The carbon cycle

Table 3.2 showed that the ocean is a net sink for carbon dioxide, while it is a net source for most of the other greenhouse gases. In this section we shall explore the reason for this behaviour, but first we will examine the role of this oceanic sink in the full carbon cycle.

3.3.1 The carbon cycle

Carbon cycles between the Earth, biosphere, ocean and atmosphere, as shown schematically in Fig. 3.4. Carbon dioxide is mixed through the troposphere within a few months. It takes much longer, probably hundreds of years, for carbon dioxide to equilibrate in the ocean. The terrestrial store also acts on several scales, an annual recycling, but a longer growth and decay cycle for longer lived organisms such as trees.

The variation in atmospheric CO2 is shown in Fig. 3.5, a monthly time series taken at Mauna Loa in Hawaii. A rapid seasonal cycle is superimposed on a longer term more gradual increase. The latter trend is due to anthropogenic activity. The annual modulation is due to the biosphere. During the summer large amounts of carbon are fixed in the terrestrial and oceanic biosphere by photosynthesis (see §4.1). The limited photosynthesis of winter leads to net respiration and release of CO2 to the atmosphere. Over the globe the seasonal exchange of CO2 between the oceans and atmosphere is roughly in balance between the two hemispheres. However, the greater land mass of the Northern Hemisphere, as compared to the Southern, leads to a greater terrestrial uptake of carbon during the northern summer than is released over land in the winter of the Southern Hemisphere. An annual cycle in global atmospheric CO2 concentration results, with a peak during the northern winter's respiration phase.

The carbon cycle, and the changes in distribution of carbon between the different reservoirs, form an important component of the climate system. Much carbon appears to be sequestered into very long-term stores, in the interior of the Earth, sediments and deep oceans. Chapter 6 will show, however, that large changes in atmospheric, and therefore climatically active, carbon dioxide can occur naturally. A near-doubling of CO2 occurred from 20 000 years BP, during the height of the last glacial period, to ad 1750 (i.e. 200 BP), before man's influence became noticeable. Substantial natural redistribution of CO2 between the atmosphere, ocean and biosphere is therefore possible. Chapter 4 will consider some of the biosphere's contribution to this exchange; Chapters 6 and 7 will discuss natural and anthropogenic changes to the carbon cycle in more detail. A very important component of this exchange, however, is the storage mechanism in the ocean. This is largely of chemical origin and discussed in detail in the next section.

Fig. 3.4. Global carbon reservoirs and annual fluxes. Numbers underlined represent net annual CO2 accumulation due to human action. Units are gigatons (109 metric tons, or 1012 kg) of carbon in the reservoirs and GtCyr-1 for fluxes. [Adapted from Fig. 1.1 of Houghton et al. (1990), with updated figures from Fig. 3.1 of Houghton et al. (2001).]

Fig. 3.4. Global carbon reservoirs and annual fluxes. Numbers underlined represent net annual CO2 accumulation due to human action. Units are gigatons (109 metric tons, or 1012 kg) of carbon in the reservoirs and GtCyr-1 for fluxes. [Adapted from Fig. 1.1 of Houghton et al. (1990), with updated figures from Fig. 3.1 of Houghton et al. (2001).]

Fig. 3.5. Monthly average atmospheric CO2 concentrations, observed continuously at Mauna Loa, Hawaii.

1955 1965 1975 1985 1995 2005

Year

1955 1965 1975 1985 1995 2005

Year

3.3.2 Oceanic control of carbon dioxide - principal processes

Carbon dioxide is a major product of the combustion of carbon-containing substances, such as wood, coal, natural gas and oil. Since the dramatic increase in combustion initiated by the Industrial Revolution in the late eighteenth century the atmospheric concentration of CO2 has increased by about a third (see Fig. 7.6). Substantially less carbon dioxide has remained in the atmosphere than has been emitted into it since 1750, perhaps as little as 50%. A significant part of this extra carbon dioxide has been absorbed by the ocean, as a result of the

Table 3.3. Dissociation constants for the principal reactions in (3.3), at a pressure of 1 atmosphere and a salinity of 35 psu K1 is multiplied by the constant [H2O]

Source: After Bracker and Peng, 1982.

gas's high solubility. The principal reason for the ocean being a sink of carbon dioxide is shown in (3.3):

CO2(gas) + H2O (^ H2COb) ^ H+ + HCO3- ^ 2H+ + CO32- (3.3)

The component reactions of (3.3) are fast, and the system exists in equilibrium. Thus, additional CO2 will be quickly swept into this system, distributing carbon atoms between the gas phase and the carbonate (CO32-) and bicarbonate (HCO3-) ions in solution. The reason for the rapidity of reaction (3.3) can be seen by examining the reaction rates, or dissociation constants, K1 and K2 of two equations of the full equilibrium in (3.3), namely

for the gas to bicarbonate equilibrium in the left half of (3.3), and

for the bicarbonate to carbonate equilibration to the right of (3.3). In (3.4) and (3.5) the square brackets represent the concentration of the respective chemical species. The dissociation constants depend on temperature, pressure and salinity. Table 3.3 shows values for K1 and K2 at different temperatures, for a pressure of 1 atmosphere and a salinity of 35 psu. The greater magnitude of K1 means that most CO2 is converted to HCO3- and only a very small portion is then converted to the carbonate ion as K2 is roughly a thousand times smaller than K1. The net effect of the reactions in (3.3) therefore leads to the summary reaction

The bicarbonate/carbonate species are not produced solely from the equilibrium with carbon dioxide, but also have a source from deposition, by rivers or wind-blown dust, of the weathering products of rocks containing calcium carbonate (CaCO3). Limestone is one common source of such material. This

'pre-existing' oceanic carbonate source from calcium carbonate weathering permits a greater absorption of carbon dioxide than would otherwise occur. This can be seen from considering the dissociation constant summarizing the entire process in (3.6):

For a fixed temperature and pressure K is a constant, so weathering-enhanced concentrations of carbonate, and hydrogen ions, through the two right-hand equilibria of (3.1), must draw more carbon dioxide into solution, as [H2O] is too large to be affected by the reactions of (3.3).

In surface waters, biological processes involving the interaction of phosphate, nitrate and carbon lead to acidity and total dissolved carbon amounts being approximately constant over the globe (these processes will be pursued in Chapter 4). The numerator of (3.7) will therefore be constant, as is [H2O]. However, the reaction rate, K, will depend on the temperature, becoming slower with warmer temperatures. If the atmospheric carbon dioxide is in equilibrium with the surface water, then

where SCO2 is the solubility of carbon dioxide and pCO2 is the partial pressure of carbon dioxide in the overlying atmosphere. This means that the partial pressure can be expressed by

Equation (3.9) suggests that, in equilibrium conditions, the atmospheric partial pressure of carbon dioxide should be a strong function of temperature. The relative changes of the solubility and K with temperature imply that the partial pressure should decrease by about a factor of three between 24°C and 0°C. We have, however, already seen that the atmospheric carbon dioxide concentration is essentially uniform over the globe. The atmosphere mixes fast enough for this thermally-driven poleward gradient not to be established.

There cannot, therefore, be a general equilibrium of the carbon dioxide exchange between the ocean and atmosphere. Fig. 3.2 shows that the CO2 levels in the polar surface waters are lower, while the tropical water CO2 partial pressure is higher, than the equilibrium value for the overlying atmosphere. There must be a net flux of atmospheric carbon dioxide into the polar waters, and out of the tropical oceans. This flux does not push the ocean towards equilibrium with the atmosphere, because the temperature of the water is changing due to the march of the seasons and changes in the surface ocean currents, and also because there are exchanges of the surface water with deeper waters in both regions, as discussed in Chapters 1, 2 and 5. There is a competition between the equilibration time for carbon dioxide and the timescale for surface water stagnation which ensures that equilibrium conditions are not always achieved.

The reaction sequence in (3.3) is the principal way in which carbon dioxide intake into the ocean is chemically enhanced. However, there are other reactions which might be significant. These are reactions of CO2 with marine minerals, either in bottom sediments or in solution within the water column. The breakdown of calcium carbonate fragments, which can include the shells of crustaceans, by the reaction

is one such process. This is the same reaction as is involved in the weathering of limestone on land. It is not thought to be widely important as a carbon dioxide sink because surface waters are generally super-saturated with respect to solid phase CaCO3. This inhibits (3.10) from occurring. However, in bottom waters in coastal zones, where anthropogenic carbon dioxide is easily able to penetrate, the water appears to be under-saturated with respect to impure forms of calcium carbonate. It is estimated that perhaps 1.5% of the carbon dioxide produced by man each year could be absorbed by the oceans in coastal regions by reaction (3.10).

We saw in §1.5 that the oceans are net absorbers of the additional carbon dioxide added to the atmosphere each year by human activity. Fig. 3.4 shows that much of the CO2 that enters the ocean from the atmosphere does not remain in the surface waters but is transferred by mixing, and formation of deep water, into the intermediate and deep ocean. Much of the additional anthropogenic carbon dioxide will also follow this route, forcing the surface ocean to move towards equilibration with the increasing carbon dioxide of the atmosphere at an even slower rate. The long timescales of deep water overturning - several centuries to a thousand years - coupled with the greater solubility of carbon dioxide in water under pressure, means that the ocean acts as a giant reservoir of carbon dioxide, and a drag on climatic change associated with CO2 increase. The converse of this also exists. If atmospheric CO2 levels were suddenly decreased, then the ocean would slowly leak carbon dioxide back to the atmosphere, acting to push the climate back towards its previous state.

The ocean also exerts another, biological, control on atmospheric carbon dioxide. As we shall see in Chapter 4, input of CO2 to the ocean encourages biological activity. Some of this carbon is then stored within marine life-forms, where its return to the atmosphere is delayed. If the marine organisms add to the sediment on the ocean floor upon their death, then the carbon is added to the geosphere's reservoir. This may not be re-cycled to the atmosphere for millions of years.

3.3.3 Oceanic control of carbon dioxide - geographical variations

The poleward decrease of upper ocean pCO2 is not without caveats (Fig. 3.2). In Fig. 3.6 latitudinal sections of ocean surface pCO2 obtained in the early 1970s for the Atlantic andPacific Oceans show both similarities and differences. Both oceans show distinct peaks near the equator, with the Atlantic peak being rather broader. Both also show the ocean and atmosphere in near-equilibrium into the mid-latitudes. However, their behaviour at high latitudes is distinctly different. The Pacific pCo2 remains near equilibrium all the way to Siberia and the Antarctic. In the sub-polar Atlantic, however, there is a distinct decrease in both hemispheres.

Fig. 3.6. Cross-section of surface Pco2 in the Atlantic and Pacific Oceans, observed on GEOSECS cruises of the 1970s. The units are microatmospheres, or parts per million, of CO2. The broken line shows the atmospheric concentration at that time.

Fig. 3.6. Cross-section of surface Pco2 in the Atlantic and Pacific Oceans, observed on GEOSECS cruises of the 1970s. The units are microatmospheres, or parts per million, of CO2. The broken line shows the atmospheric concentration at that time.

The poleward difference between the two great ocean basins is a direct result of the oceanic circulation. In the far North Atlantic, and the Weddell Sea off Antarctica, bottom and intermediate waters are formed. This takes surface water into the deep ocean, and continually renews the water exposed to the surface. This cold, carbon-dioxide-poor water then takes up CO2 from the atmosphere, but is mixed down before reaching equilibrium. The Pacific circulation does not have significant regions of deep water formation, even in the Southern Ocean.1 There is also less formation of intermediate water in the North Pacific gyre. The oceanic carbon dioxide flux is therefore closer to equilibrium in the slower vertical mixing regime of the sub-polar Pacific. The sub-polar Atlantic Ocean is thus the major oceanic sink of atmospheric carbon dioxide.

The 20% increase in surface oceanic carbon dioxide content at the equator is also due to the circulation. In §2.11.3 we saw that Ekman divergence in the upper ocean caused upwelling along the equator. This upwelled water is colder than characteristic tropical sea surface temperatures, giving it a higher solubility. It has also come from regions of higher pressure; this means that the source water for the upwelling has a compressed, and so higher, quantity of carbon dioxide. The effect of decreasing the pressure as the water rises to the surface is to let the CO2 expand. The water itself, however, is only slightly compressible, compared to a gas, and so it changes its volume little. The net effect is to increase the upwelling water's carbon dioxide concentration. Fig. 3.7 shows the potential pco2 if water of different depths and locations in the eastern Pacific were raised to the surface. Upwelling of water from as close to the surface as 100 m would be sufficient to give pCO2 values well above those observed. These are not at their potential level because carbon dioxide is given up to the atmosphere rapidly in these regions, and much carbon is gathered from the mixed layer by plankton. The upwelling occurs too quickly for equilibrium with the atmosphere to be

1 Some deep water formation occurs in the Ross Sea in the Antarctic sector of the South Pacific and north of the Aleutian Islands in the North Pacific.

Fig. 3.7. Potential CO2 pressure as a function of depth in the equatorial Pacific. Potential CO2 pressure is the CO2 partial pressure a sample of water would achieve if it were depressurized to 1 atmosphere pressure and warmed to 25°C, without changing the chemical composition of the sample. [Fig. 3.25 of Broecker and Peng (1982). Reproduced with permission of W. S. Broecker.]

Fig. 3.7. Potential CO2 pressure as a function of depth in the equatorial Pacific. Potential CO2 pressure is the CO2 partial pressure a sample of water would achieve if it were depressurized to 1 atmosphere pressure and warmed to 25°C, without changing the chemical composition of the sample. [Fig. 3.25 of Broecker and Peng (1982). Reproduced with permission of W. S. Broecker.]

achievable, and for a significant anomaly in pCO2 to be maintained. Upwelling rates of several tens of metres per year have been estimated to be required to maintain the observed imbalance.

The relative stability of the tropical atmospheric, and hence surface oceanic, circulation in the Pacific leads to the large but equatorially-confined pCO2 anomaly. By contrast, in the Atlantic there is considerable latitudinal shifting of the ITCZ through the year. This causes the dynamics of upwelling to vary seasonally in strength, and leads to a less dramatic, but more extensive, anomaly.

While the mean latitudinal variation of pCO2 is as shown in Fig. 3.6 there are further local variations of climatic significance. Upwelling regions in general will display properties similar to the equatorial zone, and thus have high carbon dioxide contents. This is evident in Fig. 3.2 in the east Atlantic and northwestern Indian Ocean, for instance.

The warm western boundary currents are also strikingly anomalous regions. The warm water, while at equilibrium with the atmosphere when it begins its poleward travel, is poor in dissolved carbon dioxide, because of the low solubility at such temperatures. As it moves polewards the water is cooled, allowing more carbon dioxide to enter solution. The rapid poleward flow tends to lead to faster cooling than the equilibration processes can keep up with. There should therefore be considerable CO2 uptake in areas of the oceans such as the Gulf Stream and the Kuroshio Current. This is consistent with Fig. 3.8, a map of the upper ocean total carbon content for the Atlantic. Total carbon is defined as the sum of dissolved carbon dioxide, bicarbonate and carbonate ions.

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