Development of Coupled Ocean Physical-Biogeochemical- Ecosystem Model

Transcription

1 Present and Future of Modeling Global Environmental Change: Toward Integrated Modeling, Eds., T. Matsuno and H. Kida, pp by TERRAPUB, Development of Coupled Ocean Physical-Biogeochemical- Ecosystem Model Yasuhiro YAMANAKA Graduate School of Environmental Earth Science, Hokkaido University, Japan INTRODUCTION TO BIOGEOCHEMICAL GENERAL CIRCULATION MODELS Biogeochemical general circulation models (BGCMs) with simplified biological processes were developed in the 1990s (e.g., Bacastow and Maier-Reimer, 1990; Najjar et al., 1992; Yamanaka and Tajika, 1996, 1997). These BGCM s have been used to estimate the oceanic uptake of anthropogenic carbon dioxide and to predict future atmospheric carbon dioxide levels (e.g., Maier-Reimer et al., 1996; Sarmiento et al., 1998). These models estimated that oceanic uptake of carbon is about 2 GtC/yr. An important issue addressed in several studies using BGCM s during the 1990 s was the export of biological production as dissolved organic matter (DOM). The accepted estimates for total export by particulate organic matter (POM) and by DOM have changed several times. Some estimates for total export by POM and DOM, respectively, are: 4 GtC/yr and 0 GtC/yr (IPCC, 1990), 2 GtC/yr and 8 GtC/yr (Sarmiento and Siegenthaler, 1992), 4 GtC/yr and 6 GtC/ yr (IPCC, 1995), and 8 GtC/yr and 2 GtC/yr (Yamanaka and Tajika, 1997). These models share common features: low horizontal resolution (about 4 4 degrees), annual mean forcing, classical horizontal/vertical mixing, and so on. Figure 1 illustrates typical biogeochemical processes in BGCMs. Prognostic variables in a typical BGCM (Yamanaka and Tajika, 1996) are concentrations of atmospheric CO 2, oceanic total CO 2 (TCO 2 ), total alkalinity (TALK), phosphate, and dissolved oxygen. Carbon dioxide gas exchange through the sea surface is assumed to be proportional to the difference of partial pressure of CO 2 (pco 2 ) between the atmosphere and the ocean surface. The partial pressure of carbon dioxide at the ocean surface is calculated from the TALK, TCO 2, temperature, and salinity, assuming instantaneous chemical equilibrium during each time step. Export production from the sea surface layer is a function of phosphate concentration in the surface water and the light factor. POM and calcite are assumed to be remineralized with the observed vertical profiles instantaneously below the euphotic zone. Sedimentation processes and river input due to the continental weathering are not included. Yamanaka and Tajika (1996, 1997) successfully reproduced observed tracer distributions on the global scale. Figure 2 shows the meridional distributions of phosphate along the Geochemical Ocean Sections Study (GEOSECS) Western 195

2 196 Y. YAMANAKA Fig. 1. Chemical and biological processes considered in the biogeochemical general circulation models (BGCMs). Pacific Sector. There is a vertical maximum of phosphate at 1000 m depth, located in the northern North Pacific. Bacastow and Maier-Reimer (1991) suggested that effects of DOM are important in determining phosphate distribution, because phosphate distributions, simulated by models including DOM, are closer to the observations than those from models without DOM (Fig. 2(c)). They estimated export productions by DOM to be 8 GtC/yr. However, their simulations of DOM were based on the high DOM concentrations measured by Sugimura and Suzuki (1988), which were later withdrawn by Suzuki (1993). Yamanaka and Tajika (1996) reproduced the observed phosphate distribution fairly well (Fig. 2(c)), even though their model did not include a representation of DOM. Yamanaka and Tajika (1997), with a model including a representation of DOM based on more recent observations (e.g., Peltzer and Hayward, 1996), reproduced well the observed phosphate distribution (Fig. 2(d)). In their simulations, they also found that the effect of DOM on the simulated phosphate distribution was minor compared to that of POM. Figure 3 shows the distribution of pco 2 which represents the difference in pco 2 between the atmosphere and the ocean. Figures 3(a) and (b) are for the observation by Tans et al. (1990), and Fig. 3(c) is for those by Yamanaka and Tajika (1996). The model successfully reproduced the following general features of the measured values. In the equatorial Pacific, pco 2 in the ocean is >100 µatm higher than that in the atmosphere due to the upwelling of deep water with high pco 2. In the subtropical regions, the oceanic pco 2 is lower than the atmospheric pco 2 because of the biological pump. Figure 3(d) shows the pco 2 difference between 1760 and 1985 (oceanic uptake of anthropogenic CO 2 occurs in the positive regions), which is regared as the anthropogenic component of pco 2.

3 Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 197 Fig. 2. Phosphate distributions along the Geochemical Ocean Sections Study (GEOSECS) sections in the western Pacific. (a) observations, (b) (e) model results. (b) and (c) are after Bacastow and Maier-Reimer (1991), (d) is after Yamanaka and Tajika (1996), (d) is after Yamanaka and Tajika (1997). Contour interval is 0.2 µmol/kg.

4 198 Y. YAMANAKA Fig. 3. Horizontal distribution of pco 2 from (a) observation (JAN APR), (b) observation (JUN OCT), (c) model result in 1985, and (d) difference of pco 2 between 1760 and Observations are after Tans et al. (1990). Contour interval is 20 µatm in (a) (c) and 2 µatm in (d).

5 Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 199 The difference of anthropogenic component of pco 2 is about ten times smaller than natural pco 2 (Note C.I. = 20 µatm in Figs. 3(a) (c), C.I. = 2 µatm in Fig. 3(d)). The globally averaged anthropogenic pco 2 is 8 µatm, which corresponds to a calculated oceanic uptake of 2 GtC/yr. In the equatorial region, although natural pco 2 in the ocean is much higher than that in the atmosphere, anthropogenic CO 2 is taken up by the ocean because anthropogenic pco 2 in the ocean is lower than that in the atmosphere. The Ocean Carbon-Cycle Modeling Intercomparison Project 2 (OCMIP2) was begun in Thirteen groups (Europe 7 groups, USA 4 groups, Japan 1 group, and Australia 1 group) participated in OCMIP2, and conducted four experiments as follows: (1) CFC Experiments: comparing effects of the different circulation fields in the different models on the simulated distribution of CFC s in the ocean, using common CFC gas exchange processes. Completed in (2) Abiotic Experiments: Advection and diffusion of Total CO 2 and Alkalinity as a result of oceanic circulation and air-sea gas exchange only. Completed in (3) Biotic Experiments: Abiotic experiments plus simple biological processes based on the restoring to observed phosphate concentrations at the sea surface. Ongoing, (4) Injection Experiments: CO 2 disposal at depths of 800, 1500, 3000 m near seven major cities. Completed in 2000 (European groups only). Two experiments, (2) and (3), consist of several time-series runs: Historical run, IS92a run, S650 run, Pulse run. The estimates of oceanic uptake of anthropogenic CO 2 from these experiments will be included in the next IPCC reports. TOWARD THE NEXT GENERATION OF BGCMS The previous studies using BGCMs, especially Yamanaka and Tajika (1996, 1997), successfully reproduced observed tracer distributions on the global scale. However, this success simulating the present does not guarantee accurate predictions of future atmospheric CO 2 levels and global warming. In previous BGCMs, export production is a function of phosphate concentration in the surface water, and is not based on the dynamics of marine ecosystems. For their predictions of future atmospheric CO 2 levels, those models usually assume that marine ecosystems do not change as a result of global warming. However, changes in ocean temperature and circulation may change the functioning of oceanic ecosystems, and such changes can affect the predicted uptake of anthropogenic CO 2 (Siegenthaler and Sarmiento, 1993). To illustrate the difficulty of such predictions, we will discuss the value of the biological efficiency parameter, r, which has a large uncertainty. Because the biological new production is proportional to the phosphate concentration in the euphotic layer, with coefficient of proportionality r, this is an essential parameter in the marine biological cycle. Yamanaka and Tajika (1996) estimated r as about one year, which is too long a time to be explained directly by the typical time scale of a marine ecosystem. One

6 200 Y. YAMANAKA reasonable explanation is that r is the effective throughput time of phosphate in the euphotic layer resulting from the complex biological activity (nutrient recycling) in the euphotic layer. Although we obtained this empirical value of r using the BGCM to simulate the present ocean, we cannot predict how much global warming might change r. We cannot directly observe the value of r, but only micro-scale values in the ecosystem at any given local station (e.g., stock sizes of phytoplankton, etc.). Looking toward the next generation models, we must use not only the BGCM but also the ecosystem model, which represents explicitly the phytoplankton and recycling of nutrients. In such a model, r is determined by the ecosystem model. This next generation model, coupling a marine ecosystem and a biogeochemical model, has the potential to simulate the effect of climate change on the marine biogeochemical cycle. OUR DEVELOPING OF AN ECOSYSTEM MODEL Now we are developing a new model (Yamanaka et al., 2001) in which the ecosystem model extended in Kawamiya et al. (1995, 1997) and Kishi et al. (2001) (hereafter KM) is coupled with the biogeochemical model in Yamanaka and Tajika (1997). Figure 4 illustrates the interactions among the fifteen compartments in the model of biological processes. We divide phytoplankton and zooplankton into two and three categories, respectively: large phytoplankton (PL), small phytoplankton (PS), large zooplankton (ZL), small zooplankton (ZS) and predatory zooplankton (ZP). ZL represents copepods with seasonally vertical migration (ascending to shallower depths in the spring, maturing while grazing an other plankton at shallow depth, and returning to deeper waters in the fall). The treatment of ZL is the same as in KM. ZS represents the others. PL represents diatoms that make siliceous shells. Therefore, the rate of photosynthesis by PL is limited by both nitrogen and silicate. PS represents the other phytoplankton (nondiatoms and flagellates). Some of PS and ZS are regarded as cocolithophorids and foraminifera, respectively, which have calcareous shells. Predatory zooplankton (ZP) represents zooplankton such as euphausiids that graze an other plankton. ZP is expected to be important for linking the lower trophic levels of this model to higher trophic levels such as fish (as discussed at the PICES modeling workshop in Nemuro Japan, January, 2000 (Eslinger et al., 2000)). The model includes three nutrients and three kinds of settling particles: nitrate (NO 3 ), ammonium (NH 4 ), and silicate (Si(OH) 4 ), particulate organic matter (POM), opal, and calcium carbonate (CaCO 3 ). Dissolved organic matter (DOM) is also included in the model. A mass balance is included for calcium (Ca) as well, so that total alkalinity (TALK) may be calculated using the concentrations of nitrate and calcium following the TALK definition. Total carbon dioxide (TCO 2 ) is calculated assuming that all phytoplankton have a Redfield C/N ratio (106/16). The cycling of silicate and nitrogen affect the ecosystem dynamics through nutrient limitation. The carbon and calcium cycles do not affect the ecosystem dynamics, because both carbon and calcium are always plentiful in the ocean and do not limit production. The ecological model described above is coupled with a vertical one-

7 Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 201 Fig. 4. Schematic view of interactions among the fifteen model compartments (Yamanaka et al., 2001).

8 202 Y. YAMANAKA Fig. 5. Vertical distributions from 1991 through 1996: (a) logarithm of vertical diffusive coefficient (C.I. = 1.0), (b) nitrate concentration (C.I. = 2.0 µmol/l), and (c) chlorophyll-a concentration (C.I. = 0.5 µg/l). Shaded area in (a) represents 3.0$ (i.e., 1000 cm 2 /sec) (Yamanaka et al., 2001). All quantities are plotted as daily averages. dimensional physical model which has the same mixed layer closure scheme used in KM. Our model has 28 vertical layers spanning a model domain from the surface to 330 m depth. The space step is 5 m for all layers in the upper 100 m, and increases to 60 m for the deepest layers. Boundary conditions for SST, SSS, wind stress and solar radiation are the same as those of KM. Hourly wind stress and solar radiation from 1991 to 1996 including realistic fluctuations from weather events can be explicitly included in the boundary conditions. A spin up integration was performed by repeating the 1991 forcing for ten years, after which the actual forcing for 1991 through 1996 was applied for the simulations presented here.

9 Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 203 Fig. 6. Time series from 1991 to 1996: primary productions by PL (thick line) and PS (thin line), and partial pressure of carbon dioxide (dotted line) (Yamanaka et al., 2001). All quantities are plotted as daily averages. Figure 5 shows vertical distributions of the vertical diffusive coefficient, nitrate concentration, and chlorophyll-a concentration from 1991 through The areas of high vertical diffusive coefficient (shaded area in Fig. 5(a)) represent the mixed layer. The simulation reproduces seasonal changes in the mixed layer depth and interannual variations in deep convection. Convection in winter penetrates deeper than 160 m in 1992 and Nitrate-rich waters (about 20 µmoln/l) are supplied to the upper ocean by this deep convection, resulting in the strong blooms of diatoms observed (and simulated) after winter. In 1992, the maximum simulated chlorophyll-a concentration reaches 4.6 µgchla/l, which compares well with what is often observed during the spring bloom at Station A- 7. On the other hand, in 1993 and 1994, convection only penetrates to around 70 m depths, and the simulated spring bloom is quite weak. In summer, when nutrients in the near surface waters are depleted, the chlorophyll-a maximum concentration is located around the 30 m depths. The model successfully reproduces the observed seasonal variations of nitrate, silicate, and chlorophyll-a from 1991 through 1996, despite the inability of this model to represent the effects of horizontal advection/diffusion due to meso-scale eddy on nutrient and chlorophylla concentrations. Figure 6 shows primary production by PS and PL, and the partial pressure of CO 2. The green line, primary production by diatoms, has its highest peak in spring and its second highest peak in fall. Simulated primary production during the spring bloom ranges from 800 through 600 mgc/m 2 day in 1992, 1995, and 1996, which compares well with observations. As found in KM, the spring bloom of diatoms has large interannual variations: the spring bloom is strong in 1992, 1995, and 1996, weak in 1991 and 1994, and intermediate in The spring bloom of diatoms ceases because of increasing grazing pressure by copepods, although both nitrate and silicate in the surface water still remain over 10 µmol/l, (larger than the half-saturation constants for their uptake) at the end of the bloom. After the diatom bloom, primary production by PS increases, in the often-observed transition from a diatom-dominated bloom to a flagellate-

10 204 Y. YAMANAKA dominated bloom that exhausts the nitrate in the surface waters by late summer. Many frequent downward spikes appear in the line representing primary production in Fig. 6. These represent decreases in primary production resulting from decreases in solar radiation during rainy or cloudy days. The partial pressure of carbon dioxide has its maximum associated with winter convection, rapidly decreases during the spring bloom, and remains almost constant from summer through fall, during which time the effect of increasing temperature (increasing pco 2 ) cancels the effect of the biological pump (depressing pco 2 ). REFERENCES Bacastow, R. and E. Maier-Reimer (1990) Ocean-circulation model of the carbon cycle, Clim. Dyn., 4, Eslinger, D. V., M. B. Kashiwai, M. J. Kishi, B. A. Megrey, D. M. Ware, and F. E. Werner (2000) Report of the 2000 MODEL Workshop on lower trophic level modeling, PICES Scientific Report, 15, IPCC (1990) Climate Change: The IPCC Scientific Assessment, edited by J. T. Houghton, G. J. Jenkins, and J. J. Ephraums, Cambridge Univ. Press, 365 pp. IPCC (1995) Climate Change 1995: The Science of Climate Change, edited by J. T. Houghton, L. G. Meira Filho, B. A. Callander, N. Harris, A. Kattenberg, and K. Maskell, Cambridge Univ. Press, 572 pp. Kawamiya, M., M. Kishi, Y. Yamanaka, and N. Suginohara (1995) An ecological-physical coupled model applied to Station Papa, J. Oceanogr., 51, Kawamiya, M., M. Kishi, Y. Yamanaka, and N. Suginohara (1997) Procuring reasonable results in different oceanic regimes with the same ecological-physical coupled model, J. Oceanogr., 53, Kishi, M. J., H. Motono, M. Kashiwai, and A. Tsuda (2001) An ecological-physical coupled model with ontogenentic vertical migration of zooplankton in the northwestern Pacific, J. Oceanogr. (in press). Maier-Reimer, E., U. Mikolajewics, and A. Winguth (1996) Future ocean uptake of CO 2 : interaction between oceancirculation and biology, Clim. Dyn., 12, Najjar, R. G., J. L. Sarmiento, and J. R. Toggweiler (1992) Downward transport and fate of organic matter in the ocean: Simulations with a general circulation model, Global Biogeochem. Cycles, 6, Peltzer, E. T. and N. A. Hayward (1996) Spatial and temporal variability of total organic carbon along 140W in the equatorial Pacific Ocean in 1992, Deep-Sea Res. II, 43, Sarmiento, J. L. and U. Siegenthaler (1992) New production and the global carbon cycle, in Primary Productivity and Biogeochemical Cycles in the Sea, edited by P. G. Falkowski and A. D. Woodhead, Plenum Press, pp Sarmiento, J. L., T. M. C. Hughes, R. J. Stouffer, and S. Manabe (1998) Simulated response of the ocean carbon cycle to anthropogenic climate warming, Nature, 393, Siegenthaler, U. and J. L. Sarmiento (1993) Atmospheric carbon dioxide and the ocean, Nature, 365, Sugimura, Y. and Y. Suzuki (1988) A high-temperature catalytic oxidation methods for determination of nonvolatile dissolved organic carbon in seawater by direct injection of a liquid sample, Mar. Chem., 24, Suzuki, Y. (1993) On the measurement of DOC and DON in seawater, Mar. Chem., 41, Tans, P. P., I. Y. Fung, and T. Takahashi (1990) Observational constraints on the global atmospheric CO 2 budget, Science, 247, Yamanaka, Y. and E. Tajika (1996) The role of the vertical fluxes of particulate organic matter and calcite in the oceanic carbon cycle: Studies using an ocean biogeochemical general circulation model, Global Biogeochem. Cycles, 10,

11 Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 205 Yamanaka, Y. and E. Tajika (1997) Role of dissolved organic matter in the marine biogeochemical cycle: Studies using an ocean biogeochemical general circulation model, Global Biogeochem. Cycles, 11, Yamanaka, Y., N. Yoshie, M. Fujii, M. Aita-Noguti, and M. J. Kishi (2001) An ecosystem model coupled with Nitrogen-Silicon-Carbon cycles applied to Station A-7 in the Northwestern Pacific, J. Oceanogr. (submitted). Y. Yamanaka (