Oceanic carbon cycle
The oceanic carbon cycle (or marine carbon cycle) is composed of processes that exchange carbon between various pools within the ocean as well as between the atmosphere, Earth interior, and the seafloor. The carbon cycle is a result of many interacting forces across multiple time and space scales that circulates carbon around the planet, ensuring that carbon is available globally. The Oceanic carbon cycle is a central process to the global carbon cycle and contains both inorganic carbon (carbon not associated with a living thing, such as carbon dioxide) and organic carbon (carbon that is, or has been, incorporated into a living thing). Part of the marine carbon cycle transforms carbon between non-living and living matter.
Three main processes (or pumps) that make up the marine carbon cycle bring atmospheric
Earth's plants and algae (
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Marine carbon
Carbon compounds can be distinguished as either organic or inorganic, and dissolved or particulate, depending on their composition. Organic carbon forms the backbone of key component of organic compounds such as – proteins, lipids, carbohydrates, and nucleic acids. Inorganic carbon is found primarily in simple compounds such as carbon dioxide, carbonic acid, bicarbonate, and carbonate (CO2, H2CO3, HCO3−, CO32− respectively).
Marine carbon is further separated into particulate and dissolved phases. These pools are operationally defined by physical separation – dissolved carbon passes through a 0.2 μm filter, and particulate carbon does not.
Inorganic carbon
There are two main types of inorganic carbon that are found in the oceans. Dissolved inorganic carbon (DIC) is made up of bicarbonate (HCO3−), carbonate (CO32−) and carbon dioxide (including both dissolved CO2 and carbonic acid H2CO3). DIC can be converted to particulate inorganic carbon (PIC) through precipitation of CaCO3 (biologically or abiotically). DIC can also be converted to particulate organic carbon (POC) through photosynthesis and chemoautotrophy (i.e. primary production). DIC increases with depth as organic carbon particles sink and are respired. Free oxygen decreases as DIC increases because oxygen is consumed during aerobic respiration.
Particulate inorganic carbon (PIC) is the other form of inorganic carbon found in the ocean. Most PIC is the CaCO3 that makes up shells of various marine organisms, but can also form in whiting events. Marine fish also excrete calcium carbonate during osmoregulation.[14]
Some of the inorganic carbon species in the ocean, such as bicarbonate and carbonate, are major contributors to alkalinity, a natural ocean buffer that prevents drastic changes in acidity (or pH). The marine carbon cycle also affects the reaction and dissolution rates of some chemical compounds, regulates the amount of carbon dioxide in the atmosphere and Earth's temperature.[15]
Organic carbon
Like inorganic carbon, there are two main forms of organic carbon found in the ocean (dissolved and particulate). Dissolved organic carbon (DOC) is defined operationally as any organic molecule that can pass through a 0.2 µm filter. DOC can be converted into particulate organic carbon through heterotrophy and it can also be converted back to dissolved inorganic carbon (DIC) through respiration.
Those organic carbon molecules being captured on a filter are defined as particulate organic carbon (POC). POC is composed of organisms (dead or alive), their fecal matter, and detritus. POC can be converted to DOC through disaggregation of molecules and by exudation by phytoplankton, for example. POC is generally converted to DIC through heterotrophy and respiration.
Marine carbon pumps
Solubility pump
Full article: Solubility pump
The oceans store the largest pool of reactive carbon on the planet as DIC, which is introduced as a result of the dissolution of atmospheric carbon dioxide into seawater – the solubility pump.[15] Aqueous CO2, carbonic acid, bicarbonate ion, and carbonate ion concentrations comprise dissolved inorganic carbon (DIC). DIC circulates throughout the whole ocean by Thermohaline circulation, which facilitates the tremendous DIC storage capacity of the ocean.[16] The chemical equations below show the reactions that CO2 undergoes after it enters the ocean and transforms into its aqueous form.
-
(1)
Carbonic acid rapidly dissociates into free hydrogen ion (technically, hydronium) and bicarbonate.
-
(2)
The free hydrogen ion meets carbonate, already present in the water from the dissolution of CaCO3, and reacts to form more bicarbonate ion.
-
(3)
The dissolved species in the equations above, mostly bicarbonate, make up the carbonate alkalinity system, the dominant contributor to seawater alkalinity.[9]
Carbonate pump
The carbonate pump, sometimes called the carbonate counter pump, starts with marine organisms at the ocean's surface producing particulate inorganic carbon (PIC) in the form of calcium carbonate (calcite or aragonite, CaCO3). This CaCO3 is what forms hard body parts like shells.[15] The formation of these shells increases atmospheric CO2 due to the production of CaCO3[9] in the following reaction with simplified stoichiometry:[17]
-
(4)
Coccolithophores, a nearly ubiquitous group of phytoplankton that produce shells of calcium carbonate, are the dominant contributors to the carbonate pump.[15] Due to their abundance, coccolithophores have significant implications on carbonate chemistry, in the surface waters they inhabit and in the ocean below: they provide a large mechanism for the downward transport of CaCO3.[19] The air-sea CO2 flux induced by a marine biological community can be determined by the rain ratio - the proportion of carbon from calcium carbonate compared to that from organic carbon in particulate matter sinking to the ocean floor, (PIC/POC).[18] The carbonate pump acts as a negative feedback on CO2 taken into the ocean by the solubility pump. It occurs with lesser magnitude than the solubility pump.
Biological pump
Full article: Biological pump
Particulate organic carbon, created through biological production, can be exported from the upper ocean in a flux commonly termed the biological pump, or respired (equation 6) back into inorganic carbon. In the former, dissolved inorganic carbon is biologically converted into organic matter by photosynthesis (equation 5) and other forms of
-
(5)
-
(6)
Inputs
Inputs to the marine carbon cycle are numerous, but the primary contributions, on a net basis, come from the atmosphere and rivers.[1] Hydrothermal vents generally supply carbon equal to the amount they consume.[15]
Atmosphere
Before the
A 2020 study found significantly higher net flux of carbon into the oceans compared to previous studies. The new study used satellite data to account for small temperature differences between the surface of the ocean and the depth of a few meters where the measurements are made.[35][36]
Carbon dioxide exchange rates between ocean and atmosphere
Ocean-atmospheric exchanges rates of CO2 depend on the concentration of carbon dioxide already present in both the atmosphere and the ocean, temperature, salinity, and wind speed.[37] This exchange rate can be approximated by Henry's law and can be calculated as S = kP, where the solubility (S) of the carbon dioxide gas is proportional to the amount of gas in the atmosphere, or its partial pressure.[1]
Revelle factor
Since the oceanic intake of carbon dioxide is limited, CO2 influx can also be described by the Revelle factor.[32][9] The Revelle Factor is a ratio of the change of carbon dioxide to the change in dissolved inorganic carbon, which serves as an indicator of carbon dioxide dissolution in the mixed layer considering the solubility pump. The Revelle Factor is an expression to characterize the thermodynamic efficiency of the DIC pool to absorb CO2 into bicarbonate. The lower the Revelle factor, the higher the capacity for ocean water to take in carbon dioxide. While Revelle calculated a factor of around 10 in his day, in a 2004 study data showed a Revelle factor ranging from approximately 9 in low-latitude tropical regions to 15 in the southern ocean near Antarctica.[38]
Rivers
Rivers can also transport organic carbon to the ocean through weathering or erosion of aluminosilicate (equation 7) and carbonate rocks (equation 8) on land,
-
(7)
-
(8)
or by the decomposition of life (equation 5, e.g. plant and soil material).[1] Rivers contribute roughly equal amounts (~0.4 GtC/yr) of DIC and DOC to the oceans.[1] It is estimated that approximately 0.8 GtC (DIC + DOC) is transported annually from the rivers to the ocean.[1] The rivers that flow into Chesapeake Bay (Susquehanna, Potomac, and James rivers) input approximately 0.004 Gt (6.5 x 1010 moles) DIC per year.[39] The total carbon transport of rivers represents approximately 0.02% of the total carbon in the atmosphere.[40] Though it seems small, over long time scales (1000 to 10,000 years) the carbon that enters rivers (and therefore does not enter the atmosphere) serves as a stabilizing feedback for greenhouse warming.[41]
Outputs
The key outputs of the marine carbon system are particulate organic matter (POC) and calcium carbonate (PIC) preservation as well as reverse weathering.[1] While there are regions with local loss of CO2 to the atmosphere and hydrothermal processes, a net loss in the cycle does not occur.[15]
Organic matter preservation
Sedimentation is a long-term sink for carbon in the ocean, as well as the largest loss of carbon from the oceanic system.
Fate of sinking organic carbon
Historically, sediments with the highest organic carbon contents were frequently found in areas with high surface water productivity or those with low bottom-water oxygen concentrations.[44] 90% of organic carbon burial occurs in deposits of deltas and continental shelves and upper slopes;[45] this is due partly to short exposure time because of a shorter distance to the seafloor and the composition of the organic matter that is already deposited in those environments.[46] Organic carbon burial is also sensitive to climate patterns: the accumulation rate of organic carbon was 50% larger during the glacial maximum compared to interglacials.[47]
Degradation
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POC is decomposed by a series of microbe-driven processes, such as methanogenesis and sulfate reduction, before burial in the seafloor.[48][49] Degradation of POC also results in microbial methane production which is the main gas hydrate on the continental margins.[50] Lignin and pollen are inherently resistant to degradation, and some studies show that inorganic matrices may also protect organic matter.[51] Preservation rates of organic matter depend on other interdependent variables that vary nonlinearly in time and space.[52] Although organic matter breakdown occurs rapidly in the presence of oxygen, microbes utilizing a variety of chemical species (via redox gradients) can degrade organic matter in anoxic sediments.[52] The burial depth at which degradation halts depends upon the sedimentation rate, the relative abundance of organic matter in the sediment, the type of organic matter being buried, and innumerable other variables.[52] While decomposition of organic matter can occur in anoxic sediments when bacteria use oxidants other than oxygen (nitrate, sulfate, Fe3+), decomposition tends to end short of complete mineralization.[53] This occurs because of preferential decomposition of labile molecules over refractile molecules.[53]
Burial
Organic carbon burial is an input of energy for underground biological environments and can regulate oxygen in the atmosphere at long time-scales (> 10,000 years).[47] Burial can only take place if organic carbon arrives to the sea floor, making continental shelves and coastal margins the main storage of organic carbon from terrestrial and oceanic primary production. Fjords, or cliffs created by glacial erosion, have also been identified as areas of significant carbon burial, with rates one hundred times greater than the ocean average.[54] Particulate organic carbon is buried in oceanic sediments, creating a pathway between a rapidly available carbon pool in the ocean to its storage for geological timescales. Once carbon is sequestered in the seafloor, it is considered blue carbon. Burial rates can be calculated as the difference between the rate at which organic matter sinks and the rate at which it decomposes.
Calcium carbonate preservation
The precipitation of calcium carbonate is important as it results in a loss of alkalinity as well as a release of CO2 (Equation 4), and therefore a change in the rate of preservation of calcium carbonate can alter the partial pressure of CO2 in Earth's atmosphere.[15] CaCO3 is supersatured in the great majority of ocean surface waters and undersaturated at depth,[9] meaning the shells are more likely to dissolve as they sink to ocean depths. CaCO3 can also be dissolved through metabolic dissolution (i.e. can be used as food and excreted) and thus deep ocean sediments have very little calcium carbonate.[15] The precipitation and burial of calcium carbonate in the ocean removes particulate inorganic carbon from the ocean and ultimately forms limestone.[15] On time scales greater than 500,000 years Earth's climate is moderated by the flux of carbon in and out of the lithosphere.[55] Rocks formed in the ocean seafloor are recycled through plate tectonics back to the surface and weathered or subducted into the mantle, the carbon outgassed by volcanoes.[1]
Human impacts
Oceans take up 15 – 40% of anthropogenic CO2,[56][57] and so far roughly 40% of the carbon from fossil fuel combustion has been taken up into the oceans.[58] Because the Revelle factor increases with increasing CO2, a smaller fraction of the anthropogenic flux will be taken up by the ocean in the future.[59] Current annual increase in atmospheric CO2 is approximately 4 gigatons of carbon.[60] This induces climate change that drives carbon concentration and carbon-climate feedback processes that modifies ocean circulation and the physical and chemical properties of seawater, which alters CO2 uptake.[61][62] Overfishing and the plastic pollution of the oceans contribute to the degraded state of the world's biggest carbon sink.[63][64]
Ocean acidification
Full article: Ocean acidification
The pH of the oceans is declining due to uptake of atmospheric CO2.[65] The rise in dissolved carbon dioxide reduces the availability of the carbonate ion, reducing CaCO3 saturation state, thus making it thermodynamically harder to make CaCO3 shell.[66] Carbonate ions preferentially bind to hydrogen ions to form bicarbonate,[9] thus a reduction in carbonate ion availability increases the amount of unbound hydrogen ions, and decreases the amount of bicarbonate formed (Equations 1–3). pH is a measurement of hydrogen ion concentration, where a low pH means there are more unbound hydrogen ions. pH is therefore an indicator of carbonate speciation (the format of carbon present) in the oceans and can be used to assess how healthy the ocean is.[66]
The list of organisms that may struggle due to ocean acidification include
Iron fertilization
Full article: Iron Fertilization
Iron fertilization is a facet of geoengineering, which purposefully manipulates the Earth's climate system, typically in aspects of the carbon cycle or radiative forcing. Of current geoengineering interest is the possibility of accelerating the biological pump to increase export of carbon from the surface ocean. This increased export could theoretically remove excess carbon dioxide from the atmosphere for storage in the deep ocean. Ongoing investigations regarding artificial fertilization exist.[68] Due to the scale of the ocean and the fast response times of heterotrophic communities to increases in primary production, it is difficult to determine whether limiting-nutrient fertilization results in an increase in carbon export.[68] However, the majority of the community does not believe this is a reasonable or viable approach.[69]
Dams and reservoirs
There are over 16 million
See also
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