Ocean carbon sink

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About 90–100 Pg C was released to the atmosphere from fossil fuel combustion and cement production during the decade 2016–2025. Part of this CO2 is absorbed by the oceans through physical and chemical processes, while another part is taken up by photosynthesis in terrestrial and marine ecosystems. However, only a small fraction of this carbon remains sequestered from the atmosphere on multi-decadal timescales.

Global gross primary production (GPP) is estimated at roughly 200–220 Pg C yr⁻¹, but net carbon sequestration represents only a few percent of this total[1][2]. Current estimates of net uptake are about 3 Pg C yr⁻¹ on land and about 2.4 Pg C yr⁻¹ in the ocean[3].. Terrestrial uptake is partly offset by emissions from land-use change (about 1–1.5 Pg C yr⁻¹)[4].

Atmospheric CO2 concentration reached about 423 ppm in 2024. The fraction of emitted CO2 remaining in the atmosphere (the airborne fraction) has remained relatively stable at about 45% on average since 1958.

For reference, storage of 1 g C corresponds to the sequestration of 3.67 g CO2. Units commonly used in the global carbon budget are petagrams of carbon (Pg C), where 1 Pg C = 1 gigaton C = 1015 g C.


Global carbon stocks

The major global pools of potentially available carbon (C) include the atmosphere, oceans, fossil fuels, and the terrestrial biosphere. The terrestrial pool comprises vegetation and soils, with additional large carbon stores in permafrost and detrital organic matter.

The oceans are the largest active carbon reservoir, containing about 38,000 Pg C, mainly as dissolved inorganic carbon (DIC) [5]. The terrestrial biosphere stores roughly 3000–3500 Pg C when including vegetation, soils, and permafrost. Fossil fuel reserves (coal, oil, and natural gas) contain on the order of 1000–1500 Pg C.

The atmosphere contains about 900 Pg C, a quantity comparable to the carbon stored in the upper ocean (mixed layer)[6][7].

Oceanic carbon occurs predominantly as dissolved inorganic carbon (DIC), consisting of about 90% bicarbonate (HCO3−), about 9% carbonate (CO32−), and only a small fraction as dissolved CO2 and carbonic acid (H2CO3).


CO2 sequestration by the solubility pump

It is estimated that the ocean absorbed 2.9 ± 0.7 Pg C yr⁻¹ of atmospheric CO2 in 2023, about 30% of current anthropogenic emissions[8].

Oceanic carbon uptake through air–sea exchange involves three steps: gas transfer across the ocean–atmosphere interface, mixing within the ocean mixed layer, and transfer to the deep ocean. Each process constrains the overall rate of carbon uptake.

The air–sea CO2 flux [math]F[/math] is driven by the concentration difference across a very thin skin layer (molecular boundary layer < 1 mm) at the ocean surface[9]:

[math]F = K_{600} (Sc/600)^{-0.5} \left( \alpha_{subskin} fCO_{2,subskin} - \alpha_{topskin} fCO_{2,topskin} \right) \, . \qquad (1)[/math]

Here, [math]fCO_{2,subskin}[/math] and [math]fCO_{2,topskin}[/math] are the CO2 fugacities at the bottom and top of the skin layer, [math]K_{600}[/math] [m s⁻¹] is the gas transfer velocity, [math]Sc \approx 660[/math] is the Schmidt number, and [math]\alpha[/math] [mol m⁻³ Pa⁻¹] is the aqueous solubility of CO2. By convention, flux from the atmosphere to the ocean is negative.

The gas transfer velocity depends on turbulence in the surface layer and is commonly parameterized as a function of wind speed. A widely used formulation is[10]:

[math]K_{600} = 10^{-7} \left( 6.17 \, U_{10}^2 + 9.25 \, U_{10} \right) \, , \qquad (2)[/math]

where [math]U_{10}[/math] [m s⁻¹] is the wind speed at 10 m height. Additional processes such as wave breaking, bubbles, buoyancy fluxes, and surface films introduce uncertainty; globally, this is on the order of ±10%[11]. The CO2 partial pressure [math]pCO_2[/math] at the base of the skin layer is controlled by mixing within the ocean mixed layer (which is a few tens to about one hundred meter thick).

Uptake of atmospheric CO2 is limited by carbonate chemistry: an increase in dissolved inorganic carbon (DIC) generates a much stronger increase in [math]pCO_2[/math] (see Ocean acidification). This sensitivity is described by the Revelle factor [math]R[/math],

[math]\dfrac{\Delta pCO_2}{pCO_2} = R \, \dfrac{\Delta DIC}{DIC} \, , \qquad (3)[/math]

where [math]DIC = [CO_{2,aq}] + [HCO_3^-] + [CO_3^{2-}] \qquad[/math] (sum of dissolved CO2, bicarbonate and carbonate ).

The Revelle factor is typically of order 10.[12].

In the absence of ventilation, the ocean mixed layer would approach equilibrium with the atmosphere on timescales of months to about a year. Continued uptake therefore requires ventilation (subduction) of surface waters into the ocean interior[13]. Transport to the deep ocean occurs on timescales of hundreds of years, and individual carbon atoms may remain sequestered for hundreds to thousands of years[14].

Cold waters take up more CO2 because solubility increases at lower temperatures, and deep-water formation transports carbon-rich surface waters into the ocean interior. Intermediate and deep waters are further enriched in DIC by remineralization of sinking organic matter. In contrast, outgassing occurs where these carbon-rich waters are upwelling into warmer regions, particularly in tropical and eastern-boundary upwelling systems, where lower solubility and high subsurface DIC favor CO2 release to the atmosphere[15][16]. The solubility pump is therefore a major mechanism of ocean carbon uptake, particularly through cooling, deep-water formation, and ventilation at high latitudes. We will see in the next section that the biological pump also contributes significantly to long-term carbon sequestration.

There is increasing agreement on multi-year variability in the ocean carbon sink, including a stagnation in the 1990s and strengthening in extra-tropical regions during the 2000s. Proposed drivers include changes in atmospheric circulation (notably upper-ocean overturning), volcanic activity causing cooling of the ocean surface, and variability in atmospheric CO2 increase rates[17][18].


CO2 sequestration by the biological pump

The ocean takes up atmospheric CO2 through physical, chemical and biological processes, including the solubility pump and the biological carbon pump. The biological pump is the set of biologically mediated processes that transform dissolved inorganic carbon into organic matter in surface waters and transfer part of this carbon to the ocean interior, where it can remain isolated from the atmosphere for decades to millennia.

Marine phytoplankton fix dissolved inorganic carbon during photosynthesis. Nitrogen-fixing microorganisms (diazotrophs), such as Trichodesmium and the unicellular symbiont UCYN-A, can enhance primary production in nitrogen-limited regions of the open ocean (see Nutrient conversion in the marine environment)[19]. By lowering surface-ocean pCO2, photosynthesis enhances the air-sea CO2 gradient and promotes additional uptake of atmospheric CO2. However, long-term sequestration requires that part of the fixed carbon be exported below the surface mixed layer or euphotic zone before it is respired back to CO2.

Organic carbon is transferred from the photic zone (ocean surface layer, ̴100 m) to greater depths through multiple pathways, including gravitational sinking of particulate organic carbon (POC; e.g. dead plankton, aggregates and fecal pellets), physical mixing and advection, and active transport by diel vertically migrating animals[20][21]. Sinking of small organic particles occurs through the aggregation into large-sized aggregates ('marine snow') comprised of tens to hundreds of cells, large/dense enough to sink. Such aggregation is further promoted through association with diatoms, ingestion by grazers and incorporation into large fecal pellets[22].

More than 90% of the organic matter sinking below the photic zone is mineralized before it reaches a depth of 1000 m. Only a small fraction (on the order of 0.2 Pg C/yr) reaches the ocean floor and is buried over longer timescales, as organic carbon or as calcium carbonate[23]. The formation and export of particulate inorganic carbon (PIC), mainly calcium carbonate, is sometimes described as part of the carbonate pump; unlike organic-carbon export, calcification consumes alkalinity and can increase local CO2, partly offsetting CO2 uptake.

Using models constrained by satellite and in situ observations, Nowicki et al. (2022[21]) estimated global carbon export from the photic zone at about 10 Pg C yr−1. They estimated that gravitational sinking accounts for about 70% of this export, most of which is associated with zooplankton fecal pellets, while migrant transport and physical mixing account for smaller fractions. The average sequestration time differs among pathways, with values on the order of decades to centuries. Carbon transported below about 1000 m is generally isolated from the atmosphere for longer, roughly century to millennium timescales[24].


CO2 sequestration by coastal wetlands

Coastal wetlands (salt marshes, mangroves), seagrass meadows, and some macroalgal (seaweed) systems can act as net carbon sinks through a combination of:

  • long-term burial of organic carbon in sediments,
  • accumulation of biomass,
  • lateral export (“outwelling”) of dissolved and particulate carbon to the coastal ocean.

These processes differ substantially in their climatic significance and residence times. Sediment burial can represent carbon storage on centennial to millennial timescales, whereas exported dissolved or particulate carbon may be remineralized and returned to the atmosphere on much shorter timescales.

The global rate of long-term organic carbon burial in these “blue carbon” ecosystems is commonly estimated to be on the order of 0.05–0.08 Pg C yr⁻¹. Estimates of the total climatic effect may increase when lateral export processes are included, although the fraction of exported carbon that results in durable atmospheric CO2 removal remains uncertain.

Mangroves and salt marshes export substantial amounts of dissolved inorganic carbon (DIC) and total alkalinity (see Ocean acidification). Exported total alkalinity (TA) can enhance the buffering capacity of coastal waters and may contribute indirectly to long-term atmospheric CO2 uptake by increasing seawater capacity to store inorganic carbon. However, alkalinity export does not necessarily correspond directly to net atmospheric CO2 sequestration, because the climatic effect depends on air–sea equilibration, carbonate chemistry, circulation, and the fate of the exported carbon. The relative balance between DIC and TA export strongly influences coastal carbonate chemistry. When TA export is sufficiently large relative to DIC export, coastal acidification can be mitigated and oceanic CO2 uptake promoted. Conversely, when DIC export exceeds TA export, acidification and CO2 outgassing may be enhanced. [25]

The outwelling fluxes represent only a fraction of the carbon delivered to the coastal ocean by rivers. Global fluvial carbon export is estimated at approximately 0.52 ± 0.17 Pg C yr⁻¹ as DIC, 0.30 ± 0.14 Pg C yr⁻¹ as dissolved organic carbon (DOC), and 0.18 ± 0.04 Pg C yr⁻¹ as particulate organic carbon (POC).[26]

A substantial fraction of fluvial DOC and POC is remineralized in estuaries, particularly in high-turbidity zones and coastal wetlands. As a result, many estuaries and nearshore wetlands are net sources of atmospheric CO2. In contrast, continental shelves are estimated to constitute a net sink for atmospheric CO2 at the global scale, although strong regional variability exists. This uptake is often associated with high primary production in river plumes and shelf waters, which can lead to pCO2 undersaturation.[27][28][29]

In addition to mangroves and salt marshes, tidal freshwater wetlands also contribute substantially to long-term carbon storage.[30] Forested tidal freshwater systems (e.g. Melaleuca and Casuarina) can store aboveground and soil carbon stocks comparable to those of mangrove forests.

Part of the climatic benefit associated with aquatic carbon uptake is offset by methane (CH4) emissions from aquatic ecosystems, estimated globally at approximately 33 Tg CH4 yr⁻¹.[31] This corresponds to approximately 0.025 Pg C yr⁻¹ as methane carbon, or roughly 0.9 Pg CO2-equivalent yr⁻¹ using a 100-year global warming potential. The magnitude of this offset varies strongly among ecosystems and timescales, particularly because methane exerts a stronger warming effect over shorter time horizons.

The long-term stability of blue carbon storage is sensitive to disturbance. Sea-level rise, erosion, drainage, eutrophication, warming, storm damage, and land-use change can accelerate remineralization of previously buried carbon and reduce the persistence of sediment carbon stocks.


Methods for increasing the ocean carbon sink

Several methods have been proposed to artificially enhance the ocean carbon sink. Three approaches are briefly outlined:
Large-scale cultivation of seaweed
Seaweed can be cultivated in nutrient-rich coastal and upwelling regions. The harvested biomass could be transported and sunk into the deep ocean, where the carbon may remain isolated from the atmosphere for decades to centuries or longer, depending on depth, remineralization, and ocean circulation[32]. However, large-scale aquaculture may disrupt ecosystems and divert nutrients from natural food webs (see also Seaweed (macro-algae) ecosystem services).

Ocean iron fertilization
Soluble iron salts or ferrous dust are added to iron-limited (high-nutrient, low-chlorophyll) regions such as the Southern Ocean, the equatorial Pacific, and the subarctic North Pacific to stimulate phytoplankton growth and enhance the biological carbon pump. Field experiments have generally shown smaller and shorter-lived carbon export than initially expected. Several possible causes have been suggested, such as rapid conversion of soluble ferrous sulphate to rapidly precipitating ferric hydroxide, primary production limited by other nutrients and trace metals, competition between picocyanobacteria and diatoms, where the former will not reach the deep ocean due to low sedimentation rates and grazing by microzooplankton[33]. Potential unintended impacts on marine ecosystems remain uncertain[34].

Ocean alkalinity enhancement
This approach increases ocean alkalinity to promote the conversion of dissolved CO2 into bicarbonate and carbonate ions, thereby enhancing CO2 uptake and reducing acidification. Natural alkalinity is supplied mainly by weathering of rocks. For example, dissolution of olivine (a ubiquitous mineral in mafic rocks) can be represented schematically as

[math]Mg_2SiO_4 + 4CO_2 + 4H_2O \rightarrow 2Mg^{2+} + 4HCO_3^- + H_4SiO_4 \, , \qquad (5)[/math]

which increases alkalinity and converts CO2 primarily into bicarbonate in seawater[35]. The feasibility, scalability, and environmental impacts of large-scale alkalinity enhancement remain uncertain[36][37].


Climate change impact on the carbon sink

Climate change affects the ocean carbon sink in several, sometimes opposing, ways. Increased surface warming enhances stratification and can reduce the upward supply of nutrients to the euphotic zone. This may reduce primary production in nutrient-limited regions and can favor smaller phytoplankton, which are often less efficiently exported to depth, potentially weakening the biological pump[38][39].

Ocean acidification increases the Revelle factor and thereby reduces seawater buffer capacity. All else being equal, this causes surface-ocean pCO2 to rise more rapidly as dissolved inorganic carbon accumulates, reducing the ocean’s capacity to absorb additional atmospheric CO2 [40].

Other changes may enhance sequestration in some regions. For example, expansion of oxygen minimum zones may reduce remineralization or particle attenuation in some settings, allowing a larger fraction of sinking organic matter to reach greater depths, although this effect is regionally variable[41][42].

Ocean acidification is expected to reduce calcification by many marine calcifying organisms. Because calcification consumes alkalinity and is associated with local CO2 release, reduced calcification may provide a partial negative feedback by maintaining higher alkalinity and allowing additional uptake of atmospheric CO2 (see Ocean acidification).


Related articles

Blue carbon sequestration
Ocean acidification
Ecosystem services
Governance policies for a bio-based blue economy
Greenhouse gas regulation


References

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The main author of this article is Job Dronkers
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