Review on the physical chemistry of iodine transformations in the oceans

Abstract

The transformation between iodate (), the thermodynamically stable form of iodine, and iodide (I^-), the kinetically stable form of iodine, has received much attention because these species are often dependent on the oxygen concentration, which ranges from saturation to non-detectable in the ocean. As suboxic conditions in the ocean’s major oxygen minimum zones indicate that is minimal or non-detectable, the incorporation of into carbonate minerals has been used as a redox proxy to determine the O₂ state of the ocean. Here, I look at the one and two electron transfers between iodine species with a variety of oxidants and reductants to show thermodynamics of these transformations. The to conversion is shown to be the controlling step in the reduction reaction sequence due to thermodynamic considerations. As reduction to is more favorable than reduction to at oceanic pH values, there is no need for nitrate reductase for reduction as other reductants (e.g. Fe²⁺, Mn²⁺) and dissimilatory reduction by microbes during organic matter decomposition can affect the transformation. Unfortunately, there is a dearth of information on the kinetics of reductants with ; thus, the thermodynamic calculations suggest avenues for research. Conversely, there is significant information on the kinetics of I^- oxidation with various oxygen species. In the environment, I^- oxidation is the controlling step for oxidation. The oxidants that can lead to are reactive oxygen species with O₃ and •OH being the most potent as well as sedimentary oxidized Mn, which occurs at lower pH than ocean waters. Recent work has shown that iodide oxidizing bacteria can also form . I^- oxidation is more facile at the sea surface microlayer and in the atmosphere due to O₃.

1 Introduction

The thermodynamically favorable form of iodine in seawater is iodate (). However, iodide (I^-) is present in oxic, suboxic and anoxic waters. The one electron transfer reaction of I^- with molecular oxygen, ³O₂, to form the iodine atom (I•) and superoxide (O₂^-) is thermodynamically unfavorable as is the reaction of two I^- with ³O₂ to form I₂ and H₂O₂ (Luther et al., 1995; Luther, 2011). Thus, other oxidants are required to initiate abiotic iodide oxidation, and I^- is a known sink for O₃. Biotic iodide oxidation has received much interest with one report showing conversion of iodide to iodate (Hughes et al., 2021). Iodate reduction can occur with common reductants (e.g., sulfide, Fe²⁺), and various organisms that decompose organic matter using iodate as the electron acceptor. Nitrate reductase and dimethyl sulfoxide reductase enzymes from these and planktonic microbes are considered important mediators for biotic iodate reduction (e.g., Hung et al., 2005; Amachi, 2008). Thus, there has been extensive interest in the chemistry of these two iodine species and the possible intermediates that form during their 6-electron redox interconversion ever since the element, iodine, was first discovered as I₂ during the study of brown kelp algae of the Laminariales (kelps/seaweeds) by Courtois in the early 1800s (Wong, 1991; Küpper et al., 2008; Küpper et al., 2011).

Possible chemical species that form during the interconversion are given in eqn. (1a, b). The loss of an O atom is equivalent to a two-electron transfer resulting in the reduction of the iodine from +5 in to +3 for (HOIO) to +1 for HOI (IO^-) and to -1 for I^-. The acid-base species in parentheses are minor species at seawater pH as the pK_a values for HOI and HOIO are 10.7 (K_a = 2 x 10^-11) and 4.49 (K_a = 3.2 x 10^-5), respectively.

Equation 1b shows the interconversion between HOI and I^- as HOI undergoes one-electron transfer to I₂ followed by another one-electron reduction per I atom to I^-.

This work considers the thermodynamics of these transformations during the reduction of , which occurs in anoxic systems, during organic matter decomposition and by phytoplankton, as well as the oxidation of I^-, which occurs by the direct oxidation of iodide by iodide oxidizing bacteria, oxidized metals and reactive oxygen species (ROS) that are produced by certain microbes, (macro)algae and abiotic processes including photochemistry. In a previous work (Luther, 2011), the chemistry and thermodynamics of chloride, bromide and iodide oxidation were compared; however, I(+3) species (HOIO, iodous acid, and , iodite) and stepwise iodate reduction were not considered. Here, stepwise reactions of the iodine species in equations 1a and 1b with environmentally important reactants (including transient ROS species) are considered for both the oxidation of I^- and the reduction of to affect their interconversion. The kinetics of these stepwise reactions are also considered. Kinetic data for the first step(s) in iodide oxidation are available, but less kinetic information is available for iodate reduction.

2 Methods

2.1 Calculations of aqueous redox potentials from half-reactions

Table 1 gives several equations for redox half-reactions that include the pH dependence for the reaction considered. These are pϵ(pH) relationships based on the balanced chemical equations and the thermodynamics of each chemical species. The basic mathematical approach has been fully developed in standard textbooks (Stumm and Morgan, 1996; Luther, 2016) and used in previous publications (Luther, 2010; Luther, 2011). Aqueous thermodynamic data to calculate the pϵ(pH) or logK(pH) relationships in Table 1 (at 25°C and 1 atm) are from Stumm and Morgan (1996) and other sources (Bard et al., 1985; Stanbury, 1989). The value used for the Gibbs free energy for Fe²⁺ (-90.53 kJ/mole) is that discussed in Rickard and Luther (2007). Values of the free energy for HOIO and are from Schmitz (2008).

The calculated pϵ value from each half-reaction is given as a function of pH as in the examples in Table 1, and these half reactions can be used for simple calculations of the pE values of full reactions (see next section). When H⁺ or OH^- is not in a balanced equation for a half-reaction, there is no pH dependence on the half-reaction. The pε calculated is termed pϵ(pH) which provides a log K for each half-reaction at a given pH. Concentration dependence for the other reactants are not considered in the calculation; thus, these are considered standard state calculations. When concentration dependence is considered, the calculated pϵ value can vary as in the following example for the O₂/H₂O couple (O1 in Table 1).

Using the balanced half reaction and the Gibbs free energy of formation of each species at 25°C, the Gibbs free energy of the reaction and the equilibrium constant are calculated.

The standard state ΔG° for the reaction = - 490.68 kJ/2 moles H₂O or 4 moles of electrons. The equilibrium constant () is given in eqn. 2a where {} indicates activity for each chemical species and the activity of H₂O is defined as 1.

On expanding, eqn. 2b results.

is for 4 mole of electrons or 21.50 for 1 mole of electrons.

For a one-electron half-reaction, we have

And equation 2b becomes equation 2c

or

From the Nernst Equation, (the standard state value), which on substitution gives equation O1 (see Table 1) where concentration is used for O_2(aq).

At ocean surface conditions of 211 µM O₂ (211 x 10^-6 M; 100% saturation at 25°C and salinity of 35), this expression becomes

and at a pH of 8, pε = 12.58.

At 1µM O₂ (10^-6 M) which occurs in oxygen minimum zones, this expression becomes

and at a pH of 7.5, pε = 12.50.

At unit activity for all reagents including H⁺, pε = pε°. At unit activity of all reagents other than the H⁺, equation O1a results, which is used for many calculations in this paper.

Note that the above equations show a 1.50 log unit change for an O₂ concentration range from 1 µM to unity activity (O1a) so the calculations could vary an order of magnitude or more in either direction when concentration dependence is included. However, comparisons can be more easily made when combining different half-reactions at a given pH. This permits an assessment of which combined half-reactions are thermodynamically favorable and thus more likely to occur in a given environmental setting.

2.2 Coupling half-reactions

As an example of coupling two half reactions to determine whether a reaction is favorable, I use the data in Table 1 for the reduction of (Io5b) by (N1) in equation 3.

Equation 4 is used to calculate a complete reaction’s pε or ΔlogK_reaction value. All values of ΔlogK_reaction > 0 indicate a favorable reaction and all values of ΔlogK_reaction< 0 indicate an unfavorable reaction.

At a pH of 7, the pε_red values for and are 7.91 and 7.28, respectively. As is the reductant, it is oxidized; thus, the sign for pε_red (7.28) is reversed to become pε_oxid (-7.28).

For reaction 3, there is no pH dependence as the pH dependence of each half-reaction is similar so cancels.

For this work, Table 1 lists the pϵ(pH) values for Mn, Fe, oxygen, nitrogen, sulfur and iodine species for the relevant iodine redox reactions considered. Dissolved Fe(II) and Mn(II) are primarily hexaaquo species until the pH is > 7, where hydroxo complexes start to become important. As most reactions occur via one and two-electron transfers, the calculations will permit assessment of a thermodynamically unfavorable step along a reaction coordinate of six-electrons as in the reduction of iodate to iodide and the oxidation of iodide to iodate. From surface waters to decomposition zones, seawater pH values range from 8 down to 7; thus, the following discussion will emphasize this pH range.

3 Results and discussion: Iodate reduction

3.1 Iodate and iodide speciation at different seawater oxygen conditions

In the oxic environment, the oxidizing condition of the environment or pε is set by the 4-electron transfer reaction of the O_2(aq)/H₂O couple [reaction O1 in Table 1]. At a pH of 8, temperature 25⁰C and a salinity of 35, 100% O_2(aq) saturation is 211 μM, which gives a pε of 12.58 (Figure 1). As the couple has a pε of 10.56 at is the thermodynamically favored iodine species.

Figure 1

Entering the pε value for a given [O_2(aq)] into equation Io6 allows the determination of the iodide to iodate ratio and the actual concentration of each assuming a total iodine concentration of 450-470 nM (Elderfield and Truesdale, 1980). Figure 1 shows the iodate and iodide concentrations are equivalent at a pε of 10.56. The vertical lines indicate the environmental pε for [O_2(aq)] of 1, 10, 100 nM, and 1, 50 and 211 μM. As oxygen minimum zones (OMZ) of the Arabian Sea and the equatorial Pacific Ocean have [O_2(aq)] concentrations in the 1-100 nM range (Revsbech et al., 2009; Lehner et al., 2015), calculations show that is the thermodynamically preferred iodine species even at 1 nM O_2(aq), which gives a pε of 11.25 for the O_2(aq)/H₂O couple. However, I^- is the dominant iodine species detected in OMZ waters (Wong and Brewer, 1977; Luther and Campbell, 1991; Rue et al., 1997; Farrenkopf and Luther, 2002; Cutter et al., 2018). At [O_2(aq)] concentrations ≤ 1 μM, , and Mn²⁺ concentrations are now similar or higher in concentration and should determine the pε of the water.

As most reactions occur by 1- or 2-electron transfers, Figure 2 shows the redox sequence for two electron transfer redox couples , MnO₂/Mn²⁺ (Mn1) and the one-electron redox couple Fe(OH)₃/Fe²⁺ (Fe3) over a wide range of pH. The redox sequence at pH = 8 is as expected for the N, Mn and Fe systems. As (up to 12 μM) and Mn²⁺ (up to 8 μM) are formed at OMZ oxic-anoxic transition zones (e.g., Arabian Sea, Black Sea, Equatiorial Pacific, see Lewis and Luther, 2000; Trouwborst et al., 2006; Cutter et al., 2018, respectively) and are in higher concentration than O_2(aq), the and MnO₂ → Mn²⁺ (Mn1) couples can be chosen to set the environmental pε. At pH = 8, the to pε is 6.15 and the MnO₂ to Mn²⁺ pε is 4.80. At pH = 7, the to pε is 7.28 and the MnO₂ to Mn²⁺ pε is 6.80. At these pε values, O_2(aq) is below 1 nM, and I^- is now the thermodynamically favored iodine species when comparing these data with the couple (pε of 10.56 at pH = 8).

Figure 2

Figure 2 also shows that the couple (Io5b) should be the first step in the reaction sequence of iodate to iodide (eqn. 1a). As the pε of the couple has a more positive pε value than the N, Mn and Fe couples in Figure 2, reduction is more favorable than these couples even though it is very close to the couple. Thus, is predicted to reduce before , and biological activity (e.g., nitrate reductase activity) is not necessary to reduce (see section 3.2). Because of the strong pH dependence for the Mn and Fe couples, they cross the and couples at lower pH, which have similar slopes. Thus, is predicted to reduce to (eqn. 3). Interestingly, Mn²⁺ and Fe²⁺ should be poorer reductants than for conversion of to at a pH< 6 and pH< 1, respectively, but are more favorable to reduce than above those pH values (see sections 3.2 and 3.3).

Although the to I^- conversion occurs at higher pε, it is a 6-electron transfer (IO6), which is not a facile process. Thus, the intermediates ( and HOI) will dictate the reactivity sequence via a combination of thermodynamic and kinetic considerations.

As shown in Figure 2, reduction to HOI and HOI reduction to I^- are also more favorable at higher pε values than the to I^- couple. At a pH = 8, the to HOI couple has a pε value of 12.06 corresponding to 2 μM O₂ (see Figure 1). Similarly, the HOI to I^- couple has a pε value of 12.66 corresponding to 250 μM O₂. At a pH of 7, both couples have pε values greater than 13 indicating that, even at O₂ saturation, I^- is the dominant species predicted when these intermediates form. At a pH = 7.5 (that is found in many OMZ waters), both to HOI and HOI to I^- couples have pε values greater than 12.6; also indicating that at O₂ saturation, I^- is the dominant species predicted. Thus, the intermediates and HOI are not predicted to be stable in marine waters; thus, the conversion of to is a key step. Interestingly, Hardisty et al. (2021) found in situ reduction in the oxycline where [O_2(aq)] was 11 μM, but not at [O_2(aq)]< 2 μM. Lastly, the O₃ to O₂ + H₂O couple is highly oxidizing indicating that all iodine couples should lead to formation. O₃ reactions will be discussed in more detail below (sections 4.2, 4.7).

In the next sections (3.2 – 3.5), the thermodynamics for the conversion of iodate to iodide via the intermediates outlined in equations 1a and 1b by environmental reductants are considered to show what step, if any, in the reduction of iodate to iodide may be unfavorable over a wide range of pH. Iodate reduction is well known in the marine environment (e.g., Wong and Brewer, 1977; Wong et al., 1985; Luther and Campbell, 1991; Rue et al., 1997; Farrenkopf and Luther, 2002; Cutter et al., 2018) and occurs via chemical reductants like sulfide (Zhang and Whitfield, 1986) and via microbes like Shewanella putrfaciens (Farrenkopf et al., 1997) and Shewanella oneidensis (Mok et al., 2018) during dissimilatory reduction coupled with decomposition (oxidation) of organic matter as well as phytoplankton mediated processes (e.g., Chance et al., 2007).

3.2 Iodate reduction by NO₂^-

Although has not yet been shown to be a reductant for in aqueous lab studies (eqn. 3), HNO₂ is a reductant for MnO₂ (Luther and Popp, 2002) and Mn(III)-pyrophosphate (Luther et al., 2021). Figure 3A shows the thermodynamic calculations for the stepwise conversion of to I^- by reduction ( oxidizes to ). All 2-electron transfer reactions, which involve O atom loss for iodine, are favorable over the pH range. For seawater pH (7-8), the least favorable reaction is the to reaction whereas the to HOI and HOI to I^- reactions are more favorable. Thus, the to conversion appears to be the controlling step in the reaction sequence. The 1-electron transfer reaction of HOI to I₂ is the most favorable, but the second 1-electron transfer reaction of I₂ to I^- is only favorable at pH > 4. Thus, reduction of to I^- by is predicted via 1-electron or 2-electron transfer reactions at seawater pH values. The data plotted in Figure 2 indicate that once forms there is no thermodynamic barrier to I^- formation.

Figure 3

The reaction with has been reported to produce I₂ in ice by Kim et al. (2019), but not in solution. The pH in the ice was 3 where HNO₂ and H₂ONO⁺ exist and are the likely reductants. Thus, polar areas may be locales for reduction. At seawater pH, the reaction seems to be hindered by kinetics in the transition state as each reactant ( and ) is an anion, which will repel each other.

3.3 Biological iodate reduction

The marine literature has many reports on the uptake of (with or without ) by phytoplankton with the iodine released as I^- (e.g., Elderfield and Truesdale, 1980; Wong, 2001; Wong et al., 2002; Chance et al., 2007; Bluhm et al., 2010). As a result of this iodate uptake, reductase was presumed by some researchers to be a key process for reduction to I^-. Also, Bluhm et al. (2010) and Carrano et al. (2020) showed that algal senescence enhanced I^- release, and Hepach et al. (2020) showed that there is a considerable lag between uptake and I^- release due to senescence.

reductase appears to reduce in some phytoplankton (Hung et al., 2005). However, de la Cuesta and Manley (2009) showed that I^- can be up taken by phytoplankton, and that different phytoplankton uptake I^- whereas other phytoplankton uptake . Thus, there is no need for nitrate reductase for reduction as I^- can be up taken by some phytoplankton rather than form from reduction. Moreover, Waite and Truesdale (2003) showed that nitrate reductase was not important for reduction by Isochrysis galbana. The latter study is consistent with the thermodynamics of the reduction to being more favorable than the reduction to .

Furthermore, under anaerobic conditions, dissimilatory reduction occurs without nitrate reductase for the denitrifying bacterium, Pseudomonas stutzeri, (Amachi et al., 2007; Amachi, 2008). Reyes-Umana et al. (2022) and Yamazaki et al. (2020) showed that iodate reductase is in the periplasmic space of Pseudomonas sp SCT. Also, Mok et al. (2018) showed that dissimilatory reduction by Shewanella oneidensis does not involve nitrate reductase. Recently, Shin et al. (2022) showed that Shewanella oneidensis requires extracellular dimethylsulfoxide (DMSO) reductase involving a molybdenum enzyme center for reduction. Guo et al. (2022) studied bacterial genomes in a variety of environments and documented that Shewanella oneidensis are ubiquitous in all fresh and marine waters; they concluded that reduction is a major biogeochemical process. Thus, nitrate reductase (also an O atom transfer reaction) is not a requirement for bacterial reduction to I^-.

The interconversion of dimethylsulfoxide with dimethylsulfide during dissimilatory reduction is another 2-electron O-atom transfer reaction. Moreover, the reactions of DMS to reduce HOI, and are thermodynamically favorable (Figure 3B, DMSO reduction is in S3, Table 1 and occurs at a lower pε than and reduction). The reaction of DMS with HOI has been suggested by Müller et al. (2021) to be a sink for DMS based on the rapid reaction of DMS with HOBr.

3.4 Iodate reduction by Mn²⁺ and Fe²⁺

Figure 4 shows the thermodynamics for the stepwise conversion of to I^- by reduction with Mn²⁺ and Fe²⁺. Concentrations of Mn²⁺ and Fe²⁺ range from several nM to μM in OMZs (e.g., Trouwborst et al., 2006; Moffett and German, 2020) and in suboxic porewaters (e.g., Oldham et al., 2019; Owings et al., 2021) to mM in waters emanating from hydrothermal vents (e.g., Estes et al., 2022); in these cases, Mn²⁺ and Fe²⁺ are normally higher in concentration than the total iodine concentration. For the 2-electron transfer reactions with Mn²⁺, only the to HOI reaction is favorable over the entire pH range. The to reaction is favorable only at pH > 6 whereas the other reactions are favorable at pH > 3. Thus, the to conversion is the controlling step in the reaction sequence when Mn²⁺ is the reductant. Using high resolution porewater profiles of I^- and Mn²⁺ obtained by voltammetric microelectrodes, Anschutz et al. (2000) showed that a I^- maximum occurred at the depth where upward diffusing Mn(II) was being removed and proposed that I^- formed by the reaction of with Mn²⁺ under suboxic conditions. The reaction has not been investigated in laboratory studies.

Figure 4

For the reaction sequence with Fe²⁺, all iodine species reductions are favorable over the entire pH range except for the I₂ to I^- reaction, which is favorable at pH > 2.5. Thus, there is no thermodynamic inhibition to reduction to I^- by Fe²⁺, and this abiotic reaction at a pH of 7 was reported to be 92% complete after 2 hours using initial concentrations of 2 mM Fe²⁺ and 0.1 mM (Councell et al., 1997). Because the Fe(OH)₃ to Fe²⁺ couple is a 1-electron transfer, two Fe²⁺ are required in each step of the sequence. Again, the to conversion is the least favorable and likely controlling step in this reaction sequence.

Comparing Figures 3, 4 indicates that the Mn²⁺ and reactions with iodine species have a similar range of ΔlogK_reaction values whereas the Fe²⁺ reactions with iodine species are more favorable (higher ΔlogK_reaction values).

3.5 Iodate reduction by sulfide

In sulfidic waters and porewaters, does not exist as sulfide reacts readily with it (Zhang and Whitfield, 1986), and S(0) forms as the initial sulfur product. Figure 5 shows the thermodynamics for the stepwise conversion of to and to HOI by sulfide where S(0) forms as an intermediate leading to S₈. As the Gibbs free energy of formation for HSOH is unknown, HSOH could not be evaluated as an intermediate, which on continued oxidation would form . The reaction of sulfide with I₂ and HOI is well known as the iodometric titration, so calculations were not performed. The only unfavorable iodine reduction reactions are the 1-electron reductions that lead to the formation of the HS radical (HS• or HS rad). The conversion of to HOI is more favorable as it has the larger ΔlogK_reaction values. Again, the to conversion is the least favorable and likely controlling step in this reaction sequence.

Figure 5

3.6 Iodate reduction by NH₄⁺

Figure 6 shows the thermodynamics for the stepwise conversion of to by where hydrazine (N₂H₄) and hydroxylamine (NH₂OH) as well as their protonated forms could form as the first N intermediates. The thermodynamic calculations for these 2 electron transfers indicate that these reactions are not favorable. However, the reaction of the intermediates, if they could form by other processes, with to form N₂ is very favorable. Thus, the to conversion is the controlling step in the reaction sequence with .

Figure 6

4 Results and discussion: Iodide oxidation

4.1 Iodide oxidation by NO₃^-, MnO₂ and Fe(OH)₃

Figures 3–6 showed the reduction of with reductants. All values of ΔlogK_reaction > 0 indicate a favorable reaction and all values of ΔlogK_reaction< 0 indicate an unfavorable reaction. These figures can be used to discuss the reverse reaction of I^- oxidation with oxidants. For reverse reactions, when a ΔlogK_reaction< 0, then I^- oxidation is favorable, but when ΔlogK_reaction > 0, I^- oxidation is unfavorable.

In Figure 3, is not an oxidant for I^- (reverse of the and I₂ reaction) except for the formation of I₂ at a pH< 4.

In Figure 4, MnO₂ oxidizes I^- to HOI (reverse of the Mn²⁺ and HOI reaction) at a pH< 3 and I^- to I₂ (reverse of the Mn²⁺ and I₂ reaction) at a pH< 5. A couple of laboratory studies showed I^- oxidation with synthetic birnessite (δ-MnO₂). First, Fox et al. (2009) showed that I₂ was produced over the pH range 4.50 – 6.25, and that formed in smaller amounts. The kinetics of the reaction were slower at higher pH by 1.5 log units (> 30-fold) and were slower when smaller amounts of MnO₂ were added (Table 2). Allard et al. (2009) investigated the same reactants to a pH of 7.5 and found I₂ and as products; above pH = 7 the reaction is very slow. Iodate was found mainly in lower pH waters. Both I₂ and adsorb to the birnessite surface. Similar results have been found over the pH range 4-6 for Mn(III) solids (Szlamkowicz et al., 2022). These MnO_x reactions with I^- are much slower that the reactions with reactive oxygen species (Table 2). Nevertheless, these are important as Kennedy and Elderfield (1987a, 1987b) showed that the conversion of iodide to iodate occurred in marine sediments.

Figure 4 also shows that I^- oxidation by Fe(OH)₃ to I₂ (reverse of the Fe²⁺ and I₂ reaction) should occur only at a pH< 2.5. The I^- to HOI conversion (reverse of the Fe²⁺ and HOI reaction) is favorable at pH ≤ 0.5.

4.2 Iodide oxidation by oxygen species

Figure 7 shows the thermodynamics of I^- oxidation to I₂ by oxygen species. Figure 7A shows that the one-electron process for I^- oxidation with ³O₂ is thermodynamically unfavorable over all pH whereas Figure 7B shows that the two-electron process is favorable at a pH< 3. Figure 7A shows that the successive 1-electron oxidations of I^- where superoxide () is reduced to hydrogen peroxide (H₂O₂), which is reduced to hydroxyl radical (•OH). Only •OH is thermodynamically favorable over the pH range considered. and H₂O₂ show favorable reactions at pH< 9 and pH< 6, respectively.

Figure 7

By contrast, Figure 7B shows that the reactions of I^- with the 2-electron oxidants H₂O₂, ¹O₂ and O₃ are all thermodynamically favorable. The likely reaction pathway is the loss of 2-electrons to produce I⁺, which then reacts with I^- to form I₂. Note that H₂O₂ reacts to form H₂O not •OH in Figure 7B. The 2-electron reaction with O₃ (Figure 7B) is more favorable than the 1-electron reaction (Figure 7A).

Wong and Zhang (2008) showed that H₂O₂ oxidizes I^- in artificial seawater from pH 7-9, which is consistent with Figure 7B. However, I^- oxidation does not lead to iodate. In fact, I^- reforms. They proposed that I₂ formed and was reduced back to I^-, but they did not provide a mechanism. The reverse reaction of I₂ with •OH (ΔlogK_reaction< 0 in the plot) is favorable to reform H₂O₂ and I^- at pH > 6 whereas the reverse reaction of H₂O₂ with I₂ to reform and I^- is favorable at a pH > 9 (ΔlogK_reaction< 0 in the plot). These thermodynamic data indicate that H₂O₂ can form I₂ in a 2-electron transfer (Figure 7B) and then reduce I₂ to I^- in a 1-electron transfer (Figure 7A).

At seawater pH, superoxide, , can oxidize I^- to I₂ and the reaction occurs with a rate constant of 10⁸ M^-1s^-1 (Bielski et al., 1985; Table 2). Because I₂ is a good electron acceptor, the subsequent reaction of with I₂ leads to (Schwarz and Bielski, 1986). As to be discussed in section 4.3, I₂ reacts with organic matter to form organo-iodine compounds. Extracellular is generated by Roseobacter sp. AzwK-3b (Li et al., 2014) and results in the oxidation of Mn²⁺ to Mn(III,IV) oxides. However, Li et al. (2014) found that also oxidized I^-. Considering that extracellular formation is a widespread phenomenon among marine and terrestrial bacteria, this could represent an important first step in the pathway for iodide oxidation in some environments. The Mn oxides formed by Roseobacter sp. AzwK-3b are not the oxidant as MnO₂ kinetics is slower (Table 2).

To obtain , further oxidation of I₂ to HOI must occur, and •OH is one candidate with a rate constant of 1.2 x 10¹⁰ (Buxton et al., 1988; Table 2). Also, O₃ has a rate constant of 1.2 x 10⁹ (Liu et al., 2001).

I₂ is a prominent intermediate in I^- oxidation yet HOI is needed to form . HOI can form directly from I^- and I₂ oxidation or from hydrolysis of I₂ (reverse of eqn. 5), which is fast at basic pH (Wong, 1991). Figure 8A shows that of the successive 1-electron oxidants (starting from O₂) for I₂ oxidation, only •OH is thermodynamically favorable over all pH to form HOI whereas O₃ is favorable at pH > 6, and is favorable at pH< 6. H₂O₂ as a 1-electron oxidant cannot oxidize I₂ to form HOI, but H₂O₂ can reduce HOI to I₂ (reverse of the and I₂ reaction). Figure 8B indicates that, as 2-electron oxidants, H₂O₂ and O₃ oxidation can lead to HOI formation. Comparing Δlog K values in Figures 7, 8 indicates that oxidation of I₂ to HOI is less favorable than the oxidation of I^- to I₂.

Figure 8

These data also indicate why the comproportionation reaction of HOI with I^- to form I₂ can occur (eqn. 5, Carpenter et al., 2013).

Although disproportionation of HOI to and I^- (eqn. 6) is fast in strongly basic solution, it is not detectable at seawater pH (Wong, 1991).

Figure 9 shows the successive 2-electron oxidation reactions of I^-, HOI and with ³O₂, ¹O₂, H₂O₂ and O₃. ³O₂ cannot affect the oxidation at any pH. Figure 9 shows that O₃ oxidation reactions with I^-, HOI and are favorable; thus, O₃ can affect the complete oxidation of I^- to . Also, the H₂O₂ oxidation reactions of I^-, HOI and are favorable and can lead to formation; however, the kinetics of H₂O₂ oxidation can be slow. Haloperoxidase enzymes from organisms enhance the kinetics (Butler and Sandy, 2009) as does the reaction of H₂O₂ with carboxylic acids secreted by microbes to form peroxy carboxylic acids, which in turn oxidize I^- to I₂ (Li et al., 2012). The reactive oxygen species ¹O₂ can oxidize I^- at pH< 10, oxidize HOI at pH > 5, and over all pH. Thus, ¹O₂ can be an oxidant of I^- to at seawater pH. These data indicate that HOI oxidation leads to formation.

Figure 9

Interestingly, ΔlogK values in Figure 9A show that the thermodynamics of I^- oxidation by the 2-electron oxidants O₃ and H₂O₂ to form HOI is slightly less favorable than I₂ formation (Figure 7B). Conversely, thermodynamics of I^- oxidation by H₂O₂ as a 2-electron oxidant to form HOI (Figure 9A) is more favorable than I₂ formation (Figure 7A).

As shown in Figure 10, a potentially potent oxidant for I^- is N₂O, which is an O atom transfer oxidant like O₃. However, the N₂O concentration in seawater is minor, but the largest reported values are 90 and 250 nmol kg^-1 for the OMZs of the Arabian Sea (Freing et al., 2012) and the Eastern Tropical North Pacific (Damgaard et al., 2020), respectively. These values are smaller than the total iodine concentration in seawater. The N₂O concentration in the atmosphere is 335 ppbv (August 2022, https://www.n2olevels.org), which is equivalent to 0.0331 Pa or 7.8 nM dissolved in surface seawater (salinity of 35) at 20 ⁰C using the solubility data from Weiss and Price (1980).

Figure 10

4.3 ROS in seawater

Reactive oxygen species exist in marine waters, but at low concentrations. O₃ penetrates a few micrometers through the water-air interface at surface iodide concentrations (Carpenter et al., 2013). Powers and Miller (2014) showed that solar-induced processes with organic matter in freshwater and seawater are a major source of ROS (as , H₂O₂, and •OH) with the inventory and production rates for H₂O₂ in surface seawater being highest of the ROS. Also, Sutherland et al. (2020) report that dark, extracellular production is prolific among marine heterotrophic bacteria, cyanobacteria, and eukaryotes. In surface ocean waters, the concentration of H₂O₂ ranges from 20 - 80 nM (Yuan and Shiller, 2001), biological production gives a total concentration of ~ 0.07 to 0.30 nM (Sutherland et al., 2020), •OH concentration is ~10^-18 M (Mopper and Zhou, 1990), and ¹O₂ concentration ranges from 10^-13 to 10^-14 M (Sunday et al., 2020). However, these ROS concentrations are typically smaller than I^- concentrations, which range from 10 to 200 nM (Chance et al., 2019). Thus, I^- oxidation in seawater samples should be difficult to observe experimentally. Hardisty et al. (2020) tracked the addition of stable isotopes of iodide in sample incubations and report the rate of I^- oxidation to be 118–189 nM yr^-1, which is similar to rates reported by mass balance approaches (Campos et al., 1996a; Truesdale et al., 2001; Žic and Branica, 2006; Žic et al., 2008). Hardisty et al. (2020) report that the product is likely HOI that results in the formation of organic-iodine compounds (see section 4.6) which on decomposition can release I^-.

As the surface concentrations of ROS are smaller than the I^- concentration, the question is how does I^- get oxidized toin seawater? Microbial processes and the oxidation of I species in the atmosphere by ROS are likely candidates. These are now discussed.

4.4 Iodide oxidation in brown kelp

Brown kelp are the strongest accumulators of iodine as I^- among living organisms (up to 100 mM, Küpper et al., 2008). The element iodine was discovered by the formation of I₂ during exposure of brown kelp to concentrated sulfuric acid, which oxidized I^- to I₂. Kelp releases I^- on the thallus surface and in the apoplast when undergoing oxidative stress during the partial emersion of the brown kelp forest at low tide; e.g., by exposure to high irradiance, desiccation, and atmospheric O₃. Kelp contain vanadium haloperoxidases (Colin et al., 2003; Küpper et al., 2008) that enhance I^- oxidation by H₂O₂. Whereas the nonenzymatic reaction of I^- with H₂O₂ is slow, the reactions with O₃, , ¹O₂, and •OH are very fast (> 10⁸ M^-1s^-1, Table 2); they are also faster than MnO₂ oxidation of I^-. Küpper et al. (2008) consider I^- as the simplest antioxidant known.

4.5 Iodide oxidation by microbes

Hughes et al. (2021) report that production occurs in cultures of the ammonia-oxidizing bacteria Nitrosomonas sp. and Nitrosococcus oceani supplied with I^-, but not in cultures of three different nitrite oxidizing bacteria. Information on the enzymes mediating the oxidation were not studied. Nevertheless, oxidation via nitrification occurs via NH₂OH formation which is a 2-electron reaction. Further reaction of NH₂OH via metalloenzymes (e.g., Mo and W oxidases that transfer O atoms) leads to (a 4-electron transfer) and . I^- likely goes through the intermediates and HOI to form , but these intermediates are reactive and not detectable by present analytical methods unlike .

Consistent with the Hughes et al. (2021) report, Kennedy and Elderfield (1987a, 1987b) showed that the conversion of iodide to iodate occurred in marine sediments. Microbial intervention is likely, but reaction of I^- with oxidized Mn is possible depending on the pH.

Amachi and Iino (2022) reviewed the genus Iodidimonas, which was originally found in brines, but was also cultured from seawater enriched with I^-. I₂ is the first oxidation product. Iodidimonas contains the iodide oxidizing enzyme (IOX), which is an extracellular protein that contains multicopper oxidases. Iodidimonas requires O₂, not H₂O₂, as the electron acceptor. Other oxidants are required to oxidize I₂ to with O₃ and •OH being the most effective.

4.6 HOI and I₂ formation leads to organic iodine

Competing with the inorganic interconversion between iodide and iodate is the formation of organic iodine compounds. Formation of C-I bonds can occur during the reduction of and oxidation of I^-. Complete reduction of to I^- does not need to occur intercellularly and can lead to HOI and I₂ formation as in Figure 4. The first step in I^- oxidation also leads to I₂ and HOI. I₂ is neutral and adds to organic compounds such as olefins, which are not very reactive in seawater, whereas I⁺ in HOI reacts with α-keto compounds and peptides through keto-enol isomerization (Truesdale and Luther, 1995). Both I₂ and HOI lead to volatile and nonvolatile organic-iodine (R-I) compounds with C-I or N-I bonds, and Harvey (1980) showed that N-iodo amides were the main organic iodine components in marine sediments. On decay of organic compounds, the C(N)-I bond breaks leading to I^- release, which mimics the senescence pathway outlined by Bluhm et al. (2010) and Hepach et al. (2020). Recently, Ooki et al. (2022) showed that CH₃I and CH₃CH₂I formed in sediments from polar and subpolar seas and was related to increased phytodetritus at the seafloor after the spring bloom.

Allard and Gallard (2013) showed that the oxidation of I^- by birnessite in the presence of organic matter also led to CH₃I over the pH range 4-5.

As total iodine in surface ocean waters is lower by a few percent compared to deep waters (Wong, 1991), the decomposition of organic-iodine leads to some I^- release, which may be oxidized to by ammonia-oxidizing bacteria (Hughes et al., 2021). This is similar to release and oxidation of to from particulate organic matter in deep waters that results in an increase of concentration with depth (recycled element profile). Deep waters contain mainly , so not much I^- is released to the deep-water column by in situ water column processes, and most organic-iodine gets to the sediments where it is released as I^- (Kennedy and Elderfield, 1987a, Kennedy and Elderfields 1987b; Luther et al., 1995). Kennedy and Elderfield (1987a, 1987b) and Shimmield and Pedersen (1990) report that the molar I/C ratio in planktonic organisms is 10^-4 whereas it is typically >10^-3 in sediments. Decomposition of sedimentary organic-I releases I^- to porewaters and the overlying water column where it can be transported hundreds of kilometers offshore along isopycnal surfaces in OMZs (Farrenkopf and Luther, 2002; Cutter et al., 2018).

4.7 Surface seawater and atmospheric formation of IO₃^-, and iodine speciation in the atmosphere

There is significant literature showing that coastal and oceanic regions are sources of iodine emissions to the atmosphere, and I note some important aspects of this air-sea connection. I^- reacts with O₃ to form IO^-, which at seawater pH forms HOI. Carpenter et al. (2013) showed that this reaction occurs in the first few micrometers below the air-water interface and that HOI is ten-fold greater than I₂ above the sea surface. HOI contributes 75% of the observed iodine oxide aerosol levels over the tropical Atlantic Ocean, and these iodine emissions to the atmosphere have increased 3-fold over the last century due to the increase in anthropogenic O₃ (Carpenter et al., 2021). O₃ reacts stepwise with this gaseous HOI (IO^-) and gaseous to form , which can attach to aerosols.

Formation and release of gaseous I₂ from seawater to air permits photochemical breaking of the I-I bond to form gaseous I atoms, •I, which are reactive radicals. Similarly, release of volatile organic-iodine compounds leads to the homolytic cleavage of the C-I bond to form I•. O₃ reacts readily with I• to form gaseous IO• in the marine boundary layer (Whalley et al., 2010). Further stepwise oxidation of gaseous IO•/HOI leads to . In laboratory experiments using mass spectrometry detection, Teiwes et al. (2019) showed that hydrated iodide, I(H₂O)^-, reacts with gaseous O₃ to form directly without formation of gaseous HOI or IO^-; thus, can form in a two-step reaction sequence in the atmosphere.

Using mass spectrometry to evaluate atmospheric I_xO_y cluster and (nano)particle formation above seabed macroalgae, Sipilä et al. (2016) showed the stepwise formation of HIO₃via HOI and IO•, which leads to (I₂O₅)_x clusters (x=2-5) containing HIO₃ that result in iodine rich aerosol particles. These data on the formation of I₂O₅ aerosols agree with the exothermic ΔH_reaction values of iodine oxide species reacting with O₃ and each other calculated using quantum mechanics (Kaltsoyannis and Plane, 2008). Sipilä et al. (2016) also showed that cluster formation increased as a burst at low tide indicating significant I₂ release from the macroalgae (and subsequent oxidation) as found by Küpper et al. (2008). Hydration of I₂O₅ leads to two . In mass spectrometry laboratory studies, Martín et al. (2022) showed that new iodine containing (nano)particles and also form in the presence of and provide ΔH_reaction data for the gas phase reactions involved. Experiments using the CERN CLOUD (Cosmics Leaving Outdoor Droplets) chamber documented the formation of HIO₃via iodooxy hypoiodite, IOIO, as an intermediate (Finkenzeller et al., 2022) and the fast growth of HIO₃ as (nano)particles (He et al., 2021).

In recent atmospheric campaigns, Koenig et al. (2020) showed that is the main iodine reservoir as it forms on aerosols in the stratosphere with iodine being responsible for 32% of the halogen induced O₃ loss. Cuevas et al. (2022) also showed that iodine can dominate (∼73%) the halogen-mediated lower stratospheric ozone loss during summer and early fall, when the heterogeneous reactivation of inorganic chlorine and bromine reservoirs is reduced.

The information in the preceding paragraphs along with the thermodynamic data from Martín et al. (2022), Figure 2 (the half reaction for O₃ to O₂ and H₂O) and Figure 9 predict that should be the dominant species in the atmosphere. Although reduction of is not predicted in an oxidizing atmosphere, analyses of rainwater (Campos et al., 1996b; Truesdale and Jones, 1996; Baker et al., 2001; Hou et al., 2009), aerosols (Gilfedder et al., 2008; Droste et al., 2021) and snow (Gilfedder et al., 2008) in the marine boundary layer indicate that aqueous iodide and iodate coexist. Hou et al. (2009) reviewed wet iodine speciation data and reported that predominates over I^- from marine sources/air masses whereas I^- predominates from continental air masses.

There are several ways that I^- (or reduced I) can form in rainwater and aerosols. The interconversion between and I^- at the pH of wet deposition also leads to HOI and I₂, which can react with organic material forming C-I bonds that can release I^- (section 4.6). This material has been given the term soluble organically bound iodine and can be larger than the sum of the concentrations of and I^- in aerosols (Gilfedder et al., 2008; Droste et al., 2021). Soluble organically bound iodine can form from release of natural organic iodine from land and sea (a primary source) or from the reaction of natural organic material with HOI or I₂ in the atmosphere (a secondary source). On photolysis of C-I, I• forms and reacts with O₃, and on C-I reaction with nucleophiles, I^- forms. During a study on the formation of cloud condensation nuclei, Huang et al. (2022) also showed that natural gaseous organic material in the marine boundary layer reacts with in aerosols resulting in gaseous I₂, which can be reoxidized to (catalysis) or react to form organic-I compounds. Lastly, Cuevas et al. (2022) reported that photolysis of particles in the stratosphere at a wavelength of about 260 nm can lead to gaseous I• and O₂ during transport from the tropics to the Antarctic region. Thus, there are several pathways for reduction of in the atmosphere.

5 Conclusions

The reduction of to I^- in solution is a facile process by biotic and abiotic reactions. The intermediates and HOI dictate the reactivity sequence via a combination of thermodynamic and kinetic considerations. The to conversion is the least favorable and likely controlling step in this reaction sequence, but there is no need for nitrate reductase for reduction based on numerous studies. The data from this study indicate that once forms there is no thermodynamic barrier to I^- formation. Chemical reduction of all iodine species (not iodide) by sulfide, Fe²⁺ and Mn²⁺ are favorable at seawater and sedimentary pH values, but only sulfide has been studied in the laboratory at oceanic pH values. Dissimilatory reduction during organic matter decomposition seems to be a key process as the couple is more favorable than the couple.

However, the oxidation of I^- back to via³O₂ has a major thermodynamic barrier in solution, and the disproportionation of HOI at seawater pH values is not measurable. Thus, ROS, oxidized Mn and microbes are important for I^- oxidation to due to favorable thermodynamics and kinetics (Table 2). Recent reports of microbial oxidation have not documented the entire six-electron oxidation in a stepwise manner so further work on this topic is necessary. Oxidation of I^- by oxidized Mn is a pH dependent reaction and less likely at seawater pH values but could occur in sedimentary environments. The reactions of O₃ and •OH with iodine species (not ) are thermodynamically favorable over all pH. However, ROS are not normally in significant concentration in seawater to influence formation. Notable exceptions are for (1) sea surface microlayer, which adsorbs atmospheric O₃, and (2) the reaction of Fe²⁺ with O₂ that leads to Fenton chemistry with •OH production. Systems where Fenton chemistry can occur are at/near hydrothermal vents (Shaw et al., 2021), submarine groundwaters (Burns et al., 2010), and sediments or water columns where O₂ and Fe²⁺ concentration profiles overlap including ancient earth (Chan et al., 2016).

I^- is a major sink for O₃ in the sea surface microlayer and the atmosphere. formation in the atmosphere and redeposition to surface seawater may be major iodine processes with the latter being similar to the deposition of trace metals from wet and dry deposition to the surface ocean (e.g., Chance et al., 2015; Meskhidze et al., 2019). Most atmospheric iodine originates from marine sources where I^- oxidation to I₂ and homolytic cleavage of C-I bonds occurs; thus, gaseous iodine emissions from the ocean are reduced. forms from these sources during oxidation by O₃ in the atmosphere. An estimate of atmospheric deposition of to the ocean surface could be made by using the amount of in rainwater and aerosols that would be returned to the ocean surface, but more information on iodine speciation in rainwater and aerosols is needed as global spatial coverage appears limited. Despite major advances in iodine geochemistry over the last two decades, significant research is still needed on the processes that affect I^- oxidation to in the atmosphere, seawater and ocean sediments.

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