**2. Materials and methods**

response of upper ocean water to typhoons can be conventionally divided into two stages, the forcing stage and the relaxation stage in [1, 2]. In the first stage, potential energy is injected into the surface ocean by strong typhoon winds. Two significant physical phenomena caused by typhoons are the mixed layer deepening and sea surface temperature cooling in the wake of the storm. Warm water in the ocean surface layers is transported outward from the typhoon center and downward to depths ranging from tens of meters to beyond a hundred meters; cold water upwells from the deeper ocean along the typhoon's passage [3]. The current velocity in the surface mixed layer can reach 2 m s−1 or more, responding to the intensity of the tropical storm winds [2]. The relaxation stage following a typhoon's passage is primarily due to the inertial gravity oscillations excited by the storm in its wake, where the ocean adjusts towards a new geostrophic equilibrium state [1, 4–6]. The work of [7] concluded that the inertial oscillations are predominantly locally generated and the surface winds account for a large part of the energy and variability of such oscillations near the ocean surface. The theory of geostrophic adjustment was further reviewed in [8]. The latter theory, which describes the nature of the difference between the baroclinic and barotropic responses of the ocean to a moving storm as caused by the difference in the gravity wave speed, was first pointed out in [9]. The storm‐induced oscillating wake is formed by the slow propagating, near‐inertial gravity baroclinic waves, while the fast propagating barotropic waves produce a broad area of convergent depth‐averaged currents with no discernible wake; the latter is determined entirely by the wind stress curl, with negligible effects due to the earth's rotation and ocean stratification [10]. The work of [11] confirmed that the mixed layer dynamics is associated with shear‐induced entrainment mixing, and forced by near‐inertial motions up to the third day

68 Recent Developments in Tropical Cyclone Dynamics, Prediction, and Detection

Tropical storms also exert a strong influence on the oceanic chlorophyll *a* field and primary production in the ocean. Another important phenomenon triggered and enhanced by tropical cyclones is the phytoplankton bloom accompanied by nutrient pumping into the oligotrophic surface layer. Concentration of nitrate, phosphate, and chlorophyll *a* is observed to significantly increase after the occurrence of cyclonic disturbances [12]. In [13], after Typhoon Fengwong and Typhoon Sinlaku passed over the southern East China Sea in 2008, the in‐situ particulate organic carbon flux was observed to experience a significant increase of about 1.7‐fold and 1.5‐ fold, respectively, compared to the recorded (140–180 mg cm−2 d−1) pre‐typhoon concentration. The phytoplankton population growth was constrained by the light limitation and the grazing pressure. This increase of the surface chlorophyll *a* concentration might last 2–3 weeks before relaxing to pre‐typhoon levels [14]. Because of the limitations imposed by in‐situ point observations from ships or moored buoys along a typhoon's track, studies of the associated biological responses have become to more and more depend on the satellite observations. In the work of [15], satellite data are used as new evidence to quantify the contribution of tropical cyclones to enhance the ocean primary production. It was found that the peak of the chloro‐ phyll *a* concentration enhancement tended to occur several days after the sea surface temper‐ ature cooling had achieved the maximum amplitude, after the typhoon's passage, see [16, 17]. The extended region and concentration of primary production bloom tend to vary in response to the translation speed and intensity of typhoon. Weak slow‐moving typhoons can cause enhanced concentrations of chlorophyll *a*, while strong fast‐moving typhoon tends to cause

after the passage of the storm.

ROMS is utilized to simulate the processes of Typhoon Cimaron, which influenced the South China Sea from October 29 to November 7, 2006. The model region covers the whole South China Sea including 0°–30°N, 99°–130°E. There are 220 × 100 orthogonal curvilinear grids with the horizontal resolution varying at Δ*x* (5.5–40 km) and Δ*y* (3.6–37 km) and the minimum and maximum depths are 5 m and 5000 m, respectively (**Figure 1**). There are 80 layers in the vertical direction using the *s*‐coordinate formulation. The western boundary is closed and the other three open boundaries are defined by the radiation boundary condition.

**Figure 1.** Model domain and grids, as well as bathymetry.

The initial temperature and salinity conditions are taken from the 1/4° grid climatological temperature and salinity analyses of October from WOA01 to represent the pre‐typhoon conditions in the South China Sea, which was discussed by Carton and Giese [31]. The climatological monthly data have 24 standard levels with depths varying from 0 to 1500 m and the seasonal data have 33 standard levels with depths from 0 to 5500 m. As the maximum depth in the model is 5000 m, the climatological monthly data are applied in the upper 1500 m and the climatological seasonal data (autumn) are applied from 1500 to 5000 m. The initial current velocity is set to zero in this study.

The lateral boundary conditions for temperature, salinity, sea level and current velocities are obtained from the 5‐day averages from the global simulations of the Simple Ocean Data Assimilation (SODA) dataset with horizontal resolution of 0.5° × 0.5° and 40 vertical layers [31]. The Kuroshio Current transport can be identified on the eastern boundary. The tidal ampli‐ tudes and phases used in this model are obtained from the TPXO Global Inverse solution database [32] with eight primary tide constituents (M2, S2, N2, K2, K1, O1, P1, and Q1) and two long‐period constituents (Mf and Mm).

The daily wind stresses are obtained from QuikSCAT satellite data. The effect of heat flux is considered as the surface boundary for momentum, although the heat flux can be neglected under the extreme meteorological phenomena such as tropical cyclones. The daily heat fluxes are obtained from the Objectively Analyzed air‐sea Fluxes (OAFlux) project which is an ongoing research and development project for global air‐sea fluxes [33]. To present the climatological heating or cooling trends, we computed the surface boundary conditions for temperature and salinity involving relaxation to its observed values. Corrections are made for the net surface heat flux in the model simulations, as discussed in [34].

For the biological model, the surface chlorophyll *a* field data are estimated from the SeaWiFS climatological seasonal data, and the nitrate (NO3) and oxygen are estimated from the monthly climatological database of WOA09 [35, 36]. The vertical structure of chlorophyll *a* is extrapo‐ lated from the surface chlorophyll *a* field using the Morel and Berthon [37] parameterization method. High concentrations of chlorophyll *a* are mostly distributed in coastal regions along the coastline of southern China and the Vietnam marginal continent, with values of over 1 mg cm−3. By comparison, in the deep sea area, the concentration is less than 0.1 mg cm−3. In most areas of the South China Sea, the NO3 content is much smaller, less than 0.1 mmol Nm−3 in the deep ocean surface layer. The concentration of NO3 increases with ocean depth. In the deep sea, with depths in excess of 800 m, the concentration of NO3 can reach 40 mmol Nm−3.

Both the total inorganic carbon (TIC) and total alkalinity are obtained from the Carbon Dioxide Information Analysis Center. The detailed data information is given in work et al. [38]. The climatological seasonal dataset is used to generate the initial and boundary conditions. Ammonium (NH4), large and small detritus, and N‐concentration are taken from the NO3 estimates and multiplied by the respective ratios of 18/62, 0.1 and 0.1, respectively, while phytoplankton, zooplankton, large and small detritus, and C‐concentration are taken from the chlorophyll *a* concentration multiplied by the respective ratios of 0.5, 0.2, 0.1, and 0.1, respectively.
