Air-sea gas exchange is a physico-chemical process, primarily controlled by the air-sea difference in gas concentrations and the exchange coefficient, which determines how quickly a molecule of gas can move across the ocean-atmosphere boundary (see Le Quere and Metzl, Chapter 12, this volume). It takes about one year to equilibrate CO2 in the surface ocean with atmospheric CO2, so it is not unusual to observe large air-sea differences in CO2 concentrations (Colorplate 2). Most of the differences are caused by variability in the oceans due to biology and ocean circulation. The oceans contain a large reservoir of carbon that can be exchanged with the atmosphere (see Colorplate 1). Since air-sea exchange can only occur at the surface, however, the rate at which carbon exchanges between the surface and the ocean interior ultimately regulates how well the atmosphere equilibrates with the ocean as a whole.
Two basic mechanisms control the natural distribution of carbon in the ocean interior: the solubility pump and the biological pump. The solubility pump is driven by two principle factors. First, more CO2 can dissolve into cold polar waters than in the warm equatorial waters. As major ocean currents (e.g., the Gulf Stream) move waters from the tropics to the poles, they are cooled and can take up more CO2 from the atmosphere. Second, the high-latitude zones are places where deepwater is formed. As the water cools, it becomes denser and sinks into the ocean's interior, taking with it the CO2 accumulated at the surface.
The biological pump also transports CO2 from the surface to the deep ocean. Growth of phytoplankton uses CO2 and other chemicals from the seawater to form plant tissue. Roughly 70 percent of the CO2 taken up by phytoplankton is recycled near the surface, and the remaining 30 percent sinks into the deeper waters before being converted back into CO2 by marine bacteria (Falkowski et al. 1998). Only about 0.1 percent of the organic carbon fixed at the surface reaches the seafloor to be buried in the sediments. The carbon that is recycled at depth is transported large distances by currents to areas where the waters return to the surface (upwelling regions). When the waters regain contact with the atmosphere, the CO2 originally taken up by the phytoplankton is returned to the atmosphere. This exchange helps to control atmospheric CO2 concentrations over decadal and longer time scales.
The amount of organic carbon that is formed and sinks out of the surface ocean is limited by the availability of light and nutrients (mainly nitrate, phosphate, silicate, and iron) and by temperature. The plankton types present in the water also play a role. Plankton that bloom create favorable conditions for the formation of fast-sinking particles, particularly when they have shells of calcium carbonate or silicate (Klaas and Archer 2002). The formation of calcium carbonate shells affects carbon chemistry in such a way that it works to counteract the drawdown of CO2 by soft tissue production. The biological pump also removes inorganic nutrients from surface waters and releases them at depth. Since productivity is limited by the availability of these nutrients, the large-scale thermohaline circulation (THC) of the oceans has a strong impact on global ocean productivity by regulating the rate at which nutrients are returned to the surface.
Up to now, humans have had a relatively small direct impact on the global-scale ocean carbon cycle. This is primarily because humans generally only transit across the ocean and because the ocean naturally contains orders of magnitude more carbon than the atmosphere and the terrestrial biosphere. The ocean does, however, act as a significant sink for CO2 ultimately derived from anthropogenic activities (Table 2.1). Because biology is not limited by carbon in the oceans, it is thought that increasing CO2 levels have not significantly affected ocean biology. The current distribution of anthropogenic CO2 is assumed to result from physico-chemical equilibration of the surface ocean with rising atmospheric CO2 and slow mixing of the anthropogenic CO2 into the ocean's interior. The long residence time for the deep oceans means that most of the deep ocean waters have not been exposed to the rising atmospheric CO2 concentrations observed over the past couple of centuries. Although the oceans have the potential to absorb 85 percent of the anthropogenic CO2 released to the atmosphere, today's oceans are only at about 15 percent capacity (Le Quere and Metzl, Chapter 12, this volume). Average penetration depth for anthropogenic CO2 in the global ocean is only about 800 m (Sabine et al. 2002). There is growing evidence, however, that changes in ocean mixing and biology may be occurring as a result of climate change.
Since the solubility of CO2 is a function of temperature, warming of the ocean will decrease its ability to absorb CO2. Furthermore, changes in temperature and precipitation may lead to significant alterations of ocean circulation and the transport of carbon and nutrients to and from the surface. Because of its effect on ocean carbon distributions and biological productivity, changes in the THC have been used to help explain past excursions in climate and atmospheric CO2, including glacial-interglacial and Dansgard-Oeschger events (Joos and Prentice, Chapter 7, this volume).
Ocean productivity can also be affected by atmospheric inputs that may change as a result of human activity. Iron in oceanic surface waters originates from terrestrial dust deposited over the ocean, deep ocean waters, continental shelves, and to a lesser extent river inflow. Because of the spatial distribution of dust deposition and other iron sources, large regions of the ocean show a deficit in iron (and to a smaller extent in silicate), although other nutrients are plentiful. These are called high-nutrient low-chlorophyll (HNLC) regions. There is a potential for enhanced biological productivity in these regions if the ocean can be "fertilized" by iron. The potential for CO2 reduction in surface waters and in the atmosphere through this artificial sequestration, however, depends on factors like the composition of plankton types and oceanic circulation, and can lead to undesirable side effects (Bakker, Chapter 26, this volume).
There is also increasing evidence that rising CO2 levels may directly affect ocean productivity and ecosystem structure. For example, Riebesell et al. (2000) showed a significant reduction in the ability of two different species of coccoliths to secrete calcium carbonate shells under elevated CO2 conditions. Similar reductions in calcification have been observed in corals and coralline algae. As atmospheric CO2 concentrations continue to rise, the potential for significantly altering the current balance between the amount of carbon moved into the ocean's interior by the biological pump versus the solubility pump in the ocean increases. The net effect on the ability of the ocean to act as a sink for anthropogenic CO2 is not clear.
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