# An iodine record of Paleoproterozoic surface ocean oxygenation

Dalton Hardisty, Zunli Lu, Noah Planavsky, Andrey Bekker, Pascal Philippot, Xiaoli Zhou, Timothy Lyons

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Dalton Hardisty, Zunli Lu, Noah Planavsky, Andrey Bekker, Pascal Philippot, et al. An iodine record of Paleoproterozoic surface ocean oxygenation. Geology, 2014, 42 (7), pp.619-622. 10.1130/G35439.1. hal-03888453

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# An iodine record of Paleoproterozoic surface ocean oxygenation

Dalton S. Hardisty¹*, Zunli Lu², Noah J. Planavsky³, Andrey Bekker¹, Pascal Philippot⁴, Xiaoli Zhou², and Timothy W. Lyons¹

¹Department of Earth Sciences, University of California, Riverside, California 92521, USA

²Department of Geological Sciences, Syracuse University, Syracuse, New York 13244, USA

³Department of Geology and Geophysics, Yale University, New Haven, Connecticut 06511, USA

⁴Institut de Physique du Globe de Paris, Sorbonne Paris Cité, Université Paris Diderot, CNRS, 75238 Paris, France

## ABSTRACT

Constraining oxygen levels in the early Precambrian surface ocean has been a longstanding goal, but efforts have been challenged by the availability of suitable proxies. Here we present a novel approach, iodine geochemistry, which broadens our perspective by providing constraints on shallow, carbonate-dominated marine settings. Iodate ($IO_3^−$) persists exclusively in oxic waters and is the sole iodine species incorporated into carbonate minerals, allowing iodine-to-calcium ratios (I/Ca) in shallow carbonates to be used as a paleoredox indicator. Our data from a series of Mesoarchean through Paleoproterozoic carbonates deposited under shallow-marine conditions reveal a progressive surface ocean oxygenation in the early Paleoproterozoic. These data seem to indicate that a largely anoxic surface ocean extended throughout the Archean until the Great Oxidation Event (GOE) at ca. 2.4 Ga, implying that previous inferences of pre-GOE oxygen production may reflect oxygen oases, transient oxidation events, or oxygen levels below those required for $IO_3^−$ accumulation. The data suggest formation and persistence of $IO_3^−$ and, consequently, surface ocean oxygen concentrations of at least 1 µM during the GOE. Following the initial rise of oxygen, carbonate-associated iodine in globally extensive carbonate units deposited during the Lomagundi positive carbon isotope excursion at ca. 2.22–2.1 Ga suggests a widespread aerobic iodine cycle beyond that operating prior to the event, synchronous with high relative rates of organic carbon burial and apparent expansion of oxidative conditions.

## INTRODUCTION

Measurements of iodine speciation in anoxic basins demonstrate that $IO_3^−$ undergoes complete reduction to iodide ($I^−$) nearly simultaneously with $O_2$ depletion (Wong and Brewer, 1977; Emerson et al., 1979; Wong et al., 1985; Luther and Campbell, 1991). It was recently observed that during calcite precipitation, concentrations of carbonate-associated iodine increase linearly with those of $IO_3^−$ in the precipitating medium, but $I^−$ is completely excluded (Lu et al., 2010). Coupled with the well-known redox behavior of iodine, its seawater residence time of ~300 k.y. (Broecker et al., 1982), and concentrations near 450 nM in the modern ocean, this correlation suggests that I/Ca ratios in carbonates can be used as a robust indicator of the presence or absence of $IO_3^−$, and hence oxygen, in the water column. Further, post-depositional alteration of carbonates through diagenesis have been shown—if anything—to decrease the original I/Ca ratios (Loope et al., 2013), making post-depositional increases in I/Ca ratios highly unlikely.

Molybdenum and chromium geochemistry, each sensitive to aerobic manganese cycling, suggest mild and/or transient atmospheric oxygenation starting as early as 3.0 Ga (Crowe et al., 2013; Planavsky et al., 2014). Permanent accumulation of appreciable $O_2$ during the Great Oxidation Event (GOE) at roughly 2.3 Ga is further inferred from non-mass dependent sulfur isotope (NMD-S) fractionations (Bekker et al., 2004). However, constraints on surface ocean oxidation are limited both before and after the GOE, despite assumptions and mounting evidence that early oxygen was sourced from the marine photic zone following the advent of oxygenic photosynthesis 300–600 m.y. prior to the GOE (see Lyons et al., 2014). Here, we use a novel approach, I/(Ca + Mg) ratios in shallow-water carbonates, to determine the timing of the onset of marine aerobic iodine

cycling and construct the first temporal record of early oxygen production specific to the surface ocean.

## MATERIALS AND METHODS

I/(Ca + Mg) ratios were measured in 190 samples from 20 carbonate units ranging in age from 3.5 Ga to 1.9 Ga, with a focus on capturing the key periods of the earliest progressive oxygen expansion (Table 1). Magnesium is included to account for dolomitization, as all measured units are dolostones. Carbonates that experienced metamorphic temperatures greater than greenschist facies were not included in this study because previous work has shown significant iodine loss in shale at metamorphic temperatures of >400 °C (Muramatsu and Wedepohl, 1998). We generated fresh powders from carbonate samples to minimize the effects of weathering and surface contamination. Measurements of I, Ca, and Mg were performed on a quadrupole inductively coupled plasma–mass spectrometer (Bruker M90) at Syracuse University (New York, USA) according to the method of Lu et al. (2010). The precision of $^{127}$I is typically better than 1% and is not reported separately for each sample. The standard deviation in counts per second (cps) for three blanks in a row is typically below 300 cps, while the sensitivity for 1 ppb standard is 80,000–120,000 cps. The detection limit for I/(Ca + Mg) is usually better than 0.1 µmol/mol, a value that loosely corresponds to 10 nM $IO_3^−$ based on carbonate precipitation experiments by Lu et al. (2010).

## RESULTS AND DISCUSSION

### Archean

All carbonate samples predating ca. 2.45 Ga record iodine concentrations analytically indistinguishable from blank values (Fig. 1; Table 1),

<table><caption>TABLE 1. CARBONATES USED IN THIS STUDY AND PLOTTED IN FIGURE 1</caption><thead><tr><th>Geologic Formation</th><th>Age (Ga)</th><th>n</th><th>Percent containing iodine</th></tr></thead><tbody><tr><td>Dresser Fm., Warrawoona Group</td><td>ca. 3.5</td><td>8</td><td>0</td></tr><tr><td>Steep Rock Group</td><td>2.80–2.76</td><td>14</td><td>0</td></tr><tr><td>Tumbiana Fm., Fortescue Group</td><td>2.73–2.72</td><td>18</td><td>0</td></tr><tr><td>Cheshire Fm., Ngezi Group</td><td>2.65–2.55</td><td>3</td><td>0</td></tr><tr><td>Wittenoom Fm., Hamersley Group</td><td>2.56–2.50</td><td>16</td><td>0</td></tr><tr><td>Campbellrand Subgroup</td><td>2.56–2.54</td><td>6</td><td>0</td></tr><tr><td>Tongwane Fm., Chunniespoort Group</td><td>2.49–2.32</td><td>5</td><td>20</td></tr><tr><td>Duitschland Fm., Pretoria Group</td><td>2.49–2.32</td><td>2</td><td>0</td></tr><tr><td>Espanola Fm., Huronian Supergroup</td><td>2.48–2.32</td><td>8</td><td>0</td></tr><tr><td>Carbonates of Turee Creek Group</td><td>2.45–2.40</td><td>29</td><td>28</td></tr><tr><td>Gandarela Fm., Minas Supergroup</td><td>ca. 2.42</td><td>3</td><td>0</td></tr><tr><td>Gordon Lake Fm., Huronian Supergroup</td><td>2.32</td><td>9</td><td>11</td></tr><tr><td>Bad River Dolomite, Chocolay Group</td><td>2.32–2.22</td><td>2</td><td>0</td></tr><tr><td>Saunders Fm., Chocolay Group</td><td>2.32–2.22</td><td>1</td><td>0</td></tr><tr><td>Cercadinho Fm., Minas Supergroup</td><td>ca. 2.22</td><td>2</td><td>0</td></tr><tr><td>Lucknow Fm., Olifantshoek Group</td><td>2.15</td><td>5</td><td>80</td></tr><tr><td>Mcheka Fm., Lomagundi Group</td><td>2.15</td><td>37</td><td>95</td></tr><tr><td>Lower Albanel Fm., Mistassini Group</td><td>2.1</td><td>10</td><td>80</td></tr><tr><td>Aquas Claras Fm., Serra Pelada Sequence</td><td>2.1</td><td>8</td><td>50</td></tr><tr><td>Aluminum River Fm., Amer Group</td><td>1.95–1.91</td><td>4</td><td>0</td></tr></tbody></table>

Note: References to ages and data are available in Tables DR1 and DR2 (see text footnote 1).

<footnote>*E-mail: dhard003@ucr.edu.</footnote>
<figure><img src="image_2.png" /><figcaption>**Figure 1.** Compilation of all available I/(Ca + Mg) ratios through time. All ages are represented by the midpoint for ranges presented in Table 1. Data from Lu et al. (2010; diamonds), Loope et al. (2013; triangles), and this study (circles). Values from Lu et al. (2010) do not include Mg concentrations in the ratio, as these are for calcites. Oxygen curve generalized from Lyons et al. (2014). LE—Lomagundi Event; GOE—Great Oxidation Event; PAL—present atmospheric levels.</figcaption></figure>

indicating I⁻ as the dominant inorganic iodine species in the water column. A lack of appreciable IO₃⁻ accumulation in the Archean shallow-marine carbonates suggests deposition under anoxic conditions. Overall, such a notion agrees well with significant NMD-S isotope signatures in the rock record until ca. 2.4 Ga (Bekker et al., 2004), constraining atmospheric oxygen concentrations to less than 10⁻⁵ of present atmospheric levels (Pavlov and Kasting, 2002). Though the sample set predating 2.5 Ga only includes six out of the 20 carbonates measured (Table 1), our work focuses on key units directly leading up to the GOE and some, such as the Tumbiana Formation, used for previous inferences of aerobiosis (Thomazo et al., 2011). Nevertheless, this limited sample scheme leaves open the possibility of surface ocean oxygen production beyond the resolution of our data. Regardless, the I/(Ca + Mg) record does not, at face value, support the idea of widespread, persistent surface ocean oxygen prior to the GOE.

Reservoir controls seem unlikely for the lack of carbonate-associated iodine prior to the GOE. Overall, the largest iodine sink in the modern ocean is organic matter (OM) burial, particularly via marine algae (Lu et al., 2010). Without effective sequestration with marine algae and other organisms, iodine would be more abundant in the ocean, as we observe for other hydrophilic halogens such as Br and Cl (Sharp and Draper, 2013). We propose that the absence of advanced marine algae and fewer biological sinks in the Archean would yield a seawater iodine reservoir much larger than today. Considering this, the first appearance of carbonate-associated iodine would be solely dependent on redox conditions allowing for IO₃⁻ production and accumulation from either I⁻ or intermediate forms.

For the Neoarchean, initially low I/(Ca + Mg) ratios, as would be expected if IO₃⁻ accumulated at only trace levels, may have been particularly vulnerable to diagenetic loss, limiting the record of rising I/(Ca + Mg) to intervals of sufficient O₂ production and concomitant IO₃⁻ accumulation. Similarly, post-GOE carbonates lacking iodine (Table 1) may have been deposited under locally anoxic conditions despite a general rise in surface ocean oxygen availability, but without a more comprehensive geological context, diagenesis cannot yet be ruled out. Further work is needed that

specifically addresses IO₃⁻ incorporation into primary dolomite and its retention during dolomitization of primary calcite. However, retention of primary fabrics in Precambrian dolostones compared to fabric destruction in most Phanerozoic dolostones suggests early dolomite formation in equilibrium with seawater, as opposed to late-stage dolomitization replacing primary calcite in reducing pore fluids (Tucker, 1982). Given this, Precambrian dolostones may be well suited to preserve primary seawater trends, analogous to calcites precipitated in Phanerozoic calcite seas (Stanley and Hardie, 1998). Regardless, the observation that I/(Ca + Mg) ratios are all below the detection limit from the carbonates ranging in age from 3.5 Ga, well before most estimates of initial O₂ production and accumulation, to 2.5 Ga (Table 1) is unlikely solely a relic of local diagenetic loss, arguing against appreciable and persistent iodine redox cycling throughout the Archean within the resolution of our data.

## Great Oxidation Event

We find the first convincing evidence for an oxidative marine iodine cycle in the ca. 2.45–2.4 Ga carbonates of the Turee Creek Group of the Hardey and Duck Creek synclines in Western Australia. The studied carbonates in the Hardey syncline are located above the glacial diamictite of the Meteorite Bore Member (Krapež, 1996; Bekker et al., 2002). The Meteorite Bore Member is correlated to the Ramsay Lake Formation of the Huronian Supergroup, Canada (Swanner et al., 2013), deposited prior to the permanent loss of NMD-S signal at ca. 2.32 Ga (Bekker et al., 2004). An absence of the Meteorite Bore Member in the Duck Creek syncline makes the age of these rocks less certain relative to glacial events, but these carbonates are inferred to precede deposition of the Meteorite Bore Member (Krapež, 1996). Though the timing and stratigraphic position of the loss of NMD-S signal in the Turee Creek Group is still unconstrained (Swanner et al., 2013), the correlation to other intervals with multiple sulfur isotope data suggests the potential for aerobic iodine cycling prior to the permanent loss of NMD-S signal. The timing of the increase in I/(Ca + Mg) is generally consistent with increasing enrichments in redox-sensitive trace metals such as Mo and Re in Neoarchean organic-rich shales linked to at least transient oxygen accumulation (Anbar et al., 2007). Surface ocean IO₃⁻ formation prior to the permanent loss of NMD-S signal could be explained with two scenarios: (1) dissolved O₂ in shallow seawater was sufficiently high, at least locally (perhaps in oxygen oases), for IO₃⁻ accumulation, while atmospheric O₂ was low enough for continued generation and preservation of the NMD-S signal (Pavlov and Kasting, 2002), or (2) both the atmosphere and surface ocean accumulated significant amounts of O₂, but the NMD-S fractionations found in deposits of this age were inherited from earlier atmospheric conditions through crustal recycling of sulfur (Reinhard et al., 2013). Regardless, the I/(Ca + Mg) record indicates surface ocean IO₃⁻ formation, implying a switch to more oxidizing surface ocean conditions at 2.45–2.4 Ga, broadly coincident with the beginning of the GOE (Lyons et al., 2014).

Resolving the significance of the observed rise in I/(Ca + Mg) during the GOE demands an understanding of modern iodine cycling. The main controls on [IO₃⁻] in surface waters are IO₃⁻ reduction either microbially or photochemically (Küpper et al., 2011), IO₃⁻ advection from oxic deep waters, and *in-situ* IO₃⁻ production (Luther et al., 1995). Independent geochemical evidence suggests a redox-stratified ocean before and after the GOE, with sulfidic anoxic conditions prominent along continental margins and Fe-rich anoxic conditions prevailing in the deeper ocean (Plavansky et al., 2011). A redox-stratified ocean would imply a GOE iodine cycle conceptually similar to that in the modern Black Sea where advection of iodine from deep waters is in the form of I⁻ (Luther and Campbell, 1991; Wong and Brewer, 1977), and with IO₃⁻ formation and accumulation up to 150 nM (Wong and Brewer, 1977; Luther and Campbell, 1991) via oxidation outpacing photochemical and biotic IO₃⁻ reduction in the surface waters (Truesdale et al., 2001). Oxidation of I⁻ to IO₃⁻ is biologically mediated in the modern ocean (Küpper et al., 2011), and oxidants
stronger than O₂, such as O₃ and H₂O₂, are considered to be essential to produce IO₃⁻ abiotically but cannot account for the abundance of IO₃⁻ in the modern ocean (Luther et al., 1995). Though the pathways of marine IO₃⁻ formation are still debated, IO₃⁻ is exclusively observed in well-oxygenated waters and is completely reduced to I⁻ with the onset of anoxic conditions (Wong and Brewer, 1977; Emerson et al., 1979; Wong et al., 1985; Luther and Campbell, 1991). These observations point to O₂ accumulation in the surface ocean and related IO₃⁻ stabilization as the driver for the first increase in I/(Ca + Mg) ratios during the GOE.

We place a tentative threshold for an aerobic iodine cycle in the surface ocean at 1 µM O₂ during the first rise of I/(Ca + Mg) during the GOE. Observations in the oxygen-minimum zone (OMZ) along the eastern coast of the north Pacific Ocean indicate oxygen concentrations ranging from ~225 µM near the ocean surface to as low as 1 µM in the core of the OMZ, with IO₃⁻ quantitatively reduced in the same interval (Rue et al., 1997). Importantly, these oxygen concentrations are not representative of the levels at which IO₃⁻ could form but rather the minimum at which we observe IO₃⁻ accumulation. Given this threshold of O₂ needed to stabilize IO₃⁻ in marine systems, the apparent absence of a fully oxic and persistent iodine cycle prior to the GOE does not conflict with other recent evidence for local oxidative processes and oxygenic photosynthesis in the Archean, albeit at still low levels. Recent studies have argued for aerobic nitrogen cycling in the surface ocean prior to the GOE (e.g., Thomazo et al., 2011) and local Mn oxidation as early as 3.0 Ga (Crowe et al., 2013; Planavsky et al., 2014), providing early evidence for some surface ocean and even atmospheric O₂ accumulation. The higher reduction potential (pE) for IO₃⁻/I⁻ compared to NO₃⁻/N₂ and MnO₂/Mn⁺² (Rue et al., 1997), the generally lower Gibbs free energy of reaction (ΔG<sub>rxn</sub>°) for IO₃⁻ reduction compared to Mn reduction and denitrification (Farrenkopf et al., 1997), and recent observations indicating modern aerobic Mn and N cycling at O₂ concentrations as low as sub-micromolar levels (Clement et al., 2009; Thamdrup et al., 2012) allow for the possibility of aerobic N and Mn cycles in the Archean without widespread IO₃⁻ or O₂ accumulation. In this light, the shallow Archean ocean might have been functionally anoxic, with its redox state lying between that required for coupled nitrification/denitrification and Mn oxidation and the somewhat higher state required for IO₃⁻ accumulation. A case consistent with this hypothesis comes from carbonates of the well-preserved 2.7 Ga Tumbiana Formation analyzed in this study, where previous work using nitrogen isotopes has led some workers to infer an aerobic nitrogen cycle (Thomazo et al., 2011), yet iodine concentrations are below detection.

## Lomagundi Event

The Lomagundi Event (LE) spans the ca. 2.22–2.1 Ga period and is marked by the largest and longest-lived positive carbonate carbon isotope ($\delta^{13}C_{carb}$) excursion in Earth history, with $\delta^{13}C_{carb}$ rising above +10‰ (Bekker and Holland, 2012). This exceptional $\delta^{13}C_{carb}$ trend has traditionally been interpreted to reflect an increase in relative rates of OM production and burial globally, causing the estimated release of 12–22× the modern atmospheric inventory of O₂ (Karhu and Holland, 1996). Our global record of individual carbonate sections with significant I/(Ca + Mg) ratios deposited during the LE stands in stark contrast to the older data and can be reasonably attributed to a shift to more oxidizing conditions (Table 1; Fig. 1). Other redox proxies, such as marine sulfate evaporites, carbonate-associated sulfate (CAS), and trace metal enrichments recorded in marine sediments, indicate increasing atmospheric and marine oxygenation during the same interval when our evidence points to surface ocean oxygenation (Schröder et al., 2008; Planavsky et al., 2012; Partin et al., 2013).

Starting during the LE, another potential control on I/(Ca + Mg) ratios is the size of the marine iodine reservoir during enhanced OM burial. For example, relatively low I/(Ca + Mg) ratios from carbonate successions deposited during Oceanic Anoxic Events (OAEs), in association with positive $\delta^{13}C_{carb}$ excursions, are interpreted to be the result of a shrinking

marine iodine reservoir during the times of enhanced OM burial—in combination with redox shifts and associated IO₃⁻ reduction (Lu et al., 2010). The Mcheka Formation shows a negative correlation between $\delta^{13}C$ values and I/(Ca + Mg) ratios during the final stage of the LE (Fig. 2). Similar to the OAE, increasing I/(Ca + Mg) ratios in tandem with a decrease in $\delta^{13}C_{carb}$ could result from an increasing marine iodine reservoir following the decline in OM burial. There is scatter and even blank values within the profile, suggesting currently unconstrained diagenetic controls. Nevertheless, the retention of CAS concentration in the same samples (Planavsky et al., 2012), which is a proxy extremely susceptible to diagenetic loss (e.g., Gill et al., 2008), and a lack of correlation between I/(Ca + Mg) and Mg/Ca ratios (Fig. DR1 in the GSA Data Repository<fnref n="1" />) suggest that the overall elevated shift in I/(Ca + Mg) ratios is primary.

<figure><img src="image_3.png" /><figcaption>**Figure 2.** $\delta^{13}C_{carb}$ values from Planavsky et al. (2012) and I/(Ca + Mg) ratios from this study for Mcheka Formation of Zimbabwe (see Table DR2 [see footnote 1]). Note apparent shift (delineated by arrows) to generally higher I/(Ca + Mg) ratios coincident with falling limb of Lomagundi Event. See text for further details. VPDB—Vienna Peedee belemnite.</figcaption></figure>

## CONCLUSION

Our results suggest the onset of an oxidative surface ocean iodine cycle during the GOE. Prior to this, redox gradients may have existed in the Archean that allowed for trace and/or transient oxygen levels to support aerobic N and Mn cycles, but the present data suggest an environment dominantly hostile to IO₃⁻ accumulation and thus one that was mostly anoxic. During the Lomagundi excursion, a general increase in carbonate units with nonzero I/(Ca + Mg) ratios implies widespread IO₃⁻ stability in the surface ocean related to O₂ concentrations above 1 µM. Although large spatiotemporal gaps still exist for the I/(Ca + Mg) record, our results provide a proof of concept, highlighting the proxy's promise even in very old carbonate rocks with varying degrees of alteration. The importance of this approach is elevated significantly by its relevance to shallow, carbonate-dominated settings, which are not well suited to interrogation by most other paleoredox proxies.

## ACKNOWLEDGMENTS

We thank Woodward Fischer for providing samples crucial to this work. Funding to Lyons from the NASA Exobiology Program, the NASA Astrobiology Institute, and the National Science Foundation (NSF) supported this research. Lu is supported by NSF grant OCE 1232620. Bekker's participation was made possible by funding provided via a Natural Sciences and Engineering Research Council of Canada Discovery Grant. Philippot acknowledges support from UnivEarthS LabEx

<footnote n="1">¹GSA Data Repository item 2014228, age references, data, and additional information regarding Turee Creek and Mcheka carbonates, is available online at www.geosociety.org/pubs/ft2014.htm, or on request from editing@geosociety.org or Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301, USA.</footnote>
(grants ANR-10-LABX-0023 and ANR-11-IDEX-0005-02). This manuscript also benefited greatly from the insightful comments of three anonymous reviewers.

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