carbon dioxide – The Ocean: Explained

The three carbon pumps of the ocean: biological, carbonate, and physical

Introduction

Carbon is the most critical component of all biological compounds and is exchanged around the Earth through a biogeochemical cycle (Archer, 2010). Although carbon is part of natural planetary systems, current concentrations of carbon dioxide (CO₂) are the highest they have been in 14 million years, and this increase is attributed to anthropogenic activity, specifically from the Industrial Revolution of the 1700s (Etheridge et al., 1996; Falkowski et al., 2000; Zhang, Pagani, Liu, Bohaty, & DeConto, 2013). The oceanic carbon cycle is comprised of processes that cycle carbon around different areas of the ocean, the seafloor, the Earth’s interior, and the atmosphere. In pre-Industrial Revolution times, the ocean provided a net source of CO₂ to the atmosphere, whereas now most of the carbon that enters the ocean comes from anthropogenic, atmospheric CO₂ (Raven et al., 2005). According to Falkowski et al. (2000), the ocean is a reservoir for ~38,400 gigatons (Gt) of carbon, a vast amount when compared to the terrestrial biosphere (~2,000 Gt) and the atmosphere (~720 Gt). CO₂ is diffused into the ocean’s surface waters and dissolves, now ready to enter the oceanic carbon cycle through three pumps: the biological pump, the carbonate pump, or the physical pump (Duan & Sun, 2003). The biological pump utilises autotrophy, such as photosynthesis by phytoplankton, to export carbon from the upper, sunlit ocean to the ocean interior or seafloor sediments and respire organic carbon into inorganic carbon (Emerson & Hedges, 2008). The carbonate pump is a process of ocean carbon sequestration driven by calcifying plankton, which releases CO₂ back into the atmosphere but sequesters it by sinking to the seafloor (Smooth & Key, 1975); this is why this process is also referred to as the carbonate counter pump. The physical pump is the physio-chemical process whereby carbon is transported from the ocean surface to its interior, where it can be stored for hundreds of years (Ito & Follows, 2003; Toggweiler, Murnane, Carson, Gnanadesikan, & Sarmiento, 2003).

The Biological Pump

The biological pump is a process of oceanic carbon sequestration that is driven mainly by autotrophic phytoplankton that inhabits the surface waters. This method of autotrophy, photosynthesis, converts CO₂ (dissolved inorganic carbon (DIC)) into organic biomass (particulate organic carbon (POC)) (Passow & Carlson, 2012; Sigman & Hain, 2012). Photosynthesis is the initial method of bringing carbon into the biological pump. It is further moved throughout the ocean by entering the food web after phytoplankton, which are primary producers at the lowest trophic level, are eaten by consumers. Carbon can then stay in the food web as higher trophic levels continuously consume organisms, or it can be released from the food web in the form of defecation or dead tissue (Passow & Carlson, 2012). This carbon sequestration process by primary production accounts for a vast majority of carbon fixation on Earth (Christina & Passow, 2007; De La Rocha, 2003).

Carbon is also moved into deep ocean currents or seafloor sediments by sinking organic matter. Organic material is formed by phytoplankton in the euphotic zone located at the surface level of the ocean. When plankton or other marine organisms eat, defecate, die, and decompose, this material, known as marine snow, begins to sink downwards (Passow & Carlson, 2012). One phytoplankton cell sinks at a rate of approximately 1 metre per day, meaning that, with an average depth of 4,000m, it can take over ten years for one carbon-carrying phytoplankton to reach the seafloor. Organic and inorganic matter, as well as expulsion of faecal matter from larger predators, aggregate to form marine snow that has a greater sinking velocity and can complete its journey to the seafloor in days (Heinze et al., 2015; Turner, 2015). Once this sinking organic biomass reaches deep-sea levels, it can enter the food web by becoming metabolic fuel for organisms that live there, including benthic organisms and deep-sea fish (Turner, 2015). Sinking matter transports an estimated 5–20 Gt of carbon to the deep ocean annually, where between 200 million–500 million tonnes of carbon is sequestered for thousands of years in seafloor sediment (Giering et al., 2020; Guidi et al., 2015; Henson et al., 2011). Any global warming-induced change on the integrity or function of phytoplanktonic populations will alter the efficiency at which POC is transported to ocean depths, with feedbacks on the rate of climate change.

With less than 0.5 Gt of sinking carbon reaching sequestration in seafloor sediment, between 44.5 Gt – 54.5 Gt of carbon is remineralised in the euphotic zone (Ducklow, Steinberg, & Buesseler, 2001) and between 5 Gt – 6 Gt of carbon is remineralised in midwater processes during particle sinking (Feely, Sabine, Schlitzer et al., 2004). Remineralisation occurs in the biological pump when heterotrophic organisms utilise organic matter produced by autotrophic organisms. They recycle the compounds from the organic form back to the inorganic form through respiration, making them available for reuse in primary production (Guidi et al., 2015). Remineralisation usually occurs with dissolved organic carbon (DOC) rather than POC because particles must typically be smaller than the organism taking it up for remineralisation (Lefevre, Denis, Lambert, & Miquel, 1996; Schulze & Mooney, 2012).

The particles that make it to the seafloor sediment may remain there for millions of years, trapping the carbon with them. Together, the processes that make up the biological pump ultimately remove carbon in its organic form from the ocean’s surface and return it to DIC in the deeper ocean. The thermohaline circulation (THC) returns deep-ocean DIC to the atmosphere on timescales that exceed millennia (the topic of THC will be explained further in “The Physical Pump”).

The Carbonate Pump

The carbonate pump is an extension of the biological pump but instead sequesters particulate inorganic carbon (PIC) and is driven by calcifying organisms (organisms that produce calcium carbonate (CaCO₃) shells). The leading contributor to the carbonate pump is the calcifying plankton known as coccolithophores due to the vast quantity of their global population. Coccolithophores are eukaryotic, unicellular phytoplankton that produces overlapping calcite platelets called coccoliths and are currently one of the most significant contributors to carbonate sediments in the deep sea (Hofmann et al., 2010; Renaud, Ziveri, & Broerse, 2002). Coccolithophore production of coccoliths through the uptake of dissolved inorganic carbon and calcium produces CaCO₃and CO₂, hence the alternate name of carbonate counter pump.

Ca²⁺ + 2HCO_3- → CaCO₃ + CO₂ + H₂O

However, some of the CO₂ released in calcification can be used in photosynthesis (Mackinder, Wheeler, Schroeder, Riebesell, & Brownlee, 2012), and over extended periods coccolithophores contribute to decreased levels of atmospheric CO₂. It is currently unknown as to the function of the coccolith. However, many theories have been proposed, including protection from predators or grazing zooplankton (Young, Andruleit, & Probert, 2009) or ballasting the cell for vertical migration into deeper water (Raven & Waite, 2004). The latter would be of considerable advantage to the carbonate pump in getting the carbon trapped in coccoliths to the seafloor. The most abundant species of coccolithophore is Emiliania huxleyi. It is likely to be the greatest global producer of calcite, meaning this species is an essential organism in transporting carbon from the ocean to be buried in marine sediment; they play a crucial role in the global biochemical carbon cycle (Balch, Holligan, & Kilpatrick, 1992).

The production of CaCO₃ shells in calcifying organisms such as molluscs, foraminifera, coccolithophores, crustaceans, echinoderms, and corals (Zondervan, Zeebe, Rost, & Riebesell, 2001) is the central part of the carbonate pump. When CO₂ dissolves in the surface layer of the ocean, it combines with water molecules. It enters into a series of chemical reactions that result in ions that calcifying organisms combine with calcium ions (Ca²⁺) to form CaCO₃ (Zeebe & Wolf-Gladrow, 2001). Even though one CO₂ molecule is released in calcification, one carbon atom becomes trapped within the CaCO₃ molecule used in calcification and becomes part of the sediment once it sinks to the bottom of the ocean. This means calcification takes two carbon atoms from the environment and only releases one back into it, even though the formation of calcium carbonate shells is a source of CO₂ (Mackie, McGraw, & Hunter, 2011) over the long-term calcifying organisms provide a sink for CO₂.

Calcifying organisms provide a large mechanism for the downward transport of CaCO₃ (Smith & Key, 1975). Dead organisms sink to the seafloor and dissolve on the way down and release carbon into deep-sea currents or reach the seafloor and build up to form CaCO₃ sediments stored for large timescales. The scale at which CaCO₃ makes its way down varies from species to species. For example, calcifying zooplankton (pteropods, ostracods and foraminifera) promote fast particulate inorganic carbon sequestration to the deep ocean due to the relatively large mass of their shells, which makes them sink rapidly. In comparison, calcifying phytoplankton such as coccolithophores will hardly sink individually, and even still, they have a broad range in sinking rates when they assimilate into biological aggregates. The burial of CaCO₃ in deep-ocean sediment is one of the primary mechanisms to reduce atmospheric CO₂ on geological timescales related to silicate weathering processes (Cartapanis, Galbraith, Bianchi, & Jaccard, 2018). Eventually, tectonic processes, including heat and pressure, transform seafloor sediments containing CaCO₃ into limestone; this process locks carbon away for millions of years (Folk, 1980). Over time, these sediment layers eventually return carbon to the oceans by weathering and erosion (Gibb, 1978).

The Physical Pump

The physical pump uses different processes to transport DIC from the ocean surface to its interior. Firstly, the solubility of CO₂ in water is the initial process of getting carbon into the ocean as DIC (Duan & Sun, 2003). Carbon dioxide dissolves in oceanic water and, unlike many other gases, it reacts with water to form a collective of ionic and non-ionic species (DIC), which include dissolved free carbon dioxide (CO₂ (aq)), carbonic acid (H₂CO₃), bicarbonate (HCO₃₋), and CO₃²⁻ (Weiss, 1974). There is a strong inverse function of seawater temperature on the solubility of CO2, as solubility is greater in colder water (Toggweiler et al., 2003). As sea surface temperature (SST) increases, less CO₂ can be taken up by the ocean; the progressive warming of the oceans releases CO₂ in the atmosphere because of its lower solubility in warmer seawater.

The THC is part of a global, oceanic conveyor belt, driven by heat and freshwater fluxes, where thermo refers to temperature, and haline refers to salinity, which together determines seawater’s density (Rahmstorf, 2003). The model of THC was first described by Stommel and Arons (1959), where they explored how temperature and salinity moved ocean water around the globe. Seawater with higher temperatures expands and is less dense than seawater at lower temperatures (Millero, Gonzalez, & Ward, 1976). Seawater with higher salinity is denser than seawater with lower salt content (Millero et al., 1976). Seawater with lower density floats over denser seawater; this is known as stable stratification (Maiti, Gupta, & Bhattacharyya, 2008). When dense seawater masses are initially formed, they are not stably stratified and will seek to correctly locate themselves vertically by their density and become stably stratified. This stratification process is the main driving force behind deep ocean currents, which carry carbon that has sunk from surface layers on the global conveyor belt, essentially sequestering it for hundreds of years.

The process of denser seawater joining the global conveyor belt is called downwelling. Higher density water accumulates and sinks below lower density water at places within the ocean where warm rings spin clockwise and create surface convergence, pushing the surface water downwards (Rao, Joshi, & Ravichandran, 2008; Yang, 2009). These areas of downwelling bring large amounts of carbon from the surface waters to be sequestered down below. Alternately, upwelling brings dense, cooler water to the surface to replace the warmer surface water. This water is usually nutrient-rich, including dissolved CO₂, meaning that these areas of upwelling provide an ideal location for an abundance of phytoplankton to carry out primary production and ultimately recycle the carbon brought to the surface from the deep ocean (Sarhan, Lafuente, Vargas, Vargas, & Plaza, 2000)

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The effect of ocean acidification on calcareous phytoplankton due to human-induced climate change

Introduction

Humans have begun to drastically alter the atmospheric environment since the Industrial Revolution of the 1700s, where the burning of fossil fuels released exponential amounts of carbon dioxide into the air. Along with increasing the global temperature, affecting life cycles of organisms, and disassembling natural species interactions, this release initiated a process known as ocean acidification, whereby the pH of the ocean has decreased from 8.2 to 8.1 (Feely, Doney, & Cooley, 2009). The effect that ocean acidification has on marine organisms starts at the base of the food chain with primary producers, such as phytoplankton, and repercussions can then be seen throughout entire food webs. Calcareous phytoplankton (i.e., coccolithophores) may be at high risk of impact from ocean acidification due to the calcium carbonate plates that they produce through calcification. Coccolithophores are key players in the global biogeochemical cycle, the pelagic food chain, and oxygen production through photosynthesis, so any negative impacts that ocean acidification may have on them need to be understood and mitigated for the health and wellbeing of humankind.

Climate change

Climate change refers to global climatic shifts that were intensified in the mid-20th century through the burning of fossil fuels (Griffin, 2018). This anthropogenic influence has resulted in an increasing amount of carbon dioxide (CO2) in the atmosphere with a >90% probability that the observed average global temperature increase is due to human-induced greenhouse gas concentrations and around half of this increase occurring in the last three decades (Feely et al., 2009; Rosenzweig et al., 2008). Anderson, Hawkins, and Jones (2016) describe the greenhouse effect as infrared energy that has been re-emitted from solar radiation is absorbed by water vapour and CO2 to create a ‘blanket’ around planet Earth. As CO₂ concentrations increase, this greenhouse effect is enhanced, and the planet warms further beyond its natural average temperature.

Regional warming exhibits observable biological changes in terrestrial systems. Included amongst these changes are increases in coastal erosion (Beaulieu & Allard, 2003; Forbes, Parkes, Manson, & Ketch; 2004; Orviku, Jaagus, Kont, Ratas, & Rivis, 2003), melting permafrost (Frauenfeld, Zhang, Barry, & Gilichinsky, 2004), and glaciers shrinking in all seven continents (Oerlemans, 2005). Natural increases and decreases in Earth’s temperature have been recorded over millions of years (Savin, 1977), which can add scepticism to the idea that current global heating is due to anthropogenic activity rather than part of a natural cycle. A study by Reichert, Bengtsson, and Oerlemans (2002) into Swiss and Norwegian glacial retreat showed that the retreat could not be due to natural climatic change because it exceeds glacial fluctuations derived from the general circulation model, meaning that another force must be acting on the glaciers.

Climate change’s effects also reach the ocean. Edwards and Richardson (2004) describe phenological changes that occur in the wake of a warming ocean. An increase in sea surface temperature (SST) is used as an indicator for climate change, and it affected the seasonal development and phenology of plankton species. Massive phenological changes occurred at a 0.9oC increase of SST. The extent of these changes was more significant than those of terrestrial studies (Root et al., 2003), which may indicate that marine communities have heightened sensitivity to climate change. Furthermore, Edwards and Richardson’s (2004) study showed there was a different rate of response in communities to ocean warming, creating a “mismatch between successive trophic levels and a change in the synchrony of timing between primary, secondary and tertiary production” (p. 883). This mismatch will directly affect populations at higher trophic levels, including commercial fish species, marine mammals, and seabirds, and adaption will need to take place to realign these trophic levels with primary production. Not only are higher trophic levels at risk in response to changing phenology of plankton due to climate change, but other essential ecosystem services will be impacted, including the production of oxygen that we breathe, the sequestration of CO2, and the biogeochemical cycling of nitrogen, phosphorus, and silica (Richardson & Schoeman, 2004). Beaugrand and Reid (2003) study provided evidence for climate change-induced, long-term changes to the three trophic levels of phytoplankton, zooplankton, and salmon. Various declines and increases of phytoplankton and zooplankton ensued due to regional temperature increases over time, which caused the ultimate decline of salmon stocks in 1988, and this decline is expected to continue as the climate proceeds to change.

Ocean acidification

The ocean has absorbed approximately one-quarter of anthropogenically emitted CO2 over the industrial era, causing chemical reactions that reduce oceanic pH, concentrations of carbonate ions (CO₃^2-), and saturation states of two calcium carbonate (CaCO₃) minerals calcite and aragonite; this is the process of ocean acidification (OA) (Feely et al., 2009). The ocean absorbs CO₂in two ways: through photosynthesis undertaken by marine phytoplankton (Rost, Riebesell, & Burkhardt, 2003) and through the dissolving of CO₂ in seawater (Feely et al., 2009). When CO₂ reacts with seawater, it creates carbonic acid (H₂CO₃), which dissociates to hydrogen ions (H⁺) that combine with carbonate to form bicarbonate ions (HCO_3-). As atmospheric CO₂ increases, the ocean continually absorbs greater amounts of CO₂, and as temperature increases, CO₂ leaks out of the ocean back into the atmosphere. As CO₂ is absorbed, carbonate gets used up and must be replaced by stocks from the deeper ocean. Currents bring water with fresh carbonate to the surface and circulate water carrying the captured carbon into the ocean. As SST increases, this circulation process becomes more difficult, and the ocean stratifies. The surface water begins to saturate with CO₂, decreasing support for phytoplankton, and photosynthetic CO₂ uptake slows. Feely et al. (2009) reported the average pH of the ocean since the industrial era to have decreased by 0.11 pH units (29% acid concentration increase) and predicted a further decrease of 0.3 pH units (150% acid concentration increase) by 2100.

Marine organisms (e.g. coral, bivalves) extract bicarbonate ions from seawater to form skeletons or shells in a process called calcification:

Ca²⁺ + 2HCO_3- → CaCO₃ + CO₂ + H₂O

Calcification is a source of CO2 and is balanced by weathering, a process where rainwater reacts with carbonate rocks and consumes atmospheric CO2 on its way to rivers:

CaCO₃+ CO₂ + H₂O → Ca²⁺ + 2HCO_3-

A decline in the availability of carbonate ions will affect the degree of carbonate polymorph saturation in seawater, therefore compromising marine organisms’ ability to construct their skeletons/shells (Gledhill, Wanninkhof, & Eakin, 2009). A study by Gattuso et al. (2009) examined OA’s ability to affect the calcification of calcareous plankton, realising that any changes experienced by calcified taxa may alter the oceans’ ability to act as a global carbon sink. Calcareous, planktonic foraminifera observations from the Southern Ocean showed a decrease in shell weight compared to older records, with the implication that OA is the causing factor (Moy, Howard, Bray, & Trull, 2009). Suppose OA can change the morphology of a species and consequently affect multiple species up the trophic levels. In that case, the question is whether these species will cope with the effects of acidification through adaptation. This was highlighted in an experiment by Bibby, Cleall-Harding, Rundle, Widdicombe, and Spicer (2007) who observed the effects of acidification on the defence mechanisms of the periwinkle Littorina littorea toward its predator the green shore crab Carcinus meanus. L. littorea can reinforce their calcified shells when experiencing extensive predation pressure, with snails exposed to C. meanus for 15 days producing shells that were 30% thicker than those that were not under predation pressure. At reduced seawater pH, L. littorea could not thicken their shell and compensated by altering their behaviour to avoid the crabs. This shows that adaptation is possible, but at what cost? Behavioural alteration in L. littorea could mean they spend less time feeding, or it could affect their interactions with other species with unknown consequences.

Coccolithophores

*Umbilicosphaera sibogae* coccolithophore conglomeration taken with ZEISS MERLIN Scanning Electron Microscope. By ZEISS Microscopy – CC BY-NC-ND 2.0, https://www.flickr.com/photos/zeissmicro/6908938729

Coccolithophores are unicellular, eukaryotic, calcareous phytoplankton that produces overlapping calcite platelets called coccoliths and are presently one of the main primary producers of the open ocean and significant contributors to carbonate sediments in the deep sea. (Hofmann et al., 2010; Renaud, Ziveri, & Broerse, 2002). Coccolithophores produce coccoliths through the uptake of dissolved inorganic carbon and calcium, and CaCO₃ and CO₂are produced. Some of the CO₂ released in calcification can be used in photosynthesis (Mackinder, Wheeler, Schroeder, Riebesell, & Brownlee, 2012). It is possible that over long periods, coccolithophores may contribute to decreased levels of atmospheric CO₂. Even though one CO₂ molecule is released in calcification, one carbon atom becomes trapped within the CaCO₃ molecule used to make coccoliths and becomes part of the sediment once it sinks to the bottom of the ocean. This means calcification takes two carbon atoms from the environment and only releases one back into it. Even though the formation of calcium carbonate shells is a source of CO₂, over the long-term coccolithophores provide a sink for CO₂ (Mackie, McGraw, & Hunter, 2011). It is currently unknown as to the function of the coccolith, but many theories have been proposed. These include protection from predators or grazing zooplankton (Young, Andruleit, & Probert, 2009), ballasting the cell for vertical migration into deeper, more nutrient-rich water (Raven & Waite, 2004), protection from virus infection (Raven & Waite, 2004), or aiding in the filtering of non-photosynthetically active radiation at the surface to aid in photosynthesis or focussing light to the chloroplasts in deeper water (Raven & Crawfurd, 2012).

The most abundant species of coccolithophore is Emiliania huxleyi. It is likely to be the greatest global producer of calcite, meaning this species is an essential organism in transporting carbon from the ocean to be buried in marine sediment; they play a crucial role in the global biochemical carbon cycle (Balch, Holligan, & Kilpatrick, 1992). As the most abundant species, this also means that E. huxleyi is a key organism in the marine pelagic system, where the base of the food web is comprised of over 5,000 species of phytoplankton (Rost & Riebesell, 2004). Among this vast array of species, only a few select taxonomic groups are responsible for most of the pelagic systems primary production, one of these is coccolithophores and therefore reinforces their importance in the oceanic ecosystem. Coccolithophores are sensitive to nitrogen, phosphorus, and silicate ratios in the water, inducing competitive dominance between coccolithophores and other phytoplanktonic communities such as diatoms, microflagellates, and dinoflagellates. Anthropogenic interference with these ratios comes in the form of agricultural runoff leading to eutrophication and increasing nitrogen in seawater, causing coccolithophores to form blooms in these favourable environments and outcompete other species (Yunev et al., 2007). As agriculture and the use of nitrogen-based pesticides increases, more runoff makes its way to the ocean, which could see a tip in the balance of conditions to favour coccolithophore species of phytoplankton. From there, will other phytoplankton species be able to adapt in time, or will they be outcompeted and eventually become extinct?

As OA decreases carbonate saturation in seawater, coccolithophores ability to produce coccoliths may be inhibited as the increase in atmospheric CO₂ may affect their calcification mechanisms. When environments change, some organisms adapt to suit their changing environment, so as ocean pH decreases, will coccolithophores have an evolutionary response? Beaufort et al. (2011) found that coccolithophores have channels to pump out H⁺ ions during calcification to avoid acidosis. Furthermore, a feedback loop is created because when the function of these channels is disrupted, calcification is halted. This study provided evidence that increased oceanic CO^₂ concentrations decreased coccolith mass as OA impairs the normal function of ion channels and places selection pressure on coccolithophore’s calcification rates (Tyrell, Holligan, & Mobley, 1999). Flynn, Clark, and Wheeler (2016) found that coccolithophores put under OA conditions showed selection for lower calcification rates to avoid the risk of acidosis at higher H⁺ concentrations. They predicted coccolithophore calcification would decrease by 25% as OA continued and atmospheric CO^₂ reached 750 ppm (an increase of 335.62 ppm from current atmospheric CO₂ levels). Conversely, a long-term study of an E. huxleyi population that was allowed to reproduce for 700 generations under conditions similar to those predicted for the year 2100 showed their population could adapt and increase calcification and CaCO₃ content (Benner et al., 2013). This contrasts with other short-term experiments and shows that long-term OA exposure could alter calcification responses in E. huxleyi, and potentially other calcareous phytoplankton as well. Similarly to this experiment, Smith et al. (2012) found naturally forming populations of highly calcified coccolithophores in water with low CaCO₃ saturations.

The Paleocene-Eocene Thermal Maximum (PETM) occurred approximately 55.5 million years ago and saw a global temperature increase of 5–8 °C and ~12,000 gigatons of carbon released over 50,000 years (0.24 gigaton per year) (McInherney & Wing, 2011). Today, the anthropogenic release of carbon is equal to 10 gigatons of carbon per year; therefore, it will only take 1,200 years for 12,000 gigatons of carbon to be released as in the PETM. It has been shown that there was no change in coccolithophore distribution attributed to acidification in the PETM (Beaufort et al., 2011; Iglesias-Rodriguez et al., 2008), meaning that they were likely able to adapt over the 50,000 years of exposure to increasing carbon levels. Will 1,200 years be enough time for them to adapt again? Results from studies about the effect OA has on coccolithophores contradict one another. Different experiment environments show different outcomes, including short-term experiments vs long-term experiments, different coccolithophore species, and different regional populations of coccolithophores (Beaufort et al., 2011; Benner et al. 2013; Flynn et al., 2016; Smith et al., 2012). An alternative way to look at the discrepancies seen between experiments is that studies about the effects of OA are often hypothesised to have a negative outcome. If it is looked at that a decreased pH is a more favourable condition to coccolithophores, rather than the alternative, it can explain why highly calcified coccolithophores can be found in conditions of low CaCO₃. In this scenario, they would have bioengineered their higher pH environment to a more favourable, lower one through calcification and the release of CO₂, creating highly calcified individuals. Once they “created” an environment with favourable pH that now had low concentrations of CaCO₃, they would no longer need the CaCO3. In some studies, at low pH environments, coccolith growth appeared to be inhibited. However, perhaps the coccolithophores were satisfied with the pH in their environment and therefore did not need to alter it further through calcification. But, I digress. What are your thoughts on the matter?

Conclusion

Anthropogenically induced climate change has a clear impact on natural, global processes, specifically the ocean’s role in balancing CO₂ levels in the atmosphere. Coccolithophores play a part in this role by producing vast amounts of calcite and acting as a carbon sink. The effects of ocean acidification on coccolithophores have been explored through previous studies and experimentation, with multiple conclusions drawn. This makes it difficult to understand how catastrophic the effects of ocean acidification will be in the future as atmospheric CO₂ continues to rise, primarily because any changes to coccolithophores will have direct and indirect consequences for other taxa and ecosystem processes. Alternatively, coccolithophores might be environmental bioengineers and use calcification to alter oceanic pH levels to suit their environment. Further research is essential to narrow down coccolithophores precise role in the ecosystem, determine the reason for their coccoliths, discover if they can adapt to a changing environment and how quickly they can do this, and establish the effect ocean acidification has had on them and can potentially have on them in the future, and what will this mean for the world.

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