of the organisms capable of using a stronger electron acceptor coupled with the oxidation of a limited pool of carbon sources, while organisms using an electron acceptor with a lesser redox potential are outcompeted. The redox zonation is thus a result of microbial ecology, favouring organisms that can gain more energy and accordingly produce more adenosine triphosphate (ATP) per electron transferred. In sediments, the competitive advantage of iron reducers thus also depends on organic carbon supply, reactivity of organic molecules, sediment properties, and sedimentation rate, interacting with the microbial community.
One factor that may modify the sequence of electron acceptors in marine sediments is the poor solubility of iron and manganese oxides, which prevail as minerals under neutral pH. Two different effects play a role for the availability of iron as an electron acceptor for microbial metabolisms:
1 Dissolution may be accelerated by chelating agents in the solution. Chelators are capable of maintaining iron in an oxidized state in solution, and they can be transported to organisms. However, chelation also strongly depends on the reactivity of the mineral.
2 Reduction of the iron may occur while the iron atom is still at its site in the crystal lattice, and it is then released as Fe2+ to the surrounding solution. Reductive dissolution may occur abiotically via sulfide oxidation or biotically using extracellular electron transport systems (see below).
According to Afonso and Stumm (1992), iron reduction by reaction with sulfide essentially depends on the surface area of the mineral. Canfield (1989) showed that reactivity towards sulfide strongly depends on the mineralogy of the iron phase. Poorly ordered iron oxides, but also iron oxides such as hematite and magnetite, are rather rapidly reduced, while some sheet silicates, such as illite or mica, are essentially insoluble over millions of years. This has also been observed in the ratio of pyrite over reactive iron content, showing a steep increase in the uppermost few metres below seafloor but only a minor increase below this depth in sediments of the Peruvian continental shelf (Meister et al. 2019). However, in sediments where sulfide is depleted, reactive iron may survive over a longer period if the sediment becomes rapidly buried into the methanogenic zone, or if microbial activity is generally too low to produce significant amounts of sulfide (suboxic zones). In methanogenic zones, iron (III) can persist over geological timescales (Riedinger et al., 2017) and thereby give rise to cryptic sulfur cycling, if sulfur becomes available (Holmkvist et al. 2011) or it can serve as an electron acceptor driving the anaerobic oxidation of methane (Beal et al. 2009; Wankel et al. 2012). Indeed, methanogenic zones can occur at very shallow depth (ca. 1 m, Thang et al. 2013; or even less: Oni et al. 2015), in which case solid iron phases are likely to bypass the sulfate reduction zone.
Figure 3.1 Concentrations and activities of reduced iron and sulfide in seawater in equilibrium with sedimentary FeS precipitate [log (IAP/Ks) = 0]. The solid line indicates concentrations of total dissolved Fe2+ and ∑H2S. The dashed line indicates the concentrations of free Fe2+ and HS–. The dotted line indicates the activities of free Fe2+ and HS–. Values are corrected for ion activity using the extended Debye–Hückel method and using parameters from the wateq4f.dat database in PhreeqC (Parkhurst and Appelo, 2013). Note that the solubilities of mackinawite (FeS), greigite (Fe3S4), and pyrite (FeS2) are magnitudes lower than bulk sedimentary FeS.
However, even within the zone where free iron is present in solution, it cannot be excluded that reductive dissolution via sulfide is driving iron reduction (Picard et al. 2016). Although the most soluble forms of iron sulfide, such as mackinawite or finely dispersed and poorly crystalline FeS in the sediment, allow for sulfide concentrations only in the nanomolar level in the presence of Fe2+ (Figure 3.1 shows the sulfide concentrations and activities for any given iron concentration /activity at saturation with respect to sedimentary and crystalline FeS), reduced sulfur species may always be available as polysulfide or organosulfur complexes. Thus, a zone with iron in solution, as shown for example in Figure 3.2a, is not necessarily an indication for microbially mediated iron reduction. Hence, in these zones it is difficult to reveal whether or not microbial iron reduction truly takes place. Evidence that could support that iron reduction is unrelated to sulfur cycling is provided by the oxygen isotope value (δ18O) of sulfate. The δ18O in sulfate in seawater (~10‰ relative to standard mean ocean water) is not in equilibrium with seawater, and equilibration occurs upon reoxidation of sulfide, as new S–O bonds are formed (Wortmann et al. 2007). Therefore, δ18O in sulfate is a good indicator for sulfur cycling, or the lack of it, in a zone of iron reduction. Furthermore, iron reduction has been linked to the consumption of several different organic substrates in incubation experiments using sediments from Skagerrak (Baltic Sea; Jensen et al. 2003). A quantitative consumption of organic substrate with iron as electron acceptor has also been shown by Canfield et al. (1993) in the Skagerrak by mass balance calculation.
Figure 3.2 Schematic porewater profiles and dominant microbial groups through the uppermost sediment showing two different types of redox zonation: (a) With low and/or poorly reactive organic matter (TOC) and high reactive ferric iron content, a suboxic zone with free Fe2+ in the porewater is established. Within the suboxic zone, organisms may increasingly use sulfur metabolism as iron becomes more and more limited. (b) If organic matter is more abundant and/or reactive iron content is lower, dissimilatory iron reduction cannot keep up with sulfate reduction and the suboxic zone disappears. The dashed line represents the oxic front.
Although the above‐mentioned studies cannot entirely exclude the contribution of sulfur cycling, Reyes et al. (2016) discussed a hypothesis for how microbial iron reduction is controlled. Iron reduction is limited by the reactivity of the mineral surface, as shown by Afonso and Stumm (1992), and depends on the mineralogy of the iron oxide. Iron reduction has a competitive advantage over sulfate reduction, as it provides more energy (sensu Froelich et al. 1979), but it is limited by the reactivity of the solid phase. Hence, iron reduction would only outcompete sulfate reduction to the level at which the reactivity of the iron mineral can keep up with the rate of sulfate reduction. In surface sediments with abundant reactive organic matter, sulfate reduction would outcompete iron reduction (Figure 3.2b). However, in somewhat less productive areas, or perhaps in zones with more refractory organic matter (e.g. organic matter that has already experienced partial degradation under aerobic conditions, or organic matter derived from land plant material) decomposing more sluggishly, at a level that iron reduction rates can keep up, iron reduction may dominate over sulfate reduction. The ecological role of iron reducers in the marine biome may thus depend on a delicate balance of factors, such as reactivity of iron minerals, reactivity of organic matter, availability of sulfur species and the microbe’s ability to mediate each of the different processes.
3.3. BIOCHEMICAL PATHWAYS OF IRON REDUCTION
3.3.1. The Specificity of Microbial Pathways with Respect to Iron
Coupling of iron reduction to the oxidation of specific organic molecules does not occur spontaneously under Earth surface conditions. Microbes can catalyse
