Chapter 5. Biogeochemical cycle of carbon

Table of Contents

5.1. The natural carbon cycle
5.1.1. Terrestrial processes
5.1.2. Oceanic processes
5.1.3 Geological processes
5.2 Methane
5.2.1 Sources and sinks of atmospheric methane
5.3 Human disruption of carbon cycle
5.4 Carbon cycle research in Hungary

Carbon, a substantial element for the biosphere, is often referred to as the "building block of life" because living organisms are all based on carbon. Carbon compounds can exist in solid (e.g. diamonds or coal), liquid (e.g. crude oil), or gas (e.g. carbon dioxide) forms. There are three naturally occurring isotopes, with 12C and 13C being stable, while 14C is radioactive, decaying with a half-life of about 5,730 years.

Carbon is stored in and exchanged among four most relevant pools (or reservoirs). In the carbon cycle carbon moves between the hydrosphere, the atmosphere, the geosphere and the biosphere (terrestrial ecosystems). In the non-living pools, carbon is stored as carbonate (CaCO3, Figure 5.1) rocks, dead organic matter, such as humus in soil, fossil fuels from dead organic matter, carbon dioxide (CO2, Figure 5.2), carbon monoxide (CO, Figure 5.3) carbon dioxide dissolved in water to form HCO3.

Chemical structure of carbonate ion.

Figure 5.1: Chemical structure of carbonate ion.

Chemical structure of carbon dioxide.

Figure 5.2: Chemical structure of carbon dioxide.

Large continuous carbon fluxes are present among these pools. In atmospheric chemistry the fluxes are considered from the point of view of the atmosphere, usually positive fluxes mean emission to the atmosphere (i.e. the pool where the carbon originates from acts as source), negative fluxes mean carbon uptake (i.e. the pool acts as a sink). The carbon pools can act both as sources or sinks of carbon, i.e. releasing or absorbing carbon to/from the atmosphere. Usually pools are sources and sinks simultaneously, i.e. carbon is emitted from the pool to the atmosphere in a certain process, and taken up via another process. The system is considered to be in the state of dynamic equilibrium, if the positive fluxes match negative fluxes on longterm average so that the pool size remain constant.

Chemical structure of carbon monoxide.

Figure 5.3: Chemical structure of carbon monoxide.

5.1. The natural carbon cycle

Both carbon dioxide (CO2) and methane (CH4, Figure 5.4) play an important role in carbon cycle representing enormous ocean-atmosphere and surface-atmosphere carbon fluxes which had been constant around 280 ppm in the preindustrial era, up until 1750.

Chemical structure of methane.

Figure 5.4: Chemical structure of methane.

Carbon dioxide is a long-lived GHG (LLGHG) in the atmosphere. Prehistoric concentrations of CO2 reconstructed from ice cores showed that previous elevation in CO2 concentrations during interglacial periods happened gradually by only a few Gt C per decade. The overall variation in CO2 concentration between glacial interglacial periods barely exceeded 100 ppm. Current high concentrations of CO2 have not been reached in the last 15 million years.

The carbon budget can be described as the balance or imbalance between sources and sinks of carbon. The most important natural sources of carbon are the ocean, biosphere respiration, geological sources. The gross terrestrial carbon flux transfers ca. 120 Gt C yr−1, the ocean-atmosphere flux is around 90 Gt C yr−1 under natural circumstances (black arrows on Figure 5.5). The balance between sources and sinks, i.e. the carbon budget (longterm net flux averaged for a decade or longer time frame) is less than 0.1 Gt C yr−1 under undisturbed natural conditions.

The global carbon cycle.

Figure 5.5: The global carbon cycle for the 1990s, showing the main annual fluxes in GtC yr–1: pre-industrial ‘natural’ fluxes in black and ‘anthropogenic’ fluxes in red (modified from Sarmiento and Gruber, 2006, with changes in pool sizes from Sabine et al., 2004a). The net terrestrial loss of –39 GtC is inferred from cumulative fossil fuel emissions minus atmospheric increase minus ocean storage. The loss of –140 GtC from the ‘vegetation, soil and detritus’ compartment represents the cumulative emissions from land use change (Houghton, 2003), and requires a terrestrial biosphere sink of 101 GtC (in Sabine et al., given only as ranges of –140 to –80 GtC and 61 to 141 GtC, respectively; other uncertainties given in their Table 1). Net anthropogenic exchanges with the atmosphere are from Column 5 ‘AR4’ in Table 7.1. Gross fluxes generally have uncertainties of more than ±20% but fractional amounts have been retained to achieve overall balance when including estimates in fractions of GtC yr–1 for riverine transport, weathering, deep ocean burial, etc. ‘GPP’ is annual gross (terrestrial) primary production. Atmospheric carbon content and all cumulative fluxes since 1750 are as of end 1994. Source: IPCC AR4 Fig. 7.3.

5.1.1. Terrestrial processes

Carbon is taken up by the biosphere through autotrophs, which are organisms capable of synthesizing their own nutrients from inorganic substances. There are two types of autotrophs between which the main difference lies in the source of energy they use for the synthesis. Photoautotrophs (most autotrophs, such as green plants, certain algae, and photosynthetic bacteria) use light for energy. Chemoautotrophs (e.g chemosynthetic bacteria) use energy from chemical reactions, e.g. from oxidization of electron donors.)

Most of the terrestrial CO2 flux takes place through terrestrial vegetation that absorbs CO2 via the photosynthesis, although most of that amount is respired back to the atmosphere (Figure 5.5). The most important land carbon pools are the terrestrial vegetation, soil and detritus. The carbon fluxes between the atmosphere and these pools take place on relatively small time scales, therefore this part of the carbon cycle is sometimes referred to as the fast carbon cycle. On the shortest time scales of seconds to minutes, plants take carbon out of the atmosphere through photosynthesis and release it back into the atmosphere via respiration.

The amount of CO2 that is converted to carbohydrates in the photosynthesis is known as gross primary production (GPP). Part of this carbon is assimilated to support plant growth and functioning, the other part is respired. The annual difference between GPP and autotrophic respiration (Ra), i.e. the annual plant growth, is called net primary production (NPP, equation (5.1)). Terrestrial photosynthesis is estimated to be around 120 Gt C yr−1 (see Figure 5.5)



On longer time scales, most dead biomass moves to the soil organic matter and detritus pools, where it can be stored for years, decades or centuries. Several soil organic matter pools with different residence times can be defined. Eventually this carbon is respired back to the atmosphere by soil microbes and fungi, decomposed at a rate depending on their chemical composition and environmental circumstances (moisture, temperature etc.). When oxygen is present aerobic respiration occurs, which releases carbon dioxide. In the oxygen limited environment anaerobic respiration occurs, producing methane instead of CO2.

Carbon entering the terrestrial biosphere via photosynthesis is emitted back via (i) autotrophic respiration, (ii) heterotrophic respiration (decomposers and herbivores) (iii) fires (combustion). The net ecosystem production shows the net amount of carbon removed from (or released to) the atmosphere i.e. gained (or lost) by the ecosystem when no other disturbances (carbon loss of the ecosytem) are considered.



Taking into account other carbon losses, e.g. transport of biomass from agricultural lands, fires etc. shows us the carbon sequestered by the terrestrial biosphere, the net biome production (NBP). In a system being in equilibrium, NBP should be zero.

5.1.2. Oceanic processes

The ocean is another major sink of carbon, under natural circumstances it removes CO2 equivalent to ca. 70 Gt C yr-1 from the atmosphere (Figure 5.5). Before the industrial revolution, the ocean contained about 60 times as much carbon as the atmosphere and 20 times as much carbon as the terrestrial biosphere/soil compartment.

There are different processes involved in the removal of CO2 transporting carbon to different depths associated with different residence times. Because of the speed of the participating processes, this part is sometimes referred as the slow carbon cycle. Oceanic carbon exists in several forms: as dissolved inorganic carbon (DIC), dissolved organic carbon (DOC), and particulate organic carbon (POC) (living and dead). Basically three main processes govern carbon absorption: the solubility pump (CO2 exchange driven by solubility of atmospheric CO2), the organic carbon pump (driven by photosynthetic uptake by marine biota and sinking of this carbon as organic particles to deeper layers) and the CaCO3 counter pump (driven by release of CO2 during the formation of CaCO3 shells). The organic carbon pump and carbonate pump processes are the so-called biological pumps, while the solubility pump is sometimes referred as the physical pump (Figure 5.6).

Physical pump and biological pump.

Figure 5.6: Physical pump and biological pump. Three main ocean carbon pumps govern the regulation of natural atmospheric CO2 changes by the ocean (Heinze et al., 1991): the solubility pump, the organic carbon pump and the CaCO3 ‘counter pump’. The oceanic uptake of anthropogenic CO2 is dominated by inorganic carbon uptake at the ocean surface and physical transport of anthropogenic carbon from the surface to deeper layers. For a constant ocean circulation, to first order, the biological carbon pumps remain unaffected because nutrient cycling does not change. If the ocean circulation slows down, anthropogenic carbon uptake is dominated by inorganic buffering and physical transport as before, but the marine particle flux can reach greater depths if its sinking speed does not change, leading to a biologically induced negative feedback that is expected to be smaller than the positive feedback associated with a slower physical downward mixing of anthropogenic carbon. Source IPCC AR4 Fig. 7.10

The solubility pump

The marine carbonate buffer system allows the ocean to take up CO2 far in excess of its potential uptake capacity based on solubility alone, and in doing so controls the pH of the ocean. This control is achieved by a series of reactions that transform carbon added as CO2 into bicarbonate (HCO3) and carbonate (CO32–). These three dissolved forms are collectively known as DIC. CO2 is a weakly acidic gas and the minerals dissolved in the ocean have over geologic time created a slightly alkaline ocean (surface pH 7.9 to 8.25). When it dissolves, it reacts with water to form carbonic acid, which dissociates into a hydrogen ion (H+) and a HCO3– ion, with some of the H+ then reacting with CO32– to form a second HCO3 ion

CO2 + H2O → H+ + HCO3→ 2H+ + CO32–


CO2 + H2O + CO32– → HCO3+ H+ + CO32– → 2HCO3.


The air-sea exchange of CO2 is determined largely by the air-sea gradient in pCO2 between atmosphere and ocean. Equilibration of surface ocean and atmosphere occurs on a time scale of roughly one year. Gas exchange rates increase with wind speed and depend on other factors such as precipitation, heat flux, sea ice and surfactants. The magnitudes and uncertainties in local gas exchange rates are maximal at high wind speeds. In contrast, the equilibrium values for partitioning of CO2 between air and seawater and associated seawater pH values are well established.

Cold, CO2 rich waters near the poles during local winters (the lower temperature the higher the solubility) and more dense, therefore sink to greater depths with the Meridional Overurning Circulation creating the solubility pump.

Organic carbon pump

The organic carbon pump (Figure 5.6) is the process in which CO2 is incorporated to organic matter in the photosynthesis. Part of the particulate organic carbon sinks to deeper layers where it gets decomposed and respired by bacteria and transported back to surface layers by marine circulation as DIC. Other part of POC sinks deeper and settles on the bottom of the ocean to be stored there with thousand years of residence time. The efficient organic pump performes a net carbon removal from the atmosphere also decreasing the total carbon content in the surface layers by developing organic matter in the photosynthesis

CO2 + H2O → CH2O + O2.


The efficiency of the organic pump is limited via the photosynthesis by the availability of light, nutrients and oxygen. Figure 5.7 shows an example of ocean primary productivity in a period of year 2001 based on MODIS imagery.

The ocean primary productivity.

Figure 5.7: The global carbon cycle is greatly influenced by ocean primary productivity, the rate of carbon dioxide uptake via marine plant photosynthesis minus the rate that carbon dioxide is put back into the ocean's carbon reservoir through respiration. This MODIS composite from May 9 to June 9, 2001 shows how variable the rates of carbon exchange are across the Earth’s oceans. Productivity tends to be high at northern and southern latitudes, where mixing from deep ocean waters brings up nutrients, and at the margins of continents, where currents draw up nutrients in the shallower waters of the continental shelves. Black areas indicate regions where productivity could not be calculated, typically because of clouds or sea ice. Image credit: MODIS Ocean Team/Ocean Primary Productivity, Wayne Esaias, Principal Investigator, NASA-GSFC; Kevin Turpie, SAIC/GSC.

Carbonate pump

The other component of the biological pump is based on the formation of carbonate and silicate shells, the ions are used by diatoms (silica-secreting organisms, e.g. Figure 5.8) and coccolithophorids (carbonate-secreting organisms, e.g. Figure 5.9) The carbonate ions produced in chemical weathering of carbonate rocks and the solution of CO2 in surface seewater play important role in the biological pump as well. The production of solid CaCO3 (that is, “carbonate precipitation”) occurs in the surface waters of the ocean, both organically - by organisms that build their shells from CaCO3 - and inorganically according to the chemical equilibrium in the oceans according to the following chemical equation:

Ca2+ + 2HCO3 → CaCO3 + CO2 + H2O


In the reaction above to form one CaCO3, one CO2 is produced, which means that altough the DIC concentration is decreased in this process, CO2 partial pressure will increase affecting solubility of atmospheric CO2 in the surface ocean. In contrary, the formation of silica shells does not produce CO2, hence being a more effective process regarding carbon sequestration.


Figure 5.8: Diatom


Figure 5.9: Coccolitophorid. Gephyrocapsa oceanica Kamptner from Mie Prefecture, Japan. SEM:JEOL JSM-6330F. Scale bar = 1.0 μm. (Source: wikipedia)

5.1.3 Geological processes

On still longer time scales, organic matter that became buried in deep sediments (and protected from decay) is slowly transformed into deposits of coal, oil and natural gas, the fossil fuels we use today. When we burn these substances, carbon that has been stored for millions of years is released once again to the atmosphere in the form of CO2.

Silicate weathering and atmospheric CO2

A small amount of carbon is absorbed in the process of weathering and carried to the ocean by surface water and stored there for longer term. Chemical weathering of rocks are represented by reactions (R5.5) and (R5.6) for carbonates, and reactions (R5.7) and (R5.8) for silicates

CO2 +H2O + CaCO3 → Ca2+ + 2HCO3


2CO2 + 2H2O +CaMg(CO3)2 → Ca2+ + Mg2+ + 4HCO3


2CO2 +3H2O + CaSiO3 → Ca2+ +2HCO3 +H4SiO4


2CO2 + 3H2O +MgSiO3 → Mg2+ + 2HCO3 + H4SiO4


In longterm control of atmospheric CO2 weathering plays an important role (Fig 5.5). CO2 dissolved in surface waters participates in weathering of rocks (R5.5 – R5.8). The ions produced in the weathering process are carried away to the ocean by rivers where they are used by marine biota to form carbonate and silicate tissues. After sinking to the ocean sediment the weathering products are built in carbonate rocks in the see crust where the carbon can be stored for several thousand years.