Ocean carbon sink

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About 10 Pg C is released annually (2017-2019) into the atmosphere by anthropogenic emissions of carbon dioxide. Part of the released CO2 is transferred to the ocean by physical processes. Another part is incorporated in biomass through photosynthesis in the terrestrial and marine environment. However, the carbon stored on land and in the ocean that will not return as CO2 to the atmosphere over multi-decadal periods is only a small part of the global gross primary production (~100-300 Pg C/yr [1][2]). Estimates of the sequestered carbon are: about 3 Pg C/yr on land and about 2.4 Pg C/yr in the sea [3]. However, a substantial fraction (order 40%) of terrestrial C storage is offset by land-use change emissions[4]. Storage of 1 g C is equivalent to sequestration of 3.67 g CO2; 1 Pg C = 1 petagram C = 1 megatonne C = 1015 g C.


Global carbon stocks

The major global pools of potentially available C (carbon) include the atmosphere, oceans, fossil fuels, and - collectively - vegetation, soils, and detritus. The oceans are the largest C pool, encompassing an estimated 38 000 petagrams of dissolved inorganic C [5]. The geological C pool, composed primarily of fossil fuels, is the next largest pool, estimated at 2000-4000 Pg C. Vegetation (mostly terrestrial, above and below ground) and detritus hold around 2000 Pg C, followed by the atmosphere, which holds about 800 Pg C - which is comparable to the carbon stock in the ocean surface layer[6]. The ocean carbon pool consists mainly of dissolved CO2 , bicarbonate (HCO3-), carbonate (CO32-) and carbonic acid (H2CO3).

CO2 flux from the atmosphere to the ocean

It is estimated that the ocean has absorbed 2.9 ±0.7 Pg C/yr of atmospheric CO2 in 2023, which is about 30% of the current human emissions[7].

The air-sea CO2 flux [math]F[/math] is determined by the concentration difference across a very thin skin layer (molecular boundary layer << 1 mm) at the top of the ocean[8],

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

In this equation, [math] fCO_{2, subskin}, \, fCO_{2, topskin}[/math] are the partial pressures (or, more precisely, the fugacities) of CO2 at the bottom and top of the skin layer respectively, [math]K_{600} \; [m/s][/math] is the gas transfer velocity across the skin layer, [math]Sc \approx 660[/math] is the CO2 Schmidt number (the ratio of kinematic viscosity and diffusivity) and [math]\alpha \; [mol/m^3\, Pa][/math] is the aqueous-phase solubility of CO2.

The gas transfer velocity [math]K_{600}[/math] depends on the thickness of the skin layer and is usually parameterized as a function of wind speed. An often used parameterization is[9] [math]\quad K_{600} = 6.17 \, U_{10}^2 + 9.25 \, U_{10} \; [10^{-7} \, m/s] \, ,[/math]

where [math]U_{10} \; [m/s][/math] is the wind speed 10 m above the sea surface. There is significant uncertainty in this estimate, since other mechanisms can affect turbulence at the interface, such as air bubbles, buoyancy fluxes, fetch, surface slicks and rain. A realistic estimate of uncertainties due to the gas transfer is ±10% applied to global data[10].

The CO2 solubility [math]\alpha \; [mol/m^3\, MPa][/math] strongly decreases with temperature, from about 50 at 5°C to about 25 at 30°C . [11]

Large differences occur between the CO2 fugacities at the bottom and the top of the surface skin layer as it takes on the order of one year to equilibrate CO2 in the surface ocean with atmospheric CO2. Most of the differences are caused by variability in the oceans due to water circulation and biotic processes. The greatest uncertainty in the estimate of the air-sea CO2 exchange flux [math]F[/math] comes from estimating [math]fCO_{2, subskin}[/math]. This value has a high spatial variability (currently between about 300 and 440 Pa), with a uncertainty on the order of 40 Pa. [7]

CO2 sequestration by the biological pump

The oceans’ role as a sink for CO2 is driven by two processes: the solubility pump and the biological pump.

The biological C pump refers to the uptake of CO2 by marine plankton from the surface waters through photosynthesis. Nitrogen-fixing bacteria (so-called diazotrophs, in particular Trichodesmium and the unicellular symbiont UCYN-A) feed the biological pump by enhancing primary production in the nitrogen-poor surface layer of the open ocean. Particulate organic material (POC, e.g. dead plankton cells, faecal pellets), and inorganic PIC (mainly calcium carbonate) is transferred from the ocean surface to deeper ocean layers through several processes: sinking by gravity, advection by downwelling currents and diurnal vertical migration of grazing organisms[12]. Sinking of small organic particles is possible 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[13]. More than 90% of the organic matter sinking below the euphotic zone is mineralized before it reaches a depth of 1000 m and a much smaller part (on the order of 0.2 Pg C/yr) reaches the ocean floor and is buried in the sediment, partly as detritus and partly as calcium carbonate[14].

Nowicki et al. (2022[15]) used computer models consistent with data from satellite-based sensors and ocean observations to estimate the carbon export from the ocean photic zone (~100 m) by the biological pump. They found an export on the order of 10 Pg C/yr with a sequestration time in the ocean on the order of 150 years. The carbon captured in the deep ocean (below 1000 m) has a turnover time of at least hundreds of years and can thus be considered sequestered[16].

A small fraction of the carbon sink is offset by ocean methane emissions, estimated at about 33 Tg CH4 [17]. This is equivalent to an offset of the ocean carbon sink of about 0.25 Pg C/yr.

CO2 sequestration by the solubility pump

The solubility pump refers to CO2 transfer through the ocean-atmosphere interface discussed above, with subsequent mixing in the upper ocean layer and transport to deeper ocean layers by large-scale ocean circulation currents[18]. The transfer of dissolved CO2 from the surface to deeper ocean layers by the solubility pump is an extremely slow process that takes hundreds to thousands of years. However, once dissolved in the ocean, a carbon atom will be sequestered for hundreds to thousands of years. [19]

The name 'solubility pump' points to the strong dependency on the CO2 solubility in seawater and the thermal stratification of the ocean. Cold, deep waters are generally rich in dissolved inorganic C because of the increased solubility of CO2. Outgassing to the atmosphere occurs when deep water wells up to warmer equatorial regions along the eastern ocean boundaries, where the solubility of CO2 is reduced[20].

There is growing evidence and consistency among methods with regard to the patterns of the multi-year variability of the ocean carbon sink, with a global stagnation in the 1990s and an extra-tropical strengthening of the sink in the 2000s. Explanations for this multi-year variability range from the ocean’s response to changes in atmospheric circulation (especially the variations in the upper ocean overturning), to external forcing through surface cooling associated with volcanic eruptions and variations in atmospheric CO2 growth rate[21][22]. Fossil fuel CO2 emissions have increased to 10 Pg C /yr in 2019 and the atmospheric CO2 concentration has reached an unprecedented level of 415 parts per million in 2022. However, the fraction of emitted CO2 remaining in the atmosphere has been fairly stable at about 45% on average since 1958. The ocean has sequestered about 25% of cumulative CO2 emissions in the period 2010-2019.

CO2 sequestration by coastal wetlands

Coastal wetlands (salt marshes, mangroves), seagrass meadows and seaweed (macroalgae) forests are net carbon sinks. The estimated global carbon sequestration capacity of 0.06-0.1 Pg C/yr is likely an underestimate when carbon outwelling is accounted for (see Blue carbon sequestration). Mineralization of organic matter discharged by rivers generates CO2 emissions, especially in estuarine high-turbidity zones. However, high primary production in offshore river plumes results in undersaturation of pCO2 and an associated carbon sink on coastal shelves that offsets a large fraction of estuarine emissions (net global emission not much greater or less than 0.1 Pg C/yr).[23][24][25]

Methods for increasing the ocean carbon sink

Several methods have been proposed to increase artificially the ocean carbon sink. Three proposals are briefly mentioned:

  1. Large scale cultivation of seaweed in nutrient-rich coastal upwelling zones. The biomass produced is then sunk into the deep ocean where the carbon remains trapped for hundreds to thousands of years[26]. However, scaling up seaweed aquaculture can have negative consequences, including disruption of the natural ecosystem and the diversion of nutrients from wild food webs. See also Seaweed (macro-algae) ecosystem services.
  2. Ocean iron fertilization. Soluble iron salts or ferrous dust are added to surface waters where it is currently lacking, for example in mid-ocean gyres and the Southern Ocean. This should increase primary production and boost the ocean carbon pump in these ocean regions. However, experiments with iron fertilization have been less successful than 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[27]. Moreover, undesirable side effects on remote ecosystems cannot be excluded[28]. Better understanding is needed of the biochemistry of ocean fertilization and its impact on the global carbon cycle and marine ecosystem.
  3. Ocean alkalinization. The concentration of carbonate or hydroxide ions in surface water is artificially raised to shift the associated chemical equilibria in seawater (see ocean acidification), thus increasing oceanic uptake of atmospheric CO2 and reducing ocean acidification. The feasibility and effectiveness of adding alkalinity at the required scale are questionable and the effects are highly uncertain[29].

Climate change impact on the carbon sink

Climate change affects the ocean carbon sink in various contrasting ways. Enhanced temperature stratification due to global warming will reduce the mixing of nutrients into the euphotic surface layer of the ocean. High temperatures and low nutrient concentrations give a competitive advantage to small phytoplankton groups that are more labile and less likely to sink than larger plankton species. The carbon sequestration by the biological pump may therefore decrease[30][31]. However, other climate change effects can strengthen carbon sequestration. For example, the expansion of oxygen minimum zones is likely to reduce microbial remineralization rates and will therefore allow sinking detrital particles to reach greater depth[32]. Ocean acidification will reduce the contribution to the carbon sink by calcifying organisms, but it will also lower ocean alkalinity due to reduced calcification, allowing additional dissolution of CO2 and a concomitant decrease of atmospheric CO2 (as explained in the article Ocean acidification).


Related articles

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


References

  1. Jump up ↑ Westberry, T.K., Silsbe, G.M. and Behrenfeld, M.J. 2023. Gross and net primary production in the global ocean: An ocean color remote sensing perspective. Earth-Science Reviews 237, 104322
  2. Jump up ↑ Wild, B., Teubner, I., Moesinger, L., Zotta, R.-M., Forkel, M., van der Schalie, R., Sitch, S. and Dorigo, W. 2022. VODCA2GPP – a new, global, long-term (1988–2020) gross primary production dataset from microwave remote sensing. Earth Syst. Sci. Data, 14: 1063–1085
  3. Jump up ↑ Le Quere, C., Andrew, R.M., Friedlingstein, P., Sitch, S., Pongratz, J. et al. 2018. Global carbon budget 2017. Earth Syst. Sci. Data 10: 405–48
  4. Jump up ↑ Global Carbon Project (2023)
  5. Jump up ↑ Hedges, J.I., 1992. Global biogeochemical cycles: progress and problems. Marine Chemistry 39: 67-93
  6. Jump up ↑ McLeod, E., Chmura, G.L., Bouillon, S., Salm, R., Bjork, M., Duarte, C.M., Lovelock, C.E.,Schlesinger, W.H., Silliman, B.R., 2011. A blueprint for blue carbon: toward an improved understanding of the role of vegetated coastal habitats in sequestering CO2. Front. Ecol. Environ. 9: 552–560
  7. ↑ Jump up to: 7.0 7.1 Ford, D. J., Blannin, J., Watts, J., Watson, A. J., Landschützer, P., Jersild, A. and Shutler, J. D. 2024. A comprehensive analysis of air‐sea CO2 flux uncertainties constructed from surface ocean data products. Global Biogeochemical Cycles 38, e2024GB008188
  8. Jump up ↑ Wanninkhof, R., Asher, W.E., Ho, D.T., Sweeney, C. and McGillis, W.R. 2009. Advances in Quantifying Air-Sea Gas Exchange and Environmental Forcing. Annu. Rev. Mar. Sci. 1: 213–244
  9. Jump up ↑ Nightingale, P.D., Malin, G., Law, C.S., Watson, A.J., Liss, P.S., Liddicoat, M.I., Boutin, J. and Upstill-Goddard, R.C. 2000. In situ evaluation of air-sea gas exchange parameterizations using novel conservative and volatile tracers. Global Biogeochem. Cycles 14: 373-387
  10. Jump up ↑ Woolf, D.K., Land, P.E., Shutler, J.D., Goddijn-Murphy, L.M. and Donlon, C.J. 2016. On the calculation of air-sea fluxes of CO2 in the presence of temperature and salinity gradients. J. Geophys. Res. Oceans 121: 1229–1248
  11. Jump up ↑ Weiss, R. F. 1974. Carbon dioxide in water and seawater: The solubility of a non‐ideal gas. Marine Chemistry 2: 203–215
  12. Jump up ↑ Boyd, P., Claustre, H., Levy, M., Siegel, D.A. and Weber, T. 2019. Multi-faceted particle pumps drive carbon sequestration in the ocean. Nature 568: 327–335
  13. Jump up ↑ Bonnet, S., Benavides, M., Le Moigne, F.A.C., Camps, M., Torremocha, A., Grosso, O., Dimier, C., Spungin, D., Berman-Frank, I., Garczarek, L. and Cornejo-Castillo, F.M. 2022. Diazotrophs are overlooked contributors to carbon and nitrogen export to the deep ocean. The ISME Journal 17: 47–58
  14. Jump up ↑ Cartapanis, O., Galbraith, E. D., Bianchi, D. and Jaccard, S. L. 2018. Carbon burial in deep-sea sediment and implications for oceanic inventories of carbon and alkalinity over the last glacial cycle. Clim. Past. 14: 1819–1850
  15. Jump up ↑ Nowicki, M., DeVries, T. and Siegel, D. A. 2022. Quantifying the carbon export and sequestration pathways of the ocean's biological carbon pump. Global Biogeochemical Cycles 36, e2021GB007083
  16. Jump up ↑ Siegel, D.A., DeVries, T., Doney, S.C. and Bel, T. 2021. Assessing the sequestration time scales of some ocean-based carbon dioxide reduction strategies. Environ. Res. Lett. 16, 104003
  17. Jump up ↑ Rosentreter, J.A., Borges, A.V., Deemer, B.R., Holgerson, M.A., Liu, S., Song, C., Melack, J., Raymond, P.A., Duarte, C.M., Allen, G.H., Olefeldt, D., Poulter, B., Battin, T.I. and Eyre, B.D. 2021. Half of global methane emissions come from highly variable aquatic ecosystem sources. Nat. Geosci. 14: 225–230
  18. Jump up ↑ Volk, T. and Hoffert, M. 1985. Ocean carbon pumps: Analysis of relative strengths and efficiencies in ocean-driven atmospheric CO2 changes, in The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present (eds. E. T. Sundquist and W. S. Broecker). Geophys. Monogr. Ser. 32: 99–110, AGU, Washington, D. C.
  19. Jump up ↑ Jones, D. C., Ito, T., Takano, Y. and Hsu, W.-C. 2014. Spatial and seasonal variability of the air-sea equilibration timescale of carbon dioxide. Global Biogeochem. Cycles 28: 1163–1178
  20. Jump up ↑ Lefèvre, N., Veleda, D. and Hartman, S.E. 2023. Outgassing of CO2 dominates in the coastal upwelling off the northwest. African coast. Deep Sea Research Part I: Oceanographic Research Papers 200, 104130
  21. Jump up ↑ DeVries, T., Le Querec, C., Andrews, O., Berthet, S., Hauck, J., Ilyina, T., Landschutzer, P., Lenton, A., Limak, I.D., Nowicki, M., Schwinger, J. and Seferian, R. 2019. Decadal trends in the ocean carbon sink. PNAS 116: 11646–11651
  22. Jump up ↑ Hauck, J., Zeising, M., Le Quere, C. ,Gruber, N., Bakker, D.C.E., Bopp, L., Chau, T.T.T., Gurses, O., Ilyina, T., Landschützer, P., Lenton, A., Resplandy, L., Rödenbeck, C., Schwinger, J. and Seferian, R. 2020. Consistency and Challenges in the Ocean Carbon Sink Estimate for the Global Carbon Budget. Front. Mar. Sci. 7: 571720
  23. Jump up ↑ Chen, C.-T.A. and Borges, A.V. 2009. Reconciling opposing views on carbon cycling in the coastal ocean: Continental shelves as sinks and near-shore ecosystems as sources of atmospheric CO2. Deep Sea Research Part II: Topical Studies in Oceanography 56: 578-590
  24. Jump up ↑ Chen, C.-T., A., Huang, T.-H., Fu, Y.-H., Bai, Y. and He, X. 2012. Strong sources of CO2 in upper estuaries become sinks of CO2 in large river plumes. Current Opinion in Environmental Sustainability 4: 179-185
  25. Jump up ↑ Noriega, C. and Araujo, M. 2013. Carbon dioxide emissions from estuaries of northern and northeastern Brazil. Scientific reports 4, 6164
  26. Jump up ↑ Froehlich, H.E., Afflerbach, J.C., Frazier, M. and Halpern, B.S. 2019. Blue Growth Potential to Mitigate Climate Change through Seaweed Offsetting. Current Biology 29: 3087–3093
  27. Jump up ↑ Jiang, H-B., Hutchins, D.A., Zhang, H-R., Feng, Y-Y., Zhang, R-F., Sun, W-W., Ma, W., Bai, Y., Wells, M., He, D., Jiao, N., Wang, Y. and Chai, F. 2024. Complexities of regulating climate by promoting marine primary production with ocean iron fertilization. Earth-Science Reviews 249, 104675
  28. Jump up ↑ Aumont, O. and Bopp, L. 2006. Globalizing results from ocean in situ iron fertilization studies. Global Biogeochem. Cycles 20: GB2017
  29. Jump up ↑ Gattuso, J-P., Magnan, A.K., Bopp, L., Cheung, W.W.L., Duarte, C.M., Hinkel, J., Mcleod, E., Micheli, F., Oschlies, A., Williamson, P., Billé, R., Chalastani, V.I., Gates, R.D., Irisson, J-O., Middelburg, J.J., Pörtner, H-O. and Rau, G.H. 2018. Ocean Solutions to Address Climate Change and Its Effects on Marine Ecosystems. Front. Mar. Sci. 5: 337
  30. Jump up ↑ Ani, C.J. and Robson, B. 2021. Responses of marine ecosystems to climate change impacts and their treatment in biogeochemical ecosystem models. Marine Pollution Bulletin 166: 112223
  31. Jump up ↑ Bindoff, N.L., Cheung, W.W.L., Kairo, J.G., Arístegui, J., Guinder, V.A., Hallberg, R., Hilmi, N,. Jiao, N., Karim, M.S., Levin, L., O’Donoghue, S., Purca Cuicapusa, S.R., Rinkevich, B., Suga, T., Tagliabue, A. and Williamson, P. 2019. Changing Ocean, Marine Ecosystems, and Dependent Communities. Ch. 5 in IPCC Special Changing Ocean, Marine Ecosystems, and Dependent Communities. In: IPCC Special Report on the Ocean and Cryosphere in a Changing Climate. Cambridge University
  32. Jump up ↑ Stukel, M.R., Irving, J.P., Kelly, T.B., Ohman, M.D., Fender, C.F. and Yingling, N. 2023. Carbon sequestration by multiple biological pump pathways in a coastal upwelling biome. Nature Communications 14: 2024


The main author of this article is Job Dronkers
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Citation: Job Dronkers (2025): Ocean carbon sink. Available from http://www.coastalwiki.org/wiki/Ocean_carbon_sink [accessed on 1-04-2025]
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