This CO 2 reduction in the form of organic matter uses
photons provided by the sun. The quality of organic matter is
very diverse, but it has the major characteristic of storing
reduced carbon and energy. Respiration is the reverse process: i.e., the oxidation of organic matter which converts the
energy stored in the chemical links between organic
molecules into energy which can be used by living cells.
The CO 2 fluxes derived from continental photosynthesis (or
Gross Primary Production*, GPP) and respiration are 120
and 119.6 Pg C year
À1 , respectively (Fig. 14.5). Half of GPP
is respired by plants themselves (autotrophic respiration*),
leaving a Net Primary Production* (NPP) of about
60 Pg C year
À1 . A large fraction of the NPP is used by
heterotrophic respiration (Bacteria, Archaea, Fungi, and
primary consumers), leaving a Net Ecosystem Production*
(NEP) of the order of 10 Pg C year
À1 . Finally, the NEP is
affected by other carbon losses (fire, harvest, erosion, export
of particulate, and dissolved carbon into the ocean), leaving
the net biome production (NBP) which represents the net
amount of carbon accumulated in the continental biosphere.
The NBP is subject to important inter-annual variations
(0.2 and 1.4 Pg C year
À1 in the years 1980 and 1990,
respectively). The continental biosphere is thus currently a
source of carbon sink, despite large releases of CO 2 from
deforestation and, more generally, from changes in land use.
Ocean–Atmosphere Fluxes
Three pumps* allow the ocean to store carbon:
1. The chemical pump is the property, described above, of
the seawater carbonate system to absorb inorganic carbon
in the form CO 2 and store it as bicarbonate and carbonate;
2. The solubility pump is based on the fact that the CO 2
solubility increases with decreasing temperature, so that
cold and CO 2 -rich waters at high latitudes sink and contribute to the deep circulation, allowing the ocean surface
to continuously absorb atmospheric CO 2 ;
The biological pump* also transfers carbon in the deep
ocean. Only part of the primary production of phytoplankton
is respired in the surface waters, the remainder is exported in
the deep ocean in the form of Particulate Organic Matter
(POM)* and Dissolved Organic Matter (DOM)*. This
exported fraction is either respired in the deep waters and
transported back to surface water by oceanic circulation
sooner or later, or is stored in the sediments. The combined
action of the solubility and biological pumps is to maintain a
concentration of Dissolved Inorganic Carbon (DIC) about
10 % lower in surface than in deep water.
In addition to these three pumps, the so-called carbonate
pump produces the opposite effect since it is a source of CO 2
during the precipitation calcium carbonate (CaCO 3 ):
Ca
2þ
þ2HCO 3
À
! CaCO 3 þ CO 2 þ H 2 O
In the short term, the ocean–atmosphere fluxes are of the
order of 90 year Pg C year
À1 in both directions (Fig. 14.5).
They depend on the gradient of CO 2 between the ocean and
atmosphere. The direction and intensity of the air-sea fluxes
vary considerably depending on temperature, wind forcing,
oceanic circulation, and the strength of the biological pump. In
contrast to warm waters, cold waters tend to absorb CO 2 . The
most important control is exercised, however, by the vertical
movements of the water masses, primarily upwellings. The
gross primary production is of the same order of magnitude in
the ocean and on land (103 vs.120 Pg C year
À1
). Only 11 Pg C
year
À1 escape respiration in the surface layer and is exported
in the deep ocean where the bulk is oxidized. Indeed, only
0.01 Pg C year
À1 is stored in the sediments for long-life time
carbon cycling.
Continent-Ocean Fluxes
Rivers transfer directly, that is without passing through the
atmosphere, a small amount of inorganic carbon from the
continents to the ocean. This contribution, about 0.4 Pg C
year
À1 , is roughly offset by the storage of calcium carbonate
in sediment (0.2 Pg C year
À1 ) and by a CO 2 emission to
atmosphere (0.1 Pg C year
À1 ). In addition to inorganic
carbon, rivers also provide organic carbon up to 0.3–0.5 Pg
C year
À1 .
14.2.2.4 Anthropogenic Changes in the Carbon
Cycle and the Feedbacks on the Climate
Emissions from the Fossil Carbon Combustion
The natural fluxes mentioned in the previous section do not
alter the total amount of carbon flowing in the short-term
carbon cycle. They only cause a redistribution of carbon
between its reservoirs. In contrast, the oxidation of fossil
fuels represents a net gain for the short-term carbon cycle.
Until the 1800s, the contribution of the fossil reservoir to this
cycle was marginal. This is no longer the case since the
industrial revolution as the CO 2 emissions produced by
human activities have increased exponentially (Fig. 14.7a)
to reach 6.4 Pg C year
À1 between 1990 and 2000, and 7.2 Pg
C year
À1 between 2000 and 2005. Although these emissions
are responsible for most of the increase in atmospheric
carbon (Fig. 14.5), they are small compared with the natural
fluxes from photosynthesis and respiration (~120 Pg C
year
À1 ) or from the exchanges between ocean and atmosphere (~90 Pg C year
À1 ). Emissions related to the changes
in land use are about 1.6 Pg C year
À1 for the period
1990–2000. The cumulative amount of this transfer of carbon from the long- to the short-term cycle is about 245 Pg C
since 1750. CO 2 emissions from human activities can accumulate in the atmosphere, or be absorbed by the terrestrial
biosphere or by the ocean.
14 Biogeochemical Cycles
519
photons provided by the sun. The quality of organic matter is
very diverse, but it has the major characteristic of storing
reduced carbon and energy. Respiration is the reverse process: i.e., the oxidation of organic matter which converts the
energy stored in the chemical links between organic
molecules into energy which can be used by living cells.
The CO 2 fluxes derived from continental photosynthesis (or
Gross Primary Production*, GPP) and respiration are 120
and 119.6 Pg C year
À1 , respectively (Fig. 14.5). Half of GPP
is respired by plants themselves (autotrophic respiration*),
leaving a Net Primary Production* (NPP) of about
60 Pg C year
À1 . A large fraction of the NPP is used by
heterotrophic respiration (Bacteria, Archaea, Fungi, and
primary consumers), leaving a Net Ecosystem Production*
(NEP) of the order of 10 Pg C year
À1 . Finally, the NEP is
affected by other carbon losses (fire, harvest, erosion, export
of particulate, and dissolved carbon into the ocean), leaving
the net biome production (NBP) which represents the net
amount of carbon accumulated in the continental biosphere.
The NBP is subject to important inter-annual variations
(0.2 and 1.4 Pg C year
À1 in the years 1980 and 1990,
respectively). The continental biosphere is thus currently a
source of carbon sink, despite large releases of CO 2 from
deforestation and, more generally, from changes in land use.
Ocean–Atmosphere Fluxes
Three pumps* allow the ocean to store carbon:
1. The chemical pump is the property, described above, of
the seawater carbonate system to absorb inorganic carbon
in the form CO 2 and store it as bicarbonate and carbonate;
2. The solubility pump is based on the fact that the CO 2
solubility increases with decreasing temperature, so that
cold and CO 2 -rich waters at high latitudes sink and contribute to the deep circulation, allowing the ocean surface
to continuously absorb atmospheric CO 2 ;
The biological pump* also transfers carbon in the deep
ocean. Only part of the primary production of phytoplankton
is respired in the surface waters, the remainder is exported in
the deep ocean in the form of Particulate Organic Matter
(POM)* and Dissolved Organic Matter (DOM)*. This
exported fraction is either respired in the deep waters and
transported back to surface water by oceanic circulation
sooner or later, or is stored in the sediments. The combined
action of the solubility and biological pumps is to maintain a
concentration of Dissolved Inorganic Carbon (DIC) about
10 % lower in surface than in deep water.
In addition to these three pumps, the so-called carbonate
pump produces the opposite effect since it is a source of CO 2
during the precipitation calcium carbonate (CaCO 3 ):
Ca
2þ
þ2HCO 3
À
! CaCO 3 þ CO 2 þ H 2 O
In the short term, the ocean–atmosphere fluxes are of the
order of 90 year Pg C year
À1 in both directions (Fig. 14.5).
They depend on the gradient of CO 2 between the ocean and
atmosphere. The direction and intensity of the air-sea fluxes
vary considerably depending on temperature, wind forcing,
oceanic circulation, and the strength of the biological pump. In
contrast to warm waters, cold waters tend to absorb CO 2 . The
most important control is exercised, however, by the vertical
movements of the water masses, primarily upwellings. The
gross primary production is of the same order of magnitude in
the ocean and on land (103 vs.120 Pg C year
À1
). Only 11 Pg C
year
À1 escape respiration in the surface layer and is exported
in the deep ocean where the bulk is oxidized. Indeed, only
0.01 Pg C year
À1 is stored in the sediments for long-life time
carbon cycling.
Continent-Ocean Fluxes
Rivers transfer directly, that is without passing through the
atmosphere, a small amount of inorganic carbon from the
continents to the ocean. This contribution, about 0.4 Pg C
year
À1 , is roughly offset by the storage of calcium carbonate
in sediment (0.2 Pg C year
À1 ) and by a CO 2 emission to
atmosphere (0.1 Pg C year
À1 ). In addition to inorganic
carbon, rivers also provide organic carbon up to 0.3–0.5 Pg
C year
À1 .
14.2.2.4 Anthropogenic Changes in the Carbon
Cycle and the Feedbacks on the Climate
Emissions from the Fossil Carbon Combustion
The natural fluxes mentioned in the previous section do not
alter the total amount of carbon flowing in the short-term
carbon cycle. They only cause a redistribution of carbon
between its reservoirs. In contrast, the oxidation of fossil
fuels represents a net gain for the short-term carbon cycle.
Until the 1800s, the contribution of the fossil reservoir to this
cycle was marginal. This is no longer the case since the
industrial revolution as the CO 2 emissions produced by
human activities have increased exponentially (Fig. 14.7a)
to reach 6.4 Pg C year
À1 between 1990 and 2000, and 7.2 Pg
C year
À1 between 2000 and 2005. Although these emissions
are responsible for most of the increase in atmospheric
carbon (Fig. 14.5), they are small compared with the natural
fluxes from photosynthesis and respiration (~120 Pg C
year
À1 ) or from the exchanges between ocean and atmosphere (~90 Pg C year
À1 ). Emissions related to the changes
in land use are about 1.6 Pg C year
À1 for the period
1990–2000. The cumulative amount of this transfer of carbon from the long- to the short-term cycle is about 245 Pg C
since 1750. CO 2 emissions from human activities can accumulate in the atmosphere, or be absorbed by the terrestrial
biosphere or by the ocean.
14 Biogeochemical Cycles
519