The biological pump, also known as the marine carbon pump, is, in its simplest form, the ocean's biologically driven sequestration of carbon from the atmosphere and land runoff to the ocean interior and seafloor sediments.[1] It is the part of the oceanic carbon cycle responsible for the cycling of organic matter formed mainly by phytoplankton during photosynthesis (soft-tissue pump), as well as the cycling of calcium carbonate (CaCO3) formed into shells by certain organisms such as plankton and mollusks (carbonate pump).[2]
Budget calculations of the biological carbon pump are based on the ratio between sedimentation (carbon export to the ocean floor) and remineralization (release of carbon to the atmosphere).
The biological pump is not so much the result of a single process, but rather the sum of a number of processes each of which can influence biological pumping. Overall, the pump transfers about 11 gigatonnes of carbon every year into the ocean's interior. This takes carbon out of contact with the atmosphere for several thousand years or longer. An ocean without a biological pump would result in atmospheric carbon dioxide levels about 400 ppm higher than the present day.
Overview
The biological pump depends on the fraction of primary produced organic matter that survives degradation in the euphotic zone and that is exported from surface water to the ocean interior, where it is mineralized to inorganic carbon, with the result that carbon is transported against the gradient of dissolved inorganic carbon (DIC) from the surface to the deep ocean. This transfer occurs through physical mixing and transport of dissolved and particulate organic carbon (POC), vertical migrations of organisms (zooplankton, fish) and through gravitational settling of particulate organic carbon.[4][5]: 526 [6]
The biological pump can be divided into three distinct phases, the first of which is the production of fixed carbon by planktonic phototrophs in the euphotic (sunlit) surface region of the ocean. In these surface waters, phytoplankton use carbon dioxide (CO2), nitrogen (N), phosphorus (P), and other trace elements (barium, iron, zinc, etc.) during photosynthesis to make carbohydrates, lipids, and proteins. Some plankton, (e.g. coccolithophores and foraminifera) combine calcium (Ca) and dissolved carbonates (carbonic acid and bicarbonate) to form a calcium carbonate (CaCO3) protective coating.[7]
Once this carbon is fixed into soft or hard tissue, the organisms either stay in the euphotic zone to be recycled as part of the regenerative nutrient cycle or once they die, continue to the second phase of the biological pump and begin to sink to the ocean floor. The sinking particles will often form aggregates as they sink, greatly increasing the sinking rate. It is this aggregation that gives particles a better chance of escaping predation and decomposition in the water column and eventually making it to the sea floor.[7]
The fixed carbon that is decomposed by bacteria either on the way down or once on the sea floor then enters the final phase of the pump and is remineralized to be used again in primary production. The particles that escape these processes entirely are sequestered in the sediment and may remain there for millions of years. It is this sequestered carbon that is responsible for ultimately lowering atmospheric CO2.[7]
The diagram immediately above illustrates the components of the biological pump. Biology, physics and gravity interact to pump organic carbon into the deep sea. The processes of fixation of inorganic carbon in organic matter during photosynthesis, its transformation by food web processes (trophodynamics), physical mixing, transport and gravitational settling are referred to collectively as the biological pump.[8]
The biological pump is responsible for transforming dissolved inorganic carbon (DIC) into organic biomass and pumping it in particulate or dissolved form into the deep ocean. Inorganic nutrients and carbon dioxide are fixed during photosynthesis by phytoplankton, which both release dissolved organic matter (DOM) and are consumed by herbivorous zooplankton. Larger zooplankton - such as copepods - egest fecal pellets which can be reingested and sink or collect with other organic detritus into larger, more-rapidly-sinking aggregates. DOM is partially consumed by bacteria (black dots) and respired; the remaining refractory DOM is advected and mixed into the deep sea. DOM and aggregates exported into the deep water are consumed and respired, thus returning organic carbon into the enormous deep ocean reservoir of DIC. About 1% of the particles leaving the surface ocean reach the seabed and are consumed, respired, or buried in the sediments. There, carbon is stored for millions of years. The net effect of these processes is to remove carbon in organic form from the surface and return it to DIC at greater depths, maintaining the surface-to-deep ocean gradient of DIC. Thermohaline circulation returns deep-ocean DIC to the atmosphere on millennial timescales.[8]
Primary production
Adapted from Simon et al., 2002.[10]
The first step in the biological pump is the synthesis of both organic and inorganic carbon compounds by phytoplankton in the uppermost, sunlit layers of the ocean.[11] Organic compounds in the form of sugars, carbohydrates, lipids, and proteins are synthesized during the process of photosynthesis:
CO2 + H2O + light → CH2O + O2
In addition to carbon, organic matter found in phytoplankton is composed of nitrogen, phosphorus and various trace metals. The ratio of carbon to nitrogen and phosphorus varies from place to place,[12] but has an average ratio near 106C:16N:1P, known as the Redfield ratio. Trace metals such as magnesium, cadmium, iron, calcium, barium and copper are orders of magnitude less prevalent in phytoplankton organic material, but necessary for certain metabolic processes and therefore can be limiting nutrients in photosynthesis due to their lower abundance in the water column.[7]
Oceanic primary production accounts for about half of the carbon fixation carried out on Earth. Approximately 50–60 Pg of carbon are fixed by marine phytoplankton each year despite the fact that they comprise less than 1% of the total photosynthetic biomass on Earth. The majority of this carbon fixation (~80%) is carried out in the open ocean while the remaining amount occurs in the very productive upwelling regions of the ocean. Despite these productive regions producing 2 to 3 times as much fixed carbon per area, the open ocean accounts for greater than 90% of the ocean area and therefore is the larger contributor.[7]
Forms of carbon
- Dissolved and particulate carbon
Phytoplankton supports all life in the ocean as it converts inorganic compounds into organic constituents. This autotrophically produced biomass presents the foundation of the marine food web.[13] In the diagram immediately below, the arrows indicate the various production (arrowhead pointing toward DOM pool) and removal processes of DOM (arrowhead pointing away), while the dashed arrows represent dominant biological processes involved in the transfer of DOM. Due to these processes, the fraction of labile DOM decreases rapidly with depth, whereas the refractory character of the DOM pool considerably increases during its export to the deep ocean. DOM, dissolved organic matter.[13][14]
Calcium carbonate
of the plates of buried coccolithophores ( see below ↓ )
Particulate inorganic carbon (PIC) usually takes the form of calcium carbonate (CaCO3), and plays a key part in the ocean carbon cycle.[16] This biologically fixed carbon is used as a protective coating for many planktonic species (coccolithophores, foraminifera) as well as larger marine organisms (mollusk shells). Calcium carbonate is also excreted at high rates during osmoregulation by fish, and can form in whiting events.[17] While this form of carbon is not directly taken from the atmospheric budget, it is formed from dissolved forms of carbonate which are in equilibrium with CO2 and then responsible for removing this carbon via sequestration.[18]
CO2 + H2O → H2CO3 → H+ + HCO3−
Ca2+ + 2HCO3− → CaCO3 + CO2 + H2O
While this process does manage to fix a large amount of carbon, two units of alkalinity are sequestered for every unit of sequestered carbon.[2][19] The formation and sinking of CaCO3 therefore drives a surface to deep alkalinity gradient which serves to raise the pH of surface waters, shifting the speciation of dissolved carbon to raise the partial pressure of dissolved CO2 in surface waters, which actually raises atmospheric levels. In addition, the burial of CaCO3 in sediments serves to lower overall oceanic alkalinity, tending to raise pH and thereby atmospheric CO2 levels if not counterbalanced by the new input of alkalinity from weathering.[1] The portion of carbon that is permanently buried at the sea floor becomes part of the geologic record. Calcium carbonate often forms remarkable deposits that can then be raised onto land through tectonic motion as in the case with the White Cliffs of Dover in Southern England. These cliffs are made almost entirely of the plates of buried coccolithophores.[20]
Oceanic carbon cycle
Three main processes (or pumps) that make up the marine carbon cycle bring atmospheric carbon dioxide (CO2) into the ocean interior and distribute it through the oceans. These three pumps are: (1) the solubility pump, (2) the carbonate pump, and (3) the biological pump. The total active pool of carbon at the Earth's surface for durations of less than 10,000 years is roughly 40,000 gigatons C (Gt C, a gigaton is one billion tons, or the weight of approximately 6 million blue whales), and about 95% (~38,000 Gt C) is stored in the ocean, mostly as dissolved inorganic carbon.[21][22] The speciation of dissolved inorganic carbon in the marine carbon cycle is a primary controller of acid-base chemistry in the oceans.
Solubility pump
The biological pump is accompanied by a physico-chemical counterpart known as the solubility pump. This pump transports significant amounts of carbon in the form of dissolved inorganic carbon (DIC) from the ocean's surface to its interior. It involves physical and chemical processes only, and does not involve biological processes.[23]
The solubility pump is driven by the coincidence of two processes in the ocean:
- The solubility of carbon dioxide is a strong inverse function of seawater temperature (i.e. solubility is greater in cooler water)
- The thermohaline circulation is driven by the formation of deep water at high latitudes where seawater is usually cooler and denser
Since deep water (that is, seawater in the ocean's interior) is formed under the same surface conditions that promote carbon dioxide solubility, it contains a higher concentration of dissolved inorganic carbon than might be expected from average surface concentrations. Consequently, these two processes act together to pump carbon from the atmosphere into the ocean's interior. One consequence of this is that when deep water upwells in warmer, equatorial latitudes, it strongly outgasses carbon dioxide to the atmosphere because of the reduced solubility of the gas.[24]
Carbonate pump
The carbonate pump is sometimes referred to as the “hard tissue” component of the biological pump.[25] Some surface marine organisms, like coccolithophores, produce hard structures out of calcium carbonate, a form of particulate inorganic carbon, by fixing bicarbonate.[26] This fixation of DIC is an important part of the oceanic carbon cycle.
Ca2+ + 2 HCO3− → CaCO3 + CO2 + H2O
While the biological carbon pump fixes inorganic carbon (CO2) into particulate organic carbon in the form of sugar (C6H12O6), the carbonate pump fixes inorganic bicarbonate and causes a net release of CO2.[26] In this way, the carbonate pump could be termed the carbonate counter pump. It works counter to the biological pump by counteracting the CO2 flux into the biological pump.[27]
Continental shelf pump
The continental shelf pump is proposed as operating in the shallow waters of the continental shelves as a mechanism transporting carbon (dissolved or particulate) from the continental waters to the interior of the adjacent deep ocean.[28] As originally formulated, the pump is thought to occur where the solubility pump interacts with cooler, and therefore denser water from the shelf floor which feeds down the continental slope into the neighbouring deep ocean.[28] The shallowness of the continental shelf restricts the convection of cooling water, so the cooling can be greater for continental shelf waters than for neighbouring open ocean waters. These cooler waters promote the solubility pump and lead to an increased storage of dissolved inorganic carbon. This extra carbon storage is further augmented by the increased biological production characteristic of shelves.[29] The dense, carbon-rich shelf waters then sink to the shelf floor and enter the sub-surface layer of the open ocean via isopycnal mixing.[28] As the sea level rises in response to global warming, the surface area of the shelf seas will grow and in consequence the strength of the shelf sea pump should increase.[30]