Ocean Uptake Potential for Carbon Dioxide Sequestration

Geochemical Journal, Vol. 39, pp. 29 to 45, 2005

Ocean uptake potential for carbon dioxide sequestration

MASAO SORAI* and TAKASHI OHSUMI**

Research Institute of Innovative Technology for the Earth (RITE), 9-2 Kizugawadai, Kizu-cho, Soraku-gun, Kyoto 619-0292, Japan

(Received June 26, 2003; Accepted May 7, 2004)

For the assessment of the long-term consequences of the carbon dioxide ocean sequestration, the CO2 injection into the middle depth parts of the ocean was simulated using a geochemical box model of the global carbon cycle. The model consists of 19 reservoir boxes and includes all the essential processes in the global biogeochemical cycles, such as the ocean thermohaline circulation, the solubility pump, the biological pump, the alkalinity pump and the terrestrial ecosys-

tem responses. The present study aims to reveal the effectiveness and consequences of the direct ocean CO2 sequestration P in relation to both lowering the atmospheric transient CO2 peak and reduction in future CO2 uptake potential of the ocean. We should note that the direct ocean injection of CO2 at the present time means the acceleration of the pH lowering in the middle ocean due to the eventual and inevitable increase of CO2 in the atmosphere, if the same amount of CO2 is added into the atmosphere-ocean system. The minimization of impact to the whole marine ecosystem might be attainable by the

direct ocean CO2 sequestration through suppressing a decrease in the pH of the surface ocean rich in biota. The geochemical implication of the ocean sequestration is such that the maximum CO2 amount to invade into the ocean, i.e., the oceanic P CO2 uptake potential integrated with time until the end of fossil fuel era, is only dependent on the atmospheric CO2 value in the ultimate steady state, whether or not the CO2 is purposefully injected into the ocean; we gave the total potential P capacity of the ocean for the CO2 sequestration is about 1600 GtC in the case of atmospheric steady state value ( CO2 ) of 550 ppmv.

P Keywords: ocean CO2 sequestration, global carbon cycle, CO2 uptake potential, box model, future atmospheric CO2

tralization effect provides a further oceanic potential to INTRODUCTION uptake excess CO2 (Nozaki, 1991). The detailed outline To mitigate the impact of the future atmospheric car- of the ocean CO2 sequestration has been presented both bon dioxide increase, the ocean CO2 sequestration has from technical and scientific aspects (Ohsumi, 1995). been proposed (Marchetti, 1977). The concept is built on The assessment of its long-term consequences requires the perspective that it would be hard to reduce drastically the comprehensive understandings on the global the future anthropogenic CO2 emissions without innova- biogeochemical cycle of the carbon. The modeling ap- tive technological developments. From a simple proach is effective in evaluation of each process govern- geochemical viewpoint, the ocean CO2 sequestration must ing the cycle. There have been so far several attempts to be one of the most reasonable options, because the ocean investigate the effects of the ocean CO2 sequestration and 2– originally contains both the carbonate ion CO3 and the its interaction with the biogeochemical processes. Hoffert – B(OH)4 ion enough to react with about 70% of the un- et al. (1979) showed clearly the “peak-shaving” effect of tapped fossil fuel carbon to form the bicarbonate ion the CO2 oceanic injection on the atmospheric CO2 con- – HCO3 and the boric acid B(OH)3 (Broecker, 2001). More- centration time profile in the future. In their work, some over, in the case of the CO2 sequestration to the bottom essential processes of the carbon cycle, such as the oce- ocean, the sedimentary calcium carbonate is expected to anic thermohaline circulation, the biological activities, – neutralize the injected CO2 to form HCO3 , and this neu- and the role of terrestrial ecosystem, were missing in the treatment. On the other hand, the EU-funded GOSAC (Global Ocean Storage of Anthropogenic Carbon) project, *Corresponding author (e-mail: [email protected]) which focused on improving the predictive capacity of *Presently at Mitsubishi Research Institute, Inc., 3-6, Otemachi 2- global-scale three-dimensional ocean carbon-cycle mod- Chome, Chiyoda-ku, Tokyo 100-8141, Japan. els, conducted a series of simulation runs where the an- **On leave from Central Research Institute of Electric Power Indus- thropogenic CO2 is injected into the several depth ranges try. of the selected points and revealed the relationship be- Copyright © 2005 by The Geochemical Society of Japan. tween the injection site and the subsequent holding ca-

29 Fig. 1. Global carbon cycle box model. The model consists of 19 reservoirs; the surface ocean is defined as 0–200 m depth, the middle part as 200–2,000 m depth, and the deep part as below 2,000 m. The fluxes between ocean boxes represent the thermohaline 6 3 –1 circulation and the numerical unit is sverdrups (1 Sv = 10 m sec ). Several key processes, such as the CO2 exchange between the atmosphere and the ocean, Fair-sea, the oceanic biological production, ONP, the terrestrial net primary production, TNP, the decomposition of the plant and soil, Fplant and FluxsoilT, are also shown by the arrows.

pacity in the ocean. Although these simulation studies MODEL DESCRIPTION provided several useful implications on the ocean CO2 sequestration, the fundamental role of the ocean has not The three-dimensional box model consists of 19 boxes yet presented explicitly: that is, what is the amount of as shown in Fig. 1: atmosphere, the Arctic Ocean (surface and middle), the North and South Atlantic Oceans CO2 to remain in the ocean? The CO2 uptake potential in the ocean is the key to understanding the effectiveness of (surface, middle and deep for each), the Indian Ocean (surface, middle and deep), the Pacific Ocean (surface, the oceanic CO2 sequestration. To address these questions, the global analysis like the Hoffert’s study is useful, al- middle and deep) and the Antarctic Ocean (surface and though the model should be revised based on the latest middle), and the land (plant and soil). We defined the knowledge on the carbon cycle system in the surface of surface ocean as 0–200 m depth, the middle part as 200– the earth. 2,000 m depth, and the deep part as below 2,000 m. The In this study, we give an estimate of the potential ca- exceptions are the Arctic and the Antarctic oceans, where the middle part is set from 200 m depth to the seafloor. pacity of the ocean for the CO2 sequestration based on the geochemical box model of the global carbon cycle. The thermohaline circulations between the ocean boxes Our present work owed much to the Wigley’s work, in are also given in Fig. 1. The solubility pump, the biologi- which the modeling approach allowed to explain the past cal pump, and the alkalinity pump are considered as the mechanism of carbon cycling in the atmosphere-ocean changes in atmospheric CO2 concentration and to predict its future trends (Wigley, 1993). His successful method- system. The contribution of terrestrial ecosystem, which ology including the treatment of the terrestrial carbon includes both CO2 fertilization and global warming ef- cycle system was taken into account in our previous model fects, is also taken into account. (Sorai et al., 1997): it enabled us to analyze the response For an n-box system, the increase rate of carbon amount M in the box i is expressed as of the ocean CO2 sequestration qualitatively and straight- forwardly. Box model approach is efficient in analyzing the effects on the atmospheric CO concentration and the dM n n 2 QQ, 1 influences on the oceanic carbon concentrations by the dT =-ÂÂijÆ + jiÆ () j =11j = purposeful injection of CO2 into the ocean. jiπ jiπ

30 M. Sorai and T. Ohsumi where Q stands for the carbon flux from the box i to triple all the other fluxes in accordance with the mass iÆj the box j, and Q , the carbon flux from the box j to the balance of the system. As a result, the “apparent” mean jÆi box i. The first term on the right hand side of Eq. (1) is residence time was lowered to about 1,000 years in our the total carbon efflux from the box i, whereas the sec- model. This modification is justified also by the repre- ond term is the total carbon influx to the box i. In our sentation of DIC in the pre-industrial steady state: a good model, the change in carbon amount in each reservoir is agreement was obtained by using our tripled flow pattern given by Eq. (1): the variables related to the carbon as discussed later. amount include the atmospheric CO2 concentration For additional justification of our modification, we ( P ), the carbon amounts of the terrestrial plant and performed sensitivity analyses changing the magnitude CO2 soil, and the concentrations of the dissolved inorganic of oceanic circulation, which enabled us to realize the carbon (DIC) and the dissolved organic carbon (DOC), relationship between the oceanic potential and the circu- and the alkalinity in each ocean boxes. An inverse Euler lation (see Appendix II). In fact, it was found that such a method was used to avoid the solution instabilities difference of the oceanic circulation had a little influence (Walker, 1991). The detailed model description includ- on the oceanic potential. ing the balance equations for each variable had been re- ported elsewhere (Sorai et al., 1997). In this section, the Solubility pump key processes of our model are presented (also see Ap- The solubility pump is defined as the carbon exchange pendix I). mediated by physical processes such as heat flux, advection, and diffusion. Net CO2 exchange flux across –2 –1 Thermohaline circulation the atmosphere-ocean interface, Fair-sea (mol m yr ), is The mass balance of water between the boxes were expressed as follows: established based on Schmitz’s compilation of the interbasin circulation of seawater (Schmitz, 1995). His bottom and deep layers in each oceanic basin are combined FKkPPair-sea= wa CO - CO ()22sea ()air to one box and the exchange of water mass between the {} surface Arctic and North Atlantic boxes is added in the EPCO PCO ,2 = 22sea - air () present treatment. {}()() It should be noticed that the diffusion process actu- where ( P ) and ( P ) are the CO partial pres- ally works for the mass and heat transports in the ocean CO2 sea CO2 air 2 in addition to the advection on the thermohaline circula- sures (ppmv) in the surface-water and in the atmosphere, –1 tion. Moreover, the wind-driven circulation enhances kw is the gas transfer rate (cm hr ), a is the solubility of –1 –1 these transports in the ocean on the time scale of several CO2 in seawater (mol L atm ), K converts units to mol –2 –1 decades. In our model, however, these processes between m yr , and E is the gas exchange coefficient (mol –2 –1 –1 the boxes are not taken into account, and hence the mass m yr matm ). Here, the solubility of CO2 is a function transport is somewhat weakened; the consideration of only of both temperature T (K) and salinity S (‰) (Weiss, Schmitz’s circulation provides the unrealistic distribution 1974): of DIC concentration, such that it becomes very high in deep region. The enhancement of the model oceanic cir- lna =-58 . 0931 + 90 . 5069◊ 100TT+ 22 . 2940 ln 100 culation would help to incorporate all these processes on () () 2 the time scale of several decades. In this case, one crite- +-STT0..027766 0 025888◊() 100+ 0 . 0050578◊() 100 . rion for the extent of the enhancement is the “apparent” [] residence time of the ocean, which is defined as the oce- ()3 anic volume divided by the total downward fluxes. At the North Atlantic Ocean in Schmitz’s model the downward The gas transfer rate is a function of wind speed and flux of 14 Sv (1 Sv = 106 m3sec–1), which corresponds to temperature. Several expressions for the relationship be- the “apparent” mean residence time of deep oceanic wa- tween gas transfer rate and wind speed have been used so ter more than 3,000 years, seems to be too low, if we fol- far. These studies have shown that gas transfer rates mea- low the generally accepted oceanic mean residence time sured over long time periods with variable winds are of about 1,000 years (Broecker and Peng, 1982). Consid- higher than those measured instantaneously or under ering that oceanic potential to hold the dissolved CO2 steady winds of the same average wind speed. For exam- species have the essential role in our simulations, the oce- ple, the low gas transfer value of Liss and Merlivat (1986) anic residence time should be primarily consistent with is based on 1–2 day measurements on a lake, whereas the the general view. Therefore, we tripled this downward higher value of Wanninkhof (1992) is based on ocean 14C flux to 42 Sv: this improvement inevitably made us to gas transfer data and long-term climatological winds. We

Ocean uptake potential for CO2 sequestration 31 use the expression for gas transfer rates by Wanninkhof in photosynthesis is transferred to the interior of the ocean (1992) because we mainly intend to predict long-term resulting in a temporary or permanent sequestration of trends. The Wanninkhof (1992) formulation is carbon. This process is controlled by the new production, where the organic carbon photosynthesized on the sur- 12 face layers of the ocean settles down to be decomposed kuSc039.,2 Sc - 4 w = ()0 () and regenerates the inorganic carbon again in the middle depth part of the box. We assumed that the magnitude of –1 where u is average wind speed (m s ), Sc is Schmidt the new production ONP (mol m–3yr–1) is limited by its number of CO2, the dimensionless quantity defined as the –3 phosphate concentration [PO4] (mol m ) and is control- kinematic viscosity divided by the diffusion coefficient, led by Michaelis-Menten kinetics: and the subscript 0 denotes a value at 20∞C. For a given gas, the Schmidt number varies with water temperature, decreasing as the temperature increases. We can use the RRcPO◊◊ 4 ONP= LC ◊ [], ()7 following relation for the temperature dependence of HRcPO+ ◊[]4 Schmidt number of CO2 (Liss and Merlivat, 1986): where LC is the latitudinally varying incident light factor 23 4(yr–1), Rc is the carbon to phosphate classical Redfield Sc=-+1860 120 T 4.... 3208 T - 0 09 T + 0 00079167 T ratio (=106), and R and H are adjustable parameters 5 () (Bacastow and Maier-Reimer, 1990). The ratio of the particulate organic carbon (POC) to the dissolved organic In fact, since it is difficult to set the average wind speed carbon (DOC) was adjusted to represent the DOC con- in each ocean, kw was adjusted to produce E values concentrations in each ocean consistent with the GEOSECS sistent with the estimations based on the satellite wind data. speed measurements (Etcheto et al., 1991). Consequently, There have been no attempts to incorporate the oce- the pre-industrial E values were 0.073 at the Arctic, 0.086 anic CO2 fertilization effect into a carbon cycle model so at the Antarctic, and 0.064 at the Atlantic, Indian and far, because the oceanic production rate is generally re- Pacific Oceans. garded to be nutrient-limited. However, some recent Furthermore, the global warming effect was also in- works revealed that at least several kinds of marine corporated in the solubility pump of our model. The tem- phytoplankton increased with an increase in P CO2 perature increase due to the greenhouse effect could be (Riebesell et al., 1993). Considering that our simulation related to the atmospheric P , as defined by CO2 treats a very large range of the atmospheric P , from CO2 280 ppmv to more than 1000 ppmv, it would be appropri- log P 280 ate to incorporate the oceanic CO2 fertilization effect; the CO2 DT =¥25. (), ()6 effect was taken into account in the model just in the same log 2 manner as in the case of the terrestrial ecosystem discussed later. We introduced the enhancement factor r0, where T (K) is the mean temperature increment relative which is defined as the ratio of ONP at P =680 ppmv D CO2 to the pre-industrial level (Stocker and Schmittner, 1997). to ONP at P =340 ppmv. Our incorporation of the fer- CO2 Equation (6) is based on the assumption that when the tilization effect in the model might represent the other atmospheric P reaches 560 ppmv, which is twice the CO2 anthropogenic impacts on the oceanic production; in the pre-industrial value, the global mean surface temperature coastal region near the big cities, the continual nutrient will increase by 2.5∞C (Houghton et al., 1995). The tem- input after the Industrial Revolution raises the oceanic peratures of only surface ocean boxes are uniformly in- production rate higher than that at the past steady state. creased by DT. It might be well justified that although the simulation studies so far had revealed that the global tem- Alkalinity pump perature increase is heterogeneous over the earth surface, There exists another driving force for the carbon trans- i.e., more enhanced on the high-latitudinal areas in both port into the oceanic interior via both production and dis- the north and south hemispheres, our pre-checking showed solution of calcium carbonate (CaCO3). This mechanism that the heterogeneity of the temperature increase due to is generally called the alkalinity pump. In our model, the the global warming is likely to have little influences on alkalinity pump works only in the uppermost and the calculated P . CO2 lowermost boxes in each oceanic basin: the reaction, 2+ – Ca + 2HCO3 = CaCO3 + CO2 + H2O, goes to the right Biological pump hand side in the euphotic layer by photosynthesis and to The biological pump is the process by which CO2 fixed the left hand side in the ocean bottom instantaneously,

32 M. Sorai and T. Ohsumi while the calcium carbonate (CaCO3) particles travel without any dissolution in the middle depth part of the ocean, PP1 bPP CO22- CO+ CO 2- CO 2 because the lysocline which is defined as the uppermost ()bb()0 () TNPC = TNP0 (){}() depth of the calcite dissolution is generally below 2,000 PP1 bPP CO22- CO+- CO 22 CO m depth. ()()0 ()bb{}()() The rain ratio, i.e., the CaCO3 production divided by ()9 the new production, is the key parameter which determines the strength of the alkalinity pump. Broecker and Peng (1982) defined this ratio as 0.25, but their model 680- PrPCO--l 340 CO included no effect of the difference of the dissolution ()22b ()b b = ()(), 10 depth between CaCO and POC produced by the new pro- () 3 rPl -1 680- CO 340 - P CO duction. On the other hand, Yamanaka and Tajika (1996) ()()()22bb()() revealed that in fact the ratio was considerably reduced to approximately 0.09, when the difference of the disso- where ( P ) is the atmospheric P at the CO com- CO2 b CO2 2 lution depth was taken into account. Since our model as- pensation point (=31 ppmv: Gifford, 1993) where net sumed that POC and CaCO3 dissolved in the middle and daytime photosynthetic CO2 fixation is zero, and rl is deep depth parts of box respectively, we adopted the ra- defined as TNP(680 ppmv)/TNP(340 ppmv). tio of 0.09. The organic carbon in the terrestrial plant will be de- As is the case in the biological pump, the CO2 effect composed. The decomposition rate Fplant was divided into which reduces the calcification of the marine plankton in two fluxes: one represents the fraction to the atmospheric response to the P increase is taken into account in the CO2 CO2 and the rest to the soil, while the ratio of these two formulation of the alkalinity pump in the model: fluxes was defined by an adjusting parameter k in the model. Similarly, the soil organic carbon will decompose P 280 and return to the atmospheric CO2. These decomposition Ê CO2 - ˆ rain= rain0 ◊ Á1 - krain ◊ ˜, ()8 fluxes, Fplant and FluxsoilT, were assumed to be propor- Ë 750- 280 ¯ tional to the plant and soil masses, respectively. Moreover, the global warming effect was applied for where rain is the rain ratio, krain is the normalized de- the net primary production and the decomposition of or- crease index of the ratio of calcification to POC produc- ganic carbon in the soil, respectively as, tion, and subscript 0 denotes the value in the pre- industrial level. Equation (8) is based on the experiment TNP TNP111 k T that the rain ratio decreased up to 52.5% with an increase CT, =+ C()NPPD () in P from 280 ppmv to 750 ppmv (Riebesell et al., CO2 and 2000). This mechanism is easily understood, because the P increase in the surface water makes the above reac- CO2 tion to go to the left hand side. FluxsoilT =+ Flux soil()112 k soilD T , ()

Terrestrial ecosystem where TNPC,T and FluxsoilT are the net primary produc- The terrestrial carbon cycle system was built up by tion (GtC yr–1) and the soil decomposition fluxes (GtC –1 simplifying Wigley’s treatment (Wigley, 1993), that is, yr ) as a function of the warming effect, and kNPP and the terrestrial components are only plant and soil in our ksoil are the warming factors on TNP and on Fluxsoil, re- model. The land biota works as the CO2 sink through the spectively (Gifford, 1993). –1 net primary production TNPC (GtC yr ) enhanced by the atmospheric P increase. This CO fertilization effect CO2 2 MODEL EVALUATION is generally formulated by either the logarithmic form (Bacastow and Keeling, 1973) or the Michaelis-Menten Initial condition and input data form (Gates, 1985) with respect to the atmospheric P . The pre-industrial oceanic variables, such as the DIC CO2 However, Gates (1985) suggests that plant photosynthe- concentrations, the DOC concentrations, the alkalinities sis and ultimately growth actually does not follow the and the temperatures, were set so as to meet the require- ment as follows: 1) the pre-industrial atmospheric P logarithmic response to the CO2 increase, but that the CO2 Michaelis-Menten form behaves more realistically on the is 280 ppmv, and 2) the pre-industrial carbon amounts of point such that it leads to zero net primary production at terrestrial plant and soil are 610 and 1560 GtC, respectively (Siegenthaler and Sarmiento, 1993). The model was the low PCO and has a limiting value at the high PCO . 2 2 started with the global data sets of GEOSECS (Takahashi Thus, we adopted the Michaelis-Menten form for TNPC:

Ocean uptake potential for CO2 sequestration 33 Table 1. Pre-industrial variables for each ocean box

Ocean DIC Alkalinity DOC Temperature (mol m–3)* (mol m–3) (mol m–3) (°C)

Arctic surface 2.219 2.366 0.062 3.6 middle 2.270 2.375 — 0.7 North Atlantic surface 2.155 2.366 0.062 17.1 middle 2.244 2.351 — 8.1 deep 2.266 2.375 — 2.2 South Atlantic surface 2.148 2.366 0.058 22.5 middle 2.946 2.351 — 9.7 deep 2.330 2.441 — 1.8 Indian surface 2.139 2.394 0.048 26.6 middle 2.480 2.434 — 11.9 deep 2.461 2.437 — 1.8 Pacificsurface 2.171 2.408 0.039 22.6 middle 2.528 2.426 — 10.8 deep 2.470 2.445 — 1.5 Antarctic surface 2.261 2.394 0.048 3.6 middle 2.465 2.435 — 0.7

*The conventional unit in oceanography of mol/(kg ocean water solution) was converted into the unit of mol m–3 by using the average density of seawater 1.025 g cm–3 (Broecker and Peng, 1982).

production for the period 1751–1999 and those from the land-use change for the period 1850–2000. The CO2 emissions from the land-use change include the effects of not only the initial removal and oxidation of the carbon in the vegetation, but also the subsequent re-growth and the changes in the soil carbon. Here, it should be noticed that the land-use change should actually have started before the year 1850. Hence, we assumed that the CO2 emission from the land-use change started in the year 1700, which agreed with the start of the abrupt increase in the world population, and that it increased linearly until the year 1850. Since the human impacts on CO2 emission before the year 1700 were negligible, we regard that before the year 1700 the anthropogenic CO2 emission was zero. This Fig. 2. Anthropogenic CO2 emissions both from the industrial leads to the assumption that the pre-industrial period be- sources and from the net land-use change. The CO2 emission fore 1700 was a steady state (Ver et al., 1999), which is from land-use change was assumed to start in the year 1700 also supported by the measurements from ice cores such and to linearly increase until the year 1850. that the atmospheric P had been almost constant CO2 within a range of 10 ppmv in this period. The anthropogenic CO2 emissions were directly input to the atmos- et al., 1981), except for Ymer80 data on the Arctic Ocean pheric CO2 reservoir in our model after the year 1700. (Anderson and Dyrssen, 1981). The spin-up run under the above boundary conditions over a million year made the Comparisons between calculated and observed values model to the steady state, so that the oceanic variables For validation of the model, ideally, all the carbon drifted to the pre-industrial steady state values, which are components should be evaluated on their mass changes specific to our model. The model pre-industrial values by the comparison between the calculated results and the for each ocean box are shown in Table 1. observed data. Of all the carbon reservoirs in the model, The historical trends of anthropogenic CO2 emissions however, only the atmospheric CO2 has been known on are given by Online Trends, year-by-year data sets by the time course of concentration change from the pre- Carbon Dioxide Information Analysis Center (Online industrial level to the present time. Therefore, several pa- TREND). As shown in Fig. 2, these include the global rameters depicted in the section “Model Description” were emissions from the fossil fuel combustion and the cement adjusted primarily aiming at the representation of atmos-

34 M. Sorai and T. Ohsumi Table 2. Adjustable parameters in the model

Parameter Definition Adopted value Literature value

ro CO2 fertilization factor on NP 1.1 1) rain0 Rain ratio in steady state 0.09 0.08–0.10 2) krain Normalized decrease index of the ratio of calcification to POC production 0.525 0.21–0.525 d Decomposition fraction of POC to DOC 0.15 3) rl CO2 fertilization factor on NPP 1.14 1.1–1.4 k Fraction of direct decomposition to air in plant degradation 0.87 0.47–0.873) –1 3) –1 kNPP Warming factor on NPP 0.01 (K ) 0.01–0.05 (K ) –1 3) –1 ksoil Warming factor on Fluxsoil 0.03 (K ) 0.03–0.05 (K )

1)Yamanaka and Tajika, 1996. 2)Riebesell et al., 2000. 3)Gifford, 1993.

pheric P changes. These parameters were finally 380 CO2 Calculated resutls in this work evaluated as shown in Table 2. In this study, our model Siple 360 was checked with respect to several fundamental vari- Mauna Loa

ables: the time-course profile of atmospheric P , the (ppmv) CO2 340

DIC concentration and the alkalinity in 1970’s, and the CO2 oceanic CO2 uptake rate in 1990’s. 320 (1) Time-course profile of CO2 concentration The time- 300 course profile of the atmospheric CO2 concentration was simulated using the values of reservoirs, fluxes and ad- 280 justable parameters defined in the section “Model De- Atmospheric P scription”. The anthropogenic CO2 emissions data de- 260 scribed above was applied. The calculated result is shown 1650 1700 1750 1800 1850 1900 1950 2000 in Fig. 3, with the data from the Siple ice core and the air Date (A.D.) measurements in Mauna Loa (Online TREND). Our re- P Fig. 3. The time-course change in the atmospheric CO2 . The sults after the year 1970 are relatively well consistent with calculated result shows a good agreement with the observation the observations, while some deviations exist especially especially after the year 1970, whereas there exist some devia- in the period of the year 1850 to 1950. We found that the tions in the period of the year 1850 to 1950. revised form of Eq. (9), in which the TNP is assumed to be proportional to the amount of the plant, gave a considerably better agreement with the observation. Although this modification appears to be more reasonable for the ited qualitatively a good agreement with the GEOSECS past change in the atmospheric P , it certainly accom- CO2 data. Especially, the alkalinity in each part of oceans panies the uncertainties on the future trend, that is, the agreed to well within 2%. The exceptions are in the sur- TNP and thus the amount of plant continue to increase face and middle parts of the Indian Ocean, but even in synergistically by both effects of CO2 fertilization and these boxes the deviation was within 3%. The DIC con- global warming; in turn, this enhanced TNP causes the centrations exhibited relatively large deviations, mostly abrupt drop of the atmospheric P to the unrealistic within about 10%, especially at all surface parts and a CO2 lower level. Therefore, the modified form of Eq. (9) was part of middle depth regions, whereas the deviations of not used in this study. Nevertheless, considering that our the deep boxes were reduced to 3 to 4%. Extremely higher study mainly focuses on the future CO2 trend, it would value was obtained in the middle part of the South be more appropriate that the model is tuned to the recent Atlantic Ocean; this would be the result of the active de- change of CO2 concentration rather than to the past un- compositions of the settling-down POC, because the large certain trend. upwelling flow in this area transports the abundant phos- (2) DIC concentration and alkalinity The GEOSECS data phate to the surface layer and causes the higher new pro- were used for checking of both the DIC concentration duction using them. These deviations are likely related and the alkalinity. Table 3 shows the comparison between to the magnitude of the oceanic thermohaline circulation the GEOSECS data and the model values at the time cor- in our model. In other words, the active oceanic circula- responding to the GEOSECS expeditions. All calculated tion would stir around the whole ocean and dilute a high results of the DIC concentration and the alkalinity exhib- DIC concentration in the middle parts of oceans. As dis-

Ocean uptake potential for CO2 sequestration 35 Table 3. Comparison between the GEOSECS data and the model values

Ocean DIC (mol m–3)Alkalinity (mol m–3)

This work Observation This work Observation

Arctic surface — — middle 2.272 2.240* 2.375 2.389* North Atlantic surface 2.176 1.993 2.369 2.344 middle 2.252 2.177 2.352 2.360 deep 2.269 2.236 2.375 2.386 South Atlantic surface 2.166 2.010 2.369 2.352 middle 2.948 2.196 2.350 2.373 deep 2.331 2.267 2.440 2.407 Indian surface 2.165 1.983 2.396 2.340 middle 2.480 2.240 2.434 2.370 deep 2.461 2.372 2.437 2.469 Pacificsurface 2.192 2.027 2.409 2.366 middle 2.528 2.241 2.426 2.395 deep 2.470 2.391 2.445 2.483 Antarctic surface 2.279 2.226 2.395 2.413 middle 2.466 2.313 2.435 2.427

*These data were obtained at the Ymer80 expedition.

cussed in the subsection “Biological pump”, we tripled pump. Our estimate of the new production in 1980’s is all flows of Schmitz’s circulation in order to represent 9.9 GtC yr–1 and this is well consistent with the estimate more realistic residence time of deep water. Neverthe- of 11 GtC yr–1 in IPCC (2001). less, our oceanic circulation might still possibly be lower. Although this point should be considered in more detail, MODEL IMPLICATION ON OCEAN SEQUESTRATION considering that our main interest of the oceanic potential is independent on the magnitude of the oceanic cir- Scenarios for the future culation, namely on the representation of the DIC distri- (1) Future emission scenario The B2 case of the Special bution, the above deviations in the DIC concentration Report on Emissions Scenarios (SRES) in IPCC (2001), would be within tolerable limits. which assumed the world where the emphasis is on local (3) Oceanic CO2 uptake rate The analysis of the oceanic solutions to economic, social, and environmental and terrestrial CO2 uptake rates provides an important sustainability, was applied for the anthropogenic CO2 criterion on the evaluation of the model. Intergovernmen- emissions from the year 2010 to 2100 (Houghton et al., tal Panel on Climate Change (IPCC) summarized the glo- 2001). After the year 2100 the emission scenario is rather bal CO2 budget: the CO2 uptake rates in 1990s as 1.7 ± arbitrary. One effective approach is to set the future emis- –1 –1 0.5 GtC yr in the ocean and as 1.4 ± 0.7 GtC yr in the sion scenario limited by the total natural resources avail- land (Houghton et al., 2001). Considering that the terres- able to us. The global fossil fuel energy resource summa- trial system still has a lot of uncertainties and that our rized by Rogner (1997) was 4911 GtC. If the industrial main purpose is to examine the oceanic CO2 sequestra- CO2 emission rate is assumed to decrease linearly after tion, we focused on the oceanic CO2 uptake rate in this the year 2100, the final year when the fossil fuel con- study. Since the solubility pump determines the magni- sumption is completed becomes the year 2652. Another tude of the CO flux from the atmosphere to the ocean, criterion is to fix the atmospheric P level in the fu- 2 CO2 the oceanic CO2 uptake rate was estimated using Eq. (2) ture steady state. Once a steady state P is set to any CO2 and each oceanic area. The averaged value of this uptake target value in the stabilization scenarios, we can esti- rate in 1990’s was 1.3 GtC yr–1. This rate corresponds to mate the whole carbon amount in the steady state; the almost the lowermost value in the compilation of IPCC difference between the total carbon amount in the year (2001). In fact, we found that more active oceanic circu- 2100 and that in the steady state equals to the carbon relation produced larger CO2 uptake rate in the ocean, but sources allowed for us to use. In this study, the steady this modification would be beyond the purpose in this state P is set to 450, 550, 650, 750 and 1000 ppmv; CO2 study. these P values agreed with those in the stabilization CO2 The new production in the ocean becomes another scenarios in IPCC (1995) (Houghton et al., 1995). Also important criterion for validation of the model biological in these cases, we assumed that the CO2 emission from

36 M. Sorai and T. Ohsumi (a) 14 of the ocean CO2 sequestration was set to the year 2015. 450ppmv The duration for the oceanic CO2 sequestration was as- 12 550ppmv 650ppmv sumed until the end year of the fossil fuel consumption. 10 750ppmv The amount of the CO sequestered was varied from 20 1000ppmv 2 8 Land-use change to 100% of the annual CO2 emission from the fossil fuel combustion. GOSAC project simulated the ocean CO 6 2 sequestration in seven offshore sites (GOSAC, 2002). We 4

2 assigned these sites to our box oceans: Bay of Biscay and

CO emissions 2 New York belong to the North Atlantic Ocean, Rio de Janeiro to the South Atlantic Ocean, Bombay and Jakarta 0 to the Indian Ocean, and San Francisco and Tokyo to the -2 2000 2100 2200 2300 2400 2500 2600 2700 Pacific Ocean. Based on this ratio of 2:1:2:2, the pre- Date (A.D.) scribed amount of CO2 in the form of DIC was injected into the middle parts of these oceans. (b) 2000 In fact, the ocean sequestration will cause the addi- 1000ppmv 750ppmv tional CO2 emission, because the CO2 capture process 650ppmv requires the extra energy consumption. David and Herzog 1500 550ppmv 450ppmv (2000) analyzed the economics of capturing CO at Inte- (ppmv) 2 grated coal Gasification Combined Cycle (IGCC) power CO2 1000 plants, Pulverized Coal (PC) power plants, and Natural Gas Combined Cycle (NGCC) power plants, and estimated the additional energy requirements in the year 2012 500 as 9% for IGCC, 15% for PC, and 10% for NGCC, respectively. Thus, we assumed that the extra 10% of the Atmospheric P sequestered CO , which is though a little bit lower than 0 2 2000 2500 3000 3500 4000 15% at PC, is added to the original emission scenario as Date (A.D.) defined by the following expression,

P Fig. 4. Future estimates of the atmospheric CO2 . (a) Future CO2,seq ErB2 13 anthropogenic CO2 emission scenarios. The B2 case of the SRES =¥ () scenario in IPCC (2001) was applied from the year 2010 to

2100. After the year 2100, the anthropogenic CO2 emissions and were linearly decreased to satisfy the total carbon emissions P reaching to the steady state CO2 values of 450, 550, 650, 750 EErext=¥ B2 ¥ 01., 14 and 1000 ppmv, respectively. In all cases, the CO2 emission () () from land-use change was assumed to become zero after the year 2200. (b) Calculated results of the change in the future where CO2,seq is the annual amount of CO2 sequestered, E is the industrial CO emission rate defined above, r atmospheric PCO . The atmospheric PCO were calculated us- B2 2 2 2 is the ratio of CO sequestered to the scenario emission, ing the above anthropogenic CO2 emission scenarios. All cases 2 once increased to the peak value and then decreased to the and Eext is the extra CO2 emission due to CO2 capturing. steady state. In the case of the ocean sequestration, the year when fossil fuel consumption is completed was quickened to match the total carbon amount used. the industrial sources linearly decreased after the year Future P trend 2100 to satisfy the total carbon emission defined above. CO2 The calculated trends of the atmospheric PCO greatly Thus, the end year when the industrial CO2 emission is 2 completed varies dependent on each scenario: the end depend on the future scenario of the anthropogenic CO2 years are the year 2163, 2274, 2371, 2457, and 2646, re- emissions. Figure 4(b) is the calculated results based on spectively. The last 1000 ppmv-stabilized scenario is various scenarios shown in Fig. 4(a). Each emission sce- roughly consistent with the case based on the fossil fuel nario had a different total amount of CO2 emission, and thus resulted in the different atmospheric P level at resources. CO2 the ultimate steady state. Here, we have no definite evi- On the other hand, the net CO2 emission from the land- use change was assumed to become zero after the year dence that the system will finally attain to the steady state. 2200. These emission scenarios are shown in Fig. 4(a). But our model assumes that no burial processes of POC (2) Ocean sequestration scenario In all cases, the start and CaCO3 occur in the ocean, so that once the CO2 input

Ocean uptake potential for CO2 sequestration 37 from anthropogenic sources are stopped, the system in- 1000 evitably attains to the steady state and thus the atmos- 100% Air 900 20% (2015-2241) pheric P is stabilized at a constant value. Moreover, 40% (2015-2227) CO2 60% (2015-2217) 800 the climate change, which might perturb the system and 80% (2015-2209) (ppmv) force to change over to another steady state, is not thor- 100% (2015-2202) CO2 700 oughly incorporated in the model, except for the global warming effects resulting in the activity change of ter- 600 restrial ecosystem and the intensity of solubility pump. 500 Such simple treatments are useful to understand the CO2 fate more explicitly. 400 Atmospheric P In all cases in Fig. 4(b), the atmospheric P once CO2 300 increased over the steady state value and then decreased 2000 2500 3000 3500 4000 gradually. It was found that the rate of approach towards Date (A.D.) the steady state and the peak P value both depend CO2 Fig. 5. The dependence of sequestered portion of fossil-fuel mainly on the magnitude of oceanic circulation: the slower CO2 on the atmospheric PCO . In all cases, the steady state circulation which causes the higher peak value of the at- 2 PCO was set to 550 ppmv and the prescribed portion of total mospheric P requires the longer time to attain to the 2 CO2 industrial emissions were injected in the ratio of 2:1:2:2 to the steady state and vice versa (Appendix II). However, what- middle depth parts of the following four oceans: the North and ever the strength of oceanic circulation, the atmospheric South Atlantic Oceans, the Indian Oceans and the Pacific P shows the essentially analogous profiles, and thus Ocean. The sequestered CO2 was varied between 20% and CO2 provides less influence on the following discussions. 100%. The CO2 was injected from the year 2015 to the end It should be noticed that all the peak P reached year of fossil fuel consumption (the end year was different de- CO2 pendent on the amount of CO sequestered). about the twice of the corresponding steady state values. 2 This implies that both oceanic and terrestrial capacities of CO2 absorption will not catch up the anthropogenic emission rate. In order to keep the atmospheric P un- understand the reason as a result of the reduction of the CO2 der the steady state value, in other words, for the realiza- CO2 emissions to the atmosphere. Therefore, this effect tion of the P time course with no transient peaks, we is more enhanced by the larger-amount and longer- CO2 should mitigate the burden to the atmosphere. The sim- period sequestration. The oceanic CO2 injection with the portion less than 80% caused the peak over 550 ppmv, plest solution for this mitigation is either direct CO2 in- while the CO2 injection more than 80% exhibited no peak. jection to the ocean or reduction of CO2 emission itself. Considering that there still exist a lot of difficulties to The DIC concentrations in the middle Pacific are also the large-scale reduction of our future emissions, direct depicted in Fig. 6(a). The other CO2-sequestered oceans, such as the middle parts of the North and South Atlantic CO2 injection to the ocean might be reasonable in a sense such that the reduction of a CO release to the atmos- Oceans and of the Indian Ocean, also have the analogous 2 time scale of the change in the DIC concentration, al- phere is achieved by acceleration of oceanic CO2 absorption. though the variation range is dependent on the volume of each ocean. Direct CO2 injection to the middle ocean abruptly increased the DIC concentration just after the Ocean CO2 sequestration Next, we examined the effect of the ocean CO se- starting of sequestration and then slowly changed to the 2 final steady state. Such a stabilization of the DIC con- questration on the atmospheric P based on the 550 CO2 centration is caused by the thermohaline circulation, ppmv stabilized case in Fig. 4(b). Figure 5 shows the cal- which transports the high-DIC water produced by the culated results of the CO sequestration, where the se- 2 ocean sequestration to the whole ocean. As expected from questered portion of the total industrial emissions was Fig. 1, most of the CO injected to the middle parts of varied from 20 to 100%. In each case, the starting year of 2 oceans is once transported to the deep parts, and then the CO sequestration was set to the year 2015. The pre- 2 spread to the whole ocean on the horizontal or vertical scribed portion of the fossil fuel CO was injected to the 2 flows. The only path to the surface layer is the South middle parts of the oceans during the CO sequestration 2 Atlantic Ocean and the Pacific Ocean. The high-CO water and the rest was released to the atmosphere. Since the 2 reaching to the surface layer raises the PCO in the sur- CO2 sequestration was assumed to proceed until the end 2 year of fossil fuel consumption, the end year was differ- face ocean, and consequently reduces the CO2 uptake rate. Especially, in the case more than 80% of sequestration, it ent dependent on the amount of CO2 sequestered. As shown in Fig. 5, ocean sequestration obviously decreases was found that the net release of CO2 from the ocean oc- a peak value of the atmospheric P . We can simply curs. CO2

38 M. Sorai and T. Ohsumi 6000 (a) 2.8 Tot al 100% Air 20% (2015-2241) 5000 Ocean 40% (2015-2227) 2.7 60% (2015-2217) 80% (2015-2209) 4000 100% (2015-2202)

3 3000 2.6

Uptake potent 2000 (mol/m )

2.5 1000

0 Dissolved inorganic carbon 2.4 400 500 600 700 800 900 1000 2000 2500 3000 3500 4000 Date (A.D.) Steady-state PCO2 (ppmv)

(b) 2.5 Fig. 7. CO2 uptake potential in the ocean after the year 2015. 100% Air The difference between the final and the present carbon amount 20% (2015-2241) in the ocean correspond to the oceanic CO uptake potential. 40% (2015-2227) 2 2.4 60% (2015-2217) P This potential is determined by the atmospheric CO2 value in 80% (2015-2209) 100% (2015-2202) the ultimate steady state. The factors to affect the CO2 uptake

3 potential are the temperature of seawater, the new production, 2.3 and the rain ratio. (mol/m )

2.2

if no CO2 reaction with solid earth is taken into account. Therefore, when we set any P value at the future

Dissolved inorganic carbon CO 2.1 2 2000 2500 3000 3500 4000 steady state, the final carbon amount in the ocean will be Date (A.D.) also fixed. The difference between the final and the present carbon amount corresponds to the maximum CO2 Fig. 6. The change in the DIC concentration due to the ocean amount to invade into the ocean that is, the oceanic CO CO sequestration. The changes in the DIC concentration were 2 2 uptake potential integrated with time until the end of our estimated (a) in the middle depth part of the Pacific Ocean, and (b) in the surface Pacific Ocean. Each calculation corre- fossil fuel era. Figure 7 shows this potential at the various steady state P . sponds to the cases in Fig. 5. CO2 It is interesting to show the other aspect of the potential calculation. The oceanic uptake potential after the year 2015 is 1623 GtC in the case of steady state P = 550 Owing to the future increase in the atmospheric P , CO2 CO2 ppmv. Hence, we conducted two additional simulation however, the DIC concentration is also raised even with- runs, in which CO2 equivalent to 1623 GtC was injected out direct CO2 injection. In other words, the CO2 seques- to the middle and deep parts of the four oceans in the tration only brings about the future increase in the DIC duration from the year 2015 to 2145. It was found that concentration earlier, at least in the case less than the in- the calculated P does not exceed 550 ppmv level in jection of 80% portion. If the biological impact to this CO2 the sequestration to the deep ocean (Fig. 8). But in the oceanic domain is the same for these curves including middle depth sequestration, P rose slightly over 550 the business-as-usual scenario (100% Air), the mitigation CO2 ppmv level until the year 3500. This unexpected phenom- of the land and surface ocean impact is the net benefit of enon is caused by the CO transport from the ocean. Here the ocean sequestration. We can see these benefits in Fig. 2 it should be noticed that this CO transport to the air is 6(b), where the corresponding DIC concentration profiles 2 not completely equal to the leakage of the sequestered in the surface Pacific are shown. It is noteworthy that since CO itself. The CO injection raises the HCO – concen- most of marine organisms are concentrated within the 2 2 3 tration in the ocean and thus also increases the P in depth range shallower than 200 m, the DIC increase in CO2 surface ocean should be suppressed preferentially. the surface water before they were homogeneously di- luted in the whole ocean. This will weaken the oceanic CO uptake potential CO2 absorption, because the solubility pump is propor- 2 tional to the P difference between in the air and in the As discussed above, the total carbon amount on the CO2 ocean. If the CO injection is further continued, the P surface domain of our planet is decided uniquely in the 2 CO2 in the surface water will become larger than the atmos- future steady state of the atmospheric CO2 concentration,

Ocean uptake potential for CO2 sequestration 39 700 3.0

Middle ocean 650 Deep ocean

2.5 600 (ppmv)

CO2 550 2.0

500 increment CO2

450 P 1.5

400 Atmospheric P

350 1.0 2000 2500 3000 3500 4000 400 500 600 700 800 900 1000

Date (A.D.) Steady-state P CO2 (ppmv) P P Fig. 8. Calculated CO2 in the case of 1623 GtC CO2 injec- Fig. 9. The ratio of the CO2 in the middle depth parts of the tions in the middle or deep depth parts in the four oceans. The Pacific Ocean in the ultimate steady state to that in the year 1623 GtC CO , which correspond to the oceanic uptake poten- P 2 1999. The basal CO2 value in the middle Pacific Ocean is about tial at the steady state P = 550 ppmv, was injected in the P CO2 1200 ppmv. CO2 in the seawater will be raised owing to the middle or deep depth parts in the four oceans. It should be no- P future increase in the atmospheric CO2 level even without di- ticed that, in the middle depth sequestration, P rose slightly CO2 rect CO2 injection. over 550 ppmv level until the year 3500.

Broecker (2001) estimated the total buffer capacity of pheric P and the net CO flux from the ocean to the CO2 2 the current ocean as more than 2500 GtC based on the air will be caused. The CO injection to the deep ocean 2 amounts of both CO 2– and B(OH) – which react with CO also produces the CO transport to the air, but its amount 3 4 2 2 to form HCO – and B(OH) , respectively. Although the is slight. Our simulation reveals that both the site and the 3 3 CO 2– concentration shows a good estimate of the CO scenario are also important factors on the ocean CO se- 3 2 2 carrying capacity of seawater (Cole et al., 1993), this only questration. provides the maximum oceanic potential, because the CO By simple consideration, we expect that the oceanic 2 injection to the ocean increases not only the HCO – con- CO uptake potential might be uniquely determined by 3 2 centration but also the P in the ocean simultaneously. the solubility of CO in seawater. However, our sensitiv- CO2 2 Therefore, even if all the CO 2– ions react with the in- ity analyses revealed that the uptake potential is regu- 3 jected CO to form HCO –, the resultant P increase lated by various factors, not only by the temperature of 2 3 CO2 seawater, but also by the magnitude of the new produc- will cause the CO2 release to the air. In other words, the tion and the rain ratio. These were related to each pump oceanic capacity of the CO2 absorption is essentially regu- mechanism explained in the section “Model Description”. lated by the CO2 exchange across the atmosphere-ocean interface. At lower temperature, the solubility of CO2 in seawater increases and thus the solubility pump enhances the oce- The ocean CO2 sequestration still leaves room for con- anic CO uptake. It should be noticed that the gas ex- cerns on the irreversible impact on the marine ecosys- 2 tem. Specifically, the P increase in the ocean, rather change coefficient E in Eq. (2) and the magnitude of the CO2 oceanic circulation do not affect the uptake potential, than the DIC increase (that is pH decrease), is one of the most important concerns on the assessment of the bio- because they only change the rate of the CO2 exchange between the air and the ocean (see also Appendix II). The logical impacts and some recent works have focused on photosynthesis in the surface ocean also plays an impor- this problem (e.g., Shirayama, 1997; Kikkawa et al., 2004). However, it should be emphasized again that the tant role in the transport of the atmospheric CO2 into the P in the seawater will be raised owing to the future ocean. Although the new production is generally control- CO2 increase in the atmospheric P level even without di- led by the nutrient concentration, it is possible that any CO2 rect CO injection and that the P in the ocean increases future perturbation such as the CO2 fertilization effect as 2 CO2 in this study affects the CO uptake potential. The alka- with an increase in the steady state P . Figure 9 shows 2 CO2 linity pump works as an inverse sense, that is, the higher the ratio of the P in the middle depth parts of the CO2 rain ratio reduces the uptake potential, because the CaCO3 Pacific Ocean in the ultimate steady state to that in the production raises the partial pressure of CO2 as discussed year 1999; the base P value in the middle Pacific CO2 in the subsection “Alkalinity pump”. Ocean is about 1200 ppmv as shown in the GEOSECS

40 M. Sorai and T. Ohsumi data. For instance, in the case of the steady state P of maximum CO amount to invade into the ocean by way CO2 2 550 ppmv, the oceanic P will reach up to about 1.7 of either natural sink processes or purposeful injection. CO2 times of the present value. Here the most important point The present study showed that the oceanic is whether this range of the P increase is acceptable thermohaline circulation should be parameterized in more CO2 to the marine organisms or not. Naturally, we can think detail to represent the more realistic distribution of the DIC concentration. These improvements may include the that the ocean CO2 sequestration will accelerate the oceanic P increase and the rate of this P increase must partitions of the model reservoirs both horizontally and CO2 CO2 give some effect on the marine organisms. Nevertheless, vertically. It should be noticed that the present model does it should be noted that if we know the limit condition of not treat the following well-known processes of carbon the P increase on the impact to the marine organisms, cycle. The sedimentation processes of the organic matter CO2 and the CaCO particles, and thus the carbon reservoir of the maximum atmospheric P and also the maximum 3 CO2 the deep sea sediment itself, were not taken into account CO amounts allowed to be injected to the ocean can be 2 in the model. It is easily expected that especially the determined from the biological viewpoint. CaCO3 has an essential role on the neutralization of the increased CO2 in the ocean. Such an effect encourages CONCLUSIONS the ocean CO2 sequestration at least on the point of the mitigation of the biological impacts, because the marine In order to assess the implication of the ocean CO 2 CaCO will probably enhance the neutralization of the sequestration, we conducted a simulation study using our 3 high-CO waters due to the CO injection. In addition, global carbon cycle box model consisting of 19 reservoirs. 2 2 Archer et al. (1998) revealed that the neutralization by The model incorporates all the essential processes in the the CaCO in the deep sea sediment occurs on a timescale global biogeochemical carbon cycles, such as the ocean 3 of about 5000 years and that this mechanism has an abil- thermohaline circulation, the solubility pump, the biologi- ity just enough to neutralize the 60–70% of the fossil fuel cal pump, the alkalinity pump and the terrestrial ecosys- release. Therefore, it is possible that the ocean holds the tem. The present major modifications from our previous much larger potential to uptake the excess CO than the model include the inclusion of the terrestrial ecosystem, 2 present estimate, when the neutralization effect was also in which the negative feedbacks both of the CO fertili- 2 considered in the simulation. zation and of the global warming on the net primary pro- On the other hand, our study assumed that climate duction and the positive feedbacks of the global warm- change brings about only the temperature increase resulting on the decomposition process of the soil organic car- ing in the activity change of terrestrial ecosystem and the bon are taken into account. Moreover, we also introduced intensity of solubility pump. In the long-term range more the CO increase effects on the oceanic new production 2 than hundreds of years, however, it is possible that the and on the marine CaCO production. These improve- 3 climate change has the significant influences to the bio- ments enabled us to simulate the future CO fate in the 2 logical systems; these include the positive feedbacks due long term using the more realistic carbon cycle system. to the global warming of the oceanic new production and Model parameters were validated with respect to the time- of the mineralization of the organic carbon (Klepper and course profile of the atmospheric P , the DIC concen- CO2 Haan, 1995). Moreover, several simulation studies have tration and the alkalinity in 1970’s, and the oceanic CO 2 pointed out that the atmospheric P increase causes uptake rate in 1990’s. CO2 the reduction or the stopping of the oceanic thermohaline Using this model, the CO injection into the middle 2 circulation (Manabe and Stouffer, 1993; Haywood et al., part of the ocean was simulated varying the sequestra- 1997; Stocker and Schmittner, 1997). If such changes tion period and the amount of CO sequestered. It was 2 occurred in the future, the global carbon cycle system will found that if we accept the 550 ppmv stabilization level be drastically changed (Mackenzie et al., 2000), not only as a future steady state, the CO injection more than 80% 2 in the oceanic environments but also in the terrestrial sys- of total anthropogenic emission is required after the year tem, and the ocean CO sequestration might show the 2015 not to exceed this level. In addition, we can show 2 another aspect entirely different from the present work. the direct ocean injection of CO2 at the present time means the acceleration of the pH lowering in the middle ocean Acknowledgments—This work was performed as a part of the due to the eventual and inevitable increase of CO2 in the Research and Development on CO2 Ocean Sequestration Project atmosphere. Moreover, for the minimization of the influ- supported by Ministry of Economy, Trade and Industry (METI). ences on a marine ecosystem, the ocean sequestration has We are grateful to Dr. Masamichi Ishikawa for many useful an important role to suppress a decrease in the pH value comments on the construction of the global carbon cycle box in the surface ocean. We estimated the CO2 uptake po- model. We thank Dr. Shigeo Murai for helpful comments on tential in the ocean based on total carbon amount in vari- the ocean sequestration scenario, and Prof. Abraham Lerman ous steady state P . These estimates provide us the and an anonymous reviewer for useful comments on a draft CO2

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42 M. Sorai and T. Ohsumi Ver, L. B., Mackenzie, F. T. and Lerman, A. (1999) dCi n Q n Q ONP rain Biogeochemical responses of the carbon cycle to natural i ijÆ Ci jiÆ Ci ◊ .A3 dt =-ÂÂV i +V j + V () and human perturbations: past, present, and future. Amer. j =11i j = i i J. Sci. 299, 762–801. jiπ jiπ Walker, J. C. G. (1991) Numerical Adventures with Geochemical Cycles. Oxford University Press, 192 pp. In this study, the DOC was defined only in the surface Wanninkhof, R. (1992) Relationship between wind speed and layer, where a part of POC decomposes to DOC: gas exchange over the ocean. J. Geophys. Res. 97, 7373– 7382. o n Q n Q Weiss, R. F. (1974) Carbon dioxide in water and seawater: the dCi ijÆ o jiÆ o ONP◊ dprato solubility of a non-ideal gas. Mar. Chem. 2, 203–215. =-Ci +Cj + .A4() dt ÂÂVi Vi Vi Wigley, T. M. L. (1993) Balancing the carbon budget: implica- j =11j = jiπ jiπ tions for projections of future carbon dioxide concentration changes. Tellus 45B, 409–425. Yamanaka, Y. and Tajika, E. (1996) The role of the vertical Also on the alkalinity, the processes responsible for fluxes of particulate organic matter and calcite in the oce- the budget are essentially identical to the above DIC cases. anic carbon cycle: Studies using an ocean biogeochemical We assumed that the alkalinity, Alk, increases by 0.17 eq general circulation model. Global Biogeochem. Cycles 10, per mol production of the organic carbon and decreases 361–382. by 2 eq per mol production of CaCO3. Therefore the balance equation of Alk in the surface layer is expressed as follows: APPENDIX I. BALANCE EQUATIONS The DIC in the surface layer of the ocean is balanced dAlk by the CO2 exchange with the atmosphere, and the bio- i logical productions of the POC and CaCO3, in addition dt to the advection: n Q n Q ONP017. 2 rain ijÆ Alki jiÆ Alki ()-¥ . =-ÂÂV i +V j + V j =11i j = i i jiπ jiπ dCi i A5 dt () n n Q Q Fair-sea ONP1 rain ijÆ Ci jiÆ Ci ()i ()+ , Conversely, in the middle and deep parts of oceans, Alk =-ÂÂV i +V j + V - V j =11i j = i iidecreases by 0.17 eq per mol decomposition of the or- ji ji π π ganic carbon and increases by 2 eq per mol dissolution of A1 ()CaCO3. In the middle part, the equation is, where Ci is the concentration of DIC, V is the volume of the ocean, rain is the rain ratio, and the subscript i refers dAlki to the box number. In the middle part of the ocean, the dt POC settling down and the DOC transported by the n Q ijÆ Alki advection decompose to the DIC: =-Â V i j =1 i jiπ n Q 017. ONP 1 dprato dCi jiÆ Alki 017. Co ¥-(), i +-¥Â V ()j j - V dt j =1 i i jiπ n Q n Q ONP1 dprato ijÆ Ci jiÆ CCi o ()- , ()A6 =-ÂÂV i +V j + j + V j =11i j = i () i jiπ jiπ ()A2 and in the deep part, where Co is the concentration of DOC and dprato is the dAlk n Q n Q ONP2 rain i ijÆ Alki jiÆ Alki ¥¥ . decomposition ratio of POC into DOC. Similarly the dt =-ÂÂV i +V j + V j =11i j = i i CaCO3 from the surface layer dissolves to the DIC in the jiπ jiπ deep part: ()A7

Ocean uptake potential for CO2 sequestration 43 (a) 2000 dPO n Q 4 i ijÆ 1000ppmv [] PO =- 4 i 750ppmv dt Âj 1 Vi [] 650ppmv = jiπ 1500 650ppmv 450ppmv n o

(ppmv) Q C ONP1 dprato jiÆ Ê PO i ˆ ()- . + 4 i + - CO2 Â V Á[]Rc˜ Rc V 1000 j =1 i Ë ¯ ◊ i jiπ ()A9 500

Atmospheric P On the other hand, in the deep part of the ocean, no reac- 0 2000 3000 4000 5000 6000 7000 8000 tion relates to the phosphate and its budget is expressed Date (A.D.) only based on the advection:

(b) 2000 dPO n Q n Q 4 i ijÆ jiÆ 1000ppmv [] PO PO .A10 =- 4 i + 4 i () 750ppmv dt ÂÂV [] V [] j =1 i j =1 i 650ppmv ji ji 1500 550ppmv π π 450ppmv

(ppmv) The terrestrial carbon reservoirs include the plant and

CO2 1000 the soil. These reservoirs are balanced by the net primary production and the decompositions of themselves:

500 dP TNP F , A11 dt =-CT, plant () Atmospheric P 0 2000 2500 3000 3500 4000 Date (A.D.) dS 1 kF F. A12 dt =-()◊ plant- soil () P Fig. A1. Calculated results of the future atmospheric CO2 varying the magnitude of the oceanic thermohaline circulation: (a) the original Schmitz’s circulation, (b) the five times as fast Finally, the atmospheric CO2 budget is expressed by as the case (a). the interactions with oceans and lands, in addition to the anthropogenic CO2 emissions:

The balance equation of the phosphate is expressed dCO2 based on the above DIC and alkalinity budgets by using dt the carbon to phosphate Redfield ratio Rc. In the surface 6 layer, the phosphate is consumed on the new production: FNPP kF F fossil landuse =-Â()air-seai -- plant - soil ++() i=1 () A13 n n () dPO4 Q Q ONP []i ijÆ PO jiÆ PO , dt =-ÂÂV []4 i + V []4 i - Rc V j =1 i j =1 i ◊ i where fossil and landuse are the CO emissions from fos- jiπ jiπ 2 sil fuel consumption and land-use change, respectively. ()A8

APPENDIX II. SENSITIVITY ANALYSES where [PO4] is the concentration of the phosphate. Both the POC and DOC reached to the middle part of the ocean As discussed in the subsection “Thermohaline circu- decompose there and regenerate the phosphate: lation”, our model required to triple the magnitude of the

44 M. Sorai and T. Ohsumi Schmitz’s thermohaline circulation so as to produce the P and shortens the time required to attain to the steady CO2 realistic distribution of the DIC in the ocean. For the as- state. However, it is noticed that the oceanic circulation sessment of this modification, we performed sensitivity affects little influence on the essence of the CO2 cycle. In analyses varying the magnitude of oceanic circulation. other words, the future profile of the atmospheric P CO2 The first run is based on the original Schmitz’s is essentially analogous to each other. This is also sup- thermohaline circulation, which has an apparent residence ported by the fact that the ocean CO2 sequestration based time of 3,100 years. The second run is the five times as on these modified circulation gave the analogous profile fast as the first one and its residence time is 620 years. of the atmospheric P . We estimated the CO uptake CO2 2 These series of results including Fig. 5 certainly provide potential for the above modified cases of the thermohaline us a good implication to understand the relationship be- circulation as in the section “Model Implication on Ocean tween the magnitude of ocean circulation and the oce- Sequestration”. The results showed the potentials thor- anic CO uptake rate. The change in the oceanic circula- 2 oughly consistent with Fig. 7. This means that the CO2 tion drastically affects the atmospheric P with respect CO2 uptake potential is independent of the magnitude of the to the peak P and the rate of approach to the steady CO2 oceanic thermohaline circulation. state (Fig. A1). The faster circulation reduces the peak

Ocean uptake potential for CO2 sequestration 45