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The biological pump (or ocean carbon biological pump or marine biological carbon pump) is the ocean's biologically driven sequestration of carbon from the atmosphere and land runoff to the ocean interior and seafloor sediments.^[1] In other words, it is a biologically mediated process which results in the sequestering of carbon in the deep ocean away from the atmosphere and the land. The biological pump is the biological component of the "marine carbon pump" which contains both a physical and biological component. It is the part of the broader 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 10.2 gigatonnes of carbon every year into the ocean's interior and a total of 1300 gigatonnes carbon over an average 127 years.^[3] 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 element carbon plays a central role in climate and life on Earth. It is capable of moving among and between the geosphere, cryosphere, atmosphere, biosphere and hydrosphere. This flow of carbon is referred to as the Earth's carbon cycle. It is also intimately linked to the cycling of other elements and compounds. The ocean plays a fundamental role in Earth's carbon cycle, helping to regulate atmospheric CO₂ concentration. The biological pump is a set of processes that transfer organic carbon from the surface to the deep ocean, and is at the heart of the ocean carbon cycle.^[5]

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.^{[6][7]: 526 [8]}

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.^[9]

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, which greatly increases 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.^[9]

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.^[9]

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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.^[10]

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.^[10]

Primary production

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.^[13] 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,^[14] 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.^[9]

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 account for 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.^[9]

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.^[15] 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.^[15][16]

Ocean carbon pools

The marine biological pump depends on a number of key pools, components and processes that influence its functioning. There are four main pools of carbon in the ocean.^[5]

• Dissolved inorganic carbon (DIC) is the largest pool. It constitutes around 38,000 Pg C ^[18] and includes dissolved carbon dioxide (CO₂), bicarbonate (HCO⁻
  ₃), carbonate (CO²⁻
  ₃), and carbonic acid (H₂CO₃). The equilibrium between carbonic acid and carbonate determines the pH of the seawater. Carbon dioxide dissolves easily in water and its solubility is inversely related to temperature. Dissolved CO₂ is taken up in the process of photosynthesis, and can reduce the partial pressure of CO₂ in the seawater, favouring drawdown from the atmosphere. The reverse process respiration, releases CO₂ back into the water, can increase partial pressure of CO₂ in the seawater, favouring release back to the atmosphere. The formation of calcium carbonate by organisms such as coccolithophores has the effect of releasing CO₂ into the water.^{[19][20][21][5]}

• Dissolved organic carbon (DOC) is the next largest pool at around 662 Pg C.^[22] DOC can be classified according to its reactivity as refractory, semi-labile or labile. The labile pool constitutes around 0.2 Pg C, is bioavailable, and has a high production rate (~ 15−25 Pg C y⁻¹).^[23] The refractory component is the biggest pool (~642 Pg C ± 32;^[22] but has a very low turnover rate (0.043 Pg C y⁻¹).^[23] The turnover time for refractory DOC is thought to be greater than 1000 years.^[24][25][5]

• Particulate organic carbon (POC) constitutes around 2.3 Pg C,^[26][27] and is relatively small compared with DIC and DOC. Though small in size, this pool is highly dynamic, having the highest turnover rate of any organic carbon pool on the planet.^[28] Driven by primary production, it produces around 50 Pg C y⁻¹ globally.^[29][30][31] It can be separated into living (e.g. phytoplankton, zooplankton, bacteria) and non-living (e.g. detritus) material. Of these, the phytoplankton carbon is particularly important, because of its role in marine primary production, and also because it serves as the food resource for all the larger organisms in the pelagic ecosystem.^[5]

• Particulate inorganic carbon (PIC) is the smallest of the pools at around 0.03 Pg C.^[32] It is present in the form of calcium carbonate (CaCO₃) in particulate form, and impacts the carbonate system and pH of the seawater. Estimates for PIC production are in the region of 0.8–1.4 Pg C y⁻¹, with at least 65% of it being dissolved in the upper water column, the rest contributing to deep sediments.^[33] Coccolithophores and foraminifera are estimated to be the dominant sources of PIC in the open ocean.^[34][33] The PIC pool is of particular importance due to its role in the ocean carbonate system, and in facilitating the export of carbon to the deep ocean through the carbonate pump, whereby PIC is exported out of the photic zone and deposited in the bottom sediments.^[35][5]

Calcium carbonate

Particulate inorganic carbon (PIC) usually takes the form of calcium carbonate (CaCO₃), and plays a key part in the ocean carbon cycle.^[36] 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.^[37] 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.^[38]

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][39] 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.^[40]

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.^[41][42] 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.^[43]

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.^[44]

Carbonate pump

The carbonate pump is sometimes referred to as the "hard tissue" component of the biological pump.^[45] Some surface marine organisms, like coccolithophores, produce hard structures out of calcium carbonate, a form of particulate inorganic carbon, by fixing bicarbonate.^[46] 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.^[46] 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.^[47]

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.^[48] 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.^[48] 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.^[49] 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.^[48] 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.^[50]

Processes in the biological pump

In the diagram on the right, phytoplankton convert CO₂, which has dissolved from the atmosphere into the surface oceans (90 Gt yr⁻¹), into particulate organic carbon (POC) during primary production (~ 50 Gt C yr⁻¹). Phytoplankton are then consumed by copepods, krill and other small zooplankton grazers, which in turn are preyed upon by higher trophic levels. Any unconsumed phytoplankton form aggregates, and along with zooplankton faecal pellets, sink rapidly and are exported out of the mixed layer (< 12 Gt C yr⁻¹ 14). Krill, copepods, zooplankton and microbes intercept phytoplankton in the surface ocean and sinking detrital particles at depth, consuming and respiring this POC to CO₂ (dissolved inorganic carbon, DIC), such that only a small proportion of surface-produced carbon sinks to the deep ocean (i.e., depths > 1000 m). As krill and smaller zooplankton feed, they also physically fragment particles into small, slower- or non-sinking pieces (via sloppy feeding, coprorhexy if fragmenting faeces),^[52] retarding POC export. This releases dissolved organic carbon (DOC) either directly from cells or indirectly via bacterial solubilisation (yellow circle around DOC). Bacteria can then remineralise the DOC to DIC (CO₂, microbial gardening).^[51]

The biological carbon pump is one of the chief determinants of the vertical distribution of carbon in the oceans and therefore of the surface partial pressure of CO₂ governing air-sea CO₂ exchange.^[53] It comprises phytoplankton cells, their consumers and the bacteria that assimilate their waste and plays a central role in the global carbon cycle by delivering carbon from the atmosphere to the deep sea, where it is concentrated and sequestered for centuries.^[54] Photosynthesis by phytoplankton lowers the partial pressure of CO₂ in the upper ocean, thereby facilitating the absorption of CO₂ from the atmosphere by generating a steeper CO₂ gradient.^[55] It also results in the formation of particulate organic carbon (POC) in the euphotic layer of the epipelagic zone (0–200 m depth). The POC is processed by microbes, zooplankton and their consumers into fecal pellets, organic aggregates ("marine snow") and other forms, which are thereafter exported to the mesopelagic (200–1000 m depth) and bathypelagic zones by sinking and vertical migration by zooplankton and fish.^[56] Although primary production includes both dissolved and particulate organic carbon (DOC and POC respectively), only POC leads to efficient carbon export to the ocean interior, whereas the DOC fraction in surface waters is mostly recycled by bacteria.^[57] However, a more biologically resistant DOC fraction produced in the euphotic zone (accounting for 15–20% of net community productivity), is not immediately mineralized by microbes and accumulates in the ocean surface as biologically semi-labile DOC.^[58] This semi-labile DOC undergoes net export to the deep ocean, thus constituting a dynamic part of the biological carbon pump.^[59] The efficiency of DOC production and export varies across oceanographic regions, being more prominent in the oligotrophic subtropical oceans.^[60] The overall efficiency of the biological carbon pump is mostly controlled by the export of POC.^[57][61]

Marine snow

Most carbon incorporated in organic and inorganic biological matter is formed at the sea surface where it can then start sinking to the ocean floor. The deep ocean gets most of its nutrients from the higher water column when they sink down in the form of marine snow. This is made up of dead or dying animals and microbes, fecal matter, sand and other inorganic material.^[63] A single phytoplankton cell has a sinking rate around one metre per day. Given that the average depth of the ocean is about four kilometres, it can take over ten years for these cells to reach the ocean floor. However, through processes such as coagulation and expulsion in predator fecal pellets, these cells form aggregates. These aggregates, known as marine snow, have sinking rates orders of magnitude greater than individual cells and complete their journey to the deep in a matter of days.^[9]

In the diagram on the right, phytoplankton fix CO₂ in the euphotic zone using solar energy and produce particulate organic carbon (POC). POC formed in the euphotic zone is processed by microbes, zooplankton and their consumers into organic aggregates (marine snow), which is thereafter exported to the mesopelagic (200–1000 m depth) and bathypelagic zones by sinking and vertical migration by zooplankton and fish. Export flux is defined as the sedimentation out of the surface layer (at approximately 100 m depth) and sequestration flux is the sedimentation out of the mesopelagic zone (at approximately 1000 m depth). A portion of the POC is respired back to CO₂ in the oceanic water column at depth, mostly by heterotrophic microbes and zooplankton, thus maintaining a vertical gradient in concentration of dissolved inorganic carbon (DIC). This deep-ocean DIC returns to the atmosphere on millennial timescales through thermohaline circulation. Between 1% and 40% of the primary production is exported out of the euphotic zone, which attenuates exponentially towards the base of the mesopelagic zone and only about 1% of the surface production reaches the sea floor.^[61][64][56]

Of the 50–60 Pg of carbon fixed annually, roughly 10% leaves the surface mixed layer of the oceans, while less than 0.5% of eventually reaches the sea floor.^[9] Most is retained in regenerated production in the euphotic zone and a significant portion is remineralized in midwater processes during particle sinking. The portion of carbon that leaves the surface mixed layer of the ocean is sometimes considered "sequestered", and essentially removed from contact with the atmosphere for many centuries.^[64] However, work also finds that, in regions such as the Southern Ocean, much of this carbon can quickly (within decades) come back into contact with the atmosphere.^[65]

Budget calculations of the biological carbon pump are based on the ratio between sedimentation (carbon export) and remineralization (release to the atmosphere).^[62] It has been estimated that sinking particles export up to 25% of the carbon captured by phytoplankton in the surface ocean to deeper water layers.^[66] About 20% of this export (5% of surface values) is buried in the ocean sediments ^[67] mainly due to their mineral ballast.^[68] During the sinking process, these organic particles are hotspots of microbial activity and represent important loci for organic matter mineralization and nutrient redistribution in the water column.^[69][70][62]

Biomineralization

Ballast minerals

Part of a series related to
Biomineralization
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Observations have shown that fluxes of ballast minerals (calcium carbonate, opal, and lithogenic material) and organic carbon fluxes are closely correlated in the bathypelagic zones of the ocean.^[68] A large fraction of particulate organic matter occurs in the form of marine snow aggregates (>0.5 mm) composed of phytoplankton, detritus, inorganic mineral grains, and fecal pellets in the ocean.^[71] Formation and sinking of these aggregates drive the biological carbon pump via export and sedimentation of organic matter from the surface mixed layer to the deep ocean and sediments. The fraction of organic matter that leaves the upper mixed layer of the ocean is, among other factors, determined by the sinking velocity and microbial remineralisation rate of these aggregates. Recent observations have shown that the fluxes of ballast minerals (calcium carbonate, opal, and lithogenic material) and the organic carbon fluxes are closely correlated in the bathypelagic zones of the ocean. This has led to the hypothesis that organic carbon export is determined by the presence of ballast minerals within settling aggregates.^{[72][73][74][68]}

Mineral ballasting is associated with about 60% of the flux of particulate organic carbon (POC) in the high-latitude North Atlantic, and with about 40% of the flux in the Southern Ocean.^[75] Strong correlations exist also in the deep ocean between the presence of ballast minerals and the flux of POC. This suggests ballast minerals enhance POC flux by increasing the sink rate of ballasted aggregates. Ballast minerals could additionally provide aggregated organic matter some protection from degradation.^[76]

It has been proposed that organic carbon is better preserved in sinking particles due to increased aggregate density and sinking velocity when ballast minerals are present and/or via protection of the organic matter due to quantitative association to ballast minerals.^[72][73][74] In 2002, Klaas and Archer observed that about 83% of the global particulate organic carbon (POC) fluxes were associated with carbonate, and suggested carbonate was a more efficient ballast mineral as compared to opal and terrigenous material. They hypothesized that the higher density of calcium carbonate compared to that of opal and the higher abundance of calcium carbonate relative to terrigenous material might be the reason for the efficient ballasting by calcium carbonate. However, the direct effects of ballast minerals on sinking velocity and degradation rates in sinking aggregates are still unclear.^[74][68]

A 2008 study demonstrated copepod fecal pellets produced on a diet of diatoms or coccolithophorids show higher sinking velocities as compared to pellets produced on a nanoflagellate diet.^[77] Carbon-specific respiration rates in pellets, however, were similar and independent of mineral content. These results suggest differences in mineral composition do not lead to differential protection of POC against microbial degradation, but the enhanced sinking velocities may result in up to 10-fold higher carbon preservation in pellets containing biogenic minerals as compared to that of pellets without biogenic minerals^[77][68]

Minerals seem to enhance the flocculation of phytoplankton aggregates ^[78][79] and may even act as a catalyst in aggregate formation.^[80] However, it has also been shown that incorporation of minerals can cause aggregates to fragment into smaller and denser aggregates.^[81] This can potentially lower the sinking velocity of the aggregated organic material due to the reduced aggregate sizes, and, thus, lower the total export of organic matter. Conversely, if the incorporation of minerals increases the aggregate density, its size-specific sinking velocity may also increase, which could potentially increase the carbon export. Therefore, there is still a need for better quantitative investigations of how the interactions between minerals and organic aggregates affect the degradation and sinking velocity of the aggregates and, hence, carbon sequestration in the ocean.^[81][68]

Remineralisation

Remineralisation refers to the breakdown or transformation of organic matter (those molecules derived from a biological source) into its simplest inorganic forms. These transformations form a crucial link within ecosystems as they are responsible for liberating the energy stored in organic molecules and recycling matter within the system to be reused as nutrients by other organisms.^[7] What fraction does escape remineralisation varies depending on the location. For example, in the North Sea, values of carbon deposition are ~1% of primary production^[82] while that value is <0.5% in the open oceans on average.^[83] Therefore, most of nutrients remain in the water column, recycled by the biota. Heterotrophic organisms will utilize the materials produced by the autotrophic (and chemotrophic) organisms and via respiration will remineralise the compounds from the organic form back to inorganic, making them available for primary producers again.

For most areas of the ocean, the highest rates of carbon remineralisation occur at depths between 100–1,200 m (330–3,940 ft) in the water column, decreasing down to about 1,200 m (3,900 ft) where remineralisation rates remain pretty constant at 0.1 μmol kg⁻¹ yr⁻¹.^[84] This provides the most nutrients available for primary producers within the photic zone, though it leaves the upper surface waters starved of inorganic nutrients.^[85] Most remineralisation is done with dissolved organic carbon (DOC). Studies have shown that it is larger sinking particles that transport matter down to the sea floor^[86] while suspended particles and dissolved organics are mostly consumed by remineralisation.^[87] This happens in part due to the fact that organisms must typically ingest nutrients smaller than they are, often by orders of magnitude.^[88] With the microbial community making up 90% of marine biomass,^[89] it is particles smaller than the microbes (on the order of 10⁻⁶) that will be taken up for remineralisation.^[90]

Key role of phytoplankton

Phytoplankton

Glass shell (frustule) of a diatom

Coccolithophores help power the carbonate pump by producing hard structures out of calcium carbonate which eventually sink to the ocean floor

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Other cycles aluminum arsenic boron bromine cadmium chlorine chromium copper fluorine gold iodine iron lead lithium manganese mercury ozone–oxygen selenium vanadium zinc
Related topics Biogeochemistry geochemical cycle chemical cycling environmental chemistry Biosequestration Deep biosphere Ocean acidification acid rain Biogeochemical planetary boundaries
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