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Staging Life in an Early Warm 'Seltzer' Ocean - Elements
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Staging Life in an Early Warm ‘Seltzer’ Ocean

The stage for the origin of life may have been set during a period that was as short as 20 million years within the first 100 million years after the formation of the Moon (at ~4.5 Ga). The atmosphere at that time contained more carbon dioxide than at any other period thereafter. Carbon dioxide sustained greenhouse conditions, accelerated the weathering of a primitive crust, and may have led to conditions conducive to forming the building blocks of life. The conversion of inorganic carbon and nitrogen to the essential building blocks of life may have been facilitated by clays, zeolites, sulfides, and metal alloys that had been formed as the crust reacted with a warm and carbonated (seltzer) ocean. Geochemical modeling constrains the conditions favorable for the formation of these potential mineral catalysts.

DOI: 10.2113/gselements.12.6.395

Keywords: Hadean, origin of life, water–rock interaction, carbon dioxide, catalysis, clay minerals

Introduction

One of the most challenging geological problems is the composition of the atmosphere, ocean, and solid Earth during the period preceding the origin of life. Lacking a rock record reaching back that far in time, we are left to infer the conditions on the prebiotic Earth by applying what can be gleaned from comparative planetary science, zircons that formed as early as 4.4 Ga, and theoretical approaches. The conditions preceding the origin of life constrain how the proverbial “building blocks of life” might have formed via abiotic reactions. Within this context, there has been interest in evaluating possible mineral-assisted or mineral-catalyzed pathways to make amino acids, to form and polymerize nucleotides, and to form other relevant organic compounds as an alternative to delivering organic molecules by meteorites and asteroids (Schoonen et al. 2004; Cleaves et al. 2012). Formed as weathering products, clays may have provided a mechanism to concentrate, stabilize, and protect biomolecules via intralayer sorption (Lyon et al. 2010). It is important to note that mineral-based reactions may also decompose life’s building blocks—a consideration that has attracted far less attention than their formation (Schoonen et al. 2004).

Before proposing a role for minerals in forming, transforming, concentrating, stabilizing, protecting, or decomposing the building blocks of life, it is necessary to evaluate what type of minerals could have been present on the prebiotic Earth. There may have been a range of environments (e.g. various types of submarine hydrothermal vents, surficial rock exposures, and/or hot springs) during this early period, referred to as the Hadean Eon. At first glance, the problem seems intractable and the parameter space too broad to develop meaningful constraints. However, the analyses of zircons formed during the Hadean, the development of advanced models of the atmosphere during the first billion years after the formation of Earth, and experimental work now provide useful insights and much tighter constraints on the origin of life than hypotheses suggested a few decades ago. This article provides an assessment of the factors that determine the composition of secondary mineral assemblages and associated fluids that may have primed early Earth for the emergence of life, as well as an assessment of the environments most likely to lead to formation of reduced carbon and nitrogen compounds.

Conditions on the Early Earth

Earth was too hot for about 2 million years after the Moon-forming event (which took place at ~4.5 Ga) for liquid water to be stable. Once sufficiently cooled, Earth settled into a period of perhaps as long as 20 My to 100 My in which liquid water was present along with a CO2-dominated atmosphere (Zahnle 2006). The partial pressure of CO2 in the atmosphere was initially perhaps as high as 100 atm, subsequently decreasing to levels below 1 atm (Sleep et al. 2001; Zahnle 2006). The rate of CO2 loss would have been governed by the kinetics of water–rock interaction with dissolved CO2 as a reactant.

Elevated CO2 during this period (the early to mid Hadean) may have created favorable conditions for the origin of life. Carbon dioxide is a greenhouse gas and would have contributed to high ocean temperatures, perhaps exceeding 100 °C, with atmospheric pressure well in excess of 1 atm, which would suppress boiling. Water–rock reactions would have consumed CO2 over a period estimated to be about 100 My. During the last 20 My of this transition, with 5–25 atm of CO2 remaining, the temperature of the ocean would have been between 60 °C and 110 °C (Sleep et al. 2001; Zahnle 2006). In essence, this early ocean was a warm carbonated water that reacted with the crust and in so doing lost its fizz. This ocean transition (essentially from fizzy to nonfizzy) shaped the composition of the earliest oceans, the fluids that were entrained in the crust, and the subsequent secondary mineral assemblages. In addition, field studies and experimental work show that, as part of the interaction of CO2-containing water with crustal materials, a range of organic compounds form via abiotic reactions (McCollum 2013). Hence, given the high concentrations of dissolved CO2 in oceans, this relatively short period of time may have produced a significant amount of organic compounds that set the stage for the origin of life.

Apart from the presence of elevated CO2 in the atmosphere, a major control on the outcome of any water–rock process is the chemical and mineralogical composition of the rocks involved. There is no Hadean rock record that would allow one to infer what type of rocks might have made up the Earth’s primitive crust. The only relicts of the Hadean are zircons that formed 4.4–3.5 Ga (Valley 2006). Zircons are common primary accessory minerals in granitoid rocks and can withstand physical and chemical weathering, allowing them to be deposited and redeposited for billions of years. In essence, zircons are small windows into Hadean conditions. Those found in the Jack Hills of Western Australia suggest that liquid water may have been present on Earth as early as 4.2 Ga. While the zircons are invaluable in helping to constrain Hadean conditions, they are likely derived from small vestiges of rocks formed by partial melting of a Hadean crust that was largely ultramafic to mafic in composition (Taylor and McLennan 2009). Ultramafic and mafic igneous rocks are characterized by their high content of magnesium and iron-rich minerals and relatively low silica content (ultramafic rocks contain less silica than mafic rocks).

On the basis of a higher heat flow during the Hadean, it is a popular notion that most of the earliest crust was komatiitic in composition, komatiites being an igneous rock high in Mg and dominated by olivine (Nisbet and Sleep 2001; Meunier et al. 2010). Others have argued that the Hadean crust had a composition similar to mid-ocean ridge basalts (MORB) or tholeiitic basalts (Taylor and McLennan 2009). Peridotite, a mantle material found in modern settings (e.g. the mid-Atlantic off-axis submarine hydrothermal system known as Lost City), may also have been exposed during the Hadean. Rather than choosing one particular rock type, it is prudent to consider all three ultramafic to mafic compositions, as well as a composition that accounts for the small vestiges of continent crust that must have been present to account for the Hadean zircons. The most realistic rock type to represent early vestiges of the continental crust is one of tonalitic composition, with major Na-rich plagioclase and minor quartz and orthoclase (Nutman 2006).

Having settled on a range of rocks to consider and assuming that there was liquid water from 4.2 Ga on, the next step is to decide whether water–rock reactions proceeded under conditions that favor rapid exchange with the atmosphere (hereafter referred to as “open systems”) or conditions that inhibit such exchange (hereafter referred to as “closed systems”). Surficial or shallow environments are expected to have sufficiently rapid gas–water exchange that equilibrium between the fluid and the atmosphere is retained throughout the water–rock interaction. By contrast, in deep environments or in rock-dominated environments (e.g. fluids in cracks) (Fig. 1), the water infiltrating into the rock is in equilibrium with the atmosphere; upon deeper infiltration, however, the fluid evolves out of equilibrium with the atmosphere. In a modern analog, the difference is analogous to modeling reactions in a littoral zone versus reactions in a deep aquifer.

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Figure 1. (Left) Schematic of the geological settings and interactions that need to be evaluated in understanding conditions in, above, and below a carbonated ocean on the ­prebiotic early Earth. Open system is in equilibrium with the atmosphere; closed system is out of equilibrium with the atmosphere. Note gas exchanges between ocean and atmosphere. (Right) The different variables that need to be evaluated when geochemically modeling early Earth carbonated ocean, rock, and atmospheric interactions. The inputs are the different rock types, CO2, and temperature (T); the outputs are the different types of secondary mineral assemblages and fluid compositions.

Figure 1 provides a schematic representation of the settings explored and the variables that must be evaluated when investigating prebiotic early Earth conditions. The purpose of such geochemical calculations is to illustrate the conditions needed to form secondary minerals (clays, zeolites and sulfides) that act as potential catalysts in the synthesis of organic molecules. Besides these minerals, there is an emerging body of literature that shows that metallic alloys form as part of the alteration of ultramafic and mafic rocks (Filippidis 1985; Smirnov et al. 2008). Although small in abundance, these metallic alloys might have played a critical role in the conversion of CO2 and N2 on the early Earth into reactive compounds that could be important to the formation of the building blocks of life. The geochemical calculations also constrain the composition of the resulting fluids (Schoonen et al. 2004; Cleaves et al. 2012).

Modeling Water–Rock Interaction

Building on the decades-long efforts of by various geochemists—Harold Helgeson (deceased) at University of California-Berkeley (USA); Mark Reed at the University of Oregon (USA); Thomas Wolery at the Lawrence Livermore National Laboratory in California (USA); Everett Shock at Washington University and Arizona State University (both USA); Craig Bethke at University of Illinois (USA)—it is now possible to efficiently conduct computer-based water–rock simulations (Bethke 2008). We stress the importance of conducting an ensemble of model calculations so that a range of conditions, such as different rock, gas and solution compositions, can be explored. This is particularly important because the problem being tackled is poorly constrained. For this article, we expand on earlier work in which we explored the interaction of komatiite and tonalite with water in a CO2-rich atmosphere (Schoonen et al. 2004) using Geochemist’s Workbench® (Bethke 2008). The emphasis in these water–rock interaction simulations is on the weathering of major rock-forming minerals. The weathering process governs the composition of the water it reacts with and dictates the suite of secondary minerals formed. It should be kept in mind, however, that these calculations do not take into account any kinetic effects that may lead to nonequilibrium assemblages. The potential problem of nonequilibrium assemblages is commonly addressed in thermodynamic calculations by suppressing the formation of minerals that are known not to form during weathering processes. An example is the formation of quartz. Although quartz is predicted to form on the basis of its thermodynamic stability, other forms of silica, such as chalcedony, are often formed instead when basaltic rocks weather.

The key results of several representative simulations are presented in the form of “heat maps,” a visualization tool widely used in biology to display data. Figure 2 shows the results of simulations with four rock types at constant PCO2 (5 atm) and T (75 °C). For each simulation, the secondary mineral composition after 50 g of rock reacted with a liter of water was converted to relative abundances in terms of weight percentage of major classes of secondary minerals. The molar ratio of Ca + Mg over Na + K (recast as a logarithm) and pH are also shown in Figure 2. In Figure 3, the influence of PCO2 on reactions with one rock type, MORB, is explored, while in Figure 4 the temperature dependence on reactions with MORB is explored at a constant PCO2 (2 atm).

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Figure 2. Color-coded table illustrating the geochemically modelled secondary mineral assemblages and fluid compositions derived from four early Earth rock types (MORB, komatiite, peridotite, and tonalite) under both open and closed systems. Calculations assume 5 atm CO2 in early Earth atmosphere, a temperature of 75 °C, and reaction of 50 g of rock with 1 kg of water. See text and Figure 1 for details on “open” and “closed” conditions.

For the simulations presented here, the reduction of CO2 to methane and other reduced forms of carbon is blocked. This is a reasonable first-order approximation because these reactions do not go forward without a catalyst and are slow even in the presence of a suitable catalyst at the temperatures considered here. Even if the atmosphere contained substantial amounts of molecular hydrogen—perhaps episodically due to meteorite impacts (Zahnle et al. 2010)—CO2 is unlikely to be reduced by hydrogen in open systems without a suitable catalyst and higher temperatures. In closed ultramafic or mafic systems, CO2 reduction to methane is likely (explained below). However, the amount of CO2 in closed systems is limited, and its reduction to methane formation will not materially change the secondary mineral assembly. Hence, as a first approximation, we consider CO2 to be inert, even in closed systems.

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Figure 3. Color-coded table illustrating the geochemically modeled influence of atmospheric CO2 pressure on secondary mineral assemblages and fluid compositions under open and closed system conditions. All calculations assume a MORB rock type, a temperature of 75 °C, and reaction of 50 g of rock with 1 kg of water. See text and Figure 1 for details on “open” and “closed” conditions.

The results of the calculations in which the water–rock interaction takes place at 75 °C while maintaining equilibrium with CO2 at 5 atm (i.e. open system) show that none of the secondary mineral assemblages contain zeolites or clays. The secondary mineral assemblages are instead dominated by a combination of SiO2 and carbonates. The formation of carbonates effectively locks up most of the divalent cations; only iron is present as a metal sulfide. These modeling results are consistent with experimental work on water–basalt interaction at 10 atm CO2 (Gysi and Stefánsson 2012). The secondary mineral assemblage for all starting materials changes considerably if the initial water that had been equilibrated with the CO2-rich atmosphere is not kept in equilibrium with the atmosphere during reaction with the rock. These closed conditions lead to secondary mineral assemblages that have far less carbonates, are essentially devoid of SiO2 and sulfides, and contain abundant clays and/or zeolites. The composition of the resulting fluid also changes depending on whether the system maintains equilibrium with the atmosphere or not. Equilibrium with CO2 throughout the water–rock interaction buffers the pH to values slightly below neutral, while in closed systems the resulting solution is alkaline. Fluids in closed systems tend to be strongly enriched in monovalent cations compared to divalent cations when compared to open systems. In fact, the open system simulation with peridotite suggests that the fluid would be enriched in divalent ions compared to monovalent ions. Solutions dominated by divalent ions may have destabilized primitive membrane vesicles (Monnard et al. 2002) and impeded a critical step toward cellular life (Deamer et al. 2002).

The level of CO2 concentration in the atmosphere is an important factor in determining the composition of the secondary mineral assemblage in addition to the composition of the crust. As CO2 falls from 5 atm to 0.1 atm in open systems (Fig. 3), the secondary mineral assemblage begins to include zeolite at the expense of SiO2 and clays, while the amount of carbonates is also reduced. In closed systems, changes in CO2 are not as consequential.

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Figure 4. Color coded table illustrating the geochemically modeled influence of temperature on secondary mineral assemblages and fluid compositions under open and closed system conditions. All model calculations assume a MORB rock type, a pressure of 2 atm CO2 in the atmosphere, and reaction of 50 g of rock with 1 kg of water. See text and Figure 1 for details on “open” and “closed” conditions.

A decrease in temperature from 75 °C to 0 °C induces some modest changes in the secondary mineral assemblage. The amount of carbonate in closed systems increases with a decrease in temperature. Perhaps the most important change as temperature decreases in open systems is the switch from fluids that are dominated by Na + K to fluids that are dominated by Ca + Mg (Fig. 4). Given that waters dominated by divalent cations destabilize membrane vesicles (Monnard et al. 2002), the warm conditions during this early stage of Earth history would have been more favorable for the assembly of membrane vesicles.

The major implication illustrated by Figures 2–4 is that water–rock interactions during the warm seltzer stage of the early Earth would have produced a range of fluids with contrasting compositions. The ocean waters would have been slightly acidic as a result of the high carbon dioxide content of the atmosphere, while fluids that infiltrated into ultramafic or mafic rocks would have been alkaline (not withstanding some compositional differences). Mixing of these types of fluids would have set up steep chemical gradients and chemical disequilibrium. Russell and coworkers (Russell et al. 2010) have argued that life arose in these mixing zones as a result of the contrasting chemical composition and the resulting chemical disequilibrium.

Formation of Metallic Catalysts

The modeling approach illustrated above allows one to constrain the overall fluid conditions and major secondary mineral assemblages from rock weathering; however, this approach does not address the fate of minor and trace elements incorporated into rock-forming minerals. Some of these minor and trace elements, such as the transition metals, may have played a key role in setting the stage for life. A case in point is the fate of nickel incorporated in olivine, a major mineral component in ultramafic to mafic rocks (e.g. peridotites and basalts). These rocks can contain up to 0.5 wt. % NiO substituted in the structures of olivines and other rock-forming minerals. The forsterite–fayalite–liebenbergite solid solution—commonly referred to as olivine (Mg,Fe,Ni)Si2O4—is particularly reactive because the SiO4 tetrahedra in its structure are not polymerized.

Molecular hydrogen and heat are generated along with metallic secondary minerals when olivine reacts with water, a process referred to as serpentinization (Fig. 5). The molecular hydrogen formation is driven by a reaction between metallic iron and water. Metallic iron itself is a secondary mineral that forms when divalent (ferrous) iron is released from the olivine. The release of heat increases the buoyancy of the resulting solution and triggers an upward movement. The buoyancy can drive solutions out of the rock and into the overlying waters, establishing a circulation pattern independent of tectonic setting. Any nickel contained in the olivine is essentially caught up in this serpentinization process and forms native nickel metal or nickel–iron alloys. The formation of metallic Ni (Ni0) and Ni–Fe alloys, especially awaruite (Ni3Fe to Ni2Fe), has been well documented both in natural (modern and ancient) and in experimentally simulated serpentinization systems (Filippidis 1985; Mevel 2003). Serpentinization of olivine is not unique as a source of mineral-based hydrogen production: such production has been documented for a number of ferrous-iron-containing minerals, including spinel (Mayhew et al. 2013). In the context of staging life, metallic nickel and iron, as well as their alloys, are of particular interest because they facilitate the reduction and hydrogenation reactions of both CO2 and N2 (Horita and Berndt 1999; Smirnov et al. 2008).

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Figure 5. Schematic diagram illustrating the plausible formation of native nickel and nickel–iron alloys as a result of the reaction of olivine with water. After Schrauzer and Guth (1976) and Janecky and Seyfried (1986). The yellow ellipses highlight the fate of the metals involved and the formation of molecular hydrogen. Molecular hydrogen under these conditions reacts with inorganic carbon and nitrogen (Schrenk et al. 2013). The presence of the Ni–Fe alloys, such as awaruite, catalyze these processes. (Inset) Scanning electron microscope image of olivine (large dull crystals covered with a crust secondary Fe minerals) with small (50–100 nm) Ni–Fe–S phases (bright phases) formed during serpentinization in the presence of H2S. From an unpublished experimental study by Smirnov.

The fate of nickel contained in olivine is further influenced by the availability of hydrogen sulfide. In a system free of hydrogen sulfide, Ni0 and awaruite are stable under alkaline to mildly acidic pH conditions, even at temperatures as high as 100 °C. In the presence of hydrogen sulfide, nickel–iron alloys are only stable at alkaline pH, and they are replaced by sulfides of varying Ni:Fe:S ratios at low pH. While the presence of sulfur decreases abundances of pure Fe–Ni alloys/metals, sulfides such as heazlewoodite (Ni3S2), pentlandite [(Fe,Ni)9S8], pyrite (FeS2) or mackinawite (FeS) can also facilitate the formation of prebiotically useful reduced carbon and nitrogen compounds (Schoonen and Xu 2001; Cody et al. 2004; Singireddy et al. 2012; Gordon et al. 2013).

Minerals can participate in the reduction reactions mentioned above either as catalysts or as reactants (Schoonen et al. 1998). A catalyst, by definition, only facilitates reactions without net decrease in its mass or reactive surface area. Therefore, even accessory amounts of Ni–Fe metals and alloys, considered negligible on a planetary scale, could have provided globally important amounts of reaction products, such as ammonia and reduced carbon compounds (Smirnov et al. 2008). Conversely, if Ni–Fe metals and alloys participated as reactants, the yields of prebiotically useful products would be proportional to their spatial and temporal availability, and there should then be a strong correlation between prebiotic products and the rates of reactant formation (e.g. via serpentinization) and destruction (e.g. consumed in reactions). While the complexity and local variability in conditions make it challenging to address which of these modes would have been more prevalent on the early Earth, both the metal catalyst and reactant mechanisms have potential for introducing locally, or even globally, significant concentrations of prebiotically relevant compounds.

Overall, the warm seltzer ocean phase of Earth’s history may have been short-lived, but it could have been crucial in setting the stage for the origin of life, aided by a combination of open and closed systems, each with their own suite of major and minor secondary minerals and fluids. During this stage, the Earth may have been endowed with an assortment of organic molecules, primitive vesicles, and perhaps the first metabolic systems thriving on steep chemical gradients at the interfaces between open and closed systems. The reactions that drive the formation of secondary minerals in closed systems and in mafic rock systems still operate today, but the amount of CO2 dissolved in the infiltrating solutions is significantly lower. Ironically, the closest analogs to the deep past are provided by modern attempts to curb the current buildup of atmospheric CO2 by injecting this greenhouse gas into basalts, where some of it dissolves into water and reacts (Gysi and Stefánsson 2012).

Acknowledgments

NASA’s Exobiology and Astrobiology program provided more than a decade of funding to Schoonen’s group at Stony Brook University to investigate the role of minerals, particularly sulfides, in shaping the conditions during the Hadean through theoretical and experimental approaches. Many students, including Alexander Smirnov, contributed to this research effort. Collaborations, discussions, and student exchanges with Scott McLennan (Stony Brook University), Daniel Strongin (Temple University), John Peters (University of Montana), Hiroshi Ohmoto and Jim Kasting (Penn State), Tom McCollum (University of Colorado), George Cody (Carnegie), and Nita Sahai (Akron University) helped shape our work and thinking on this topic. Alexander Smirnov would like to thank Francis McCubbin (Johnson Space Center) for providing ongoing petrological perspectives into his work. Schoonen is particularly thankful for the mentorship by Hu Barnes (Penn State) and the interaction with the late Dick Holland, who instilled an interest in the geochemistry of the early Earth. Jan Schoonen, Scott McLennan, two reviewers and Element’s editors are thanked for reviewing an earlier draft of the paper. Work on this paper was supported in part by the RIS4E node of the NASA Solar System Exploration Research Virtual Institute (SSERVI). This is SSERVI publication SERVI-2016-032. Brookhaven National Laboratory is supported by the Department of Energy, Office of Science.

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