Carbon Dioxide Interactions in Air Sea Water System

The interaction between carbon dioxide in the atmosphere and the hydrosphere is the principal factor for understanding large carbon biogeochemical cycles. As it has been mentioned above, the gases of the troposphere and the surface layer of the ocean persist in a state of kinetic equilibrium.

Compared with the atmosphere, where most carbon is presented by CO2, oceanic carbon is mainly present in four forms: dissolved inorganic carbon (DIC), dissolved organic carbon (DOC), particulate organic carbon (POC), and the marine biota itself.

DIC concentrations have been monitored extensively since the appearance of precise analytical techniques. When CO2 dissolves in water it may hydrate to form H2CO3(aq), which, in turn, dissociates to HCO- and CO2-. This process depends on pH and specification is shown in Figure 5.

The conjugate pairs responsible for most of the pH buffer capacity in marine water are HCO-/CO3- and B(OH)3/B(OH)4-. Although the predominance of HCO-at the oceanic pH of 8.2 actually places the carbonate system close to a pH buffer minimum, its importance is maintained by the high DIC concentration (~2 mm). Ocean water in contact with the atmosphere will, if the air-sea gas exchange rate is short compared to the mixing time with deeper water, reach equilibrium according to Henry's law. At the pH of oceanic water around 8.2, most of the DIC is in the form of HCO3- and CO3 - with a very small proportion of H2CO3. Although H2CO3 changes in proportion to CO2 (g), the ionic form changes little as a result of various acid-base equilibrium.

Figure 5 Distribution of dissolved carbon species in seawater as a function of pH at 15 °C and a salinity of 35. Average oceanic pH is about 8.2.

Figure 5 Distribution of dissolved carbon species in seawater as a function of pH at 15 °C and a salinity of 35. Average oceanic pH is about 8.2.

From chemical aqueous carbon specification, the alkalinity, Alk, representing the acid-neutralizing capacity of the solution, is given by the following equation:

Alk = [OH"] - [H+]+ [B(OH)4- ] + [B(OH)3] + 2[CO3 - ]

Average DIC and Alk concentrations for the world's oceans are shown in Figure 6.

With an average DIC of 2.35 mmol kg" seawater and the world oceanic volume of 1370 x 106km3, the DIC carbon reservoir is estimated to be 37 900 x 1091 C. The surface waters of the world's oceans contain a minor part of DIC, ^700 x 109tC. However, these waters play an important role in air-deep water exchange (see above).

Oceanic surface water is supersaturated everywhere with respect to the two solid calcium carbonate species calcite and aragonite. Nevertheless, calcium precipitation is exclusively controlled by biological processes, specifically the formation of hard parts (shells, skeletal parts, etc.). The very few existing amounts ofspontaneous inorganic precipitation of CaCO3(s) come from the Bahamas region of the Caribbean.

The detrital rain of carbon-containing particles can be divided into two groups: the hard parts comprised of

Alkalinity (|iEq kg-1)

Alkalinity (|iEq kg-1)

DIC (|imol kg-1)
Figure 6 The vertical distribution of alkalinity (a) and dissolved inorganic carbon in the world's ocean (b). Ocean regions are shown as NA (North Atlantic), SA (South Atlantic), AA (Antarctic), SI (South Indian), NI (North Indian), SP (South Pacific), and NP (North Pacific).

calcite and aragonite, and the soft tissue containing organic carbon. The composition of the soft tissue shows the average ratio of biophils as P:N:C:Ca:S = 1:15:131:26:50, with Cc:Co ratio as 1:4.

The estimation of Cc and Co mass annually eliminated from the biogeochemical cycles in ocean is a very uncertain task (see above). The carbonate-hydrocarbonate system includes the precipitation of calcium carbonate as a deposit:

Atmosphere CO2

Surface ocean layer H2O $ H2CO3 $ H+ + HCO3- $ H+ + CO2- + Ca2+

Deep ocean water CaCO3

The binding ofcarbon into carbonates is related to the activity of living organisms. However, the surface runoff of Ca2+ ions from the land determines the formation of carbonate deposits to a significant degree. The Ca2+ ion stream is roughly 0.53 x 109tyr-1, which can provide for a CaCO3 precipitation rate of 1.33 x 109tyr-1. This would correspond to the loss of 0.57 x 10 tCO2, or 0.16 x 10 tC from the carbonate-hydrocarbonate system.

The surface runoff from the world's land plays an important role in the global carbon mass exchange. The

Table 4 Fluxes of carbon in the biosphere



continental runoff supply of HCO3 is 2.4 :

109 tyr~

, that is, 0.47 x 109 tyr-1 for carbon. Besides, stream waters contain dissolved organic matter at 6.9 mgl-1, which make up to an annual loss of 0.28 x 109tyr-1. The average carbon concentration ofsuspended insoluble organic matter in the stream discharge is 5 mgl-1, which gives the loss of about 0.2 x 10 tyr- . Most of this mass fails to reach the open ocean and becomes deposited in the shelf and the estuar-ine delta of rivers. One can see that equal amounts of Cc and Co (0.5 x 109t for each) are annually lost from the world's land surface.

The formation of carbonates and the accumulation of organic matter are not confined solely to oceans; these processes occur also on land. The mass of carbonates annually produced in the soils of arid landscapes appears to be high enough.

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