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 ecosystem responses. The present study aims to reveal the effectiveness and consequences of the direct ocean CO2 sequestration in relation to both lowering the atmospheric transient PCO2 peak and reduction in future CO 2 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 CO2 uptake potential integrated with time until the end of fossil fuel era, is only dependent on the atmospheric PCO2 value in the ultimate steady state, whether or not the CO2 is purposefully injected into the ocean; we gave the total potential capacity of the ocean for the CO2 sequestration is about 1600 GtC in the case of atmospheric steady state value ( PCO2 ) of 550 ppmv. Keywords: ocean CO 2 sequestration, global carbon cycle, CO 2 uptake potential, box model, future atmospheric PCO2
tralization effect provides a further oceanic potential to uptake excess CO2 (Nozaki, 1991). The detailed outline of the ocean CO2 sequestration has been presented both from technical and scientific aspects (Ohsumi, 1995). The assessment of its long-term consequences requires the comprehensive understandings on the global biogeochemical cycle of the carbon. The modeling approach is effective in evaluation of each process governing the cycle. There have been so far several attempts to investigate the effects of the ocean CO2 sequestration and its interaction with the biogeochemical processes. Hoffert et al. (1979) showed clearly the “peak-shaving” effect of the CO2 oceanic injection on the atmospheric CO2 concentration time profile in the future. In their work, some essential processes of the carbon cycle, such as the oceanic thermohaline circulation, the biological activities, 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, which focused on improving the predictive capacity of global-scale three-dimensional ocean carbon-cycle models, conducted a series of simulation runs where the anthropogenic CO2 is injected into the several depth ranges of the selected points and revealed the relationship between the injection site and the subsequent holding ca-
INTRODUCTION To mitigate the impact of the future atmospheric carbon dioxide increase, the ocean CO2 sequestration has been proposed (Marchetti, 1977). The concept is built on the perspective that it would be hard to reduce drastically the future anthropogenic CO2 emissions without innovative technological developments. From a simple geochemical viewpoint, the ocean CO2 sequestration must be one of the most reasonable options, because the ocean originally contains both the carbonate ion CO32– and the B(OH)4– ion enough to react with about 70% of the untapped fossil fuel carbon to form the bicarbonate ion HCO3– and the boric acid B(OH)3 (Broecker, 2001). Moreover, in the case of the CO2 sequestration to the bottom ocean, the sedimentary calcium carbonate is expected to neutralize the injected CO2 to form HCO3–, and this neu-
*Corresponding author (e-mail:
[email protected]) *Presently at Mitsubishi Research Institute, Inc., 3-6, Otemachi 2Chome, Chiyoda-ku, Tokyo 100-8141, Japan. **On leave from Central Research Institute of Electric Power Industry. Copyright © 2005 by The Geochemical Society of Japan.
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 circulation and the numerical unit is sverdrups (1 Sv = 106 m3sec–1). 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 provided several useful implications on the ocean CO2 sequestration, the fundamental role of the ocean has not yet presented explicitly: that is, what is the amount of CO2 to remain in the ocean? The CO2 uptake potential in the ocean is the key to understanding the effectiveness of the oceanic CO2 sequestration. To address these questions, the global analysis like the Hoffert’s study is useful, although the model should be revised based on the latest knowledge on the carbon cycle system in the surface of the earth. In this study, we give an estimate of the potential capacity of the ocean for the CO2 sequestration based on the geochemical box model of the global carbon cycle. Our present work owed much to the Wigley’s work, in which the modeling approach allowed to explain the past changes in atmospheric CO2 concentration and to predict its future trends (Wigley, 1993). His successful methodology including the treatment of the terrestrial carbon cycle system was taken into account in our previous model (Sorai et al., 1997): it enabled us to analyze the response of the ocean CO2 sequestration qualitatively and straightforwardly. Box model approach is efficient in analyzing the effects on the atmospheric CO2 concentration and the influences on the oceanic carbon concentrations by the purposeful injection of CO2 into the ocean. 30
M. Sorai and T. Ohsumi
MODEL DESCRIPTION The three-dimensional box model consists of 19 boxes as shown in Fig. 1: atmosphere, the Arctic Ocean (surface and middle), the North and South Atlantic Oceans (surface, middle and deep for each), the Indian Ocean (surface, middle and deep), the Pacific Ocean (surface, middle and deep) and the Antarctic Ocean (surface and middle), and the land (plant and soil). We defined the surface ocean as 0–200 m depth, the middle part as 200– 2,000 m depth, and the deep part as below 2,000 m. The exceptions are the Arctic and the Antarctic oceans, where the middle part is set from 200 m depth to the seafloor. The thermohaline circulations between the ocean boxes are also given in Fig. 1. The solubility pump, the biological pump, and the alkalinity pump are considered as the mechanism of carbon cycling in the atmosphere-ocean system. The contribution of terrestrial ecosystem, which includes both CO2 fertilization and global warming effects, is also taken into account. For an n-box system, the increase rate of carbon amount M in the box i is expressed as n n dM = - Â Qi Æ j + Â Q j Æ i , dT j =1 j =1 j πi
j πi
(1)
where QiÆj stands for the carbon flux from the box i to the box j, and QjÆi, the carbon flux from the box j to the box i. The first term on the right hand side of Eq. (1) is the total carbon efflux from the box i, whereas the second term is the total carbon influx to the box i. In our model, the change in carbon amount in each reservoir is given by Eq. (1): the variables related to the carbon amount include the atmospheric CO 2 concentration ( PCO 2 ), the carbon amounts of the terrestrial plant and soil, and the concentrations of the dissolved inorganic carbon (DIC) and the dissolved organic carbon (DOC), and the alkalinity in each ocean boxes. An inverse Euler method was used to avoid the solution instabilities (Walker, 1991). The detailed model description including the balance equations for each variable had been reported elsewhere (Sorai et al., 1997). In this section, the key processes of our model are presented (also see Appendix I). Thermohaline circulation The mass balance of water between the boxes were 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 to one box and the exchange of water mass between the surface Arctic and North Atlantic boxes is added in the present treatment. It should be noticed that the diffusion process actually works for the mass and heat transports in the ocean in addition to the advection on the thermohaline circulation. Moreover, the wind-driven circulation enhances these transports in the ocean on the time scale of several decades. In our model, however, these processes between the boxes are not taken into account, and hence the mass transport is somewhat weakened; the consideration of only Schmitz’s circulation provides the unrealistic distribution of DIC concentration, such that it becomes very high in deep region. The enhancement of the model oceanic circulation would help to incorporate all these processes on the time scale of several decades. In this case, one criterion for the extent of the enhancement is the “apparent” residence time of the ocean, which is defined as the oceanic volume divided by the total downward fluxes. At the North Atlantic Ocean in Schmitz’s model the downward flux of 14 Sv (1 Sv = 10 6 m3sec–1), which corresponds to the “apparent” mean residence time of deep oceanic water more than 3,000 years, seems to be too low, if we follow the generally accepted oceanic mean residence time of about 1,000 years (Broecker and Peng, 1982). Considering that oceanic potential to hold the dissolved CO2 species have the essential role in our simulations, the oceanic residence time should be primarily consistent with the general view. Therefore, we tripled this downward flux to 42 Sv: this improvement inevitably made us to
triple all the other fluxes in accordance with the mass balance of the system. As a result, the “apparent” mean residence time was lowered to about 1,000 years in our model. This modification is justified also by the representation of DIC in the pre-industrial steady state: a good agreement was obtained by using our tripled flow pattern as discussed later. For additional justification of our modification, we performed sensitivity analyses changing the magnitude of oceanic circulation, which enabled us to realize the relationship between the oceanic potential and the circulation (see Appendix II). In fact, it was found that such a difference of the oceanic circulation had a little influence on the oceanic potential. Solubility pump The solubility pump is defined as the carbon exchange mediated by physical processes such as heat flux, advection, and diffusion. Net CO2 exchange flux across the atmosphere-ocean interface, Fair-sea (mol m–2yr–1), is expressed as follows:
{(
Fair -sea = Kkwa PCO 2
{(
= E PCO 2
)
sea
)
(
sea
- PCO 2
(
)
- PCO 2
air
)
},
air
}
(2 )
where ( PCO 2 )sea and ( PCO 2 )air are the CO2 partial pressures (ppmv) in the surface-water and in the atmosphere, kw is the gas transfer rate (cm hr–1), a is the solubility of CO2 in seawater (mol L–1atm–1), K converts units to mol m –2 yr –1, and E is the gas exchange coefficient (mol m–2yr–1 m atm–1). Here, the solubility of CO2 is a function of both temperature T (K) and salinity S (‰) (Weiss, 1974): ln a = -58.0931 + 90.5069 ◊ (100 T ) + 22.2940 ln(T 100)
[
]
+ S 0.027766 - 0.025888 ◊ (T 100) + 0.0050578 ◊ (T 100) . 2
(3)
The gas transfer rate is a function of wind speed and temperature. Several expressions for the relationship between gas transfer rate and wind speed have been used so far. These studies have shown that gas transfer rates measured over long time periods with variable winds are higher than those measured instantaneously or under steady winds of the same average wind speed. For example, the low gas transfer value of Liss and Merlivat (1986) is based on 1–2 day measurements on a lake, whereas the higher value of Wanninkhof (1992) is based on ocean 14C 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 (1992) because we mainly intend to predict long-term trends. The Wanninkhof (1992) formulation is kw = 0.39u 2 ( Sc Sc0 )
-1 2
( 4)
,
where u is average wind speed (m s–1), Sc is Schmidt number of CO2, the dimensionless quantity defined as the kinematic viscosity divided by the diffusion coefficient, 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 following relation for the temperature dependence of Schmidt number of CO2 (Liss and Merlivat, 1986):
Sc = 1860 - 120T + 4.3208T - 0.09T + 0.00079167T . 2
3
4
(5)
In fact, since it is difficult to set the average wind speed in each ocean, kw was adjusted to produce E values consistent with the estimations based on the satellite wind speed measurements (Etcheto et al., 1991). Consequently, the pre-industrial E values were 0.073 at the Arctic, 0.086 at the Antarctic, and 0.064 at the Atlantic, Indian and Pacific Oceans. Furthermore, the global warming effect was also incorporated in the solubility pump of our model. The temperature increase due to the greenhouse effect could be related to the atmospheric PCO 2 , as defined by
DT = 2.5 ¥
(
log PCO 2 280 log 2
),
(6 )
where DT (K) is the mean temperature increment relative to the pre-industrial level (Stocker and Schmittner, 1997). Equation (6) is based on the assumption that when the atmospheric PCO 2 reaches 560 ppmv, which is twice the pre-industrial value, the global mean surface temperature will increase by 2.5∞C (Houghton et al., 1995). The temperatures of only surface ocean boxes are uniformly increased by DT. It might be well justified that although the simulation studies so far had revealed that the global temperature increase is heterogeneous over the earth surface, i.e., more enhanced on the high-latitudinal areas in both the north and south hemispheres, our pre-checking showed that the heterogeneity of the temperature increase due to the global warming is likely to have little influences on the calculated PCO 2 . Biological pump The biological pump is the process by which CO2 fixed 32
M. Sorai and T. Ohsumi
in photosynthesis is transferred to the interior of the ocean resulting in a temporary or permanent sequestration of carbon. This process is controlled by the new production, where the organic carbon photosynthesized on the surface layers of the ocean settles down to be decomposed and regenerates the inorganic carbon again in the middle depth part of the box. We assumed that the magnitude of the new production ONP (mol m–3yr–1) is limited by its phosphate concentration [PO4] (mol m–3) and is controlled by Michaelis-Menten kinetics: ONP = LC ◊
R ◊ Rc ◊ [ PO4 ] , H + Rc ◊ [ PO4 ]
(7)
where LC is the latitudinally varying incident light factor (yr–1), Rc is the carbon to phosphate classical Redfield ratio (=106), and R and H are adjustable parameters (Bacastow and Maier-Reimer, 1990). The ratio of the particulate organic carbon (POC) to the dissolved organic carbon (DOC) was adjusted to represent the DOC concentrations in each ocean consistent with the GEOSECS data. There have been no attempts to incorporate the oceanic CO2 fertilization effect into a carbon cycle model so far, because the oceanic production rate is generally regarded to be nutrient-limited. However, some recent works revealed that at least several kinds of marine phytoplankton increased with an increase in PCO 2 (Riebesell et al., 1993). Considering that our simulation treats a very large range of the atmospheric PCO 2 , from 280 ppmv to more than 1000 ppmv, it would be appropriate to incorporate the oceanic CO2 fertilization effect; the effect was taken into account in the model just in the same manner as in the case of the terrestrial ecosystem discussed later. We introduced the enhancement factor r0, which is defined as the ratio of ONP at PCO 2 =680 ppmv to ONP at PCO 2 =340 ppmv. Our incorporation of the fertilization effect in the model might represent the other anthropogenic impacts on the oceanic production; in the coastal region near the big cities, the continual nutrient input after the Industrial Revolution raises the oceanic production rate higher than that at the past steady state. Alkalinity pump There exists another driving force for the carbon transport into the oceanic interior via both production and dissolution of calcium carbonate (CaCO3). This mechanism is generally called the alkalinity pump. In our model, the alkalinity pump works only in the uppermost and lowermost boxes in each oceanic basin: the reaction, Ca2+ + 2HCO3– = CaCO3 + CO2 + H2O, goes to the right hand side in the euphotic layer by photosynthesis and to the left hand side in the ocean bottom instantaneously,
while the calcium carbonate (CaCO3) particles travel without any dissolution in the middle depth part of the ocean, because the lysocline which is defined as the uppermost depth of the calcite dissolution is generally below 2,000 m depth. The rain ratio, i.e., the CaCO3 production divided by 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 included no effect of the difference of the dissolution depth between CaCO3 and POC produced by the new production. On the other hand, Yamanaka and Tajika (1996) revealed that in fact the ratio was considerably reduced to approximately 0.09, when the difference of the dissolution depth was taken into account. Since our model assumed that POC and CaCO3 dissolved in the middle and deep depth parts of box respectively, we adopted the ratio of 0.09. As is the case in the biological pump, the CO2 effect which reduces the calcification of the marine plankton in response to the PCO 2 increase is taken into account in the formulation of the alkalinity pump in the model: PCO 2 - 280 ˆ Ê rain = rain0 ◊ Á1 - krain ◊ ˜, 750 - 280 ¯ Ë
(8)
where rain is the rain ratio, krain is the normalized decrease index of the ratio of calcification to POC production, and subscript 0 denotes the value in the preindustrial level. Equation (8) is based on the experiment that the rain ratio decreased up to 52.5% with an increase in PCO 2 from 280 ppmv to 750 ppmv (Riebesell et al., 2000). This mechanism is easily understood, because the PCO 2 increase in the surface water makes the above reaction to go to the left hand side. Terrestrial ecosystem The terrestrial carbon cycle system was built up by simplifying Wigley’s treatment (Wigley, 1993), that is, the terrestrial components are only plant and soil in our model. The land biota works as the CO2 sink through the net primary production TNPC (GtC yr–1) enhanced by the atmospheric PCO 2 increase. This CO 2 fertilization effect is generally formulated by either the logarithmic form (Bacastow and Keeling, 1973) or the Michaelis-Menten form (Gates, 1985) with respect to the atmospheric PCO 2 . However, Gates (1985) suggests that plant photosynthesis and ultimately growth actually does not follow the logarithmic response to the CO2 increase, but that the Michaelis-Menten form behaves more realistically on the point such that it leads to zero net primary production at the low PCO 2 and has a limiting value at the high PCO 2 . Thus, we adopted the Michaelis-Menten form for TNPC:
TNPC = TNP0
(P - (P ) ){1 + b((P ) - (P ) )} ((P ) - (P ) ){1 + b(P - (P ) )} CO 2
CO 2
CO 2
CO 2
0
CO 2
b
b
0
CO 2
b
CO 2
CO 2
b
( 9)
b=
(680 - (P ) ) - r (340 - (P ) ) , (r - 1)(680 - ( P ) )(340 - ( P ) ) CO 2
l
b
CO 2
l
b
CO 2
CO 2
b
(10)
b
where ( PCO 2 )b is the atmospheric PCO 2 at the CO2 compensation point (=31 ppmv: Gifford, 1993) where net daytime photosynthetic CO2 fixation is zero, and r l is defined as TNP(680 ppmv)/TNP(340 ppmv). The organic carbon in the terrestrial plant will be decomposed. The decomposition rate Fplant was divided into two fluxes: one represents the fraction to the atmospheric CO2 and the rest to the soil, while the ratio of these two fluxes was defined by an adjusting parameter k in the model. Similarly, the soil organic carbon will decompose and return to the atmospheric CO2. These decomposition fluxes, Fplant and FluxsoilT, were assumed to be proportional to the plant and soil masses, respectively. Moreover, the global warming effect was applied for the net primary production and the decomposition of organic carbon in the soil, respectively as, TNPC, T = TNPC (1 + kNPP DT )
(11)
FluxsoilT = Fluxsoil (1 + ksoil DT ),
(12)
and
where TNPC,T and FluxsoilT are the net primary production (GtC yr–1) and the soil decomposition fluxes (GtC yr–1) as a function of the warming effect, and kNPP and ksoil are the warming factors on TNP and on Fluxsoil, respectively (Gifford, 1993). MODEL EVALUATION Initial condition and input data The pre-industrial oceanic variables, such as the DIC concentrations, the DOC concentrations, the alkalinities and the temperatures, were set so as to meet the requirement as follows: 1) the pre-industrial atmospheric PCO 2 is 280 ppmv, and 2) the pre-industrial carbon amounts of terrestrial plant and soil are 610 and 1560 GtC, respectively (Siegenthaler and Sarmiento, 1993). The model was started with the global data sets of GEOSECS (Takahashi Ocean uptake potential for CO2 sequestration
33
Table 1. Pre-industrial variables for each ocean box Ocean
Arctic North Atlantic
South Atlantic
Indian
Pacific
Antarctic
surface middle surface middle deep surface middle deep surface middle deep surface middle deep surface middle
DIC (mol m– 3 )*
Alkalinity (mol m– 3 )
DOC (mol m– 3 )
Temperature (°C)
2.219 2.270 2.155 2.244 2.266 2.148 2.946 2.330 2.139 2.480 2.461 2.171 2.528 2.470 2.261 2.465
2.366 2.375 2.366 2.351 2.375 2.366 2.351 2.441 2.394 2.434 2.437 2.408 2.426 2.445 2.394 2.435
0.062 — 0.062 — — 0.058 — — 0.048 — — 0.039 — — 0.048 —
3.6 0.7 17.1 8.1 2.2 22.5 9.7 1.8 26.6 11.9 1.8 22.6 10.8 1.5 3.6 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).
Fig. 2. Anthropogenic CO2 emissions both from the industrial sources and from the net land-use change. The CO2 emission from land-use change was assumed to start in the year 1700 and to linearly increase until the year 1850.
et al., 1981), except for Ymer80 data on the Arctic Ocean (Anderson and Dyrssen, 1981). The spin-up run under the above boundary conditions over a million year made the model to the steady state, so that the oceanic variables drifted to the pre-industrial steady state values, which are specific to our model. The model pre-industrial values for each ocean box are shown in Table 1. The historical trends of anthropogenic CO2 emissions are given by Online Trends, year-by-year data sets by Carbon Dioxide Information Analysis Center (Online TREND). As shown in Fig. 2, these include the global emissions from the fossil fuel combustion and the cement 34
M. Sorai and T. Ohsumi
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 leads to the assumption that the pre-industrial period before 1700 was a steady state (Ver et al., 1999), which is also supported by the measurements from ice cores such that the atmospheric PCO 2 had been almost constant within a range of 10 ppmv in this period. The anthropogenic CO2 emissions were directly input to the atmospheric CO2 reservoir in our model after the year 1700. Comparisons between calculated and observed values For validation of the model, ideally, all the carbon components should be evaluated on their mass changes by the comparison between the calculated results and the observed data. Of all the carbon reservoirs in the model, however, only the atmospheric CO2 has been known on the time course of concentration change from the preindustrial level to the present time. Therefore, several parameters depicted in the section “Model Description” were adjusted primarily aiming at the representation of atmos-
Table 2. Adjustable parameters in the model Parameter
Definition
Adopted value
ro rain0 k rain d rl k k NPP k soil
CO2 fertilization factor on NP Rain ratio in steady state Normalized decrease index of the ratio of calcification to POC production Decomposition fraction of POC to DOC CO2 fertilization factor on NPP Fraction of direct decomposition to air in plant degradation Warming factor on NPP Warming factor on Fluxsoil
1.1 0.09 0.525 0.15 1.14 0.87 0.01 (K– 1) 0.03 (K– 1)
Literature value 0.08–0.101 ) 0.21–0.5252 ) 1.1–1.43 ) 0.47–0.873 ) 0.01–0.053 ) (K– 1) 0.03–0.053 ) (K– 1)
1)
Yamanaka and Tajika, 1996. Riebesell et al., 2000. 3) Gifford, 1993. 2)
380
Atmospheric PCO2 (ppmv)
pheric PCO 2 changes. These parameters were finally evaluated as shown in Table 2. In this study, our model was checked with respect to several fundamental variables: the time-course profile of atmospheric PCO 2 , the DIC concentration and the alkalinity in 1970’s, and the oceanic CO2 uptake rate in 1990’s. (1) Time-course profile of CO2 concentration The timecourse profile of the atmospheric CO2 concentration was simulated using the values of reservoirs, fluxes and adjustable parameters defined in the section “Model Description”. The anthropogenic CO2 emissions data described above was applied. The calculated result is shown in Fig. 3, with the data from the Siple ice core and the air measurements in Mauna Loa (Online TREND). Our results after the year 1970 are relatively well consistent with the observations, while some deviations exist especially in the period of the year 1850 to 1950. We found that the 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 past change in the atmospheric PCO 2 , it certainly accompanies the uncertainties on the future trend, that is, the TNP and thus the amount of plant continue to increase synergistically by both effects of CO2 fertilization and global warming; in turn, this enhanced TNP causes the abrupt drop of the atmospheric PCO 2 to the unrealistic lower level. Therefore, the modified form of Eq. (9) was not used in this study. Nevertheless, considering that our study mainly focuses on the future CO2 trend, it would be more appropriate that the model is tuned to the recent change of CO2 concentration rather than to the past uncertain trend. (2) DIC concentration and alkalinity The GEOSECS data were used for checking of both the DIC concentration and the alkalinity. Table 3 shows the comparison between the GEOSECS data and the model values at the time corresponding to the GEOSECS expeditions. All calculated results of the DIC concentration and the alkalinity exhib-
Calculated resutls in this work Siple Mauna Loa
360
340
320
300
280
260 1650
1700
1750
1800
1850
1900
1950
2000
Date (A.D.)
Fig. 3. The time-course change in the atmospheric PCO2 . The calculated result shows a good agreement with the observation especially after the year 1970, whereas there exist some deviations in the period of the year 1850 to 1950.
ited qualitatively a good agreement with the GEOSECS data. Especially, the alkalinity in each part of oceans agreed to well within 2%. The exceptions are in the surface and middle parts of the Indian Ocean, but even in these boxes the deviation was within 3%. The DIC concentrations exhibited relatively large deviations, mostly within about 10%, especially at all surface parts and a part of middle depth regions, whereas the deviations of the deep boxes were reduced to 3 to 4%. Extremely higher value was obtained in the middle part of the South Atlantic Ocean; this would be the result of the active decompositions of the settling-down POC, because the large upwelling flow in this area transports the abundant phosphate to the surface layer and causes the higher new production using them. These deviations are likely related to the magnitude of the oceanic thermohaline circulation in our model. In other words, the active oceanic circulation would stir around the whole ocean and dilute a high DIC concentration in the middle parts of oceans. As disOcean uptake potential for CO2 sequestration
35
Table 3. Comparison between the GEOSECS data and the model values DIC (mol m– 3 )
Ocean
Arctic North Atlantic
South Atlantic
Indian
Pacific
Antarctic
surface middle surface middle deep surface middle deep surface middle deep surface middle deep surface middle
Alkalinity (mol m– 3 )
This work
Observation
This work
Observation
2.272 2.176 2.252 2.269 2.166 2.948 2.331 2.165 2.480 2.461 2.192 2.528 2.470 2.279 2.466
— 2.240* 1.993 2.177 2.236 2.010 2.196 2.267 1.983 2.240 2.372 2.027 2.241 2.391 2.226 2.313
2.375 2.369 2.352 2.375 2.369 2.350 2.440 2.396 2.434 2.437 2.409 2.426 2.445 2.395 2.435
— 2.389* 2.344 2.360 2.386 2.352 2.373 2.407 2.340 2.370 2.469 2.366 2.395 2.483 2.413 2.427
*These data were obtained at the Ymer80 expedition.
cussed in the subsection “Biological pump”, we tripled all flows of Schmitz’s circulation in order to represent more realistic residence time of deep water. Nevertheless, our oceanic circulation might still possibly be lower. Although this point should be considered in more detail, considering that our main interest of the oceanic potential is independent on the magnitude of the oceanic circulation, namely on the representation of the DIC distribution, the above deviations in the DIC concentration would be within tolerable limits. (3) Oceanic CO2 uptake rate The analysis of the oceanic and terrestrial CO2 uptake rates provides an important criterion on the evaluation of the model. Intergovernmental Panel on Climate Change (IPCC) summarized the global CO2 budget: the CO2 uptake rates in 1990s as 1.7 ± 0.5 GtC yr–1 in the ocean and as 1.4 ± 0.7 GtC yr–1 in the land (Houghton et al., 2001). Considering that the terrestrial system still has a lot of uncertainties and that our main purpose is to examine the oceanic CO2 sequestration, we focused on the oceanic CO2 uptake rate in this study. Since the solubility pump determines the magnitude of the CO2 flux from the atmosphere to the ocean, the oceanic CO2 uptake rate was estimated using Eq. (2) and each oceanic area. The averaged value of this uptake rate in 1990’s was 1.3 GtC yr –1. This rate corresponds to almost the lowermost value in the compilation of IPCC (2001). In fact, we found that more active oceanic circulation produced larger CO2 uptake rate in the ocean, but this modification would be beyond the purpose in this study. The new production in the ocean becomes another important criterion for validation of the model biological
36
M. Sorai and T. Ohsumi
pump. Our estimate of the new production in 1980’s is 9.9 GtC yr–1 and this is well consistent with the estimate of 11 GtC yr–1 in IPCC (2001). MODEL I MPLICATION ON OCEAN SEQUESTRATION Scenarios for the future (1) Future emission scenario The B2 case of the Special Report on Emissions Scenarios (SRES) in IPCC (2001), which assumed the world where the emphasis is on local solutions to economic, social, and environmental sustainability, was applied for the anthropogenic CO2 emissions from the year 2010 to 2100 (Houghton et al., 2001). After the year 2100 the emission scenario is rather arbitrary. One effective approach is to set the future emission scenario limited by the total natural resources available to us. The global fossil fuel energy resource summarized by Rogner (1997) was 4911 GtC. If the industrial CO2 emission rate is assumed to decrease linearly after the year 2100, the final year when the fossil fuel consumption is completed becomes the year 2652. Another criterion is to fix the atmospheric PCO 2 level in the future steady state. Once a steady state PCO 2 is set to any target value in the stabilization scenarios, we can estimate the whole carbon amount in the steady state; the difference between the total carbon amount in the year 2100 and that in the steady state equals to the carbon resources allowed for us to use. In this study, the steady state PCO 2 is set to 450, 550, 650, 750 and 1000 ppmv; these PCO 2 values agreed with those in the stabilization scenarios in IPCC (1995) (Houghton et al., 1995). Also in these cases, we assumed that the CO2 emission from
(a)
14 450ppmv
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Fig. 4. Future estimates of the atmospheric PCO2 . (a) Future 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 were linearly decreased to satisfy the total carbon emissions reaching to the steady state PCO2 values of 450, 550, 650, 750 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 atmospheric PCO2 . The atmospheric PCO2 were calculated us-
of the ocean CO2 sequestration was set to the year 2015. The duration for the oceanic CO2 sequestration was assumed until the end year of the fossil fuel consumption. The amount of the CO2 sequestered was varied from 20 to 100% of the annual CO 2 emission from the fossil fuel combustion. GOSAC project simulated the ocean CO2 sequestration in seven offshore sites (GOSAC, 2002). We assigned these sites to our box oceans: Bay of Biscay and New York belong to the North Atlantic Ocean, Rio de Janeiro to the South Atlantic Ocean, Bombay and Jakarta to the Indian Ocean, and San Francisco and Tokyo to the Pacific Ocean. Based on this ratio of 2:1:2:2, the prescribed amount of CO2 in the form of DIC was injected into the middle parts of these oceans. In fact, the ocean sequestration will cause the additional CO2 emission, because the CO2 capture process requires the extra energy consumption. David and Herzog (2000) analyzed the economics of capturing CO2 at Integrated coal Gasification Combined Cycle (IGCC) power plants, Pulverized Coal (PC) power plants, and Natural Gas Combined Cycle (NGCC) power plants, and estimated the additional energy requirements in the year 2012 as 9% for IGCC, 15% for PC, and 10% for NGCC, respectively. Thus, we assumed that the extra 10% of the sequestered CO2, which is though a little bit lower than 15% at PC, is added to the original emission scenario as defined by the following expression, CO 2 ,seq = EB2 ¥ r
(13)
Eext = ( EB2 ¥ r ) ¥ 0.1,
(14)
and
ing the above anthropogenic CO2 emission scenarios. All cases once increased to the peak value and then decreased to the steady state.
where CO2,seq is the annual amount of CO2 sequestered, EB2 is the industrial CO2 emission rate defined above, r is the ratio of CO2 sequestered to the scenario emission, and Eext is the extra CO2 emission due to CO2 capturing. 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 2100 to satisfy the total carbon emission defined above. Thus, the end year when the industrial CO2 emission is completed varies dependent on each scenario: the end years are the year 2163, 2274, 2371, 2457, and 2646, respectively. The last 1000 ppmv-stabilized scenario is roughly consistent with the case based on the fossil fuel resources. On the other hand, the net CO2 emission from the landuse change was assumed to become zero after the year 2200. These emission scenarios are shown in Fig. 4(a). (2) Ocean sequestration scenario In all cases, the start
Future PCO2 trend The calculated trends of the atmospheric PCO 2 greatly depend on the future scenario of the anthropogenic CO2 emissions. Figure 4(b) is the calculated results based on various scenarios shown in Fig. 4(a). Each emission scenario had a different total amount of CO2 emission, and thus resulted in the different atmospheric PCO 2 level at the ultimate steady state. Here, we have no definite evidence that the system will finally attain to the steady state. But our model assumes that no burial processes of POC and CaCO3 occur in the ocean, so that once the CO2 input Ocean uptake potential for CO2 sequestration
37
Ocean CO2 sequestration Next, we examined the effect of the ocean CO2 sequestration on the atmospheric PCO 2 based on the 550 ppmv stabilized case in Fig. 4(b). Figure 5 shows the calculated results of the CO2 sequestration, where the sequestered portion of the total industrial emissions was varied from 20 to 100%. In each case, the starting year of the CO2 sequestration was set to the year 2015. The prescribed portion of the fossil fuel CO2 was injected to the middle parts of the oceans during the CO2 sequestration and the rest was released to the atmosphere. Since the CO2 sequestration was assumed to proceed until the end year of fossil fuel consumption, the end year was different dependent on the amount of CO2 sequestered. As shown in Fig. 5, ocean sequestration obviously decreases a peak value of the atmospheric PCO 2 . We can simply 38
M. Sorai and T. Ohsumi
1000
Atmospheric PCO2 (ppmv)
from anthropogenic sources are stopped, the system inevitably attains to the steady state and thus the atmospheric PCO 2 is stabilized at a constant value. Moreover, the climate change, which might perturb the system and force to change over to another steady state, is not thoroughly incorporated in the model, except for the global warming effects resulting in the activity change of terrestrial ecosystem and the intensity of solubility pump. Such simple treatments are useful to understand the CO2 fate more explicitly. In all cases in Fig. 4(b), the atmospheric PCO 2 once increased over the steady state value and then decreased gradually. It was found that the rate of approach towards the steady state and the peak PCO 2 value both depend mainly on the magnitude of oceanic circulation: the slower circulation which causes the higher peak value of the atmospheric PCO 2 requires the longer time to attain to the steady state and vice versa (Appendix II). However, whatever the strength of oceanic circulation, the atmospheric PCO 2 shows the essentially analogous profiles, and thus provides less influence on the following discussions. It should be noticed that all the peak PCO 2 reached about the twice of the corresponding steady state values. 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 PCO 2 under the steady state value, in other words, for the realization of the PCO 2 time course with no transient peaks, we should mitigate the burden to the atmosphere. The simplest solution for this mitigation is either direct CO2 injection to the ocean or reduction of CO2 emission itself. Considering that there still exist a lot of difficulties to the large-scale reduction of our future emissions, direct CO2 injection to the ocean might be reasonable in a sense such that the reduction of a CO2 release to the atmosphere is achieved by acceleration of oceanic CO2 absorption.
100% Air 20% (2015-2241)
900
40% (2015-2227) 60% (2015-2217)
800
80% (2015-2209) 100% (2015-2202)
700 600 500 400 300 2000
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Fig. 5. The dependence of sequestered portion of fossil-fuel CO2 on the atmospheric PCO2 . In all cases, the steady state PCO2 was set to 550 ppmv and the prescribed portion of total industrial emissions were injected in the ratio of 2:1:2:2 to the middle depth parts of the following four oceans: the North and South Atlantic Oceans, the Indian Oceans and the Pacific Ocean. The sequestered CO 2 was varied between 20% and 100%. The CO2 was injected from the year 2015 to the end year of fossil fuel consumption (the end year was different dependent on the amount of CO2 sequestered).
understand the reason as a result of the reduction of the CO2 emissions to the atmosphere. Therefore, this effect is more enhanced by the larger-amount and longerperiod sequestration. The oceanic CO2 injection with the portion less than 80% caused the peak over 550 ppmv, while the CO2 injection more than 80% exhibited no peak. The DIC concentrations in the middle Pacific are also depicted in Fig. 6(a). The other CO2-sequestered oceans, such as the middle parts of the North and South Atlantic Oceans and of the Indian Ocean, also have the analogous time scale of the change in the DIC concentration, although 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 starting of sequestration and then slowly changed to the final steady state. Such a stabilization of the DIC concentration is caused by the thermohaline circulation, which transports the high-DIC water produced by the ocean sequestration to the whole ocean. As expected from Fig. 1, most of the CO2 injected to the middle parts of oceans is once transported to the deep parts, and then spread to the whole ocean on the horizontal or vertical flows. The only path to the surface layer is the South Atlantic Ocean and the Pacific Ocean. The high-CO2 water reaching to the surface layer raises the PCO 2 in the surface ocean, and consequently reduces the CO2 uptake rate. Especially, in the case more than 80% of sequestration, it was found that the net release of CO2 from the ocean occurs.
6000
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Fig. 7. CO2 uptake potential in the ocean after the year 2015. The difference between the final and the present carbon amount in the ocean correspond to the oceanic CO2 uptake potential. This potential is determined by the atmospheric PCO2 value in the ultimate steady state. The factors to affect the CO2 uptake potential are the temperature of seawater, the new production, and the rain ratio.
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Fig. 6. The change in the DIC concentration due to the ocean CO 2 sequestration. The changes in the DIC concentration were estimated (a) in the middle depth part of the Pacific Ocean, and (b) in the surface Pacific Ocean. Each calculation corresponds to the cases in Fig. 5.
Owing to the future increase in the atmospheric PCO 2 , however, the DIC concentration is also raised even without direct CO2 injection. In other words, the CO2 sequestration only brings about the future increase in the DIC concentration earlier, at least in the case less than the injection of 80% portion. If the biological impact to this oceanic domain is the same for these curves including the business-as-usual scenario (100% Air), the mitigation of the land and surface ocean impact is the net benefit of the ocean sequestration. We can see these benefits in Fig. 6(b), where the corresponding DIC concentration profiles in the surface Pacific are shown. It is noteworthy that since most of marine organisms are concentrated within the depth range shallower than 200 m, the DIC increase in surface ocean should be suppressed preferentially. CO2 uptake potential As discussed above, the total carbon amount on the surface domain of our planet is decided uniquely in the future steady state of the atmospheric CO2 concentration,
if no CO2 reaction with solid earth is taken into account. Therefore, when we set any PCO 2 value at the future steady state, the final carbon amount in the ocean will be also fixed. The difference between the final and the present carbon amount corresponds to the maximum CO2 amount to invade into the ocean that is, the oceanic CO 2 uptake potential integrated with time until the end of our fossil fuel era. Figure 7 shows this potential at the various steady state PCO 2 . 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 PCO 2 = 550 ppmv. Hence, we conducted two additional simulation runs, in which CO2 equivalent to 1623 GtC was injected to the middle and deep parts of the four oceans in the duration from the year 2015 to 2145. It was found that the calculated PCO 2 does not exceed 550 ppmv level in the sequestration to the deep ocean (Fig. 8). But in the middle depth sequestration, PCO 2 rose slightly over 550 ppmv level until the year 3500. This unexpected phenomenon is caused by the CO2 transport from the ocean. Here it should be noticed that this CO2 transport to the air is not completely equal to the leakage of the sequestered CO2 itself. The CO2 injection raises the HCO3– concentration in the ocean and thus also increases the PCO 2 in the surface water before they were homogeneously diluted in the whole ocean. This will weaken the oceanic CO2 absorption, because the solubility pump is proportional to the PCO 2 difference between in the air and in the ocean. If the CO2 injection is further continued, the PCO 2 in the surface water will become larger than the atmosOcean uptake potential for CO2 sequestration
39
3.0 Middle ocean
650
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PCO2 increment
Atmospheric PCO2 (ppmv)
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700
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Fig. 8. Calculated PCO2 in the case of 1623 GtC CO2 injections in the middle or deep depth parts in the four oceans. The 1623 GtC CO2, which correspond to the oceanic uptake potential at the steady state PCO2 = 550 ppmv, was injected in the middle or deep depth parts in the four oceans. It should be noticed that, in the middle depth sequestration, PCO2 rose slightly over 550 ppmv level until the year 3500.
Fig. 9. The ratio of the PCO2 in the middle depth parts of the Pacific Ocean in the ultimate steady state to that in the year 1999. The basal PCO2 value in the middle Pacific Ocean is about 1200 ppmv. PCO2 in the seawater will be raised owing to the future increase in the atmospheric PCO2 level even without direct CO2 injection.
pheric PCO 2 and the net CO2 flux from the ocean to the air will be caused. The CO2 injection to the deep ocean also produces the CO2 transport to the air, but its amount is slight. Our simulation reveals that both the site and the scenario are also important factors on the ocean CO2 sequestration. By simple consideration, we expect that the oceanic CO2 uptake potential might be uniquely determined by the solubility of CO2 in seawater. However, our sensitivity analyses revealed that the uptake potential is regulated by various factors, not only by the temperature of seawater, but also by the magnitude of the new production and the rain ratio. These were related to each pump mechanism explained in the section “Model Description”. At lower temperature, the solubility of CO2 in seawater increases and thus the solubility pump enhances the oceanic CO2 uptake. It should be noticed that the gas exchange coefficient E in Eq. (2) and the magnitude of the oceanic circulation do not affect the uptake potential, because they only change the rate of the CO2 exchange between the air and the ocean (see also Appendix II). The photosynthesis in the surface ocean also plays an important role in the transport of the atmospheric CO2 into the ocean. Although the new production is generally controlled by the nutrient concentration, it is possible that any future perturbation such as the CO2 fertilization effect as in this study affects the CO2 uptake potential. The alkalinity pump works as an inverse sense, that is, the higher rain ratio reduces the uptake potential, because the CaCO3 production raises the partial pressure of CO2 as discussed in the subsection “Alkalinity pump”.
Broecker (2001) estimated the total buffer capacity of the current ocean as more than 2500 GtC based on the amounts of both CO32– and B(OH)4– which react with CO2 to form HCO3– and B(OH)3, respectively. Although the CO32– concentration shows a good estimate of the CO2 carrying capacity of seawater (Cole et al., 1993), this only provides the maximum oceanic potential, because the CO2 injection to the ocean increases not only the HCO3– concentration but also the PCO 2 in the ocean simultaneously. Therefore, even if all the CO32– ions react with the injected CO2 to form HCO3–, the resultant PCO 2 increase will cause the CO2 release to the air. In other words, the oceanic capacity of the CO2 absorption is essentially regulated by the CO2 exchange across the atmosphere-ocean interface. The ocean CO2 sequestration still leaves room for concerns on the irreversible impact on the marine ecosystem. Specifically, the PCO 2 increase in the ocean, rather than the DIC increase (that is pH decrease), is one of the most important concerns on the assessment of the biological impacts and some recent works have focused on this problem (e.g., Shirayama, 1997; Kikkawa et al., 2004). However, it should be emphasized again that the PCO 2 in the seawater will be raised owing to the future increase in the atmospheric PCO 2 level even without direct CO2 injection and that the PCO 2 in the ocean increases with an increase in the steady state PCO 2 . Figure 9 shows the ratio of the PCO 2 in the middle depth parts of the Pacific Ocean in the ultimate steady state to that in the year 1999; the base PCO 2 value in the middle Pacific 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 PCO 2 of 550 ppmv, the oceanic PCO 2 will reach up to about 1.7 times of the present value. Here the most important point is whether this range of the PCO 2 increase is acceptable to the marine organisms or not. Naturally, we can think that the ocean CO2 sequestration will accelerate the oceanic PCO 2 increase and the rate of this PCO 2 increase must give some effect on the marine organisms. Nevertheless, it should be noted that if we know the limit condition of the PCO 2 increase on the impact to the marine organisms, the maximum atmospheric PCO 2 and also the maximum CO2 amounts allowed to be injected to the ocean can be determined from the biological viewpoint. CONCLUSIONS In order to assess the implication of the ocean CO2 sequestration, we conducted a simulation study using our global carbon cycle box model consisting of 19 reservoirs. The model incorporates all the essential processes in the global biogeochemical carbon cycles, such as the ocean thermohaline circulation, the solubility pump, the biological pump, the alkalinity pump and the terrestrial ecosystem. The present major modifications from our previous model include the inclusion of the terrestrial ecosystem, in which the negative feedbacks both of the CO2 fertilization and of the global warming on the net primary production and the positive feedbacks of the global warming on the decomposition process of the soil organic carbon are taken into account. Moreover, we also introduced the CO2 increase effects on the oceanic new production and on the marine CaCO3 production. These improvements enabled us to simulate the future CO2 fate in the long term using the more realistic carbon cycle system. Model parameters were validated with respect to the timecourse profile of the atmospheric PCO 2 , the DIC concentration and the alkalinity in 1970’s, and the oceanic CO2 uptake rate in 1990’s. Using this model, the CO2 injection into the middle part of the ocean was simulated varying the sequestration period and the amount of CO2 sequestered. It was found that if we accept the 550 ppmv stabilization level as a future steady state, the CO2 injection more than 80% of total anthropogenic emission is required after the year 2015 not to exceed this level. In addition, we can show 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. Moreover, for the minimization of the influences on a marine ecosystem, the ocean sequestration has an important role to suppress a decrease in the pH value in the surface ocean. We estimated the CO2 uptake potential in the ocean based on total carbon amount in various steady state PCO 2 . These estimates provide us the
maximum CO2 amount to invade into the ocean by way of either natural sink processes or purposeful injection. The present study showed that the oceanic thermohaline circulation should be parameterized in more detail to represent the more realistic distribution of the DIC concentration. These improvements may include the partitions of the model reservoirs both horizontally and vertically. It should be noticed that the present model does not treat the following well-known processes of carbon cycle. The sedimentation processes of the organic matter and the CaCO3 particles, and thus the carbon reservoir of the deep sea sediment itself, were not taken into account in the model. It is easily expected that especially the CaCO3 has an essential role on the neutralization of the increased CO2 in the ocean. Such an effect encourages the ocean CO2 sequestration at least on the point of the mitigation of the biological impacts, because the marine CaCO3 will probably enhance the neutralization of the high-CO2 waters due to the CO2 injection. In addition, Archer et al. (1998) revealed that the neutralization by the CaCO3 in the deep sea sediment occurs on a timescale of about 5000 years and that this mechanism has an ability just enough to neutralize the 60–70% of the fossil fuel release. Therefore, it is possible that the ocean holds the much larger potential to uptake the excess CO2 than the present estimate, when the neutralization effect was also considered in the simulation. On the other hand, our study assumed that climate change brings about only the temperature increase resulting in the activity change of terrestrial ecosystem and the intensity of solubility pump. In the long-term range more than hundreds of years, however, it is possible that the climate change has the significant influences to the biological systems; these include the positive feedbacks due to the global warming of the oceanic new production and of the mineralization of the organic carbon (Klepper and Haan, 1995). Moreover, several simulation studies have pointed out that the atmospheric PCO 2 increase causes the reduction or the stopping of the oceanic thermohaline circulation (Manabe and Stouffer, 1993; Haywood et al., 1997; Stocker and Schmittner, 1997). If such changes occurred in the future, the global carbon cycle system will be drastically changed (Mackenzie et al., 2000), not only in the oceanic environments but also in the terrestrial system, and the ocean CO2 sequestration might show the another aspect entirely different from the present work. Acknowledgments—This work was performed as a part of the Research and Development on CO2 Ocean Sequestration Project supported by Ministry of Economy, Trade and Industry (METI). We are grateful to Dr. Masamichi Ishikawa for many useful comments on the construction of the global carbon cycle box model. We thank Dr. Shigeo Murai for helpful comments on the ocean sequestration scenario, and Prof. Abraham Lerman and an anonymous reviewer for useful comments on a draft Ocean uptake potential for CO2 sequestration
41
version of this paper. We also appreciate helpful discussions with Prof. Naohiro Yoshida.
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APPENDIX I. BALANCE EQUATIONS The DIC in the surface layer of the ocean is balanced by the CO2 exchange with the atmosphere, and the biological productions of the POC and CaCO3, in addition to the advection: dCii dt n
Qi Æ j
j =1 j πi
Vi
= -Â
n
Qj Æi
j =1 j πi
Vi
Cii + Â
C ij +
( Fair -sea )i Vi
-
ONP(1 + rain) , Vi
(A1)
where Ci is the concentration of DIC, V is the volume of the ocean, rain is the rain ratio, and the subscript i refers to the box number. In the middle part of the ocean, the POC settling down and the DOC transported by the advection decompose to the DIC:
n Q n Q dCii ONP ◊ rain iÆ j j Æi Cii + Â C ij + = -Â . dt V V Vi i i j =1 j =1 j πi
n
Qi Æ j
j =1 j πi
Vi
= -Â
n
Qj Æi
j =1 j πi
Vi
Cii + Â
(C
i j
)
+ C oj +
ONP(1 - dprato) , Vi
(A2)
where Co is the concentration of DOC and dprato is the decomposition ratio of POC into DOC. Similarly the CaCO3 from the surface layer dissolves to the DIC in the deep part:
j πi
In this study, the DOC was defined only in the surface layer, where a part of POC decomposes to DOC: n Q n Q dCio ONP ◊ dprato j Æi iÆ j Cio + Â = -Â . C oj + dt V V Vi i i j =1 j =1 j πi
(A4)
j πi
Also on the alkalinity, the processes responsible for the budget are essentially identical to the above DIC cases. We assumed that the alkalinity, Alk, increases by 0.17 eq per mol production of the organic carbon and decreases by 2 eq per mol production of CaCO3. Therefore the balance equation of Alk in the surface layer is expressed as follows: dAlki dt n
Qi Æ j
j =1 j πi
Vi
= -Â
n
Qj Æi
j =1 j πi
Vi
Alkii + Â
Alk ij +
ONP(0.17 - 2 ¥ rain) . Vi
(A5)
Conversely, in the middle and deep parts of oceans, Alk decreases by 0.17 eq per mol decomposition of the organic carbon and increases by 2 eq per mol dissolution of CaCO3. In the middle part, the equation is,
dAlki dt n
Qi Æ j
j =1 j πi
Vi
n
Qj Æi
j =1 j πi
Vi
= -Â
+Â
dCii dt
(A3)
Alkii
( Alk
i j
)
- 0.17 ¥ C oj -
0.17 ¥ ONP(1 - dprato) , Vi
(A6)
and in the deep part, n Q n Q dAlki ONP ¥ 2 ¥ rain iÆ j j Æi = -Â . Alkii + Â Alk ij + dt V V Vi i i j =1 j =1 j πi
j πi
Ocean uptake potential for CO2 sequestration
(A7) 43
(a)
d[ PO4 ]i
2000
Atmospheric PCO2 (ppmv)
1000ppmv
dt
750ppmv 650ppmv
1500
650ppmv
Atmospheric PCO2 (ppmv)
Vi
[ PO4 ]i
Qj Æi Ê C o ˆ ONP(1 - dprato) . PO4 ]i + i ˜ [ Á Rc ¯ Rc ◊ Vi j =1 Vi Ë
1000
j πi
(A9)
500
3000
4000
5000
6000
7000
8000
dt
750ppmv 650ppmv
1500
On the other hand, in the deep part of the ocean, no reaction relates to the phosphate and its budget is expressed only based on the advection: d[ PO4 ]i
2000 1000ppmv
550ppmv 450ppmv
Qi Æ j
j =1 j πi
Vi
n
Qj Æi
j =1 j πi
Vi
[ PO4 ]i + Â
[ PO4 ]i .
(A10)
500
0 2000
2500
3000
3500
The balance equation of the phosphate is expressed based on the above DIC and alkalinity budgets by using the carbon to phosphate Redfield ratio Rc. In the surface layer, the phosphate is consumed on the new production: Qi Æ j
j =1 j πi
Vi
= -Â
n
Qj Æi
j =1 j πi
Vi
[ PO4 ]i + Â
[ PO4 ]i
ONP , Rc ◊ Vi
(A8)
where [PO4] is the concentration of the phosphate. Both the POC and DOC reached to the middle part of the ocean decompose there and regenerate the phosphate:
M. Sorai and T. Ohsumi
(A11)
dS = (1 - k ) ◊ Fplant - Fsoil . dt
(A12)
Finally, the atmospheric CO2 budget is expressed by the interactions with oceans and lands, in addition to the anthropogenic CO2 emissions: dCO2 dt 6
(
)
= - Â ( Fair -sea )i - NPP - kFplant - Fsoil + ( fossil + landuse) i =1
n
dP = TNPC, T - Fplant , dt 4000
Fig. A1. Calculated results of the future atmospheric PCO2 varying the magnitude of the oceanic thermohaline circulation: (a) the original Schmitz’s circulation, (b) the five times as fast as the case (a).
d[ PO4 ]i
n
= -Â
The terrestrial carbon reservoirs include the plant and the soil. These reservoirs are balanced by the net primary production and the decompositions of themselves:
1000
Date (A.D.)
44
j =1 j πi
+Â
Date (A.D.)
dt
Qi Æ j
n
450ppmv
0 2000
(b)
n
= -Â
(A13)
where fossil and landuse are the CO2 emissions from fossil fuel consumption and land-use change, respectively. APPENDIX II. SENSITIVITY ANALYSES As discussed in the subsection “Thermohaline circulation”, our model required to triple the magnitude of the
Schmitz’s thermohaline circulation so as to produce the realistic distribution of the DIC in the ocean. For the assessment of this modification, we performed sensitivity analyses varying the magnitude of oceanic circulation. The first run is based on the original Schmitz’s thermohaline circulation, which has an apparent residence time of 3,100 years. The second run is the five times as fast as the first one and its residence time is 620 years. These series of results including Fig. 5 certainly provide us a good implication to understand the relationship between the magnitude of ocean circulation and the oceanic CO2 uptake rate. The change in the oceanic circulation drastically affects the atmospheric PCO 2 with respect to the peak PCO 2 and the rate of approach to the steady state (Fig. A1). The faster circulation reduces the peak
PCO 2 and shortens the time required to attain to the steady state. However, it is noticed that the oceanic circulation affects little influence on the essence of the CO2 cycle. In other words, the future profile of the atmospheric PCO 2 is essentially analogous to each other. This is also supported by the fact that the ocean CO2 sequestration based on these modified circulation gave the analogous profile of the atmospheric PCO 2 . We estimated the CO2 uptake potential for the above modified cases of the thermohaline circulation as in the section “Model Implication on Ocean Sequestration”. The results showed the potentials thoroughly consistent with Fig. 7. This means that the CO2 uptake potential is independent of the magnitude of the oceanic thermohaline circulation.
Ocean uptake potential for CO2 sequestration
45
