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

1School of Geosciences, University of Edinburgh, West Mains Road, Edinburgh,

EH9 3JW, UK

2School 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

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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 carbonate-

rich rocks, provided that care is taken to assess the possible impact of deep burial

dolomitization.

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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

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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 FeHR, 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

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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 FeHR

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

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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 carbonate-

associated 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 Fe-

based 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

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carbonate platforms. Nevertheless, the data provide important information on Fe

systematics in a wide range of settings of differing carbonate content. Note that near-

shore 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; Canfield et al.,

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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% (2 level).σ

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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

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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

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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 and

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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 Fe T 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

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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

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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

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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

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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, FeHR

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 and

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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 FeT <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 FeT >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 FeT 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 Fe-

speciation 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%). Late

24

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 carbonate-

rich 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 pre-

screening 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 ‘Co-

evolution 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).

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Tables

Table 1. Data sources for modern Fe/Al core-top calibration.Environment Reference

Normal Oxic Marine:AtlanticEast PacificNorth AtlanticNorth PacificMediterranean

AnoxicBlack SeaCariaco BasinKau Bay

HydrothermalEast Pacific RiseMid-Atlantic Ridge

Govin et al. (2012)Gromov (1975)Dubinin and Rozanov (2001)Dubinin (2006)Mobius et al. (2010)

Raiswell and Canfield (1998)Lyons et al. (2003)Raiswell and Canfield (1998)

Lyle (1986); Dubinin (2006)Dubinin (2006); Govin et al. (2012)

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Table 2. Sample descriptions for new Fe-speciation data.Redox ID Description Locality Age Reference

Oxic Ooids Abiotic ooid Abu Dhabi Holocene This studyPecten Biogenic Shallow water, UK Holocene This studyRazor Clam

Biogenic Shallow water, UK Holocene This study

Coquina Biogenic Sharks Bay, Aus. Holocene This studyCoral Biogenic Caribbean Holocene This study

Strom. lith BiogenicLagoa Vemelha,

BrazilHolocene

Vasconcelos et al., 2006

Strom. unlith

BiogenicLagoa Vemelha,

BrazilHolocene

Vasconcelos et al., 2006

AD xxCarbonate ramp sands &

mudAbu Dhabi Holocene

Lockier and Steauber, 2008

GBR xx Inter reef carbonate sandsGreat Barrier Reef,

Aus.Holocene

Scoffin and Tudhope, 1985

N.Uist 6Temperate Carbonate

sandsNorth Uist,

ScotlandHolocene This study

905/14 Carbonate slope sediment Oman Margin Holocene This study

AzagadorTemperate shallow water

LimestoneMediterranean Miocene Weijermars, 1991

Abad Marl Pelagic Diatom/carbonate Mediterranean Miocene Weijermars, 1991Anoxic BS carb Authigenic carbonate Black Sea Holocene Bahr et al., 2009

BS ooze Unit 1 coccolith ooze Black Sea Holocene Bahr et al., 2009

YazerezDeep water Carbonate/

gypsumMediterranean Miocene Stefano et al., 2012

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Table 3. Published data-sets used for Fe speciation and Fe/Al compilations.Reference LocalityRaiswell and Berner, 1986 Robin Hoods Bay, Great Paxton,

Lillingstone Lovell, Deanshanger, Tattenhoe, Snake Hill, Oslo, Cautley, Ohio

Canfield et al., 1996 Black SeaRaiswell and Canfield, 1998 Framvaren, Black Sea, Orca Basin, Kau

BayRaiswell et al., 2001 Kimmeridge Clay, Jurassic, UKDubinin and Rozanov, 2001 Trans AtlanticLyons et al., 2003 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).wt% FeX/FeHR

Sample TOC CaCO3 % FeT Fecarb Feox Femag Fepy FeHR/FeT Fecarb Feox Femag Fepy

Oxic Razor Clam BD0

98.64 0.004 -GBR 38 BD 89.85 0.008 - - - - - - - - -

Pecten BD 94.77 0.014 - - - - - - - - -

Coquina BD 92.62 0.021 - - - - - - - - -

Strom. lith 0.16 93.31 0.023 - - - - - - - - -

Ooids BD 89.15 0.035 - - - - - - - - -Coral 0.06 92.73 0.068 0.028 0.009 0.012 0.036 1.23 0.33 0.10 0.14 0.43

GBR 41 BD 96.84 0.073 0.033 0.022 0.025 0.006 1.17 0.38 0.26 0.29 0.08AD B15 BD 85.42 0.085 0.018 0.011 0.010 0.012 0.62 0.35 0.22 0.20 0.24AD S31 BD 54.86 0.089 0.020 0.010 0.013 0.004 0.54 0.42 0.21 0.28 0.09AD S25 BD 64.51 0.117 0.026 0.003 0.016 0.043 0.75 0.29 0.03 0.19 0.49AD B8 BD 60.54 0.132 0.023 0.012 0.015 0.018 0.52 0.34 0.17 0.22 0.27

Azagador BD 90.79 0.201 0.068 0.078 0.013 0.004 0.81 0.42 0.48 0.08 0.02Strom. unlith 2.12 59.33 0.231 0.052 0.116 0.028 0.021 0.94 0.24 0.54 0.13 0.10

AD S70 0.19 33.45 0.383 0.049 0.046 0.026 0.046 0.43 0.29 0.28 0.15 0.28AD S36 0.44 30.37 0.408 0.087 0.059 0.033 0.085 0.65 0.33 0.22 0.12 0.32AD S67 0.65 52.02 0.415 0.064 0.048 0.030 0.043 0.44 0.35 0.26 0.16 0.23GBR 36 BD 93.31 0.439 0.061 0.265 0.036 0.013 0.85 0.16 0.71 0.10 0.03AD B14 BD 87.68 0.641 0.072 0.094 0.037 0.015 0.34 0.33 0.43 0.17 0.07AD S21 0.09 32.24 0.644 0.071 0.118 0.039 0.022 0.39 0.29 0.47 0.15 0.09N.Uist 6 0.12 77.84 0.648 0.063 0.127 0.023 0.019 0.36 0.27 0.55 0.10 0.08AD S64 0.31 68.66 0.737 0.101 0.107 0.044 0.012 0.36 0.38 0.40 0.17 0.05905/14 1.98 66.31 0.923 0.047 0.174 0.050 0.086 0.39 0.13 0.49 0.14 0.24

Abad Marl 0.31 41.53 2.086 0.412 0.135 0.096 0.123 0.37 0.54 0.18 0.13 0.16

Anoxic BS ooze 4.46 86.86 0.441 0.086 0.071 0.017 0.150 0.74 0.27 0.22 0.05 0.46BS cement 0.28 79.65 0.696 0.148 0.035 0.062 0.220 0.67 0.32 0.07 0.13 0.47

Yazerez 2.48 56.25 1.598 0.089 0.736 0.042 0.001 0.54 0.10 0.85 0.05 0.00

36

Table 5. Fe-speciation data for limestone-dolomite sample pairs. wt% FeX/FeHR

Sample FeT Fecarb Feox Femag Fepy FeHR/FeT Fecarb Feox Femag Fepy

Dol.1 0.364 0.065 0.196 0.020 0.002 0.78 0.23 0.69 0.07 0.01Lst. 1 0.438 0.054 0.244 0.015 0.001 0.72 0.17 0.78 0.05 0.00Dol. 2 0.350 0.060 0.187 0.019 0.001 0.76 0.23 0.70 0.07 0.01Lst. 2 0.367 0.057 0.224 0.018 0.001 0.82 0.19 0.75 0.06 0.00Dol. 3 0.422 0.276 0.138 0.017 0.005 1.03 0.63 0.32 0.04 0.01Lst. 3 0.316 0.032 0.167 0.016 0.001 0.68 0.15 0.77 0.07 0.00Dol. 4 0.826 0.663 0.206 0.000 0.018 1.07 0.75 0.23 0.00 0.02Lst. 4 0.521 0.011 0.500 0.000 0.003 0.99 0.02 0.97 0.00 0.01Dol. 5 1.050 0.930 0.163 0.000 0.013 1.05 0.84 0.15 0.00 0.01Lst. 5 0.030 0.011 0.014 0.000 0.000 0.84 0.43 0.57 0.00 0.00

37

Table 6. Summary of threshold values for the use of Fe speciation and Fe/Al as paleo-redox 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

RedoxFeT

(wt%)

TOC (wt%) FeHR/FeT Fe/Al

Oxic >0.5 NA <0.22 0.53 ± 0.11Anoxic NA >0.5 >0.38 >0.64

38