7.3.4 Ocean Carbon Cycle Processes and Feedbacks to Climate
188.8.131.52 Overview of the Ocean Carbon Cycle
Oceanic carbon exists in several forms: as DIC, DOC, and particulate organic carbon (POC) (living and dead) in an approximate ratio DIC:DOC:POC = 2000:38:1 (about 37,000 GtC DIC: Falkowski et al., 2000 and Sarmiento and Gruber, 2006; 685 GtC DOC: Hansell and Carlson, 1998; and 13 to 23 GtC POC: Eglinton and Repeta, 2004). 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.
Seawater can, through inorganic processes, absorb large amounts of CO2 from the atmosphere, because 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: Degens et al., 1984; Royal Society, 2005). 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 (Wanninkhof and McGillis, 1999; Nightingale et al., 2000) 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 (Zeebe and Wolf-Gladrow, 2001; see Box 7.3).
Box 7.3: Marine Carbon Chemistry and Ocean Acidification
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 HCO3– and CO32–. These three dissolved forms (collectively known as DIC) are found in the approximate ratio CO2:HCO3–:CO32– of 1:100:10 (Equation (7.1)). CO2 is a weak acid and 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 (Equation (7.2)).
|CO2 + H2O → H+ + HCO3– → 2H+ + CO32– ||(7.1) |
|CO2 + H2O + CO32– → HCO3– + H+ + CO32– → 2HCO3– ||(7.2) |
Therefore, the net result of adding CO2 to seawater is an increase in H+ and HCO3–, but a reduction in CO32–. The decrease in the CO32– ion reduces the overall buffering capacity as CO2 increases, with the result that proportionally more H+ ions remain in solution and increase acidity.
This ocean acidification is leading to a decrease in the saturation state of CaCO3 in the ocean. Two primary effects are expected: (1) the biological production of corals as well as calcifying phytoplankton and zooplankton within the water column may be inhibited or slowed down (Royal Society, 2005), and (2) the dissolution of CaCO3 at the ocean floor will be enhanced (Archer, 2005). Aragonite, the meta-stable form of CaCO3 produced by corals and pteropods (planktonic snails; Lalli and Gilmer, 1989), will be particularly susceptible to a pH reduction (Kleypas et al., 1999b; Hughes et al., 2003; Orr et al., 2005). Laboratory experiments under high ambient CO2 with the coccolithophore species Emiliania huxleyi and Gephyrocapsa oceanica produce a significant reduction in CaCO3 production and a stimulation of POC production (Riebesell et al., 2000; Zondervan et al., 2001). Other species and growth under other conditions may show different responses, so that no conclusive quantification of the CaCO3 feedback is possible at present (Tortell et al., 2002; Sciandra et al., 2003).
The sinking speed of marine particle aggregates depends on their composition: CaCO3 may act as an efficient ballast component, leading to high sinking speeds of aggregates (Armstrong et al., 2002; Klaas and Archer, 2002). The relatively small negative feedback of reduced CaCO3 production to atmospheric pCO2 may be compensated for by a change in the ballast for settling biogenic particles and the associated shallowing of re-mineralization depth levels in the water column for organic carbon (Heinze, 2004). On the other hand, production of extracellular organic carbon could increase under high CO2 levels and lead to an increase in export (Engel et al., 2004).
Ecological changes due to expected ocean acidification may be severe for corals in tropical and cold waters (Gattuso et al., 1999; Kleypas et al., 1999a; Langdon et al., 2003; Buddemeier et al., 2004; Roberts et al., 2006) and for pelagic ecosystems (Tortell et al., 2002; Royal Society, 2005). Acidification can influence the marine food web at higher trophic levels (Langenbuch and Pörtner, 2003; Ishimatsu et al., 2004).
Since the beginning of the industrial revolution, sea surface pH has dropped by about 0.1 pH units (corresponding to a 30% increase in the H ion concentration). The expected continued decrease may lead within a few centuries to an ocean pH estimated to have occurred most recently a few hundred million years before present (Caldeira and Wickett, 2003; Key et al., 2004; Box 7.3, Figure 1).
According to a model experiment based on the IPCC Scenarios 1992a (IS92a) emission scenario, bio-calcification will be reduced by 2100, in particular within the Southern Ocean (Orr et al., 2005), and by 2050 for aragonite-producing organisms (see also Figure 10.24). It is important to note that ocean acidification is not a direct consequence of climate change but a consequence of fossil fuel CO2 emissions, which are the main driver of the anticipated climate change
Box 7.3, Figure 1. (a) Atmospheric CO2 emissions, historical atmospheric CO2 levels and predicted CO2 concentrations from the given emission time series, together with changes in ocean pH based on horizontally averaged chemistry. The emission time series is based on the mid-range IS92a emission scenario (solid line) prior to 2100 and then assumes that emissions continue until fossil fuel reserves decline. (b) Estimated maximum change in surface ocean pH as a function of final atmospheric CO2 pressure, and the transition time over which this CO2 pressure is linearly approached from 280 ppm. A: Glacial-interglacial CO2 changes; B: slow changes over the past 300 Myr; C: historical changes in ocean surface waters; D: unabated fossil fuel burning over the next few centuries. Source: Caldeira and Wickett (2003). Reprinted with permission from Macmillan Publishers Ltd: Nature, Caldeira and Wickett (2003), copyright (2003).
In addition to changes in advection and mixing, the ocean can alter atmospheric CO2 concentration through three mechanisms (Volk and Hoffert, 1985), illustrated in Figure 7.10: (1) absorption or release of CO2 due to changes in solubility of gaseous CO2 (‘solubility pump’); (2) changes in carbon fixation to POC in surface waters by photosynthesis and export of this carbon through sinking of organic particles out of the surface layer (‘organic carbon pump’) – this process is limited to first order by availability of light and nutrients (phosphate, nitrate, silicic acid and micronutrients such as iron); and (3) changes in the release of CO2 in surface waters during formation of CaCO3 shell material by plankton (‘CaCO3 counter pump’).
Figure 7.10. 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. Reprinted with permission, copyright 1991 American Geophysical Union.
Organic particles are re-mineralized (oxidized to DIC and other inorganic compounds through the action of bacteria) primarily in the upper 1,000 m of the oceanic water column, with an accompanying decrease in dissolved O. On the average, CaCO3 particles sink deeper before they undergo dissolution: deep waters are undersaturated with respect to CaCO3. The remainder of the particle flux enters marine sediments and is subject to either re-dissolution within the water column or accumulation within the sediments. Although the POC reservoir is small, it plays an important role in keeping DIC concentrations low in surface waters and high in deep waters. The loop is closed through the three-dimensional ocean circulation: upwelling water brings inorganic carbon and nutrients to the surface again, leading to outgassing and biogenic particle production. Dissolved organic carbon enters the ocean water column from rivers and marine metabolic processes. A large fraction of DOC has a long ocean residence time (1–10 kyr), while other fractions are more short-lived (days to hundreds of years; Loh et al., 2004). The composition of dissolved organic matter is still largely unknown.
In conjunction with the global ocean mixing or overturning time of the order of 1 kyr (Broecker and Peng, 1982), small changes in the large ocean carbon reservoir can induce significant changes in atmospheric CO2 concentration. Likewise, perturbations in the atmospheric pCO2 can be buffered by the ocean. Glacial-interglacial changes in the atmospheric CO2 content can potentially be attributed to a change in functioning of the marine carbon pump (see Chapter 6). The key role for the timing of the anthropogenic carbon uptake by the ocean is played by the downward transport of surface water, with a high burden of anthropogenic carbon, into the ocean’s interior. The organic carbon cycle and the CaCO3 counter pump modulate, but do not dominate, the net marine uptake of anthropogenic carbon.