Assessing the utility of Fe/Al and Fe-speciation to record water column redox
conditions in carbonate-rich sediments
Clarkson, M.O.1, Poulton, S.W.2, Guilbaud, R.2 and Wood, R.1
1
School of Geosciences, University of Edinburgh, West Mains Road, Edinburgh,
EH9 3JW, UK
2
School of Earth and Environment, University of Leeds, Leeds, LS2 9JT, UK
Abstract
Geochemical proxies based on Fe abundance (Fe/Al) and Fe-speciation have been
widely applied to marine sediments in order to unravel paleo-depositional redox
conditions though geological time. To date, however, these proxies have only been
calibrated in relation to modern and ancient siliciclastic marine sediments. This clearly
limits their use, particularly in relation to carbonate-rich sediments and rocks. To
address this, we here explore the applicability of Fe-based redox proxies in carbonates
through three approaches. First, we have compiled Fe/Al data for modern marine
sediments to investigate variability in Fe-enrichments as a function of carbonate content
and depositional setting. Second, we have expanded this approach with a compilation of
new and existing Fe-speciation data for modern and ancient marine sediments
deposited under oxic and euxinic (anoxic and sulfidic) water column conditions. Finally,
we show new data from paired limestone and dolomite sample sets to demonstrate the
potential significance of deep burial dolomitization on the Fe/Al and Fe-speciation redox
proxies.
Modern marine sediments deposited under oxic conditions show no relationship
between Fe/Al and carbonate content. These sediments have an average Fe/Al ratio of
1
0.55 ± 0.11, with some higher values potentially being attributable to steady-state early
diagenetic remobilization of Fe towards the sediment-water interface. In contrast,
significant Fe/Al enrichments occur as a consequence of water column Fe mineral
formation and deposition, either under anoxic conditions, or due to input of anoxic
hydrothermal fluids into oxic seawater. Iron speciation data also show no direct
correlation with carbonate content, and instead three groups can be distinguished based
on total Fe (FeT) and organic C contents. Sediments deposited under oxic water column
conditions, with FeT >0.5 wt%, generally plot below the 0.38 FeHR/FeT siliciclastic
reference threshold for distinguishing oxic and anoxic environments, regardless of
organic C content. Also consistent with siliciclastic calibrations, carbonate-rich
sediments that contain significant organic matter (>0.5 wt%) and which were deposited
under anoxic water column conditions tend to have FeHR/FeT ratios >0.38, independent
of FeT content. In contrast, oxic carbonate-rich sediments with low FeT (<0.5 wt%) and
low organic C (<0.5 wt%) routinely give a spuriously high FeHR/FeT ratio, suggesting that
the use of Fe-speciation for such samples is not appropriate for evaluating water column
redox conditions. Analysis of burial dolostones suggests that the Fe-speciation proxy
may also be compromised by deep burial dolomitization, where there has been a clear
source of mobile Fe to enrich rocks during recrystallization. This new assessment
expands the utility of Fe-based redox proxies to also incorporate appropriate carbonaterich rocks, provided that care is taken to assess the possible impact of deep burial
dolomitization.
2
1. Introduction
Ancient redox reconstructions are a major focus of paleoenvironmental research
and have greatly advanced our understanding of biogeochemical cycles, key
evolutionary events, and past periods of environmental change (e.g., Canfield, 2005;
Lyons et al., 2009; Lyons and Severmann, 2006; Meyer and Kump, 2008; Poulton and
Canfield, 2011; Raiswell and Canfield, 2012). Two of the most widely utilized
geochemical proxies in the redox toolbox are built upon the environmental behavior of
Fe, through enrichments of total Fe relative to aluminum (Fe/Al), and highly reactive Fe
to total Fe (FeHR/FeT) (Lyons and Severmann, 2006; Poulton and Canfield, 2011; Poulton
and Raiswell, 2002; Raiswell and Canfield, 1998; Raiswell et al., 2001). FeHR refers to Fe
minerals that are considered highly reactive towards biological and abiological
reduction under anoxic conditions (Canfield et al., 1992; Poulton et al., 2004a), and
includes
carbonate-associated
Fe
(Fecarb;
e.g.,
ankerite
and
siderite),
ferric
(oxyhydr)oxides (Feox; e.g., goethite and hematite), magnetite Fe (Femag) and Fe sulfide
minerals (Fepy; e.g., makinawite and pyrite) (Poulton and Canfield, 2005).
Sediments may be enriched in FeHR under anoxic marine conditions due to either
export of remobilized Fe from the oxic shelf (Anderson and Raiswell, 2004; Duan et al.,
2010; Raiswell and Anderson, 2005; Severmann et al., 2008), or under more widespread
anoxia, due to upwelling of deep water Fe(II) (Poulton and Canfield, 2011). Precipitation
of this mobilized water column Fe is then potentially induced through a variety of
processes, including Fe sulfide precipitation if the Fe encounters water column sulfide,
through biogenic or abiogenic oxidation of Fe(II) to form Fe(III)-containing minerals, or
via direct precipitation of Fe(II) carbonates or phosphates (e.g., Canfield et al., 1996;
Crowe et al., 2008; Jilbert and Slomp, 2013; Raiswell and Canfield, 1998; Zegeye et al.,
2012). These processes have the consequence that FeHR/FeT ratios in deposited
3
sediments provide a particularly sensitive means to determine whether a depositional
setting was oxic or anoxic.
Calibration in modern and ancient marine environments suggests that FeHR/FeT
<0.22 indicates oxic water column conditions, while FeHR/FeT >0.38 provides a robust
indication of deposition from an anoxic water column (Poulton and Canfield, 2011;
Poulton et al., 2002; Raiswell and Canfield, 1998; Raiswell et al., 2001). Values between
0.22-0.38, however, are somewhat equivocal, and care needs to be taken to determine
whether such values are a consequence of masking of the additional anoxic water
column flux of FeHR, either due to rapid sedimentation (Lyons and Severmann, 2006;
Poulton et al., 2004b; Raiswell and Canfield, 1998), or due to post-depositional
transformation of unsulfidized FeHR minerals to less reactive sheet silicate minerals
(Cumming et al., 2013; Poulton et al., 2010; Poulton and Raiswell, 2002). By additionally
examining the ratio of Fepy/FeHR, the Fe speciation technique has the unique advantage
in that it allows the separation of anoxic settings into euxinic (sulfidic) environments
(Fepy/FeHR >0.7-0.8) and non-sulfidic (Fe-rich; ferruginous) environments (Fepy/FeHR
<0.7) (März et al., 2008; Poulton and Canfield, 2011; Poulton et al., 2004b).
Fe/Al ratios provide a bulk measurement of this enrichment in Fe HR, which can
allow anoxic and oxic depositional environments to be distinguished. The inclusion of
‘unreactive’ Fe (FeU), largely in the form of silicate-associated Fe, tends to make Fe/Al
less sensitive than FeHR/FeT and more difficult to define a normal oxic baseline level for
(this tends to vary quite considerably, dependent on the specific depositional setting and
terrestrial sediment source; van der Weijden, 2002). Nevertheless, for Fe/Al it is
common to consider that enrichments above the average oxic Phanerozoic shale value of
0.53 ± 0.11 denote anoxic conditions (Lyons and Severmann, 2006; Raiswell et al.,
2008). Furthermore, a particular advantage of the Fe/Al proxy is that it does not suffer
4
from the possibility of post-depositional transformation of unsulfidized FeHR to less
reactive minerals, and thus Fe/Al and Fe-speciation in combination provide a
particularly powerful means to evaluate water column redox conditions (Cumming et al.,
2013; Lyons and Severmann, 2006; Poulton et al., 2010).
Normalization of Fe components to Al or FeT corrects for variable dilution by
carbonate or biogenic material (Raiswell and Canfield, 1998), and also lessens the
influence of variability in grain size and source mineralogy (Poulton and Raiswell, 2005),
making the proxies more widely applicable to different sediment types. Nevertheless,
the Fe-based redox proxies were developed and tested on siliciclastic-rich marine
sediments, and whilst the utility of the technique has been demonstrated for a variety of
chemical sediments, including banded iron formations (Poulton et al. 2004a; Poulton et
al., 2010) and some carbonate-rich marine sediments (Kendall et al., 2010; März et al.,
2008; Zerkle et al., 2012), the method has not yet been calibrated for carbonates. Indeed,
Lyons et al. (2012) highlight that careful consideration of lithology is required when
applying Fe-based redox proxies.
Theoretical concerns with the application of these proxies to carbonates largely
relate to the decreased detrital contents (and hence low FeHR and FeT), which ultimately
means that carbonate sediments are much more sensitive to highly reactive Fe inputs
that may originate from sources other than the detrital and anoxic water column inputs.
These additional sources could include the incorporation of low concentrations of Fe HR
into the carbonate lattice during carbonate precipitation under oxic conditions, or a
post-depositional influx of dissolved Fe into the sediment profile, a process which is of
particular concern during deep burial dolomitization (Warren, 2000).
These possibilities are particularly important to evaluate, as the application of
Fe-based redox proxies to carbonates could potentially provide a vast increase in spatial
5
and temporal understanding of ocean redox dynamics. As a lithology, carbonates
account for ~25% of the rock record and often provide important complimentary
information in the form of 13C and Sr isotope records, REE profiles, and carbonateassociated sulfur (CAS) estimates of the isotopic composition of contemporaneous
seawater sulfate (Gill et al., 2007; Hurtgen et al., 2009; Newton et al., 2004; Planavsky et
al., 2012). Additionally, carbonates often represent shallower water environments,
which tend to be centers of biodiversity and therefore record important evolutionary
events such as radiations and extinctions (Kiessling et al., 2010).
Here, we present an assessment of the utility of Fe/Al and Fe-speciation to record
water column redox conditions across a wide range in carbonate content. Firstly, we
explore a compilation of modern marine Fe/Al data to evaluate variability in the Fe
contents of carbonates from different depositional settings. Secondly, Fe-speciation data
are presented for a selection of new and published modern and ancient carbonate-rich
sediments deposited under oxic and euxinic water column conditions. Finally, the
potential impact of deep burial dolomitization is evaluated with new data obtained
across a dolomitization front in early Triassic carbonates from Oman. Together, this
approach allows us to place preliminary constraints on the careful application of Febased redox proxies to appropriate carbonate-rich sediments.
2. Materials and Methods
2.1. A compilation of modern core-top Fe/Al data
Data were compiled for modern (Holocene) oxic water column open-ocean and
continental margin core-top sediments from the Pangaea database (Table 1). These data
tend to represent shelf to basin environments and do not sample shallow marine
6
carbonate platforms. Nevertheless, the data provide important information on Fe
systematics in a wide range of settings of differing carbonate content. Note that nearshore environments that receive unusually high inputs of highly weathered terrestrial
sediment (e.g., those close to major river systems), such as the Amazon Shelf and Congo
Fan mobile mud belts, were not included in this compilation. Although this includes
relatively few samples in the Pangaea data-base, surface sediments from such sites tend
to have unusually high Fe contents (and FeHR in particular), due to both the highly
weathered nature of the sediment (Poulton and Raiswell, 2002) and due to intense
diagenetic remobilization of Fe (Aller et al., 2004; Aller et al., 1986). We also avoided
upwelling areas and other areas that tend to exhibit sporadic water column anoxia, and
instead utilize published data from persistently anoxic basins (Lyons et al., 2003;
Raiswell and Canfield, 1998) and from sites with significant hydrothermal Fe input
(Lyle, 1986; Dubinin, 2006; Govin et al., 2012), focusing on samples for which carbonate
concentration data were also available (Table 1).
2.2. Modern and ancient Fe-speciation data-set
New data are presented for modern carbonate samples from a diverse range of
environments, including shallow marine carbonate platforms (see Table 2). These
include pure biogenic carbonates, abiotic ooids and carbonate sands from temperate
and tropical environments. New data for ancient rocks come from Miocene carbonates
from Spain, representing carbonates with a simple alteration history of uplift and
meteoric weathering (Weijermars, 1991). Additionally, we incorporate modern and
ancient Fe-speciation data (Table 3) from published calibration studies (Table 3;
7
Canfield et al., 1996; Lyons et al., 2003; Poulton and Raiswell, 2002; Raiswell and
Canfield, 1998; Raiswell et al., 2008).
The dolomitization study was performed on limestone and dolomite pairs
sampled from the same beds across an oblique dolomitization front in the Early Triassic
Maqam Formation, Oman (Richoz, 2006). The dolomitization front is clear from the
orange colouration of altered samples, representing increased Fe. Carbon and oxygen
stable isotope measurements across the front show a depletion of the original oxygen
isotope signature in the dolostones compared to the limestones, whilst the carbon
values were preserved (Atudorei, 1999) consistent with deep burial dolomitization
(Richoz et al., 2010).
2.3. Geochemical methods
Fe-speciation extractions were performed according to calibrated extraction
procedures (Poulton and Canfield, 2005), whereby FeCarb was extracted with Na-acetate
at pH 4.5 and 50°C for 48 h, FeOx was extracted via Na-dithionite at pH 4.8 for 2 h, and
FeMag was extracted with ammonium oxalate for 6 h. FeT extractions were performed on
ashed samples (8 h at 550°C) using HNO3-HF-HClO4. All Fe concentrations were
measured via atomic absorption spectrometry and replicate extractions gave a RSD of
<5% for all steps. Acid volatile sulphur (AVS) and pyrite were determined
stoichiometrically from precipitated Ag2S after HCl and chromous chloride distillation,
respectively (Canfield et al., 1986). Total inorganic carbon (TIC) was measured using a
CM 5012 Coulometer. Total organic C (TOC) was measured on a LECO® carbon analyser
after carbonate removal (two 25 % (vol/vol) HCl washes for 24 hours). Replicate
analyses gave a precision of ± 0.09 wt% (2σ level).
8
3. Results
3.1. Modern core-top compilation
Figure 1 shows the total Fe and Al contents of the modern data compilation as a
function of CaCO3 content, with the data divided into oxic normal marine, anoxic marine
(which are all euxinic environments), and hydrothermal settings. The data show the
expected overall negative correlation between carbonate and both Fe and Al for all
depositional settings, highlighting the simple dilution effect of the carbonate on major
element concentrations. This results in very low FeT and Al contents at the highest
concentrations of carbonate.
In detail, Al tends to be more variable as a function of carbonate content for both
oxic and anoxic settings, presumably due to a greater degree of variability in the
lithogenic input, relative to Fe. For normal oxic marine sediments, the carbonate dilution
effect causes the range in FeT to decrease at higher carbonate contents, due to a
reduction in the relative impact of variability in the chemical composition of the
lithogenic fraction, which is also seen to a lesser degree for Al. In contrast, for samples
deposited beneath an anoxic water column, the lower lithogenic input at higher
carbonate contents means that FeT concentrations are more significantly affected by
relative variability in the rates of deposition of Fe minerals from the water column and
rates of carbonate formation. This results in enhanced variability in FeT as carbonate
increases (Fig. 1). Hydrothermal sediments might be expected to show enhanced
variability in FeT throughout the entire range in carbonate content, relative to Al. This is
because, in contrast to Al, FeT will be controlled by a balance between the rate of
hydrothermal Fe mineral deposition and the rate of carbonate production, both of which
9
are highly variable on a global scale. There is some indication that this may be the case
in Figure 1, but our data-set is not large enough to fully evaluate this suggestion.
Figure 2 shows Fe/Al ratios for the different depositional settings as a function of
CaCO3 content. The normal marine data exhibit a range in Fe/Al from 0.30 to 0.80 (with
an average of 0.55 ± 0.11), and show no correlation with carbonate content. The
combined euxinic data-set also shows no relationship with CaCO3 content, however, the
Black Sea and Kau Bay data-sets show an overall increase in Fe/Al as carbonate content
increases (see also Canfield et al., 1996; Raiswell and Canfield, 1998). Samples from the
Black Sea and Kau Bay also tend to be significantly enriched in Fe relative to the normal
marine data (Black Sea Fe/Al = 0.79 ± 0.10; Kau Bay Fe/Al = 0.87 ± 0.08). In contrast,
euxinic samples from the Cariaco Basin have low Fe/Al ratios that are more similar to
the normal marine data (0.49 ± 0.02). Fe/Al ratios are also high for hydrothermal sites
close to mid-ocean ridges (Fe/Al = 3.03 ± 3.77), with the highest values (and largest
range) occurring at higher carbonate contents (Fig. 2).
3.2. Modern Fe-speciation data
New Fe-speciation data are presented in Table 4 and compiled with literature
data in Figures 3 and 4. The anoxic modern and ancient data-sets generally plot above
the anoxic siliciclastic FeHR/FeT threshold value of 0.38, and show an overall increase in
FeHR/FeT with increasing carbonate content (Fig. 3). A few samples, largely comprising
sediments from Kau Bay, Indonesia (Middelburg, 1991) plot below 0.38, but generally
above the 0.22 FeHR/FeT threshold that is commonly taken as an upper value for robust
identification of oxic water column conditions in ancient samples (this value is based on
10
an average FeHR/FeT ratio for oxic water column deposition during the Phanerozoic of
0.14 ± 0.08; Poulton and Canfield, 2011; Poulton and Raiswell, 2002).
In contrast to the anoxic samples, sediments deposited from oxic bottom waters
show no direct relationship with carbonate content (Fig. 3). In general, most of the oxic
samples plot below 0.38, with 78% of the Phanerozoic oxic samples falling below the
0.22 threshold, as opposed to 37% for modern oxic samples. This decrease in the
Phanerozoic average FeHR/FeT ratio, relative to the modern, likely arises due to loss of
unsulfidized FeHR to sheet silicate minerals during diagenesis, as discussed by Poulton
and Raiswell (2002). However, some oxic samples plot significantly above 0.38, and this
is particularly the case for some modern samples with high carbonate contents.
Consideration of these data in terms of total Fe, rather than carbonate content,
shows that FeHR/FeT decreases at higher FeT for anoxic samples. In contrast, oxic
samples display no direct correlation between these parameters. However, from Figure
4 it is apparent that the majority of oxic samples that plot above the 0.38 threshold have
very low FeT (<0.5 wt%).
3.3. Paired Limestones and Dolomites
The limestone samples from the Early Triassic Maqam Formation, Oman (Richoz,
2006) have low FeT (<0.52 wt%; Table 5), and consistent with the oxic modern and
ancient compilation (Fig. 4), this results in elevated FeHR/FeT ratios, despite an inferred
oxic depositional setting for these samples. Nevertheless, comparison with dolomitized
samples from the same beds is instructive in terms of evaluating the potential role of
burial dolomitization on Fe-speciation. Sample pairs 1 and 2 show little variation
(within that expected for individual samples from the same bed) between the limestones
11
and dolomites, in terms of FeT, Fe partitioning between the different Fe pools, and
FeHR/FeT (Table 5). However, sample pairs 3, 4 and 5 show an increase in FeT during
burial dolomitization, which is particularly significant for samples 4 and 5. This increase
tends to arise as a result of an increase in Fecarb (Table 5), which is consistent with Fe
addition to the system during dolomitization, although for sample pair 4, approximately
50% of the additional Fecarb appears to be sourced from the Feox fraction. As a
consequence of the additional Fe input to the system, FeHR/FeT ratios are elevated in
dolostones relative to limestones for these samples.
4. Discussion
4.1. Behaviour of the Fe/Al paleo-redox proxy in carbonate-rich sediments
Figure 2 provides the first compilation of Fe/Al as a function of carbonate
content. The variability in Fe/Al observed for the oxic marine data-set (0.30-0.80) is
larger than, but overlaps, values measured for modern siliciclastic-dominated sediments
from oxic parts of the Black Sea, Effingham Inlet and Orca Basin (0.44-0.63; Lyons and
Severmann, 2006). A potential explanation for some of the relatively high Fe/Al ratios in
the modern compilation may relate to the sampling strategy used. To obtain a sufficient
amount of data, and to be internally consistent, we utilized core-top data for our
compilation. The potential for enrichment in total Fe in normal marine surface
sediments is well-documented (e.g. Aller, 1980; Leslie et al., 1990; Trefry and Presley,
1982), and arises due to steady state remobilization of highly reactive Fe during anoxic
diagenesis, followed by upwards diffusion and precipitation at the sediment-water
interface. This process is likely particularly prevalent in organic-rich siliciclastic
12
sediments (i.e., sediments low in carbonate), which is consistent with the occurrence of
the highest Fe/Al ratios at low CaCO3 in our compilation (Figure 2).
Despite this complication, and more significantly in terms of the use of Fe/Al as a
paleo-redox proxy, the mean of 0.55 ± 0.11 is entirely consistent with the Phanerozoic
normal marine average Fe/Al ratio of 0.53 ± 0.11 (Raiswell et al., 2008). This
Phanerozoic normal marine average is, however, based on siliciclastic sediments with
average Al concentrations of 8.68 ± 2.94 wt% (Raiswell et al., 2008). Importantly, the
lack of covariation between Fe/Al and carbonate content for modern normal marine
sediments (Fig. 2) suggests that the proxy behaves in a consistent manner during
deposition under oxic water column conditions, even when carbonate is high, and Fe T
and Al concentrations are low (c.f., Fig. 1).
The positive correlation observed between Fe/Al and CaCO3 for the independent
euxinic Black Sea and Kau Bay data-sets is well-documented (Canfield et al., 1996;
Raiswell and Canfield, 1998). This relationship arises because the organic matter that
fuels
sulfate
reduction
(and
hence
sulfide
production)
is
derived
from
coccolithophorides in these settings, and thus sulfidized Fe that forms in the water
column is intimately associated with their calcareous skeletons (Canfield et al., 1996;
Raiswell and Canfield, 1998). However, as shown by Lyons and Severmann (2006) for an
expanded suite of sediments from euxinic settings (for which CaCO3 data are not
available), Fe enrichments may be decoupled from biogenic sediment inputs, and the
overall controlling factor is simply precipitation of water column Fe under anoxic
conditions. Therefore, no simple global relationship exists between Fe/Al and CaCO3 in
anoxic settings. This is further exemplified in Figure 2, which also highlights one
potential issue with the Fe/Al proxy. While the Kau Bay and Black Sea data-sets show
clear enrichments in Fe/Al across a range of carbonate contents, relative to normal oxic
13
marine sediments, the Cariaco Basin euxinic sediments do not. However, euxinic Cariaco
Basin sediments are in fact enriched in Fe/Al, relative to the lithogenic sediment
supplied to the basin (Lyons et al., 2003). This highlights that local compositional
variability in the terrestrial sediment influx into the marine environment can potentially
mask sediment Fe/Al enrichments (relative to global average Fe/Al ratios) under anoxic
conditions.
The nature of Fe/Al ratios in carbonate-rich sediments from hydrothermal
settings (Fig. 2) is also consistent with the Fe enrichment mechanism outlined above. At
active spreading centres along the East Pacific Rise (E.P.R.) and Mid-Atlantic Ridge
(M.A.R.), anoxic fluids enriched in dissolved Fe(II) are vented into oxic seawater. In
sulfidic hydrothermal vent systems, Fe sulfides precipitate rapidly and are mostly
deposited close to the vent (e.g., Feely et al., 1987; Mottl and McConachy, 1990).
However, Fe(II) can continue to precipitate as Fe (oxyhydr)oxide minerals for some time
in the neutrally buoyant plume that forms tens to hundreds of metres above the vents
(Baker et al., 1985; Lupton and Craig, 1981; Reid, 1982). These plumes may be laterally
advected from the ridge crest for hundreds of kilometres, with the result that Fe
(oxyhydr)oxide minerals may continue to be deposited a considerable distance from the
active vent (Baker et al., 1985; Klinkhammer and Hudson, 1986; Poulton and Canfield,
2006). The hydrothermal data presented in Figure 2 clearly demonstrate this process,
and as might be expected due to the low lithogenic sediment input to such sediments, Fe
enrichments may be particularly large.
The hydrothermal data also highlight an important and often under-appreciated
aspect of the Fe enrichment process. Iron enrichments under euxinic conditions require
that Fe is transported into anoxic sulfidic waters and precipitated as sulfide minerals
(e.g., Canfield et al., 1996). Similarly, under ferruginous conditions, Fe minerals may
14
form under strictly anoxic water column conditions (e.g., Zegeye et al., 2012). However,
Fe (oxyhydr)oxide enrichments can also occur under oxic conditions, provided there is
an anoxic mechanism to transport Fe(II) into the oxic setting. This may occur as a result
of hydrothermal Fe(II) inputs into the ocean, as discussed above, but could also occur
due to upwelling of anoxic Fe(II)-rich waters into shallow oxic surface waters, a process
that may have been particularly prevalent during precipitation of some Precambrian
banded iron formations (e.g., the 1.88 Ga Gunflint Formation; Pufahl et al., 2000).
Sediments formed in this latter manner tend to be particularly enriched in ferric
(oxyhydr)oxide minerals, and in these specific Fe(II) upwelling cases, enrichments
actually imply that the adjacent deeper water column was anoxic, rather than the water
column directly overlying the site of Fe enrichment.
The above observations provide a framework for extending the potential
application of the Fe/Al paleo-redox proxy to incorporate carbonate-rich sediments.
Unfortunately, Fe/Al data is sparse for ancient sediments where paleo-redox conditions
have been independently evaluated and where carbonate contents are also available.
Nevertheless, initial calibrations of the Fe/Al proxy for siliciclastic samples were based
on modern sediments deposited under different redox conditions (Lyons and
Severmann, 2006; Lyons et al., 2003). Furthermore, the average modern and
Phanerozoic Fe/Al ratios for normal oxic marine sediments are almost identical. This
suggests that the oxic Phanerozoic siliciclastic Fe/Al ratio of 0.53 ± 0.11 (Raiswell et al.,
2008) is also appropriate for carbonate-rich sediments, provided that deep burial
dolomitization has not affected primary depositional Fe/Al ratios (see discussion
below). However, the large relative standard deviation on this ratio (20%) highlights
that the lithogenic sediment supplied to a particular locality can be highly variable in
terms of chemical composition, which appears to be primarily related to enhanced
15
variability in Al contents, relative to Fe (Fig. 1). Thus, where possible, the best approach
is to define an oxic baseline Fe/Al value for a particular setting (Lyons et al., 2003;
Poulton et al., 2010). Nevertheless, our data suggest that Fe enrichments significantly
above the normal oxic range (i.e., Fe/Al >0.64) can generally be used to identify anoxic
depositional conditions in modern and ancient settings, for both siliciclastic and
carbonate-rich sediments.
4.2. Fe-speciation in carbonate-rich sediments
Figures 3 and 4 demonstrate that most sediments deposited from anoxic bottom
waters have FeHR/FeT ratios above the 0.38 siliciclastic reference threshold for
recognizing anoxia in modern and ancient sediments (Poulton and Raiswell, 2002;
Raiswell and Canfield, 1998; Raiswell et al., 2001). This is a robust relationship that also
holds for samples with very high carbonate (Fig. 3). In more detail, however, Fe HR
enrichments tend to be more pronounced at higher CaCO3 in both modern and ancient
sediments (Fig. 3). As discussed above, this feature may, in part, be due to sulfate
reducing bacteria in the water column utilizing the organic matter associated with
carbonate producers (Canfield et al., 1996; Raiswell and Canfield, 1998). However, the
highest FeHR/FeT ratios also occur at low FeT, with a decrease as FeT increases (Fig. 4).
Thus, consistent with the more recent suggestion that Fe enrichments are decoupled
from biogenic sediment inputs (Lyons and Severmann, 2006), the degree of FeHR
enrichment is perhaps better described more generally, as a balance between rates of
water column FeHR deposition relative to the flux and composition of the lithogenic
sediment fraction.
Most of the anoxic sediments that plot below the 0.38 threshold are from Kau
Bay. These sediments are unusual, in that despite low FeHR/FeT, Fe/Al ratios are high
16
and entirely consistent with deposition from an anoxic water column (Fig. 2). Masking of
anoxic water column Fe enrichments due to rapid sedimentation is well-documented in
both modern (Canfield et al., 1996) and ancient (Poulton et al., 2004b) settings, but such
a process would reduce FeHR/FeT and Fe/Al ratios. In fact, the Kau Bay water column is
not persistently euxinic, and instead alternates between euxinic and low-oxygen
conditions (Middelburg, 1991). This gives two potential explanations for the low
FeHR/FeT and high Fe/Al ratios. Firstly, under low oxygen, non-euxinic conditions, Fe
minerals such as siderite and magnetite may ultimately be enriched in the deposited
sediment (see Poulton and Canfield, 2011). This would give elevated Fe/Al, but these
minerals are not extracted by the Fe extraction technique (Raiswell et al., 1994) used by
Raiswell and Canfield (1998) to analyse these samples, and thus FeHR/FeT ratios would
be low. A second possibility is that during low oxygen, non-euxinic intervals, FeHR
minerals formed in the water column may escape sulfidation. In this case, early
diagenetic transformation of unsulfidized FeHR to sheet silicate minerals would
potentially reduce FeHR/FeT ratios, while maintaining high Fe/Al (e.g., Poulton et al.,
2010; Cumming et al., 2013).
The majority of oxic samples with FeT >0.5 wt% plot close to, or below, the 0.38
FeHR/FeT lower limit for anoxic siliciclastic sedimentation (Fig. 4). This suggests that
when FeT is greater than 0.5 wt%, the thresholds derived for siliciclastic sediments
deposited from an oxic water column are also appropriate for carbonate-rich sediments.
In contrast, all oxic sediments with <0.5 wt% FeT have FeHR/FeT ratios >0.38 (Fig. 4).
This is a generic feature of the data-set and does not solely relate to samples with very
high carbonate, although particularly high FeHR/FeT ratios are evident for the purest oxic
carbonates (Fig. 3). However, the low FeT for these samples does imply a relatively low
17
lithogenic fraction and hence a high biogenic fraction, whether it be carbonate or silica
(or abiotic gypsum in the case of the anoxic Yazerez sample; Table 2).
A review of literature data for Fe incorporation into a variety of calcifying marine
organisms (including molluscs, scleractinian corals and planktonic gastropods), suggests
that the FeT content may vary considerably in such carbonate biominerals (from 0.17 to
1540 ppm; Cravo et al., 2007; Foster and Chacko, 1995; Foster and Cravo, 2003; Keller et
al., 2007; Kumar et al., 2010; Turekian et al., 1973), which is consistent with our own
analyses (Table 4). Clearly, this concentration range is well below the lowest FeT
contents of our oxic sample suite (Fig. 4). Hence, Fe uptake by marine calcifying
organisms (which would be as FeHR) could potentially account for only a small fraction
of the FeHR in our oxic samples that have spuriously high FeHR/FeT ratios. Instead, these
samples have received FeHR from additional sources, which in some cases could include
slowly depositing Fe (oxyhydr)oxide minerals from an oxic water column (a process that
forms deep sea Fe-Mn nodules under oxic conditions), or due to Fe incorporation during
early diagenetic mobilization of Fe and associated carbonate recrystallization (see
discussion below).
The oxic data therefore suggest that when FeT is <0.5 wt%, Fe speciation should
not be used to recognize oxic sedimentation. However, Figure 4 highlights an additional
important constraint on the use of Fe speciation for carbonate-rich sediments. Modern
and ancient samples deposited from anoxic bottom waters, but with Fe T <0.5 wt% (i.e.,
mostly carbonate-rich samples; Fig. 3), give FeHR/FeT ratios that appropriately record
anoxic sedimentation. In this case, however, the ratios are consistent with the overall
negative trend of decreasing FeHR/FeT with increasing FeT, in contrast to the oxic
samples with FeT <0.5 wt% (which scatter across a wide range in FeHR/FeT; Fig. 4). This
suggests that Fe speciation may still be appropriate for recognizing anoxic depositional
18
conditions for samples with low FeT, but additional constraints are required to
distinguish such samples from oxic, low FeT samples with spuriously high FeHR/FeT
ratios.
In this context, Raiswell et al. (2001) suggest that samples used to identify anoxic
water column sedimentation via Fe speciation should generally be organic-C bearing,
although no concentration limits were suggested in this study. To assess this for
carbonate-rich sediments, Figure 5 recasts the FeHR/FeT data in terms of TOC rather than
FeT, with the oxic low FeT samples (i.e., those with misleadingly high FeHR/FeT ratios; Fig.
4) distinguished as open triangles. Thus, based on organic C contents, the low FeT, high
FeHR/FeT oxic samples can generally be distinguished from similar samples deposited
under anoxic water column conditions. In fact, almost all oxic low FeT samples plot
below 0.5 wt% TOC, whereas anoxic samples (including those with very high carbonate;
Fig. 4) tend to have considerably higher TOC (with the exception of two samples),
consistent with higher production and/or preservation of organic matter under anoxic
conditions. Figure 5 also shows that oxic samples which do behave in a consistent
manner with regard to their Fe speciation characteristics (i.e., those with Fe T >0.5 wt%),
may also have very low TOC contents. Hence, it is not appropriate to simply define a TOC
threshold for the use of Fe speciation, and instead, TOC and FeT concentrations should
be considered in tandem.
These combined constraints are demonstrated in Figure 6 and summarized in
Table 6. Oxic samples behave appropriately with regard to Fe speciation, provided that
FeT concentrations are >0.5 wt%, and no constraint on minimum TOC content is
required in this case. The main exception to this concerns the unlithified stromatolite
from Lagoa Vemelha, Brazil (Table 2), which has elevated TOC (2.12 wt%) but low FeT
(0.23 wt%). However, the lithified stromatolite sample from the same environment
19
behaves more consistently (FeT = 0.02 wt%; TOC = 0.16 wt%), suggesting that
unlithified stromatolites may be an anomaly, but such samples will not feature
prominently in the geologic record. The Fe speciation technique also works
appropriately for anoxic carbonate-rich sediments, regardless of FeT content, but these
samples require a minimum TOC content of 0.5 wt% to be distinguished. In contrast,
samples where both FeT and TOC are <0.5 wt% (open triangles; Fig. 5 and 6) do not
record appropriate FeHR/FeT ratios, and hence Fe speciation should not be used in these
cases. However, it is also interesting to note that these samples plot in their own distinct
field on Figure 6, and thus we tentatively suggest that low FeT (<0.5 wt%) and low TOC
(<0.5 wt%) contents in carbonate-rich samples may indicate oxic sedimentation,
without the requirement for Fe speciation analyses.
4.3. Diagenetic alteration
A potential caveat to the boundary conditions outlined above for the application
of Fe speciation to carbonate-rich rocks and sediments concerns the possibility for Fe
enrichment during diagenesis. Carbonate diagenesis is complex with numerous
potential stages of recrystallization and cementation. Since the partition coefficient for
Fe, with respect to calcite, is greater than unity, preferential scavenging of Fe occurs
during precipitation (Barnaby and Rimstidt, 1989). This scavenging increases as Mg
increases, because Mg and Fe have greater miscibility in a carbonate lattice than Ca and
Fe. This leads to enrichment of Fe in the carbonate lattice compared to the fluid,
although Fe can also be present between lattice planes, in lattice defects, along crystal
boundaries, or can be added through adsorption (Tucker and Wright, 1990). For Fe to be
incorporated into the carbonate lattice, however, it must be in the divalent state,
20
therefore the Eh of the pore fluid is the controlling factor on the incorporation of Fe
(Barnaby and Rimstidt, 1989). Significant Fe incorporation is therefore not likely in
primary precipitates or unconfined systems where oxygenated open marine and
meteoric waters influence recrystallization, and this is reflected in the low Fe contents of
primary abiotic and biotic calcites (Section 4.2).
Anoxic pore waters, however, may promote the build-up of Fe2+ that can then be
incorporated
into
carbonate
during
early
cementation
and recrystallization.
Furthermore, the associated decrease in pH and increase in alkalinity that occurs during
organic matter remineralization will enhance dissolution and reprecipitation of
carbonate, therefore promoting the incorporation of Fe2+ into early cements. These early
diagenetic processes would result in the transfer of FeHR between the different pools
comprising the FeHR fraction (as is also the case with siliciclastic sediments), and
particularly into carbonate phases during early recrystallization. When there is a
significant lithogenic input of Fe, overall enrichments in FeHR/FeT are unlikely to occur,
as the total FeHR pool should be conserved. However, when FeT is low, the spatial
heterogeneity of these processes is more likely to result in FeHR enrichments. Potentially,
if there is a significant external meteoric component to the anoxic diagenetic fluids, Fe
will also be added to the system. Early recrystallization is exemplified by our modern
carbonate mud and sand samples from Abu Dhabi (AD samples in Tables 2 and 4), which
were deposited in a carbonate ramp environment and have undergone early
dolomitization (see below) or have been affected by bacterial sulfate reduction within
microbial mats (Lokier and Steuber, 2008). Samples with FeT >0.5 wt% do not show
enrichments in FeHR/FeT above the 0.38 threshold (despite the fact that ratios may have
been somewhat increased during early dolomitization), but samples with low Fe T do
21
show considerable enrichments, providing support for our FeT thresholds defined
earlier.
Since the incorporation of Fe is linked to the Mg content, the stabilization of high
magnesium calcite (HMC) to low magnesium calcite (LMC) will reduce the tendency of a
carbonate to take up Fe. LMC is the more stable form, and so ultimately calcite will tend
towards this composition. This may result in a minor loss of Fe in confined systems
during conversion, but more importantly, will aid the resistance of LMC carbonates to
later Fe additions. The largest concern for Fe-speciation, however, is that if substantial
quantities of Mg2+ are available during recrystallization then dolomitization occurs.
Various kinetic problems tend to prevent the formation of dolomite as a primary
precipitate in normal seawater (Land, 1998). These problems may be overcome by
raising the Mg/Ca ratio of pore fluids (such as in sabkha settings, reflux dolomitization
or through meteoric influence), or through microbial mediation or precipitation at
higher temperatures (burial dolomites) (Warren 2000). Due to the greater stability of Fe
and Mg, the effective distribution coefficient of Fe between dolomite and water is higher
than for calcite. This leads to the preferential uptake of Fe, relative to Ca, during
dolomitization (Tucker and Wright, 1990). Hence dolomites in the field can often be
identified by their pink/orange colouration compared to limestones. However, the total
Fe available for incorporation is again limited by pore water chemistry. For burial
dolomites, Mg2+ may be supplied from a number of sources including bittern salt
dissolution and clay mineral transformations (Boles and Franks, 1979; Kahle, 1965;
Warren, 2000). This additionally supplies Fe from the unreactive fraction of
autochthonous or allochthonous shales that is mobilized in anoxic pore waters, just as
with siliciclastic diagenesis, and the total supply of Fe will be determined by the degree
of open system behavior. The increased alkalinity of pore waters in carbonate-rich
22
sediments, however, promotes the precipitation of siderite, ferroan dolomite and
ankerite, which augments the FeHR fraction (these minerals are extracted as part of the
Fecarb fraction; Poulton and Canfield, 2005). Thus, deep burial dolomitization has a
significantly greater potential to affect FeHR/FeT ratios across a wide range in FeT than
early cementation, recrystallization or shallow dolomitization.
Our data for samples collected across a deep burial dolomitization front in Early
Triassic carbonates of the Maqam Formation, Oman (Table 5), clearly highlight the
potential impact of this process on FeHR/FeT ratios. Thus, prior to any Fe speciation
study of ancient carbonates, samples should be screened for additional Fe input via deep
burial dolomitization. The extent of burial dolomitization is largely controlled by local
permeability and, as such, is a localized process, creating sharp contacts between
replaced dolomite and unaltered limestones (e.g., Carmichael and Ferry, 2008).
Dolomitization may be readily discernable in outcrops from the distinct colouration of
dolomites, but in addition, detailed petrographic distinction should be made between
early and burial dolomitization phases, and saddle (ferroan) dolomites should not be
included in Fe speciation studies. More detailed analyses are also possible, since
incorporation of Mn and Fe into diagenetic cements can preserve a record of Fe
mobilization, and this may be recognized via staining and cathodoluminescence (e.g.,
Dickson, 1965; Tobin et al., 1996). The measurement of oxygen isotopes can also give an
estimate of recrystallization temperatures, and therefore depth, which may ultimately
help to discriminate between burial dolomitization and early diagenetic alteration.
5. Summary
We present the first compilations of Fe/Al and Fe-speciation data with carbonate
content to investigate the utility of these proxies as redox indicators in carbonate-rich
23
sediments. Based on this we suggest new limits for the careful application of Fespeciation and Fe/Al, which are constrained by both FeT and TOC contents. When both
TOC and FeT are <0.5 wt%, samples commonly record elevated FeHR/FeT ratios and thus
Fe speciation cannot be used to identify paleo-depositional redox conditions in these
cases, although low FeT and low TOC values in combination may indicate oxic
sedimentation. For oxic samples, Fe-speciation behaves consistently, regardless of
lithology, when FeT is >0.5 wt%. Fe/Al ratios in oxic carbonate-rich sediments are close
to that of average Phanerozoic shale deposited under oxic water column conditions
(0.53 ± 0.11; Raiswell et al., 2008), but the high relative standard deviation, caused
primarily by enhanced variability in the Al content of the lithogenic fraction, suggests
that a local oxic baseline should be defined where possible.
Fe-speciation also allows carbonate-rich sediments deposited from anoxic waters
to be identified, regardless of FeT content or lithology. However, an important caveat
here is that the identification of water column anoxia additionally requires a TOC
content of >0.5 wt%. This pattern is particularly clear in Jurassic Kimmeridge Clay
samples (Raiswell et al., 2001), where high FeHR/FeT ratios and TOC contents are evident
for low FeT (<0.5 wt%) samples deposited under anoxic water column conditions. The
same principal is applicable to Fe/Al, where ratios >0.64 (i.e., 0.53 ± 0.11; Raiswell et al.,
2008) combined with TOC contents of >0.5 wt% suggest an anoxic depositional setting.
The impact of early diagenesis on Fe partitioning in carbonate-rich sediments is
in many ways similar to that for siliciclastics. During early diagenesis, Fe may be
transformed between the fractions that comprise the FeHR pool, but there will be a
tendency for preservation of the mobilized Fe as Fecarb in carbonate-rich sediments.
Early diagenetic recrystallization in carbonates may cause an increase in FeHR and FeT,
but this only tends to cause spuriously high FeHR/FeT ratios when FeT is low (<0.5 wt%).
24
Late stage, deep burial diagenesis may significantly enhance FeHR/FeT ratios (and hence
also Fe/Al). However, samples can be screened for these overprints, and hence careful
sample selection means that these Fe-based redox proxies can be applied to carbonaterich sediments, within the geochemical framework outlined above.
The above observations are entirely consistent with existing Fe speciation
studies on ancient carbonate-rich sediments for which independent evidence of anoxic
deposition exists, including the Toarcian (Jurassic) OAE (Raiswell et al., 2001) and the
Cretaceous Coniacian-Santonian OAE3 (März et al., 2008). Therefore, with careful prescreening of samples, Fe-based redox proxies can now routinely be applied to
carbonate-rich lithologies, opening up a rich potential archive for future reconstructions
of water column redox conditions.
Acknowledgements
MOC acknowledges funding from the Edinburgh University Principal's Career
Development Scholarship, International Centre for Carbonate Reservoirs and the Moray
Endowment Award. RW and SWP acknowledge support from NERC through the ‘Coevolution of Life and the Planet’ scheme. Sylvain Richoz and Rob Newton are thanked for
fieldwork assistance and sample collection, and we thank Rob Raiswell for compiling
and sharing data. Thanks also to Stephen Lockier, Sandy Thudope, Simone Kasemann,
Cees van der Land, Kate Darling, Simon Jung and André Bahr for supplying samples.
25
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Figure Captions
Figure 1: Plot showing Fe and Al contents as a function of CaCO3. The data are separated
into oxic normal marine, anoxic marine and hydrothermal settings.
Figure 2: Fe/Al ratios as a function of CaCO3, with the same classifications as in Fig. 1.
Dashed lines represent the normal oxic marine average Fe/Al ratio (± 1 s.d.).
Figure 3: Fe speciation data plotted against CaCO3 for modern (circles) and Phanerozoic
(squares) oxic (open) and anoxic (closed) samples. Dashed lines at 0.22 and 0.38
FeHR/FeT, and represent the oxic and anoxic interpretative thresholds (Poulton and
Canfield, 2011).
Figure 4: Fe speciation data plotted against FeT for modern (circles) and Phanerozoic
(squares) samples, with oxic (open) and anoxic (closed) settings distinguished. Dashed
lines at 0.22 and 0.38 FeHR/FeT represent the oxic and anoxic interpretative thresholds
(Poulton and Canfield, 2011).
Figure 5: FeHR/FeT plotted as a function of TOC. Dashed lines at 0.22 and 0.38 FeHR/FeT
represent the oxic and anoxic interpretative thresholds (Poulton and Canfield, 2011).
Samples with FeT <0.5 wt% are plotted as triangles.
31
Figure 6: FeT plotted as a function of TOC. Open triangles represent oxic samples with
FeT <0.5 wt% (these samples record spuriously high FeHR/FeT ratios).
32
Tables
Table 1. Data sources for modern Fe/Al core-top calibration.
Environment
Normal Oxic Marine:
Atlantic
East Pacific
North Atlantic
North Pacific
Mediterranean
Reference
Govin et al. (2012)
Gromov (1975)
Dubinin and Rozanov (2001)
Dubinin (2006)
Mobius et al. (2010)
Anoxic
Black Sea
Cariaco Basin
Kau Bay
Raiswell and Canfield (1998)
Lyons et al. (2003)
Raiswell and Canfield (1998)
Hydrothermal
East Pacific Rise
Mid-Atlantic Ridge
Lyle (1986); Dubinin (2006)
Dubinin (2006); Govin et al. (2012)
33
Table 2. Sample descriptions for new Fe-speciation data.
Redox
Oxic
ID
Ooids
Pecten
Razor
Clam
Coquina
Coral
Description
Abiotic ooid
Biogenic
Locality
Abu Dhabi
Shallow water, UK
Age
Holocene
Holocene
Reference
This study
This study
Biogenic
Shallow water, UK
Holocene
This study
Biogenic
Biogenic
Holocene
Holocene
Strom. lith
Biogenic
Strom.
unlith
Biogenic
Sharks Bay, Aus.
Caribbean
Lagoa Vemelha,
Brazil
Lagoa Vemelha,
Brazil
AD xx
Carbonate ramp sands &
mud
Abu Dhabi
Holocene
GBR xx
Inter reef carbonate sands
This study
This study
Vasconcelos et al.,
2006
Vasconcelos et al.,
2006
Lockier and
Steauber, 2008
Scoffin and
Tudhope, 1985
N.Uist 6
905/14
Azagador
Anoxic
Abad Marl
BS carb
BS ooze
Yazerez
Temperate Carbonate
sands
Carbonate slope sediment
Temperate shallow water
Limestone
Pelagic Diatom/carbonate
Authigenic carbonate
Unit 1 coccolith ooze
Deep water Carbonate/
gypsum
Great Barrier Reef,
Aus.
North Uist,
Scotland
Oman Margin
Holocene
Holocene
Holocene
Holocene
This study
Holocene
This study
Mediterranean
Miocene
Weijermars, 1991
Mediterranean
Black Sea
Black Sea
Miocene
Holocene
Holocene
Weijermars, 1991
Bahr et al., 2009
Bahr et al., 2009
Mediterranean
Miocene
Stefano et al., 2012
34
Table 3. Published data-sets used for Fe speciation and Fe/Al compilations.
Reference
Raiswell and Berner, 1986
Canfield et al., 1996
Raiswell and Canfield, 1998
Raiswell et al., 2001
Dubinin and Rozanov, 2001
Lyons et al., 2003
Locality
Robin Hoods Bay, Great Paxton,
Lillingstone Lovell, Deanshanger,
Tattenhoe, Snake Hill, Oslo, Cautley,
Ohio
Black Sea
Framvaren, Black Sea, Orca Basin, Kau
Bay
Kimmeridge Clay, Jurassic, UK
Trans Atlantic
Cariaco Basin
35
Table 4. New Fe-speciation data for carbonate-rich samples. Ordered by ascending FeT. BD = below detection (0.05 wt% for TOC).
Oxic
Anoxic
Sample
Razor Clam
GBR 38
Pecten
Coquina
Strom. lith
Ooids
Coral
GBR 41
AD B15
AD S31
AD S25
AD B8
Azagador
Strom. unlith
AD S70
AD S36
AD S67
GBR 36
AD B14
AD S21
N.Uist 6
AD S64
905/14
Abad Marl
BS ooze
BS cement
Yazerez
BD
BD
0.16
BD
0.06
BD
BD
BD
BD
BD
BD
2.12
0.19
0.44
0.65
BD
BD
0.09
0.12
0.31
1.98
0.31
CaCO3 %
98.64
89.85
94.77
92.62
93.31
89.15
92.73
96.84
85.42
54.86
64.51
60.54
90.79
59.33
33.45
30.37
52.02
93.31
87.68
32.24
77.84
68.66
66.31
41.53
FeT
0.004
0.008
0.014
0.021
0.023
0.035
0.068
0.073
0.085
0.089
0.117
0.132
0.201
0.231
0.383
0.408
0.415
0.439
0.641
0.644
0.648
0.737
0.923
2.086
Fecarb
0.028
0.033
0.018
0.020
0.026
0.023
0.068
0.052
0.049
0.087
0.064
0.061
0.072
0.071
0.063
0.101
0.047
0.412
wt%
Feox
0.009
0.022
0.011
0.010
0.003
0.012
0.078
0.116
0.046
0.059
0.048
0.265
0.094
0.118
0.127
0.107
0.174
0.135
Femag
0.012
0.025
0.010
0.013
0.016
0.015
0.013
0.028
0.026
0.033
0.030
0.036
0.037
0.039
0.023
0.044
0.050
0.096
Fepy
0.036
0.006
0.012
0.004
0.043
0.018
0.004
0.021
0.046
0.085
0.043
0.013
0.015
0.022
0.019
0.012
0.086
0.123
FeHR/FeT
1.23
1.17
0.62
0.54
0.75
0.52
0.81
0.94
0.43
0.65
0.44
0.85
0.34
0.39
0.36
0.36
0.39
0.37
Fecarb
0.33
0.38
0.35
0.42
0.29
0.34
0.42
0.24
0.29
0.33
0.35
0.16
0.33
0.29
0.27
0.38
0.13
0.54
FeX/FeHR
Feox
Femag
0.10
0.14
0.26
0.29
0.22
0.20
0.21
0.28
0.03
0.19
0.17
0.22
0.48
0.08
0.54
0.13
0.28
0.15
0.22
0.12
0.26
0.16
0.71
0.10
0.43
0.17
0.47
0.15
0.55
0.10
0.40
0.17
0.49
0.14
0.18
0.13
Fepy
0.43
0.08
0.24
0.09
0.49
0.27
0.02
0.10
0.28
0.32
0.23
0.03
0.07
0.09
0.08
0.05
0.24
0.16
4.46
0.28
2.48
86.86
79.65
56.25
0.441
0.696
1.598
0.086
0.148
0.089
0.071
0.035
0.736
0.017
0.062
0.042
0.150
0.220
0.001
0.74
0.67
0.54
0.27
0.32
0.10
0.22
0.07
0.85
0.46
0.47
0.00
TOC
BD
0
BD
0.05
0.13
0.05
36
Table 5. Fe-speciation data for limestone-dolomite sample pairs.
Sample
Dol.1
Lst. 1
Dol. 2
Lst. 2
Dol. 3
Lst. 3
Dol. 4
Lst. 4
Dol. 5
Lst. 5
FeT
0.364
0.438
0.350
0.367
0.422
0.316
0.826
0.521
1.050
0.030
Fecarb
0.065
0.054
0.060
0.057
0.276
0.032
0.663
0.011
0.930
0.011
wt%
Feox
0.196
0.244
0.187
0.224
0.138
0.167
0.206
0.500
0.163
0.014
Femag
0.020
0.015
0.019
0.018
0.017
0.016
0.000
0.000
0.000
0.000
Fepy
0.002
0.001
0.001
0.001
0.005
0.001
0.018
0.003
0.013
0.000
FeHR/FeT
0.78
0.72
0.76
0.82
1.03
0.68
1.07
0.99
1.05
0.84
Fecarb
0.23
0.17
0.23
0.19
0.63
0.15
0.75
0.02
0.84
0.43
FeX/FeHR
Feox
Femag
0.69
0.07
0.78
0.05
0.70
0.07
0.75
0.06
0.32
0.04
0.77
0.07
0.23
0.00
0.97
0.00
0.15
0.00
0.57
0.00
Fepy
0.01
0.00
0.01
0.00
0.01
0.00
0.02
0.01
0.01
0.00
37
Table 6. Summary of threshold values for the use of Fe speciation and Fe/Al as paleoredox proxies in carbonate-rich sediments that have not experienced Fe addition during
deep burial dolomitization. NA = not applicable (i.e., no threshold value required).
Water column
FeT
TOC
Redox
(wt%) (wt%) FeHR/FeT
Fe/Al
Oxic
>0.5
NA
<0.22
0.53 ± 0.11
Anoxic
NA
>0.5
>0.38
>0.64
38