Russell, MJ, Hall, AJ, and Mellersh, AR, 2003

63 Pages·2005·5.85 MB·English

Russell, M.J., Hall, A.J., and Mellersh, A.R., 2003, On the dissipation of thermal and
chemical energies on the early Earth: The onsets of hydrothermal convection,
chemiosmosis, genetically regulated metabolism and oxygenic photosynthesis, in
Ikan, R., ed., Natural and Laboratory-Simulated Thermal Geochemical Processes:
Dordrecht, Kluwer Academic Publishers, p. 325-388.
ON THE DISSIPATION OF THERMAL AND CHEMICAL ENERGIES ON THE
EARLY EARTH
The onsets of hydrothermal convection, chemiosmosis, genetically regulated metabolism
and oxygenic photosynthesis
M.J. RUSSELL
Scottish Universities Environmental Research Centre
Glasgow, G75 0QF, Scotland
A.J. HALL
Department of Archaeology
University of Glasgow, G12 8QQ, Scotland
A.R. MELLERSH
Chemistry Department
University of Derby, DE22 1GB, U.K.
"It is the inorganic elements that bring organic chemistry to life"
David Garner
1. Introduction
To this day the Earth is kept active by gravitational, radioactive and solar energies. The
convective mass transfer of heat driven by one or more of these forms of energy, from the
very core of our planet through to the upper atmosphere, eventually conduces the
interfacing of the chemical tensions appropriate to the nurturing of life. Without
convection in the Earth's mantle there would be no plate tectonics, no volcanoes, no hot
springs, no mountains — in effect, no fresh surfaces and thus no chemical potential on
Earth to drive the metabolic process. Without advection and convection in the seas and
atmosphere there would be no rain to irrigate the land. And without those spiralling
currents in the liquid iron core there would be no electromagnetic field to protect us from
lethal cosmic radiation. In particular we may say that genetically regulated metabolism and
convection are coupled on our planet and must have always been so, right back to life's
origin (Russell et al. 1988).
Just as convection is inevitable where thermal gradients are steep enough in
particular conditions, so too there is an inexorable drive, through chemiosmosis, toward
metabolism. Metabolism, the combined processes of biosynthesis of complex organic
molecules and their exergonic breakdown within a cell, results from chemical tensions on
this, and for that matter any other wet, rocky, sunlit planet in the Universe: water is
required as polar solvent and to provide hydrogen, protons, electrons and eventually
oxygen; rock to provide the trace elements and reduced molecules; sunlight to split charge.
Metabolism is a downstream corollary of hydrothermal convection and the two processes
have certain parallels. Like steady state convection, life should be recognized as an
emergent dissipative (kinetic) structure — a structure, unlike a crystal, that can bear
perturbation. Metabolism quickens, by many orders of magnitude, oxidation and reduction
reactions on our planet. Convection involves juxtaposed molecules behaving in concert as
they transfer heat to a lower temperature sink. In the metabolic case too, orderly and
commensurate flows through a membrane of electrons, protons, anions, cations as well as
uncharged small molecules, ensures that reactions will fall into step in the kind of
neighbourly co-operation that leads to the rapid dissipation of up to a volt or so of
renewable electrochemical energy.
Thus we may expect the chemical tensions responsible for the onset of life to be
quite evident and various; the components simple, obvious and readily available (Russell
and Hall 1997, 2001; Amend and Shock 2001). Availability of similar components from
more than one source, and appropriate energies derived from several mechanisms, as well
as autocatalytic feedback cycles, confers robustness to the embryonic chemical co-
operative. The way in which life was conceived should not surprise us: a hypothesis
should be economical and not resort to special pleading. More than this, it should be clear
how, given the controlling freedom conferred by genes, life was driven to evolve that it
might tap other energy sources that could contribute to metabolism from further afield and
to tolerate perturbation through the generations. In this contribution we consider the major
evolutionary jumps, from the origin of life itself up to and including the onset of oxygenic
photosynthesis, as the earliest steps are taken up the hierarchy of manageable energy
sources available to life on Earth. During this early exploration, cells enrolled pre-existing
inorganic metal clusters as active centres, became well armoured, ever more able to exploit
the more complex molecules for their operations, and, where illuminated, capable of
converting ultraviolet light to waste heat and, eventually, to use it to generate electrons and
protons from H S and then H O (Fig. 1).
2 2
The initial conditions from which life must have emerged may be gleaned from the
calculable states of the solar system and the Earth in Priscoan (~Hadean) times between
4.5 and 3.9 billion years ago (Ga). Just how the conditions for emergence might translate
to a chemical model and experimental reconstruction of key aspects are addressed. The
ultimate sources of the carbon, hydrogen and oxygen that comprise our organic molecules
was, and is, the carbon dioxide and water that comprise our 'volatisphere'. We attempt to
identify counterparts between our alkaline hydrothermal model for the origin of life with a
proposed nexus of chemical reactors containing heterogeneous catalysts, so that a
reconstruction of how life emerged might be brought about in conditions comparable to
those obtaining on young planets such as Earth or Mars (Shock 1996; Russell and Hall
1997; 1999). The emergence of oxygenic photosynthesis is a biochemical rather than a
geochemical issue. Even so, we speculate that a pre-existing inorganic CaMn complex
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may have been sequestered by proteins to enable the oxidation of water (Russell and Hall
2001). Thus, a suggestion is made as to how a reaction centre belonging to an ancestor of
a purple or green non-sulfur bacterium might chelate such a molecule.
2. The Scale of the Problem
In order to understand how life emerged we need to assess the scales of the various
processes involved. Generally chemists and biochemists, using the 'top-down' or back-
tracking method, have underestimated the spatial scale, overestimated its duration, and
considered a myriad of long-chained hydrocarbons necessary as fuel and building blocks.
The 'bottom-up' approach adopted here, on the other hand, views the initial conditions as
all-important and is open to the notion of miniaturization and sophistication of the
regulated metabolic system during the rapid evolution to the first chemoautotrophic
prokaryotes (Martin and Russell 2003). For our purposes we can ignore the fact that the
Universe itself had to have evolved for hundreds of millions of years before the elements
required for the life process had condensed in isolated giant stars (Riordan and Schramm
1991). We can note in passing the dimensions of the "habitable zone", both in regard to
our galaxy and in our solar system (Gonzalez et al. 2001; Kasting et al. 1993).
Once initial conditions had been met on Earth for life to emerge, then the spatial
dimensions to hold in mind are of a few thousand cubic kilometres to encompass
hydrothermal convection, as well as advection, of ocean water. At the loci of emergence
itself we suggest that a duration of weeks or months is all that is required for an
encapsulated, regulated metabolism to ensue. More than a million hydrothermal seepages
will be conveying energy to the deep ocean at any one time. Geochemical heterogeneities
are to be expected and fluid deliveries will pulsate (Blichert-Toft and Albarède 1994). The
large spatial scale encompasses the geochemical pathways of anabolism as well as the
delivery of the photolytic electron acceptor from the ocean surface. Yet the scale of the
carbon-bearing building blocks is C to C rather than long-chained hydrocarbons. We
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suggest that the energy for emergence is best understood in terms of a chemiosmotic
coupling to provide a protonmotive force through a semiconducting and semipermeable
inorganic membrane, a membrane that separates activated reagents in space (Williams
1961; Mitchell 1967; Russell et al. 1994). Origin-of-life experimentalists have yet to take
the protonmotive force and acid-base catalysis into account.
Fig. 1 The focusing of solar energy to produce a) photolytic iron oxidation and the potentiation of
chemosynthetic life (Cairns-Smith et al. 1992); b) reduction of ferredoxin and the onset of photo-induced
non-cyclic electron transport (Vermaas 2002; Blankenship 2002); c) photo-oxidation of calcium-
manganese bicarbonate and the generation of a precursor to the water oxidising complex (cf. ranciéite)
(Russell and Hall 2001; Dismukes et al. 2001; Sauer and Yachandra 2002); d) oxygenic photosynthesis
through reduction of Mn IV (Blankenship 2002). Iron and manganese are exhaled from hot springs at
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ocean floor spreading centres.
3. The Early Earth
An ocean had precipitated on Earth by 4.4 Ga (Wilde et al. 2001). Although it may have
been vapourized many times over by massive bolide impacts in the ensuing 500 million
years, dust clouds produced by these impacts would have effected rapid cooling of the
deep ocean, during intermissions, to less 20°C (Godderis and Veizer 2000). This cooling
would then have allowed regulated mesophilic metabolic processes to begin at around
40°C (Forterre and Philippe 1999) at a submarine seep. Once conceived in the ocean deep,
microbes could have survived further impacts as they were entrained in ocean waters
percolating to depth in the crust (Parkes et al. 1994; Summit and Baross 2001). There
were no continents (Godderis and Veizer 2000; Kamber et al. 2001) and land masses and
bodies of fresh water, if any, would have been ephemeral in this violent era.
Radioactive heat production within the Earth four billion years ago exceeded
present production more than five-fold (Turcotte 1980). Convection in the mantle
culminated in a spreading rate of ocean crust at a metre or more per year, from numerous
short active centres (Lagabrielle et al. 1997) (Fig. 2). The ocean floor also suffered
extensive submarine volcanic resurfacing. Concomitant destruction occurred over
convective down-draughts hundreds of kilometres from the spreading centres (Abbott and
Hoffman 1984). High temperature submarine hydrothermal convection cells, involving
ocean water, developed to dissipate heat from all these zones. Lower temperature springs
and seepages occurred in the somewhat restricted, quieter conditions of the deep ocean
floor. Yet although thermal energy was dissipated in this way, chemical disequilibrium was
exacerbated. The lower temperature springs and seepages in particular focused a strong
chemical disequilibrium at the ocean floor, a disequilibrium that was to be partially
resolved by the emergence of life through the self-organizing process of chemiosmotic
coupling (Russell et al. 1994). The main redox states of concern involving one or two
electron transfer in the onset of chemosynthetic life, are of iron  as Fe0 and FeII in the
crust, and FeIII from Fe2+ in the ocean, photo-oxidized at a wavelength of 350 to 400 nm,
and deposited as Fe(OH) on the seafloor (Braterman et al. 1983) (Fig. 1).
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Fig. 2 Cross-section of mantle convection cell assumed for the Earth at 4.4 to 4.3 Ga (Smith 1981;
Macleod et al. 1994). Note the warm (50°-100°C) alkaline seepage, one of a myriad in the deep ocean at
which the hydroxide/sulfide/carbonate mounds developed (cf. Kelley et al. 2001; Bounama et al. 2001;
Marteinsson et al. 2001; Geptner et al. 2002).
4. The two classes of hydrothermal convection
The convection of ocean water in fractures transferred a quantity of heat from the fresh hot
crust, in a myriad of open system convection cells, to the intermittently cool ocean where,
ultimately, it was radiated to space. Hydrothermal convection cells, sourced from the ocean
and operating in oceanic crust, self-organize into two distinct classes as described below
(Cathles 1990).
4.1. 400°C HYDROTHERMAL CONVECTION
High enthalpy, high temperature, convection cells are fed from cool ocean water and are
driven directly by magmatic heat. Such cells would have been driven by shallow magmatic
intrusion at oceanic spreading centres and other sites of high temperature intrusion (Barrie
et al. 2001), some of them a response to meteoritic impact (Whitehead et al. 1990; Ames et
al. 1998; Price 2001). Although a portion of this water interacts with invasive magmatic
dykes at up to about 800°C, overall it is physically buffered today at the two-phase
boundary to exhale at ~400°C (Von Damm 2000). These high temperature solutions are
rendered acidic (pH ~ 3) by loss of magnesium from the circulating ocean waters
(Seyfried and Bischoff 1981):
Mg2+ + 2H O → Mg(OH) + 2H+ (1)
2 2 (silicate)
Prior to self-stabilization at its present volume (Kasting and Holm 1992) the earliest ocean
may have been up to three times the present depth (Bounama et al. 2001). And, with heat
flow higher and the crust weaker then, the ocean ridges were probably less salient. If so
then temperatures of exhaling fluids may have been even higher as they tracked the two
phase boundary of seawater toward 800 atmospheres (Bischoff and Rosenbauer 1984),
though rapid mineral precipitation may have prevented such large temperature excursions
(Cathles 1990). Moreover, the intrusion of magma would have rendered the host rocks
plastic and therefore relatively impermeable, although volatiles emanating from the
crystallizing magma could have brecciated this envelope and escaped (Fournier 1999;
Fisher and Becker 2000). Even at 400°C fluids are able to convey up to 25 mM/l of
ferrous iron, and minor to trace quantities of other "biophile" elements such as Mn, Zn, Ni,
Co, Mo, Se and W, to the ocean (Goldschmidt 1937; Hemley et al. 1992; Von Damm
1990). They also carry hydrogen sulfide (generally ~10 mM/l), but in the Priscoan there
would have been no contribution from the thermal reduction of sulfate, so sulfide
concentrations would have been appreciably lower (Walker and Brimblecombe 1985).
Thus very little mineral precipitation would have taken place at, and in the immediate
surrounds of these very hot springs, especially as, unlike modern times, their pH and Eh
would not have contrasted greatly with that of the acidulous Priscoan ocean. This low
oceanic pH of between 5 and 6 (Macleod et al. 1994) was a consequence of the high CO
2
partial pressure on the early Earth (≥1 atmosphere; Walker 1985; Kasting 1993). Thus we
can expect the early ocean to have contained ferrous iron at concentrations between 10 and
20 mM/l, tenors comparable to those in the carbonic lakes with a pH of ~5.5 in Cameroon
(Sigurdsson et al. 1987; Kling et al. 1989). Nevertheless, quenching would have led to the
widespread precipitation of iron monosulfide on the ocean floor (Walker and
Brimblecombe 1985).
4.2. 75° TO 150°C HYDROTHERMAL CONVECTION
The temperature at the base of a hydrothermal cell, reached at a depth of five kilometres or
so in areas of high heat flow but in the absence of magmatic intrusion, depends on the
chemistry and mineralogy of rocks comprising the crust. Though excursions to 200°C are
possible, aqueous fluids are normally buffered at around 75° to 150°C in solid oceanic
crust by serpentinization (hydration and oxidation) and pressure solution of minerals
composing the walls of initially permeable fracture sets which can extend to depths of
several kilometres (Fehn and Cathles 1986). While mafic and ultramafic crust is
particularly prone to hydration, carbonation and oxidation, fractures would have retained
their permeability as the crust was continually flexed by active tectonics as well as the tidal
forces exerted by the close and rapidly orbiting moon (Gaffey 1997). And as we shall see,
alkaline fluids of moderate temperature emanating from the oceanic crust are the likely site
for life's emergence (Russell et al. 1988, 1994; Shock 1992).
To understand the dynamic inter-relatednesses of the processes by which these
fluids attain their chemical make-up we must first investigate the geochemistry of iron, the
most common element with a variable valency. Iron has a negligible solubility in alkaline
solutions (Macleod et al. 1994). Instead it is responsible for the contribution of hydrogen
to these low to moderate temperature hydrothermal solutions. Iron in the ferrous state was
especially concentrated in the silicate and sulfide minerals comprising the Earth's earliest
mantle and crust, to the extent of 10 weight % (Francis et al. 1999). Moreover, vestiges of
meteoritic native iron and nickel were left in the crust as the remainder gravitated to the
Earth's core (Righter et al. 1997). We can think of this Fe0/FeII couple constituting a
hydrogen electrode as hydrothermal solutions began to oxidise the iron with the emission
of hydrogen:
Fe0 + H O → FeO + H ↑ (2)
2 2
and
3Fe0 + 4H O → Fe O + 4H ↑ (3)
2 3 4 2
Of more general significance is the serpentinization of pyroxene, another constituent of the
oceanic crust which has the effect of increasing pH at these temperatures to a value of 10
or more (Neal and Stanger 1984; Macleod et al. 1994):
12Ca Mg Fe Si O + 16H O →
0.25 1.5 0.25 2 6 2
6Mg Si O (OH) + 12SiO + Fe O + 3Ca2+ + 6OH- + H ↑ (4)
3 2 5 4 2 3 4 2
But the early crust also comprised a large fraction of relatively iron-rich olivine (Francis et
al. 1999). Where carbon dioxide was introduced to this crust, hydrogen would have been
joined by methane as a reduced gas. Abrajano et al. (1990) record that methane represents
over half the gas phase emitted during the present-day serpentinization of the Zambales
ophiolite (exposed oceanic crust) in the Philippines. A notional reaction is offered in
equation 5 and geological evidence of such alteration is shown in Figure 3:
6Mg SiO + 12Fe SiO + 14H O + CO →
2 4 2 4 2 2
8Fe O + 4Mg Si O (OH) + 10SiO + CH ↑ + 4H ↑
3 4 3 2 5 4 2 4 2
(5)
We therefore surmise that waters, derived from this same ocean, exothermically
serpentinized the mafic and ultramafic crust to become the alkaline, H - and CH -bearing
2 4
convecting fluids at temperatures peaking at about 150°C (equ. 3), much of which
ultimately seeped into the still carbonic, mildly oxidized, deep ocean (Fig. 4). In theory the
alkaline fluids are capable of dissolving large concentrations of sulfide (as HS-), if
introduced to the base of the cell (as S2-) by magma degassing (Seward and Barnes 1997;
Katsura and Nagashima 1974). Otherwise, concentrations of hydrosulfide (HS-)
generated solely by water-rock reactions can reach 10 mM/l or so (Rahman 2002).
Kelley et al. (2001) have discovered just such a warm (≤75°C) alkaline (pH ~9.8)
spring emanating from 1.5 Myr old crust in the North Atlantic, though lacking in sulfide.
The main precipitates are of CaCO and Mg(OH) . Comparable also is a fresh-water,
3 2
warm (72°C) alkaline (pH 10) submarine spring discovered off the coast of Iceland
characterised by cones of Mg-rich clay tens of metres high (Geptner et al. 2002).
As downward excavation of the ancient hydrothermal cell took place so C
1
molecules previously generated in the crust (section 6.1) could be entrained at its base. For
example, gaseous magmatic emanations can be occluded in glasses and newly crystallized
rock (Kelley and Früh-Green 1999), from where they may be stripped out and entrained
when hydrothermal solutions of medium temperature gain access to the crust. Alternatively
volatiles could leak directly into the base of such hydrothermal convection cells (Gerlach
1989). Appel et al. (2001) have reported the discovery of methane in fluid inclusions
associated with what appears to be a fossil hydrothermal system at least 3.75 billion years
old in the Isua Greenstone Belt, West Greenland.
Fig. 3. Serpentinized olivine (equ. 5).
Fig. 4. Model environment for the emergence of life at a submarine alkaline ~40°C seepage on the floor
of the Priscoan ocean comprising a hydroxide/sulfide/carbonate mound (Russell and Hall 1997). The
acidulous ocean contains the electron acceptor, photolytic FeOOH, that induces chemiosmosis.
Fig. 5. Early Earth as photelectrochemical cell.
Volatiles such as formaldehyde, ammonia and cyanide also may have been present in
certain portions of the oceanic crust (section 6.1). With this chemical and physical
knowledge of the two hydrothermal fluid types as well as of the early ocean we can
investigate how they may have reacted together to produce the first living system.
5. Model for the onset of chemosynthetic life
The Priscoan Earth was a giant photoelectrochemical cell with a potential approaching one
volt, commensurate with the needs of chemosynthetic life (Fig. 5). The atmosphere
comprised ≥1 bar of CO (Walker 1985), with minor concentrations of HCl
2
(Maisonneuve 1982). The ocean was the fluid matrix to a dispersed positive electrode, γ-
FeIIIOOH, generated by the impact of UVC on Fe2+ supplied through 400°C submarine
springs (Cairns-Smith et al. 1992) (Fig. 1):
2Fe2+ + 2H+ + hν → 2FeIII + H2↑ (6)
The consequential hydrogen was lost to space. Yet a hydrogen electrode was maintained at
the cooler, deep, alkaline springs through the reduction of water, a consequence of the
oxidation of ferrous silicates and nickeliferous iron during hydrothermal convection within
the crust (Fig. 4, equations 3-5). Titration of this H -bearing hydrothermal fluid (which
2
also contained CH , HCHO, CH OH, CH COO-, HS-, CH CH S-, NH , CN-, and
4 3 3 3 2 3
simple amino acids and nucleic acid bases generated in the hydrothermal mound) with the
acidulous ocean was inhibited by the precipitation of barriers of sulfides and superposed
clays, as well as the γ-FeIIIOOH and FeII FeIII(OH) which had been eddy-pumped to the
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ocean floor (Fig. 1a). These "precipitate membranes" (cf. Beutner 1913), comprised
essentially of a myriad of nanocrystals of mackinawite (Fe S) (Fig. 6, 7), prevented
1+x
direct neutralization. However, they did permit restricted electron and proton flow as
hydrogen was oxidized by the external FeIII, a reversal of reactions 4-6, there being no
photons to disturb normal redox potentials:
1/2H2 + FeIII → H+ + Fe2+ (7)
Protons carried thus towards the outside of the barrier are partly responsible for the build
up of an inward directed proton potential, and may cause the coupled chemiosmotic
dehydration of inorganic phosphate (Baltscheffsky 1996, and see Josse 1966):
H+ + MgPO OH + PO OH2- → MgP O OH- + H O (8)
3 3 2 6 2
These two separated reactions are linked electrically through the membrane, simplified as:
1/2H2 + FeIII + 2PO3OH2- + H+[out] → Fe2+ + HOP2O63- + H2O + H+[in] (9)
Or, in Mitchell's words, "oxidoreduction (is) coupled to hydrodehydration" (Mitchell
1967), though he was referring to oxidoreduction in aerobic conditions. This (reversible)
coupling through a membrane was, in our view, the fundamental chemical process in
anaerobic oxidative phosphorylation and thus brought about the onset of metabolism. The
protonic potential drove, via the generation of pyrophosphate, organic polymerization. At
the same time a portion of H was activated on the iron monosulfide surface to react with
2
CO (later, aidedby enzymes, with CO ) (cf. Gunter et al. 1987; Bourcier et al. 1987):
2
H2 → H* + H+ + e- (10)
Also formaldehyde and cyanide, supplied from the alkaline solution, were concentrated
Fig. 6. The mackinawite structure, FeS (from Russell et al. 1998). Note that it can contain some Ni and
Co, and minor Mg and Ca, in place of Fe (Morse and Arakaki 1993). Electrons can be transported along
the metal-rich layers in the 'a' and 'b' planes (Ferris et al. 1992). Thus mackinawite could have acted as an
electron transfer agent driving chemiosmosis, Fe(III) acting as an electron acceptor (cf. Fig. 7). It could
also have been responsible for hydrogenations, nickel acting as the catalytic site (cf. Volbeda et al., 1995).
Note too that the double layer of sulfur atoms render mackinawite an insulator along the 'c' axis, perhaps a
factor in maintaining an electrochemical gradient across the inorganic "precipitate membrane".