7.3 The Carbon Cycle and the Climate System
7.3.1 Overview of the Global Carbon Cycle
126.96.36.199 The Natural Carbon Cycle
Over millions of years, CO2 is removed from the atmosphere through weathering by silicate rocks and through burial in marine sediments of carbon fixed by marine plants (e.g., Berner, 1998). Burning fossil fuels returns carbon captured by plants in Earth’s geological history to the atmosphere. New ice core records show that the Earth system has not experienced current atmospheric concentrations of CO2, or indeed of CH4, for at least 650 kyr – six glacial-interglacial cycles. During that period the atmospheric CO2 concentration remained between 180 ppm (glacial maxima) and 300 ppm (warm interglacial periods) (Siegenthaler et al., 2005). It is generally accepted that during glacial maxima, the CO2 removed from the atmosphere was stored in the ocean. Several causal mechanisms have been identified that connect astronomical changes, climate, CO2 and other greenhouse gases, ocean circulation and temperature, biological productivity and nutrient supply, and interaction with ocean sediments (see Box 6.2).
Prior to 1750, the atmospheric concentration of CO2 had been relatively stable between 260 and 280 ppm for 10 kyr (Box 6.2). Perturbations of the carbon cycle from human activities were insignificant relative to natural variability. Since 1750, the concentration of CO2 in the atmosphere has risen, at an increasing rate, from around 280 ppm to nearly 380 ppm in 2005 (see Figure 2.3 and FAQ 2.1, Figure 1). The increase in atmospheric CO2 concentration results from human activities: primarily burning of fossil fuels and deforestation, but also cement production and other changes in land use and management such as biomass burning, crop production and conversion of grasslands to croplands (see FAQ 7.1). While human activities contribute to climate change in many direct and indirect ways, CO2 emissions from human activities are considered the single largest anthropogenic factor contributing to climate change (see FAQ 2.1, Figure 2). Atmospheric CH4 concentrations have similarly experienced a rapid rise from about 700 ppb in 1750 (Flückiger et al., 2002) to about 1,775 ppb in 2005 (see Section 2.3.2): sources include fossil fuels, landfills and waste treatment, peatlands/wetlands, ruminant animals and rice paddies. The increase in CH4 radiative forcing is slightly less than one-third that of CO2, making it the second most important greenhouse gas (see Chapter 2). The CH4 cycle is presented in Section 7.4.1.
Both CO2 and CH4 play roles in the natural cycle of carbon, involving continuous flows of large amounts of carbon among the ocean, the terrestrial biosphere and the atmosphere, that maintained stable atmospheric concentrations of these gases for 10 kyr prior to 1750. Carbon is converted to plant biomass by photosynthesis. Terrestrial plants capture CO2 from the atmosphere; plant, soil and animal respiration (including decomposition of dead biomass) returns carbon to the atmosphere as CO2, or as CH4 under anaerobic conditions. Vegetation fires can be a significant source of CO2 and CH4 to the atmosphere on annual time scales, but much of the CO2 is recaptured by the terrestrial biosphere on decadal time scales if the vegetation regrows.
Carbon dioxide is continuously exchanged between the atmosphere and the ocean. Carbon dioxide entering the surface ocean immediately reacts with water to form bicarbonate (HCO3–) and carbonate (CO32–) ions. Carbon dioxide, HCO3– and CO32– are collectively known as dissolved inorganic carbon (DIC). The residence time of CO2 (as DIC) in the surface ocean, relative to exchange with the atmosphere and physical exchange with the intermediate layers of the ocean below, is less than a decade. In winter, cold waters at high latitudes, heavy and enriched with CO2 (as DIC) because of their high solubility, sink from the surface layer to the depths of the ocean. This localised sinking, associated with the Meridional Overturning Circulation (MOC; Box 5.1) is termed the ‘solubility pump’. Over time, it is roughly balanced by a distributed diffuse upward transport of DIC primarily into warm surface waters.
Phytoplankton take up carbon through photosynthesis. Some of that sinks from the surface layer as dead organisms and particles (the ‘biological pump’), or is transformed into dissolved organic carbon (DOC). Most of the carbon in sinking particles is respired (through the action of bacteria) in the surface and intermediate layers and is eventually recirculated to the surface as DIC. The remaining particle flux reaches abyssal depths and a small fraction reaches the deep ocean sediments, some of which is re-suspended and some of which is buried. Intermediate waters mix on a time scale of decades to centuries, while deep waters mix on millennial time scales. Several mixing times are required to bring the full buffering capacity of the ocean into effect (see Section 5.4 for long-term observations of the ocean carbon cycle and their consistency with ocean physics).
Together the solubility and biological pumps maintain a vertical gradient in CO2 (as DIC) between the surface ocean (low) and the deeper ocean layers (high), and hence regulate exchange of CO2 between the atmosphere and the ocean. The strength of the solubility pump depends globally on the strength of the MOC, surface ocean temperature, salinity, stratification and ice cover. The efficiency of the biological pump depends on the fraction of photosynthesis exported from the surface ocean as sinking particles, which can be affected by changes in ocean circulation, nutrient supply and plankton community composition and physiology.
Figure 7.3. 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.
In Figure 7.3 the natural or unperturbed exchanges (estimated to be those prior to 1750) among oceans, atmosphere and land are shown by the black arrows. The gross natural fluxes between the terrestrial biosphere and the atmosphere and between the oceans and the atmosphere are (circa 1995) about 120 and 90 GtC yr–1, respectively. Just under 1 GtC yr–1 of carbon is transported from the land to the oceans via rivers either dissolved or as suspended particles (e.g., Richey, 2004). While these fluxes vary from year to year, they are approximately in balance when averaged over longer time periods. Additional small natural fluxes that are important on longer geological time scales include conversion of labile organic matter from terrestrial plants into inert organic carbon in soils, rock weathering and sediment accumulation (‘reverse weathering’), and release from volcanic activity. The net fluxes in the 10 kyr prior to 1750, when averaged over decades or longer, are assumed to have been less than about 0.1 GtC yr–1. For more background on the carbon cycle, see Prentice et al. (2001), Field and Raupach (2004) and Sarmiento and Gruber (2006).