7.3.4.2 Carbon Cycle Feedbacks to Changes in Atmospheric Carbon Dioxide

Chemical buffering of anthropogenic CO₂ is the quantitatively most important oceanic process acting as a carbon sink. Carbon dioxide entering the ocean is buffered due to scavenging by the CO₃^2– ions and conversion to HCO₃^–, that is, the resulting increase in gaseous seawater CO₂ concentration is smaller than the amount of CO₂ added per unit of seawater volume. Carbon dioxide buffering in seawater is quantified by the Revelle factor (‘buffer factor’, Equation (7.3)), relating the fractional change in seawater pCO₂ to the fractional change in total DIC after re-equilibration (Revelle and Suess, 1957; Zeebe and Wolf-Gladrow, 2001):

Revelle factor (or buffer factor) = (Δ[CO₂] / [CO₂]) / (Δ[DIC] / [DIC]) (7.3)

The lower the Revelle factor, the larger the buffer capacity of seawater. Variability of the buffer factor in the ocean depends mainly on changes in pCO₂ and the ratio of DIC to total alkalinity. In the present-day ocean, the buffer factor varies between 8 and 13 (Sabine et al., 2004a; Figure 7.11). With respect to atmospheric pCO₂ alone, the inorganic carbon system of the ocean reacts in two ways: (1) seawater re-equilibrates, buffering a significant amount of CO₂ from the atmosphere depending on the water volume exposed to equilibration; and (2) the Revelle factor increases with pCO₂ (positive feedback; Figure 7.11). Both processes are quantitatively important. While the first is generally considered as a system response, the latter is a feedback process.

The ocean will become less alkaline (seawater pH will decrease) due to CO₂ uptake from the atmosphere (see Box 7.3). The ocean’s capacity to buffer increasing atmospheric CO₂ will decline in the future as ocean surface pCO₂ increases (Figure 7.11a). This anticipated change is certain, with potentially severe consequences.

Figure 7.11. (a) The Revelle factor (or buffer factor) as a function of CO₂ partial pressure (for temperature 25°C, salinity 35 psu, and total alkalinity 2,300 µmol kg^–1) (Zeebe and Wolf-Gladrow, 2001, page 73; reprinted with permission, copyright 2001 Elsevier). (b) The geographical distribution of the buffer factor in ocean surface waters in 1994 (Sabine et al., 2004a; reprinted with permission, copyright 2004 American Association for the Advancement of Science). High values indicate a low buffer capacity of the surface waters.

Increased carbon storage in the deep ocean leads to the dissolution of calcareous sediments below their saturation depth (Broecker and Takahashi, 1978; Feely et al., 2004). The feedback of CaCO₃ sediment dissolution to atmospheric pCO₂ increase is negative and quantitatively significant on a 1 to 100 kyr time scale, where CaCO₃ dissolution will account for a 60 to 70% absorption of the anthropogenic CO₂ emissions, while the ocean water column will account for 22 to 33% on a time scale of 0.1 to 1 kyr. In addition, the remaining 7 to 8% may be compensated by long-term terrestrial weathering cycles involving silicate carbonates (Archer et al., 1998). Due to the slow CaCO₃ buffering mechanism (and the slow silicate weathering), atmospheric pCO₂ will approach a new equilibrium asymptotically only after several tens of thousands of years (Archer, 2005; Figure 7.12).

Figure 7.12. Model projections of the neutralization of anthropogenic CO₂ for an ocean-only model, a model including dissolution of CaCO₃ sediment and a model including weathering of silicate rocks, (top) for a total of 1,000 GtC of anthropogenic CO₂ emissions and (bottom) for a total of 5,000 GtC of anthropogenic CO₂. Note that the y-axis is different for the two diagrams. Without CaCO₃ dissolution from the seafloor, the buffering of anthropogenic CO₂ is limited. Even after 100 kyr, the remaining pCO₂ is substantially higher than the pre-industrial value. Source: Archer (2005).

Elevated ambient CO₂ levels appear to also influence the production rate of POC by marine calcifying planktonic organisms (e.g., Zondervan et al., 2001). This increased carbon fixation under higher CO₂ levels was also observed for three diatom (siliceous phytoplankton) species (Riebesell et al., 1993). It is critical to know whether these increased carbon fixation rates translate into increased export production rates (i.e., removal of carbon to greater depths). Studies of the nutrient to carbon ratio in marine phytoplankton have not yet shown any significant changes related to CO₂ concentration of the nutrient utilisation efficiency (expressed through the ‘Redfield ratio’ – carbon:nitrogen:phosphorus:silicon) in organic tissue (Burkhardt et al., 1999).