**4. Influence on mesoscale eddies**

A mesoscale eddy is one vortex with its core surrounded by closed circulations. In the ocean, it has a time scale of 10–100 days and a dimension of 10–100 km in horizontal and 100–1000 m in vertical. It can travel a distance of 100–1000 km during its lifetime of several months. According to the rotation direction of closed circulations, mesoscale eddies are classified into cyclonic (counterclockwise) and anticyclonic (clockwise) eddies. Because the temperature inside a cyclonic eddy is usually lower than that of surrounding water, the cyclonic eddy is also called cold eddy. Inversely, an anticyclonic eddy is also called warm eddy. Sea level anomaly with respect to local mean sea level is negative in a cyclonic eddy whereas it is positive in an anticyclonic eddy.

Mesoscale eddies are ubiquitous in the open ocean and marginal seas (e.g., the SCS) of the WNP. There are two eddy‐rich regions in the open ocean of the WNP [47–49]: the Kuroshio Extension region (30°–40°N, 140°E–180°W) and the North Pacific Subtropical Countercurrent region (18°–25°N, 122°E–160°E). The SCS is a well‐known eddy‐rich marginal sea with a large area of 3,500,000 km2 and an average depth of over 2000 m [50, 51]. In these regions, typhoons appear frequently, which may influence the structure and evolution of some eddies.

Pre‐existing cyclonic eddies can be intensified after a typhoon passes over. Zheng et al. [52, 53] revealed that the SST in pre‐existing cyclonic eddies markedly decreased after the passage of Typhoon Hai‐Tang (2005) because cold deep water was easily brought up due to uplifted thermocline, large current shear just below the mixed layer, and small thermal inertia within the cyclonic eddies. Chiang et al. [8] showed that when Typhoon Kai‐Tak (2000) passed over the West Luzon Eddy (a cyclonic eddy), it produced an unusually intense SST drop of about 10.8°C and high chlorophyll *a* (Chl‐*a*) concentration in the eddy. Lai et al. [54] found that a cold eddy was intensified when Typhoon Morakot (2009) passed by the eddy, resulting in a significant cooling. Nam et al. [55] pointed out that during Typhoon Man‐Yi (2007), more distinct cooling of the SST and deepening of the mixed layer occurred within a cyclonic eddy with a thin mixed layer than within an anticyclonic eddy with a thick mixed layer. Sun et al. [56] demonstrated that four typhoons in 2004 successively enhanced and enlarged a cyclonic eddy in the Kuroshio meandering region, south of Japan, due to strong air‐sea interaction and typhoon‐induced upwelling. Although 49 super typhoons passed over 192 cyclonic eddies in the WNP during the period of 2000–2008, only about 10% of these eddies were intensified by the typhoons [57]. Thus, Sun et al. [57] concluded that the impact of typhoons on the strength of cyclonic eddies is inefficient.

Aside from the impact on the intensity of cyclonic eddies, typhoons can directly generate cyclonic eddies via strong wind stress curls and long‐time forcing. Hu and Kawamura [58] found that three TCs separately generated a cyclonic eddy in the northern SCS where they had a loop track and provided sufficient forcing time for the eddy generation. Yang et al. [59] also reported that two mesoscale cyclonic eddies were separately induced in the SCS and the WNP by long time forcing of strong wind stress curls associated with Typhoons Hagibis and Mitag in 2007, respectively. Similar phenomenon has been revealed in other studies (e.g., [57, 60]). It is interesting that these cases happened in the SCS or in the west of the WNP where typhoons often travel slowly and easily change their moving direction. The main reason may be that this area is close to or located at the western edge of the North Pacific Subtropical High around which clockwise atmospheric circulation steers the movement of typhoons in the WNP to a great degree.

Compared with cyclonic eddies, the mixed layer is usually thicker and the thermocline is deeper within anticyclonic eddies. As a result, anticyclonic eddies are not significantly impacted by typhoons. Lin et al. [49] demonstrated that Super‐typhoon Maemi (2003) caused a weak sea surface cooling of only about 0.5°C within an anticyclonic eddy, quite smaller than that (2°C) outside the eddy, while the typhoon was rapidly intensified due to the presence of the warm eddy. They reckoned that a thick mixed layer in the anticyclonic eddy prevented deep, cold water from being entrained into the surface layer and at the same time the warm water in the thick layer contributed to the intensification of the typhoon and sustained the typhoon's intensity due to a reduced negative feedback from typhoon‐induced weak sea surface cooling. The results obtained by Nam et al. [55] showed that under the influence of Typhoon Man‐Yi (2007), the entrainment within an anticyclonic eddy was weak because of the thick mixed layer, resulting in small changes of temperature profiles.

### **5. Storm surges and large waves**

**4. Influence on mesoscale eddies**

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

in an anticyclonic eddy.

area of 3,500,000 km2

of cyclonic eddies is inefficient.

A mesoscale eddy is one vortex with its core surrounded by closed circulations. In the ocean, it has a time scale of 10–100 days and a dimension of 10–100 km in horizontal and 100–1000 m in vertical. It can travel a distance of 100–1000 km during its lifetime of several months. According to the rotation direction of closed circulations, mesoscale eddies are classified into cyclonic (counterclockwise) and anticyclonic (clockwise) eddies. Because the temperature inside a cyclonic eddy is usually lower than that of surrounding water, the cyclonic eddy is also called cold eddy. Inversely, an anticyclonic eddy is also called warm eddy. Sea level anomaly with respect to local mean sea level is negative in a cyclonic eddy whereas it is positive

Mesoscale eddies are ubiquitous in the open ocean and marginal seas (e.g., the SCS) of the WNP. There are two eddy‐rich regions in the open ocean of the WNP [47–49]: the Kuroshio Extension region (30°–40°N, 140°E–180°W) and the North Pacific Subtropical Countercurrent region (18°–25°N, 122°E–160°E). The SCS is a well‐known eddy‐rich marginal sea with a large

Pre‐existing cyclonic eddies can be intensified after a typhoon passes over. Zheng et al. [52, 53] revealed that the SST in pre‐existing cyclonic eddies markedly decreased after the passage of Typhoon Hai‐Tang (2005) because cold deep water was easily brought up due to uplifted thermocline, large current shear just below the mixed layer, and small thermal inertia within the cyclonic eddies. Chiang et al. [8] showed that when Typhoon Kai‐Tak (2000) passed over the West Luzon Eddy (a cyclonic eddy), it produced an unusually intense SST drop of about 10.8°C and high chlorophyll *a* (Chl‐*a*) concentration in the eddy. Lai et al. [54] found that a cold eddy was intensified when Typhoon Morakot (2009) passed by the eddy, resulting in a significant cooling. Nam et al. [55] pointed out that during Typhoon Man‐Yi (2007), more distinct cooling of the SST and deepening of the mixed layer occurred within a cyclonic eddy with a thin mixed layer than within an anticyclonic eddy with a thick mixed layer. Sun et al. [56] demonstrated that four typhoons in 2004 successively enhanced and enlarged a cyclonic eddy in the Kuroshio meandering region, south of Japan, due to strong air‐sea interaction and typhoon‐induced upwelling. Although 49 super typhoons passed over 192 cyclonic eddies in the WNP during the period of 2000–2008, only about 10% of these eddies were intensified by the typhoons [57]. Thus, Sun et al. [57] concluded that the impact of typhoons on the strength

Aside from the impact on the intensity of cyclonic eddies, typhoons can directly generate cyclonic eddies via strong wind stress curls and long‐time forcing. Hu and Kawamura [58] found that three TCs separately generated a cyclonic eddy in the northern SCS where they had a loop track and provided sufficient forcing time for the eddy generation. Yang et al. [59] also reported that two mesoscale cyclonic eddies were separately induced in the SCS and the WNP by long time forcing of strong wind stress curls associated with Typhoons Hagibis and Mitag in 2007, respectively. Similar phenomenon has been revealed in other studies (e.g., [57, 60]). It is interesting that these cases happened in the SCS or in the west of the WNP where typhoons

appear frequently, which may influence the structure and evolution of some eddies.

and an average depth of over 2000 m [50, 51]. In these regions, typhoons

A typhoon usually induces low frequency abnormal variations in sea level (namely storm surges) and high frequency large waves (commonly called typhoon waves or storm waves). The main disaster causes accompanying with typhoons encompass storm surge, huge wave, strong wind and heavy rainfall. In low‐lying coastal areas and islands, storm surge together with high spring tide is most destructive.

In the open ocean, the abnormal sea level is mainly controlled by pressure drop in the center of a TC according to the inverted atmospheric pressure effect. Although the cyclonic wind tends to reduce sea level height by Ekman transport, a round water bulge is often produced in response to low pressure in the center of the cyclone. At the same time, some long waves are induced and propagate forward faster than the cyclone itself. When these free waves arrive at the shallow and wide shelf waters near Mainland China, their wave height rises because of shallow depth and then their energy is gradually dissipated by bottom friction.

When a TC comes close to the coastal area, strong wind forcing is dominant in the generation of storm surges. The combination of onshore wind and low pressure results in a maximum storm surge near the TC center. The system of topography, bathymetry and coastline plays an important role in the distribution of storm surges. The inertial oscillations in storm surges generated by a strong typhoon would disappear quickly due to overdamping in the ultra‐ shallow water of the Bohai Sea. Surge waves can be induced by a typhoon and propagate counterclockwise or from north to south along the coasts in the China Seas, including the Bohai Sea, Yellow Sea, ECS, and SCS [61, 62]. Edge waves may appear when a typhoon moves in parallel with the coastline. The dissipation effect of bottom friction on these coastal trapped waves is significant in the wide shelfs of the China Seas. After a typhoon makes landfall and moves away, Ekman setup can generate surges along the coasts such as Tottori coasts of the East Sea [63].

Tides affect storm surges via tide‐surge interaction in the shallow waters, which is significant in the coastal areas in the WNP [64–67]. In the Taiwan Strait, the tide‐surge interaction is intensified because bottom friction is enhanced by strong tidal currents and storm‐induced currents in the along‐strait direction [68]. As a result of the tide‐surge interaction, obvious oscillations appear in storm surges. The waves accompanying with a typhoon modulate storm surges by wave radiation stress, wave‐dependent bottom shear stress and surface wind stress [69, 70]. Wave setup due to the breaking of wind waves associated with a typhoon can directly contribute to storm surges in the nearshore.

Typhoon‐induced storm surges and huge waves often lead to loss of lives and damage to property in coastal regions and islands with a large population. Typhoon Saomai (2006) with a maximum wind of 75.8 m/s [71] and a lowest pressure of 915 hPa was the strongest one of typhoons making landfall on Mainland China in the past 50 years. When it made landfall at the coast of Zhejiang Province, the central pressure was still 920 hPa and a momentary maximum wind of 68 m/s was recorded. During its influence period, the highest wave recorded by a buoy (27.5°N,122.53°E) in the ECS was 8.6 m. Saomai induced devastating storm surges and a maximum storm surge of 4.01 m was observed at Aojiang station, Zhejiang Province. As reported in the Bulletin of Marine Disaster of China, the storm surges superposing on a high spring tide caused 230 persons dead and 96 missing, and a damage of more than 7 billion yuan to property. Another devastating typhoon, Haiyan, hit Philippines islands on 8 November 2013. Before its landfall, the maximum sustained wind was nearly 88 m/s and the momentary maximum wind was about 105 m/s. Storm surges induced by Typhoon Haiyan inundated a large coastal area in Philippines and caused a catastrophic damage [72, 73]. The inundation depth was up to more than 7 m in Tacloban city. During the typhoon, 6300 persons were killed and 1061 were still missing [72].

Storm surges, together with tides, mainly determine extreme sea levels and their spatial pattern along the coasts of the WNP [66, 74]. For some places where tidal range is small and tide‐surge interaction is weak, storm surges play a more important role in the extreme sea levels [75]. Storm surges also contribute to the local increasing trend of annual maximum water level at coastal stations [76].

#### **6. Biogeochemical and biological responses**

TCs definitely have significant impacts on biogeochemical and biological processes in the oceans. Biological responses, including associated biogeochemical aspects, to typhoons have been observed in the WNP by ship surveys, buoys and satellite remote sensing data (e.g., [77– 79, 20]), which will be discussed in the following text. Although typhoons may have direct and indirect effects on the carbon cycling in the WNP (e.g., [80–82]) and potentially contribute to global climate change, detailed contents of these effects are not included here. Many investi‐ gations show that various typical waters (open ocean waters, marginal seas, shelf waters) display some different biological responses to typhoons with unique mechanisms.

#### **6.1. In open ocean water**

moves away, Ekman setup can generate surges along the coasts such as Tottori coasts of the

Tides affect storm surges via tide‐surge interaction in the shallow waters, which is significant in the coastal areas in the WNP [64–67]. In the Taiwan Strait, the tide‐surge interaction is intensified because bottom friction is enhanced by strong tidal currents and storm‐induced currents in the along‐strait direction [68]. As a result of the tide‐surge interaction, obvious oscillations appear in storm surges. The waves accompanying with a typhoon modulate storm surges by wave radiation stress, wave‐dependent bottom shear stress and surface wind stress [69, 70]. Wave setup due to the breaking of wind waves associated with a typhoon can directly

Typhoon‐induced storm surges and huge waves often lead to loss of lives and damage to property in coastal regions and islands with a large population. Typhoon Saomai (2006) with a maximum wind of 75.8 m/s [71] and a lowest pressure of 915 hPa was the strongest one of typhoons making landfall on Mainland China in the past 50 years. When it made landfall at the coast of Zhejiang Province, the central pressure was still 920 hPa and a momentary maximum wind of 68 m/s was recorded. During its influence period, the highest wave recorded by a buoy (27.5°N,122.53°E) in the ECS was 8.6 m. Saomai induced devastating storm surges and a maximum storm surge of 4.01 m was observed at Aojiang station, Zhejiang Province. As reported in the Bulletin of Marine Disaster of China, the storm surges superposing on a high spring tide caused 230 persons dead and 96 missing, and a damage of more than 7 billion yuan to property. Another devastating typhoon, Haiyan, hit Philippines islands on 8 November 2013. Before its landfall, the maximum sustained wind was nearly 88 m/s and the momentary maximum wind was about 105 m/s. Storm surges induced by Typhoon Haiyan inundated a large coastal area in Philippines and caused a catastrophic damage [72, 73]. The inundation depth was up to more than 7 m in Tacloban city. During the typhoon, 6300 persons were killed

Storm surges, together with tides, mainly determine extreme sea levels and their spatial pattern along the coasts of the WNP [66, 74]. For some places where tidal range is small and tide‐surge interaction is weak, storm surges play a more important role in the extreme sea levels [75]. Storm surges also contribute to the local increasing trend of annual maximum water level at

TCs definitely have significant impacts on biogeochemical and biological processes in the oceans. Biological responses, including associated biogeochemical aspects, to typhoons have been observed in the WNP by ship surveys, buoys and satellite remote sensing data (e.g., [77– 79, 20]), which will be discussed in the following text. Although typhoons may have direct and indirect effects on the carbon cycling in the WNP (e.g., [80–82]) and potentially contribute to global climate change, detailed contents of these effects are not included here. Many investi‐

East Sea [63].

contribute to storm surges in the nearshore.

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

and 1061 were still missing [72].

**6. Biogeochemical and biological responses**

coastal stations [76].

In the deep ocean water of the WNP, Merritt‐Takeuchi and Chiao [83] found an obvious growth of biological substances after the passage of Typhoon Xangsane (2006) which brought nutrients from the depths to the surface layer. There is a negative correlation between SST and Chl‐*a* with the correlation coefficient of –0.67. Salyuk et al. [84] demonstrated that 81% of 123 TCs increased Chl‐*a* concentration after the TC passage, which could last about 2 weeks. Based on multiple satellite observations and numerical experiments, however, Lin [85] showed that only two cases (18%) among 11 typhoons in 2003 induced phytoplankton blooms in the WNP.

The biological response of the upper ocean depends on the translation speed, spatial size, moving track and intensity of a typhoon and pre‐existing ocean conditions. There is a positive correlation between Chl‐*a* concentration and wind speed [83, 86]. Lin [85] found that 9 of 11 typhoons in 2003 did not induce phytoplankton blooms in the WNP. Among the 9 typhoons, eight typhoons had relatively small size, fast translation speed and insufficient wind intensity, and then they only caused weak responses in the ocean with a deep nutricline/mixed layer. Owing to the presence of warm eddy, the other typhoon, Maemi, was not able to induce phytoplankton bloom although it was very strong with a maximum wind of about 77 m/s. Typhoon Haitang (2005) did not enhance Chl‐*a* in the sea area east of Taiwan because of high translation and a shallower MLD than the nutricline [87].

#### **6.2. In the SCS**

Typhoons are very active in the SCS. They often trigger phytoplankton blooms in this oligo‐ trophic sea. Sun et al. [86] found that Typhoon Hagibis (2007) with a steep turn track had a significant effect on the surface Chl‐*a* concentration in the SCS. Its long forcing time is favorable for the enhancement of Chl‐*a* concentration although it is just a category 1 typhoon. Typhoon Kai‐Tak (2000) increased surface Chl‐*a* concentration by 30‐fold on average in the SCS during three days [49]. As a result, Kai‐Tak alone contributed 2–4% of the SCS's annual new primary production. A similar enhancement of Chl‐*a* concentration appeared separately after typhoons Lingling (2001) and Damrey (2005) in the SCS [20, 88]. Zhao et al. [89] conservatively estimated that typhoons accounted for 3.5% of annual primary production in the SCS during the period of 1945–2005. Chen et al. [90] estimated that 5–15% of annual new primary production in the SCS was attributed to typhoons during 2003–2012.

Eddies are ubiquitous in the SCS. Some investigations have shown that Chl‐*a* concentration enhanced by typhoons may be associated with cold core eddies in this area (e.g., [22, 91]). Chen and Tang [92] found that in the SCS a cold eddy was generated where Typhoon Linfa (2009) hovered, and subsequently an eddy‐feature phytoplankton bloom was induced. The consis‐ tence between the bloom pattern and the cold eddy suggested that the typhoon‐induced eddy potentially brought nutrients upward to the surface water, which contributed to the bloom.

Chen et al. [90] compared typhoon‐enhanced primary production in the SCS with that in the subtropical ocean water of the WNP during the period of 2003–2012. Their results showed that the annual mean carbon fixation induced by typhoons was more in the SCS than in the ocean water. This is because the mixed layer is thicker and the nutricline depth is deeper in the latter in spite of its larger area and more super typhoons appearing there.

The biological response to typhoons can happen not only in the surface layer but also in the subsurface. Ye et al. [93] found that a Chl‐*a* bloom appeared in the subsurface layer (20–100 m depth) of the SCS after the passage of Typhoon Nuri (2008) and lasted for three weeks. This subsurface bloom was stronger and its life was longer than the synchronous surface Chl‐*a* bloom. Previous estimates of the contribution from typhoons to annual primary production in the SCS were mostly based on the results in the surface layer using remote sensing data. On this aspect, these estimates probably underestimate the actual contribution of typhoons.

#### **6.3. In continental shelf waters**

There are many wide and shallow continental shelf regions in the WNP, mostly located in the China Seas. The ecosystem in these shelf regions can become more productive after typhoons pass through [77, 78, 87].

The continental shelf of the southern ECS, northwest of Taiwan, is a typical oligotrophic and strong stratification area during summer. However, Shiah et al. [78] found that all chemical and biological parameters measured in a survey after the passage of Typhoon Herb (1996) were much larger than normal summer conditions. The typhoon caused primary production, particulate organic carbon (POC) concentration, bacterial production, and biomass to increase by at least two‐fold. Their results indicated that wind mixing, re‐suspension and terrestrial runoff associated with the typhoon were responsible for these responses. Based on multi‐ satellite observations, Chang et al. [16] showed that Typhoon Hai‐Tang (2005) induced upwelling and increased Chl‐*a* concentration in this shelf region, persisting for more than 10 days. The upwelling was likely caused by Ekman pumping due to strong typhoon wind and wind‐driven shoreward intrusion of Kuroshio water along the shelf break. Siswanto et al. [94] also demonstrated that long‐lasting southerly winds accompanying with typhoons can force Kuroshio current axis to move toward the shelf of the southern ECS, inducing upwelling. This process uplifts nutrients and then increases new productivity, which contributes 0.6–11.8% of the summer‐fall new productivity in the ECS. In addition to the effects on primary production represented by Chl‐*a*, typhoons may have influence on phytoplankton composition. Chung et al. [95] observed that a diatom bloom was induced in the southern ECS after the passage of Typhoon Morakot (2009) and the species composition was changed also.

Zhang et al. [96] observed that the surface Chl‐*a* concentration in the continental shelf southeast of Hainan Island was increased by 38.5% after the passage of tropical storm Washi (2005). Chen et al. [97] showed that the primary productivity and nitrate‐uptake‐based new production in the upstream Kuroshio close to southern Taiwan were enhanced after the passage of three typhoons in 2007 by riverine mixing associated with the typhoons. After Typhoon Malou (2010) passed over Sagami Bay in the central part of Japan, both Chl‐*a* concentration and bacterial abundance increased at an inshore station due to terrestrial runoff and sediment resuspension while only Chl‐*a* concentration rose at an offshore station due to terrestrial runoff [98].

There are some contrary cases. Zhou et al. [99] found that Typhoon Fengshen (2008) destroyed the pre‐existing upwelling and meanwhile caused freshwater plume to spread in the conti‐ nental shelf of the northeastern SCS, which led to nutrient‐limited conditions. As a result, a negative phytoplankton growth rate appeared a week after the typhoon passage. Zhao et al. [100] also found that a sharp decrease of Chl‐*a* concentration was caused by Typhoon Matsa (2005) in the nearshore area of the ECS.

Therefore, biological response to typhoons in the continental shelfs, especially nearshore areas, is very complex. It changes with different sea areas and different time periods, depending on specific circumstances such as geographical configuration, hydrological environment as well as the existence of rivers or not.
