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Biosphere on Earth

Introduction

NASA Global Biosphere:

"Life is an integral part of the Earth system. Living things influence the composition of the atmosphere by "inhaling" and “exhaling” carbon dioxide and oxygen. They play a part in the water cycle by pulling water from the soil and the air, and they help put it back again by exhaling water vapor and aerating the soil so rain can soak into the ground. They regulate ocean chemistry by taking carbon out of the atmosphere. Earth would not be the planet that it is without its biosphere, the sum of its life."

There are two aspects of interaction between biosphere and climate:

Kleidon A. How does the Earth system generate and maintain thermodynamic disequilibrium and what does it imply for the future of the planet? Philos Trans A Math Phys Eng Sci. 2012 Mar 13;370(1962):1012-40.

After Grenfell JL et al. Co-evolution of atmospheres, life, and climate. Astrobiology. 2010 Jan-Feb;10(1):77-88. (.pdf available on-line)

Sessions AL, Doughty DM, Welander PV, Summons RE, Newman DK. The continuing puzzle of the great oxidation event. Curr Biol. 2009 Jul 28;19(14):R567-74. (.pdf available on-line)

Falkowski PG, Godfrey LV. Electrons, life and the evolution of Earth's oxygen cycle. Philos Trans R Soc Lond B Biol Sci. 2008 Aug 27;363(1504):2705-16. (.pdf available on-line).

Also see "Electrons, life and the evolution of Earth's chemical cycles" (adaptation of Brian Glazer lecture (based on Falkowski and Godfrey (2008)) (.pdf available on-line) .

Thermodynamic disequilibrium as a sign of a habitable planet

Lovelock suggested that in search for habitable planets the chemical disequilibrium in the composition of a planetary atmosphere may be used as a recognizable sign for presence of widespread life on a planet. He argued that the Earth's high concentration of oxygen in combination with other gases constitutes substantial chemical disequilibrium that would quickly be dissipated by simple chemical reactions if it were not continuously replenished by processes generated and maintained by life.

It is accepted that the atmospheric O2 is the main aspect of the Earth system that is maintained far from a state of equilibrium. Photochemistry uses excited electrons that have absorbed solar photons in the visible range and generates the electronic free energy avoiding that the energy of solar photons is directly converted into heat after absorption.

 

Four major drivers are responsible for free energy generation within the Earth system.

ABIOTIC drivers in order of generated power:

  1. Solar heat engines. Solar radiation is associated with spatial and temporal variations that maintain temperature gradients. These gradients are the main driver for climate system processes. The total free energy generated from this source includes the generation of potential free energy associated with air, water vapor and aerosols, the kinetic energy associated with motion in the atmosphere, oceans and river flow, the chemical free energy generated by dehumidification and desalination of sea water, the electric free energy generated by thunderstorms and so on. All of these are fuelled by uneven heating and cooling, resulting in vertical and horizontal gradients in density and pressure. Using a mean surface heating of 170Wm−2 and typical temperatures of 288 and 255 K, Kleidon estimates that free energy generation from this vertical gradient is less than 5000TW (note that much of this power is used for vertical mixing and transport and is likely to contribute relatively little to large-scale cycling and transport). Owing to the planet's geometry and rotation, absorption of incident solar radiation results in horizontal gradients. Using a mean difference in solar radiation of about 40% between the tropics and the poles yields an upper limit of about 900TW. The temporal variation of heating in time yields an additional power of about 170TW at maximum efficiency, so that the total power generated from radiative heating gradients is of the order of 6000TW.
  2. Interior heat engines. Radioactive decay, crystallization of the core and secular cooling of the interior provide means to generate free energy within the interior. This free energy is associated with the kinetic energy of mantle convection and plate tectonics, with potential free energy generation associated with plate tectonics, with magnetic free energy generation associated with the maintenance of the Earth's magnetic field, and with geochemical free energy generation associated with metamorphosis and other geochemical transformations. Given that the heat flux from the interior is less than 0.1Wm−2 at the Earth's surface, maximum efficiency estimates by Dyke et al. yield a maximum generation rate of free energy in the various forms of about 40TW most of which are involved in the generation of kinetic energy associated with mantle convection and plate tectonics.
  3. Gravitational engines. Gravitational forces by the Moon and the Sun provide some additional free energy by generating potential free energy mostly in the ocean in the form of tides. Estimates place the total generation rate at around 5TW.

BIOTIC: free energy produced and stored by photosynthesis

Solar photochemical engines. Incident solar radiation contains wavelengths that can be used to generate chemical free energy when visible or ultraviolet radiation is absorbed by electronic absorption or photodissociation. Photodissociation can, in principle, generate radicals that are associated with free energy, but it is omitted here because those compounds have very short residence times and therefore unlikely to result in sustained free energy generation of significant magnitude. Photosynthesis is able to generate longer lasting free energy using complex photochemistry that prevents rapid dissipation. Using typical values for global gross primary productivity and typical free energy content of carbohydrates yields a generation rate of chemical free energy of about 215TW.

Summary

In summary, the total free energy generation for current conditions yield about Pgeo, a ≈6070TW of power by physical processes within the atmosphere from the exchange fluxes at the Earth–space boundary, about Pbio ≈215TW of chemical free energy by photosynthesis, and about Pgeo,b ≈40TW driven by the depletion of initial conditions in the interior. Hence, the total power generation by the planet is about Pplanet = Pgeo,a + Pgeo,b + Pbio ≈6325TW.

Only a small fraction of the total generated power Pplanet is available for geochemical transformations. Of the ≈6070TW of geophysical free energy generated within the Earth's atmosphere, most is likely to be dissipated by atmospheric convection. About 1000TW are dissipated by frictional dissipation by the large-scale atmospheric circulation, 65TW are transferred into the oceans to generate waves and maintain the wind-driven circulation, about 560TW are associated with lifting water to the height at which it condenses and precipitates to the ground, and about 27TW are associated with desalinating seawater.

In contrast to these very small generation terms of chemical free energy, biotic productivity generates 215TW of chemical free energy. Not all of this free energy is available for geochemical transformations of the environment, as the metabolic activity of organisms consumes about half of the generated free energy. This contribution is nevertheless likely to be 1–2 orders of magnitude larger than abiotic means of geochemical free energy generation. Hence, this estimate substantiates the suggestion given by Lovelock that the Earth’s planetary geochemical disequilibrium is mostly attributable to the presence of widespread life on the planet.

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Co-evolution of atmosphere, life, and climate

After Earth's origin, the Sun, was 20-25% dimmer than today. Without greenhouse-like conditions to warm the atmosphere, the early Earth would have been an ice ball and life may never have evolved. Origin of life indicates that greenhouse gases must have been present. Evidence from the geologic record indicates an abundance of the greenhouse gas CO2. CH4 should have been present as well. Also, in this regard methanogenic bacteria may have contributed to modification of the early atmosphere later on.

Molecular hydrogen was likely the third main component of Earth’s prebiotic atmosphere; In an enhanced presence of CO2 and if large amounts of hydrogen were absent in the upper atmosphere, the exosphere may have been relatively cool, which could have resulted in low escape rates of atomic hydrogen (Tian et al. 2005). Hydrogen that was released by volcanoes but not efficiently lost to space must have accumulated to levels of the order of 200 mb (Tian et al. 2005). As water was present on Earth before 4.4 Gyr, water vapor was also an important constituent of the lower atmosphere. Other atmospheric species that may have been present at this time include: CO and sulphur-bearing species like H2S released by volcanoes and possibly methane produced abiotically in hydrothermal vents.

Molecular oxygen was not present, as is indicated by the study of rocks from that era, which contain iron carbonate rather than iron oxide. The development of photosynthesis allowed the Sun's energy to be harvested directly by life forms. The resultant oxygen accumulated in the atmosphere and formed the ozone layer in the upper atmosphere. Multicellular organisms evolved originally in the ancient water bodies. Aided by the absorption of harmful UV radiation in the ozone layer, primitive life forms colonized Earth's surface.

Nitrogen on Earth was outgassed during the first hundred million years. Therefore, the atmospheric pressure was at least 0.8 bar in Earth's prebiotic atmosphere. If climate regulation via the carbonate-silicate cycle is assumed (Walker et al. 1977), then the level of CO2 is determined to a first approximation by the solar luminosity.

Early Earth's atmosphere is still a big mystery because to be conductive to the evolution of life not only significant greenhouse effect must have been present but also its chemical content should have allowed for abiotic synthesis of primordial organic molecules. If there were more CH4 (more potent greenhouse gas) and less CO2 there would be a thick organic haze which would cool the planet accompanied with rapid escape of H2 to space. If there were more CO2 and less CH4 then life would much less likely to occur because of absence of necessary chemical conditions to ensure abiotic formation of organic molecules.

Methane

Methane is a trace gas in the present Earth atmosphere (about 2 ppm), and its origin is biological except for a small fraction produced in hydrothermal systems. Methanogens are among the most primitive Archaea found in the tree of life, some of which are autotrophic (consuming CO2 and H2) and others heterotrophic (consuming organic molecules).

If a biogenic release were equal to the present day, the level of methane would have reached 100 - 1000 PAL (Present Atmospheric Level) in the absence of atmospheric O2 (Pavlov et al. 2000). As today's methanogens can only grow in very limited environments where O2 is absent and H2 or organics are present, the production of methane by the biosphere was probably much higher in the early anaerobic environment. Thus, very high levels of methane can be inferred, which lasted for more than 1 Gyr, between the emergence of methanogens (probably earlier than 3.4 Gyr) and the rise of O2 (2.3 Gyr).

Thus, it might be possible that biological methanogens contributed to the oxidation of the atmosphere and lithosphere and enhanced the loss of H2, making possible, later, for the rise of oxygen (Carling et al. 2001).

Oxygen

A naive chemist examining the atmosphere on Earth may be completely surprised that the two most abundant gases are N2 and O2. N2 behaves as a noble gas and it is virtually non-reactive. Geochemists assume that the amount of N2 in the atmosphere has remained constant since the planet was first formed about 4.6 Ga ago. Indeed, the turnover time for N2 in the atmosphere is estimated to be ca 1 Ga (Berner 2006). By contrast, O2, the second most abundant gas in Earth's atmosphere, is highly reactive, and without a continuous source would become rapidly depleted (Keeling et al. 1993). This gas exists far from thermodynamic equilibrium with a virtually infinite source of reductant in Earth's mantle. Indeed, high concentrations of gaseous diatomic oxygen are unique to this planet in our Solar System and this feature of our planetary atmosphere has not yet been found on any other planet within approximately 20 parsecs of our Solar System (Kasting 1993). The presence of high concentrations of the gas in a planetary atmosphere is presently understood to be a virtually irrefutable indication of life on other terrestrial planets.

Geological records have revealed the chemical action of free oxygen after about 2.3 Gyr ago (Bekker et al. 2004), except for some deposits from the deep ocean that remained anoxic for a few hundred million years or more (Rouxel et al. 2005).

The reactivity of oxygen is driven by electron transfer (redox) reactions, leading to highly stable products, such as H2O, CO2, HNO3, H2SO4 and H3PO4. The abiotic reactions of oxygen often involve unstable reactive intermediates such as H2O2, NO, NO2, CO and SO2. The reactions of oxygen with the other abundant light elements are almost always exergonic, meaning that, in contrast to N2, without a continuous source, free molecular oxygen would be depleted from Earth's atmosphere within a few million years.

Indeed, 2.7 Gyr old molecular fossils are interpreted as the remains of primitive cyanobacteria and eukaryotes, which are producers and consumers of O2, respectively (Brocks et al. 1999). Several reasons could explain this delay. First, the budget reaction of oxygenic photosynthesis also works in the reverse direction, since respiration and oxidation of organic sediments consume oxygen.

This occurs today at a rate of 589 Tg O2/year (Catling and Claire 2005), which means that the net release of the present O2 atmospheric content (1021g O2) takes about 2 Myr (this is about 1000 times slower than the production of O2 by photosynthesis). This rate is balanced by the oxidation of rocks, old sediments, and volcanic gases. The oxidation sinks for O2 may have been much more efficient on early Earth, partly due to the presence of large amounts of reduced iron in the ocean and the crust (Walker 1977).

Some tectonic processes may have favored the burial of reduced carbon and allowed the rise of O2 by about 2 Gyr ago (Des Marais et al. 1993). Another hypothesis has already been mentioned and is linked with the slow oxidation of Earth through the escape of hydrogen to space: in other words,

There might also be a climatic reason for the delay between the emergence of O2- producers and the rise of O2. CH4 has a very short photochemical lifetime in an O2-rich atmosphere, which means that a consequence of a build up of O2 is a decrease of the CH4 atmospheric abundance and, thus, a fall of the mean surface temperature that could lead to a global freezing event.

Therefore, in a biosphere where CH4 and O2 producers both exist, the solar luminosity might be a strong constraint on the timing of the oxygenation (Selsis 2002). Some authors argue that complex multicellular life can only develop in an oxic environment (e.g., Catling et al. 2005).

Geological records have provided us with only qualitative information about the presence or absence of oxygen in the atmosphere. After the rise of O2 and until the end of the Precambrian (about 550 Myr ago), it can only be inferred that the level of O2 was about 1% PAL (present atmospheric level)(0.2 % in abundance). During the Phanerozoic (from - 550 Myr to now), models based on the chemical and isotopic composition of sedimentary rocks allow us to trace back the evolution of the level of O2 and show that it has varied roughly between 0.7 and 1.8 PAL (Berner et al. 2003).

The principal cause of this enhanced level was the rise of large vascular land plants and the consequent increased global burial of organic matter. Higher levels of O2 are consistent with the presence of Permo-Carboniferous (300 Myr ago) giant insects.

Ozone (O3)

Ozone (O3) is produced by only one reaction: O + O2 + M -> O3 + M, where M is any compound (this is called a 3-body or association reaction) and atomic oxygen comes from the photolysis of O2. On the other hand, O3 is destroyed by photolysis and many trace compounds in the atmosphere (HOx, NOx, ClOx, …).

The variation of O2 alone during the last 550 Myr would not have changed significantly the level of O3, as ozone has only a weak dependency on the O2 level (Léger et al. 1993; Segura et al. 2003).

The abundance of O3 was more certainly affected by changes in the trace gases content. We do not know for sure whether, between the rise of O2 and the beginning of the Phanerozoic, O3 provided a UV shield for land life, but it can be inferred that it was present when the first lichens colonized the lands during the Ordovician (500 - 425 Myr ago).

Co-evolution of life and climate on Earth
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Oxygenic photosynthesis

The atmosphere has apparently been oxygenated since the 'Great Oxidation Event' ca 2.4 Ga ago, but it appears that the photosynthetic oxygen production began well before that. However, geological and geochemical evidence from older sedimentary rocks indicates that oxygenic photosynthesis evolved well before this oxygenation event. Fluid-inclusion oils in ca 2.45 Ga sandstones contain hydrocarbon biomarkers evidently sourced from similarly ancient kerogen, preserved without subsequent contamination, and derived from organisms producing and requiring molecular oxygen. Mo and Re abundances and sulphur isotope systematics of slightly older (2.5 Ga) kerogenous shales record a transient pulse of atmospheric oxygen. As early as 2.7 Ga, stromatolites and biomarkers from evaporative lake sediments deficient in exogenous reducing power strongly imply that oxygen-producing c yanobacteria had already evolved. Even at 3.2 Ga, thick and widespread kerogenous shales are consistent with aerobic photoautrophic marine plankton, and U-Pb data from 3.8 Ga metasediments suggest that this metabolism could have arisen by the start of the geological record. Hence, the hypothesis that oxygenic photosynthesis evolved well before the atmosphere became permanently oxygenated seems well supported.

The rise of atmospheric O2 was a milestone in the history of life. Although O2 itself is not a climate-active gas, its appearance would have removed a methane greenhouse present on the early Earth and potentially led to dramatic cooling. Moreover, by fundamentally altering the biogeochemical cycles of C, N, S and Fe, its rise first in the atmosphere and later in the oceans would also have had important indirect effects on Earth's climate. On the early Earth, atmospheric O2 would initially have been very low, probably <10-5 of the present atmospheric level (PAL). Around 2.45 Ga years ago, atmospheric O2 rose suddenly in what is now termed the Great Oxidation Event.

Of all the biochemical inventions in the history of life, the machinery to oxidize water — photosystem II — using sunlight is surely one of the grandest. Not only did the ability to use water as a fuel provide early cyanobacteria with the advantage of an almost limitless supply of energy, but the production of O2 as a waste product also profoundly changed the composition of the world's oceans, continents and atmosphere. The rise of atmospheric O2 also had important indirect effects on Earth's climate. An anoxic atmosphere on the early Earth would likely have contained significant amounts of methane, which is a potent greenhouse gas. Increasing pO2 would have removed this methane greenhouse, possibly triggering dramatic cooling. Indeed, a series of near-global glaciations are believed to have occurred at about the same time that atmospheric O2 first rose significantly.

It remains unclear how the evolution of the two photosystems that gave rise to oxygenic photosynthesis within a single clade of bacteria came about (Blankenship 2001; Falkowski & Knoll 2007). The prevailing hypothesis, based largely on structural studies of the reaction centers, is that oxygenic photosynthesis resulted from lateral gene transfer between a purple non-sulphur bacterium with a quinone-based reaction center and a green sulphur bacterium with an iron–sulphur-based reaction centre, giving rise to a chimeric organism. However, while both the electron transfer processes and amino acid sequences of the two reaction center complexes are significantly different, the structural homology between them is strikingly similar, suggesting that they may have evolved from a single ancestor in one organism via gene duplication events, followed by divergent evolution (Blankenship et al. 2007). Whatever process led to oxygenic photosynthesis, this energy transduction machine is undoubtedly the most complex in nature. In extant cyanobacteria, well over 100 genes are required for the construction of the protein scaffolds as well as the enzymes required for biosynthesis of the prosthetic groups (Shi et al. 2005). Consequently, oxygenic photosynthesis is, unlike any other core prokaryotic metabolic pathway, completely isolated to cyanobacteria.

Regardless of the pathway(s) that led to the evolution of oxygenic photosynthesis, the biological process was a necessary, but not sufficient, condition to allow for a net accumulation of oxygen in Earth’s atmosphere. The production and consumption of oxygen are almost always very closely balanced on local scales. Indeed, on time scales of millions of years, there is virtually no change in the oxygen concentration of the atmosphere or in its isotopic composition (Bender & Sowers 1994; Falkowski et al. 2005). These results strongly suggest that photosynthesis and respiration are extremely tightly coupled. On longer geological time scales, however, there are net changes in atmospheric oxygen. A net accumulation of oxygen in the atmosphere requires a net sink for reductants; that is, there must be an imbalance between oxygen production and its biological and abiological consumption (Holland 2006). Indeed, the very presence of oxygen in the atmosphere implies a permanent sink for reductants (hydrogen atoms). The primary biological sink for the reductants generated by oxygenic photoautotrophs is carbon dioxide, leading to the formation of reduced (organic) carbon (Berner 2004). Indeed, the counterpart of the story of O is, in large measure, the formation and sequestration of C–H bonds. Mixing of these two reservoirs would, purely from a thermodynamic perspective, consume both pools, leading to the formation of water and inorganic carbon. Hence, a permanent reservoir for organic matter is imperative if oxygen is to accumulate from water splitting coupled to carbon fixation.

By far, the largest sink for organic matter is the lithosphere (Falkowski & Raven 2007). In the contemporary ocean, the vast majority (more than 99%) of organic matter produced by photosynthetic organisms is respired in the water column or in the sediments. However, a very small fraction is buried in sediments, especially along continental margins and in shallow seas (Hedges & Keil 1995; Aller 1998). Similar processes must have operated in the Late Archaean and Early Proterozoic oceans ca 2.5 Ga ago, but some key factors differed.

It is currently thought that during the Late Archaean and Early Proterozoic, primary production in general was exclusively conducted by prokaryotes, and oxygenic photosynthesis specifically was carried out exclusively by cyanobacteria (Knoll & Bauld 1989). Grazing pressure on these organisms was almost certainly nil; there were no metazoan grazers. However, there may have been eukaryotic heterotrophic or anaerobic photosynthetic protists (Fenchel & Finlay 1995). A small fraction of the organic matter sank out of the water column (and it probably was a much smaller fraction than in the contemporary ocean) to become incorporated into sediments. The rate of delivery of sediments was also probably much slower than in the contemporary oceans; there were no soils on the continents and no terrestrial plants to help catalyse weathering reactions (Berner 1997). Abiological authogenic carbonate precipitation on biological particles may have been a major source of sedimentary materials that could have facilitated sinking of organic matter.

It is now well accepted that O2 would at first have been toxic as well as a rich potential substrate for cellular respiration and biosynthesis. A dual-edged sword, it provided the driving force for the evolution of mechanisms both to detoxify and respire it. Because oxygen detoxification mechanisms were presumably limited in an anaerobic world, a significant rise in O2 may have caused a cataclysm in those organisms that were first exposed to reactive oxygen species. Organisms that could not adapt to O2 were forced to remain in anaerobic niches, while others evolved detoxification mechanisms ranging from enzymatic (for example, catalase and superoxide dismutase) to behavorial (for example, clustering together) to survive in the newly aerobic environment.

However, aerobic respiration was such a metabolic breakthrough (it is estimated to be about 16 times more efficient in generating ATP than anaerobic fermentation) - that some have argued this permitted the development of complex eukaryotic life, although others have proposed the first eukaryotic cell arose from an anaerobic symbiosis. In any way, the eventual appearance of eukaryotic cells with the capacity to respire O2 is generally agreed to have permitted the evolution of complex multicellular life and its eventual migration into terrestrial ecosystems.

While photosynthesis is the primary source of O2, there are multiple environmental sinks, most significantly reduced carbon, iron and sulfur. The re-oxidation of biologic organic matter is the stoichiometric inverse of photosynthesis, so a first requirement for the net accumulation of O2 at the Earth's surface is the burial of reduced organic carbon for geologic timescales. This burial process may have been much less efficient in the Precambrian than today due to a lack of biomineralization, which in the modern oceans provides ballast for sinking particles. A second requirement is that the net production of O2 must also have exceeded the delivery of reduced inorganic species from deep in the Earth by geologic processes. Fluxes of reduced Fe and S from the mantle, mainly at volcanoes and mid-ocean ridges, would also have been larger on the ancient Earth and presumably limited (or prevented) the accumulation of O2.

Oxygenation timeline

Although large uncertainties surround estimates of ancient O2 levels, a general consensus now exists that the atmosphere of the earliest Earth was almost certainly anoxic, containing <0.1% of the present atmospheric level (PAL) and probably very much less (<10-5 PAL). Shortly after 2.45 Ga, atmospheric O2 rose rapidly and substantially to probably a few percent or more of PAL. Due to its apparent speed and singular nature, this time interval has come to be known as the Great Oxidation Event.

By 1.8 Ga atmospheric O2 appears to have stabilized, possibly in the range of 5-18% PAL, while the deep oceans remained anoxic. A second significant increase in atmospheric O2 is postulated at around 0.6-0.8 Ga, and was accompanied by the oxygenation of the deep oceans and emergence of multicellular animals. By around 0.5 Ga, atmospheric O2 was probably near its present level (21%), and has fluctuated around that value (15–35%) for the past half billion years.

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Nitrogen

The biogeochemical cycles of H, C, N, O and S are coupled via biologically catalysed electron transfer (redox) reactions. The metabolic processes responsible for maintaining these cycles evolved over the first ca 2.3 Ga of Earth's history in prokaryotes and, through a sequence of events, led to the production of oxygen via the photobiologically catalysed oxidation of water. Oxygen accumulation caused fundamental alterations in the nitrogen cycle. In the latter case, the presence of free molecular oxygen allowed ammonium to be oxidized to nitrate and subsequently denitrified. The interaction between the oxygen and nitrogen cycles in particular led to a negative feedback, in which increased production of oxygen led to decreased fixed inorganic nitrogen in the oceans. This feedback, which is supported by isotopic analyses of fixed nitrogen in sedimentary rocks from the Late Archaean, continues to the present. However, once sufficient oxygen accumulated in Earth’s atmosphere to allow nitrification to out-compete denitrification, a new stable electron ‘market’ emerged in which oxygenic photosynthesis and aerobic respiration ultimately spread via endosymbiotic events and massive lateral gene transfer to eukaryotic host cells, allowing the evolution of complex (i.e. animal) life forms. The resulting network of electron transfers led a gas composition of Earth’s atmosphere that is far from thermodynamic equilibrium (i.e. it is an emergent property), yet is relatively stable on geological time scales. The early coevolution of the C, N and O cycles, and the resulting nonequilibrium gaseous by-products can be used as a guide to search for the presence of life on terrestrial planets outside of our Solar System.

The production and burial of organic matter in the ocean are not simply dependent on carbon dioxide being reduced and buried. Marine microbes, be they prokaryotes or eukaryotes, are, from a biogeochemical perspective, primarily porous bags containing enzymes that allow exchange of gases with the environment. Enzymes, being proteins, require a source of fixed nitrogen in the form of NH4 for their synthesis. Ultimately, the source of NH4 is biological reduction of N2. The enzyme system responsible for this reaction, nitrogenase, is widely dispersed among bacteria and archaea (Postgate 1998) and contains 19 iron–sulphur clusters, which upon exposure to molecular oxygen become oxidized, thereby rendering the enzyme complex catalytically inert.

The enzyme can use free H2 to fix N2; however, free H2 is very rare on Earth at present and probably has not been a major gas in Earth's atmosphere for at least 3 Ga. In vitro, the enzyme can be made to behave as an H2-dependent ATPase (Mortenson 1964), but in vivo, the source of the reductant ultimately is organic carbon. Hence, the reduction of N2, a reaction that is arguably as critical to the evolution and perpetuation of life on Earth as photosynthesis, is metabolically coupled to the oxidation of organic carbon. Further, given the extremely high-energy requirements of nitrogenase (a requirement that is not well understood from a mechanistic perspective), in the absence of a high ATP flux, the enzyme would be extremely inefficient. Under strictly anaerobic conditions, heterotrophic ATP supply is coupled to substrate phosphorylation, and only approximately two ATPs can be generated per glucose oxidized. The best source of ATP is proton-coupled phosphorylation; the best supply of the energy is photosynthesis. Indeed, nitrogenase is contained in some, but not all cyanobacteria.

In modelling, the N cycle as a function of O2, Fennel et al. (2005) suggest that there was a 'node'; that is, as free molecular oxygen began to accumulate, nitrate would have initially formed from the abundant pool of ammonium, but as the oxidation state was low, denitrifiers would have efficiently removed the nitrate to form N2. The model indicates that there is a critical concentration of O2 above which nitrifiers successfully out-compete denitrifiers, allowing nitrate to remain as a stable intermediate in the dissolved phase of the oceans; based on experimental studies with extant microbes, that threshold concentration of O2 is between approximately 15 and 30 mM. Hence, until sufficient O2 accumulated in the upper ocean and atmosphere, nitrate would be rapidly depleted, causing a global 'nitrogen crisis', in which virtually all fixed inorganic nitrogen would be removed by denitrifiers. The amount of oxygen that could be produced under the most ideal conditions would have been constrained by the phosphate inventory of the ocean, while the rate of accumulation of O2 in the atmosphere would have been constrained by the rate of burial of organic matter. In practical terms, the latter is dependent on the areal extent of shallow continental shelves, which are the main repository of organic matter in ocean basins.

The superplume event that approximately coincided with the appearance of cyanobacteria may have also played a significant role in tipping the oxidation state of Earth. This event appears to have initiated the onset of 'modern style' plate tectonics, the growth of cratons that led to the creation of wide shelves and carbonate platforms (Eriksson et al. 2006), allowing for massive burial of organic C, a key step in the accumulation of oxygen in Earth's atmosphere (Fennel et al. 2005). From 2.4 to 2.6 Ga, large epeiric seas occupied the passive margins of the newly forming supercontinent Kenorland and its fragments that hosted thick successions of clastic and chemical sediments. As organic matter sank from the surface ocean and was sequestered in shelf and shallow sea sediments, two processes occurred: (i) the primary driver of denitrification, organic C, was removed and isolated from available nitrate and (ii) faster remineralization of N relative to C returned fixed N to the ocean. Increased shelf slope upwelling and recycling of nutrients within these shallow seas would have increased the production of O2 and nitrate, but rates of denitrification decreased owing to the sequestration and burial of Corg. Thus, fixed N, and nitrate in particular, could slowly start to accumulate in the upper ocean; a condition that was stabilized as O2 itself finally began to accumulate in Earth's atmosphere.

Evolution

In general, the biological evolution of Earth can be divided into two major evolutionary eons (Falkowski 2006). The first half, up until the oxidation of the atmosphere, we can classify as an 'eon of biological innovation', when the major metabolic pathways of life that facilitate all the electron transfers, which would become biogeochemical cycles, evolved. The second half of Earth's history is one of 'biological adaptation', where organisms appropriated metabolic pathways that evolved early on, but 'repackaged' them in new body plans which allowed them to be carried forward. Before discussing how these 'insurance policies' that preserved core metabolic pathways through a wide variety of planetary traumas evolved, let us consider the problem of how the oxygen cycle interacted with the metabolic pathways which evolved, for the most part, under anaerobic conditions.

So why has not nature developed a better oxygen evolving machine? The most probable answer is: 'Because it can't'. Over 2.4 Ga of evolution, mutations of every amino acid in every combination of positions possible surely occurred, yet D1 is one of the most conserved proteins in the photosynthetic apparatus; the sequences are more than 80% identical between cyanobacteria and higher plants, and approximately 90% similar. The extraordinary conservation of this protein, which is irreversibly damaged by a product of its own catalysis, suggests that the restrictions imposed by protein– protein, protein–prosthetic group and protein–lipid interactions cannot be overcome.

Are there other examples of the effects of oxygen in the electron market of life on Earth? Indeed, there are several. Rubisco is an enzyme that evolved under anoxic conditions; in the presence of oxygen, rather than carboxylating ribulose 1,5-bisphosphate to form two molecules of 3-phosphoglycerate, the enzyme can catalyse a non-productive oxygenase reaction leading to the formation of glycerate and glycolate (Tabita et al. 2008). These two products are respired, leading to an unproductive consumption of photosynthetically generated reducing power. Indeed, in the contemporary atmosphere, upwards of 40% of Rubisco activity is oxygenase in a typical C3 terrestrial plant. Many algae have developed carbon concentrating mechanisms (Raven et al. 2008), which allow the cell to increase the concentration of CO2 in the vicinity of Rubisco (Coleman 1991); this process suppresses but does not eliminate the oxygenase activity of the enzyme. The primary mechanism for reducing the impact of the oxygenase activity is to synthesize large amounts of Rubisco (Tabita et al. 2008); indeed, this is probably the most abundant protein in nature. Such a solution requires a significant investment in fixed inorganic nitrogen and as nitrogen is frequently itself a limiting element, this investment comes at the expense of growth.

Another example is the nitrogen-fixing enzyme, nitrogenase, a highly conserved heterodimer, which invariably contains Fe–S motifs (Peters et al. 1995). Fixation of nitrogen from the atmospheric N2 reservoir is critical for life on Earth, yet the enzyme is relatively rare in nature (Capone & Carpenter 1982). When exposed to O2, nitrogenase is irreversibly inactivated due to the oxidation of Fe(II ) in the Fe–S clusters. Indeed, at 21% oxygen, approximately 30% of the nitrogenase activity is lost in cyanobacteria (Berman- Frank et al. 2003). To compensate for this loss, the cells must overproduce the enzyme, at a great expense to their own nitrogen economy. Even after 2.4 Ga of evolution, the solutions allowing nitrogenases to function in an aerobic environment are invariably to reduce the partial pressure of oxygen in the immediate vicinity of the enzyme; this is achieved primarily through metabolic processes such as the Mehler reaction (a photosynthetic reaction that reduces O2 back to H2O, usually via an H2O2 intermediate), or by aerobic respiration (Milligan et al. 2007).

Thus, while the evolution of the oxygen cycle on Earth ultimately facilitated a much more efficient pathway for the oxidation of organic matter, it also posed feedbacks and constraints at a molecular level on key metabolic pathways including oxygenic photosynthesis itself, carbon fixation and nitrogen fixation. In every case, the 'fixes' for these pathways circumvented major evolutionary changes in the core pathways and, instead, built mechanisms to repair or overproduce the 'inefficient' pathways, leaving their defects permanently imprinted in life on Earth.

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