone red beds from lacustrine facies of the 1.9 Ga
Roraima Group, Guyana, South America. C) Photomicrograph in plane-polarized light from the 1.7 Ga Thelon Formation, Nunavut, Canada. Detrital quartz grains (Dq) in this eolian
facies are coated with hematite dust rims (Hr) that underlie pore-filling quartz cement giving this quartz arenite a red color. D) Outcrop showing an upturned, ripple-marked bedding
plane of the Paleoproterozoic (ca. 2.3 Ga) Lorrain Formation (fluvial facies), Huronian, Blind River, Ontario. This red bed succession was one of the examples originally used as evidence
for the GOE. Photo courtesy of Steve Beyer.

1987; Lovley et al.,1991). Thus, the appearance of red beds coincides
with the GOE because groundwater became sufficiently oxygenated
to retain Fe-(oxyhydr)oxides in the shallow burial realm and
potentially preserve them in the sedimentary record.
Most researchers agree that the evolution of oxygenic photo-
synthesis within cyanobacteria was the source of oxygen that
caused the GOE (cf. Cloud, 1973). The timing of cyanobacterial
evolution, however, is problematic since biomarkers indicate they
may have evolved as early as ca. 2.9 Ga (Nisbet et al., 2007) and
were abundant by ca. 2.7 Ga (Brocks et al., 2003, 2005; Canfield,
2005; Buick, 2008), at least 400 million years before the onset of
the GOE. This long lag likely represents a period of inertia where
oxygen-consuming chemical reactions prevented the rise of
photosynthetic oxygen by consuming oxygen in inorganic reactions
with reduced mineral phases and organic matter in the oceans
(François and Gérard, 1986; Goldblatt et al., 2006; Saito, 2009).
Recent molecular clock analyses of cyanobacteria lineages by
Blank and Sánchez-Baracaldo (2010) further suggest the earliest
cyanobacteria were restricted to freshwater environments until ca.
2.4 Ga when they diversified and exploited marine ecosystems. This
diversification, extraordinary increase in habitat, and the resulting
extensive organic carbon flux to the deep oceans could have caused
a rapid increase in oxygen during the GOE. Based on this same
molecular clock model, mat-forming cyanobacteria with 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 primary
driver of oxygenation during the GOE, a number of processes are
postulated to have played a role in changing the redox state of the-
Paleoproterozoic ocean-atmosphere system. These include: (1)
increased burial of organic matter (Des Marais et al.,1992; Melezhik
et al., 2005); (2) loss of hydrogen to space from a methane-rich
atmosphere (Kasting et al., 1979; Catling et al., 2001); (3) collapse
of atmospheric methane (Zahnle et al.,2006; 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 cyanobacterialoxygen (Papineau et al.,2009); and(6) a period
of major continental growth at the Archean-Proterozoic boundary
(Godderis and Veizer, 2000). The collapse of a methane-rich atmo-
sphere is also thought to have been an important contributor to
the onset of Paleoproterozoic ice ages (Reddy and Evans, 2009).
Little is known regarding oxygen levels immediately following
the GOE (2.0e1.8 Ga; Fig. 1). Frei et al. (2009) interpret a decrease
in oxygen, possibly dropping to pre-GOE levels, based on a change
in chromium isotopic values from a small dataset (Fig. 2D; Table 1).
Because a major interval of black shale deposition corresponds
to this interval (Fig. 1), however, it is likely that any decrease
in photosynthetic oxygen production would be at least partially
compensated for by removal of organic carbon from the ocean-
atmosphere system. Thus, this interval should be explored further
to elucidate whether there was in fact a major dip in oxygen
concentration following the GOE (Fig. 1).
Oxygen levels during the Earth’s middle age (ca. 1.85e0.85 Ga)
apparently stabilized somewhere between 1 and 10% PAL (Fig. 1;
Lyons and Reinhard, 2009). Such oxygen levels are hypothesized
to have led to oxic chemical weathering of the continents, which
oxidized sulfide minerals to produce sulfate that was delivered by
rivers to the ocean. In this model, dissolved sulfate delivered to
the oceans by rivers was transformed through bacterial sulfate
reduction into sulfide causing euxinic conditions that developed at
the end of the Paleoproterozoic (Canfield, 1998, 2005; Poulton and
Canfield, 2011; Kendall et al., 2011). By ca. 1.85 the flux of sulfate
was great enough to cause sulfidic intermediate and bottom waters
(Fig.1; Poulton et al., 2004; see also Pufahl et al., 2010). Widespread
euxinia may have been perpetuated by thriving anoxygenic
photoautotrophs that tempered oxygen production by using sulfide
as an electron donor (Johnston et al., 2009). These conditions are
hypothesized to have prevailed for nearly a billion years and also
perturbed the cycling of bioessential elements, possibly causing
a long stasis in the evolution of eukaryotes (Anbar and Knoll, 2002).
This period is often referred to as the “Boring Billion” because
biological evolution is thought to have stagnated during this pro-
tracted interval (Anbar and Knoll, 2002; Holland, 2006).
Oxygen concentrations increased to >10% PAL (>0.2% or
2000 ppm) during the Neoproterozoic ‘snowball glaciations’ (Fig. 1;
Canfield, 2005; Holland, 2006). Ice cover that shrouded the Earth
between ca. 740 and 630 Ma is thought to have slowed chemical
weathering and delivery of sulfate to the oceans, causing the
demise of widespread euxinia. This set the stage for the Earth’s
transition from its prokaryote-dominated middle age by removing
sulfide, a physiological barrier to eukaryote diversification
(Johnston et al., 2010). For the first time in Earth history the
complete dominance of oxygenic photosynthesis led to the venti-
lation of the deep ocean. By ca. 580 Ma bottom waters were
oxygenated enough to stimulate the evolution of multicellular
benthic animals (Canfield et al., 2007; Narbonne, 2010). With
continued input of photosynthetic oxygen, Phanerozoic oxygen
levels were achieved by ca. 540 Ma (Holland, 2006).
3. Bioelemental sediments and the record of Earth’s
oxygenation
The sedimentary and geochemical record of the GOE is
preserved primarily in bioelemental sediments, a relatively new
classification of sedimentary rocks that encompasses iron forma-
tion, chert, and phosphorite (Pufahl, 2010). Because bioelemental
sediments are precipitated directly or indirectly by biological
processes they are often associated with organic-rich deposits such
as black shale, which can be included in the bioelemental spectrum
since it contains biologically fixed C.
The occurrence of bioelemental sediments through time reflects
changes in ocean chemistry linked to climate change, biologic
evolution, and tectonic processes (Fig. 4). These factors have
influenced the biogeochemical cycling of Fe, Si, P and C (e.g. Logan
et al., 1995) and the types of bioelemental sediments produced
before, during, and after the GOE. Thus, the temporal distribution of
bioelemental sediments provides a framework for understanding
the nature of the GOE (Fig. 4) and associated long-term changes to
ocean-atmosphere chemistry.
Also important are changes in the stacking patterns of
bioelemental lithofacies because the redox-sensitive minerals and
chemical proxies they contain provide the most detailed informa-
tion about shifts in water column oxygenation. The best records
of seawater oxygenation come from pristine lithofacies. Pristine
sedimentary facies are generally fine-grained and accumulate in
calm environments. They are characterized by undisturbed water
column precipitates and/or in situ authigenic minerals.
In a very general sense, the occurrence of bioelemental sediments
increased after the onset of the GOE and coincides with a 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). As
chert occurs in such close affinity with iron formation, phosphorite,
and black shale it is discussed in relation to these sediments.
3.1. Iron formation
Iron formation is a predominantly Precambrian, Fe-rich, marine
chemical sedimentary rock (Figs. 1, 4e6; e.g. Gross (1983); Clout and
Simonson, 2005; Klein, 2005; Bekker et al., 2010; Pufahl, 2010). The

original definition included a requirement of at least 15 wt. % Fe
(James, 1954), but later workers have found this lower limit too
restrictive (e.g. Klein, 2005). In weakly metamorphosed iron forma-
tion common minerals include the Fe-oxides hematite and magne-
tite as well as the silicates chert, greenalite, and stilpnomelane (Klein,
2005). Oxygenation of the ocean during the GOE, with either direct or
indirect involvement of Fe-oxidizing bacteria, is believed to be
responsible for deposition of all large Paleoproterozoic iron forma-
tions (Figs. 4e6; e.g. Cloud,1973; Konhauser et al., 2002). In addition
to the importance of iron formation as a recorder of oxygen levels on
the early Earth, it is economically significant because it contains most
of the world’s iron ore.
3.1.1. Temporal distribution
The Archean is characterized by pyrite and magnetite-rich deep-
water exhalative iron formation deposited in tectonically active
areas around spreading centers associated with volcanic arcs.
The dramatic rise in iron formation at ca. 2.8 Ga may correspond to
the evolution of oxygenic photosynthesis (Nisbet et al., 2007) and
resulting precipitation of ferrous Fe from the Archean ocean (Fig. 4).
Although some evidence suggests that iron formation prior to this
time was also linked to photosynthetic oxygen (e.g. Hoashi et al.,
2009), most data indicate deposition through a combination of
anoxygenic photosynthesis, dissimilatory iron reduction, oxygen
produced via nonphototrophic sources, and episodic increases in
the input of hydrothermal Fe and Si during mantle plume events
(Isley and Abbott, 1999; Konhauser et al., 2002; Pufahl, 2010;
Bekker et al., 2010).
The iron formation peak at ca. 2.5 Ga is interpreted to signal
a shift from deep-water deposition to upwelling-driven, neritic
accumulation on the expansive, unrimmed platforms that 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 shelf
environments from an oxygen-stratified ocean born during the
GOE (Pufahl, 2010). Precipitation occurred when ferrous Fe in
upwelled, anoxic waters was either mixed with photosynthetically
oxygenated seawater or oxidized during anoxygenic, bacterial
photosynthesis (Fig. 7; Cloud, 1973; Klein, 2005; Konhauser et al.,
2002; Bekker et al., 2010; Pufahl, 2010). Chert formed abiogeni-
cally primarily in subtidal environments where evaporitic
concentration (Maliva et al., 2005) and Fe-redox pumping could
saturate bottom- and pore water with silica (Fischer and Knoll,
2009; Pufahl, 2010). A suboxic seafloor was a prerequisite for Fe-
redox pumping to saturate sediment with silica. Such conditions
are interpreted to have occurred in coastal environments where
photosynthetic oxygen oases impinged on the seafloor (Nelson
et al., 2010; Pufahl, 2010). Silica was concentrated in pore water
during burial when Fe-(oxyhydr)oxides dissolve below the 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 reflect
the increased precipitation of oxidized Fe from seawater as well as
Figure 4. Temporal distribution of iron formation (red), ironstone (purple), phosphorite (yellow) and black shale (black). Based on deposit age, resource estimates and timing of
Earth events in Glenn et al. (1994), Kholodov and Butuzova (2004), Condie et al. (2001), Klein (2005), Reddy and Evans (2009), and Bekker et al. (2010). Events: OP ¼ appearance of
oxygenic photosynthesis; GOE ¼ Great Oxidation Event; BB ¼ Boring billion; CE ¼ Cambrian Explosion. Glaciations: 1 ¼ Mesoarchean; 2 ¼ Huronian; 3 ¼ Paleoproterozoic;
4 ¼ Neoproterozoic ‘Snow Ball’; 5 ¼ Ordovician; 6 ¼ Permian; 7 ¼ Neogene. Modified from Pufahl (2010).

a reduction in the delivery of Fe and Si to the ocean. As in the
Archean, peaks in iron formation abundance through the Protero-
zoic have also been correlated to mantle plume activity (Isley and
Abbott, 1999; Abbott and Isley, 2001).
Deposition of iron formation on continental margins ceased at
ca. 1.8 Ga, possibly due to the development of widespread euxinia
(Fig. 4; Canfield, 1998; Poulton et al., 2004; Kendall et al., 2011). In
a sulfidic water column dissolved sulfide is interpreted to have
combined with ferrous Fe to form pyrite, titrating the ocean of
dissolved Fe. An important change in the Precambrian Si cycle also
occurred at this time and is marked by the end of subtidal chert
deposition (Maliva et al., 2005). This change is thought to reflect
waning hydrothermal input of Si and a decrease in Si derived from
chemical weathering.
Sulfidic ocean conditions are interpreted to have continued
for nearly a billion years (Anbar and Knoll, 2002). Bioessential trace
elements were largely removed from the oceans as sulfides asso-
ciated with organic matter-rich sediments, which is thought to
have contributed to the apparent lull in eukaryote evolution (Anbar
and Knoll, 2002). The period that followed these changes is termed
the ‘Boring Billion’ because there appears to have been little change
in the atmosphereeocean and biological systems over this pro-
tracted interval of Earth history (Fig. 4). Iron formation finally
reappears coincident with the Neoproterozic ‘snowball’ glaciations
between 740 and 630 Ma (Fig. 4; Klein, 2005; Reddy and Evans,
2009; Bekker et al., 2010).
3.1.2. Deposition and chemistry
Unfortunately, there are only a few integrated studies that
couple sedimentology, mineralogy and geochemistry of bioelemental
deposits bracketing the GOE. Most of those that do focus on the
disposition and chemistry of suboxic and anoxic lithofacies forming
the large continental margin iron formations of the Paleoproterozoic
(e.g. Beukes and Klein, 1990; Klein and Ladeira, 2000; Pickard et al.,
2004; Pufahl and Fralick, 2004; Klein, 2005; Fralick and Pufahl,
2006; Fischer and Knoll, 2009; Pecoits et al., 2009). This is because
the presence of a prominent oxygen chemocline is the primary
control on facies mineralogy (Fig. 7; Pufahl, 2010).
Deposition of suboxic lithofacies occurred along segments of the
coastline where photosynthetic cyanobacteria produced oxygen
(Figs. 5A and 7). These deposits are characterized by hematite
(Fe2O3) and chert (SiO2; Fig. 8; Klein, 2005). Anoxic lithofacies
are distinguished by the presence of magnetite (Fe3O4), greenalite
((Fe2þ
, Fe3þ
)2-3Si2O5(OH)4), or stilpnomelane (K(Fe2þ
,Mg,Fe3þ
)8
(Si,Al)12(O,OH)27$nH2O; Fig. 8; Klein, 2005). All of these minerals
contain some reduced Fe, and reflect precipitation under extremely
low oxygen concentrations (ca. 10À20
pO2-water and were likely as
low as 10À70
pO2-water; Mel’nik, 1982).
Figure 5. Iron formation lithofacies. When alteration is considered the mineralogic composition can be used to infer the redox conditions of paleo-seawater/pore water.
Hematite ¼ suboxic; magnetite ¼ anoxic (Klein, 2005; Pufahl, 2010). A) Stromatolitic Paleoproterozoic Kona Dolomite, Northern Michigan, U.S.A. Dashed line highlights large
stromatolite form. The production of oxygen by such cyanobacteria was responsible for the GOE. B) Laminated magnetite. Neoarchean Eagle Island Group, northwestern Ontario,
Canada. C) Metamorphosed, laminated hematite and magnetite. Paleoproterozoic Negaunee Iron Formation, Northern Michigan, USA. D) Granular hematite-chert grainstone.
Paleoproterozoic Sokoman Formation, Labrador, Canada. E) Granular iron formation with pebble sized rip-ups of magnetite and hematite mudstone. Paleoproterozoic Sokoman
Formation, Labrador, Canada. F) Laminated magnetite and Fe-carbonate with rare magnetite mudstone intraclasts. Paleoproterozoic Sokoman Formation, Labrador, Canada.

Indirect chemical proxies such as the Fe isotopic (d56
Fe) and REE
composition of iron formation have also been used to infer
oxygenation history (Fig. 2C and 7; Table 1; e.g. Beukes and Klein,
1990; Klein, 2005; Johnson et al., 2008). The REE systematics of
redox sensitive facies, however, seems more robust and easier to
interpret, primarily because it is a direct measure of seawater Eh
(Elderfield and Greaves, 1982) without the issues of potentially
strong and not yet understood biologic fractionations (Johnson
et al., 2008). In general, iron formation and chert facies formed in
oxygenated marine environments have negative Ce anomalies and
are enriched in heavy REE’s (HoeLu) when compared to shales
(Klein, 2005). This is because although the oxidation of Ce3þ
greatly
reduces Ce solubility, oxidative scavenging on the surface of freshly
precipitated Fe-(oxyhydr)oxides removes Ce from seawater (Ohta
and Kawabe, 2001). Since the overall concentration of Ce is low,
seawater is left depleted in Ce producing a negative Ce anomaly
(e.g. Elderfield and Greaves, 1982; Piper et al., 1988). The enrich-
ment of heavy REE’s (Byrne and Sholkovitz, 1996) is also inter-
preted to be the result of preferential oxidative removal of the other
light REE’s (LaeDy) from seawater. These processes will only
produce a “seawater pattern” if deposition occurs away from a
terrigenous clastic source since siliciclastic material has no Ce
anomaly or heavy REE enrichment (Watkins et al., 1995).
Recently, the Ni concentration in iron formation has been used
to infer both the timing and cause of the GOE (Fig. 2E; Table 1;
Konhauser et al., 2009). A significant decrease in the Ni/Fe ratio
at ca. 2.7 Ga is interpreted to correlate to a major drop in the
concentration of Ni in seawater. Because of their insatiable appetite
for Ni, this change likely limited methanogens in the Neoarchean
and led to a concomitant reduction in the generation of atmo-
spheric methane. With decreasing methane and the other
environmental changes that occurred at the end of the Archean
(Des Marais et al., 1992; Godderis and Veizer, 2000; Catling et al.,
2001; Kump et al., 2001; Holland, 2002; Papineau et al., 2009)
the stage was apparently set for the accumulation of cyanobacterial
oxygen and the GOE.
3.2. Phosphorite
Phosphorite is a bioelemental sedimentary rock rich in P, is often
associated with coastal upwelling, and occurs almost exclusively in
the Phanerozoic (Figs. 1, 4 and 9). It is defined as a rock with greater
than 18 wt. % P2O5, but P2O5 can be as great as 40 wt. %, making
these rocks an important fertilizer ore (Pufahl, 2010). Most pub-
lished accounts of Proterozoic and Neoproterozoic phosphorites do
not describe true phosphorite, but phosphatic deposits that contain
much less than 18 wt. % P2O5. This distinction is important because
uncritical reporting of phosphatic occurrences has resulted in
an over estimation of Precambrian phosphorite, which has led to
errors in assessing temporal abundance and understanding the
Precambrian P cycle (e.g. Papineau, 2010).
Phosphorite forms through phosphogenesis, the authigenic
precipitation of francolite within sediment just beneath the seafloor
(Glenn et al., 1994). Francolite is a highly substituted carbonate
fluorapatite (Ca10-a-bNaaMgb(PO4)6-x(CO3)x-y-z(CO3$F)x-y-z(SO4)zF2).
Its precipitation is microbially mediated and also controlled by
the redox potential of bottom- and pore water (Jahnke et al., 1983;
Glenn et al., 1994). Authigenic, biological, and hydrodynamic
processes work together to form phosphatic laminae, in situ nodules
or reworked granular beds (Föllmi et al., 1991; Föllmi, 1996).
Phosphorite is the most important long-term sink in the
global phosphorus cycle. In the Phanerozoic the majority of P in the
oceans is sequestered in marine sediment on continental margins
and beneath regions of active coastal upwelling (Filippelli, 2008;
Fig. 10). Phosphorus is removed from nutrient-rich surface waters
by phytoplankton and authigenically converted to francolite in
accumulating organic-rich sediment through a series of microbially
mediated redox reactions (Jahnke et al., 1983; Glenn et al., 1994).
Bacterial sulfate reduction is the most efficient of these reactions
at liberating organically bound P to pore water (Arning et al., 2009).
Precipitation of francolite occurs when pore water becomes
supersaturated with respect to calcium-phosphate (Glenn et al.,
1994). Such phosphorite co-occurs with biogenic chert and black
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) Interlaminated magnetite, chert and fine-grained, hematitic grainstone laminae. Early Paleoproterozoic
Joffre Iron Formation, Western Australia. D) Laminated magnetite and chert intercalated with thin beds of hematitic grainstone. Early Paleoproterozoic Joffre Iron Formation, Western
Australia. Grainstones are interpreted as event deposits that carried granular sediment downslope from higher energy environments that were above the oxygen chemocline.

shale forming an upwelling triad of sediments. In areas not asso-
ciated with prominent upwelling the concentration of phosphate
in sediment is regulated by Fe-redox pumping (Fig.11; Heggie et al.,
1990). Preferential adsorption and release of phosphate on
Fe-(oxyhydr)oxide is kinetically favoured in Phanerozoic seawater
because it is severely silica-undersaturated (Konhauser et al., 2007).
Phosphorite 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 in
the Paleoproterozoic coincides with the GOE and the onset of oxic
chemical weathering of the continents (Papineau, 2010; Pufahl,
2010). This relatively minor phosphatic episode was not associated
with upwelling and unlike Phanerozoic phosphorites, restricted
to shallow-water environments (Nelson et al., 2010). It occurred
between 2.2 and 1.8 Ga, beginning just after the Huronian Glaciation
and in the middle of the GOE (Papineau, 2010; Pufahl, 2010).
This episode probably records an abrupt increase in the delivery of
phosphate to the oceans. Increased phosphate likely fueled a 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-glacial
continental chemical weathering under an oxygenated atmosphere
from a long period dominated by mechanical weathering during
the Huronian Glaciation (Papineau, 2010; Pufahl, 2010). Thus, the
appearance of Paleoproterozoic phosphorite is directly linked to
the GOE and the oxygenation of the oceans (Nelson et al., 2010;
Pufahl, 2010).
Phosphogenesis during this initial episode was restricted to
segments of the shoreline that were silica undersaturated and
oxygenated through microbial photosynthesis. These conditions
permitted a combination of bacterial sulfate reduction and Fe-redox
pumping to concentrate P in coastal sediment (Fig. 11; Nelson et al.,
2010). Such shallow-water phosphorite is in stark contrast to
upwelling-related, Phanerozoic phosphorites that accumulated in
a range of shelf environments. This difference likely reflects the
dissimilarity in the oxygenation state of the seafloor (Nelson et al.,
2010). The anoxia that typified Precambrian intermediate and
bottom water prevented Fe-redox pumping from operating in
deeper settings (Fig. 11).
During the onset of sulfidic ocean conditions the Fe and P cycles
became decoupled, which led to the disappearance of phosphorite
Figure 7. Continental margin iron formation. Lithofacies formed a sedimentary wedge
that fines and thickens basinward. Coastal upwelling provided a sustained supply of
anoxic bottom water rich in dissolved Fe and Si. Precipitation occurred in an oxygen
stratified water column that was suboxic down to fair-weather wave base. Nearshore
lithofacies consist of cross-stratified grainstones that are commonly stromatolitic.
Laminated pristine lithofacies accumulated in low energy environments such as
shallow lagoons and below fair-weather wave base on the middle and distal shelf. REE
spidergrams show the behaviour of Ce across the shelf. A negative Ce anomaly is most
pronounced along segments of the paleoshoreline that were oxygenated by photo-
synthesis. It disappears offshore where bottom and intermediate waters were anoxic.
The positive Eu anomaly reflects the hydrothermal source of Fe (Klein, 2005 and
references therein). SWB ¼ storm wave base; FWB ¼ fair-weather wave base. Modified
from Pufahl (2010).
Figure 8. Paragenesis typical of pristine iron formation in suboxic and anoxic paleoenvironments. Modified from Klein (2005) and Pufahl (2010).

at ca. 1.8 Ga (Fig. 4; Pufahl, 2010). The precipitation of pyrite in
a euxinic water column decreased the potential for Fe-redox
pumping, even in nearshore oxygen oases (Nelson et al., 2010).
Widespread sulfidic conditions likely made bacterial sulfate
reduction ineffective as a driver of phosphogenesis because phos-
phate would have been released to the water column where it could
be efficiently recycled and not fixed as francolite in the sediment
(Nelson et al., 2010). Phosphorite, like iron formation, was not
deposited again until the Neoproterozoic (Fig. 4).
Although rare, Proterozoic phosphorites hold great promise for
refining what is known about changes in ocean redox structure
(Melezhik et al., 2005; Pufahl, 2010), especially when coupled with
the sedimentology and chemistry of co-occurring bioelemental
sediments. Francolite readily incorporates a variety of redox
sensitive trace elements into its crystal structure and thus, often
preserves a record of pore water and bottom water Eh during
deposition (Fig. 10; Jarvis et al., 1994). Trace elements generally
replace Ca2þ
in francolite, but can also be transferred to the sedi-
ment by absorption onto crystal surfaces, scavenging by organic
matter, or substitution in sulfides (Jarvis et al., 1994).
Enriched elements include Ag, Cu, Cr, V, Cd, Mo, Se, U, Y, and Zn,
and the REEs La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, and Lu
(e.g McArthur and Walsh 1984; Altschuler, 1980; Hiatt and Budd,
2003; Fig. 10). The most commonly used to infer redox conditions
are Cu, Cr, V, Cd, Mo, U, and Zn (Fig. 10). All but U are mobilized and
incorporated under reducing conditions. As in iron formation, the
REE content of francolite records seawater values although it
continues to absorb REE from pore water below the sedimente-
water interface (Altschuler, 1980; Piper et al., 1988). As in iron
formation, 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 composition
of francolite can be used to understand the microbial processes
releasing phosphate to pore water (d13
CCO3
) and to determine
precipitation temperature (d18
OCO3
; d18
OPO4
; Piper and Kolodny,
1987; Shemesh et al., 1988; Hiatt and Budd, 2001). Temperature
calculations are sometimes coupled with trace element analysis to
infer the redox conditions and paleooceanography of ancient seas
(e.g. Hiatt and Budd, 2003).
Figure 9. Precambrian phosphorite lithofacies. Most Precambrian phosphorites are unlike Phanerozoic phosphatic deposits because they do not form aerially extensive deposits. They
generally consist of thin pristine phosphorite inperitidal environments and granular phosphatic lags in shallow-waterlithofacies. A) Laminated pristine phosphorite. Subhedral crystals are
pyrite, black blebs are organic matter and the honey-brown mineral between organic-rich laminae is francolite. B) Francolite peloids (brown) with greenalite cement (acicular crystals)
surrounded by dolomite. Opaque square-shape is pyrite. Paleoproterozoic Ruth Formation, Labrador, Canada. Authigenic glauconite is commonly associated with such francolite grains
indicating phosphogenesis along a suboxic seafloor (Pufahl, 2010). C) Phosphatic peloids on bedding surfaces (arrows) in cross-laminated chert. Paleoproterozoic Bijiki Iron Formation,
northern Michigan, U.S.A. D) Francolite peloid (brown) cemented with ankerite. Paleoproterozoic Bijiki Iron Formation, northern Michigan, U.S.A.
Figure 10. Continental margin phosphorite and black shale. Phosphorite accumulates
within organic-rich sediment beneath the sites of coastal upwelling. A pronounced
oxygen minimum zone (OMZ) develops as benthic bacteria exhaust oxygen to degrade
organic matter. Black shale is also associated with upwelling, but can form in calm,
nutrient-rich coastal environments such as lagoons. The plots show redox-related
changes in trace element concentrations across the shelf. In the nearshore a 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 portions
of the shelf. SWB ¼ storm wave base; FWB ¼ fair-weather wave base. Modified from
Pufahl (2010).

3.3. Black shale
Black shale is a dark, thinly laminated, carbonaceous fine-
grained clastic sedimentary rock (Fig. 12) that is rich in organic
matter (>2 wt. %), sulfides (especially pyrite), and redox sensitive
trace elements (U, V, Cu, and Ni; Arthur and Sageman, 1994; Piper
and Calvert, 2009). It can form in a wide range of paleoenviron-
ments from peritidal to deep basin settings, is often associated with
phosphorite, and can be a hydrocarbon source rock (Fig. 10).
Black shales are commonly interpreted as recording deposition
beneath a highly productive surface ocean or within anoxic, sulfidic
bottom waters, or a combination of both (Piper and Calvert, 2009).
Recent work, however, suggests that high planktic productivity
is the most important control on organic matter enrichment
in marine sediment (e.g. Piper and Calvert, 2009 and references
therein). Organic matter accumulates because the rate of produc-
tion and settling is greater than the rate of degradation of organic
carbon on the seafloor (Pedersen and Calvert, 1990). Since
processes of black shale deposition can occur across the spectrum
of shelf environments, their occurrence is not always an indication
of accumulation in a deep, open ocean basin.
Processes leading to the formation of black shale are important
because they link the various pools of carbon in the ocean-
atmosphere system (Arthur and Sageman, 1994). These processes
govern carbon burial, which regulates climate, and oxygen levels by
controlling the rate reduced C is sequestered in the geologic record
(Holland, 2002; Canfield, 2005). Since P is the primary control
on productivity over geologic time scales the phosphorus cycle
ultimately determines the rate of organic matter burial and
removal of carbon dioxide from the atmosphere.
Figure 11. Extent of phosphogenesis resulting from Fe-redox pumping on Precambrian and Phanerozoic shelves. As Fe-(oxyhydr)oxides are buried beneath the Fe-redox boundary
they dissolve, liberating sorbed HPO4
2-
to pore water. Francolite precipitation is limited in the sediment by the availability of seawater-derived FÀ
. Although important in stimulating
phosphogenesis in the Phanerozoic, bacterial sulfate reduction was likely much less efficient at promoting the precipitation of francolite in the Precambrian because of the very low
seawater sulfate levels. Thus, the difference in the size of phosphogenic regions in the Precambrian and Phanerozoic is interpreted to the consequence of the disparity in the
oxygenation state of the seafloor. In the Precambrian, photosynthetically oxygenated nearshore environments possessed suboxic seafloors that facilitated Fe-redox pumping and
phosphogenesis. Phosphogenesis could not occur in the middle and distal shelf because these regions were below the oxygen chemocline. Phosphogenesis in the Phanerozoic
occurs across the entire shelf because the seafloor is generally well oxygenated. SWB ¼ storm wave base; FWB ¼ fair-weather wave base. Modified from Nelson et al. (2010).

The temporal distribution of black shale is even less well con-
strained than that of phosphorite (Figs. 1 and 4), primarily because
it forms components of other depositional systems. The record of
Precambrian black shale is also severely biased given the rarity of
preserved deep-sea sediments, and because they are easily eroded.
In general, however, the timing of black shale deposition reflects
periods when oxygen concentrations could increase in the atmos-
phereeocean system (Fig. 4; e.g. Berner, 2004). Secular changes
in black shale deposition result from changes in carbon cycling in
“active” surface ocean pools, in the atmosphere, on land, and in
marine sediment, and carbon pools that cycle on much longer
timescales (Burdige, 2006).
Such changes are partly linked to the GOE (Des Marais et al.,
1992), which apparently follows an episode of enhanced carbon
burial in the late Archean (Fig. 4). This is the first of three noticeable
peaks in black shale deposition during the Precambrian (Fig. 4;
Condie et al., 2001). It is the least prominent and occurs in the
Neoarchean between ca. 2.7 and 2.5 Ga. This initial pulse of black
shale accumulation is thought to correspond to either a mantle
plume event, which through climate warming, increased chemical
weathering and nutrient delivery to the oceans (Condie, 2004),
or a change in ocean currents (Condie et al., 2001) that resulted
in initiation of upwelling along favorably positioned cratons.
The sequestration of reducing organic matter during this episode is
interpreted to have contributed to the GOE (Des Marais et al.,1992).
The second pulse is a prominent event occurring between ca. 2.0
and 1.7 Ga (Fig. 4; Condie et al., 2001), just after the Huronian
Glaciation. As with iron formation of this age, the accumulation of
black shale is also correlated to mantle volcanism (Condie et al.,
2001). Intense chemical weathering of post-glacial landscapes
(Papineau, 2010; Pufahl, 2010) is interpreted to have increased P
fluxes to the ocean that not only stimulated primary production
and phosphogenesis (Nelson et al., 2010), but also black shale
deposition as well.
Black shale again becomes conspicuous in the Cryogenian
between ca. 800e600 Ma (Fig. 4). Organic-rich mudstones, some of
which are phosphatic, accumulated during retreat of the “snow-
ball” glaciations (Condie et al., 2001; Le Heron et al., 2009). Elevated
surface ocean productivities were likely sustained by delivery of
nutrients through glacial runoff and invigorated coastal upwelling
(Papineau, 2010). The pronounced equator-to-pole temperature
gradient that develops during glaciations leads to more energetic
atmospheric circulation and thus, coastal upwelling, resulting in
the widespread accumulation of organic-rich sediment (Vincent
and Berger, 1985).
Correspondence between peaks of black shale and those of iron
formation deposition in the Precambrian (Fig. 4) highlights the
importance that photosynthetic oxygen played in the accumulation
of iron formation. This relationship also emphasizes the connection
between ocean circulation and upwelling to deliver reduced iron
and P to the photic zone.
Like phosphorite, pulses of black shale deposition in the Phan-
erozoic are linked to ocean-climate feedback (Bluth and Kump,1991;
Arthur and Sageman, 1994). Prominent peaks are also the conse-
quence of enhanced P burial from invigorated coastal upwelling or
increased chemical weathering and delivery of phosphate to the
oceans (Fig. 4; Glenn et al., 1994; Föllmi, 1996).
Much information about fluctuations in seawater Eh is derived
from paragenetic studies of black shale-hosted, authigenic minerals
(Glenn and Arthur, 1988; Arthur and Sageman, 1994; Pufahl and
Grimm, 2003; Raiswell et al., 2011). Textural relationships between
glauconite, pyrite, francolite, and carbonate provide a high fidelity
record of the physical, chemical, and biologic processes causing
subtle shifts in redox potential (Glenn and Arthur, 1988; Pufahl
and Grimm, 2003). These minerals precipitate through a series of
microbially mediated redox reactions (Froelich et al., 1979; Glenn
et al., 1994). In order of decreasing energy yield these reactions
include oxic respiration, denitrification, transition metal oxide
reduction, sulfate reduction, and methanogenesis. Geochemical
evidence suggests that all but oxic respiration evolved by the late
Archean (Garvin et al., 2009; Lyons and Gill, 2010), and aerobic
Figure 12. Black shale. A) Pristine phosphorite associated with black shale of the Permian Meade Peak Member (M), which is overlain by the Rex Chert Member. Permian
Phosphoria Formation, Wyoming, U.S.A. From Pufahl (2010). B and C) Black shale from the Marra Mamba Iron Formation, Western Australia. Drill core WRL-1. Yellowish staining in
(B) is from weathered sedimentary sulfides. Organic-rich laminae in (C) are commonly scoured by very fine-grained, thinly bedded sandstone layers. Dashed line highlights a scour
surface. Arrows denote younging direction. D) Black shale from Joffre Iron Formation. Drill core SPD-50. Minute light-coloured specks within certain laminae are pyrite crystals.

heterotrophs evolved by ca. 2.1 Ga in response to increasing oxygen
levels (Papineau et al., 2005). Because of widespread ocean anoxia
bacteriaedriven reactions that produce and consume organic matter
were not confined to below the seafloor, but also occurred within the
water column.
Pyrite precipitates below the sulfate redox interface through the
conversion of monosulfides formed during bacterial sulfate
reduction (Schieber, 2002). Sedimentary pyrite is often framboidal
and finely disseminated (Raiswell, 1982; Wilkin and Arthur, 2001;
Schieber, 2002), but discrete layers have been interpreted as
recording precipitation and suspension settling through a euxinic
water column (Poulton et al., 2004). Francolite precipitates in
association with the microbial reduction of nitrate, Mn-oxides,
Fe-oxides, and sulfate (Pufahl, 2010). Unlike the formation of
pyrite, however, phosphogenesis is not a redox-controlled reaction,
but is regulated only by the concentration of phosphate in pore
water (Glenn et al., 1994).
In addition to pyrite and francolite, other authigenic
minerals that may precipitate in black shales include glauconite
((K,Na,Ca)1.2-2.0(Fe3þ
,Al,Fe2þ
,Mg)4.0[Si7-7.6Al1.0-0.4O20](OH)4$n(H2O)),
calcite (CaCO3), dolomite (CaMg(CO3)2, and siderite (FeCO3), all of
which can be used to further constrain pore water Eh. Glauconite
forms first, within suboxic pore water at the Fe-redox interface,
followed by pyrite, and then carbonate at progressively deeper
levels in the sediment. Glauconite occurs as authigenic peloids,
coatings, or cement. Carbonate minerals precipitate within the
alkalinity maximum that develops during intense microbial respi-
ration. The type of carbonate mineral produced depends on the
availability of Ca2þ
, Fe2þ
and Mg2þ
. Calcite forms from pore water
with little Fe2þ
and Mg2þ
, whereas siderite precipitates in anoxic
pore water with abundant Fe2þ
(François and Gérard, 1986; Klein,
2005). Dolomite is created in pore water enriched in Ca2þ
and
Mg2þ
when the bacterial reduction of SO2À
4 , a kinetic inhibitor to
dolomite precipitation, is converted to H2S (Baker and Kastner,
1981; Kastner, 1984; Wright and Wacey, 2005). These carbonate
minerals are generally a microcrystalline cement that binds detrital
grains and earlier authigenic phases together, but can also form
displacive concretionary horizons (Kholodov and Butuzova, 2004).
The bulk trace element content of black shales and the sulfur
isotopic composition of co-occurring pyrite provide evidence of
changing seawater Eh over longer timescales. An increase in the
concentration of redox-sensitive trace elements and an increase in
the fractionation of d34
S roughly correspond to the onset of the GOE
(Fig. 2A; Table 1). The restricted range of d34
S values in pyrite prior to
ca. 2.5 Ga is interpreted to reflect low seawater SO2À
4 concentrations
of the Archean; the consequence of negligible SO2À
4 delivery from
anoxic chemical weathering (Canfield, 2001). Low SO2À
4 levels
restrict bacterial sulfate reduction and produce little variation in d34
S
values (Canfield, 2001). After ca. 2.5 Ga fractionations increase
dramatically to values expected for bacterial sulfate reduction,
which is not limited by low sulfate concentrations. Higher sulfate
levels were produced by weathering of pyrite under an oxygenated
atmosphere (Canfield, 2001).
MIF of sulfur isotopes (D33
S; Figs. 1 and 2B; Table 1) in pyrite
provides the best evidence of Precambrian atmospheric oxygen
levels and the timing of the GOE (Farquhar et al., 2000; Farquhar
and Wing, 2003). The nature of its onset is preserved in the
record of multiple sulfur isotope distributions (d34
S-D33
S), which
suggests that oxygen levels began to fluctuate ca. 150 million years
prior to the permanent rise at ca. 2.4 Ga (Partridge et al., 2008; cf.
Wille et al., 2007). The decreased variability and appearance of
positive pyrite d56
Fe values after ca. 2.3 Ga corroborate D33
S data
(Fig, 2; Rouxel et al., 2005), but it is unclear whether these changes
reflect seawater composition or diagenesis (Johnson et al., 2008).
Molybdenum isotopes from FeeMoeS precipitates in black shale
provide further clues (Lyons et al., 2009; Severmann and Anbar,
2009; Voegelin et al., 2010). d98/95
Mo values corroborate the rise
in oxygen levels ca. 150 million years before the accepted onset of
the GOE (Wille et al., 2007; Voegelin et al., 2010). Molybdenum
isotope data also suggest that although euxinic conditions may
have eventually developed in the late Paleoproterozoic, the ocean
was probably not a “global Black Sea” (Lyons et al., 2009).
Another approach that is increasingly being used is the analysis
of carbonate-associated-sulfur (CAS; e.g. Guo et al., 2009). Because
CAS can acquire the isotopic composition of pore water and
seawater (Burdett et al., 1989) it is particularly attractive as a pale-
oredox proxy in Precambrian limestones. Sulfur isotope data from
associated pyrite also allows potential calculation of the offset
between SO2À
4 and H2S during bacterial sulfate reduction (Lyons
and Gill, 2010), further constraining the nature of redox sensitive
microbial processes in the sediment and water column.
4. Reading the record of Earth’s oxygenation:
diagenetic and metamorphic effects
Diagenesis and metamorphism can significantly alter sediment
chemistry, especially in deposits as old as the Precambrian (e.g.
Hayes et al., 1983; Ayalon and Longstaffe, 1988; Crusius and
Thomson, 2000; Shields and Stille, 2001; Petsch et al., 2005;
Gonzalez Alvarez and Kerrich, 2010; Hiatt et al., 2010). Thus, it is
difficult to reconcile why so few studies of Precambrian depositional
systems attempt to understand alteration of what are primarily
metamorphic rocks and minerals before interpreting geochemical
data. This is especially true since techniques exist to assess whether
observed anomalies reflect conditions at the time of deposition,
alteration, or a combination of both (e.g. Kendall et al., 2009).
Unfortunately, technological breakthroughs that have allowed the
rapid analysis of samples (Watson, 2008) also lead to the brisk
publication of data without full assessment of potential alteration.
4.1. Sedimentology, basin evolution, and alteration
The current level of understanding bioelemental sediment
alteration is at about the same stage as understanding limestone
diagenesis was 30 years ago. A basic tenet that has important
implications for understanding geochemical proxies is that the most
desirable deposits for geochemical analysis are pristine lithofacies.
Like lime mudstones, fine-grained bioelemental facies usually
represent seawater and authigenic precipitates with low hydraulic
conductivities that tend to fix pore water chemistry close to the
time of deposition (e.g. Kyser et al., 1998). This fact is commonly
overlooked when extrapolating the chemistry of authigenic minerals
to the overlying water column. Coarser facies have higher fluid/rock
ratios and experience greater fluid fluxes during burial that often
results in chemical compositions that are significantly different from
original ones (e.g. Reeckman, 1981).
Sedimentologic and basin evolution context are both paramount
when interpreting geochemical data. A properly constrained depo-
sitional and post-depositional framework provides information
on how oceanography, depositional environment, seawater and
pore water chemistry, microbial biology, and alteration influence
the chemical composition of bioelemental sediments. Without
this perspective it is a challenge to interpret whether geochemical
anomalies through the GOE are the consequence of local environ-
mental factors or global in character (Lyons et al., 2009; Pufahl
et al., 2010). In most cases ascertaining the nature of anomalies in
Precambrian sedimentary rocks is especially difficult given the
rarity of preserved deep-sea sediments (Pufahl et al., 2010). Unfor-
tunately, this issue is often overlooked resulting in conclusions that
are not fully supported by sedimentologic data. For example, key

stratigraphic units in the Paleoproterozoic Pretoria Group, where
much of the geochemical data is derived (e.g. Bekker et al., 2004;
Bau and Alexander, 2006), are interpreted as epeiric sea deposits
(Eriksson and Reczko, 1998; Eriksson et al., 2009). Sedimentologic
evidence suggests that epeiric seas are typified by restricted circu-
lation patterns that produce water masses with compositions that
differ substantially from the open ocean (e.g. Hiatt and Budd, 2001;
Piper, 2001; Algeo and Heckel, 2008).
In all sedimentary basins diagenetic hydrostratigraphy is
controlled primarily by lithofacies and resulting diagenetic 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 and
grain size of lithofacies (e.g. Hiatt and Budd, 2003). Widespread
lateral flow of diagenetic fluids can occur over distances of hundreds
of kilometers and at temperatures >200 C (Hiatt and Kyser, 2007;
Kyser, 2007; Alexandre et al., 2009; Hiatt et al., 2010). Diagenetic
fluids flow most intensively within coarse-grained facies that
have not experienced intense cementation and are situated above
unconformities or along internal discordances such as parasequence
boundaries (Hiatt et al., 2003; Hiatt et al., 2007; Hiatt and Kyser,
2007). Stratigraphic boundaries often allow preferential diagenetic
and metamorphic fluid flow that can reset relatively robust
geochemical proxies, such as d34
S (Pufahl et al., 2010). This long-term
diagenetic hydrostratigraphy can involve multiple diagenetic events
and can persist into deep burial settings (5 km) over long periods of
basin evolution, which in the case of Proterozoic basins can extend
over 500 Ma (Holk et al., 2003; Hiatt et al., 2007, 2010; Kyser, 2007;
Alexandre et al., 2009; Hiatt et al., 2010).
Such heterogeneous alteration does not occur gradually through
time, but during discrete episodes of pronounced diagenesis and
metamorphism. The paragenesis of diagenetic and metamorphic
mineral assemblages in Proterozoic sedimentary basin-hosted
uranium and PbeZn deposits demonstrate that periods of elevated
fluid flow and concomitant alteration are driven by tectonic events
that changed basin hydrology (Kotzer et al., 1992; Polito et al., 2004,
2011; Alexandre and Kyser, 2005; Kyser, 2007; Alexandre et al.,
2009; Hiatt et al., 2010; Polito et al., 2011). In these systems the
punctuated recrystallization of iron oxides (Kotzer et al., 1992) and
uraninite (Polito et al., 2004, 2011; Alexandre and Kyser, 2005; Polito
et al., 2011) as well as the precipitation of diagenetic illite (Polito
et al., 2004; Alexandre et al., 2009; Hiatt et al., 2010; and Polito
et al., 2011) indicate fluid/rock ratios increased during regional-
scale tectonic events, which created the hydraulic gradients neces-
sary for fluid flow. All successions used to interpret the nature of the
GOE have been subjected to these conditions. Thus, careful assess-
ment of alteration using petrographic techniques should comple-
ment geochemical analyses to fully evaluate whether sedimentary
successions contain chemical proxies that reflect paleoenvironment.
4.2. Mineralogy, chemistry, and alteration
Diagenetic and metamorphic changes to bioelemental sediments
are not only critical to fully understanding and interpreting the
sedimentary record of the GOE, but also other important events in
ocean-atmosphere evolution. Metamorphic mineral transformations
in iron formation, primarily because of the detailed thermodynamic
and paragenetic work of Klein (e.g. Klein, 2005; Figs. 5, 6 and 12), are
the best understood aspect of bioelemental sediment alteration.
During burial, authigenic greenalite and stilpnomelane change to
minnesotaite ((Fe2þ
,Mg)3Si4O10(OH)2; Fig. 8; Klein, 2005), a common
alteration mineral. With increasing metamorphic grade amphiboles,
pyroxenes, and fayalite are high-temperature reaction products.
These relationships can be used to infer the original mineralogies of
iron formation, thus providing information regarding the paleoredox
structure of the Precambrian ocean (Pufahl, 2010). Trace element
concentration data can then be interpreted in terms of mineralogical
changes, but little is known about the potential fractionation of
isotopes within systems that are increasingly employed to interpret
oceanographic conditions associated with the GOE, such as
d56
Fe d53
Cr, d97/95
Mo, and d98/95
Mo. There are potentially significant
fractionations of these isotopes between phases such as greenalite,
stilpnomelane, and minnesotaite.
Because REE ratios in iron formation are usually not significantly
modified by alteration (Derry and Jacobsen, 1990) concentration
patterns of REE’s are potentially useful trace element proxies. Coupling
the stratigraphic correlation of metamorphosed Fe- and Si-rich facies
to their REE composition further constrains ocean redox conditions
(e.g. Derry and Jacobsen, 1990), and REE patterns could provide
a potential method to evaluate other geochemical proxies.
The trace element composition of phosphorite is more suscep-
tible to diagenesis and metamorphism because elements within
the “open” crystal structure of francolite can be mobilized and
can fractionate (Bonnot-Courtois and Flicoteaux, 1989). Oxygen
and carbon isotopes must also be used with caution. Isotopic
exchange can affect d18
O values from both the carbonate (d18
OCO3
)
and phosphate (d18
OPO4
) sites, although the d18
OCO3Àfrancolite is
more vulnerable to exchange with surrounding pore waters
(e.g. McArthur and Herczeg, 1990). The result of post-burial
exchange of carbon isotopes is best observed when d18
OCO3
and
d13
CCO3
values are plotted against each other. As in altered lime-
stones, such a plot is constrained at one end by seawater and the
other by diagenetic francolite compositions (Jarvis et al., 1994).
Because of the relative ease with which isotopic exchange occurs in
francolite, caution should be exercised, especially when interpret-
ing the stable isotopic composition of Precambrian phosphorite.
The effects of diagenesis and metamorphism on the d56
Fe,
d98/95
Mo, d34
S, D33
S composition of authigenic phases in bio-
elemental sediments are generally unknown. This is especially true
for CAS, where sulfur does not sit within structural sites of carbonate
minerals (Morse and Mackenzie,1990; Marenco et al., 2008). Further
work is also required to understand the full range of processes
controlling isotopic fractionations prior to burial. Thus, much
research is required before the true nature of geochemical anomalies
(both spatial and stratigraphic) through the GOE can be fully
assessed.
5. Integrated approach and future research
Although the number of studies that combine sedimentology
and geochemistry to understand the GOE and Earth’s subsequent
oxygenation has increased in recent years (e.g. Beukes and Klein,
1990; Klein and Ladeira, 2000; Pickard et al., 2004; Klein, 2005;
Fralick and Pufahl, 2006; Schröder and Grotzinger, 2007; Schröder
et al., 2008; Fischer and Knoll, 2009; Pecoits et al., 2009; Poulton
et al., 2010; Pufahl et al., 2010), most are geochemical 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; Siebert
et al., 2005; Johnston et al., 2006, 2009; Bottrell and Newton,
2006; Bau and Alexander, 2006; Frei et al., 2009; Guo et al.,
2009; Johnson et al., 2008; Kendall et al., 2009; Konhauser et al.,
2009; Lyons and Reinhard, 2009; Lyons et al., 2009; Planavsky
et al., 2009, 2010; Severmann and Anbar, 2009; Lyons and Gill,
2010; Papineau, 2010; Voegelin et al., 2010; Basta et al., 2011).
None use a fully integrated approach incorporating sedimentology,
stratigraphy, alteration, and basin analysis to constrain the depo-
sitional and post-depositional context of bioelemental sediments

(Fig. 13). Such a method mitigates the potential for incorrectly
interpreting geochemical data because it not only provides infor-
mation on how oceanography and depositional environment
influenced sediment chemistry, but also the effects of seawater and
pore water, microbial biology, and alteration.
Central to this approach is the description of outcrop and drill
core to understand lithofacies associations and stratal architecture.
This allows the construction of a sequence stratigraphic framework
to understand the evolution of paleoenvironments and ocean
current systems through time (Catuneanu et al., 2009). Although
documenting the sequence record in the Precambrian is difficult
because of poor preservation, especially of deeper water lithofacies,
and structural deformation (Miall, 2005), it can be done (Nelson
et al., 2010). As in the Phanerozoic, attention must be given to
the identification of lithofacies stacking patterns and breaks in
sedimentation since each genetic unit, or systems tract, is defined
by specific correlation of vertical and lateral facies trends and
bounding surfaces (Catuneanu et al., 2009).
Sampling for petrography and geochemistry should be
lithofacies specific and interpreted in a sequence stratigraphic
framework. Doing so permits the interpretation of petrographic and
geochemical data in paleoenvironmental context and provides the
backdrop for understanding the effects of post-depositional fluid
flow and alteration on sediment composition (e.g. Kyser, 2007; Hiatt
et al., 2010). Any geochemical analyses should be superseded
by petrographic investigations aimed at understanding mineral
paragenesis (Fig. 13). Clarification of primary and secondary
textures dictate what samples should be analyzed for their chemical
composition. Once these relationships are understood, anomalies in
high-resolution geochemical data across individual lithofacies can
be properly assessed, elucidating any connection to alteration and if
primary, whether they are of local or regional extent.
6. Conclusions
The GOE marks the beginning of the most significant change
in Earth history, setting the stage for wholesale changes in ocean
chemistry and the evolution of multicellular life. It is the utmost
expression of co-evolution between the geosphere and biosphere.
The geosphere provided the chemical building blocks and ecolog-
ical niches for early life, and the biosphere provided oxygen,
which changed the nature of weathering, nutrient cycling, mobility
of redox sensitive elements such as iron and uranium, and in
turn provided environmental stresses that pushed life along new
evolutionary pathways.
Early understanding of the GOE was based on temporal trends
in bioelemental sediments, changes in mineralogy such as iron
mineral abundances (hematite and magnetite in iron formation),
the disappearance of detrital phases (uraninite and pyrite), and
the appearance of red beds in the continental sedimentary rock
record. Knowledge of the GOE has been enhanced and refined
using geochemical proxies derived from bioelemental sediments
that span this 100 million year interval. These proxies paint
a picture using broad brush-strokes that show the oxygenation of
the atmosphereeocean system was more gradual than previously
surmised and not a simple linear rise.
We demonstrate in this review that the fine lines necessary to
further focus this picture can only be attained by interpreting
geochemical data in a sedimentologic and oceanographic framework
that incorporates an understanding of diagenetic reactions. Although
basin diagenetic hydrostratigraphy is rarely, if ever, considered when
interpreting the geochemistry of sedimentary and metamorphic
rocks, it is a prominent control on diagenesis. What becomes obvious
is that geochemical trends often shift along lithofacies changes
and sequence stratigraphic bounding surfaces because they have
contrasting hydrologic properties. The holistic method advocated
in this review mitigates the potential for incorrectly interpreting
geochemical trends because it not only considers paleoenvironment
and oceanography, but also assesses the effects of fluid flow on
alteration. Such an approach should help determine whether trends
are local, regional, or truly related to the GOE.
Surprisingly, there are currently no studies that interpret high-
resolution geochemical data in a sequence stratigraphic frame-
work to understand the subtle nuances of Earth’s oxygenation. The
need to make such connections and understand the data in their full
geologic context is imperative as technological advances continue to
increase the rate at which geochemical data are generated. Caution
should be exercised so that our ability to measure chemical anom-
alies keeps pace with our capacity to understand them. Developing
a detailed appreciation of how chemical proxies respond to
alteration should be a central focus of future work. Only once the
alteration of bioelemental sedimentary rocks is better understood
can the GOE and Earth’s subsequent oxygenation history be fully
interpreted.
Acknowledgements
This paper was improved through critical review by P.G. Eriksson
and four anonymous reviewers. We are grateful to N.P. James, T.K.
Kyser, F. Pirajno, T. Clarke, G. Broadbent, D. Rossell, and P.W. Fralick
Figure 13. Conceptual framework for integrating sedimentologic and geochemical
studies of bioelemental sedimentary systems. TL ¼ transmitted light microscopy;
RL ¼ reflected light microscopy; CL ¼ cathodoluminescent microscopy.

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