Ferruginous Conditions:
A Dominant Feature of the
Ocean through Earth’s History
Banded iron formation
from the Hamersley
Group, Western Australia
Simon W. Poulton1 and Donald E. Canfield2
1811-5209/11/0007-00107$2.50 DOI: 10.2113/gselements.7.2.107
T
he reconstruction of oceanic paleoredox conditions on Earth is essential
for investigating links between biospheric oxygenation and major
periods of biological innovation and extinction, and for unravelling
feedback mechanisms associated with paleoenvironmental change. The occurrence of anoxic, iron-rich (ferruginous) oceanic conditions often goes unrecognized, but refined techniques are currently providing evidence to suggest
that ferruginous deep-ocean conditions were likely dominant throughout
much of Earth’s history. The prevalence of this redox state suggests that a
detailed appraisal of the influence of ferruginous conditions on the evolution
of biogeochemical cycles, climate, and the biosphere is increasingly required.
rally and spatially), from various
levels of oxygen depletion, through
anoxic and non-sulfidic conditions, to fully euxinic. It is during
anoxic, non-sulfidic conditions,
which can essentially be considered an intermediate state between
the oxic and euxinic end-members,
that the environmental behavior
of iron takes central stage. In this
situation, iron may be mobilized
as ferrous iron [Fe(II)] resulting in
the development of ferruginous
water column conditions. Here, we
Keywords : ocean chemistry, ferruginous, euxinic, iron speciation, paleoredox
highlight recent research which
suggests that ferruginous condiINTRODUCTION
tions may have been a dominant feature of anoxic oceanic
The reconstruction of ocean redox conditions in the past intervals throughout much of Earth’s history. Understanding
the driving mechanisms and broader influence of such
has long been a focus of paleoenvironmental research and
is critical for examining links between ocean chemistry conditions is currently a major goal.
and the evolution of the biosphere and atmosphere. For
example, there is widespread (but not universal) support
SOURCES AND SINKS
for the suggestion that delayed oxygenation of the ocean OF WATER COLUMN IRON
after the first rise of atmospheric oxygen (the Great
Our understanding of the processes that may lead to ferruOxidation Event, GOE; Holland 1984) ~2.4 billion years
ginous water column conditions is firmly rooted in studies
ago (Ga) restricted eukaryote evolution and limited primary
of the modern global iron cycle. Over the last decade or
productivity (potentially restricting organic carbon burial
so, considerable effort has been targeted at quantifying the
and hence oxygen production; Anbar and Knoll 2002). In
operation of the global iron cycle. Major advances have
this view, it was only after a second rise in oxygen about
resulted from studying iron minerals that are biogeochemi580 million years ago that oxygenation of the deep ocean
cally reactive in surface and near-surface environments
occurred, paving the way for the subsequent rapid evolu(Raiswell and Canfield 1998; Poulton and Raiswell 2002;
tion of the first large multicellular animals (Canfield et al.
Raiswell et al. 2006). Here, we briefly discuss the pathways
2007, 2008). At the opposite end of the spectrum, periods
for the generation of water column iron that were likely
of ocean anoxia in the Paleozoic have been implicated in
significant during ancient anoxic periods. In this context,
several prominent episodes of mass extinction (see Meyer
dissolved ferrous iron may enter the ocean from two main
and Kump 2008), while Mesozoic ocean anoxic events
sources: from hydrothermal vents and from sediments
(OAEs) have been linked to rapid radiation and turnover
during anoxic diagenesis (or potentially also from particuin marine plankton (Leckie et al. 2002).
lates during transport and deposition in anoxic settings)
(Poulton and Raiswell 2002; Raiswell 2006).
In this regard, our understanding of the links between
ocean chemistry and evolution of the biosphere relies on
In the modern oxic ocean, the vast majority of Fe(II)
detailed knowledge of the precise nature of ocean oxidareleased from deep sea hydrothermal vents either reacts
tion–reduction (redox) conditions. In many cases, largely
rapidly with dissolved sulfide and is deposited close to the
as a result of the available geochemical toolbox, reconstrucvents as pyrite (FeS2) or is oxidized to ferric (oxyhydr)oxide
tions of ocean redox conditions have tended to focus on
minerals. Iron(II) generated by reductive dissolution of
the extreme case of euxinia (anoxic and sulfidic). However,
reactive ferric (oxyhydr)oxide minerals during marine sedithe redox state of the ocean can vary widely (both tempoment diagenesis may also be released to the overlying water
column. As with the hydrothermal source, a large fraction
of this Fe(II) is rapidly reoxidized and deposited, but for
both sources, some Fe(II) may escape precipitation and
1 School of Civil Engineering and Geosciences, Newcastle University
hence contribute to the oceanic dissolved-iron inventory
Drummond Building, Newcastle upon Tyne NE1 7RU, UK
(Raiswell 2006; Tagliabue et al. 2010).
E-mail: [email protected]
2 Nordic Centre for Earth Evolution and Institute of Biology
University of Southern Denmark, 5230 Odense M, Denmark
E-mail: [email protected]
E lements , V ol . 7,
pp.
107–112
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Under anoxic oceanic conditions, however, the role of
hydrothermal and diagenetic dissolved-Fe(II) sources is
much more significant. Reduced iron released by these
pathways is essentially stable until transported to an environment where precipitation and deposition can occur. In
the case of euxinic settings, Fe(II) is rapidly titrated when
it encounters water column sulfide, resulting in an additional flux (and hence enrichment) of pyrite to the deposited sediment (Raiswell and Canfield 1998). Lyons and
Severmann (2006) suggested that sediment iron enrichments following water column precipitation in anoxic
environments are a unique feature of euxinic depositional
conditions. However, recent work on ancient sediments
(Poulton et al. 2004, 2010; Canfield et al. 2007, 2008; März
et al. 2008; Johnston et al. 2010; Li et al. 2010) has demonstrated that sediment iron enrichments also readily occur
during anoxic ferruginous deposition; in this case, however,
the additional iron flux occurs due to water column formation of unsulfidized minerals, such as ferric (oxyhydr)oxides (e.g. ferrihydrite) and/or ferrous carbonates (e.g.
siderite). These minerals may readily react with dissolved
sulfide and are often partially sulfidized during diagenesis.
Nevertheless, the tendency is for euxinic iron enrichments
to be dominated by pyrite, while under ferruginous conditions, a significant proportion of the authigenic iron
minerals may remain unsulfidized (e.g. Poulton et al. 2004;
Canfield et al. 2008).
RECOGNIZING FERRUGINOUS CONDITIONS
In contrast to euxinic conditions, where a range of trace
metal, isotope, and biomarker proxies have the potential
to detect free sulfide in the water column (albeit with
varying levels of success), the toolbox for identifying ferruginous conditions is much more limited. From the above
discussion, it is clear that the identification and mineralogical characterization of sediment iron enrichments are
the key factors for recognizing ferruginous deposition. For
some periods of Earth history this is relatively straightforward. For example, at certain times during the Precambrian,
extensive deposition of marine chemical sediments
enriched in iron (banded iron formation, BIF; Fig. 1) points
to at least episodic ferruginous oceanic conditions (see
below). However, the spatially and temporally restricted
Banded iron formation outcrop from the Hamersley
Group, Western Australia, and a close-up of mesoscale
banding (over 15 cm) in an early Paleoproterozoic BIF from Ontario,
Canada
Figure 1
E lements
nature of these deposits severely limits their use for the
identification of ferruginous conditions over much of
Earth’s history.
Instead, a geochemical technique for recognizing iron
enrichments in normal marine shales is more appropriate.
The use of Fe/Al ratios has significant potential for identifying authigenic enrichments in sediment iron and has
been successfully used in a number of cases (see Lyons and
Severmann 2006). However, while this technique is useful
for detecting anoxic conditions, particularly when an oxic
baseline Fe/Al ratio can be readily defined for a particular
setting (Lyons and Severmann 2006; Poulton et al. 2010),
it does not allow euxinic conditions to be distinguished
from ferruginous conditions. In this regard, Poulton and
Canfield (2005) developed a new sequential extraction
technique for iron, based on a refinement of existing redox
proxies centered on the mineralogy and biogeochemical
behavior of iron (e.g. Raiswell and Canfield 1998). This
technique allows full evaluation of the iron minerals that
are considered to be highly reactive (FeHR) towards biological and abiological reduction under anoxic conditions.
The iron is broadly apportioned into four different pools:
(1) iron associated with carbonate minerals (Fecarb ; siderite,
ankerite); (2) ferric (oxyhydr)oxide minerals (Feox; ferrihydrite, lepidocrocite, goethite, hematite); (3) magnetite
(Femag); and (4) Fe sulfide (Fepy; pyrite). The sum of these
pools gives the total concentration of highly reactive Fe
(FeHR = Fecarb + Feox + Femag + Fepy).
Extensive studies of modern (Raiswell and Canfield 1998)
and Phanerozoic (Poulton and Raiswell 2002) marine sediments have demonstrated that FeHR/FeT ratios do not exceed
0.38 during normal deposition under oxic water column
conditions (FeHR is normalized to total Fe, FeT, to allow for
variable dilution of absolute FeHR concentrations by organic
matter, silica and/or carbonate precipitation, and to minimize variations associated with different grain-size distributions). By contrast, values above this threshold are
commonly found when deposition occurs under anoxic
conditions. In fact, the 0.38 ratio represents an extreme
upper value for oxic deposition, and work on Phanerozoic
shales (Poulton and Raiswell 2002) has shown that ancient
sediments deposited from an oxic ocean have average FeHR/
FeT ratios of 0.14 ± 0.08. These observations allow anoxic
(FeHR /FeT ≥ 0.38) and oxic (FeHR /FeT ≤ 0.22) depositional
conditions to be distinguished (Fig. 2). Intermediate FeHR /
FeT values may still reflect anoxic depositional conditions,
but in these cases, water column iron enrichments may
have been masked by rapid sedimentation, as is common
during turbidite deposition (Raiswell and Canfield 1998;
Poulton et al. 2004), or the Fe HR pool may have been
depleted due to conversion of a proportion of the unsulfidized minerals to less reactive sheet silicate minerals
during burial diagenesis or metamorphism (Poulton and
Raiswell 2002; Lyons and Severmann, 2006; Poulton et al.
2010). This latter case can, in part, be evaluated by an
additional iron extraction aimed at determining sheet silicate minerals that are poorly reactive towards dissolved
sulfide (see Poulton et al. 2010). It should also be noted
that, unless reactive iron enrichments are solely a consequence of an enhanced hydrothermal iron flux, sediment
iron enrichments cannot occur throughout an anoxic
ocean (i.e. simple mass balance constraints require that
some sediments have to be depleted in reactive iron to
account for enrichments elsewhere). However, because
marine sediments display a wide variability in Fe HR /FeT
below 0.38 (i.e. below the lower limit for anoxia), depletions in FeHR /FeT due to mobilization of FeHR are difficult
to recognize. This is likely further enhanced by the fact
that reactive iron enrichments occur in specific environ-
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Early Archean to Late Paleoproterozoic
(>3.8 to ~1.9 Ga)
The widespread deposition of BIF early in Earth’s history
is a dominant feature of the Precambrian rock record. There
is no modern equivalent for these enigmatic deposits, and
considerable effort is still being made to understand the
mechanisms involved in their genesis and also to extract
the environmental information recorded by their geochemistry. Nevertheless, the one factor that has received nearuniversal accord concerns their formation following
transport of dissolved Fe(II) under anoxic oceanic conditions (e.g. Holland 1984). It is commonly assumed, based
on indirect evidence from rare earth element and Nd
isotope data (e.g. Derry and Jacobsen 1990), that the
majority of the Fe(II) required for BIF formation originated
from hydrothermal sources, but a significant contribution
from terrestrial runoff (Holland 1984) and diagenetic iron
cycling remains a possibility (see Raiswell 2006).
The distribution of BIF deposition from the early Archean
to the late Paleoproterozoic is shown in Figure 3. With the
exception of a temporally restricted episode of deposition
associated with extensive glaciations in the late
Neoproterozoic, major BIF deposition ceased at ~1.8 Ga
(Isley and Abbott 1999). It is thus tempting to suggest that
anoxic ferruginous conditions reigned supreme from the
early Archean to the late Paleoproterozoic, but in fact BIF
deposition occurred in pulses that were associated with
periods of enhanced volcanism (Isley and Abbott 1999).
Iron enrichments in late Archean marine shales provide
some evidence for persistent ferruginous conditions
between these pulses of BIF formation (Reinhard et al.
2009; Kendall et al. 2010), but in general, little is known
about the nature of ocean chemistry between many of the
major episodes of BIF deposition.
Recent high-resolution reconstructions of ocean redox
conditions have also documented a gradual evolution from
an ocean characterized by ferruginous conditions
throughout to an ocean exhibiting a more detailed redox
structure (Fig. 4). Based on a number of redox proxies, it
is increasingly apparent that dissolved oxygen began to
accumulate in the ocean (possibly transiently) prior to the
Great Oxidation Event and that oxygenated or partially
oxygenated surface waters were present by ~2.7 Ga (e.g.
Reinhard et al. 2009; Kendall et al. 2010). This early
oxygenation also had a major impact on the sulfur cycle,
increasing the oxidation of pyrite (probably during transE lements
0.2
1
0.8
1
anoxic
Fepy/FeHR
0.4
0.6
ferruginous
0
0.8
euxinic
THE PRECAMBRIAN
oxic
Ferruginous conditions can be distinguished from euxinic
conditions by considering the extent of pyritization of the
highly reactive iron pool (Fepy/FeHR) in sediments showing
clear evidence of anoxic deposition (Fig. 2). Observations
on the extent of highly reactive iron pyritization in Black
Sea sediments suggest that Fepy/FeHR = 0.8 characterizes
the upper limit for ferruginous deposition (Anderson and
Raiswell 2004). However, this value was defined using an
extraction scheme for iron which does not fully quantify
Fecarb or Femag, and recent work on Phanerozoic sediments
suggests that an upper threshold of 0.7 may be more appropriate for distinguishing ferruginous conditions (März et
al. 2008; Fig. 2).
FeHR/FeT
0.4
0.6
0.2
possibly euxinic
0
possibly anoxic
ments that promote water column iron precipitation
(which commonly happens during upwelling onto continental slopes and shelves). Hence, only minor mobilization
(i.e. loss) of reactive iron from global oceanic sediments is
required to explain significant spatially restricted enrichments elsewhere.
Conceptualization of the iron-speciation parameters
for evaluating ocean redox conditions. FeHR = highly
reactive iron; FeT = total iron; Fepy = pyrite iron
Figure 2
port and deposition) and hence the supply of sulfate to the
ocean. In some continental slope environments, this apparently led to the formation of transient, spatially restricted
episodes of water column euxinia (Reinhard et al. 2009),
a concept which we explore in more detail below. However,
although the redox structure of the ocean began to evolve
in the late Archean, it remains likely that this immense
period of time was dominantly characterized by deepwater, ferruginous conditions (Fig. 4).
Late Paleoproterozoic to Late Neoproterozoic
(~1.9 to 0.58 Ga)
The notion of widespread euxinic conditions in the midProterozoic, as invoked by Canfield (1998), has received
considerable attention, with geochemical studies providing
abundant evidence for euxinia at this time (e.g. Poulton
et al. 2004; Scott et al. 2008). Such conditions would have
titrated dissolved Fe(II) from the ocean (as pyrite), thus
spelling an end to the global deposition of BIF. Recently,
however, a detailed reconstruction of ocean redox structure
at a time coincident with the global termination of BIF
deposition at ~1.8 Ga suggests that although euxinic conditions were prevalent in continental margin/slope settings
(extending at least 100 km from the paleoshoreline), the
deep ocean remained anoxic and ferruginous (Fig. 4;
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Height
0.4
THE PHANEROZOIC
0.3
Taken as a whole, the Phanerozoic ocean has been generally
well oxygenated. However, numerous examples of anoxic
oceanic intervals have been documented throughout the
Paleozoic and Mesozoic. Many of these intervals provide
strong evidence for euxinia, although ferruginous conditions have been reported for the early Cambrian (Canfield
et al. 2008). Furthermore, as reviewed by Meyer and Kump
(2008), Paleozoic sections are heavily biased towards
shallow, epicontinental sea environments, and thus the
spatial and temporal extents of euxinia are often unclear.
In addition, low-resolution studies tend to overestimate
the significance of euxinic conditions. For example, rapid
redox changes, from euxinic through to oxic depositional
conditions, have recently been advocated for late Paleozoic
(Pennsylvanian) strata from the Midcontinent US, challenging the assumption of persistent euxinia as indicated
by bulk geochemical proxies (see Schultz 2004). Thus it
appears likely that euxinic intervals were short-lived, and
rapid redox cycling resulting in ferruginous intervals
remains a distinct possibility.
0.2
0.1
0
1.5
2.0
2.5
Time (Ga)
3.0
3.5
4.0
Temporal history of banded iron formation deposition
prior to 1.5 Ga. The y-axis reflects the weighted abundance of BIF deposition. Modified after Isley and A bbott (1999)
Figure 3
Poulton et al. 2010). Such conditions would have nevertheless resulted in the termination of BIF deposition, as
upwelling Fe(II) would have been precipitated (as pyrite)
on the euxinic continental margins.
At present, no other studies have evaluated the redox structure of the mid-Proterozoic ocean in such detail. It thus
remains possible that at certain times in the mid-Proterozoic, euxinic conditions were even more widespread
(Fig. 4). However, Scott et al. (2008) suggested that euxinic
conditions were increasingly difficult to maintain through
this period as a result of a nutrient-limiting feedback associated with molybdenum removal under euxinic conditions, while Canfield (2004) suggested that the temporal
extent of such conditions would have been limited by the
subduction of sedimentary sulfides and hence a reduction
in the size of the Earth-surface sulfur pool. Indeed, in a
global study of late Neoproterozoic marine sediments,
Canfield et al. (2008) found overwhelming evidence for
widespread ferruginous deep-ocean conditions (consistent
with a return to BIF deposition associated with
Neoproterozoic glaciations) and only occasional evidence
for euxinic conditions, prior to widespread deep-ocean
oxygenation at ~0.58 Ga (Canfield et al. 2007). Building
upon these observations, Johnston et al. (2010) and Li et
al. (2010) suggested that deep-water ferruginous conditions
were dominant at this time, with some areas experiencing
either persistent or transient mid-depth euxinia. The similarity in redox structure proposed for the early and late
stages of this period (Poulton et al. 2010; Johnston et al.
2010; Li et al. 2010) suggests that this may have been the
dominant structure of the ocean throughout the period,
although further detailed studies are required to evaluate
the evolution of oceanic redox structure over this interval.
Nevertheless, the available data suggest that ferruginous
deep-water conditions were likely a major feature of this
time period.
Early to Late Archean
(>3.8 to ~2.7 Ga)
This suggestion finds support in recent studies of Mesozoic
oceanic anoxic events (OAEs). The well-documented
Toarcian (early Jurassic) and Cretaceous OAEs display abundant, but not ubiquitous, evidence for euxinic deposition
(see Meyer and Kump 2008). However, a recent high-resolution study of Cretaceous OAE3 sediments from Demerara
Rise found rapid and cyclic changes in deep-ocean redox
conditions (März et al. 2008). In this case, euxinic conditions dominated, but after ~90–100 ky these conditions
were punctuated by ferruginous bottom waters that lasted
for ~15–20 ky. These observations suggest that ferruginous
intervals during Phanerozoic anoxic periods may have been
more prevalent than generally recognized, but the significance of these conditions is only just beginning to be
appreciated.
CONTROLS ON FERRUGINOUS
VERSUS EUXINIC CONDITIONS
The above observations highlight the hypothesis that ferruginous conditions were likely a major feature of the ocean
throughout many periods of water column anoxia. But
what controls the temporal and spatial extents of ferruginous and euxinic conditions? At the broadest scale, the
oceanic budgets of S and FeHR determine whether a specific
redox condition is possible (or favorable) when the ocean
is in an anoxic state, and thus quantification of these
budgets provides primary insight into potential controls
on redox variability through time. The global fluxes of S
and FeHR to the modern ocean (Table 1) are dominated by
continental-weathering inputs and by hydrothermal
Late Archean to Late Paleoproterozoic
(~2.7 to ~1.9 Ga)
O2
H2S?
Fe(II)
Late Paleoproterozoic to Late Neoproterozoic
(~1.9 to ~0.58 Ga)
O2
O2
H2S
Fe(II)
Fe(II)
O2
H2S
?
Fe(II)
Fe(II)
Fe
Fe
Fe
Fe
Fe
Cartoon depicting the evolution of ocean redox conditions in the Precambrian, prior to deep-ocean oxygenation 580 million years ago. The balance between euxinic and
ferruginous conditions after ~1.9 Ga may have been highly variable.
Figure 4
E lements
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Fe
activity (Poulton and Raiswell 2002; Raiswell et al. 2006).
In terms of FeHR sinks, a large proportion of riverine FeHR
(~85%) is trapped in flood plains and lagoons (Poulton and
Raiswell 2002). The remaining riverine FeHR is deposited
in the ocean and is augmented by Fe HR from aeolian,
glacial, and hydrothermal inputs (Poulton and Raiswell
2002; Raiswell et al. 2006).
Although internal dynamics exert a major influence on
the precise redox structure of the ocean, this simple oceanic
budget suggests that minor variations in the relative fluxes
of sulfur and FeHR through time would favor either euxinic
or ferruginous conditions. Prior to the onset of oxidative
weathering early in Earth’s history, sulfur was primarily
locked in sulfide minerals, and thus oceanic sulfate fluxes
and concentrations were very low (Canfield 1998). Coupled
with enhanced hydrothermal iron inputs to the deep
ocean, low seawater sulfate would promote widespread
ferruginous conditions, as evident from the rock record.
With the onset of oxidative weathering, increased oceanic
sulfate fluxes would allow the potential development of
euxinic conditions in continental margin settings, but
removal of sulfate in these environments would also
restrict the spatial extent of euxinia (Poulton et al. 2010).
Transient increases in the flux of sulfate to the ocean have
also been implicated in the potential for development of
euxinic conditions during Cretaceous OAEs (Adams et al.
2010), and the fine balance between the oceanic FeHR and
sulfur budgets would readily drive cycling between euxinic
and ferruginous conditions (e.g. März et al. 2008) as
oceanic sulfate concentrations were progressively depleted
under euxinic conditions.
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E lements
FeHR
S†
Continental sources*#
6.5 ± 1.7
2.6 ± 0.6
Hydrothermal*#
0.3 ± 0.1
0.5 ± 0.4
-
0.2 ± 0.1
6.8 ± 1.7
3.3 ± 0.7
Fluxes (×
In quantifying the inner-shore S sink, we assume that FeHR
in these environments reflects the global riverine particulate composition (1.9 wt%; Poulton and Raiswell 2002) and
that average sulfur contents are similar to adjacent oxic
shallow-shelf concentrations (0.2 wt%; Berner 1982).
Combining these FeHR and S contents yields a molar S:FeHR
ratio of 0.18:1, and combining this with the depositional
flux of FeHR in inner-shore environments yields a S flux of
~1.0 × 1012 mol y–1 (Table 1). The remaining sulfur flux
(2.3 × 1012 mol y–1) deposits on the continental margins
and in the deeper ocean.
The resulting molar S:FeHR ratios (Table 1) show whether
euxinic or ferruginous conditions are favorable. Therefore,
with S:FeHR > 1.8 (the typical molar S:FeHR ratio observed
during euxinic deposition; Poulton et al. 2004, 2010), there
is sufficient S to overwhelm the FeHR flux and potentially
drive an anoxic system to euxinia. By contrast, sulfur is
limiting with S:FeHR < 1.8, favoring the development of
ferruginous conditions. Our budget calculations yield a
S:Fe HR ratio of 1.8, suggesting that when the ocean is
anoxic, the global ocean system is precariously balanced
between ferruginous and euxinic conditions.
GLOBAL HIGHLY REACTIVE IRON (Fe HR ) AND SULFUR (S) BUDGETS FOR THE MODERN OCEAN
Table 1
1012
mol
Volcanic #
Total
Sinks (×
1012
mol
S:FeHR
y-1)
y-1)
Inner-shore areas*
5.5 ± 0.6
1.0
0.2
Global ocean*
1.3 ± 0.3
2.3
1.8
* FeHR budgets from Poulton and Raiswell (2002) and Raiswell et al. (2006)
# For sulfur fluxes see Canfield and Farquhar (2009).
† Uncertainties on S sinks are not reasonably definable.
CONCLUDING REMARKS
AND FUTURE DIRECTIONS
There is a growing body of evidence to suggest that ferruginous conditions were a major feature of anoxic oceanic
intervals throughout Earth’s history. However, the temporal
and spatial records are currently sparse, and despite decades
of research, detailed (preferably high-resolution) reconstructions of paleoceanic redox conditions are still needed.
In many regards, the consequences of ferruginous conditions are poorly understood. Implications for the evolution
of the biosphere are profound and require careful consideration, as do feedback mechanisms associated with the
potential for enhanced macro- and micronutrient drawdown under ferruginous conditions (e.g. März et al. 2008).
In this context, a focus on biogeochemical cycling in
modern ferruginous environments is required in order to
evaluate process controls in ancient settings. Researchers
often bemoan the rarity of modern-day euxinic environments for evaluating ancient euxinic intervals—the Black
Sea and the Cariaco Basin are the most significant examples
of modern euxinic environments. The situation is even
more severe with regard to modern ferruginous environments. Nevertheless, ferruginous settings do exist, and
ongoing research is providing valuable new insights into
the nature of biogeochemical cycling under such conditions (e.g. Crowe et al. 2008).
ACKNOWLEDGMENTS
We are grateful to Rob Raiswell for many inspiring discussions on the modern and ancient Fe cycle over the years.
We thank the Natural Environment Research Council (UK)
and the Danish Grundforskningsfond for funding various
aspects of our research into the Fe cycle. Lee Kump and
Ariel Anbar are thanked for thoughtful reviews.
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