327
values of alkalinity and total dissolved inorganic
carbon in intermediate and deep waters are the result of dissolving calcite particles settling through
the water column, or of sedimentary dissolution
above the lysocline, would require better knowledge of the total amount of calcite production in
the surface ocean and of the dissolution rates in
the sediment. While the first is very difficult to
measure, the second may be reached by improving
data sets of inorganic carbon flux rates from the
sediment to ocean at intermediate and deep water
levels. The next Section 9.4.2. describes the approach of estimating the fluxes of total carbon and
inorganic carbon resulting from calcite dissolution
in more detail.
9.4.2
Inorganic and Organic Carbon
Release from Deep Sea Sediments
Although only a small fraction of carbon arriving
at the seafloor is finally buried, over geological
time scales deep sea sediments have formed the
largest reservoir of carbon on earth mainly
consisting of biogenic carbonates (~50 Mio. Gt)
and organic detritus (~12 Mio. Gt; i.e. de Baar &
Suess, 1993). This huge storage potential of
marine carbonates is a major factor for maintaining
reasonably low atmospheric CO 2 -levels in earth
history. However, the ratio of accumulation versus
recycling of carbon at the seafloor reflects a
dynamic equilibrium depending on various external forcing factors. Hence, the question for the
“how much is being recycled” is an important
issue, but quite difficult to explore.
Diagenetic mineralization and dissolution
processes in deep sea sediments have been
recognized as important control factors and
numerous efforts have been made to estimate the
contribution by benthic reflux to the ocean
budget. The diagenetic reactions are most intense
at the sediment-water interface where the most
labile components become rapidly mineralized.
Hence, this is the place where the determination
between burial and recycling is made. The driving
forces of carbon release from the sediments are
the degradation of organic matter and the
dissolution of calcium carbonate, which are in turn
dependent on the supply by the remnants of
biological production in the surface waters.
However, the dissolution of calcium carbonate in
deep-sea sediments is controlled by two major
factors, the degree of undersaturation of the deep
ocean waters with respect to calcite and aragonite,
and the reaction with carbon dioxide from
respiration processes. Two of the key parameters
are thus, the “rain ratio” (C CaCO3 /C POC ) of the
sinking material and the water depth of final
deposition, above or below the CO 3
2saturation
horizon (lysocline). Because the solubility of
calcium carbonate increases with increasing
pressure, the ocean is usually super-saturated at
shallow to intermediate depths and undersaturated in the deep basins.
Another important effect, which has been
outlined in Chapter 6.2, is the ageing, or better the
respiratory enrichment in CO 2 of the deep waters
along the flow path from the North Atlantic to the
North Pacific causing large differences in saturation state above the seafloor in the Atlantic and
the Pacific (Fig. 9.6).
While calcite dissolution below saturation
horizons in the water column can be described by
the equations reported in section 9.3.2., dissolution due to CO 2 release from organic carbon
respiration is expressed by the following equations:
9.4
Carbonate Reservoir Sizes and Fluxes Between Particulate and Dissolved Reservoirs
Fig. 9.6 ∆CO 3
2- (giving the difference between the saturation concentration of CO 3
2- and the ambient concentration of
CO 3
2- at a specific depth) based on GEOSECS-data (Takahashi
et al., 1980) of two locations in the Atlantic and the Pacific
ocean. The intersection of the data points with the 0 µM line
denotes the position of the lysocline (from Archer, 1996).
values of alkalinity and total dissolved inorganic
carbon in intermediate and deep waters are the result of dissolving calcite particles settling through
the water column, or of sedimentary dissolution
above the lysocline, would require better knowledge of the total amount of calcite production in
the surface ocean and of the dissolution rates in
the sediment. While the first is very difficult to
measure, the second may be reached by improving
data sets of inorganic carbon flux rates from the
sediment to ocean at intermediate and deep water
levels. The next Section 9.4.2. describes the approach of estimating the fluxes of total carbon and
inorganic carbon resulting from calcite dissolution
in more detail.
9.4.2
Inorganic and Organic Carbon
Release from Deep Sea Sediments
Although only a small fraction of carbon arriving
at the seafloor is finally buried, over geological
time scales deep sea sediments have formed the
largest reservoir of carbon on earth mainly
consisting of biogenic carbonates (~50 Mio. Gt)
and organic detritus (~12 Mio. Gt; i.e. de Baar &
Suess, 1993). This huge storage potential of
marine carbonates is a major factor for maintaining
reasonably low atmospheric CO 2 -levels in earth
history. However, the ratio of accumulation versus
recycling of carbon at the seafloor reflects a
dynamic equilibrium depending on various external forcing factors. Hence, the question for the
“how much is being recycled” is an important
issue, but quite difficult to explore.
Diagenetic mineralization and dissolution
processes in deep sea sediments have been
recognized as important control factors and
numerous efforts have been made to estimate the
contribution by benthic reflux to the ocean
budget. The diagenetic reactions are most intense
at the sediment-water interface where the most
labile components become rapidly mineralized.
Hence, this is the place where the determination
between burial and recycling is made. The driving
forces of carbon release from the sediments are
the degradation of organic matter and the
dissolution of calcium carbonate, which are in turn
dependent on the supply by the remnants of
biological production in the surface waters.
However, the dissolution of calcium carbonate in
deep-sea sediments is controlled by two major
factors, the degree of undersaturation of the deep
ocean waters with respect to calcite and aragonite,
and the reaction with carbon dioxide from
respiration processes. Two of the key parameters
are thus, the “rain ratio” (C CaCO3 /C POC ) of the
sinking material and the water depth of final
deposition, above or below the CO 3
2saturation
horizon (lysocline). Because the solubility of
calcium carbonate increases with increasing
pressure, the ocean is usually super-saturated at
shallow to intermediate depths and undersaturated in the deep basins.
Another important effect, which has been
outlined in Chapter 6.2, is the ageing, or better the
respiratory enrichment in CO 2 of the deep waters
along the flow path from the North Atlantic to the
North Pacific causing large differences in saturation state above the seafloor in the Atlantic and
the Pacific (Fig. 9.6).
While calcite dissolution below saturation
horizons in the water column can be described by
the equations reported in section 9.3.2., dissolution due to CO 2 release from organic carbon
respiration is expressed by the following equations:
9.4
Carbonate Reservoir Sizes and Fluxes Between Particulate and Dissolved Reservoirs
Fig. 9.6 ∆CO 3
2- (giving the difference between the saturation concentration of CO 3
2- and the ambient concentration of
CO 3
2- at a specific depth) based on GEOSECS-data (Takahashi
et al., 1980) of two locations in the Atlantic and the Pacific
ocean. The intersection of the data points with the 0 µM line
denotes the position of the lysocline (from Archer, 1996).