The Anthropogenic Carbon Cycle
The Anthropogenic carbon Cycle as described in the Intergovernmental Panel on Climate Change’s 4th Assessment Report by Working Group One (IPCC AR4 WG1):
natural and perturbed carbon cycle between the land, atmosphere and
ocean is shown in IPCC AR4WG1 Figure 7.3. Carbon dioxide cycles
between the atmosphere, oceans and land biosphere. Its removal from
the atmosphere involves a range of processes with different time
scales. About 50% of a CO2 increase will be removed from the
atmosphere within 30 years, and a further 30% will be removed within
a few centuries. The remaining 20% may stay in the atmosphere for
many thousands of years. Atmospheric carbon dioxide (CO2)
concentration has continued to increase and is now almost 100 ppm
above its pre-industrial level. The annual mean CO2 growth rate was
significantly higher for the period from 2000 to 2005 (4.1 ± 0.1 GtC
yr–1) than it was in the 1990s (3.2 ± 0.1 GtC yr–1). Annual
emissions of CO2 from fossil fuel burning and cement production
increased from a mean of 6.4 ± 0.4 GtC yr–1 in the 1990s to 7.2 ±
0.3 GtC yr–1 for 2000 to 2005.
IPCC AR4WG1 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., 2004). 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.
IPCC AR4WG1 Table 7.1. The global carbon budget (GtC yr–1); errors represent ±1 standard deviation uncertainty estimates and not interannual variability, which is larger. The atmospheric increase (first line) results from fluxes to and from the atmosphere: positive fluxes are inputs to the atmosphere (emissions); negative fluxes are losses from the atmosphere (sinks); and numbers in parentheses are ranges. Note that the total sink of anthropogenic CO2 is well constrained. Thus, the ocean-to-atmosphere and land-to-atmosphere fluxes are negatively correlated: if one is larger, the other must be smaller to match the total sink, and vice versa.
||3.3 ± 0.1
||3.3 ± 0.1
||3.2 ± 0.1
||3.2 ± 0.1||4.1 ± 0.1|
|Emissions (fossil + cement)c
||5.4 ± 0.3||5.4 ± 0.3||6.4 ± 0.4||6.4 ± 0.4||7.2 ± 0.3|
|Net ocean-to-atmosphere fluxd
||-1.9 ± 0.6||-1.8 ± 0.8||-1.7 ± 0.5||-2.2 ± 0.4||-2.2 ± 0.5|
|Net land-to-atmosphere fluxe
||-0.2 ± 0.7||-0.3 ± 0.9||-1.4 ± 0.7||-1.0 ± 0.6||-0.9 ± 0.6|
|Partitioned as follows
| Land use change flux
(0.6 to 2.5)
(0.4 to 2.3)
(0.5 to 2.7)
| Residual terrestrial sink
(-3.8 to -0.3)
(-3.4 to 0.2)
(-4.3 to -0.9)
a TAR values revised according to an ocean heat content correction for ocean oxygen fluxes (Bopp et al., 2002) and using the Fourth Assessment Report (AR4) best estimate for the land use change flux given in Table 7.2.
b Determined from atmospheric CO2 measurements (Keeling and Whorf, 2005, updated by S. Piper until 2006) at Mauna Loa (19°N) and South Pole (90°S) stations, consistent with the data shown in Figure 7.4, using a conversion factor of 2.12 GtC yr-1 = 1 ppm.
c Fossil fuel and cement emission data are available only until 2003 (Marland et al., 2006). Mean emissions for 2004 and 2005 were extrapolated from energy use data with a trend of 0.2 GtC yr-1.
d For the 1980s, the ocean-to-atmosphere and land-to-atmosphere fluxes were estimated using atmospheric O2:N2 and CO2 trends, as in the TAR. For the 1990s, the ocean-to-atmosphere flux alone is estimated using ocean observations and model results (see Section 188.8.131.52.1), giving results identical to the atmospheric O2:N2 method (Manning and Keeling, 2006), but with less certainty. The net land-to-atmosphere flux then is obtained by subtracting the ocean-to-atmosphere flux from the total sink (and its errors estimated by propagation). For 2000 to 2005, the change in ocean-to-atmosphere flux was modelled (Le Quéré et al., 2005) and added to the mean ocean-to-atmosphere flux of the 1990s. The error was estimated based on the quadratic sum of the error of the mean ocean flux during the 1990s and the root mean square of the five-year variability from three inversions and one ocean model presented in Le Quéré et al. (2003).
e Balance of emissions due to land use change and a residual land sink. These two terms cannot be separated based on current o
How much anthropogenic carbon will the terrestrial biosphere take up over this next century and beyond?
The terrestrial biosphere interacts strongly with the climate, providing both positive and negative feedbacks due to biogeophysical and biogeochemical processes. Some of these feedbacks, at least on a regional basis, can be large (e.g. Chagnon et al., 2004). Surface climate is determined by the balance of fluxes, which can be changed by radiative (e.g., albedo) or non-radiative (e.g., water cycle related processes) terms. Both radiative and non-radiative terms are controlled by details of vegetation. High-latitude climate is strongly influenced by snow albedo feedback, which is drastically reduced by the darkening effect of vegetation. In semi-arid tropical systems, such as the Sahel or northeast Brazil, vegetation exerts both radiative and hydrological feedbacks (e.g. Oyama and Nobre, 2004). Surface climate interacts with vegetation cover, biomes, productivity, respiration of vegetation and soil, and fires, all of which are important for the carbon cycle. Various processes in terrestrial ecosystems influence the flux of carbon between land and the atmosphere. Terrestrial ecosystem photosynthetic productivity changes in response to changes in temperature, precipitation, CO2 and nutrients. If climate becomes more favorable for growth (e.g., increased rainfall in a semi-arid system), productivity increases, and carbon uptake from the atmosphere is enhanced. Organic carbon compounds in soils, originally derived from plant material, are respired (i.e., oxidized by microbial communities) at different rates depending on the nature of the compound and on the microbial communities; the aggregate rate of respiration depends on soil temperature and moisture. Shifts in ecosystem structure in response to a changing climate can alter the partitioning of carbon between the atmosphere and the land surface. Migration of boreal forest northward into tundra would initially lead to an increase in carbon storage in the ecosystem due to the larger biomass of trees than of herbs and shrubs, but over a longer time (e.g., centuries), changes in soil carbon would need to be considered to determine the net effect. A shift from tropical rainforest to savannah, on the other hand, would result in a net flux of carbon from the land surface to the atmosphere.
Interannual and inter-decadal variability in the growth rate of atmospheric CO2 is dominated by the response of the land biosphere to climate variations. Evidence of decadal changes is observed in the net land carbon sink, with estimates of 0.3 ± 0.9, 1.0 ± 0.6, and 0.9 ± 0.6 GtC yr–1 for the 1980s, 1990s and 2000 to 2005 time periods, respectively. (See references in Table 7.1 of Denman et al., 2007)
A combination of techniques gives an estimate of the flux of CO2 to the atmosphere from land use change of 1.6 (0.5 to 2.7) GtC yr–1 for the 1990s. A revision of the Third Assessment Report (TAR) estimate for the 1980s downwards to 1.4 (0.4 to 2.3) GtC yr–1 suggests little change between the 1980s and 1990s, and continuing uncertainty in the net CO2 emissions due to land use change. (See references in Table 7.1 of Denman et al., 2007)
Fires, from natural causes and human activities, release to the atmosphere considerable amounts of radiatively and photochemically active trace gases and aerosols (e.g. Hoffman et al., 2002). If fire frequency and extent increase with a changing climate, a net increase in CO2 emissions is expected during this fire regime shift (e.g. Flannigan et al., 2005).
The functioning of ocean ecosystems depends strongly on climatic conditions including near-surface density stratification, ocean circulation, temperature, salinity, the wind field and sea ice cover. In turn, ocean ecosystems affect the chemical composition of the atmosphere (e.g. CO2, N2O, oxygen (O2), dimethyl sulphide (DMS) and sulphate aerosol). Most of these components are expected to change with a changing climate and high atmospheric CO2 conditions. Marine biota also influence the near-surface radiation budget through changes in the marine albedo and absorption of solar radiation (biooptical heating). Feedbacks between marine ecosystems and climate change are complex because most involve the ocean’s physical responses and feedbacks to climate change. Increased surface temperatures and stratification should lead to increased photosynthetic fixation of CO2, but associated reductions in vertical mixing and overturning circulation may decrease the return of required nutrients to the surface ocean and alter the vertical export of carbon to the deeper ocean. The sign of the cumulative feedback to climate of all these processes is still unclear. Changes in the supply of micronutrients required for photosynthesis, in particular iron, through dust deposition to the ocean surface can modify marine biological production patterns. Ocean acidification due to uptake of anthropogenic CO2 may lead to shifts in ocean ecosystem structure and dynamics, which may alter the biological production and export from the surface ocean of organic carbon and calcium carbonate (CaCO3).
Improved estimates of ocean uptake of CO2 suggest little change in the ocean carbon sink of 2.2 ± 0.5 GtC yr–1 between the 1990s and the first five years of the 21st century. Models indicate that the fraction of fossil fuel and cement emissions of CO2 taken up by the ocean will decline if atmospheric CO2 continues to increase. (See references in Table 7.1 of Denman et al., 2007)
Ocean CO2 uptake has lowered the average ocean pH (increased acidity) by approximately 0.1 since 1750 (Caldeira and Wickett., 2003). Consequences for marine ecosystems may include reduced calcification by shell-forming organisms (Orr et al., 2005), and in the longer term, the dissolution of carbonate sediments (Archer, 2005).
The first-generation coupled climate-carbon cycle models indicate that global warming will increase the fraction of anthropogenic CO2 that remains in the atmosphere. This positive climate-carbon cycle feedback leads to an additional increase in atmospheric CO2 concentration of 20 to 224 ppm by 2100, in models run under the IPCC (2000) Special Report on Emission Scenarios (SRES) A2 emissions scenario (Friedlingstein et al., 2006).
At GFDL, we are developing and implementing the state of the art in terrestrial carbon models by including our best scientific understanding of the processes that control carbon cycling under ecological and climate dynamics, fire, land use, and other factors to determine the net balance of uptake and release of CO2 by the terrestrial biosphere. On the ocean side, GFDL is developing and implementing state of the art ocean biogeochemical models with state of the art ocean physical models to better characterize the ocean circulation and coupled carbon cycle responses to global climate change and acidification in order to reduce uncertainties in the ocean uptake of atmospheric CO2.
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