A review of the mechanisms of mineral-
based metabolism in early Earth analog 
rock-hosted hydrothermal ecosystems
Authors: Maximiliano J. Amenabar and Eric S. Boyd
This is a postprint of an article that originally appeared in World Journal of Microbiology & 
Biotechnology on Faebruary 2019. The final version can be found at DOI: 10.1007/
s11274-019-2604-2.
Amenabar, Maximiliano J., and Eric S. Boyd. "A review of the mechanisms of mineral-based 
metabolism in early Earth analog rock-hosted hydrothermal ecosystems." World Journal of 
Microbiology & Biotechnology 35, no. 29 (February 2019): 10-16. DOI:10.1007/
s11274-019-2604-2.
Made available through Montana State University’s ScholarWorks 
scholarworks.montana.edu 
1 
A review of the mechanisms of mineral-based metabolism in early Earth analog 
rock-hosted hydrothermal ecosystems 
Maximiliano J. Amenabar and Eric S. Boyd 
Department of Microbiology and Immunology, Montana State University, Bozeman, Montana. 
*Author of correspondence:  Eric S. Boyd (eboyd@montana.edu)
Department of Microbiology & Immunology 
Montana State University 
PO Box 173520 
Bozeman, MT 59717 
Phone: (406) 994-7046 
Fax: (406) 994-4926 
Keywords: Elemental sulfur, iron oxide, ferrous iron, sulfide, Archaea, thermophile, hot springs, 
Yellowstone 
2 
ABSTRACT 
Prior to the advent of oxygenic photosynthesis ~ 2.8 Ga, life was dependent on chemical 
energy captured from oxidation-reduction reactions involving minerals or substrates generated 
through interaction of water with minerals. Terrestrial hydrothermal environments host abundant 
and diverse non-photosynthetic communities and a variety of minerals that can sustain microbial 
metabolism. Minerals and substrates generated through interaction of minerals with water are 
differentially distributed in hot spring environments which, in turn, shapes the distribution of 
microbial life and the metabolic processes that support it. Emerging evidence suggests that 
terrestrial hydrothermal environments may have played a role in supporting the metabolism of 
the earliest forms of microbial life. It follows that these environments and their microbial 
inhabitants are increasingly being studied as analogs of early Earth ecosystems. Here we review 
current understanding of the processes that lead to variation in the availability of minerals or 
mineral-sourced substrates in terrestrial hydrothermal environments. In addition, we summarize 
proposed mechanisms of mineral substrate acquisition and metabolism in microbial cells 
inhabiting terrestrial hydrothermal environments, highlighting the importance of the dynamic 
interplay between biotic and abiotic reactions in influencing mineral substrate bioavailability. An 
emphasis is placed on mechanisms involved in the solubilization, acquisition, and metabolism of 
sulfur- and iron-bearing minerals, since these elements were likely integrated into the 
metabolism of the earliest anaerobic cells.  
3 
1. INTRODUCTION
Microorganisms and their activities drive global biogeochemical cycles and influence the 
formation and fate of numerous minerals, both today and throughout the geological past. In many 
cases, the formation and fate of minerals is an indirect consequence of microbial activity. For 
example, the formation of iron monosulfide [mackinawite (FeS)] often involves an abiotic 
reaction between ferrous iron [Fe(II)] and sulfide, either or both of which can be produced by 
ferric iron Fe(III) or sulfate (SO42-) reducing organisms, respectively. Likewise, sulfide, 
produced by microorganisms growing via reduction of oxidized sulfur compounds [e.g., SO42-, 
elemental sulfur (S°)], can drive the subsequent abiotic oxidation of FeS to yield pyrite (FeS2) 
(Drobner et al. 1990; Rickard 1997), the predominant reservoir of iron in the crust (Berner and 
Berner 1996). Similarly, at low temperature and pressures, microbial exopolymers and activities 
have been suggested to be the primary means by which calcite or aragonite (CaCO3) can be 
altered to dolomite (MgCa(CO3)2) (Roberts et al. 2013; Vasconcelos et al. 1995; Zhang et al. 
2012; Zhang et al. 2015). Indeed, it has been estimated that ~2800 of the ~5000 known minerals 
on the planet were either directly or indirectly formed because of microbial activity (Hazen et al. 
2008). 
Of equal importance is the direct use of minerals in oxidation-reduction reactions 
involving minerals which often results in their oxidative or reductive dissolution. For example, 
numerous chemolithotrophic microorganisms can couple the oxidation of hydrogen (H2) to the 
reduction of ferric oxyhydroxides or S° and capture and store energy in the form of ATP (Kletzin 
et al. 2004; Weber et al. 2006). The ferric oxyhydroxides and S° are reductively dissolved to 
4 
Fe(II) and sulfide, respectively, thereby increasing the bioavailability of iron and sulfur for use in 
cellular biosynthesis or as electron donors for taxa in adjacent and more oxidized environments. 
Microbial metabolisms that are dependent on mineral-sourced nutrients are common in 
both marine and terrestrial environments, surface and subsurface locales, and oxic and anoxic 
systems (Gadd 2010). Non-photosynthetic contemporary ecosystems are key analogs for 
understanding the processes that supported life on early Earth prior to the emergence of 
photosynthetic organisms (Amils et al. 2004; Rothschild and Mancinelli 2001). However, 
accessible surface environments where photosynthetic communities are excluded and where 
chemosynthetic communities flourish are rare. These characteristics combine to make the study 
of the use of mineral-sourced energy in contemporary ecosystems that are minimally influenced 
by photosynthesis logistically difficult. Fortunately, photosynthetic life is excluded from 
terrestrial hydrothermal environments with temperatures that exceed 70°C in alkaline systems 
and 54°C in acidic systems (Boyd et al. 2012; Boyd et al. 2010; Brock 1967; Cox et al. 2011; 
Hamilton et al. 2012). Despite the absence of photosynthetic production, abundant and dominant 
filamentous and sessile communities (Fig. 1) develop in close association with minerals in the 
higher temperature transects of these springs, suggesting that minerals or products of mineral-
water interaction support their metabolism. Moreover, despite hot spring waters being exposed to 
the atmosphere, the inverse relationship between the solubility of oxygen and temperature 
(Amend and Shock 2001) allows for anoxic conditions to develop in hot springs and anaerobic 
microorganisms and processes to flourish (Canganella and Wiegel 2014). Thus, these systems 
are often regarded as analogs of anoxic hydrothermal systems on early Earth (Djokic et al. 2017; 
Rothschild and Mancinelli 2001). 
5 
Phylogenetic reconstructions of chemosynthetic microbial inhabitants of high 
temperature hot spring environments reveal that they often form short lineages that branch near 
the base or the root of universal trees, indicating that they are among the earliest, yet least 
evolved of extant microorganisms (Pace 1997; Schwartzman and Lineweaver 2004). Moreover, 
anaerobic chemotrophs are thought to pre-date aerobic chemotrophs based on both physiological 
and paleo-geochemical and geological considerations (Farquhar et al. 2000; Lyons et al. 2014; 
Morris 1975; Thauer et al. 1977). Consistent with these considerations, phylogenetic 
reconstructions of extant organisms indicate that anaerobic chemotrophs typically pre-date 
aerobic chemosynthetic chemotrophs, as has been recently demonstrated for the Archaea 
(Colman et al. 2018). Together, such observations have been interpreted as evidence that 
thermophilic, chemosynthetic organisms and the processes that sustain them may be reminiscent 
of those that supported life on early Earth (Djokic et al. 2017; Rothschild and Mancinelli 2001). 
Moreover, such data have been extended to suggest that high temperature hydrothermal 
environments might have played an important role in the origin of life (Baross and Hoffman 
1985; Corliss et al. 1981; Miller and Bada 1988; Nisbet and Sleep 2001; Shock et al. 1995), a 
hypothesis that dates to as early as 1924 (Harvey 1924). 
High temperature (>70°C) terrestrial hydrothermal environments exhibit colors 
attributable to precipitated minerals that are often generated by the activity of microorganisms 
(Langner et al. 2001; Macur et al. 2013). These include a variety of non-redox active silicate 
minerals including feldspar, tridymite, kaolinite, opal, and christobalite (Colman et al. 2016) as 
well as numerous redox active minerals (Allen and Day 1935). Several of the more abundant 
6 
groups of redox active minerals commonly detected in high temperature terrestrial hydrothermal 
environments include carbonates (aragonite, calcite, siderite), sulfur compounds, (S°, FeS, FeS2), 
iron oxyhydroxides (ferrihydrite, hematite, limonite), amorphous and crystalline arsenic sulfides 
(realgar, orpiment), iron arsenate (scorodite), mercury sulfide (cinnabar), manganese oxides 
(birnessite, hollandite), and sulfate minerals (alunite, jarosite), among others. This review will 
address current understanding of the geobiological controls on the distribution of select minerals 
in terrestrial hydrothermal environments and the mechanisms that allow cells to acquire these 
minerals as substrates to drive their energy metabolisms and to support biosynthesis. We will 
focus on the Yellowstone National Park (YNP) geothermal system given the extensive history of 
coupled geological, geochemical, and microbiological investigations conducted on springs that 
span a wide range of pH (1.5 to 10), temperature (ambient to 93°C), and geochemical and 
mineralogical compositions [e.g., (Allen and Day 1935)]. The emphasis of this review, as it 
relates to mechanisms of mineral substrate acquisition by microbial life, will be placed on sulfur- 
and iron-bearing minerals, since transformation of these elements were likely integrated into the 
metabolism of the earliest anaerobic cells (Philippot et al. 2007; Russell and Martin 2004; 
Wachtershauser 1988; Wachtershauser 1992; Wachtershauser 2006). 
2. Geological settings of terrestrial magmatic hot springs
Terrestrial hydrothermal environments are often localized at plate boundaries including 
those that are divergent (e.g., Iceland) or that are convergent (e.g., Kamchatka, Russia; El Tatio, 
Chile; Taupo, New Zealand). Hydrothermal environments can also be located away from plate 
boundaries in areas that are known as hot spots, which are thought to be generated through either 
7 
crustal extension that can allow mantle melts to approach the surface more easily or through 
mantle plumes (e.g., YNP and Hawaii, United States of America) (Christiansen et al. 2002; 
Smith and Braile 1994). Regardless of the geological setting associated with terrestrial magmatic 
hydrothermal systems, they all share three common requirements for their formation: a source of 
water, a source of heat, and permeable rock strata overlying the heat source (Heasler et al. 2009). 
Moreover, all hydrothermal systems must be continuously recharged with water. 
A schematic of the functioning of continental, volcanic hydrothermal environment is 
illustrated in Fig. 2, using YNP as the model system. Importantly, non-volcanic continental 
hydrothermal systems (e.g., Great Basin, Nevada) are thought to function similarly, with the 
major differences being the heat source and input of magmatic volatiles. Not included in the 
schematic illustrated in Fig. 2 is the hot spot magma plume, estimated to be ~ 40 km beneath the 
surface of YNP (Huang et al. 2015). This magma plume is transferring heat (and volatiles) to a 
basalt magma partial melt (~46,000 km3 in volume) at ~ 45-20 km depth which is transferring 
heat to an upper rhyolitic magma partial melt (~10,000 km3 in volume). The upper chamber is 
depicted in Fig. 2. 
Based on current understanding, meteoric water is thought to infiltrate the geothermal 
system in a recharge area, which in the case of YNP is largely concentrated just outside of the 
North and Eastern Park boundary (Kharaka et al. 2002; Rye and Truesdell 2007). Water 
percolates through fractures and/or pores in rocks and into the subsurface to form a single deep 
“parent” water body. The temperature of this parent water body in YNP has been estimated to be 
up to 360°C with a chloride (Cl-) concentration of 310 mg L-1 (McKenzie and Truesdell 1977) 
8 
and a SO42- concentration of 73 to 98 mg L-1 (Nordstrom et al. 2009). The parent water body is 
infused with oxidized volcanic gases, including sulfur dioxide (SO2) and hydrochloric acid (HCl) 
which, upon contact with water at high temperature, are thought to contribute the SO42- estimated
above via SO2 disproportionation (Reaction 1) and the Cl- estimated above via HCl ionization,
respectively (McKenzie and Truesdell 1977; Truesdell and Fournier 1976; White et al. 1971).
4SO2 + 4H2O → 3H2SO4 + H2S    Reaction 1 
Heat-induced pressure or density differences allow these fluids to ascend to the surface along a 
fluid pathway that is controlled by fractures and faults. As such, the majority of hot spring 
activity in YNP is concentrated along the edge of the main ring fracture associated with the 0.6 
mya caldera forming eruption and at the margins of the two resurgent domes, the Mallard Lake 
and the Sour Creek.  
3. Geological controls on hot spring water geochemical composition
As hydrothermal fluids ascend to near surface local aquifers that feed hot springs, 
geysers, and fumaroles, they can be altered by processes such as water-rock interaction, boiling, 
and mixing with near-surface groundwater (Lowenstern et al. 2012; Nordstrom et al. 2009). A 
primary control on the composition of hot spring waters is the process of decompressional 
boiling which can separate fluids into a liquid and a vapor phase in what is oftentimes referred to 
as phase separation (Fournier 1989; Nordstrom et al. 2009; Truesdell and Fournier 1976; White 
et al. 1971). Most evidence from YNP hot springs suggests that phase separation is a multiple-
stage process (incomplete separation of liquid and vapor) that produces results intermediate 
between single-stage and continuous phase separation (Truesdell and Fournier 1976). Volatiles 
9 
in hydrothermal fluids are likely to be partitioned to the vapor phase during phase separation 
(Fournier 1989; Nordstrom et al. 2009). These volatiles include H2, helium (He), methane (CH4) 
and hydrogen sulfide (H2S), among other gases (Lowenstern et al. 2015). As described in more 
detail below, the behavior and fate of H2S during phase separation exerts significant influence on 
the geochemical composition of hot spring waters (Lindsay et al. 2018; Nordstrom et al. 2009), 
and of particular relevance to this review, the availability of minerals to support microbial 
metabolism. It follows that an important consideration of the hydrothermal fluids undergoing 
phase separation is the pH and temperature of parent fluids, given that the pKa of H2S/HS- is 6.5 
at 100°C (Amend and Shock 2001). Thus, depending on the pH and temperature of the parent 
hydrothermal fluids in question, a fraction of the total H2S/HS- may partition during initial 
boiling events, potentially leaving substantial HS- in solution that is then available for processing 
during subsequent boiling episodes (multi-stage phase separation model). Subsequent phase 
separation events would be expected to progressively deplete the H2S/HS- remaining in the liquid 
phase. Acidification of fluids as they ascend to the surface (described below) or as they mix with 
acidified meteoric fluids in the form of perched aquifers (Nordstrom et al. 2009), which 
themselves could be acidic in particular in areas around the Norris Geyser Basin (White et al. 
1988), may affect the speciation of H2S/HS-. This in turn may alter the amount of H2S that is 
volatized in subsequent phase separation events, potentially leading to its enrichment and thus 
acidification of waters in certain geographic locations or springs. It is generally thought that the 
waters that boil in deeper portions of the YNP system are near neutral (Fournier 1989; White et 
al. 1971). 
10 
Following phase separation, the vapor is free to migrate toward the surface where it can 
condense and interact with oxygen (O2)-rich meteoric fluids. The near surface oxidation of 
H2S/HS- with O2 can generate acidity in the form of sulfuric acid (Nordstrom et al. 2009), a 
process that is thought to be mediated by microorganisms (Colman et al. 2018). Thus, springs 
that are sourced with waters that are influenced by vapor phase input tend to be acidic and 
enriched in SO42- (Fournier 1989; Nordstrom et al. 2005; Xu et al. 1998). Such systems are
common in the north to north central (Norris Geyser Basin, Central Plateau) and eastern (Hot
Spring Basin, Mud Volcano, Crater Hills) edges of the caldera (Lowenstern et al. 2015). In
contrast, systems receiving substantial liquid phase input tend to be alkaline and rich in Cl-.
These types of systems tend to be concentrated in the Lower, Midway, and Upper Geyser Basin
on the western side of the caldera (Lowenstern et al. 2015). Thus, phase separation of
hydrothermal fluids results in the concentration of spring types geographically but that fit within
a bimodal distribution of hot spring pH types, with the two spring types centered on sulfuric acid
(vapor phase influenced) and bicarbonate (liquid phase influenced) buffered springs (Amenabar
et al. 2015). Variable mixing and dilution of spring waters can further affect their composition
and differences in the underlying bedrock and the interaction of these rocks with hydrothermal
fluids can further alter spring chemistry (Fournier 1989).
4. Geological controls on the distribution of redox active minerals in hot springs
The processes described above, in particular phase separation, are also likely to influence 
the availability of minerals in hot springs, including those that play a prominent role in 
supporting the microbial metabolism of thermophiles such as S° and iron (hydr)oxides. The 
11 
deposition of S° in hot springs most likely begins with oxidation of H2S, which is thought to be 
primarily sourced to the parent hydrothermal aquifer by SO2 disproportionation (Reaction 1). For 
reasons discussed above, H2S can be enriched in vapor phase influenced hot spring waters 
(Nordstrom et al. 2009). The kinetics of abiotic oxidation of H2S with O2 are several orders of 
magnitude slower than biotic oxidation (Chen and Morris 1972; Luther et al. 2011; Zhang and 
Millero 1993), specifically in solutions with pH <6.0 (Chen and Morris 1972) where the 
predominant, uncharged H2S species is less reactive (Millero 1986). However, these experiments 
have only been performed at low temperature (<45°C) and the role of abiotic reactions in the 
oxidation of H2S at higher temperatures is not known, especially at lower pH. The Arrhenius law 
would predict an increase in the rate of abiotic H2S oxidation with increasing temperature, to the 
extent that O2 in solution does not become limiting given its decreasing solubility at elevated 
temperature (Shock et al. 2010). Trace metals (e.g., vanadate, Fe(III), cupric copper) have been 
shown to enhance rates of abiotic H2S oxidation (Steudel 1996), and thus must also be 
considered when assessing the relative contribution of biotic and abiotic processes in H2S 
oxidation.  
Regardless of whether oxidation is biotically or abiotically catalyzed, the process 
typically proceeds through a series of intermediate sulfur species that includes thiosulfate (S2O32-
), in particular when O2 is limiting (Chen and Morris 1972; Xu et al. 1998). Sulfide oxidation has 
also been suggested to directly result in formation of S° (Fuseler and Cypionka 1995). At 
circumneutral pH, S2O32- is thermodynamically stable and can serve as an oxidant or reductant
for chemical reactions. However, at acidic pH (pH <4.0), S2O32- is unstable (Xu and Schoonen
1995) and disproportionates to form solid-phase S° and sulfite (SO32-), the latter of which is also
12 
unstable and hydrolyzes to form SO42- and protons (Nordstrom et al. 2004). Irrespective of the 
mechanism by which it is formed, the S° formed from H2S oxidation accumulates in many acid 
(< 4.0) geothermal springs due to its slow reactivity with water below 100°C (Fournier 1989; 
Kamyshny 2009) and thus represents a key electron donor or acceptor capable of supporting 
microbial metabolism (Boyd et al. 2009). 
Protons generated from H2S/S° oxidation in vapor-phase-influenced environments can 
also contribute to the release of metals from altered host rock (Ball et al. 2010), including Fe(II) 
that can then be oxidized leading to formation of soluble Fe(III) ions and Fe(III) containing solid 
phases. The iron (hydr)oxide mats that form at Chocolate Pots (pH 5.8 to 6.0 at source), a vapor 
influenced mildly acidic spring just south of Norris Geyser Basin (Allen and Day 1935), is 
perhaps the most well studied iron (hydr)oxide mat in a hot spring in YNP if not the world. This 
spring has been studied extensively from both the mineralogical [e.g., (Wade et al. 1999; Wu et 
al. 2013)] and the biological perspectives [e.g., (Fortney et al. 2016; Fortney et al. 2018a; 
Fortney et al. 2018b; Pierson and Parenteau 2000; Trouwborst et al. 2007)]. Current 
understanding suggests that Fe(II) in the source vent waters of Chocolate Pots is rapidly oxidized 
by O2 that is in supersaturating concentrations due to the activity of oxygenic Cyanobacteria. The 
precipitates at Chocolate Pots comprise a variety of iron (hydr)oxides including goethite (α-
FeOOH) and hematite (α-Fe2O3), and the carbonate mineral siderite (FeCO3). The iron 
(hydr)oxides can serve as electron acceptors in microbial metabolism whereas siderite could 
serve as an electron donor or potentially even an inorganic carbon source. However, to the 
authors knowledge, siderite has not been extensively examined as an electron donor or inorganic 
carbon source in this or other YNP springs. 
13 
In more acidic hot springs (pH < 4.0) such as “Dragon Spring” in the Norris Geyser 
Basin (Langner et al. 2001), abiotic oxidation of Fe(II) with O2 is kinetically inhibited (Edwards 
et al. 2000; Edwards et al. 1999; Johnson and Hallberg 2005). In these systems, Fe(II) oxidation 
is catalyzed by the activity of thermoacidophilic Archaea and Bacteria (Kozubal et al. 2008; 
Kozubal et al. 2012), which form iron oxide containing mineralized microbial mats (Fig. 1). In 
addition to microbial cells, the mats can be comprised of amorphous iron oxyhydroxides, 
ferrihydrite, goethite, and hematite (Kozubal et al. 2012; Langner et al. 2001). Importantly in 
acidic springs, such as “Dragon Spring”, Fe oxide mats tend to not form until H2S is oxidized 
with concomitant S° deposition or volatilized from the system, likely due to the more favorable 
energetics associated with microbial oxidation of H2S or S° with O2 when compared to Fe(II) 
with O2 (Amend and Shock 2001). These Fe(III) oxides or ions can serve as a source of oxidant 
for Fe(III) reducing taxa.  
While Fe(II) is more effectively leached from bedrock by acidic spring waters than 
circumneutral to alkaline spring waters and is typically in higher concentrations in the former 
spring type (Ball et al. 2010), Fe has been detected in springs of the latter type. This includes 
aqueous Fe(III) in the source waters of Perpetual Spouter (pH 7.1, 84°C) (Ball et al. 2010) and in 
the solid phase precipitates in the source of the spring (Inskeep et al. 2005). In this spring, 
concentrations of Fe(III) were higher than what is expected based on the solubility of a variety of 
iron oxides under these conditions, indicating that the Fe(III) may be complexed by organic 
ligands thereby increasing its solubility (Kraemer 2004). Moreover, given that Fe(II) reacts 
quickly with O2 in solutions with circumneutral to alkaline pH and that are high temperature 
14 
(Morgan and Lahav 2007), it is likely that oxidation and formation of Fe(III) oxides in such 
systems is abiotic. Regardless, these Fe(III) oxides can serve as electron acceptors to support 
microbial metabolism. 
5. Metabolism of S° in hydrothermal environments
5.1 Geomicrobiology of S° 
S°, formed from the incomplete oxidation of H2S, is a common redox active mineral in 
hot springs (Amenabar and Boyd 2018; Boyd and Druschel 2013; Boyd et al. 2007; Colman et 
al. 2016). Precipitation of S° often leads to an S° depositional zone of characteristic yellow color 
(Amenabar and Boyd 2018; Langner et al. 2001) (Fig. 1). S° is generally more abundant in 
acidic hot springs (Boyd et al. 2007; Inskeep et al. 2013), since these springs are influenced to a 
greater degree by vapor phase input that is enriched in H2S relative to liquid phase input 
(Nordstrom et al. 2009) and the pH and Eh of these springs is within the stability field of S° 
(Rickard and Luther 2007). Indeed, a taxonomically diverse array of thermoacidophiles have 
been shown to use S° in their energy metabolism, either as an electron donor (sulfur oxidation) 
(Amenabar et al. 2017; Friedrich et al. 2001; Huber and Prangishvili 2006) or as an electron 
acceptor (sulfur reduction) (Amenabar et al. 2017; Boyd and Druschel 2013; Boyd et al. 2007; 
Boyd et al. 2009; Huber and Prangishvili 2006). In many cases, organisms that integrate S° into 
their energy metabolism have been shown to be numerically dominant in communities inhabiting 
acidic hot springs communities (Boyd and Druschel 2013; Boyd et al. 2007; Boyd et al. 2009; 
Colman et al. 2016; Colman et al. 2018; Friedrich et al. 2001; Giaveno et al. 2013; Muyzer and 
Stams 2008; Power et al. 2018; Segerer et al. 1986; Urschel et al. 2016; Yoshida et al. 2006). 
15 
Thermoacidophiles within the archaeal order Sulfolobales, which are physiologically versatile 
and grow as facultative aerobes with the ability to oxidize or reduce S° (Amenabar et al. 2018; 
Giaveno et al. 2013; Huber and Stetter 2015), are commonly detected in acidic thermal 
environments. In anoxic acidic thermal springs, organisms within the archaeal order 
Desulfurococcales, which typically use S° as an oxidant, are commonly detected.  
Despite the lower occurrence of S° in neutral/alkaline environments, Archaea and 
Bacteria have been suggested to respire S° under these conditions using either CO2 
(chemolithoautotrophs) or organic substrates (chemoorganotrophs) as carbon source and electron 
donor (Amenabar et al. 2015; Canfield and Raiswell 1999; Ghosh and Dam 2009). However, 
when compared to acidophiles, far less is known about the mechanisms of S° reduction in 
organisms inhabiting circumneutral to alkaline hot springs (Amenabar et al. 2015; Eder and 
Huber 2002; Ghosh and Dam 2009). As described in more detail in section 5.2 below, it is 
possible, if not likely, that organisms in these environments are utilizing a more soluble source of 
S° such as polysulfide (Sx2- where X is the number of sulfur atoms). Sx2- are formed by reaction 
of HS- with S° (Steudel and Eckert 2003) and are stable under sulfidic, circumneutral (pH > 6.0) 
conditions (Schauder and Muller 1993).  
Reduction of S° (or Sx2-) has been suggested to be an ancient process (Canfield and 
Raiswell 1999; Stetter et al. 1990; Stetter and Gaag 1983) and it is worth exploring the basis of 
this argument for the purposes of this review. Fractionation of sulfur isotopes (34S/32S) during 
microbial S° (or Sx2-) reduction to H2S/HS- is minimal or close to zero (Gest and Hayes 1984; 
Kaplan and Rittenberg 1962; Surkov et al. 2012). This makes it difficult to use sulfur isotopes in 
16 
minerals preserved in the rock record to date when this metabolic process may have emerged. 
Yet, isotopic evidence from marine chert-barite deposits from the Dresser Formation, Australia 
suggest that S° was at least available as a substrate at 3.5 Ga, with evidence indicating that it was 
disproportionated not necessarily reduced (Philippot et al. 2007). Importantly, the barite veins 
that were analyzed in the aforementioned study have been suggested to represent conduits for 
hydrothermal fluid circulation and may have been deposited in a volcanic caldera like 
environment (Van Kranendonk 2006). These observations, combined with the widespread 
prevalence of S° reducing thermophilic organisms in both the bacterial and archaeal domain 
(Amenabar et al. 2015; Canfield and Raiswell 1999; Ghosh and Dam 2009), including deeply 
rooted methanogens (Stetter and Gaag 1983), suggests that energy conservation via S° reduction 
is likely to be a primitive process. Moreover, the prevalence of hyperthermophiles that branch 
near the root of the universal tree of life, and that support their energy metabolism through S° 
reduction, may suggest that this process emerged in a hyperthermophile inhabiting a 
hydrothermal environment early in Earth history (Schauder and Kröger 1993; Stetter et al. 1990). 
S° can also serve as a reductant in the energy metabolism of a variety of microorganisms, 
including those that inhabit hot springs (Amenabar et al. 2017). While the majority of work on 
S° oxidation has focused on aerobic organisms (Ghosh and Dam 2009; Kletzin et al. 2004), S° 
oxidation has also been demonstrated in anaerobes, in particular those that can reduce soluble or 
insoluble Fe(III). Among characterized organisms from high temperature habitats, the coupling 
of S° oxidation with Fe(III) reduction has thus far only been demonstrated in members of the 
archaeal genus Acidianus (Amenabar et al. 2017; Giaveno et al. 2013). Like the biological 
reduction of S° to H2S/HS-, the biological oxidation of S° to SO42- imparts a minimal
17 
fractionation on sulfur isotopes (Zerkle et al. 2009), again, making it difficult to use sulfur 
isotopes preserved in minerals in the rock record to gauge when this process might have 
emerged. However, evidence for S° in environmental systems as early as 3.5 Ga (e.g., (Philippot 
et al. 2007; Van Kranendonk 2006)) and Fe(III) in ancient banded iron formations, including 
those as old as 3.8 Ga in the Isua of West Greenland (e.g., as reviewed in (Gole and Klein 
1981)), indicates that the substrates to support anaerobic S° oxidation coupled with Fe(III) 
reduction were available at these times, albeit not necessarily in the same location. Moreover, the 
prevalence of archaeal and bacterial organisms that support their energy metabolism through the 
reduction of Fe(III), including those that grow at extremely high temperature (Kashefi and 
Lovley 2003), has been used to suggest that Fe(III) reduction is among the most ancient of 
metabolic processes, perhaps predating the process of S° reduction (Vargas et al. 1998). This 
conclusion is partially supported by isotopic data from banded iron and other iron-containing 
formations that together indicate that the process of Fe(III) reduction has an early origin, with 
current data suggesting this process to be at least as old as 3.1 Ga (as reviewed in (Johnson et al. 
2008)). Importantly, it is not yet clear from these observations what the electron donor 
supporting Fe(III) reduction would have been during this time. 
S° can also simultaneously serve as an electron donor and acceptor in a process called 
sulfur disproportionation (Muyzer and Stams 2008). This process is thought to be ancient, with 
34S/32S isotopic fractionation data indicative of this process in hydrothermal barite deposits dated 
to 3.5 Ga (Philippot et al. 2007). Despite the prevalence of S° in hydrothermal environments, 
both today and in Earth’s geologic past, only two thermophilic organisms with the ability to 
disproportionate S° have been described (Amenabar and Boyd 2018; Slobodkin et al. 2012). The 
18 
lack of cultured representatives capable of S° disproportionation from hydrothermal 
environments may be due to difficulties in obtaining cultures of these organisms due to 
thermodynamic constraints (Amenabar and Boyd 2018) imposed on this type of metabolism by 
high levels of H2S/HS- or SO42- that commonly associate with acidic thermal environments 
(Amenabar et al. 2015). Nonetheless, a recently described archaeal strain from a S° rich acidic 
hydrothermal environment in YNP was shown to oxidize, reduce, and disproportionate S° 
(Amenabar and Boyd 2018), pointing to the key role of S° in sustaining the energy metabolism 
of thermoacidophiles in these environments. Importantly, microorganisms with the ability to 
incorporate S° into their energy metabolism are common outside of hydrothermal environments 
as well. As will be described in the following section, the mechanisms that allow for the use of 
S°, whether it be oxidatively, reductively, or disproportionatively, are unlikely to be dependent 
on temperature. Thus, discussion will focus on those organisms that are the target of this review, 
thermophiles.  
5.2 Mechanism of S reduction, oxidation, and disproportionation in thermophiles 
S° can be effectively described as insoluble in water (478 nM at 80°C) (Kamyshny 2009). 
Thus, in order to use this mineral for use as a substrate in energy metabolism, microorganisms 
have to access it by either direct contact or by using a more “hydrophilic” or “colloidal” form 
(Schauder and Kroger 1993). It has been shown that aqueous solutions containing H2S/HS- can 
dissolve S° through the formation of polysulfides (Sx2−) (Giggenbach 1972; Kleinjan et al. 2005), 
with the rate of Sx2− formation dependent on the pH of the solution; pH influences the stability of 
19 
Sx2−, as described below. This is an important consideration since Sx2− can act as an auto-catalyst 
accelerating the dissolution of S° (Giggenbach 1972; Kleinjan et al. 2005). 
Sx2− ions are unbranched chains of sulfur atoms, with the number and structure of those 
linear sulfur chains being dependent on the pH of the solution (Kamyshny et al. 2007; Steudel 
and Eckert 2003). At neutral to intermediate pH conditions, S62-, S52-, and S42- ions dominate the 
solution. However, under acidic conditions, polysulfide ions are not stable and rapidly 
disproportionate (Boyd and Druschel 2013; Giggenbach 1972; Kamyshny et al. 2003; Kleinjan et 
al. 2005). The concentration of Sx2- in aqueous solutions is dependent on pH, the concentration of
sulfide, and the availability of S°. In aqueous solutions with pH > 6.0, measurements of Sx2− 
formed from S° in the presence of abundant sulfide indicates that Sx2− is present at sufficient 
concentration to support sulfur reduction at the growth conditions of sulfur reducing Bacteria and 
Archaea (Schauder and Kröger 1993; Schauder and Muller 1993). Growth assays indicate that 
Sx2− supports the growth of sulfur reducing neutrophilic microorganisms such as the bacterium 
Wolinella succinogenes (Klimmek et al. 1991) and the archaeon Pyrococcus furiosus 
(Blumentals et al. 1990). These observations, in addition to the mechanisms of Sx2− formation 
and Sx2− stability outlined above, suggest that S° reducing microorganisms inhabiting 
circumneutral pH environments are likely using Sx2− as terminal electron acceptor (TEA) (Fig. 
3A). Indeed, a recent study identified a membrane bound oxidoreductase, termed MBX (Schut et 
al. 2007; Wu et al. 2018), that is involved in reduction of sulfane bonds during Sx2− reduction in 
the hyperthermophile P. furiousus. Thus, Sx2− is likely the soluble intermediate for S° reduction 
for microorganisms growing at pH >6.0 (Findlay 2016; Schauder and Muller 1993).  
20 
Intriguingly, like S° reducers inhabiting circumneutral pH environments, S°-reducing 
thermoacidophiles also appear to be reducing a soluble source of S°. This was originally reported 
for the S°-reducing thermoacidophilic crenarchaeote Acidilobus sulfurireducens 
(Desulfurococcales), which was shown to not associate with bulk S° during S° dependent growth 
(Boyd et al. 2007). However, the amount of S° dissolved as Sx2− decreases dramatically with 
decreasing pH, due to the instability of Sx2− in acidic solutions (Schauder and Muller 1993). In 
aqueous solutions with pH < 6.0, Sx2− (formed via reactions between sulfide (H2S/HS-) and S°) 
disproportionate to yield H2S and soluble nanoparticulate S° rings (S8) according reaction 2: 
2 S52- + 4 H+ → α-S8° + 2 H2S Reaction 2 
The decreased stability of Sx2− in acidic solutions relative to alkaline solutions suggests 
that microorganisms inhabiting acidic environments are unlikely to be using Sx2− as TEA but 
rather are likely using nanocrystalline S° formed through Sx2− disproportionation (reaction 2) 
(Boyd and Druschel 2013). This was demonstrated by sequestering bulk S° powder inside 
dialysis membranes of different pore sizes to limit the access of cells to the surface of S° during 
growth. Under this growth condition, the thermoacidophile A. sulfurireducens did not require 
direct access to S° to use it as a TEA (Boyd and Druschel 2013), indicating that these cells were 
using a more soluble form of S° (nanoparticulate S°). Moreover, decreased S° reduction activity 
and cell production were observed when S° was sequestered in dialysis membranes with 
increasingly small pore sizes, an observation that was attributed to the nanoparticulate S° 
existing in a range of particle sizes (Boyd and Druschel 2013). Nanoparticulate S° rapidly 
aggregates/coarsens into larger particles due to hydrophilic interactions (Boyd and Druschel 
2013; Garcia and Druschel 2014) and thus was suggested to be the soluble form of S° that served 
as electron acceptor. 
21 
Founded on an extensive literature database describing the inorganic chemistry of Sx2- 
(Giggenbach 1972; Kamyshny 2009; Kamyshny et al. 2003; Schauder and Muller 1993; Steudel 
1996; Steudel and Eckert 2003), a model involving a series of biological/abiological feedbacks 
was put forth to describe the formation and fate of nanoparticulate S° in cultures of A. 
sulfurireducens (Boyd and Druschel 2013). Here, H2S/HS- resulting from the biological 
reduction of S°, generates soluble linear chains of Sx2- through abiotic nucleophilic attack on
bulk (stacked) S° rings. The thermodynamic instability of Sx2- in the pH 3.0 cultivation medium
used to cultivate A. sulfurireducens leads to its rapid, abiotic disproportionation yielding soluble
S8° molecular rings that rapidly aggregate or coarsen, reaching average particle diameters of 400
nm within several minutes of their disproportionation (Boyd and Druschel 2013; Garcia and
Druschel 2014). These reactions were confirmed to take place in the absence of cells. Thus, the
abiotic rates of Sx2- production and disproportionation, the former of which is dependent on the
rate of H2S produced by biological activity, and the kinetics of coarsening are likely to generate a
distribution of soluble S° particle sizes during the incubation period.
Consistent with this model, similar growth experiments with the thermoacidophile 
Acidianus strain DS80 indicated that the electron acceptor (i.e., nanoparticulate S°) that 
supported growth also exhibited an apparent size dependence (Amenabar and Boyd 2018). These 
results suggest that a similar mechanism involving biological and abiological feedbacks (Fig. 
3A) is likely supporting S° dependent growth in strain DS80 (Sulfolobales) and A. 
sulfurireducens (Desulfurococcales) (Boyd et al. 2007) and may extend to other 
thermoacidophiles such as the thermoacidophilic bacterium Thermovibrio ammonificans (Jelen 
22 
et al. 2018). Importantly, application of environmental voltametric techniques by Lorenson and 
colleagues (Lorenson 2006) and chromatographic methods by Kamyshny and colleagues 
(Kamyshny et al. 2014) reveal the presence of soluble S° colloids or nanoparticles in a variety of 
sulfur rich, hydrothermal environments in YNP including Evening Primrose (pH 6.5) and Cinder 
Pool (pH 4.0 – 5.0). These environments also feature elevated abundances of organisms affiliated 
with Acidilobus and Acidianus (Colman et al. 2016; Urschel et al. 2016; Urschel et al. 2015), 
suggesting that the process that allows for solubilization and use of S° in pure cultures may also 
be operational in sulfur rich hot springs inhabited by these strains.  
Based on the intermediates involved in the proposed mechanism of S° reduction (Fig. 
3A) and their thermodynamic stabilities, the requirement for direct access to S° is going to differ 
depending on whether cells are using the mineral as an oxidant, a reductant, or both as a 
reductant and oxidant (disproportionation). This was demonstrated using Acidianus strain DS80 
as a model (Amenabar and Boyd 2018). Like Acidianus cells grown via S° reduction, those 
grown via S° disproportionation did not require direct access to the mineral for it to be used in its 
energy metabolism. This suggests that the H2S formed through disproportionation of S° 
functioned to solubilize S° through a mechanism like what is described above for S° reducing 
cells. However, unlike DS80 cells reducing or disproportionating S°, cells that were grown via 
S° oxidation required direct access to the mineral. The requirement for direct contact to S° for 
use as an electron donor is consistent with microscopic observations of cultures of Acidianus 
spp. (Amenabar, unpublished data) and other S° oxidizing crenarchaeotes such as Sulfolobus 
spp. that were shown to be attached to S° crystals in both lab cultures (both genera) and in 
samples collected from various hot spring environments (Sulfolobus spp. only) (Brock et al. 
23 
1972). Further evidence that direct access to bulk S° is required for S° to serve as electron donor 
comes from a study of Acidianus strain DS80, whereby growth was not observed when access to 
bulk S° was restricted by placing it inside of dialysis membranes (Amenabar and Boyd 2018). 
This observation further substantiates the proposed mechanism that allows cells access to S° 
under reducing conditions since the H2S/HS- (the product of S° reduction and disproportionation) 
needed to initiate the series of abiotic reactions that solubilize S° (Boyd and Druschel 2013) was 
incapable of being produced under oxic growth conditions (Figs. 3A, B). 
S° oxidation in thermoacidophiles involves the cytoplasmic sulfur oxygenase reductase 
(SOR) complex which disproportionates S° to yield SO32- and H2S. The SO32- produced can then 
be assimilated via APS reductase or further oxidized to SO42- via a membrane bound
sulfide:acceptor oxidoreductase. Through a series of non-enzymatic steps involving S° as an
intermediate, H2S is thought to be converted to S2O32- and oxidized via a thiosulfate:quinone
oxidoreductase (Kletzin et al. 2004). SOR is common among thermoacidophiles within the
Sulfolobales genera Acidianus and Sulfolobus. Little is known of putative mechanisms involved
in the oxidation of S° in organisms inhabiting neutral to alkaline environments (Ghosh and Dam
2009; Sorokin et al. 2001). This is likely due to the predominant form of S° being Sx2- under
these conditions since the intermediate that is largely responsible for S° precipitation during
H2S/HS- oxidation, S2O32-, is stable at pH > 4.0 (Xu et al. 1998).
Two primary mechanisms of S° reduction that involve membrane-bound enzyme 
complexes have been described to date. In chemolithotrophic S° reducing organisms that use H2 
as electron donor (e.g, Thermoproteus tenax, A. ambivalens or Acidianus strain DS80), a 
24 
respiratory system involving a membrane-bound sulfur reductase complex (SRE) and a 
membrane-bound [NiFe]-hydrogenase, similar to those found in the mesophilic S° reducing 
bacterium W. succinogenes (Amenabar et al. 2018; Hedderich et al. 1998; Laska et al. 2003), is 
involved. Electron transfer between the membrane bound [NiFe]-hydrogenase involved in 
oxidizing H2 and SRE involved in S° reduction is thought to involve the hydride carrier 
sulfolobusquinone. Specificity of the hydrogenase or SRE for sulfolobusquinone has been 
suggested to constrain these organisms to obligately reduce S° with only H2 as electron donor 
(Amenabar et al., 2018). A second cytoplasmic SRE (SRE2) complex has been identified in S° 
reducing organisms that use organic substrates as electron donor [(Amenabar et al. 2018), and 
references therein]. The cytosolic location of SRE2 is thought to allow for reduced electron 
carriers from organic carbon oxidation via the Kreb’s cycle (e.g., NADPH) to reduce S°.  
A second prevalent mechanism of S° (SX2-) reduction has been described in heterotrophic 
hyperthermophilic Archaea, primarily within the Thermococcales (Euryarchaeota) order. The 
proteins comprising this multisubunit oxidoreductase complex, termed MBX, resemble 
components of energy converting [NiFe]-hydrogenases and NADH dehydrogenase (Complex I) 
(Schut et al. 2013). This has led to the suggestion that MBX represents an evolutionary 
“intermediate” between early evolving H2 based respiratory metabolisms and those using 
oxidants with much higher potential (Boyd et al. 2014; Schut et al. 2016). Neutrophilic 
hyperthermophiles, such as Pyrococcus and Thermococcus, use MBX to reduce the sulfane 
bonds in SX2- (Wu et al. 2018). Homologs of MBX have also been identified in a variety of
crenarchaeotes that commonly inhabit hot spring environments, including a variety of
neutrophilic members of the Desulfurococcales (Schut et al. 2013). The apparent lack of MBX
25 
homologs among acidophiles and the detection of MBX homologs among neutrophiles is 
consistent with the prevalence and thermodynamic stability of SX2- at these pH values. This
suggests that the stabilities of sulfur compounds may have played a role in the emergence and
evolution of enzymatic mechanisms to utilize these substrates in energy metabolism.
6. Metabolism of Fe in hydrothermal environments
6.1 Geomicrobiology of Fe 
In addition to S°, common minerals capable of supporting microbial metabolism in hot 
springs include a variety of iron (hydr)oxides (Fig.1) (Fortney et al. 2016; Inskeep and 
McDermott 2005; Kashefi et al. 2008b; Langner et al. 2001; Lovley et al. 2004; Slobodkin 2005; 
Slobodkina et al. 2012), which result from the oxidation of Fe(II) with oxygen (Inskeep et al. 
2004; Kozubal et al. 2008). Dissimilatory iron reduction (DIR) plays a key role on the 
biogeochemistry of iron in both moderate and higher temperature environments (Lovley et al. 
2004) and has been suggested to be one of the earliest forms of microbial respiration (Vargas et 
al. 1998). Microorganisms with the capacity to conserve energy from DIR are phylogenetically 
dispersed throughout the bacterial and archaeal domains and this metabolism has been suggested 
to be a highly conserved trait among hyperthermophiles (Lovley 2004; Lovley et al. 2004). See 
text in section 5.1 for additional discussion on the primitive nature of DIR and the organisms that 
catalyze this process. 
Iron availability is low in most oxic and circumneutral environments. This is not 
necessarily due to low total iron content but rather is due to the low solubility and the slow 
26 
dissolution kinetics of iron bearing minerals (Kraemer 2004). Therefore, in order to use iron 
(hydr)oxides in DIR, microorganisms had to evolve mechanisms to transfer metabolic electrons 
to insoluble solid phases. Perhaps the most well studied bacterial models for understanding the 
mechanisms of iron mineral reduction/oxidation in microbial metabolism under mesophilic 
conditions are Geobacter spp. and Shewenalla spp. (Fredrickson et al. 2008; Lovley et al. 2004; 
Nealson and Scott 2006; Shi et al. 2007; Shi et al. 2009). However, thermophiles with the ability 
to reduce iron bearing minerals have also been described (Amenabar and Boyd 2018; Kashefi 
and Lovley 2000; Kashefi et al. 2008b; Slobodkin et al. 1997; Slobodkin et al. 1999), yet the 
mechanisms involved in the respiration of these solid minerals are not well understood. Most of 
our knowledge on iron (hydr)oxide respiration in hyperthermophiles comes from studies of the 
neutrophilic genus Pyrobaculum, members of which are metabolically flexible and are widely 
distributed in circumneutral to alkaline terrestrial hydrothermal environments [(Feinberg and 
Holden 2006), and references therein].  
Under high temperature and acidic conditions, our understanding of iron (hydr)oxide 
reduction is less complete, likely because thermoacidophilic iron reducers have only just been 
reported (Amenabar et al. 2017; Giaveno et al. 2013; Kozubal et al. 2012; Reysenbach et al. 
2006; Yoshida et al. 2006) and detailed characterizations of the mechanisms of insoluble iron 
reduction are not yet available. Arguably, members from the thermoacidophilic archaeal genus 
Acidianus could be considered as a counterpart to the hyperthermophilic, neutrophilic 
Pyrobaculum due to their widespread distribution in terrestrial hydrothermal environments 
(Huber and Stetter 2015), their metabolic flexibility (Amenabar et al. 2018), and several 
physiological studies reporting the capacity of some members of this genus to respire iron 
27 
(hydr)oxide minerals (Amenabar and Boyd 2018; Giaveno et al. 2013; Yoshida et al. 2006). 
Similar to what has been described for many mesophilic iron reducers (Lovley 1991), 
thermophilic DIR organisms tend to prefer poorly crystalline Fe(III) (hydr)oxides (e.g., 
ferrihydrite) over more crystalline Fe(III) (hydr)oxides such as goethite or hematite (Amenabar 
and Boyd 2018; Greene et al. 1997; Kashefi and Lovley 2000). This suggests that constraints 
imposed by the low solubility and bioavailability of iron (hydr)oxide mineral and/or the 
mechanisms of DIR could be similar between mesophilic and thermophilic organisms.  
6.2 Mechanism of Fe reduction during microbial metabolism 
Two general mechanisms of DIR have been described in mesophilic organisms 
(Hernandez and Newman 2001; Weber et al. 2006), including the use of chelators or 
siderophores that solubilize minerals facilitating their transport into the cell (Gralnick and 
Newman 2007) and extracellular electron transfer (EET) based mechanisms (Marsili et al. 2008; 
Reguera et al. 2005; Shi et al. 2016). Although data indicating secretion of compounds that might 
act as siderophores or chelators in thermophilic DIR organisms have not been reported, the 
presence of these types of compounds in the environment might stimulate the reduction of iron 
bearing minerals by increasing the bioavailability of Fe(III) (Gralnick and Newman 2007; 
Kraemer 2004). Here, we focus on EET mechanisms of DIR. 
EET mechanisms vary in whether direct access to the surface of the solid is required and 
if so, the mechanisms by which cells transfer electrons to the mineral surface (Fig. 4). Three 
different EET mechanisms have been proposed in mesophilic iron reducers, several of which 
28 
appear to be shared with their thermophilic counterparts. The first mechanism involves a direct 
electron transfer between a membrane bound electron carrier, such as outer-membrane c-type 
cytochromes, and the solid iron bearing mineral (Fig. 4A) (Beliaev et al. 2001; Magnuson et al. 
2001; Myers and Myers 2001; Richardson 2000). Although this mechanism has not been directly 
reported in thermophilic iron reducers, it is possible that some (hyper)thermophiles, such as 
Pyrobaculum islandicum (Kashefi et al. 2008a), rely on a similar mechanism where direct 
contact is required. This is suggested by electron micrographs showing intimate contact between 
P. islandicum and poorly crystalline Fe(III) oxides during DIR-dependent growth. Similarly,
physiological experiments with iron minerals sequestered inside dialysis membranes indicate that 
Pyrobaculum calidifontis likely requires direct contact with insoluble iron (hydr)oxide for DIR-
dependent growth (Feinberg et al. 2008). Interestingly, unlike mesophilic iron reducers, where c-
type cytochromes are integral components of cellular membranes during electron transport to the 
solid mineral, P. islandicum is not thought to code for c-type cytochromes indicating a different 
mechanism of electron transport during DIR (Childers and Lovley 2001). In contrast to P. 
islandicum, P. calidifontis does code for a putative membrane-bound polyheme c-type 
cytochrome similar to the polyheme c-type cytochromes found in the mesophilic bacterial genera 
Shewanella and Geobacter (Feinberg et al. 2008). 
The second EET mechanism of DIR involves soluble electron shuttles such as flavins and 
quinones (Fig. 4B), where direct contact with the mineral is not required for Fe reduction, such 
as the case of Shewanella (Newman and Kolter 2000; von Canstein et al. 2008). Similar to 
Shewanella, P. aerophilum and Pyrobaculum arsenaticum also can reduce iron (hydr)oxides 
without direct contact to the iron mineral (Feinberg et al. 2008). To this end, P. aerophilum and 
29 
P. arsenaticum might use a similar mechanism involving an electron shuttle as an extracellular 
mediator of iron reduction. Indeed, this mechanism of EET has been suggested for P. aerophilum 
(Feinberg and Holden 2006). 
The third EET mechanism of DIR involves the use of an extracellular component such as 
pili as nanowires that could transfer electrons from the bacterial cell to the solid phase (Fig. 4C), 
such as has been widely described in Geobacter (Gorby et al. 2006; Reguera et al. 2005). 
Although the physical properties conferring conductivity to microbial nanowires have been 
controversial, it has been recently suggested that overlap of the pi-pi orbitals of aromatic amino 
acids in the structural protein PilA from Geobacter imparts metallic-like conductivity to the 
structures, independent of the presence of polyheme c-type cytochrome (Malvankar et al. 2015). 
Recently it has been shown that the thermophilic bacterium Carboxydothermus ferrireducens 
also required direct contact to reduce poor crystalline Fe(III) oxides at moderately elevated 
temperatures (65°C) (Gavrilov et al. 2012). A putatively physiological role of pili-like 
appendages in the extracellular electron transfer to the iron mineral was suggested for this strain. 
Homologs of putative porin-cytochrome complexes, which have been hypothesized to be 
involved in EET and iron reduction in Geobacter (Liu et al. 2014), have been identified in iron 
oxide mats recovered from Chocolate Pots hot spring in YNP (Fortney et al. 2016; Fortney et al. 
2018a; Fortney et al. 2018b). These homologs were related to divergent Ignavibacterium- and 
Thermodesulfovibrio-like populations, among others.  
Most of the work focused on DIR in thermophiles has been conducted on neutrophilic 
strains, where iron (hydr)oxides can represent prevalent sources of electron acceptors. However, 
30 
the otherwise low solubility of iron (hydr)oxide minerals at neutral pH, a condition that likely 
drove the evolution of numerous mechanisms of EET or Fe(III) acquisition, might be overcome 
in acidic conditions since iron solubility is inversely proportional to pH (Kraemer 2004). 
Consistent with this notion, it has been shown that Fe(II) and Fe(III) are more soluble in acidic 
waters (Johnson et al. 2012) and are enriched in many acidic hot springs (Ball et al. 2010). As 
such, many acidic springs have iron (hydr)oxide depositional zones resulting from extensive 
oxidation of Fe(II) and precipitation of Fe(III) (Beam et al. 2016; Inskeep and McDermott 2005; 
Inskeep et al. 2004; Kozubal et al. 2008; Langner et al. 2001). Despite these observations, only a 
few thermoacidophiles have been shown to couple Fe(III) reduction to growth in high-
temperature acidic conditions (Amenabar et al. 2017; Giaveno et al. 2013; Kozubal et al. 2012; 
Yoshida et al. 2006). Moreover, if these thermoacidophiles reduce solid phase iron (hydr)oxides 
or if they reduce Fe(III) released by acid-promoted dissolution of these minerals is not known for 
all described strains. It has been suggested that iron oxide reducing thermoacidophilic 
microorganisms [e.g., Acidicaldus sp. (Kozubal et al. 2012)] might reduce solubilized Fe(III) 
ions or ion complexes generated during acid leaching of the solid phase.  
Reports that Acidianus strain DS80 can grow via DIR without direct access to the mineral 
surface, albeit with slower rates and slower generation times than when provided with access to 
the surface, suggests that these cells can either directly reduce solid phase iron minerals or 
promote their dissolution leading to greater availability of Fe(III) ions for reduction (Amenabar 
and Boyd 2018). Consistent with the notion of cells reducing a solubilized form of Fe and not the 
bulk mineral, microscopic imaging of DIR DS80 cells grown with ferrihydrite reveals that they 
are not intimately associated with the mineral (Amenabar et al. 2017). In addition to ferrihydrite, 
31 
this was found to be true for DIR DS80 cells grown with a variety of Fe(III) minerals as TEA, 
including goethite, and hematite (Amenabar and Boyd 2018). Importantly, however, rates of DIR 
in DS80 cultures varied depending on the Fe(III) mineral that served as TEA. Higher rates of 
reduction were observed with poorly crystalline ferrihydrite than with goethite, and higher rates 
of reduction were observed with goethite than crystalline hematite. These observations are 
consistent with the equilibrium solubilities of the different iron sources tested: ferrihydrite (Ks = 
3.55) > goethite (Ks = 0.36) > hematite (Ks = -0.53) (Kraemer 2004). However, with these data it 
is difficult to deconvolute whether the decrease in Fe reduction rate with increased mineral 
crystallinity is due to an increased ease by which electrons can be deposited into poorly 
crystalline mineral lattices, as has been shown for a Shewanella alga strain (Roden and Zachara 
1996), or if this is due to increased reduction of soluble Fe(III) ions due to differences in mineral 
solubility. Nevertheless, DS80 cells grew via DIR with ferrihydrite sequestered in dialysis 
membranes, albeit at slower rates than when provided direct access to the mineral (Amenabar 
and Boyd 2018). Taken together, these data indicate that under acidic conditions, such as those 
present in the environments where Acidianus DS80 or other thermoacidophiles inhabit, 
microorganisms are poised to reduce soluble Fe(III) ions leached from precipitated iron minerals 
through a proton-promoted dissolution mechanism (Kraemer 2004). It is possible that intimate 
contact with the surface of the mineral allows for electron exchange between Fe(II), Fe(III) ions, 
and/or the Fe(III) solid phase, which may influence the rate of iron oxide dissolution. 
7. Concluding remarks
32 
Numerous lines of evidence from a variety of perspectives point to a role for high 
temperature, hydrothermal conditions in supporting the earliest forms of microbial life. In its 
early stages, microbial life was likely dependent on chemical sources of energy in the form of 
oxidation-reduction reactions involving minerals or substrates generated by interaction of 
minerals with water, the extent of which can be enhanced at high temperature depending on the 
reaction. Terrestrial, volcanic hydrothermal systems integrate subsurface and surface processes 
that together result in extensive geochemical variation, including variable distributions of 
minerals and mineral-sourced substrates capable of supporting microbial metabolism. These 
environments also host non-photosynthetic communities that comprise early evolving microbial 
lineages. To this end, terrestrial, volcanic hydrothermal environments represent convenient 
analogs for developing an understanding of the role of minerals in supporting early forms of 
microbial life. 
Microorganisms and their activities have a profound influence on the formation and fate 
of minerals in hydrothermal environments. The microbial influence on mineral formation and 
dissolution can take place through indirect (abiotic) or direct (biotic) mechanisms, both of which 
are important to take into consideration when thinking about global element cycles or more 
localized element cycling such as within a given hydrothermal system. This is certainly the case 
when considering the iron and sulfur cycles, where feedbacks between abiotic and biotic 
processes not only influence the formation and fate of minerals, they also influence element 
availability for use in microbial metabolism. These considerations, in turn, influence the 
distribution of microorganisms in hydrothermal environments in contemporary settings and may 
33 
have influenced the origin of these metabolic processes and the taxa supported by them in early 
Earth environments. 
8. Acknowledgements
This work was supported by a NSF grant (EAR-1820658) to ESB. MJA acknowledges 
support from the CONICYT Becas-Chile fellowship program. 
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Figures and Figure Legends 
Figure 1. Filamentous and sessile microbial communities growing in close association with 
minerals in a sulfur-rich, acidic hot spring in Yellowstone National Park, Wyoming, U.S.A. The 
source of the hot spring (73°C) is depicted by the grey geyserite minerals at the lower right-hand 
corner of the image. The spring water cools as it flows away from its source, creating distinct 
thermal and chemical zones corresponding to differences in microbiological activity and mineral 
precipitation. The yellow color represents precipitated elemental sulfur (S°) that coats the 
microbial filaments. The brown colored minerals represent iron (hydr)oxides while the green 
color is due to pigments produced by photosynthetic algal communities. The temperature of the 
transition between the yellow filamentous communities and the brown iron (hydr)oxides is 
~65°C whereas the temperature of the transition between the brown iron (hydr)oxides and the 
phototrophic mats is ~52-54°C. Thus, the transition from chemosynthetic to photosynthetic 
communities takes place at ~ 52-54°C in this spring. Photo by Maximiliano Amenabar. Inset, 
scanning electron microscopic image of filamentous sulfur communities. Note the presence of 
rhombic S° crystals. The scale bar is 30 µm. Image by Eric Boyd. 
49 
Figure 2. A model describing the development of geochemical variation in continental, volcanic 
environments such as Yellowstone National Park (YNP), Wyoming, U.S.A. Oxygen-rich 
meteoric water (precipitation) can infiltrate the subsurface where it interacts with hot rocks. Less 
dense heated waters then can ascend to the surface through flow paths along fractured and 
faulted bedrock. During its ascent to the surface, hot fluids interact with bedrock, allowing for 
leaching of minerals and acquisition of solutes. Ascending water can undergo decompressional 
boiling that can separate fluids into a liquid and a vapor phase (phase separation). Ions such as 
Cl- behave conservatively and remain in the liquid phase whereas volatiles, including hydrogen 
sulfide (H2S), partition into the vapor phase. This vapor can also ascend to the surface and 
condense with oxygen-rich meteoric water that can drive the production of sulfuric acid and the 
acidification of hydrothermal waters. Phase separation is thought to be a multi-phase process that 
may affect the amount of sulfide in vapor- or liquid phase-influenced waters and thus the extent 
of possible acidification of those waters. In the schematic, multi-stage phase separation is 
depicted by arrow color with blue depicting circumneutral, Cl- rich waters and green, yellow, 
orange, and red depicting progressively more acidic waters. Condensation of vapor with perched 
aquifers that are progressively more acidified could favor speciation of H2S/HS- (pKa = 6.5 at 
100°C) to the protonated form, thereby increasing the amount that can be volatized in subsequent 
phase separation events. If sulfide input is the primary driver of acidification of spring waters, as 
has been suggested, this could potentially lead to acidification of waters in certain geographic 
locations or springs, as is has beensuggested for the Norris Geyser Basin in YNP (see text for 
details). 
50 
Figure 3. Schematic describing the primary reactions involved in generating bioavailable 
nanoparticulate elemental sulfur (S°) or polysulfide (Sx2-) as a substrate for reduction (R), 
oxidation (Ox), or disproportionation (D) of bulk S° (depicted by image of α-S° powder) as a 
function of environmental pH. (A) Schematic illustrating S° reduction or disproportionation 
under acidic (pH <6.0) or circumneutral to alkaline (pH > 6.0) conditions. Note that reduction or 
disproportionation is not thought to require direct access to bulk S° due to a series of reactions 
that involve H2S/HS- that generate soluble nanoparticulate S° (pH < 6.0) or Sx2- (pH > 6.0) (B) 
Schematic illustrating S° oxidation at acidic pH. Acidophilic organisms that oxidize S° are 
thought to require direct contact with the bulk mineral since substrates capable of driving its 
solubilization via reactions described in (A) (e.g., H2S/HS-) tend to not accumulate in oxidized 
environments. Oxidation of sulfur is not depicted at circumneutral to alkaline conditions since it 
is unlikely that S° will accumulate under these conditions given that the intermediates that lend 
to S° formation via H2S/HS- oxidation (e.g., S2O32-) are stable. Cells are depicted as grey vibrio 
shapes.  
51 
Figure 4. Schematic illustrating the four described mechanisms for electron transport to iron 
(hydr)oxides (illustrated by red substrate) during dissimilatory iron reduction (DIR). Cells are 
depicted by grey vibrio shapes. The four describing mechanisms of DIR include (A) direct 
electron transfer to iron (hydr)oxides, (B) the involvement of compounds (e.g., flavins) that 
shuttle electrons to iron (hydr)oxides, (C) the involvement of extracellular appendages (e.g., 
nanowires) as a conductive matrix for electron transport to iron (hydr)oxides, and (D) the 
involvement of chelators (green circles) that solubilize iron (hydr)oxides for subsequent 
reduction.