Response of primary production and calcification to changes of pCO2 during experimental blooms of the coccolithophorid Emiliania huxleyi
28 July 2004
13 March 2005
Abstract
[1] Primary production and calcification in response to different partial pressures of CO2 (PCO2) (“glacial,” “present,” and “year 2100” atmospheric CO2 concentrations) were investigated during a mesocosm bloom dominated by the coccolithophorid Emiliania huxleyi. The day‐to‐day dynamics of net community production (NCP) and net community calcification (NCC) were assessed during the bloom development and decline by monitoring dissolved inorganic carbon (DIC) and total alkalinity (TA), together with oxygen production and 14C incorporation. When comparing year 2100 with glacial PCO2 conditions we observed: (1) no conspicuous change of net community productivity (NCPy); (2) a delay in the onset of calcification by 24 to 48 hours, reducing the duration of the calcifying phase in the course of the bloom; (3) a 40% decrease of NCC; and (4) enhanced loss of organic carbon from the water column. These results suggest a shift in the ratio of organic carbon to calcium carbonate production and vertical flux with rising atmospheric PCO2.
1. Introduction
[2] In the context of rising PCO2 in the atmosphere and concomitant increase of pCO2 in the oceans, the response of marine organisms and ecosystems to elevated pCO2 has received little attention compared to terrestrial plants and ecosystems. This is partly due to the fact that photosynthesis by marine phototrophs is, generally, not considered to be carbon limited due to the large pool of DIC in seawater, mostly in the form of bicarbonate. Indeed, if some studies have shown that marine autotrophic communities are often insensitive to pCO2 changes, several studies have shown that some seagrasse [Zimmerman et al., 1997], macroalgae [Gao et al., 1993], diatom [Riebesell et al., 1993; Chen and Durbin, 1994], coccolithophorid [Nimer and Merrett, 1993; Hiwatari et al., 1995; Riebesell et al., 2000; Zondervan et al., 2001], and cyanobacteria [Qiu and Gao, 2002] species exhibit higher rates of photosynthesis under CO2 enrichment. In their review of the effects of CO2 concentration on marine plankton, Wolf‐Gladrow et al. [1999] pointed out the apparent discrepancy between ample CO2 supply from the bulk medium combined with the capacity for direct utilization of HCO3− in many marine phytoplankton on the one hand and the sensitivity of both phytoplankton growth rate and elemental composition to CO2 concentration on the other. They argue that one of the factors to be considered when trying to resolve this discrepancy is the low affinity for CO2 of the primary carboxylating enzyme RuBisCO and that the sensitivity of marine phytoplankton to CO2 is best viewed as a co‐limitation of CO2 in concert with light availability and other limiting factors such as nutrients. Irrespective of which mechanism is responsible for the sensitivity of some phytoplankton species to pCO2, the topic has received comparatively little attention considering its potential importance for carbon export and sequestration and its potential negative feedback to rising atmospheric CO2.
[3] The response of calcifying organisms and communities to elevated pCO2 appears to be more straightforward. Biogenic precipitation of calcium carbonate is generally described by the following equation:
[4] Thus calcification acts as a source of CO2 to the water column and counteracts the photosynthetic uptake of CO2. The net effect of these two antagonistic processes on CO2 is dependent on the ratio of NCC to NCP (C:P). Moreover, the release of CO2 by NCC is modulated by the buffering capacity of the carbonate system [Frankignoulle, 1994]. Any variation of the C:P ratio will have significant implications on biogeochemical fluxes, and particularly on the sign and strength of the overall feedback of coccolithophorids to rising atmospheric PCO2. Furthermore, changes of the C:P ratio affect the density and sinking rate of coccolithophorid cells and debris and, therefore, the magnitude of carbon export. Hence Buitenhuis et al. [2001] suggested that a decrease of the C:P ratio induces a decrease of carbon export and affects the overall uptake of CO2 by coccolithophorids.
[5] Generally, the rate of calcification decreases with rising pCO2 and diminishing CaCO3 saturation state (Ω). This response is nowadays well documented for reef‐building corals and coralline algae [Gattuso et al., 1999]. Mesocosm experiments have recently established the link between the decrease in calcification rate of coral reefs with rising CO2, and the concomitant drop of aragonite Ω [Langdon et al., 2000; Leclercq et al., 2002; Langdon et al., 2003]. In the same way, Bijma et al. [1999] observed that the decrease of Ω causes a decrease in the calcification rate by foraminifera. Riebesell et al. [2000] and Zondervan et al. [2001, 2002] observed a similar response with coccolithophorids, which may be the largest contributors to marine pelagic calcification. These authors showed, in batch cultures of Emiliania huxleyi and Gephyrocapsa oceanica, that the production of particulate organic carbon (POC) increases with increasing CO2 and is additionally depending on the total irradiance and the photoperiod length. Also, the production of particulate inorganic carbon (PIC) decreases, leading to the decrease of the C:P ratio. It therefore seems that depression of calcification or/and C:P ratio at elevated pCO2 is a general feature among marine calcifying organisms.
[6] Previous studies of the response of coccolithophorids to increasing pCO2 were most often made in batch cultures by manipulating the carbonate system through the addition of acid or base [Nimer and Merrett, 1993; Buitenhuis et al., 1999; Zondervan et al., 2001, 2002]. The aim of identifying the influence of increasing CO2 on calcification has driven these authors to eliminate environmental interactions and to grow cultures under optimal conditions which do not perfectly reflect in situ conditions. Furthermore, the manipulation of the carbonate system through the addition of acid and base also alters TA, whereas oceanic TA will not change significantly in the next decades. For instance, an increase of pCO2 induced by acid addition leads to a decrease in calcite and aragonite saturation states 20% higher and an increase in HCO3− concentration 60% lower than would be observed for the same increase of pCO2 induced by CO2 addition. However, calcification appears to be controlled by Ω than by pCO2 per se [Langdon et al., 2000; Leclercq et al., 2002]. Moreover, HCO3− was suggested to be the substrate for calcification of E. huxleyi [Buitenhuis et al., 1999]. Therefore the control of the carbonate system using gas addition may be more suited to reproduce the future changes of carbonate chemistry than acid/base addition techniques.
[7] Building on the pioneering batch culture experiments, the aim of the present study was to follow the development and decline of a bloom of a natural plankton community dominated by the coccolithophorid E. huxleyi exposed to various pCO2 under more natural conditions in large seawater volumes. This allowed us to investigate the pCO2 related effects at the community level and to examine their impacts on the dynamics of the bloom. This was achieved by employing large seawater enclosures containing natural assemblages of bacterioplankton, phytoplankton, and micro‐zooplankton kept under ambient light and temperature conditions and subjected to atmospheric CO2 concentrations simulating the “glacial,” “present,” and “year 2100” atmospheric PCO2 conditions (respectively, 180, 370, and 700 ppmV). Inorganic nutrient concentrations and the carbonate chemistry were adjusted prior to the onset of the bloom, and were allowed to evolve without further regulation as would occur in the mixed surface layer of a stratified water column during the course of a bloom. The day‐to‐day response of inorganic and organic carbon production by the enclosed communities was assessed using O2 production and 14C uptake during incubations together with the monitoring of daily changes in TA and DIC.
2. Material and Methods
2.1. Overall Description of the Experiment
[8] The experiment was carried out between 31 May and 25 June 2001 at the Marine Biological Field Station (Raunefjorden, 60.3°N, 5.2°E) of the University of Bergen, Norway. Nine enclosures made of polyethylene bags of 2 m diameter and volume of 11 m3 were used. The bags were secured to the sides of a raft equipped with a small laboratory. Each bag was filled with unfiltered nutrient poor (post‐spring bloom) water pumped at a depth of 2 m in the fjord on 1 June. The following provides a short description of the experimental setup, for more details we refer to a companion paper by Engel et al. [2004a].
[9] The tops of the mesocosms were covered with tetrafluoroethylene films (95% transmission for photosynthetically active radiation) forming a tent over more than 90% of the mesocosm surface area. The atmospheric PCO2 underneath the tents was controlled by injecting a continuous stream of gases with a known CO2 content. Three levels of pCO2 (180, 370, and 700 ppmV) were used with three replicates each; they will be referred to as glacial (Mesocosm (M)7, M8, and M9), present (M4, M5, and M6) and year 2100 (according to the Intergovernmental Panel on Climate Change “business as usual” scenario IS92a; M1, M2, and M3). In contrast to similar laboratory experiments carried out on phytoplankton cultures, in which seawater pCO2 was maintained at constant values throughout the experiment, in this study, seawater CO2 was manipulated only at the start of the experiment before initiation of the bloom. This was achieved by bubbling CO2‐free, ambient, or CO2‐enriched air at the bottom of the mesocosms until 6 June (hereinafter referred to as “day 0”). From day 0 until the end of the experiment, the carbonate system was allowed to evolve naturally while maintaining only the atmosphere underneath the tents covering each mesocosm at glacial, present, and year 2100 atmospheric PCO2 conditions. The seawater pCO2 values reached on 6 June are given in Table 1. To promote the development of the coccolithophorid bloom, nitrate and phosphate were added to each mesocosm on day 0 at initial concentrations of about 17 μmol L−1 NO3− and 0.5 μmol L−1 PO43−. The water enclosed in the mesocosms was gently homogenized using an airlift system consisting of a plastic tube in which a gas stream, having a PCO2 identical to that of the atmosphere confined above the mesocosm, produced an upward motion of water in the tube. The flow rate was very low in order to avoid significant gas exchange between the air stream and seawater.
| Mesocosm | d0 | d10 | d15 | |||
|---|---|---|---|---|---|---|
| pCO2, ppmV | [CO2], μmol kg−1 | pCO2, ppmV | [CO2], μmol kg−1 | pCO2, ppmV | [CO2], μmol kg−1 | |
| Year 2100 | ||||||
| M1 | 710 | 31.7 | 542 | 23.1 | 293 | 12.3 |
| M2 | 709 | 31.6 | 557 | 23.8 | 323 | 13.6 |
| M3 | 720 | 32.1 | 604 | 25.1 | 317 | 13.3 |
| Present | ||||||
| M4 | 407 | 18.2 | 360 | 15.4 | 217 | 9.1 |
| M5 | 426 | 19.1 | 344 | 14.7 | 228 | 9.7 |
| M6 | 408 | 18.3 | 341 | 14.5 | 205 | 8.6 |
| Glacial | ||||||
| M7 | 188 | 8.4 | 176 | 7.5 | 118 | 4.9 |
| M8 | 192 | 8.6 | 185 | 7.9 | 125 | 5.3 |
| M9 | 190 | 8.5 | 185 | 7.9 | 128 | 5.4 |
[10] Each mesocosm was sampled every day between 0900 and 1100 local time (LT), while daylight lasted from around 0430 to 2300 LT, to measure the abundances of various phytoplankton groups, bacteria, and viruses and the concentration of nutrients, as well as parameters of the carbonate system (TA and pCO2). Elapsed time is referred to as “dx” where x is the number of days since “d0”.
2.2. Seawater Partial Pressure of CO2 and Temperature
[11] Measurements of pCO2 were carried out using an equilibrator coupled to an infrared gas analyzer (IRGA, Li‐Cor® 6262). Seawater flows into the equilibrator (3 L min−1) from the top, and a closed air loop (3 L min−1) ensures circulation through the equilibrator (from the bottom to the top), a desiccant (Drierite®), and the IRGA [Frankignoulle et al., 2001]. The barometric pressure inside the equilibrator was kept equal to atmospheric pressure. Both the barometric pressure and temperature were monitored in the air loop. The IRGA was calibrated daily with air standards with nominal mixing ratios of 0, 350, and 800 ± 0.3 ppmV of CO2 supplied by Air Liquide Belgium® and Hydrogas®. The equilibration time of the system was less than 3 min [Frankignoulle et al., 2001]. The system was kept running twice this time before recording and averaging the values given by the IRGA and temperature sensors over a 30‐s period.
[12] In situ and equilibrator temperatures were measured simultaneously using Li‐Cor® sensors. Differences in temperature were less than 0.5°C. TA measurements were made with each measurement of pCO2 and used to temperature correct pCO2 using dissociation constants of Roy et al. [1993]. The uncertainty of pCO2 is estimated to ±3 ppmV.
2.3. Total Alkalinity, Dissolved Inorganic Carbon, and Salinity
[13] TA was measured using the classical Gran potentiometric method [Gran, 1952] on 100‐mL GF/C filtered samples. The reproducibility of measurements was ±3 μ mol kg−1.
[14] Dissolved inorganic carbon (DIC) was calculated from pCO2 and TA. CO2 speciation was calculated using the CO2SYS Package [Lewis and Wallace, 1998], the CO2 acidity constants of Roy et al. [1993], the CO2 solubility coefficient of Weiss [1974], and the borate acidity constant of Dickson [1990]. The total borate molality was calculated using the Uppström [1974] boron to salinity ratio. The uncertainty on the DIC computation is estimated to ±5 μmol kg−1.
[15] TA was corrected for the drawdown of nitrate and phosphate associated with phytoplankton nutrient utilization. According to the classical Redfield‐Ketchum‐Richards reaction of biosynthesis [Redfield et al., 1963; Richards, 1965],
[16] DIC changes were corrected daily for air‐sea exchange of CO2 using the air‐sea gradient of CO2, the volume of the bags, and assuming that water in the bags was well homogenized and that there was zero wind under the tents. Air‐sea exchange of CO2 from the enclosed atmosphere to the water was computed using the algorithm for stagnant boundary layer thickness from Smith [1985], molecular diffusivity from Jähne et al. [1987], and chemical enhancement model from Hoover and Berkshire [1969]. The formulation given by Smith [1985] was established using the stagnant film model and measurements from wind tunnels at low wind speed and corresponds better to our experimental setup than other relations derived from in situ measurements affected by additional turbulent processes such as currents and rain. Correction for air‐sea exchange accounted for less than 5% of changes in DIC.
[17] Normalized TA and DIC at a constant salinity (S = 31) are denoted as TA31 and DIC31. Salinity was measured using a conductivity‐temperature‐pressure sensor (CTD SAIV A/S, model SD204).
2.4. Flow Cytometry Sample Processing and Analysis
[18] Analyses were performed with a FACSCalibur flow cytometer (Becton Dickinson®) equipped with an air‐cooled laser providing 15 mW at 488 nm and with standard filter setup. The algae were analyzed from fresh samples at a high flow rate (∼70 μL min−1) with the addition of 1 μm fluorescent beads (Molecular Probes®). Autotrophic groups were discriminated on the basis of their forward or right angle light scatter (FALS, RALS) and chlorophyll (and phycoerythrin for Synechococcus and cryptophyte populations) fluorescence. Enumeration of viruses was carried out on samples fixed with glutaraldehyde (0.5% final concentration) and frozen (in liquid nitrogen). Once thawed at 37°C, samples were diluted 10 to 100 times in Tris‐EDTA (pH = 8) buffer and heated for 10 min at 80°C after staining with the DNA dye SYBR® Green I (1/20,000 final concentration, Molecular Probes®, [Marie et al., 1999]). Counts were performed at medium rate (∼30 μL min−1). Viruses were discriminated on the basis of their RALS versus green DNA‐dye fluorescence. Listmode files were analyzed using CYTOWIN [Vaulot, 1989] (available at http://www.sb-roscoff.fr/Phyto/cyto.html#cytowin) and WinMDI (version 2.7, Trotter, available at http://www.bio.umass.edu/mcbfacs/flowcat.html#winmdi).
2.5. Primary Production From 14C Incubations
[19] Subsurface seawater for incubation experiments was sampled in M1, M4, and M9 before sunrise. All water samples were pre‐sieved through a 200‐μm nylon mesh to remove large zooplankton. All incubations were carried out in 60‐mL flasks inoculated with H14CO3 (∼20 μCi per 500 mL) and incubated in situ for 24 hours at a depth of 1.5 m in the water of the fjord adjacent to the mesocosms. Concomitant incubations were made in dark bottles. After incubation, samples were filtered on Whatman® GF/F filters under gentle vacuum. Duplicate filters were collected for each sample incubated and rinsed with 0.2‐m‐filtered seawater in order to remove excess DI14C. One set of filters was treated with 100 μL HCl (0.01 N) to eliminate the radiocarbon incorporated into CaCO3. Primary production was estimated from 14C measurements after exposure of the filters to HCl while calcification was estimated by subtracting primary production from the total 14C collected on untreated filters.
2.6. Net Community Production and Respiration (O2 Technique)
[20] Samples were collected in mesocosms M1, M2, M4, M5, M8, and M9 before sunrise and immediately distributed into 60‐mL BOD bottles (overflowing > 150 mL). For each sampled mesocosm, four bottles were fixed immediately with Winkler reagents, three sets of three bottles were incubated in situ nearby the mesocosms at 0.5, 1.5, and 4 m, and four bottles were incubated in the laboratory in darkness at in situ temperature. The bottles incubated in situ were fixed at sunset, and duration of the incubations was about 18 hours (from about 0510 LT until 2300 LT). The dark bottles incubated in the laboratory were fixed the next day between 0800 and 1000 LT, and the duration of the incubations was 27 to 29 hours.
[21] The concentration of dissolved oxygen was determined using an automated Winkler titration technique with a potentiometric end‐point detection [Anderson et al., 1992] using an Orion® 9778‐SC electrode. Reagents and standardizations were similar to those described by Knap et al. [1994].
2.7. Net Community Production and Net Community Calcification From DIC and TA Changes
[22] NCC can be estimated from the time course of TAcorrected according to
3. Results
3.1. E. Huxleyi Abundance, pCO2, DIC31, and TA31
[23] A succession of distinct phytoplankton assemblages took place in the course of the experiment. The assemblage was first dominated by Synechococcus sp. and nanoflagellates (S. Jacquet, unpublished data, 2001), and subsequently by E. huxleyi [Engel et al., 2004a]. From d0 to d7, both Synechococcus sp. and nanoflagellate abundances increased to reach a maximum between d5 and d7, depending on the mesocosm, with abundances ranging from 10 to 15 × 103 cells mL−1 and 30 to 170 × 103 cells mL−1 for Synechococcus sp. and nanoflagellates, respectively (S. Jacquet, unpublished data, 2001). This maximum cell abundance corresponded to chlorophyll‐a (Chl a) concentrations ranging from 0.3 to 0.5 μg L−1 [Engel et al., 2004a]. The abundance of Synechococcus sp. and nanoflagellates subsequently decreased sharply to less than 3.2 and 7.0 × 103 cell mL−1, respectively on d10, while E. huxleyi abundance remained below 2.7 × 103 cell mL−1 from d0 to d10 (Figure 1). During this coccolithophorid pre‐bloom phase, small decreases of DIC31 and PCO2 were observed while TA31 remained unchanged with similar values in all mesocosms.
[24] On d10, the decreases of DIC31 and pCO2 were larger and concomitant with the sharp increase of the abundance of E. huxleyi. The minimum pCO2 was observed on d16. The magnitude of changes in pCO2 exhibited larges differences depending on the pCO2 conditions from d0 to d14 ranging from 386 ± 16 ppmV in the year 2100 mesocosms to 187 ± 13 ppmV in the present mesocosms and 61 ± 2 ppmV in the glacial mesocosms. However, until d14, the differences in DIC31 patterns between mesocosms were not significant. From d12 onward, TA31 decreased sharply in all mesocosms indicating the onset of calcification by coccolithophorids. Decreases of TA31 ranged from 30 to 95 μmol kg−1 between d13 and d14.
[25] Between d14 and d16, the overall patterns of most parameters indicate the transition toward the decline of the bloom. On d14, nutrients (NO3− and PO43−) were exhausted in all mesocosms [Engel et al., 2004a]. From d16 onwards, pCO2 remained constant or increased slightly whereas DIC31 continued to decrease in most mesocosms.
[26] DIC31 evolutions became confusing after d14. DIC31 reached rapidly a plateau on d14 in M5, M6 and M8 (glacial mesocosm) whereas it continued to decrease significantly in M1 and M3, M4, M7, and M9 (year 2100 mesocosms). Concomitantly, TA31 also began to differ greatly between mesocosms. In most cases, TA31 reached a plateau after a large and continuous drop. However, it should be noted that TA31 decreased at a high rate until the end of the experiment in M1 and M3 (year 2100 conditions).
[27] Interestingly, within these two mesocosms, viruses specific to E. huxleyi were either absent or present in low abundance. The collapse of nutrient‐induced E. huxleyi blooms, as we observed in 7 of the 9 mesocosms, has also been commonly reported in similar experiments. It has been attributed to viral lysis by a virus identified as EhV [Bratbak et al., 1996; Jacquet et al., 2002; Castberg et al., 2001] which belongs to the genus Coccolithovirus proposed by Schroeder et al. [2002] within the family of algal viruses Phycodnaviridae.
3.2. Net Community Production of Organic Carbon
[28] NCP (Figure 2a) assessed from oxygen incubations 
and 
normalized against Chl a concentrations (Figure 2c), i.e., the net community productivity 
exhibited similar patterns for all the mesocosms during the pre‐bloom period (d0 to d10) and the peak of the bloom (d10 to d15). They increased sharply from d7 onward to reach maximum values on d12 coinciding with the marked increase of abundance of E. huxleyi (Figure 1). From d16 onward, mesocosms showed different 
patterns which do not seem to be related to PCO2 conditions.
(b) net primary production estimated from 14C incubations 
(c) net community production normalized against chlorophyll a from oxygen incubations 
(d) net primary production normalized against chlorophyll a from 14C 
and (e) community respiration (CR) based on oxygen incubations in mesocosms 1 (solid squares, solid line), 2 (open squares, solid line), 4 (solid triangles, dashed line), 5 (open triangles, dashed line), 8 (solid circles, dotted line), and 9 (open circles, dotted line).
[29] Net primary production (Figure 2b) estimated by the uptake of 14C 
exhibited a similar pattern to those of 
From d14 onward, 
appears to be lower in M1 (year 2100) compared to M4 (present) and M9 (glacial). However, similarly to 
(Figure 2d) was similar in all the mesocosms from d9 onward.
[30] Community respiration (Figure 2e) assessed from oxygen incubations decreased during the pre‐bloom period, then increased during the peak of the bloom to reach a maximum value ranging from 40 to 75 mmol O2 m−2 d−1 on d14. It subsequently remained between 30 and 60 mmol O2 m−2 d−1. No obvious differences of community respiration were found in the various PCO2 conditions.
3.3. Ca14CO3 Production Rate
[31] Calcification by E. huxleyi estimated from 14C in situ incubations started on d9 in M4 and M9 and on d11 in M1 (Figure 3a), suggesting that the onset of calcification was delayed under year 2100 PCO2 conditions. The highest values were reached on d13 in M4 and M9 (16 and 20 μmol CaCO3 kg−1 d−1, respectively). Calcification was higher in M9 than in M4, while it remained low in M1. This is consistent with the larger decreases of TA31 observed in M4 and M9 compared to M1. During the peak of the bloom, M1 exhibited lower rates of calcification normalized to Chl a than M4 and M9 (Figure 3b), suggesting a lower competence of E. huxleyi to calcify under year 2100 PCO2 conditions, while net primary productivity seemed unaffected (see section 3.2).
3.4. Molar Respiration Ratio
[32] We compared net community production values obtained from O2 incubations with values estimated from DIC and TA changes by plotting 
versus NCPDIC (Figure 4). Model II linear regression gives a slope of 1.45 ± 0.12 and a correlation coefficient of 0.68 (p < 0.0001, n = 54, 4). The slope of the linear regression corresponds to the so‐called molar respiration ratio (R.R.). Indeed, if we consider the respiration as the reverse of the Redfield‐Ketchum‐Richard equation (equation (2)), the complete oxic degradation of phytoplankton theoretically requires 138 moles of dissolved O2/106 moles of organic carbon (C) leading to a molar respiration ratio (O2/C) of 1.30. The R.R. obtained in the present study agrees well with the estimates of Hedges et al. [2002] based on phytoplankton elemental composition using nuclear magnetic resonance which ranges from 1.41 to 1.47, depending on the geographic area considered. The consistency of these results validates the use of NCPDIC derived from DIC and TA and enables therefore a comprehensive day‐to‐day comparison of NCP and NCC.
versus NCP computed from DIC and TA (NCP DIC) with a model II regression line.
3.5. Timing of Organic and Inorganic Carbon Production
[33] NCPDIC as a function of NCC are shown in the Figure 4 where each point corresponds to a daily measurement. Connecting day‐to‐day estimates provides an overview of the temporal evolution of NCPDIC relative to NCC. During the pre‐bloom period, NCPDIC increased steadily (upward displacement along the Y‐axis), first owing to the rising abundances of Synechococcus sp. and nano‐flagellates, and subsequently to the onset of the bloom of E. huxleyi. The increase of NCPDIC was enhanced from d10 onward in all the conditions, concomitantly with the beginning of the peak of the bloom period (Figure 1) and leads to maximum values of NCPDIC on d12 and d13. By d15 the nutrients were exhausted; NCPDIC decreased markedly. NCC increases (displacement to the right along the X‐axis) in a second phase, when the coccolithophorid bloom is well underway, proceeded from d11 to d19, and remained at a high level while NCPDIC decreased dramatically, which is consistent with observations in cultures of E. huxleyi [Dong et al., 1993]. The third phase was the collapse of the bloom with a dramatic decrease of both NCPDIC and NCC. This phase corresponds to the period during which the coccolithovirus abundance passes over a threshold value, estimated to be around 5.106 part mL−1 (dashed lines in Figure 5). At the end of the experiment, NCPDIC and NCC show negative values due to elevated respiration and CaCO3 dissolution, as suggested by Milliman et al. [1999]. This is consistent with the increase of both CR (Figure 2c) and bacterial abundance determined by flow cytometry (S. Jacquet, unpublished data, 2001).
[34] Under glacial conditions (M7, M8 and M9), NCC started at the onset of the peak of the bloom (d10–d11) and then increased steadily in parallel to NCPDIC leading to an almost simultaneous maximum (only 1 day time lag). In contrast, in the year 2100 conditions (M1, M2 and M3), NCC began later (d12 to d13), and suddenly, while NCPDIC had already reached its maximum. NCC subsequently increased very rapidly while NCPDIC was decreasing. The present conditions exhibited an intermediate behavior between the year 2100 and glacial conditions: The maximum level of NCPDIC was reached when NCC was already substantial but had not reached its maximum value.
[35] Thus, if the overall pattern of NCPDIC prior to viral lysis is similar for all the conditions, the onset of NCC occurs sooner in the glacial and present conditions than in the year 2100 conditions. This is consistent with the calcification rates measured with the 14C in situ incubations. Furthermore, under glacial conditions, NCC increases steadily from the very beginning of the peak of the bloom in parallel to the exponential rise of NCPDIC, while in the year 2100, NCC occurs suddenly at the maximum of NCPDIC.
3.6. Mean Community Production and Calcification Rates
[36] Changes in the standings stocks of total organic carbon (TOC) and PIC, associated with the bloom of E. huxleyi, were assessed during the peak of the bloom (d10–d15) by integrating, respectively, daily NCPDIC and NCC over time (Figure 6). TOC standing stocks exhibit an almost linear evolution prior to the exhaustion of PO43− and NO3− allowing the computation of mean TOC production rates. During this period, bacterial abundance was low and no phytoplankton species other than E. huxleyi were present in significant numbers (S. Jacquet, unpublished data, 2001), so TOC changes are mainly due to primary production by coccolithophorids. Likewise, since changes of PIC are linear from d11 until the coccolithovirus passed a threshold value of about 5 106 part. mL−1, it is possible to compute the mean PIC production rates (Figure 7) prior to viral lysis.
[37] TOC increased steadily from d10 to d15 in all the mesocosms at rates ranging from 21.4 to 25.9 μmolC kg−1 d−1 (Figure 6). TOC production rates are similar under the three PCO2 conditions (Figure 7). In contrast, the mean rate of PIC production (Figure 7) was conspicuously lower in the year 2100 conditions (10.3 ± 3.2 μmol C kg−1 d−1) than in the present (17.9 ± 4.4 μmolC kg−1 d−1) and in the glacial (17.9 ± 2.5 μmol C kg−1 d−1) conditions. The mean PIC/TOC production ratio (C:P ratio) of E. huxleyi is similar in the glacial and present conditions between 0.73 and 0.78, but fell to 0.45 in the year 2100 conditions.
3.7. Carbon Losses
[38] Carbon losses were calculated as the difference between TOC produced by photosynthesis (estimated from the integration of NCPDIC) and the accumulation of POC in the water column (data from Engel et al. [2004a]). Carbon losses during the peak of the bloom were conspicuously higher in the year 2100 (48 ± 10 μmol C kg−1 d−1) than under the glacial conditions (25 ± 16 μmol C kg−1 d−1) while TOC production remained similar (Figure 7). If integrated over the d1–d15 period, this trend is enhanced, carbon losses being more than twice as high in the year 2100 (74 ± 14, μmol C kg−1 d−1) than under glacial conditions (34 ± 16 μmol C kg−1 d−1)(data not shown).
4. Discussion
4.1. The pCO2 Changes and Buffering Effect of the Carbonate System
[39] Until d14, changes in DIC31 followed the same trend and the same magnitude in all mesocosms while changes in pCO2 were conspicuously different between the three conditions right from the beginning of the experiment (Figure 1). Indeed, the magnitude of changes in pCO2 from d0 to d14 was 6 times higher in the year 2100 conditions than in the glacial conditions, while until d14, differences in DIC31 patterns between mesocosms were not significant since NCPDIC was roughly similar under the three PCO2 conditions and NCC was negligible. Therefore differences in the magnitudes of pCO2 changes with regard to the PCO2 conditions are not ascribable to biological processes and must be explained taking into account the buffering effect of the carbonate system.
[40] The chemical buffer factor (β = ΔpCO2/ΔDIC) describes the change in pCO2 relative to the DIC change induced by an input/output of dissolved CO2. It results from equilibrium dissociation reactions of the carbonate system and is a function of several physico‐chemical conditions, among them on the pCO2 itself. The evolution of the buffer factor, calculated for the initial conditions of the experiment, is given in Figure 8. The β increases from 9.6 in the glacial conditions to 16.6 in the year 2100 conditions. Subsequently, for the same removal of CO2 by primary production, the consequent decrease of pCO2 is 6 times higher in the year 2100 condition (ΔpCO2 = 116 ppmV for ΔDIC = 20 μmol kg−1 and β = 16.6 at pCO2 = 700 ppmV, and DIC = 2000 μmol kg−1) than under the glacial conditions (ΔpCO2 = 20 ppmV for ΔDIC = 20 μmol kg−1 and β = 9.6 at pCO2 = 180 ppmV and DIC = 1740 μmol kg−1).
[41] Thus, owing to thermodynamic interactions of the carbonate system, the change in pCO2 is significantly higher in the year 2100 condition than in the other conditions, even though the process originally responsible of these pCO2 changes, i.e., the uptake of CO2 by photosynthesis, appears to be roughly similar under the three PCO2 conditions. Hence one can note that in the future CO2 rich world, other processes, such as temperature oscillations, upwelling of CO2 rich waters, or precipitation of calcium carbonate [Frankignoulle et al., 1994] among others, will also contribute to the thermodynamic enhancement of the amplitude of pCO2 changes from daily to seasonal timescales. In the same way, a higher spatial heterogeneity of pCO2 can be expected from local to global scales.
[42] At the end of the experiment, from d16 onward, pCO2 remained constant or slightly increased whereas DIC31 continued to decrease in most mesocosms. The decoupling of pCO2 and DIC31, associated to a decrease of TA31, indicates the larger influence of NCC compared to NCP on pCO2 as observed in natural coccolithophorid blooms and mesocosm experiments [Robertson et al., 1994; Purdie and Finch, 1994; Buitenhuis et al., 1996].
4.2. Primary Production and Carbon Export
[43] The effect of pCO2 on growth, productivity and calcification of the coccolithophorid E. huxleyi is still a matter of debate. In our study, no conspicuous changes in primary productivity (both 
and 
related to PCO2 conditions were observed during the peak of the E. huxleyi bloom. Differences in the NCP and NPP observed during the bloom decline should be ascribed to the occurrence or not of viral lysis rather than to some PCO2 related effects.
[44] On the other hand, evidence of a higher loss of particulate organic carbon from the water column under year 2100 conditions (Figure 7) emerges surprisingly. Since DOC concentrations were similar under all PCO2 conditions [Rochelle‐Newall et al., 2004], while grazing was negligible, enhanced carbon losses observed under year 2100 conditions are likely due to a higher rate of particle settling. In fact, when normalized to E. huxleyi cell concentration, TEP production was highest under year 2100 conditions [Engel et al., 2004a], consistent with previous observations of enhanced TEP formation under elevated CO2 [Engel, 2002]. TEP formation is typically seen as the result of carbon overproduction leading to exudation of polysaccharides by algal cells [Passow, 2002]. Aggregation of dissolved polysaccharides through a cascade of aggregation processes from the molecular scale up to the size of fast settling particles can lead to an enhancement of particle aggregation and subsequent export [Engel et al., 2004b]. The possible involvement of carbon‐rich TEP as a mediator of enhanced particle settling in the year 2100 mesocosms is further supported by higher C:N ratio of suspended particles under high CO2 conditions [Engel et al., 2004a].
4.3. Overview of the Response of the C:P Ratio to Rising CO2
[45] In contrast to NCPDIC, NCC exhibited a conspicuous decrease with increasing PCO2. These responses lead to a drastic decrease of the PIC/TOC production ratio (C:P ratio) under elevated PCO2. Also, calcification is delayed in the year 2100 condition, acting to reduce the overall amount of CaCO3 produced during the experiment.
[46] During the pre‐bloom period (d0 to d10) and the peak of the bloom (d10 to d15), NO3− and PO43− decreased rapidly from about 10 μmol L−1 and 0.7 μmol L−1, respectively, on d10 to below 0.4 μmol L−1 and 0.3 μmol L−1 on d15 with a mean photon flux density (PFD) around 650 μmol m−2 s−1 for 18:6 light period and a light attenuation coefficient at the bottom of the mesocosm (4 m) of about 80%. Therefore the mesocosms were not light limited but were in an intermediate state toward NO3− and PO43− depletion. When attempting to reconcile the results of the present study to experiments reported in the literature and to draw a comprehensive picture of the response of primary production and calcification of E. huxleyi to elevated pCO2, it is worth noting that the present knowledge is based on a mosaic of experiments with different growth conditions and involving several E. huxleyi ecotypes which give in some cases very different results (Table 2). However, it emerges that when irradiance is not drastically reduced [Zondervan et al., 2002], elevated pCO2 appears to be detrimental to calcification [Riebesell et al., 2000; Zondervan et al., 2001, 2002; Sciandra et al., 2003] (also this study), [Riebesell et al., 2000; Zondervan et al., 2001; Zondervan et al., 2002; Sciandra et al., 2003] and this generally leads to a decrease of the C:P ratio (Table 2). Such response of calcification to changes in seawater carbonate chemistry has also been observed in corals and foraminifera [Gattuso et al., 1998; Wolf‐Gladrow et al., 1999; Langdon et al., 2000; Leclercq et al., 2002; Langdon et al., 2003; Reynaud et al., 2003]. The cause of such a decrease of calcification by E. huxleyi in response to elevated pCO2 remains unclear. If it can be almost intuitive that a decrease of Ωcalc concomitant to an increase of oceanic pCO2 can reduce biogenic calcification, i.e., an environmental control of the calcification as it has been reported for corals reefs, some authors have rather suggested an internal control of calcification by E. huxleyi. For instance, several studies have suggested that calcification could support photosynthesis [Sikes et al., 1980; Nimer and Merrett, 1992; Anning et al., 1996; Buitenhuis et al., 1999], acting as an effective low‐cost energy pathway to directly supply the chloroplast with CO2 in addition to direct CO2 diffusion into the cell, and then raise the concentration of CO2 in the chloroplast at the site of photosynthesis. However, recent studies have severely questioned this hypothesis [Sekino and Shiraiwa 1994; Herfort et al., 2002, 2004; Rost and Riebesell, 2004]. Some other metabolic benefits from calcification have been suggested like the “trash‐can” function facilitating the use of HCO3− in photosynthesis (see Paasche [2002] and Rost and Riebesell [2004] for reviews) and serving to remove excess Ca2+ [Berry et al., 2002]. Furthermore, calcification could rid the cell of excess energy and therefore prevent damage of the photosynthetic machinery [Rost and Riebesell, 2004].
| E. huxleyi Strain | Nitrate Concentration, μmol L−1 | Irradiance, μmol m−2 s−1 | Organic Production | CaCO3 Production | C:P Ratio | |
|---|---|---|---|---|---|---|
| Riebesell et al. [2000] | subarctic North Pacific natural assemblages | in situ conditions | 30% of surface irradiance | 0[−] | − | − |
| Zondervan et al. [2002] after Riebesell et al. [2000] and Zondervan et al. [2001] | Plymouth | 100 | 150 (24/0), 150 (16/8) | + | − | − |
| Marine | 100 | 80 (24/0) | + | − | − | |
| Laboratory | 100 | 30 (24/0) and 80 (16/8) | + | 0 | − | |
| B92/11A | 100 | 15 (24/0) | 0 | 0 | 0[−] | |
| Sciandra et al. [2003] | Caen University TW1 | 0.5 | 170 (24/0) | − | − | 0 |
| This study | Norwegian natural assemblages | 0.3 to 10 | 150 to 650 (16/8) | 0 | − | − |
- a In the irradiance column, numbers in brackets denote the daily light/day period in hours. Here “+”,“−”, and “0” denote, respectively, increase, decrease, and no significant changes, while brackets denote a trend not statistically significant. We limit our comparison to experiments which addressed the response of E. huxleyi to increase of PCO2 or [CO2] and the concomitant decrease of the calcite saturation state (Ωcalc), within ranges of similar magnitude as predicted changes during the next hundred years.
[47] Some ecological implications have been also proposed. Hence coccolith production could protect the integrity of the cell and maintain a suitable environment around the cell surface [Young, 1994; Paasche, 2002]. However, among these ecological benefits, the ability of coccospheres to prevent viral lysis [Young, 1994] is not supported by our experiment. Indeed coccolithoviruses were detected in the course of the exponential growth phase in seven of the nine mesocosms, but virus‐induced lysis was not detected in the two mesocosms where calcification rates were the lowest (M1 and M3: year 2100 conditions). Thus the benefits of calcification for coccolithophores still remain an open question.
[48] Published data on the effect of elevated CO2 on organic carbon production and C:P ratio by E. huxleyi are even less clear (Table 2). Some papers report a decline of primary production at elevated pCO2 and others an increase, while no conspicuous change was observed in the present study. It is obvious that environmental parameters such as light and nutrients interact with pCO2 since they play a major role in the energy status and/or metabolism. Furthermore, two experiments carried out in similar conditions can show opposite trends regarding changes of POC and PIC production with pCO2 [Nimer and Merrett, 1992; Buitenhuis et al., 1999] depending on the strain used. This underlines that attention must be paid to the ecotypes used in experiments or encountered in natural assemblages as also noted in a recent review on E. huxleyi physiology [Paasche, 2002].
[49] Among this mosaic of contrasting results, it must be pointed out that the two experiments carried out with natural communities under irradiance and nutrient concentrations close to environmentally realistic conditions [Riebesell et al., 2000] (and the present study) [Riebesell et al., 2000] converge to show that PIC production and the C:P ratio decrease markedly while POC production remains roughly constant with rising pCO2.
[50] Previous experiments were carried out in batch or continuous cultures and provide little information on the dynamics of calcification in natural conditions. Following the development and decline of a bloom demonstrates that the onset of calcification was delayed by 24 to 48 hours in the year 2100 compared to glacial CO2 conditions. Unfortunately, since the bloom prematurely collapsed owing to a massive viral infection, it is not possible to assess the overall duration of the calcification phase. However, we surmise that the delay in the onset of calcification under high PCO2 could act to decrease the overall duration of the calcification phase. Such a reduction would lower the overall production of CaCO3 in the full course of a coccolithophorid bloom.
4.4. Implications of the Observed pCO2 Related Effects on Biogeochemical Fluxes
[51] The net effect of reduced calcification on air‐sea CO2 gradients and fluxes is the balance between two counteracting processes. First, the decrease in calcification reduces CO2 release. Second, the changes in seawater carbonate chemistry induced by rising pCO2 lead to an increase of the molar ratio of released CO2 over calcium carbonate precipitation [Frankignoulle et al., 1995]. These two antagonistic processes seem to be balanced in coral reefs [Gattuso et al., 1999], but Zondervan et al. [2001] suggested that the response of pelagic calcification leans toward a negative feedback that increases the retention of CO2 in the ocean. However, the response of pelagic biogenic CaCO3 fluxes to rising CO2 also needs to consider export processes. Blooms of E. huxleyi may be more efficient carbon sinks than other phytoplankton blooms as a result of the higher density of sedimenting cells and zooplankton fecal pellets, due to the high density of calcite [Buitenhuis et al., 2001]. This is supported by the compilation of sediment trap data below 1000 m depth which shows that ballast minerals, and in particular calcium carbonate, drives the sinking of organic carbon to the deep ocean [Armstrong et al., 2001; Klaas and Archer, 2002]. Hence, a lower C:P ratio of coccolithophorids under year 2100 conditions could lead to a smaller ballast effect and to a subsequent reduction of carbon export, thereby acting as a positive feedback to rising atmospheric CO2.
[52] In the present study, however, we actually observed the opposite, since carbon export by the E. huxleyi community, estimated as carbon losses, was higher under year 2100 conditions. Enhancement of carbon export through TEP production conspicuously overcomes the diminution of the ballast effect and turns the overall response of export of E. huxleyi to rising CO2 concentration toward a negative feedback. This comes in addition to the decrease of the production of CO2 as a consequence of the reduction of both rate and duration of calcification. Finally, if the enhancement of carbon export driven by higher TEP production under elevated pCO2 is as significant for other phytoplanktonic groups as for coccolithophorids, then it would potentially represent a major negative feedback on rising atmospheric CO2.
5. Conclusions
[53] No conspicuous change of both net community productivity and net primary productivity of E. huxleyi was detected during the peak and the decline of a bloom of the coccolithophorid E. huxleyi for pCO2 ranging from 175 to 600 ppmV. In contrast, the rate of net community calcification declined at elevated pCO2, corroborating the observations of Riebesell et al. [2000], Zondervan et al. [2001, 2002], and Sciandra et al. [2003] on cultures of E. huxleyi. Furthermore, the onset of calcification is delayed by 24 to 48 hours in the year 2100 conditions compared to glacial conditions. The decrease of calcification rate combined with a rather constant organic carbon production led to a significant decrease of the C:P ratio.
[54] When comparing previous reports on the response of organic and inorganic carbon production of E. huxleyi to increasing pCO2, it appears that in nonsaturating light and nutrient replete conditions, the increase in pCO2 promotes organic carbon production [Riebesell et al., 2000; Zondervan et al., 2001, 2002] unless pCO2 becomes too high (above 1000 ppmV) [Nimer et al., 1994]. In contrast, under nutrient‐limiting conditions, organic carbon production remains constant [Riebesell et al.,. 2000] (also the present study) or even decreases [Sciandra et al., 2003] with increasing pCO2.
[55] Since the changes of both organic and CaCO3 production with rising pCO2 are strongly influenced by light and nutrient conditions as well as by the level of pCO2, one could expect even more complex response of the C:P ratio. On the whole, a decrease of the C:P ratio with increasing pCO2 appears to be a general trend with the exceptions under severe light or nutrient limitations. However, it must be pointed out that in most studies, little attention has been paid to the ecotype of E. huxleyi or to species other than E. huxleyi. Information available to date [Zondervan et al., 2001, 2002] indicates differential responses depending on the species considered. Also, the interaction of increased temperature, which is significant in corals [Reynaud et al., 2003], has not been investigated in coccolithophorids. Any prediction of the future response of ocean biogeochemistry to elevated pCO2 must therefore take into consideration the composition of the community as well as the interaction with various other environmental parameters which are also predicted to change. In this sense, it is relevant to note that increased surface temperature will lead to a higher stratification and lower nutrient inputs, so that the evolution of marine communities and ecosystems should be envisaged in a “high CO2, low nutrient and warmer ocean” context.
[56] Many open questions need to be settled to predict reliably the response of coccolithophorids to rising CO2. However, it is worth noting that the two experiments carried out in environmentally realistic conditions converge satisfactorily and suggest that organic carbon production remains roughly constant with rising CO2 while inorganic carbon production decreases drastically, reducing concomitantly the C:P production. Moreover, a delay in the onset of calcification under elevated PCO2 conditions superimposed on a decrease of CaCO3 production rate is likely to reduce the overall production of CaCO3 in the course of a coccolithophorid bloom.
Acknowledgments
[57] This work is dedicated to our invaluable colleagues and friends Roland Wollast and Michel Frankignoulle who left us on 28 July 2004 and 13 March 2005, respectively. We are grateful to Anja Engel, Emma Rochelle‐Newall, and Antoine Sciandra for fruitful discussions, to Marie‐Dominique Pizay for technical assistance, to Jorun Egge for her welcome at the marine biological station of the University of Bergen, and to Daniel Delille and two anonymous reviewers for their pertinent comments on the manuscript. Access to installations from the University of Bergen was funded by the Improving Human Potential Programme from the European Union (contract HPRI‐CT‐1999‐00056 “Bergen Marine”). A. V. B. and M. F. were, respectively, a post‐doctoral researcher and a senior research associate at the Fonds National de la Recherche Scientifique. B. D and J. H. were supported by the Belgian Federal Office for Scientific, Technical and Cultural Affairs (contracts EV/12/7E and EV/11/5A, respectively). This is MARE contribution 062.
Notes :
- 1 28 July 2004
- 2 13 March 2005
- a In the irradiance column, numbers in brackets denote the daily light/day period in hours. Here “+”,“−”, and “0” denote, respectively, increase, decrease, and no significant changes, while brackets denote a trend not statistically significant. We limit our comparison to experiments which addressed the response of E. huxleyi to increase of PCO2 or [CO2] and the concomitant decrease of the calcite saturation state (Ωcalc), within ranges of similar magnitude as predicted changes during the next hundred years.
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