The surface ocean and the atmosphere are closely coupled systems within the fast carbon cycle, and CO₂ levels in both rise and fall essentially in parallel. As anthropogenic emissions increase the concentration of carbon in fast cycle reservoirs, more and more CO₂ flows from the atmosphere to the surface ocean each year [NOAA GFDL Earth System Model] through direct absorption, which is then recycled through respiration and photosynthesis. The ocean is estimated to have absorbed at least 25-30% of the total CO₂ emitted since the beginning of the industrial era [Friedlingstein et al., 2022] ,with some estimates as high as 40% [Carroll et al., 2022].
Seawater is stratiﬁed by density gradients, which are formed by variations in salinity and temperature, and suppress vertical mixing of the water, contributing to the slow turnover of ocean bottom water with the atmosphere [Sprintall, 2009]. Therefore, once carbon is transported to the deep ocean, it remains trapped there for hundreds to thousands of years – becoming part of the slow carbon cycle.
Organic Carbon Pathways
Biological carbon fixation occurs mainly in the ocean surface layer, where marine organisms, primarily phytoplankton, photosynthetically fix dissolved inorganic carbon, allowing the surface ocean to absorb more CO₂ from the atmosphere [Heinze et al., 2015]. Approximately 85-90% of the carbon fixed through photosynthesis in the euphotic zone (the uppermost layer of the ocean) will be remineralized and recycled [Fox et al., 2020], remaining in the fast carbon cycle at the surface ocean or re-releasing into the atmosphere. However, a small percentage of plankton sink into the deep ocean after they die, transporting the carbon acquired at the surface along with them. Grazing on phytoplankton by higher trophic members of the ecosystem, such as zooplankton and fin fish, packages this carbon into fecal pellets which sink rapidly, thereby removing biologically fixed carbon to the deep sea. The process of photosynthetic marine organisms moving carbon from the fast carbon cycle in the surface ocean to the slow carbon cycle in the deep ocean is known as the “biological pump” [De La Rocha et al., 2014].
Once biomass sinks into the deep ocean, it is subject to several sequestration fates: labile organic carbon is likely to be metabolized by microbes and benthic macrofauna (e.g. amphipods) and remineralized into deep ocean waters [Gao et al., 2021], while biomass that is naturally resistant to degradation in marine environments (such as wood) is likely to be buried in ocean sediments, where much of it will be stored for geological timescales [Burdige, 2007]. Depending on location, ultimate depth of settlement, and chemical fate of the biomass, this process results in carbon storage for a minimum of centuries, to upwards of geologic timescales (i.e., hundreds of thousands to millions of years). The ocean is a massive reservoir that, by current estimates, stores approximately 39,000 gigatons of dissolved carbon [Rackley, 2010] — roughly 50 times more than the amount contained in the atmosphere [Kayler et al., 2017].
Worldwide, the biological pump transfers approximately ten gigatons of carbon from the atmosphere to the deep ocean each year [Sarmiento, 2002], where it is projected to remain, on average, for longer than 1,000 years [Orr, 1992].
Terrestrial forest systems photosynthetically fix carbon in the fast cycle, storing it in aboveground biomass (wood), belowground biomass (root structures and soils), and their supported secondary biomass, such as mycelium networks. Oceans and rivers naturally transport fixed carbon from these terrestrial systems (e.g. fallen trees transported to the ocean through river systems) and carry this biomass to the deep ocean [Kandasamy et al., 2016], thereby transferring carbon from the fast to the slow cycle.
Growing and sinking macroalgae at the surface of the ocean differs from coastal blue carbon seaweed afforestation, as the biomass growth — and thus the photosynthetic fixation — occurs in the open ocean. Thus, when intact macroalgae sinks, it is transferred to the slow carbon cycle in the deep ocean. This contrasts with organic carbon that sinks in coastal waters on the shallow continental shelf, where warm temperatures, wave action, intense bioturbation, and more rapid currents support remineralization of most of that organic carbon to CO₂, which is released back to the fast carbon cycle. Because macroalgae species generally have more favorable Redfield ratios (the ability to fix carbon per available nutrients) than phytoplankton [Martiny, 2013], macroalgae cultivation and sinking in the open ocean may increase the amount of carbon transported to the deep ocean by the biological pump.
The exact durability of carbon stored is dependent on the depth, location, and chemistry where biomass sinks [LaRowe et al., 2020]. Ideal sinking locations are deep, stable, and characterized by relatively high rates of sedimentation, which increases the proportion of biomass that is buried and preserved for millennia.
Inorganic Carbon Pathways
Greater total volumes of CO₂ within the fast carbon cycle manifest as higher concentrations of CO₂ in the atmosphere and higher concentrations of dissolved inorganic carbon in the surface ocean. Ocean acidification is the result of CO₂ dissolving in the ocean and reducing the pH of surface waters. This acidification of seawater makes it harder for many types of marine organisms, like oysters, corals, and scallops, to form their protective calcium carbonate shells and skeletons [Doney et al. 2009], which has the potential to cause catastrophic ecosystem impacts over the coming decades.
Photosynthetic organisms — primarily phytoplankton — transform dissolved CO₂, a form of inorganic carbon, into organic carbon in the same way that trees transform atmospheric CO₂ into terrestrial biomass. This process helps mitigate CO₂-induced ocean acidification by decreasing CO₂ concentrations in the surface layer of the ocean. However, ocean acidification, warming, and pollution affect photosynthesizing organisms in the ocean in complex ways, potentially diminishing the ocean's capacity to fix inorganic carbon through photosynthesis. Without positive interventions, the capacity of the ocean to sequester and store atmospheric CO₂ may be impaired by warming and acidification [Chikamoto 2021].
As atmospheric CO₂ dissolves in the surface ocean, carbonic acid is formed and subsequently dissociates into protons and bicarbonate ions, the balance of which is governed in a given region of water by temperature, pressure, salinity, and pH. The result is a “buffer pool” of dissolved inorganic carbon, which allows a volume of seawater to dissolve many times more carbon than might be otherwise expected [Sarmiento et al., 2006]. When and where surface waters mix and sink into deeper water, inorganic carbon that has been taken from the atmosphere and dissolved in surface water is carried to the slow carbon cycle of the deep ocean.
Dissolution of alkaline minerals into the surface ocean — i.e., “alkalinity enhancement” — alters the balance of the carbonate system in seawater, increasing surface seawater’s capacity to convert dissolved CO₂ to dissolved bicarbonate ions, and thus take up more CO₂ from the atmosphere. Alkalinity enhancement achieves carbon removal by reallocating fast cycle carbon (aqueous CO₂) to the larger and more stable bicarbonate reservoir.
Because the magnitude of the bicarbonate reservoir is so large compared to the aqueous CO₂ reservoir [Zeebe and Wolf-Gladrow, 2001], the residence time of carbon within the bicarbonate reservoir is comparably longer than the residence time of carbon in the CO₂ reservoir, rendering the bicarbonate reservoir functionally consistent with a slow carbon reservoir rather than a fast one. Simultaneously, this process amplifies the fast cycle transfer of atmospheric CO₂ into surface waters, as the dissolved CO₂ that was reallocated to the bicarbonate reservoir is replaced by novel atmospheric CO₂ exchanged across the air-sea boundary [Campbell et al., 2022].
Notably, this differs from an atmospheric-centric framing of carbon removal for alkalinity enhancement, which is primarily concerned with the fast cycle transfer of atmospheric CO₂ that “replaces” the associated reduction in the partial pressure of CO₂ resulting from a CO₂ transfer to the bicarbonate reservoir.
The gross mass transfer ratio of ocean alkalinity enhancement will vary seasonally and regionally between 0.26-0.95 with the physico-chemical conditions of surface seawater and the molecular weight of the alkaline source mineral [He et al., 2023].