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Present and Future of Modeling Global Environmental Change: Toward Integrated Modeling,Eds., T. Matsuno and H. Kida, pp. 195–205.© by TERRAPUB, 2001.

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 biologicalprocesses 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 beenused to estimate the oceanic uptake of anthropogenic carbon dioxide and topredict future atmospheric carbon dioxide levels (e.g., Maier-Reimer et al., 1996;Sarmiento et al., 1998). These models estimated that oceanic uptake of carbon isabout 2 GtC/yr. An important issue addressed in several studies using BGCM’sduring the 1990’s was the export of biological production as dissolved organicmatter (DOM). The accepted estimates for total export by particulate organicmatter (POM) and by DOM have changed several times. Some estimates for totalexport 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. Prognosticvariables in a typical BGCM (Yamanaka and Tajika, 1996) are concentrations ofatmospheric CO2, oceanic total CO2 (TCO2), total alkalinity (TALK), phosphate,and dissolved oxygen. Carbon dioxide gas exchange through the sea surface isassumed to be proportional to the difference of partial pressure of CO2 (pCO2)between the atmosphere and the ocean surface. The partial pressure of carbondioxide at the ocean surface is calculated from the TALK, TCO2, temperature,and salinity, assuming instantaneous chemical equilibrium during each time step.Export production from the sea surface layer is a function of phosphateconcentration in the surface water and the light factor. POM and calcite areassumed to be remineralized with the observed vertical profiles instantaneouslybelow the euphotic zone. Sedimentation processes and river input due to thecontinental weathering are not included.

Yamanaka and Tajika (1996, 1997) successfully reproduced observed tracerdistributions on the global scale. Figure 2 shows the meridional distributions ofphosphate along the Geochemical Ocean Sections Study (GEOSECS) Western

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Pacific Sector. There is a vertical maximum of phosphate at 1000 m depth, locatedin the northern North Pacific. Bacastow and Maier-Reimer (1991) suggested thateffects of DOM are important in determining phosphate distribution, becausephosphate distributions, simulated by models including DOM, are closer to theobservations than those from models without DOM (Fig. 2(c)). They estimatedexport productions by DOM to be 8 GtC/yr. However, their simulations of DOMwere 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 andTajika (1997), with a model including a representation of DOM based on morerecent observations (e.g., Peltzer and Hayward, 1996), reproduced well theobserved phosphate distribution (Fig. 2(d)). In their simulations, they also foundthat the effect of DOM on the simulated phosphate distribution was minorcompared to that of POM.

Figure 3 shows the distribution of ∆pCO2 which represents the difference inpCO2 between the atmosphere and the ocean. Figures 3(a) and (b) are for theobservation by Tans et al. (1990), and Fig. 3(c) is for those by Yamanaka andTajika (1996). The model successfully reproduced the following general featuresof the measured values. In the equatorial Pacific, pCO2 in the ocean is >100 µatmhigher than that in the atmosphere due to the upwelling of deep water with highpCO2. In the subtropical regions, the oceanic pCO2 is lower than the atmosphericpCO2 because of the biological pump. Figure 3(d) shows the ∆pCO2 differencebetween 1760 and 1985 (oceanic uptake of anthropogenic CO2 occurs in thepositive regions), which is regared as the anthropogenic component of ∆pCO2.

Fig. 1. Chemical and biological processes considered in the biogeochemical general circulationmodels (BGCMs).

Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 197

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Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 199

The difference of anthropogenic component of ∆pCO2 is about ten times smallerthan “natural” ∆pCO2 (Note C.I. = 20 µatm in Figs. 3(a)–(c), C.I. = 2 µatm inFig. 3(d)). The globally averaged anthropogenic ∆pCO2 is 8 µatm, whichcorresponds to a calculated oceanic uptake of 2 GtC/yr. In the equatorial region,although natural pCO2 in the ocean is much higher than that in the atmosphere,anthropogenic CO2 is taken up by the ocean because anthropogenic pCO2 in theocean is lower than that in the atmosphere.

The Ocean Carbon-Cycle Modeling Intercomparison Project 2 (OCMIP2)was begun in 1998. Thirteen groups (Europe 7 groups, USA 4 groups, Japan 1group, and Australia 1 group) participated in OCMIP2, and conducted fourexperiments as follows:

(1) CFC Experiments: comparing effects of the different circulation fieldsin the different models on the simulated distribution of CFC’s in the ocean, usingcommon CFC gas exchange processes. Completed in 1998.

(2) Abiotic Experiments: Advection and diffusion of Total CO2 andAlkalinity as a result of oceanic circulation and air-sea gas exchange only.Completed in 1999.

(3) Biotic Experiments: Abiotic experiments plus simple biologicalprocesses based on the restoring to observed phosphate concentrations at the seasurface. Ongoing, 2000.

(4) Injection Experiments: CO2 disposal at depths of 800, 1500, 3000 mnear 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 anthropogenicCO2 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 accuratepredictions of future atmospheric CO2 levels and global warming. In previousBGCMs, export production is a function of phosphate concentration in thesurface water, and is not based on the dynamics of marine ecosystems. For theirpredictions of future atmospheric CO2 levels, those models usually assume thatmarine ecosystems do not change as a result of global warming. However,changes in ocean temperature and circulation may change the functioning ofoceanic ecosystems, and such changes can affect the predicted uptake ofanthropogenic CO2 (Siegenthaler and Sarmiento, 1993). To illustrate the difficultyof such predictions, we will discuss the value of the biological efficiencyparameter, r, which has a large uncertainty. Because the biological new productionis proportional to the phosphate concentration in the euphotic layer, with coefficientof 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 atime to be explained directly by the typical time scale of a marine ecosystem. One

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reasonable explanation is that r is the effective throughput time of phosphate inthe euphotic layer resulting from the complex biological activity (nutrientrecycling) in the euphotic layer. Although we obtained this empirical value of rusing the BGCM to simulate the present ocean, we cannot predict how muchglobal warming might change r. We cannot directly observe the value of r, butonly micro-scale values in the ecosystem at any given local station (e.g., stocksizes of phytoplankton, etc.). Looking toward the next generation models, wemust use not only the BGCM but also the ecosystem model, which representsexplicitly the phytoplankton and recycling of nutrients. In such a model, r isdetermined by the ecosystem model. This next generation model, coupling amarine ecosystem and a biogeochemical model, has the potential to simulate theeffect 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 theecosystem model extended in Kawamiya et al. (1995, 1997) and Kishi et al.(2001) (hereafter KM) is coupled with the biogeochemical model in Yamanakaand Tajika (1997). Figure 4 illustrates the interactions among the fifteencompartments in the model of biological processes. We divide phytoplankton andzooplankton 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 verticalmigration (ascending to shallower depths in the spring, maturing while grazingan other plankton at shallow depth, and returning to deeper waters in the fall). Thetreatment of ZL is the same as in KM. ZS represents the others. PL representsdiatoms that make siliceous shells. Therefore, the rate of photosynthesis by PL islimited by both nitrogen and silicate. PS represents the other phytoplankton (non-diatoms and flagellates). Some of PS and ZS are regarded as cocolithophorids andforaminifera, respectively, which have calcareous shells. Predatory zooplankton(ZP) represents zooplankton such as euphausiids that graze an other plankton. ZPis expected to be important for linking the lower trophic levels of this model tohigher trophic levels such as fish (as discussed at the PICES modeling workshopin Nemuro Japan, January, 2000 (Eslinger et al., 2000)). The model includes threenutrients and three kinds of settling particles: nitrate (NO3), ammonium (NH4),and silicate (Si(OH)4), particulate organic matter (POM), opal, and calciumcarbonate (CaCO3). Dissolved organic matter (DOM) is also included in themodel. A mass balance is included for calcium (Ca) as well, so that total alkalinity(TALK) may be calculated using the concentrations of nitrate and calciumfollowing the TALK definition. Total carbon dioxide (TCO2) is calculatedassuming that all phytoplankton have a Redfield C/N ratio (106/16). The cyclingof silicate and nitrogen affect the ecosystem dynamics through nutrient limitation.The carbon and calcium cycles do not affect the ecosystem dynamics, becauseboth carbon and calcium are always plentiful in the ocean and do not limitproduction. The ecological model described above is coupled with a vertical one-

Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 201

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dimensional physical model which has the same mixed layer closure scheme usedin KM. Our model has 28 vertical layers spanning a model domain from thesurface 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 stressand solar radiation from 1991 to 1996 including realistic fluctuations fromweather events can be explicitly included in the boundary conditions. A spin upintegration was performed by repeating the 1991 forcing for ten years, after whichthe actual forcing for 1991 through 1996 was applied for the simulationspresented here.

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 cm2/sec) (Yamanaka et al.,2001). All quantities are plotted as daily averages.

Development of Coupled Ocean Physical-Biogeochemical-Ecosystem Model 203

Figure 5 shows vertical distributions of the vertical diffusive coefficient,nitrate concentration, and chlorophyll-a concentration from 1991 through 1996.The areas of high vertical diffusive coefficient (shaded area in Fig. 5(a)) representthe mixed layer. The simulation reproduces seasonal changes in the mixed layerdepth and interannual variations in deep convection. Convection in winterpenetrates deeper than 160 m in 1992 and 1996. Nitrate-rich waters (about 20µmolN/l) are supplied to the upper ocean by this deep convection, resulting in thestrong blooms of diatoms observed (and simulated) after winter. In 1992, themaximum simulated chlorophyll-a concentration reaches 4.6 µgChla/l, whichcompares 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 70m depths, and the simulated spring bloom is quite weak. In summer, whennutrients in the near surface waters are depleted, the chlorophyll-a maximumconcentration is located around the 30 m depths. The model successfully reproducesthe observed seasonal variations of nitrate, silicate, and chlorophyll-a from 1991through 1996, despite the inability of this model to represent the effects ofhorizontal advection/diffusion due to meso-scale eddy on nutrient and chlorophyll-a concentrations. Figure 6 shows primary production by PS and PL, and thepartial pressure of CO2. The green line, primary production by diatoms, has itshighest peak in spring and its second highest peak in fall. Simulated primaryproduction during the spring bloom ranges from 800 through 600 mgC/m2day in1992, 1995, and 1996, which compares well with observations. As found in KM,the spring bloom of diatoms has large interannual variations: the spring bloom isstrong in 1992, 1995, and 1996, weak in 1991 and 1994, and intermediate in 1993.The spring bloom of diatoms ceases because of increasing grazing pressure bycopepods, although both nitrate and silicate in the surface water still remain over10 µmol/l, (larger than the half-saturation constants for their uptake) at the endof the bloom. After the diatom bloom, primary production by PS increases, in theoften-observed transition from a diatom-dominated bloom to a flagellate-

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 areplotted as daily averages.

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dominated bloom that exhausts the nitrate in the surface waters by late summer.Many frequent downward spikes appear in the line representing primary productionin Fig. 6. These represent decreases in primary production resulting fromdecreases in solar radiation during rainy or cloudy days. The partial pressure ofcarbon dioxide has its maximum associated with winter convection, rapidlydecreases during the spring bloom, and remains almost constant from summerthrough fall, during which time the effect of increasing temperature (increasingpCO2) cancels the effect of the biological pump (depressing pCO2).

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Kawamiya, M., M. Kishi, Y. Yamanaka, and N. Suginohara (1997) Procuring reasonable results indifferent oceanic regimes with the same ecological-physical coupled model, J. Oceanogr., 53,397–402.

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Peltzer, E. T. and N. A. Hayward (1996) Spatial and temporal variability of total organic carbonalong 140W in the equatorial Pacific Ocean in 1992, Deep-Sea Res. II, 43, 1155–1180.

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Sarmiento, J. L., T. M. C. Hughes, R. J. Stouffer, and S. Manabe (1998) Simulated response of theocean carbon cycle to anthropogenic climate warming, Nature, 393, 245–249.

Siegenthaler, U. and J. L. Sarmiento (1993) Atmospheric carbon dioxide and the ocean, Nature, 365,119–125.

Sugimura, Y. and Y. Suzuki (1988) A high-temperature catalytic oxidation methods for determinationof nonvolatile dissolved organic carbon in seawater by direct injection of a liquid sample, Mar.Chem., 24, 105–131.

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CO2 budget, Science, 247, 1431–1438.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 circulationmodel, Global Biogeochem. Cycles, 10, 361–382.

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Yamanaka, Y. and E. Tajika (1997) Role of dissolved organic matter in the marine biogeochemicalcycle: Studies using an ocean biogeochemical general circulation model, Global Biogeochem.Cycles, 11, 599–613.

Yamanaka, Y., N. Yoshie, M. Fujii, M. Aita-Noguti, and M. J. Kishi (2001) An ecosystem modelcoupled with Nitrogen-Silicon-Carbon cycles applied to Station A-7 in the NorthwesternPacific, J. Oceanogr. (submitted).

Y. Yamanaka (e-mail: galapen@ees.hokudai.ac.jp)