719

Journal of Oceanography, Vol. 60, pp. 719 to 729, 2004

Review

Keywords:⋅⋅⋅⋅⋅ CO2 effects,⋅⋅⋅⋅⋅ carbonate chemis-try,

⋅⋅⋅⋅⋅ phytoplankton,⋅⋅⋅⋅⋅ biocalcification,⋅⋅⋅⋅⋅ photosynthesis,⋅⋅⋅⋅⋅ ecosystem re-sponses,

⋅⋅⋅⋅⋅ biological carbonpump,

⋅⋅⋅⋅⋅ biogeochemicalfeedbacks.

* E-mail address: uriebesell@ifm-geomar.de

Copyright © The Oceanographic Society of Japan.

Effects of CO2 Enrichment on Marine Phytoplankton

ULF RIEBESELL*

Leibniz Institut für Meereswissenschaften, IFM-GEOMAR, Kiel, Germany

(Received 14 October 2003; in revised form 5 May 2004; accepted 5 May 2004)

Rising atmospheric CO2 and deliberate CO2 sequestration in the ocean changeseawater carbonate chemistry in a similar way, lowering seawater pH, carbonate ionconcentration and carbonate saturation state and increasing dissolved CO2 concen-tration. These changes affect marine plankton in various ways. On the organismallevel, a moderate increase in CO2 facilitates photosynthetic carbon fixation of somephytoplankton groups. It also enhances the release of dissolved carbohydrates, mostnotably during the decline of nutrient-limited phytoplankton blooms. A decrease inthe carbonate saturation state represses biogenic calcification of the predominantmarine calcifying organisms, foraminifera and coccolithophorids. On the ecosystemlevel these responses influence phytoplankton species composition and succession,favouring algal species which predominantly rely on CO2 utilization. Increasedphytoplankton exudation promotes particle aggregation and marine snow formation,enhancing the vertical flux of biogenic material. A decrease in calcification may af-fect the competitive advantage of calcifying organisms, with possible impacts on theirdistribution and abundance. On the biogeochemical level, biological responses to CO2enrichment and the related changes in carbonate chemistry can strongly alter thecycling of carbon and other bio-active elements in the ocean. Both decreasing calcifi-cation and enhanced carbon overproduction due to release of extracellular carbohy-drates have the potential to increase the CO2 storage capacity of the ocean. Althoughthe significance of such biological responses to CO2 enrichment becomes increasinglyevident, our ability to make reliable predictions of their future developments and toquantify their potential ecological and biogeochemical impacts is still in its infancy.

driven by earth’s orbital patterns, the extent of change inclimate, CO2 and other climate-relevant gases is control-led by a variety of mechanisms, the exact balance of whichis still unknown. Among these mechanisms, biologically-driven reactions and feedbacks, involving both terrestrialand marine ecosystems, are bound to play a critical role.

In the past 200 years, man has become a distinct fac-tor of the earth system by influencing the carbon cycle,mainly through the injection of carbon dioxide by theburning of fossil fuels and by changes in land use. Thecurrent period, frequently termed the “anthropocene”, hasno direct analogue in the geological past. Although hu-man activities impact most strongly on the carbon cycle,this has repercussions for the earth system as a whole,since the carbon cycle is coupled with climate, the watercycle, nutrient cycles and photosynthesis on land and inthe oceans. It is also central to human society as it is di-rectly linked to the production of energy. Growing world

1. IntroductionThe gas composition of the atmosphere and the cli-

mate of planet earth are products of the co-evolution ofthe biospheric and climatic systems. Over earth’s historythese have evolved gradually as a result of externalforcings, such as solar irradiance and orbital patterns, andinternal feedbacks between the atmosphere and biosphere.During its latest phase this co-evolution settled into apersistent pattern of glacial-interglacial cycles, with at-mospheric CO2 fluctuating between 180 µatm in glacialand 280 µatm in interglacial times (Petit et al., 1999).For about the last 20 million years atmospheric pCO2 hasnever exceeded 300 µatm (Berner et al., 1990). Whilethe periodic cycle of glacial-interglacial changes is clearly

720 U. Riebesell

population and changing life styles put an increasing pres-sure on energy production and, hence, man-made changesin the carbon cycle proceed at an unprecedented rate.

While the impact of human existence on the carboncycle may have been unintended, attempts to regulate ormanage it, however, will require an intentional globaleffort on an unprecedented scale. No single measure islikely to provide the capacity to stabilize atmospheric CO2concentration in the presence of continued CO2 emissionsat the present-day level. Among the wide range of possi-ble mitigation strategies ocean CO2 sequestration, eitherthrough an enhancement of biological carbon fixation orthrough direct CO2 injection, is receiving increasing at-tention. Obviously, any responsible effort aiming to miti-gate the green-house problem through purposeful CO2sequestration in the ocean requires a solid understandingof marine ecosystem regulation and biogeochemical cy-cling. While the study of terrestrial ecosystem responsesto man-made environmental changes, including CO2 en-richment, has been a topic of intense research over morethan a decade, research of this kind in the marine bio-sphere is developing only recently. The limited informa-tion presently available indicates, however, that marineecosystems are indeed sensitive to CO2 enrichment. Thispaper provides examples of marine planktonic responsesto increased seawater CO2 levels and the related changesin seawater carbonate chemistry. While the initial responsetypically occurs at the organismal level, the effects cancarry on to the ecosystem level, which in turn can impactmarine biogeochemical cycling. Corresponding changesin biologically-mediated carbon cycling can have bothpositive and negative feedbacks on oceanic CO2 storageand air/sea CO2 exchange.

How does this relate to deep ocean CO2 sequestra-tion? Ideally, it should not, at least not on centennial timescales. If properly conducted, the sequestered CO2 wouldbe cut off from exchange with the surface layer for atleast several hundred years. Phytoplankton and the bulkof zooplankton thrives in or close to the sunlit surfacelayer of the ocean, i.e. outside the waters which would beconsidered for deep sea CO2 injection. Hence, the effectof deep ocean CO2 sequestration on the upper oceanplankton would be—if anything—to contribute to miti-gating adverse effects of a human-induced increase inatmospheric CO2 on marine plankton. The surface ocean,on the other hand, experiences an unintentional, gradualand global CO2 enrichment due to the continuous rise ofatmospheric CO2. Changes in atmospheric pCO2 are mir-rored by corresponding changes in the carbonate systemof surface seawater with a time lag of less than one year(Zeebe and Wolf-Gladrow, 2001). Responses of the ma-rine plankton to CO2 enrichment therefore can be expectedto occur in phase with the present increase in atmosphericCO2. The focus of this paper will therefore be on the ef-

fect of CO2-related changes in seawater carbonate chem-istry caused by the uptake of anthropogenic CO2 by thesurface ocean.

2. Effects of CO2 Enrichment on Seawater Carbon-ate ChemistryThe present rise in atmospheric CO2 and the result-

ing net flux of CO2 into the surface ocean causes a con-tinuous change in seawater carbonate chemistry. The dis-solved CO2 reacts with water according to:

CO H O H HCO2 2+

3+ ⇔ + −

HCO H CO3+

32− −⇔ + .

At a typical surface ocean pH value of 8.2, less than 1%of the added CO2 remains as dissolved CO2, while therest is converted into HCO3

– (ca. 90%) and CO33– (ca.

9%). The reaction of CO2 with water generates one pro-ton for each HCO3

– and two protons for each CO33–

formed. This acidification causes a shift of the pH-de-pendent equilibrium between CO2, HCO3

– and CO33– to-

wards higher CO2 levels and lower carbonate ion con-centrations, [CO3

3–]. The net result of CO2 addition there-fore is an increase in the concentration of dissolved CO2,[CO2], a decrease in [CO3

3–] and seawater pH (Fig. 1),and a slight increase in the concentrations of HCO3

– anddissolved inorganic carbon, DIC.

The recent increase in atmospheric CO2 from a pre-industrial level of 280 µatm to the present level of 370µatm has decreased surface ocean pH values by approxi-

Fig. 1. Seawater pH and the dissolved carbon dioxide (CO2)and carbonate ion (CO3

2–) concentrations in the surface layerof the ocean assuming a “business as usual” (IS92a) an-thropogenic CO2 emission scenario (Houghton et al., 1995).Dashed lines represent the predicted changes in carbonatechemistry if CO2 emissions are reduced according to theKyoto Protocol (modified after Wolf-Gladrow et al., 1999).

Effects of CO2 Enrichment on Marine Phytoplankton 721

mately 0.12 units. At atmospheric CO2 levels of 700 µatm,as expected for the end of this century in a “business asusual” scenario (IS92a, Houghton et al., 1995, 2001),seawater pH will decrease by an additional 0.3 units.While the concentration of dissolved CO2 will have tri-pled by this time, the carbonate ion concentration willdecrease by nearly 50%. The projected pH shift coversthe entire range of pH variations between 7.7–8.3 pres-ently observed in open ocean surface waters. The pro-jected increase in [CO2] and decrease in [CO3

3–] is out-side the natural range of the past 20 Million years.

Changes in surface ocean carbonate chemistry causedby the present rise in atmospheric CO2 are slow and mod-est compared to those occurring in the vicinity of CO2injection plumes. Depending on the method of injectionand the surface area of the hydrate formed, seawater pHin close vicinity of the injected CO2 can decrease by asmuch 2–3 pH units (Brewer et al., 2000). With the for-mation of a thin hydrate film after slow injection of liq-uid CO2, on the other hand, pH values became indistin-guishable from the in situ background level. While theseshort-term changes in seawater chemistry occur on verysmall scale, from an impact assessment point of view thelonger-term and larger-scale effects of CO2 injection, in-cluding the risk of re-surfacing of sequestered CO2, needto be considered.

Following sections outline potential effects of CO2enrichment and the related changes in seawater carbon-ate chemistry on the marine phytoplankton from theorganismal (Section 3) to the ecosystem level (Section 4)and will give examples of possible implications for ma-rine biogeochemical cycling (Section 5).

3. Effects on the Organismal Level

3.1 PhotosynthesisPhotosynthesis involves a series of reactions that start

with capturing light energy, transferring it into the en-ergy-conserving compounds NADPH and ATP, and usingthese compounds to fix CO2 in the Calvin cycle. Activelygrowing algae allocate about 50% of recent photosynthateto protein synthesis, resulting in a competition forNADPH and ATP between CO2 fixation, transport proc-esses, nitrate reduction and protein formation (Geider andMacIntyre, 2001). Thus photosynthesis and downstreamprocesses leading to cellular growth are primarily light-dependent processes. Photosynthesis-irradiance responsecurves are commonly used to assess the photosyntheticperformance.

In the dark reaction of photosynthesis CO2 is fixedby the carboxylating enzyme, ribulose-1,5-bisphosphatecarboxylase/oxygenase (RUBISCO). RUBISCO uses CO2as the only carbon substrate and has a half-saturation con-stant of 20 to 70 µmol kg–1 of CO2, depending on the

species (Badger et al., 1998). Typical seawater CO2 con-centrations, ranging between 10 and 25 µmol kg–1, aretherefore insufficient to ensure effective operation ofRUBISCO carboxylation. To overcome the low CO2 af-finity of their main carboxylating enzyme, algae havedeveloped mechanisms to actively increase the CO2 con-centration at the site of carboxylation. These include ac-tive uptake of CO2 and HCO3

– into the algal cell and/orinto the chloroplasts. In concert with active carbon trans-port, algae catalyse the conversion between HCO3

– andCO2 intra- and extracellularly with the enzyme carbonicanhydrase. Which ever the pathways for this so-calledcarbon concentrating mechanism (CCM) may be, its op-eration requires additional energy. Since marinephytoplankton live in an environment in which light (en-ergy) is one of the primary limiting resources, the energyavailable for active transport of inorganic carbon and othernutrients is often limited. Energy-limitation may be oneof the reasons why phytoplankton do not rely entirely onthe active uptake of inorganic carbon from the abundantpool of HCO3

–, but are drawing part of their carbon fromthe small and variable pool of CO2.

Fig. 2. Photosynthesis of phytoplankton species differs withrespect to CO2 sensitivity: While most species (hereSkeletonema costatum and Phaeocystis globosa) are at orclose to CO2 saturation at present day CO2 levels (8–20µmol L–1), coccolithophores such as Emiliania huxleyi havecomparatively low affinities for inorganic carbon and ap-pear to be carbon-limited in today’s ocean. This raises thepossibility that coccolithophores may benefit directly fromthe present increase in atmospheric CO2. The range in half-saturation concentrations (K1/2; in µmol L–1) for photosyn-thesis shown here reflects the degree of regulation as a func-tion of pCO2 during growth (according to Rost et al., 2003).Highest apparent affinities for CO2 were generally observedin cells which were grown under low pCO2.

722 U. Riebesell

Recent studies indicate that dominant phytoplanktonspecies differ in their CO2 requirement. Whereas somespecies preferably use CO2 as a carbon source, othersmainly draw their inorganic carbon from the large poolof HCO3

– (e.g. Elzenga et al., 2000). Group-specific dif-ferences in CO2 sensitivity also exist in the carbon me-tabolism (Fig. 2). Most notably, the photosynthetic car-bon fixation rate of all diatom species tested so far and ofthe prymnesiophyte Phaeocystis globosa are at or closeto CO2-saturation at present day CO2 levels (Burkhardtet al., 1999, 2001; Rost et al., 2003). In contrast, thecoccolithophorids Emiliania huxleyi and Gephyrocapsaoceanica are well below saturation at these levels(Riebesell et al., 2000; Rost et al., 2003). These findingssuggest large differences in CO2-sensitivity between ma-jor phytoplankton taxonomic groups. CO2-sensitive taxa,such as the calcifying coccolithophorids, should there-fore benefit more from the present increase in atmosphericCO2 compared to the non-calcifying diatoms andPhaeocystis.

3.2 CalcificationThe formation of calcareous skeletons is a wide-

spread phenomenon in various groups of marine plank-tonic organisms. Their production and subsequent sink-ing generates a continuous rain of calcium carbonate tothe deep ocean and the sediments. CO2-related changesin seawater carbonate chemistry were recently shown toaffect marine biogenic calcification. A doubling inpresent-day atmospheric CO2 concentrations is predictedto cause a 20–40% reduction in biogenic calcification ofthe predominant calcifying organisms, the corals,

coccolithophorids (Fig. 3), and foraminifera (Gattuso etal., 1998; Kleypas et al., 1999; Bijma et al., 1999;Riebesell et al., 2000). Diminished calcification incoccolithophorids was accompanied by an increased pro-portion of malformed coccoliths and incompletecoccospheres (Fig. 4) in the two bloom-forming speciesEmiliania huxleyi and Gephyrocapsa oceanica (Riebesellet al., 2000).

Although the sensitivity of the different groups ofcalcifying organisms to CO2-related changes in carbon-ate chemistry is very similar, suggesting a commonmechanism for the observed responses in each of thesegroups, it is not clear yet which is the physiologicallyrelevant factor ultimately causing the slow down in bio-genic calcification. While the pH value, carbonate ionconcentration and carbonate saturation state each maydirectly impact the calcification process, a reasonablerepresentation of the effect of carbonate chemistry onbiologically-mediated CaCO3 precipitation can be basedon the carbonate saturation state of seawater, Ω, which isexpressed as:

Ω =[ ] ×[ ]−

∗

Ca CO2+

sw 32

sw

spK

where [Ca2+]sw and [CO32–]sw are the calcium and car-

Fig. 3. Rising CO2 decreases the ratio of calcification to or-ganic carbon production (PIC/POC) in E. huxleyi. The de-crease in PIC/POC is caused by enhanced photosyntheticcarbon fixation and reduced or constant calcification. Thistrend is consistent over a range of photon flux densities(PFDs; in µmol photons m–2s–1), yet declines under severelight-limitation (modified after Zondervan et al., 2002).

Fig. 4. Scanning electron microscopy (SEM) photographs ofEmiliania huxleyi (A, C) and Gephyrocapsa oceanica (B,D) collected from cultures incubated at pCO2 levels of ca.300 µatm (upper panel) and 780–850 µatm (lower panel).Scale bars represent 1 µm. Note the difference in thecoccolith structure (including distinct malformations) andin the degree of calcification of cells grown at normal andelevated CO2 levels (adapted form Riebesell et al., 2000).

Effects of CO2 Enrichment on Marine Phytoplankton 723

bonate ion concentrations in seawater and Ksp* is the stoi-chiometric solubility product, defined as K sp* =[Ca2+]sat × [CO3

2–]sat. [Ca2+]sat and [CO32–]sat refer to the

equilibrium ion concentrations in a seawater solution satu-rated with CaCO3.

Concerning the question which is the relevant pa-rameter determining the sensitivity of biogenic calcifica-tion to seawater carbonate chemistry, it may help to takea look at the geological past. Recent evidence from fluidinclusions suggests that calcium concentrations in theCretaceous were more than twice as high as in the mod-ern ocean (Horita et al., 2002). As CaCO3 compensationtends to restore the saturation state of the ocean towardsCaCO3 saturation, higher calcium concentrations wouldcorrespond to lower carbonate ion concentrations. In spiteof lower expected CO3

2– concentration, massive accumu-lation of biogenic carbonates occurred in the Cretaceous,suggesting that the carbonate saturation state may exert astronger control on biogenic calcification than any of theother possible candidates, e.g. pH, CO2, or CO3

2– con-centrations.

What effect reduced CaCO3 production has for cal-cifying organisms ultimately depends on the role of cal-cification and calcareous skeletons in the cell physiologyand ecology of these organisms. Despite intensive re-search on coccolithophores the function of calcificationin this group is still not well understood (Young, 1994;Paasche, 2002). It has been hypothesized that the cell-covering layer of coccoliths, the coccosphere, serves toprotect against grazing and/or viral attack and thatcoccoliths could bundle and modify the light interceptedby the algal cell (Young, 1994). While these hypothesesremain speculative, there is some support for the “trash-can function” of calcification, whereby CaCO3 precipita-tion serves as a mechanism to facilitate the use of bicar-bonate in photosynthesis. Provided that HCO3

– is the car-bon source used for calcification, one possible benefit forphotosynthesis could be the release of CO2 or protons inthe course of calcification according to the following re-actions:

Ca HCO CaCO CO H O2+3 3 2 2+ →− + +2

or

Ca HCO CaCO H2+3 3

++ →− + .

Either CO2 could be used directly in photosynthesis orprotons could be used in the conversion of HCO3

– to CO2.In either case this would provide the calcifying cell witha mechanism to access HCO3

–, the largest pool of inor-ganic carbon in seawater, for photosynthetic carbon fixa-tion.

While compelling evidence for or against such amechanism is still lacking, it appears that calcification incoccolithophores is neither a prerequisite for efficientphotosynthesis nor is it particularly effective in mitigat-ing CO2 limitation. This is indicated, for example, by thefact that photosynthesis continues unaltered when calci-fication ceases, such as when cells are grown in calcium-free medium (Paasche, 1964). Likewise, non-calcifyingcells can photosynthesise as efficient or even more effi-ciently than calcifying ones, as indicated by P vs. E rela-tionships of non-calcifying compared to calcifying strainsof E. huxleyi (Rost and Riebesell, 2004). Moreover, massspectrometric measurements indicate that non-calcifyingcells of E. huxleyi are capable of direct uptake of HCO3

–

(Rost and Riebesell, 2004), which implies that HCO3–

utilisation is not tied to calcification. Further evidencethat calcification is unsuited to prevent CO2 limitation ofcoccolithophores comes from observations that the rateof photosynthesis decreases with decreasing CO2 concen-tration despite a concomitant increase in calcification rate(Riebesell et al., 2000; Zondervan et al., 2002). In es-sence, if supply of CO2 for photosynthesis is indeed aprimary role of calcification, this mechanism is ratherinefficient when compared to CO2 concentrating mecha-nisms of non-calcifying phytoplankton such as diatomsand Phaeocystis (Burkhardt et al., 2001; Rost et al., 2003).

Whatever the purpose of calcification incoccolithophores, continued acidification of surfaceseawater due to rising atmospheric CO2 will further dete-riorate the chemical conditions for biogenic calcification.Despite a wealth of information on the effect of environ-mental conditions on coccolith production, the mecha-nism of calcification by coccolithophores is not com-pletely understood. It is too early, therefore, to make anypredictions regarding the physiological or ecological con-sequences of a CO2-related slow down in biogenic calci-fication. Obviously, a better mechanistic understandingof the process of biogenic calcification will help to betterconstrain the effects of global changes, including changesin temperature and light availability, on bio-calcification.

3.3 Elemental compositionOne of the few canonical numbers in marine

biogeochemical sciences is the Redfield ratio, the ratioof C:N:P in marine particulate matter. When Redfield dis-covered an unexpected congruence of this value in nu-merous regions of the world’s ocean (Redfield et al.,1963), it soon became one of the cornerstones in biologi-cal oceanography and geochemistry, commonly used toinfer about the large-scale operation of the marinebiogeochemical system. Recent studies have challengedthe concept of a constant Redfield ratio and have high-lighted the importance to better constrain the factors de-termining phytoplankton stoichiometry (for review see

724 U. Riebesell

Sterner and Elser, 2002). Phytoplankton indeed show ahigh degree of variability in their stoichiometric compo-sition C:N:P. While a large body of literature exists onthe role of nutrients, light and temperature on algal stoi-chiometry (Sterner and Elser, 2002), comparatively littleis known about the effects of inorganic carbon and CO2availability on phytoplankton elemental composition.

Initial studies by Burkhardt and Riebesell (1997) andBurkhardt et al. (1999) showed considerable changes inC:N:P composition in response to increasing pCO2 butalso revealed large species-specific differences. Whereasin some microalgae cellular C:N and C:P increased withincreasing pCO2, other species responded with lowerC:nutrient ratios under increased CO2 availability. More-over, algal C:N:P mainly responded to altered CO2 atpCO2 levels lower than current values and remained moreor less stable at higher than present pCO2 values. It shouldbe noted, however, that these experiments were conductedunder nutrient replete conditions. Strongest deviationsfrom Redfield proportions typically occur in nutrient lim-ited phytoplankton and under conditions of unbalancednutrient supply. Clearly, more work is needed to furtherexamine the effect of CO2 availability on phytoplanktonstoichiometry.

Research on microalgal stoichiometry has commonlyfocussed on the cellular composition of phytoplankton.Although it is well-recognized that the proportion of pri-mary produced organic matter lost from the cell throughexudation can be substantial, particularly under nutrient-limitation, this fraction of primary production is fre-quently neglected in biogeochemical analyses. As dis-solved organic matter can be subject to rapid bacterialdegradation, this material is thought to contribute littleto vertical carbon transport in the ocean. However, re-cent work demonstrated the spontaneous assembly of dis-solved organic matter into polymer gels (Chin et al.,1998). Colloidal organic matter is an important precur-sor for the formation of transparent exopolymer particles(TEP) which are shown to accelerate particle aggrega-tion and thereby enhance vertical particle flux (Passow,2002). This process may therefore provide an efficientpathway to channel dissolved and colloidal organic mat-ter into the particulate pool. Analyses on the variabilityof Redfield proportions should therefore include dissolvedorganic components.

That the production of dissolved organic matter maybe sensitive to CO2 availability is indicated in a recentstudy by Engel (2002), showing that in a natural plank-ton assemblage TEP production increases with increas-ing pCO2. A similar response was also observed in mono-specific cultures of the diatom Thalassiosira weissflogiiand the coccolithophorid Emiliania huxleyi (Heemann,2002). Since transparent exopolymer particles primarilyform from exuded polymeric carbohydrates they have a

molar C:N:P ratio much higher than the Redfield ratio(Engel and Passow, 2001). CO2 sensitivity of TEP pro-duction may therefore cause significant shifts in elemen-tal composition of particulate matter towards higher thanRedfield ratios.

4. Effects on the Ecosystem LevelThe ecological success of a phytoplankton species

is ultimately determined by its ability to optimise the bal-ance between growth and loss processes. While the formeris largely controlled by the availability and optimal utili-sation of essential resources, such as light and nutrients,the latter includes factors such as cell sinking, cell mor-tality due to grazing, viral and parasite infection, as wellas autolysis. Any investment successfully reducing mor-tality therefore weighs equally strong as cellular meas-ures capable of increasing cell growth by the same mag-nitude. Although the role of CaCO3 production incoccolithophores is not clear, reduced calcification islikely to affect both resource utilization and cellular pro-tection with possible adverse effects on the competitivefitness of this group. Although it would be premature tospeculate on possible implications of reduced calcifica-tion for coccolithophore distribution and productivity, alikely scenario would be a shift in the relative distribu-tion and productivity of the calcareous primary produc-ers relative to their siliceous counterpart, the diatoms.Both groups are important components of the pelagic foodweb and major drivers of biogeochemical cycling in theocean. Due to their specific roles in plankton ecosystemregulation, a shift in the ratio of coccolithophores to dia-toms is likely to also affect higher trophic levels. A ba-sin-wide shift in the opal:carbonate ratio of sedimentingparticles has in fact occurred across the entire North At-lantic and is attributed to large-scale changes in climaticforcing (Antia et al., 2001). The observed steady decreasein this ratio since the early eighties may reflect an in-crease in the abundance and productivity ofcoccolithophores relative to diatoms.

To examine the direct effect of CO2 enrichment onthe ecosystem level, CO2 perturbation experiments wereconducted in outdoor mesocosms with natural planktoncommunities, in Raunafjorden south of Bergen, Norway.Nine 11 m3 enclosures moored to a floating raft were aer-ated in triplicate with normal air, CO2-enriched and CO2-depleted air to achieve pCO2 levels of 190 µatm, 370µatm, and 710 µatm, simulating glacial, present day, andpredicted year 2100 conditions (assuming IPCC scenario92a; Houghton et al., 1995), respectively (Fig. 5). Theenclosures were covered with gas tight tents made ofETFE foil (Foiltec, Bremen, Germany), which allowedfor 95% light transmission of the complete spectrum ofsunlight. To allow the seawater carbonate system to freelychange in response to biological activities, aeration of the

Effects of CO2 Enrichment on Marine Phytoplankton 725

enclosed water column was discontinued after target CO2values were achieved. At this point the nutrient depletedwater was fertilized with nitrate and phosphate to initiatethe development of a phytoplankton bloom. Aeration ofthe atmosphere underneath the tents at target CO2 levelscontinued throughout the investigation.

During the course of this study the development anddecline of a bloom consisting of a mixed phytoplanktoncommunity was followed. Significantly higher net com-munity production (as measured by DIC drawdown) wasobserved under elevated CO2 levels during the build-upthe bloom (Bellerby et al., unpubl.). CO2-related differ-ences in primary production continued after nutrient ex-haustion, leading to higher production of transparentexopolymer particles (TEP) under high CO2 conditions(Engel et al., unpubl.). CO2-dependent differences were

also observed with respect to the phytoplankton compo-sition. Whereas under glacial CO2 conditions diatomsdominated at the end of the bloom, the community devel-oping under CO2 conditions as expected for the end ofthis century was dominated by Emiliania huxleyi (Huonicet al., unpubl.). These findings are consistent with ob-served differences in carbon acquisition mechanisms be-tween diatoms and coccolithophorids as outlined above(Subsection 3.1).

In principle, coccolithophores may benefit from thepresent increase in atmospheric pCO2 and related changesin seawater carbonate chemistry. At pre-industrial CO2levels, rates of photosynthetic carbon fixation of E.huxleyi and Gephyrocapsa oceanica are well below CO2saturation. In comparison, photosynthesis of other bloom-forming phytoplankton groups such as diatoms andPhaeocystis are less CO2-sensitive due to their efficientCCMs (Rost et al., 2003). Although under natural condi-tions CO2 limitation is likely to be of minor importancefor the proliferation of E. huxleyi compared to other lim-iting resources and loss processes, increasing CO2 avail-ability may improve the overall resource utilisation of E.huxleyi and possibly of other fast-growingcoccolithophore species. If this provides an ecologicaladvantage for coccolithophores, rising atmospheric CO2could potentially increase the contribution of calcifyingphytoplankton to overall primary production.

CO2 effects on the taxonomic composition and nu-trient utilization were also observed in a naturalphytoplankton assemblage by Tortell et al. (2002). A de-crease in the abundance of diatoms and an increase inthat of Phaeocystis occurred when pCO2 levels wereraised from 150 µatm to 750 µatm. This CO2-dependentshift was associated with a significant change in nutrientutilization, with higher ratios of nitrate:silicate andnitrate:phosphate consumption by phytoplankton in thelow CO2 treatment. Despite these changes, total biomassand primary productivity did not differ significantly be-tween CO2 treatments. This and the mesocosm study de-scribed above indicate that CO2 sensitivity of microalgaeis not limited to species relying entirely on CO2 utiliza-tion. Even for species capable to access the large pool ofHCO3

–, either through direct uptake of HCO3– or its con-

version to CO2 via carbonic anhydrase, increased CO2availability can reduce the allocation of resources intocarbon acquisition. In cases of resource (energy and/ornutrient) limitation, the ability to regulate resource allo-cation accordingly could provide these species a competi-tive advantage over those incapable to regulate in responseto changes in CO2 availability (Rost et al., 2003).

5. Effects on the Biogeochemical LevelDespite a high degree of structural complexity, the

energy flow and elemental cycling of many ecosystems

CO2-absorber

CO2 mixing-chamber

Air

Air

Air

air-compressor

air-compressor

air-compressor

CO2

CO2analysis

compressor

Data logger

flow meter

Enclosure 7 Enclosure 8 Enclosure 9 Enclosure 1 Enclosure 2 Enclosure 3 Enclosure 4 Enclosure 5 Enclosure 6

Fig. 5. Upper panel: Mesocosm facilities of the University ofBergen Large-Scale-Facilities in the Raunefjorden, 60.3°N,5.2°E. Shown are the nine enclosures covered by tents madeof gas-tight ETFE foil. Lower panel: Sketch of the aerationsystems for CO2 manipulation of the seawater inside themesocosms. Ambient air was pumped either directly to themesocosms (4–6), enriched with CO2 in a mixing chamber(1–3), or depleted in CO2 through adsorption onto Na2CO3platelets (7–9).

726 U. Riebesell

are dominated by a comparatively small number of spe-cies. A first step in understanding possible consequencesof environmental change on ecosystem functioning, there-fore, is to determine the relevant responses of these keyspecies. This also holds true for the marine pelagic sys-tem, the largest ecosystem on our planet. While the baseof its complex food web is formed by more than 5000species of marine phytoplankton, only a few taxonomicgroups of phytoplankton, and within these groups a rela-tively small number of species, are responsible for mostof the system’s primary production, the energy transferto higher trophic levels and the vertical export of bio-genic material to the deep ocean. These key species canbe further distinguished into so-called “functionalgroups”, i.e. phytoplankton building siliceous or calcare-ous shells, such as the diatoms and coccolithophores, re-spectively, flagellates forming organic plates or mucilagi-nous colonies, and cyanobacteria fixing atmospheric ni-trogen. Each of these functional groups has a distinct ef-fect on elemental fluxes, both between the surface anddeep ocean as well as the overlying atmosphere.

With regard to the oceanic carbon cycle, the effectof calcifying plankton, including coccolithophores as thedominant calcifying phytoplankton, differs greatly fromthat of other primary producers. The fixation of inorganiccarbon via photosynthesis in the sunlit upper mixed layerand the vertical export of part of this organic materialcause a draw down of CO2 in the surface ocean.Remineralisation of the particulate organic carbon on itsway to depth releases organically-bound CO2, which thenaccumulates in deeper layers (Fig. 6). This process, termed

the organic carbon pump, thereby causes a net draw downof CO2 from the atmosphere into the ocean. In contrast,the production and export of calcium carbonate not onlybinds dissolved carbon into particulate carbon and therebyreduces total dissolved inorganic carbon, it also lowersseawater alkalinity and changes the equilibrium betweenthe different forms of dissolved inorganic carbon, caus-ing a net release of CO2 to the atmosphere. Due to itscounteracting effect on CO2 flux, this process is oftenreferred to as the carbonate counter pump. The relativestrength of the two biological carbon pumps, representedby the so-called rain ratio (the ratio of particulate inor-ganic to organic carbon in exported biogenic matter),determines to a large extent the flux of CO2 between thesurface ocean and the overlying atmosphere.

But how can a process transporting inorganic car-bon from the surface to the deep ocean release CO2 to theatmosphere? The formation of one unit of calcium car-bonate according to

Ca CO CaCO2+32

3+ → ( )− 1

lowers the dissolved inorganic carbon (DIC) concentra-tion, with

DIC = CO HCO CO2 3 32[ ] + [ ] + [ ] ( )− − 2

by one unit. Due to its double negative charge the con-centration of carbonate ions carries the factor 2 in theexpression of total alkalinity (TAlk), with

T HCO CO OH OH H3 32 +Alk B= [ ] + [ ] + ( )[ ] + [ ] − [ ] ( )− − − −2 34 .

Thus, the precipitation of one unit of calcium carbonatelowers seawater alkalinity by two units. How this affectsthe concentration of dissolved CO2 is best illustrated, fol-lowing Zondervan et al. (2001), by plotting CO2 concen-tration as a function of DIC and TAlk (Fig. 7). Starting attypical surface ocean conditions of DIC = 2.0 mmol kg–1

and TAlk = 2.3 mmol kg–1 (beginning of diagonal vectoron left side), the corresponding CO2 concentration, [CO2],is 10.5 µmol kg–1 (assuming T = 15°C, S = 35). Precipi-tating 50 µmol kg–1 of CaCO3 lowers DIC and TAlk by0.05 and 0.1 mmol kg–1, respectively. With the new DICand TAlk values of 1.95 mmol kg–1 and 2.2 mmol kg–1

(end of diagonal vector), the corresponding CO2 concen-tration now is 12.4 µmol kg–1. Assuming that the seawaterinitially was in equilibrium with the atmosphere with re-spect to CO2, it is now CO2 over-saturated. To restoreequilibrium, 0.63 units of CO2 need to be released foreach unit of inorganic carbon precipitated via calcifica-tion. In our example, 50 µmol kg–1 CaCO3 production

Fig. 6. Changes in the concentrations of DIC, total alkalinity,and CO2 as a result of calcification at pre-industrial times(left vectors) and in year 2100 (right vectors). Numbers atthe vectors are relative values. The difference in the effectof calcification on seawater carbonate chemistry in the twoscenarios results from a CO2-related change in the seawaterbuffer capacity (see text for details, modified afterZondervan et al., 2001).

Effects of CO2 Enrichment on Marine Phytoplankton 727

would require a degassing of 32 µmol kg–1 of CO2 in or-der to maintain CO2 equilibrium between our parcel ofseawater and the overlying atmosphere. Due to the de-crease in the buffer capacity with increasing [CO2] (de-creasing pH) of seawater, the ratio of CO2 released percarbonate precipitated increases with rising atmosphericCO2 (Frankignoulle et al., 1994). Under conditions ex-pected for the year 2100 (see Fig. 1), the surface oceanequilibrium CO2 concentration would be ca. 27µmol kg–1. Under these conditions, precipitation of 50µmol kg–1 CaCO3 would release ca. 40 µmol kg–1 of CO2,corresponding to 0.79 units of CO2 for each unit of CaCO3(right vectors). If global calcification were to remain con-stant over this time, increased calcification-induced CO2release results in a positive feedback to rising atmosphericCO2 (Zondervan et al., 2001). With the observed decreaseof bio-calcification with increasing pCO2 this positivefeedback turns into a negative one, leading to an increasein the CO2 storage capacity of the surface ocean (Riebesellet al., 2000).

Obviously, the ratio of coccolithophore calcificationto organic matter production, but even more so the ratioof calcareous to non-calcareous primary production,strongly determine the relative strengths of the two bio-logical carbon pumps. CaCO3 may also act as a “ballast”mineral that increases the transfer efficiency of POC fromthe surface to the deep sea (Armstrong et al., 2002; Klaasand Archer, 2002). The presence of coccoliths could thusenhance the sedimentation of POC such as in faecal pel-lets (Buitenhuis et al., 1996). On geological time-scalesvariation in coccolith abundance in the sedimentary record

is often correlated with glacial-interglacial transitions,indicating a predominance of coccolithophores duringinterglacial periods (McIntyre et al., 1972; Henrich,1989). The presumably lower contribution ofcoccolithophores relative to non-calcifying phytoplanktonmay have contributed to the lower atmospheric CO2 lev-els in glacial times (Harrison, 2000; Ridgwell et al., 2002).

Estimates of global calcium carbonate productionrange from 0.64 to 2 Gt C per year (e.g. Milliman andDroxler, 1996). Recent estimates of global annual ma-rine primary production are between 40 and 50 Gt C an-nually (Antoine et al., 1996), of which about 10 Gt C areexported to the deep sea. The estimated export ratio ofCaCO3 to particulate organic carbon (rain ratio) is there-fore relatively high, ranging from 0.17 to 0.4 (see Holliganand Robertson, 1996 for review). Since the export ofCaCO3 and POC have opposite effects on the surfaceocean CO2 partial pressure, a change in the rain ratio af-fects the partitioning of CO2 between ocean and atmos-phere. It has been suggested that the rain ratio is not con-stant through time and is an important factor in the regu-lation of global climate (Archer and Maier-Reimer, 1994).

Changes in marine production and phytoplanktonspecies composition and succession will also impact otherbiogeochemical cycles, such as the nitrogen, opal, andsulphur cycles, which in turn is bound to feedback onclimate. Despite the potential importance of globalchange-induced biogeochemical feedback, our under-standing of these processes is still in its infancy. It is nowbecoming increasingly clear, however, that the assump-tion of a constant oceanic biosphere, implicitly used inmany assessments of future global change, is no longerviable.

6. OutlookIn view of the rapid changes in environmental con-

ditions presently occurring on a global scale, a major chal-lenge in earth system sciences is to predict biosphericresponses to global change. Any attempt to forecast bio-logically-mediated changes in marine biogeochemicalcycling critically depends on our ability to explain thedistribution and succession of the dominant planktongroups in relation to biotic and abiotic environmentalconditions. The importance of a fundamental understand-ing of marine ecosystem regulation becomes even morepressing in view of recent considerations to manage ormanipulate oceanic CO2 sequestration. While the ocean’scapacity to store anthropogenic CO2 is immense, its sen-sitivity to related changes in environmental conditions islikely to be high. In contrast to the terrestrial biosphere,marine ecosystems have evolved in a comparatively ho-mogenous environment under conditions of relative sta-bility over evolutionary timescales. Even subtle changesin the environmental conditions may therefore have strong

Fig. 7. The biological carbon pumps: Photosynthetic produc-tion of organic matter in the surface layer and its subse-quent transport to depth, termed organic carbon pump, gen-erates a CO2 sink in the ocean. In contrast, calcium carbon-ate production and its transport to depth, referred to as thecarbonate counter pump, releases CO2 in the surface layer.The relative strengths of these two processes (rain ratio)largely determine the biologically-mediated ocean atmos-phere CO2 exchange.

728 U. Riebesell

effects on ecosystem functioning, with yet unforeseeableconsequences for elemental cycling.

Considerable evidence is now accumulating and in-dicating that CO2-related changes in seawater carbonatechemistry can directly affect the marine biosphere. CO2-dependent effects on the marine plankton can range fromphysiological responses on the organismal level, changesin the ecosystem structure and regulation, to large scaleshifts in biogeochemical cycling. Their net effect on car-bon cycling ultimately depends on the relative importanceand sensitivity of each of these processes to global change.In addition, possible changes at the level of primary pro-ducers, such as those outlined above, are likely to beamplified, compensated or reversed by interference fromhigher trophic levels, which themselves are bound to re-spond to environmental change. Much of the informationpresently available is based on single parameter pertur-bation studies. Global change, however, simultaneouslymodifies a multitude of environmental conditions, whosecombined effects on the biosphere are largely unknown.Moreover, global change related responses of the marineplankton as described above have so far been observedonly in short-term perturbation experiments. To what ex-tent physiological acclimatization and genetic adaptationmay counteract adverse effects related to environmentalchanges is presently unknown.

In view of the inherent complexity of biological sys-tems, including marine pelagic systems, it seems impos-sible at this point to provide a comprehensive and reli-able forecast of large-scale and long-term biological re-sponses to global environmental change. Any responsi-ble consideration aiming to regulate or manipulate theearth system in an attempt to mitigate the green-houseproblem is presently hindered by the large gaps in ourunderstanding of earth system regulation. With regard tothe marine pelagic system, understanding the complexstructure and regulation of the marine plankton, their re-sponses to global environmental changes and the conse-quences thereof for marine biogeochemical cycling willbe a major challenge for marine sciences in the comingdecades.

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