Carbon Cycle Changes and Processes
Involved
Long-term Regulation of Atmospheric CO 2
On the geological timescale, and neglecting at this
point the possible impact of sedimentary organic
carbon cycling, the sources of carbon for the ocean–
atmosphere system are the degassing of the mantle
and metamorphic processes. Carbon is injected as
CO 2 into the ocean–atmosphere system today
through the degassing of fresh basalts along midoceanic ridges (MOR) (1.5–2.2 Â 10
12 mol y
À1
), and
through plume events and arc volcanism (1.5–
5.5 Â 10
12 mol y
À1
). These various sources account
for the total degassing flux F VOL . Once released, this
carbon is rapidly (within 10
3 years) redistributed
between the atmosphere and the ocean, reaching a
steady-state repartition after a negligible time compared to the geological timescale. Carbon can leave
the system mainly through the deposition of carbonate minerals on the seafloor. The rate of carbon
removal through carbonate deposition is controlled
by the saturation state of the ocean, and thus by the
rate of supply of alkalinity by the chemical weathering of continental minerals. Carbonate and silicate minerals exposed at the continental surface
weather under the corrosive action of atmospheric
CO 2 dissolved in rain water as carbonic acid, and
eventually concentrated by the microbial respiration
in soils. Regarding the chemical weathering of carbonate minerals, the net budget of the dissolution
reaction can be written as follows eqn [I].
CaCO 3 rock
ð
ÞþH 2 CO 3 atmosphere
ð
Þ
-Ca
2þ rivers
ð
Þþ2HCO
À
3 rivers
ð
Þ
½I
In this reaction, there is a net transfer of Ca
2þ (or
Mg
2þ in the case of magnesium carbonates) from the
continental crust to the ocean, a creation of two
equivalents of alkalinity and a transfer to the river
system of two moles of carbon per mole of Ca
2þ , one
coming from the crust, the other one from the atmosphere. Once the weathering products reach the
ocean, they will increase the saturation state of surface waters with respect to calcite and induce rapidly
(within 10
3 years) the biologically driven precipitation of one mole of CaCO 3 followed by its
deposition on the seafloor. The precipitation–
deposition reaction is the reverse of reaction [1]. The
net carbon budget of the weathering of carbonate
minerals followed by deposition of sedimentary carbonate is thus equal to zero.
The chemical weathering of continental silicate
rocks is fundamentally different, since silicate rocks
do not contain carbon. The budget can be written as
CaSiO 3 rock
ð
Þþ2H 2 CO 3 atmosphere
ð
ÞþH 2 O
-Ca
2þ rivers
ð
Þþ2HCO
À
3 rivers
ð
ÞþH 4 SiO 4 rivers
ð
Þ
½II
Here CaSiO 3 stands for a ‘generic’ silicate mineral.
The weathering reaction of more realistic Ca- (or
Mg-) silicate minerals, if more complex, displays the
same budget in terms of alkalinity versus carbon
fluxes. Again, once reaching the ocean, the excess
Ca
2þ will precipitate as CaCO 3 , thus removing one
mole of carbon from the ocean per mole of weathered silicates. The net budget of this reaction, after
sedimentary carbonate precipitation, is the transfer
of exospheric carbon to the crust. Chemical weathering of continental silicate minerals thus acts as
the main sink of carbon on the geological timescale.
Today, about 6 Â 10
12 mol y
À1 of Ca
2þ and Mg
2þ
are released from silicate weathering.
The size of the exospheric carbon pool
(ocean þ atmosphere) is about 3.2 Â 10
18 mol today.
As mentioned above, the fluxes entering and leaving
this reservoir are of the order of 10
12
–10
13 mol y
À1
.
A relatively small imbalance between the input and
output of carbon of 10
12 mol y
À1
, the output being
higher, but persisting for several million years, will
result in a drastic reduction in the exospheric content. Three million years will be sufficient to remove
all the carbon from the exospheric system, thus forcing the atmospheric PCO 2 to zero. There is no
lithological, fossil, or geochemical record of such a
dramatic event during the Cenozoic, or event during
the complete Phanerozoic. To avoid the occurrence
of such events for which there is no evidence, the
perturbations of the carbon cycle had to be limited in
time and amplitude, and thus the past exospheric
carbon cycle was not strictly at, but always close to,
steady state. The same considerations apply to the
alkalinity budget. These steady-state conditions require that the amount of carbon removed from the
atmosphere–ocean system by continental silicate
weathering must always closely track the amount of
carbon released by degassing. Mathematically, these
conditions translate into eqn [1].
F SW ¼ F VOL
½1
The question is now how to physically force F SW to
follow the degassing. The answer lies in the fact that
the chemical weathering of continental silicates appears to be dependent on air temperature, the dissolution being enhanced during warmer climates.
This dependence provides a negative feedback that
524 CENOZOIC OCEANS – CARBON CYCLE MODELS
Involved
Long-term Regulation of Atmospheric CO 2
On the geological timescale, and neglecting at this
point the possible impact of sedimentary organic
carbon cycling, the sources of carbon for the ocean–
atmosphere system are the degassing of the mantle
and metamorphic processes. Carbon is injected as
CO 2 into the ocean–atmosphere system today
through the degassing of fresh basalts along midoceanic ridges (MOR) (1.5–2.2 Â 10
12 mol y
À1
), and
through plume events and arc volcanism (1.5–
5.5 Â 10
12 mol y
À1
). These various sources account
for the total degassing flux F VOL . Once released, this
carbon is rapidly (within 10
3 years) redistributed
between the atmosphere and the ocean, reaching a
steady-state repartition after a negligible time compared to the geological timescale. Carbon can leave
the system mainly through the deposition of carbonate minerals on the seafloor. The rate of carbon
removal through carbonate deposition is controlled
by the saturation state of the ocean, and thus by the
rate of supply of alkalinity by the chemical weathering of continental minerals. Carbonate and silicate minerals exposed at the continental surface
weather under the corrosive action of atmospheric
CO 2 dissolved in rain water as carbonic acid, and
eventually concentrated by the microbial respiration
in soils. Regarding the chemical weathering of carbonate minerals, the net budget of the dissolution
reaction can be written as follows eqn [I].
CaCO 3 rock
ð
ÞþH 2 CO 3 atmosphere
ð
Þ
-Ca
2þ rivers
ð
Þþ2HCO
À
3 rivers
ð
Þ
½I
In this reaction, there is a net transfer of Ca
2þ (or
Mg
2þ in the case of magnesium carbonates) from the
continental crust to the ocean, a creation of two
equivalents of alkalinity and a transfer to the river
system of two moles of carbon per mole of Ca
2þ , one
coming from the crust, the other one from the atmosphere. Once the weathering products reach the
ocean, they will increase the saturation state of surface waters with respect to calcite and induce rapidly
(within 10
3 years) the biologically driven precipitation of one mole of CaCO 3 followed by its
deposition on the seafloor. The precipitation–
deposition reaction is the reverse of reaction [1]. The
net carbon budget of the weathering of carbonate
minerals followed by deposition of sedimentary carbonate is thus equal to zero.
The chemical weathering of continental silicate
rocks is fundamentally different, since silicate rocks
do not contain carbon. The budget can be written as
CaSiO 3 rock
ð
Þþ2H 2 CO 3 atmosphere
ð
ÞþH 2 O
-Ca
2þ rivers
ð
Þþ2HCO
À
3 rivers
ð
ÞþH 4 SiO 4 rivers
ð
Þ
½II
Here CaSiO 3 stands for a ‘generic’ silicate mineral.
The weathering reaction of more realistic Ca- (or
Mg-) silicate minerals, if more complex, displays the
same budget in terms of alkalinity versus carbon
fluxes. Again, once reaching the ocean, the excess
Ca
2þ will precipitate as CaCO 3 , thus removing one
mole of carbon from the ocean per mole of weathered silicates. The net budget of this reaction, after
sedimentary carbonate precipitation, is the transfer
of exospheric carbon to the crust. Chemical weathering of continental silicate minerals thus acts as
the main sink of carbon on the geological timescale.
Today, about 6 Â 10
12 mol y
À1 of Ca
2þ and Mg
2þ
are released from silicate weathering.
The size of the exospheric carbon pool
(ocean þ atmosphere) is about 3.2 Â 10
18 mol today.
As mentioned above, the fluxes entering and leaving
this reservoir are of the order of 10
12
–10
13 mol y
À1
.
A relatively small imbalance between the input and
output of carbon of 10
12 mol y
À1
, the output being
higher, but persisting for several million years, will
result in a drastic reduction in the exospheric content. Three million years will be sufficient to remove
all the carbon from the exospheric system, thus forcing the atmospheric PCO 2 to zero. There is no
lithological, fossil, or geochemical record of such a
dramatic event during the Cenozoic, or event during
the complete Phanerozoic. To avoid the occurrence
of such events for which there is no evidence, the
perturbations of the carbon cycle had to be limited in
time and amplitude, and thus the past exospheric
carbon cycle was not strictly at, but always close to,
steady state. The same considerations apply to the
alkalinity budget. These steady-state conditions require that the amount of carbon removed from the
atmosphere–ocean system by continental silicate
weathering must always closely track the amount of
carbon released by degassing. Mathematically, these
conditions translate into eqn [1].
F SW ¼ F VOL
½1
The question is now how to physically force F SW to
follow the degassing. The answer lies in the fact that
the chemical weathering of continental silicates appears to be dependent on air temperature, the dissolution being enhanced during warmer climates.
This dependence provides a negative feedback that
524 CENOZOIC OCEANS – CARBON CYCLE MODELS