The Princeton Guide to Ecology

III.13

The Marine Carbon Cycle

Paul Falkowski

OUTLINE

1. The two carbon cycles

2. Chemoautotrophy

3. The evolution of photoautotrophy

4. Selective pressure in the evolution of oxygenic photosynthesis

5. Primary production

6. Who are the photoautotrophs?

7. Carbon burial

8. Carbon isotope fractionation in organic matter and carbonates

9. Concluding remarks

The system of scholastic disputations encouraged in the Universities of the middle ages had unfortunately trained men to habits of indefinite argumentation, and they often preferred absurd and extravagant propositions, because greater skill was required to maintain them; the end and object of such intellectual combats being victory and not the truth.

—Charles Lyell, Principles of Geology, 1830

Approximately 50% of all the primary production on Earth occurs in the oceans, virtually all by microscopic, singlecelled organisms that drift with the currents, the phytoplankton. On ecological time scales of days to years, the vast majority of the organic matter produced by phytoplankton is consumed by grazers such that the turnover time of marine organic carbon is on the order of 1 week, compared with over a decade for terrestrial plant ecosystems. On geological time scales of millions of years, however, a small fraction of the carbon fixed by phytoplankton organisms is buried in marine sediments, thereby both giving rise to oxygen in Earth’s atmosphere and providing fossil fuel in the form of petroleum and natural gas. In this chapter, we examine the factors controlling the marine carbon cycle and its role in the ecology and biogeochemistry of Earth.

GLOSSARY

acid-base reactions. A class of (bio)chemical reactions that involve the transfer of protons without electrons.

chemoautotrophy. A mode of nutrition by which an organism can reduce inorganic carbon to organic matter in the absence of light using preformed bond energy contained in other molecules.

isotopic record of carbon. The changes in the ratio of ¹³C to ¹²C over geological time in marine carbonates or in organic matter in sediments or sedimentary rocks.

net primary production. The organic carbon that is produced by photosynthetic organisms and becomes available for other trophic levels in an ecosystem.

photoautotrophy. A mode of nutrition by which an organism can reduce inorganic carbon to organic matter using light energy.

phytoplankton. Microscopic, mostly single-celled photosynthetic organisms that drift with the currents.

redox reactions. A class of (bio)chemical reactions that involve the transfer of electrons with or without protons (i.e., hydrogen atoms). Addition of electrons or hydrogen atoms to a molecule is called “reduction”; removal of electrons or hydrogen atoms from a molecule is called “oxidation.” “Redox” is a contraction of the terms reduction and oxidation.

1. THE TWO CARBON CYCLES

All life on Earth is critically dependent on the fluxes of six elements: H, C, N, O, S, and P. Of these, the flux of C is unique. Not only is C used to make the substrates of key biological polymers, such as lipids, carbohydrates, proteins, and nucleic acids, the oxidation and reduction of C provide the major conduit of energy supply for life itself. The biological carbon cycle is based on electron transfer (i.e., redox) reactions in which the formation and utilization of the bond energy of C–H and, to a lesser extent, C–C molecules provide the major driving force of life. However, and perhaps paradoxically, the overwhelming majority of C on Earth is contained in a relatively immobile pool in the lithosphere in the form of carbonate rocks (table 1). This oxidized pool of carbon contains no biologically available energy. To sustain a flux of carbon (and hence, an essential biological building block and energy supply) on geological time scales, the lithospheric, oxidized carbon in carbonates must enter one of two mobile pools, either the atmosphere or the ocean, from which biological processes can access the carbon, reduce it to organic matter, and transfer the organic matter through metabolic processes. Hence, there are two parallel carbon cycles on Earth. One cycle is slow and abiotic, and its chemistry is based on acid–base reactions. The physical processes that drive this cycle play a key role in Earth’s climate. The second is fast and biologically driven; its chemistry is based on electron transfer reactions. The biological processes that drive this cycle play a key role in sustaining ecosystems. Let us briefly consider the two carbon cycles and then focus on the unique role of the ocean as the conduit where both cycles meet.

The “Slow” Geological Carbon Cycle

The slow geological carbon cycle operates on multimillion-year time scales and is dictated by tectonics, which is itself related to the amount of radiogenic heat produced in the Earth’s interior. In this cycle, CO₂ is released from Earth’s mantle to the atmosphere and oceans via volcanism and sea floor spreading. CO₂ is a unique gas, however. In aqueous solution, it reacts with water to form carbonic acid (H₂CO₃), which is a weak acid. In the atmosphere, this acid is formed in precipitation. When rain falls on silicate-rich rocks such as granite, the carbonic acid reacts with the silicates, and Ca²⁺ and Mg²⁺ ions are extracted and solublilized. Orthosilicic acid and the two cations are carried to the sea via rivers. In the ocean, Ca²⁺ and are generally present in excess of the concentrations in equilibrium with CaCO₃, meaning that the solution is supersaturated. In the contemporary ocean, many organisms including corals, shellfish, and several taxa of plankton catalyze the precipitation of carbonates as CaCO₃ and Ca(Mg)CO₃, which ultimately are sources of marine sediments and sedimentary and metamorphic rocks (e.g., marble). On geological time scales, most of the carbonates are subsequently subducted into the mantle, where they are heated, and their carbon is released as CO₂ to the atmosphere and ocean to carry out the cycle again. This cycle would operate whether or not there were life on the planet.

Table 1. Carbon pools in the major reservoirs on Earth

Pools	Quantity (× 10¹⁵ g)
Atmosphere	720
Oceans	38,400
Total inorganic	37,400
Surface layer	670
Deep layer	36,730
Dissolved organic	600
Lithosphere
Sedimentary carbonates	>60,000,000
Kerogens	15,000,000
Terrestrial biosphere (total)	2000
Living biomass	600–1000
Dead biomass	1200
Aquatic biosphere	1–2
Fossil fuels	4130
Coal	3510
Oil	230
Gas	140
Other (peat)	250

Absent organisms to catalyze the precipitation of carbonates, ultimately the ocean would become highly supersaturated, and carbonates would spontaneously precipitate. Indeed, such a situation almost certainly occurred over the first 3 billion years of Earth’s 4.5-billion-year history. This slow carbon cycle is a critical determinant of the concentration of CO₂ in Earth’s atmosphere and oceans on time scales of tens and hundreds of millions of years. CO₂, in turn, has an infrared absorption cross section, making it one of the most important greenhouse gases on Earth.

Once carbon, derived either from volcanism or from sea floor spreading, enters the atmosphere or oceans, it becomes mobile. In the atmosphere, virtually all of the carbon is in the form of gaseous CO₂. In the ocean, however, carbonic acid (H₂CO₃) forms a buffer system, which can be described by the following equations:

The equilibrium reactions are shifted toward the right at high pH and toward the left at low pH. Specifically, in seawater at 20°C, the pK for the first deprotonation reaction is about 6, and that for the second deprotonation is about 9. Thus, in the ocean, with an average pH of 8.2, virtually all (>95%) of the inorganic carbon is present in the form of bicarbonate. This buffer system is the major determinant of the pH of the ocean.

The ionic forms of CO₂ do not contribute to the vapor pressure of the gaseous form; thus, the concentration of the sum of all the dissolved inorganic carbon (TCO₂) can greatly exceed the atmosphere/water equilibrium concentration of gaseous CO₂ (PCO₂). The vapor pressure is predicted from Henry’s law:

where [CO₂] is the concentration of CO₂ in moles/liter, K_H is the Henry’s law constant of about 10–1.5 and is a weak function of temperature and ionic strength, and PCO₂ is the partial pressure of the gas in atmospheres. In the surface ocean, for example, total dissolved inorganic carbon (i.e., TCO₂) is approximately 2 µM, whereas [CO₂] is only about 10 μM. This [CO₂] is close to that of the atmosphere (corresponding at present to approximately 380 parts per million by volume) and resulting in approximately a 50-fold higher concentration of dissolved inorganic carbon in the ocean than of CO₂ in the atmosphere (table 1). Indeed, on time scales of thousands of years, the concentration of atmospheric CO₂ is determined by oceanic processes that control the dissolved CO₂ concentrations in surface waters.

Because of the partitioning into the three phases (equation 1), the inorganic carbon system in aquatic environments has very little chance to reach equilibrium. On the left side of equation 1, CO₂ in solution tends to equilibrate with the gas phase (i.e., the CO₂ in the overlying atmosphere under natural conditions), whereas on the right side of the equation, the tends to equilibrate with the solid phase of CaCO₃ or MgCa(CO₃)₂. Furthermore, the equilibrium constants for the various inorganic carbon reactions are temperature and salinity dependent. The partitioning of CO₂ between aqueous solution and gas phase increasingly favors the gas phase as temperature or salinity increases.

Calcification leads to a loss of one Ca²⁺ for each atom of carbon precipitated. The loss of Ca²⁺ is compensated by the formation of H⁺, which shifts the equilibrium of the inorganic carbon system, described in equation 1, to the left. Thus, calcification potentiates the formation of CO₂, leading to higher PCO₂ while simultaneously reducing the concentration of total dissolved inorganic carbon. It should be noted that although the biological formation of CaCO₃ requires metabolic energy, the energy is not stored in the chemical bonds of the product; that is, calcification is not a chemical reduction of CO₂. Rather, the energy is used to reduce the entropy in formation of the crystalline carbonate.

The “Fast” Biological Carbon Cycle

The second carbon cycle is dependent on the biologically catalyzed reduction of inorganic carbon to form organic matter, the overwhelming majority of which is oxidized back to inorganic carbon by respiratory metabolism. This cycle, which is observable on time scales of days to millennia, is driven by redox reactions that evolved over about 2 billion years, first in microbes and subsequently in multicellular organisms. A very small fraction of the reduced carbon escapes respiration and becomes incorporated into the lithosphere. In the process, some of the organic matter is transferred to the slow carbon cycle.

To form organic molecules, inorganic carbon (CO₂ and its hydrated equivalents) must be chemically reduced, a process that requires the addition of hydrogen atoms (not just protons but protons plus electrons) to the carbon atoms. Broadly speaking, the biologically catalyzed reduction reactions are carried out by two groups of organisms: chemoautotrophs and photoautotrophs, which are collectively called primary producers. The organic carbon they synthesize fuels the growth and respiratory demands of the primary producers themselves and all remaining organisms in the ecosystem.

All oxidation–reduction reactions are coupled sequences. Reduction is accomplished by the addition of an electron or hydrogen atom to an atom or molecule. In the process of donating an electron to an acceptor, the donor molecule is oxidized. Hence, oxidation–reduction reactions require pairs of substrates and can be described by a pair of partial reaction, or halfcells:

The tendency for a molecule to accept or release an electron is therefore relative to some other molecule being capable of conversely releasing or binding an electron. Chemists scale this tendency, called the redox potential, E, relative to the reaction:

which is arbitrarily assigned an E of 0 at pH 0 and is designated E₀. Biologists define the redox potential at pH 7, 298 K (i.e., room temperature), and 1 atmosphere pressure (= 101.3 kPa). When so defined, the redox potential is denoted by the symbols or sometimes E_m7. The for a standard hydrogen electrode is –420 mV.

2. CHEMOAUTOTROPHY

Organisms capable of reducing sufficient inorganic carbon to grow and reproduce in the dark without an external organic carbon source are called chemoautotrophs (literally, “chemical self-feeders”). Genetic analyses suggest that chemoautotrophy evolved very early in Earth’s history and is carried out exclusively by procaryotic organisms in both the Archea and Bacteria superkingdoms.

Early in Earth’s history, the biological reduction of inorganic carbon may have been directly coupled to the oxidation of H2. At present, however, free H2 is scarce on the planet’s surface. Rather, most of the hydrogen on the surface of Earth is combined with other atoms, such as sulfur or oxygen. Activation energy is required to break these bonds in order to extract the hydrogen. One source of energy is chemical bond energy itself. For example, the ventilation of reduced mantle gases along tectonic plate subduction zones on the sea floor provides hydrogen in the form of H2S. Several types of microbes can couple the oxidation of H2S to the reduction of inorganic carbon, thereby forming organic matter in the absence of light.

Ultimately all chemoautotrophs depend on a nonequilibrium redox gradient, without which there is no thermodynamic driver for carbon fixation. For example, the reaction involving the oxidation of H₂S by microbes in deep sea vents described above is ultimately coupled to oxygen in the ocean interior. Hence, this reaction is dependent on the chemical redox gradient between the ventilating mantle plume and the ocean interior that thermodynamically favors oxidation of the plume gases. Maintaining such a gradient requires a supply of energy, either externally, from radiation (solar or otherwise), or internally, via planetary heat and tectonics, or both.

The overall contribution of chemoautotrophy in the contemporary ocean to the formation of organic matter is relatively small, accounting for less than 1% of the total annual primary production in the sea. However, this process is critical in coupling reduction of carbon to the oxidation of low-energy substrates and is essential for completion of several biogeochemical cycles.

The oxidation state of the ocean interior is a consequence of a second energy source: light, which drives photosynthesis. Photosynthesis is an oxidation reduction reaction of the general form:

where A is, for example, an S atom. In this formulation, light is specified as a substrate, and a fraction of the light energy is stored as chemical bond energy in the organic matter. Organisms capable of reducing inorganic carbon to organic matter by using light energy to derive the source of reductant or energy are called photoautotrophs. Analyses of genes and metabolic sequences strongly suggest that the machinery for capturing and utilizing light as a source of energy to extract reductants was built on the foundation of chemoautotrophic carbon fixation; i.e., the predecessors of photoautotrophs were chemoautotrophs. The evolution of a photosynthetic process in a chemoautotroph forces consideration of both the selective forces responsible (why) and the mechanism of evolution (how).

3. THE EVOLUTION OF PHOTOAUTOTROPHY

Reductants for chemoautotrophs are generally deep in the Earth’s crust. Vent fluids are produced in magma chambers connected to the Athenosphere. As a result, the supply of vent fluids is virtually unlimited. Although the chemical disequilibria between vent fluids and bulk seawater provide a sufficient thermodynamic gradient to continuously support chemoautotrophic metabolism in the contemporary ocean, in the early Earth the oceans would not have had a sufficiently large thermodynamic energy potential to support a pandemic outbreak of chemoautotrophy. Moreover, magma chambers, volcanism, and vent fluid fluxes are tied to tectonic subduction and spreading regions, which are transient features of Earth’s crust and hence only temporary habitats for chemoautotrophs. In the Archean and early Proterozoic oceans, the chemoautotrophs would have had to have been dispersed throughout the oceans by physical mixing in order to colonize new vent regions. This same dispersion process would have helped ancestral chemoautotrophs exploit solar energy near the ocean surface.

Although the processes that selected photosynthetic reactions as the major energy transduction pathway remain obscure, central hypotheses have emerged based on our understanding of the evolution of Earth’s carbon cycle, the evolution of photosynthesis, biophysics, and molecular phylogeny. Photoautotrophs are found in all three major superkingdoms; however, there are very few known Archea capable of this form of metabolism. Efficient photosynthesis requires harvesting of solar radiation and hence the evolution of a light-harvesting system. Although some Archea and Bacteria use the pigment protein rhodopsin, by far the most efficient and ubiquitous light harvests are based on chlorins; no known Archea has a chlorin-based photosynthetic metabolic pathway. The metabolic pathway for the synthesis of porphyrins and chlorins is one of the oldest in biological evolution and is found in all chemoautotrophs. It has been proposed that the chlorin-based photosynthetic energy conversion apparatus originally arose from the need to prevent UV radiation from damaging essential macromolecules such as nucleic acids and proteins. The UV excitation energy could be transferred from the aromatic amino acid residues in the macromolecule to a blue absorption band of membrane-bound chlorins to produce a second excited state, which subsequently decays to the lower-energy excited singlet. This energy dissipation pathway can be harnessed to metabolism if the photochemically produced charge-separated primary products are prevented from undergoing a back reaction but rather form a biochemically stable intermediate reductant. This metabolic strategy was selected for the photosynthetic reduction of CO₂ to carbohydrates, using reductants such as S^2- or Fe²⁺, which have redox potentials that are too positive to reduce CO₂ directly.

The synthesis of reduced (i.e., organic) carbon and the oxidized form of the electron donor permits a photoautotroph to use “respiratory” metabolic processes but operate them in reverse. However, not all of the reduced carbon and oxidants remain accessible to the photoautotrophs. In the oceans, cells tend to sink, carrying with them organic carbon. The oxidation of Fe²⁺ forms insoluble Fe³⁺ salts that precipitate. The sedimentation and subsequent burial of organic carbon and Fe³⁺ remove these components from the water column. Without replenishment, the essential reductants for anoxygenic photosynthesis would eventually become depleted in the surface waters. Thus, the necessity to regenerate reductants potentially prevented anoxygenic photoautotrophs from providing the major source of fixed carbon on Earth for eternity. Major net accumulation of reduced organic carbon in Proterozoic sediments implies local depletion of reductants such as S^2- and Fe²⁺ from the upper ocean. These limitations almost certainly provided the evolutionary selection pressure for an alternative electron donor.

4. SELECTIVE PRESSURE IN THE EVOLUTION OF OXYGENIC PHOTOSYNTHESIS

Water (H₂O) is a potentially useful biological reductant with an effectively unlimited supply on Earth. Water contains about 100 kmol/m³ of H atoms, and, given >10¹⁸ m³ of water in the hydrosphere and cryosphere, more than 10²⁰ kmol of reductant is potentially accessible. Use of H₂O as a reductant for CO₂, however, requires a larger energy input than does the use of Fe²⁺ or S^2-. Indeed, to split water by light requires 0.82 electron volts at pH 7 and 298 K. Utilizing light at such high energy levels required the evolution of a new photosynthetic pigment, chlorophyll a, which has a red (lowest singlet) absorption band that is 200 to 300 nm blue shifted relative to bacteriochlorophylls. Moreover, stabilization of the primary electron acceptor to prevent a back reaction necessitated thermodynamic inefficiency that ultimately led to the evolution of two light-driven reactions operating in series. This sequential action of two photochemical reactions is unique to oxygenic photoautotrophs and presumably involved horizontal gene transfer through one or more symbiotic events. As discussed below, oxygenic photosynthesis appears to have arisen only once in a single clade of Bacteria (the cyanobacteria).

In all oxygenic photoautotrophs, equation 3 can be modified to:

where Chl a is the pigment chlorophyll a exclusively utilized in the reaction. Equation 6 implies that somehow chlorophyll a catalyzes a reaction or a series of reactions whereby light energy is used to oxidize water:

yielding gaseous molecular oxygen. Hidden within equation 7 are a complex suite of biological innovations that have not been yet successfully mimicked by humans. At the core of the water-splitting complex is a quartet of Mn atoms that sequentially extract electrons, one at a time, from two H₂O molecules, releasing gaseous O₂ to the environment and storing the reductants on biochemical intermediates.

The photochemically produced reductants generated by the reactions schematically outlined in equation 7 are subsequently used in the fixation of CO₂ by a suite of enzymes that can operate in vitro in darkness, and, hence, the ensemble of these reactions is called the dark reactions. At pH 7 and 25°C, the formation of glucose from CO₂ requires an investment of 915 calories per mole. If water is the source of reductant, the overall efficiency for photosynthetic reduction of CO₂ to glucose is approximately 30%; i.e., 30% of the absorbed solar radiation is stored in the chemical bonds of glucose molecules.

5. PRIMARY PRODUCTION

When one subtracts the costs of all other metabolic processes by the chemoautotrophs and photoautotrophs, the organic carbon that remains is available for the growth and metabolic costs of heterotrophs. This remaining carbon is called net primary production (NPP). NPP provides an upper bound for all other metabolic demands in an ecosystem. If NPP is greater than all respiratory consumption of the ecosystem, the ecosystem is said to be net autotrophic. Conversely, if NPP is less than all respiratory consumption, the system must either import organic matter from outside its bounds, or it will slowly run down—it is net heterotrophic.

It should be noted that NPP and photosynthesis are not synonymous. On a planetary scale, the former includes chemoautrophy; the latter does not. Moreover, photosynthesis per se does not include the integrated respiratory term for the photoautotrophs themselves. In reality, that term is extremely difficult to measure directly; hence, NPP is generally approximated from measurements of photosynthetic rates integrated over some appropriate length of time (a day, a month, a season, or a year), and respiratory costs are either assumed or neglected. From satellite data used to estimate upper-ocean chlorophyll concentrations, satellite-based observations of incident solar radiation, atlases of seasonally averaged sea-surface temperature, and models that incorporate a temperature response function for photosynthesis, it is possible to estimate global net photosynthesis in the world oceans. Although estimates vary among models based on how the parameters are derived, for illustrative purposes we use a model based on empirical parameterization of the daily integrated photosynthesis profiles as a function of depth. The physical depth at which 1% of irradiance incident on the sea surface remains is called the euphotic zone. This depth can be calculated from surface chlorophyll concentrations and defines the base of the water column at which net photosynthesis can be supported. Given such information, net primary production can be calculated following the general equation:

where PP_eu is daily net primary production integrated over the euphotic zone, C_sat is the satellite-based (upper water column, derived from table 2) chlorophyll concentration, is the maximum daily photosynthetic rate within the water column, Z_eu is the depth of the euphotic zone, DL is the photoperiod, and F is a function describing the shape of the photosynthesis depth profile. This general model can be both expanded (differentiated) and collapsed (integrated) with respect to time and irradiance; however, the global results are fundamentally similar. The models predict that NPP in the world’s oceans amounts to 40–50 Pg per annum.

Table 2. Comparison of marine and terrestrial net primary productivity across biomes

6. WHO ARE THE PHOTOAUTOTROPHS?

In the oceans, oxygenic photoautotrophs are a taxonomically diverse group of mostly single-celled photosynthetic organisms that drift with currents. In the contemporary ocean, these organisms, called phytoplankton (derived from the Greek, meaning to wander), are comprised of approximately 20,000 morphologically defined species distributed among at least eight taxonomic divisions or phyla. By comparison, higher plants are comprised of more than 250,000 species, almost all of which are contained within one class in one division. Thus, unlike terrestrial plants, phytoplankton appear to be represented by relatively few morphological species, but they are phylogenetically extremely diverse. This deep taxonomic diversity is reflected in their evolutionary history and ecological function.

Within this diverse group of organisms, three basic evolutionary lineages are discernible. The first contains all prokaryotic oxygenic phytoplankton, which belong to one class of bacteria, namely the cyanobacteria. Cyanobacteria are the only known oxygenic photoautotrophs that existed before about 2.4 billion years ago. These prokaryotes numerically dominate the photoautotrophic community in contemporary marine ecosystems, and their continued success bespeaks an extraordinary adaptive capacity. At any moment in time, there are approximately 10²⁴ cyanobacterial cells in the contemporary oceans. To put that in perspective, the number of cyanobacterial cells in the oceans is two orders of magnitude more than all the stars in the sky. The other two groups are eukaryotic. One, broadly speaking, contains chlorophylls a and b and is called the “green” line. These organisms, which are the progenitors of terrestrial plants, are not as abundant as a third group, which contains chlorophylls a and c and is often called the “red” line. The red line includes diatoms, coccolithophorids, and most dinoflagellates. All three groups are extremely important players in NPP and carbon burial in the contemporary ocean.

7. CARBON BURIAL

On geological time scales, there is one important fate for NPP, namely burial in the sediments. By far the largest reservoir of organic matter on Earth is locked up in rocks. Virtually all of this organic carbon is the result of the burial of exported marine organic matter in coastal sediments over literally billions of years of Earth’s history. On geological time scales, the burial of marine NPP effectively removes carbon from biological cycles and places most (not all) of that carbon into the slow carbon cycle. A small fraction of the organic matter escapes tectonic processing via the Wilson cycle and is permanently buried, mostly in continental rocks. The burial of organic carbon effectively removes reducing equivalents from the atmosphere and ocean and thereby allows oxygen to accumulate in Earth’s atmosphere.

Carbon burial is not inferred from direct measurement but rather from indirect means. One of the most common proxies used to derive burial on geological time scales is based on isotopic fractionation of carbonates. The rationale for this analysis is that the primary enzyme responsible for inorganic carbon fixation is ribulose 1,5-bisphosphate carboxylase/oxygenase (RuBisCO), which catalyzes the reaction between ribulose 1,5-bisphosphate and CO₂ (not ), to form two molecules of 3-phosphoglycerate. The enzyme strongly discriminates against ¹³C, such that the resulting isotopic fractionation amounts to approximately 27 parts per thousand relative to the source carbon isotopic value. The extent of the actual fractionation is somewhat variable and is a function of carbon availability and of the transport processes for inorganic carbon into the cells as well as the specific carboxylation pathway. However, regardless of the quantitative aspects, the net effect of carbon fixation is an enrichment of the inorganic carbon pool in ¹³C, whereas the organic carbon produced is enriched in ¹²C.

8. CARBON ISOTOPE FRACTIONATION IN ORGANIC MATTER AND CARBONATES

The isotopic fractionation in carbonates mirrors the relative amount of organic carbon buried. It is generally assumed that the source carbon, from volcanism (so-called mantle carbon) has an isotopic value of about –5 parts per thousand. Because mass balance must constrain the isotopic signatures of carbonate carbon and organic carbon with the mantle carbon, then, in the steady state, the fraction of buried organic matter of the total carbon buried (f_org) can be calculated from the relationship,

where f_org is the fraction of organic carbon buried, δ_w is the average isotopic content of the carbon weathered, δ_car is the isotopic signature of the carbonate carbon, and Δ_B is the isotopic difference between organic carbon and carbonate carbon deposited in the ocean. Equation 9 is a steady-state model that presumes the source of carbon from the mantle is constant over geological time. This basic model is the basis of nearly all estimates of organic carbon burial rates.

Figure 1. The change in the CO₂ concentrations over the past 550 million years.

Carbonate isotopic analyses reveal positive excursions (i.e., implying organic carbon burial) in the Proterozoic and more modest excursions throughout the Phanaerozoic (figure 1). Burial of organic carbon on geological time scales implies that export production must deviate from the steady state on ecological time scales. Such a deviation requires changing one or more of (1) ocean nutrient inventories, (2) the utilization of unused nutrients in enriched areas, (3) the average elemental composition of the organic material, or (4) the “rain” ratios of particulate organic carbon to particulate inorganic carbon to the sea floor.

Over the past 200 million years, the carbon isotopic record indicates that a significant amount of organic carbon has been buried in the lithosphere. The burial of organic carbon denotes the burial of reductants. As a consequence, oxidized molecules must have accumulated in some other domain. The oxidized molecule is O₂. Hence, to a geochemist, the burial of organic matter formed by oxygenic photosynthetic organisms requires the oxidation of the atmosphere. Quantitative analysis of isotopic record of carbonates suggests that oxygen rose from about 11% 200 million years ago to the contemporary value of 21% as a result of the burial of the organic matter, largely in marine sediments. The removal of a small fraction of the buried carbon by humans to fuel the current industrialization of the world represents a reversal of this process, namely the consumption of oxygen and the reoxidation of organic matter by machines rather than by biological respiration.

9. CONCLUDING REMARKS

The evolution of primary producers in the oceans profoundly changed the chemistry of the atmosphere, ocean, and lithosphere of Earth. Primarily through the utilization of solar radiation, the biological (fast) carbon cycle allows for a disequilibrium in geochemical processes, such that Earth maintains an oxidized atmosphere and ocean. This disequilibrium prevents atmospheric oxygen from being depleted, maintains a reduced atmospheric CO₂ concentration, and simultaneously imprints the ocean interior and the lithosphere elemental compositions that reflects that of the bulk biological material from which it is derived. Although primary producers in the ocean comprise only about 1% of Earth’s biomass, their metabolic rate and biogeochemical impact rival those of the much larger terrestrial ecosystem. On geological time scales, these organisms are the little engines that are essential to maintaining life as we know it on this planet.

Over the past 200 years, the fossil remains of marine photosynthetic organisms have been extracted from the lithosphere by humans at a rate approximately 1 million times faster than they accumulated. The extraction and subsequent combustion of fossil fuels have temporarily inverted the carbon cycle; the oceans are not in equilibrium with the atmosphere, and the excess CO₂ potentially will alter Earth’s climate rapidly and dramatically. The rise of CO₂ is not debatable; it is a scientific fact. Unfortunately, however, the fundamental scientific facts pertaining to the carbon cycle on Earth are still debated, obscuring a sustainable path forward. Were he alive today, Charles Lyell might think we had not yet left the philosophy encouraged by universities in the Middle Ages.

The Marine Carbon Cycle

OUTLINE

GLOSSARY

The “Slow” Geological Carbon Cycle

The “Fast” Biological Carbon Cycle

FURTHER READING