Oceanic carbon cycle
- At the surface of the oceans towards the poles, seawater becomes cooler and more carbonic acid is formed as CO2 becomes more soluble. This is coupled to the ocean's thermohaline circulation which transports dense surface water into the ocean's interior (see the entry on the solubility pump).
- Although the deep ocean contains much more dissolved inorganic carbon than the surface ocean, the concentration is only 15% higher in the deep ocean as compared to the surface ocean due to the higher volume of the deep ocean.
- In upper ocean areas of high biological productivity, organisms convert reduced carbon to tissues, or carbonates to hard body parts such as shells and tests. These are, respectively, oxidized (soft-tissue pump) and redissolved (carbonate pump) at lower average levels of the ocean than those at which they formed, resulting in a downward flow of carbon (see entry on the biological pump).
- The flux or absorption of carbon dioxide into the world's oceans is influenced by the presence of widespread viruses within ocean water, that infect many species of bacteria. The resulting bacterial deaths spawn a sequence of events that lead to greatly enlarged respiration of carbon dioxide, enhancing the role of the oceans as a carbon sink.
- The balance of dissolved inorganic carbon (DIC) : dissolved organic carbon (DOC) : particle organic carbon is about 2000:38:1.
- The CaCO3 counter pump increases the partial pressure of CO2 in the ocean, thus leading to higher outgasing of carbon dioxide.
- Higher ocean temperatures lead to stronger layering, thus less mixing and less capacity for thermohaline circulation to bring carbon into lower ocean layers.
- The weathering of silicate rock (see carbonate-silicate cycle). Carbonic acid reacts with weathered rock to produce bicarbonate ions. The bicarbonate ions produced are carried to the ocean, where they are used to make marine carbonates. Unlike dissolved CO2 in equilibrium or tissues which decay weathering does not move the carbon into a reservoir from which it can readily return to the atmosphere.
- Much remains to be learned about the cycling of carbon in the deep ocean. For example, a recent discovery is that larvacean mucus houses (commonly known as "sinkers") are created in such large numbers that they can deliver as much carbon to the deep ocean as has been previously detected by sediment traps. Because of their size and composition, these houses are rarely collected in such traps, so most biogeochemical analyses have erroneously ignored them.
- The amount of dissolved inorganic carbon in the ocean is significantly higher in the deep layer (below 300 m depth). This is caused by the solubility pump and the biological pump.
Movement of carbon in the ocean
CO2 is absorbed from the atmosphere at the ocean's surface and converted into dissolved inorganic carbon (DIC). It is then converted in gross primary production (GPP) by phytoplankton into organic carbon. About half of the GPP is autotrophically respirated and converted back into DIC. The rest stays in the form of net primary production (NPP). Some of the organic carbon sinks into the lower ocean levels as detritus or calcium carbonate in shells. Some soft tissue is converted into particulate organic carbon or dissolved organic carbon and, from these forms, into dissolved inorganic carbon. The rest sinks to the ocean floor. Shells out of calcium carbonate are also deposited on the ocean floor as sediment, whereas the carbon can dissolve and reach the lower ocean levels again. Thermohaline circulation can bring carbon in the deep ocean levels to the upper levels, where it can again be exchanged with the atmosphere. Units are in gigatons.
In the oceans, dissolved carbonate can combine with dissolved calcium to precipitate solid calcium carbonate, CaCO3, mostly as the shells of microscopic organisms. When these organisms die, their shells sink and accumulate on the ocean floor. Over time these carbonate sediments form limestone which is the largest reservoir of carbon in the carbon cycle. The dissolved calcium in the oceans comes from the chemical weathering of calcium-silicate rocks, during which carbonic and other acids in groundwater react with calcium-bearing minerals liberating calcium ions to solution and leaving behind a residue of newly formed aluminium-rich clay minerals and insoluble minerals such as quartz.
Carbon exchange between the ocean and other systems
Carbon is readily exchanged between the atmosphere and ocean. Each year, approximately 90 gigatons of carbon are exchanged between the atmosphere and the ocean in each direction, leading to a quick equilibration of surface ocean and atmospheric carbon levels. In regions of oceanic upwelling, carbon is released to the atmosphere. Conversely, regions of downwelling transfer carbon (CO2) from the atmosphere to the ocean. Interactions with the atmosphere also influence the rate of carbon uptake from other systems. Extreme storms such as hurricanes and typhoons bury extremely large amounts of carbon, because they wash away so much sediment.
Inorganic carbon, that is carbon compounds with no carbon-carbon or carbon-hydrogen bonds, is important in its reactions within water. This carbon exchange becomes important in controlling pH in the ocean and can also vary as a source or sink for carbon. When CO2 enters the ocean, it participates in a series of reactions which are locally in equilibrium:
- CO2 (atmospheric) ⇌ CO2 (dissolved)
Conversion to carbonic acid:
- CO2 (dissolved) + H2O ⇌ H2CO3
- H2CO3 ⇌ H+ + HCO3−(bicarbonate ion)
- HCO3− ⇌ H+ + CO32−(carbonate ion)
This set of reactions, which of each has its own equilibrium coefficient, determines the form that inorganic carbon takes in the oceans. The coefficients, which have been determined empirically for ocean water, are themselves functions of temperature, pressure, and the presence of other ions (especially borate). In the ocean the equilibria strongly favor bicarbonate. Since this ion is three steps removed from atmospheric CO2, the level of inorganic carbon storage in the ocean does not have a proportion of unity to the atmospheric partial pressure of CO2. The factor for the ocean is about ten: that is, for a 10% increase in atmospheric CO2, oceanic storage (in equilibrium) increases by about 1%, with the exact factor dependent on local conditions. This buffer factor is often called the Revelle factor, after Roger Revelle.
The ocean's buffer capacity is limited and depends on the relatively slow weathering of rocks, which takes place much slower than current CO2 emissions into the atmosphere. Current emission rates exceed the supply of the necessary mineral cations for these reactions, meaning that on a time scale of millennia, the ocean's ability to absorb CO2 will decrease. A major contributing factor to this decrease in capacity is the ocean's falling pH levels. The absorption of carbon dioxide makes the ocean more acidic and given the projected rates of increases in atmospheric CO2, ocean surface water pH will decrease by about 0.2 by 2100. This will additionally slow the biological precipitation of calcium carbonates.
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