importance of tropical cyclone heat potential for … of tropical cyclone heat potential for...

427

Journal of Oceanography, Vol. 63, pp. 427 to 447, 2007

Keywords:⋅⋅⋅⋅⋅ Tropical cycloneheat potential,

⋅⋅⋅⋅⋅ tropical cycloneintensity,

⋅⋅⋅⋅⋅ sea surfacetemperature,

⋅⋅⋅⋅⋅ rapid intensifica-tion.

* Corresponding author. E-mail: [email protected]

Copyright©The Oceanographic Society of Japan/TERRAPUB/Springer

sification, SST or tropical cyclone heat potential (TCHP)?This issue continues to be controversial. According to thereply of Scharroo (2006) to the comment of Sun et al.(2006) on the importance of dynamic topography in theintensification of Hurricane Katrina (2005), high SST isa necessary but insufficient condition for hurricane in-tensification (Scharroo et al., 2005).

The ocean thermal energy is defined as TCHP (Gray,1979) and is calculated by summing the heat content in acolumn where the sea temperature is above 26°C (Leipperand Volgenau, 1972). A conventional methodology forestimating TCHP using the European Remote SensingSatellite-2 (ERS-2) and TOPEX/Poseidon satellitealtimetry observations has recently been developed (e.g.Shay et al., 2000). We can find the results of real-timemonitoring of TCHP (Goni and Trinanes, 2003) via theWeb page (http://www.aoml.noaa.gov/phod/cyclone/data/[cited 30 September 2006]).

Importance of Tropical Cyclone Heat Potential forTropical Cyclone Intensity and Intensificationin the Western North Pacific

AKIYOSHI WADA* and NORIHISA USUI

Meteorological Research Institute, Japan Meteorological Agency,Nagamine, Tsukuba, Ibaraki 305-0052, Japan

(Received 2 October 2006; in revised form 5 December 2006; accepted 11 January 2007)

Which is more important for tropical cyclone (TC) intensity and intensification, seasurface temperature (SST) or tropical cyclone heat potential (TCHP)? Investigationsusing best-track TC central pressures, TRMM/TMI three-day mean SST data, andan estimated TCHP based on oceanic reanalysis data from 1998 to 2004, show thatthe central pressure is more closely related to TCHP accumulated from TC formationto its mature stages than to the accumulated SST and its duration. From an oceanicenvironmental viewpoint, a rapid deepening of TC central pressure occurs when TCHPis relatively high on a basin scale, while composite distributions of TCHP, verticalwind shear, lower tropospheric relative humidity, and wind speed occurring in casesof rapid intensification are different for each TC season. In order to explore the influ-ence of TCHP on TC intensity and intensification, analyses using both oceanicreanalysis data and the results of numerical simulations based on an ocean generalcirculation model are performed for the cases of Typhoons Chaba (2004) and Songda(2004), which took similar tracks. The decrease in TCHP due to the passage of Chabaled to the suppression of Songda’s intensity at the mature stage, while Songda main-tained its intensity for a relatively long time because induced near-inertial currentsdue to the passage of Chaba reproduced anticyclonic warm eddies appearing on theleftside of Chaba’s track before Songda passed by. This type of intensity-sustenanceprocess caused by the passage of a preceding TC is often found in El Niño years.These results suggest that TCHP, but not SST, plays an important role in TC inten-sity and its intensification.

1. IntroductionOne of the decisive factors influencing the tropical

cyclone (TC) intensity and its intensification is ocean ther-mal energy in the upper ocean. The relationship betweenTC potential intensity and sea surface temperature (SST)has been discussed for the past decade (e.g. DeMaria andKaplan, 1994a). The intensity-SST relationship also playsan important role in statistical intensity predictionschemes such as the National Hurricane Center Statisti-cal Hurricane Intensity Prediction Scheme (SHIPS)(DeMaria and Kaplan, 1994b, 1999; DeMaria et al., 2005).However, recent studies focus on the relationship betweenthe intensity and the upper ocean heat content. Which ismore important for the potential intensity and the inten-

428 A. Wada and N. Usui

This conventional methodology for TCHP estima-tion has also been applied to case studies of hurricanes inthe North Atlantic (Shay et al., 2000; Goni and Trinanes,2003). These studies suggested that warm core eddies orLoop Current in the Gulf of Mexico were important indecreasing the impact of the negative feedback on theintensification of hurricanes. In general, a strong windaccompanied by TCs stirs and cools the underlying seawater. This leads to the reduction of sensible and latentheat fluxes from the ocean to the atmosphere (e.g. Ginis,1995). In contrast, a deep warm layer with a positive seasurface height anomaly (SSHA) is hardly cooled by astrong wind. This helps TCs maintain their intensity orintensify rapidly in places where the SSHA is relativelyhigh. This effect has been found not only in hurricanes inthe North Atlantic (Bao et al., 2000; Hong et al., 2000;Shay et al., 2000) but also in typhoons in the westernNorth Pacific (Holliday and Thompson, 1979; Lin et al.,2005). This process is one of the environmental factorsthat control the intensity and intensification of TCs(Emanuel et al., 2004; DeMaria et al., 2005).

An interesting question regarding these TC-oceaninteractions is whether we can discuss the predictabilityof TC potential intensity statistically using the abovementioned TCHP, instead of using SST as in the SHIPSapproach of DeMaria et al. (2005). The predictability isexpected to apply not only to intensification but also tothe overall life stages of TCs. Wada (2002) reported in

the case of Typhoon Rex (1998) that the evolution of asimulated TCHP was well correlated with that of the TCcentral pressure. Another interest is how and to what ex-tent warm core eddies in the western North Pacific withhigh TCHP vary following TC passage, influencing sub-sequent TC intensity. According to Bender and Ginis(2000), the intensity prediction for Hurricane Fran (1996)was significantly improved when the coupled model wasrun with the well-represented cold wake of HurricaneEdouard (1996).

This paper is organized as follows. Section 2 de-scribes the data and the specification of the ocean gen-eral circulation model used in the present study. Section3 describes the statistical relationship between minimumcentral pressures (MCPs) and TCHPs around the TCcenter. In Section 4 we investigate the relationship be-tween locations of rapid intensification and atmosphericor oceanic environmental factors such as vertical windshear, lower-tropospheric relative humidity and wind ve-locity, and TCHP. Section 5 describes the relationshipbetween TCHP and intensity or intensification in the casesof Typhoons Chaba (2004) and Songda (2004). The im-pact of TCHP on the intensification of Chaba and Songdais investigated in Section 6. Section 7 explores thebaroclinic ocean response to a moving TC, related to themaintenance of Songda’s peak intensity, from the view-point of TCHP variation. In Section 8 further discussesthe relationship between MCP and TCHP accumulated

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Development

Mature Decay

(a) (b)

(c) (d)

LongitudeLongitude

Longitude Longitude

Lat

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Lat

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Lat

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Formation

Fig. 1. Frequency and locations of tropical cyclones at the (a) formation, (b) development, (c) mature, and (d) decay stages from1998 to 2004 using RSMC tropical cyclone best-track data. The distribution expands into the western North Pacific during thetropical cyclone formation stage, while it is concentrated around Okinawa Island as the stage shifts from development todecay.

Importance of TCHP for TC Intensity and Intensification 429

from formation to the mature stages. A formation mecha-nism for warm core eddies in the western North Pacificand its impact on Songda’s intensity is also discussed inSection 8. Section 9 is devoted to an overall summary.

2. Methods

2.1 Tropical cyclone best-track dataWe employed Regional Specialized Meteorological

Center (RSMC) TC best-track data obtained from 1998to 2004. The data include the TC center positions, its cen-tral pressure, and its 10-minute-averaged maximum sus-tained wind measured four times a day, at 0600, 1200,1800, and 2400 UTC and so on. The data include theseelements taken eight times a day when a TC is locatedwithin 300 km of the coast of Japan, or every hour whena TC makes landfall in Japan.

Here, each TC central position over the ocean from1998 to 2004 is counted in 5 × 5° bins for each life stage:TC formation, development, mature, and decay stages.TC formation is defined as occurring when a TC reacheswind speeds of at least 17 m/s. The development stage isa time when the central pressure decreases continuously,including a temporary period when the central pressuredoes not change. The post stages, defined as the matureand decay stages, occur after the central pressure of theTC reaches MCP. The mature stage is defined as the pe-riod from the end of the development stage to the begin-ning of the decay stage. The decay stage is a period dur-ing which the central pressure rises monotonically.

Figure 1 illustrates the horizontal distribution of TClocations for each life stages. TCs are frequently formedaround 120°E, 20°N in the South China Sea and around130° to 140°E, 18°N (Fig. 1(a)). In Fig. 1(b), TCs de-velop west of the Philippines in the South China Sea oraround Okinawa Island. Around Okinawa Island, TCs fre-quently sustain their intensity as well (Fig. 1(c)). In fact,the horizontal distribution of TC locations in the maturestage is similar to that of the development stage. How-ever, the TC locations in the mature stage tend to shiftnorthward compared with those in the development stage.A peak in the number of TC locations in the decay stageis found along the Japan coast, partly because many ob-servations are made around Japan (Fig. 1(d)). Neverthe-less, the appearance of a peak with an eastward exten-sion at high latitudes is one of the characteristics of thedistribution of TC locations in the decay stage.

2.2 Satellite sea surface temperatureThe Tropical Rainfall Measuring Mission (TRMM)/

TRMM Microwave Imager (TMI) three-day mean SSTcovering January 1998 to December 2004 was used toreconfirm the relationships between SST and the TC maxi-mum potential intensity (MPI) derived from empiricalfunctions of the monthly SST. The TRMM/TMI three-day mean SST covers a global region extending from 40°Sto 40°N with a horizontal resolution of 0.25°. One of theimportant features of TRMM/TMI microwave retrievalsis that SST can be measured through clouds.

We also used a daily Microwave Optimally Interpo-

Fig. 2. Relationship between three-day mean SST and tropical cyclone best-track central pressure from 1998 to 2004. Crossmarks with gray shaded squares show the SST at the formation stage. Circles show the SST at the development stage. Filledsquares show the SST at the mature stage. Triangles show the SST at the decay stage. Three empirical functions (DeMaria andKaplan, 1994a; Whitney and Hobgood, 1997; Baik and Paek, 1998) are drawn in the same panel. The maximum potentialintensity derived from the three empirical functions is usually overestimated compared to the observations.


lated (OI) SST (MW OI SST), for 2004 only. The MW OISST covers the global ocean with a horizontal resolutionof 0.25°. MW OI SSTs include data from TRMM/TMIand from Aqua/Advanced Microwave Scanning Radiom-eter for Earth observing system (AMSR-E) satellite radi-ometers. Because AMSR-E was initiated to perform ob-servations beginning in May 2002, the SST data were hereonly applied to case studies in 2004.

The relationship between the three-day mean SSTand the RSMC best-track central pressure is denoted inFig. 2 with three MPI empirical functions of the monthlySST (DeMaria and Kaplan, 1994a; Whitney and Hobgood,1997; Baik and Paek, 1998). Since some of the MPI em-pirical formulas (DeMaria and Kaplan, 1994a; Whitneyand Hobgood, 1997) are functions of SST and maximumwind velocity, the MPI expressed by maximum wind ve-locity is replaced with that expressed by central pressure,using an empirical formula representing the modified ver-sion of Atkinson and Holliday’s wind-pressure relation-ship (Atkinson and Holliday, 1977) proposed by Knaffand Zehr (2006). Since the difference in sea-level pres-sure between the environment and the center is derivedfrom this formula, we assume that the environmental pres-sure is 1010 hPa.

Baik and Paek (1998) determined empirical functionsthat cap storm intensity over the western North Pacific,using only SST. They applied least-squares fitting to ex-ponential and polynomial functions of order up to 5 us-ing observed maxima and 99th intensity percentile data.In the present study, the regression curve for the 99th in-tensity percentile is used for comparison. Note that theabove mentioned MPI empirical functions utilize onlySST, with attempts made to estimate an upper bound onTC intensity for given atmospheric and oceanic condi-tions (see a review of Wang and Wu, 2004).

Figure 2 reveals that TCs are formed when the SSTis above 26.5°C, which is consistent with the previousreport from Palmén (1948). In contrast, SST variationsare not always dependent on TC intensity, especially dur-ing the development, mature and decay stages. The in-tensity rarely reaches its MPI because a TC may be af-fected by many other possible negative influences (Evans,1993). Nevertheless, the empirical function of Baik andPaek (1998) is closest to the relationship between theTRMM/TMI three-day SST and the lower bound of thebest-track central pressure. The empirical function is morereasonable, especially during the decay stage when theSST is always below 27°C. It is worth noting that thevalue of the lower bound of the best-track central pres-sure tends to be high compared to the value determinedfrom empirical functions in which SST is above 29°C.This implies that the lower bound of the best-track cen-tral pressure may not always be reliable enough to use asverification of numerically predicted central pressures

when the SST is above 29°C.The quantitative difference between the results of the

present study and those of previous studies may be causedby the representative depth of each SST. According toDonlon et al. (2002), the TRMM/TMI SST (TSST) is rep-resentative of the SST at the bottom of the skin SST, atdepths within a thin layer (nearly 500 µm). The repre-sentative depth of the TSST or SST on a low-frequency(6 to 10 GHz) microwave radiometer is shallower than 1mm. In contrast, the representative depth of bucket SST(BSST) is a few meters. The difference between BSSTand TSST is usually above 1°C during daytime.

2.3 Atmospheric reanalysis dataIn order to investigate the impact of the atmospheric

environment on rapid intensification of TCs from 1998to 2004, the National Centre for Environmental Predic-tion (NCEP)—Department of Energy Atmospheric ModelIntercomparison Project reanalysis (abbreviated NCEP R-2) (Kanamitsu et al., 2002) is used in the present study.NCEP R-2 is also used to provide the atmospheric forc-ing to perform numerical simulations of the ocean re-sponse to Typhoons Chaba (2004) and Songda (2004).The horizontal resolution is 2.5° square in a latitude-lon-gitude coordinate system, and the time interval is sixhours. The horizontal resolution of the daily atmosphericforcing data for the numerical simulation is 1.875° by lin-ear interpolation.

In order to realistically reflect the cyclonic stormcirculations of Chaba and Songda in the atmospheric forc-ing, each Rankin vortex, determined from the central po-sition, maximum sustained wind velocity, and a distanceof the radius of 50 knots wind velocity, is merged intothe JMA/Global ANALysis (GANAL) data as a TC bo-gus (Wada, 2002). The horizontal resolution of JMA/GANAL is 1.25° square in a latitude-longitude coordi-nate system and the time interval is six hours.

2.4 Oceanic reanalysis data and tropical cyclone heatpotentialOceanic reanalysis data from the North Pacific ver-

sion of the Meteorological Research Institute multivariateocean variational estimation (MOVE) system is used(Usui et al., 2006) to estimate daily TCHP from 1998 to2004. The MOVE system includes an ocean general cir-culation model (OGCM: Ishikawa et al., 2005) and a vari-ational analysis scheme to synthesize observations suchas sea temperature, salinity, and sea-surface height in theirown time and space. The horizontal resolution of MOVEis 0.5° square in a latitude-longitude coordinate system.The model domain of MOVE covers from 15°S to 65°Nand 100°E to 75°W.

Using this reanalysis data, the TCHP, based on thatof Leipper and Volgenau (1972), is calculated using the


following formula:

Q C T ZTCHP ph

H

= −( ) ( )=∑ ρ 26 1

0

∆

where ρ is the density of the sea water at each layer, Cp isthe specific heat at constant pressure, T is the sea tem-perature, and ∆Z is the thickness at each layer. When T isbelow 26°C, it is reset to 26°C. The ocean heat contentQTCHP represents TCHP in the upper ocean where the seatemperature is above 26°C. In the present study, the unitis transformed from kcal cm–2 to kJ cm–2. Figure 3presents the relationship between the TRMM/TMI three-day mean SST and the TCHP. The TCHP increasesexponentially while the SST increases linearly. However,the relationship between the SST and the TCHP is notalways unique. The second power of the correlation co-efficient between them is approximately 0.45. TCHP issometimes high (low) during low (high) SST when theSST is above 26°C. This means that TCHP is not onlydependent on the SST but also on the depth of the 26°Cisotherm.

2.5 Ocean general circulation modelThe OGCM is a version of the community ocean

model developed by the Meteorological Research Insti-tute (MRI) (Ishikawa et al., 2005). Runs of the MOVEsystem and the MRI Community Ocean Model(MRI.COM) are usually forced by daily wind stress, heatflux, and fresh water flux data provided from NCEP R-2.

In the present numerical simulations, the horizontalresolution of MRI.COM is 0.25° square in a latitude-lon-

gitude coordinate system. The model domain covers 10°to 50°N and 120° to 160°E. MRI.COM has a total of 54layers. The maximum bottom depth is set to 5625 m. Thethickness at the bottom is 250 m. There are 17 layers abovea 200-m depth at levels of 0.5 m, 1.5 m, 4.0 m, 7.0 m,12.0 m, 18.0 m, 26.0 m, 38.0 m, 50.0 m, 66.0 m, 82.0 m,100.0 m, 118.0 m, 138.0 m, 158.0 m, 178.0 m, and 200.0m. The turbulent closure scheme of Noh and Kim (1999)is used. Tuning parameters α and m in Noh and Kim(1999) are 5 and 100 in the present study. The empiricalbulk formula (Kondo, 1975) is employed to calculate sen-sible and latent heat fluxes. The water flux is correctedby restoring sea surface salinity to the climatology with arestoring time of 24 hours.

3. Relationship between Tropical Cyclone Heat Po-tential and the Central Pressure of Tropical Cy-clonesAs discussed in the previous section, the SST-derived

MPI is usually deeper than the best-track central pres-sure. This indicates that SST alone is an inadequate pre-dictor of TC intensity (Evans, 1993). This is because TCintensity is controlled not only by SST but also by thethermodynamic environment, the vertical shear, and dis-turbances in the upper troposphere (Emanuel et al., 2004).In addition, the SST-derived MPI is not influenced byocean thermal energy or the upper oceanic structure un-derneath the storms. The concern is how accurately theTCHP-derived MPI can predict the best-track central pres-sure when we use TCHP as a substitute for SST.

Parameters used for the above mentioned investiga-tion are as follows.

• Attainable TC intensity is defined as sea-levelpressure when a TC reaches its minimum central pres-sure (MCP), corresponding to the best-track MCP.

• Duration of a TC is defined as the period fromthe formation stage to the mature stage, when the centralpressure has just reached the MCP.

• Accumulated TCHP (ATCHP) and AccumulatedSST (ASST) are defined as the accumulation and integra-tion of the average of THCP or SST within a 1.5° squarearound the TC center during the six-hour duration. It isnoted that attainable TC intensity, duration, ATCHP, andASST, all are determined uniquely for each TCs.

Figure 4 presents the relationships between the at-tainable TC intensity and the ASST, between the attain-able TC intensity and the duration, and between that in-tensity and the ATCHP from 1998 to 2004. The intensity-ASST correlation coefficient denoted in Fig. 4(a) is thesmallest of the three relationships. The intensity-durationrelationship (Fig. 4(b)) is considered to be closely tied tothe intensity-ASST relationship because the SST is rela-tively high before a TC reaches the attainable TC inten-sity (Fig. 2) and its subsequent variation is relatively

Fig. 3. Relationship between TRMM/TMI three-day mean SSTand TCHP derived from the MOVE system. The regressioncurve, its formula, and the regression coefficient are de-noted in the panel. The regression curve shows the expo-nential increase of TCHP as the SST linearly increases.


small. The longer the duration is, or the longer the TCexperiences high SST during the duration, the strongerthe TC becomes. The intensity-ATCHP relationship ismore reasonable for the attainable TC intensity (Fig. 4(c)).The TC is much stronger when it undergoes high TCHPcontinuously over its duration.

The predictability of potential intensity derived fromthe ATCHP, the duration, and the ASST is examined us-ing the relationship between the attainable TC intensityand the ATCHP, the duration, and the ASST (Fig. 5). As-suming that the distribution of TCHP (SST) does notchange from TC formation to its mature stage, the ATCHP(ASST) can be calculated along the TC best track everysix hours using the TCHP (SST) distribution at the timeof TC formation. The correlation coefficients between thebest-track MCP and ATCHP, duration, and ASST are ap-proximately 0.47, 0.33, and 0.40. The correlation coeffi-cients and gradients of the regression functions shown inFig. 5 indicate that the predicted potential intensity de-rived from ATCHP is superior to that based on durationalone or derived from ASST. However, it is worth notingthat the problem of TC intensity prediction still remains,even though the potential intensity derived from ATCHPagrees with the best-track MCP better than the intensityderived from ASST and the duration. In fact, the poten-tial intensity predicted by the ATCHP tends to underesti-mate strong TCs and overestimate weak TCs. These er-

Fig. 4. Relationships between a tropical cyclone’s central pres-sure at the mature stage and (a) the average accumulatedSST (ASST), (b) the duration of the tropical cyclone, and(c) the average accumulated TCHP (ATCHP). The regres-sion line, the formula, and the coefficient are denoted ineach panel. This shows the analytical value of ATCHP forits relation to tropical cyclone intensity.

Fig. 5. Predictability of tropical cyclone intensity using theaverage accumulated SST (ASST: cross), duration (graycircles), and average accumulated TCHP (ATCHP: filledcircles). Horizontal axis represents the tropical cyclone best-track central pressure; vertical axis is the predicted centralpressure. Tropical cyclone intensity predicted using ATCHPis closer to the best-track intensity than that using SST orduration.


Year The number of rapidintensification cases

The number of TC formationin the western North Pacific

1998 3 161999 0 222000 5 232001 1 262002 4 262003 4 212004 9 29

Total 26 163

Table 1. Rapid intensification cases and tropical cyclone for-mation in the western North Pacific from 1998 to 2004.

rors may be attributed to atmospheric or oceanic envi-ronmental conditions.

4. Relationship between the Atmospheric and Oce-anic Environments and Tropical Cyclone Loca-tions of Rapid IntensificationBesides its potential intensity, the rapid intensifica-

tion of a TC is an important aspect of accurately predict-ing TC intensity. In fact, it is so difficult that rapid TCintensification is numerically predicted using sophisti-cated numerical models with a relatively coarse horizon-tal resolution. According to a statistical analysis by Kaplanand DeMaria (2003), rapid intensification occurs wherethe SST is relatively warm, the lower-tropospheric rela-tive humidity is relatively high, and the vertical shear isrelatively weak. In the present study, rapid intensifica-tion is defined as occurring when the best-track centralpressure falls by more than 10 hPa in six hours. Underthis definition, there were a total of 26 rapid intensifica-tion cases from 1998 to 2004 (Table 1). As seen in Table1, the number of rapid intensification cases is not relatedto the amount of TC formation. Interestingly, half of therapid intensifications occurred from 2003 to 2004.

The interesting point is whether or not rapid TC in-tensification is related to atmospheric and/or oceanic en-vironments, especially TCHP. In the following section wedescribe the relationship between rapid TC intensifica-tion and atmospheric and/or oceanic environmental fac-tors such as TCHP, lower tropospheric relative humidityand wind at a level of 850 hPa, as well as vertical shearbetween levels of 200 hPa and 850 hPa.

4.1 Tropical cyclone heat potentialFigure 6 presents composite maps of TCHP in cases

of rapid TC intensification from 1998 to 2004. The com-posite analysis was made using the daily TCHP observedwhen rapid TC intensification occurs. The TCHP com-posite analysis showed that the composite TCHP was high

south of the equator in the Pacific under conditions ofrapid TC intensification (Fig. 6(a)). The high-TCHP areaexpanded northwestward, probably due to easterly windssouth of a subtropical high. Therefore, from aclimatological viewpoint, rapid intensification occurs inthe western North Pacific where the TCHP is relativelyhigh. In fact, TCs intensify rapidly when the TCHP isabove 120 kJ cm–2.

The horizontal distribution of TCHP may be influ-enced by intraseasonal or intra-annual climate phenom-ena such as the Madden-Julian Oscillation or the El NiñoSouthern Oscillation (ENSO) occurring from 1998 to2004. In other words, atmospheric environments associ-ated with the rapid intensification of TCs are not alwaysuniquely determined climatologically (see the next sec-tion). In the 2000 season (Fig. 6(b)) there were four casesof rapid TC intensification. During that season, the TCHPwas not particularly high south of the equator in the Pa-cific, in contrast to the distribution of TCHP in Fig. 6(a).The TCHP was high south of the equator in the centralPacific during the 2004 season (Fig. 6(c)). In turn, theTCHP east of the Philippines was below 60 kJ cm–2. It isworth noting that the locations of rapid intensificationwere different in the 2000 and 2004 seasons. In particu-lar, the TCHP east of the Philippines during the 2000 sea-son was higher than the TCHP in 2004, although therewas less rapid intensification in the 2000 season than inthe 2004 one. This suggests that the frequency and loca-tions of rapid intensification are influenced not only byhigh TCHP but also by the atmospheric environment andthe TC track, although high TCHP is a necessary condi-tion. In other words, the frequency of rapid intensifica-tion does not increase with higher local TCHPs. In addi-tion, the location is not affected by higher local TCHPs,at least in the western North Pacific. This result is simi-lar to the conclusion reached by Chan and Liu (2004).This conclusion does not involve the ideas of Camp andMontgomery (2001), referring to Shay et al. (1992). Theuse of local TCHP does not always aid in producing anMPI formulation that is more representative of theintensities that storms actually achieve. This is becauseTC rapid intensification requires favorable atmosphericconditions.

4.2 Lower tropospheric relative humidityIn the case of rapid TC intensification, the lower

tropospheric environment was investigated by a compos-ite analysis of the NCEP R-2 winds and relative humidityat a level of 850 hPa for both the 2000 (Fig. 7(a)) and2004 (Fig. 7(b)) seasons. The composite analysis wasmade using the six-hour NCEP R-2 reanalysis of the pe-riod when rapid TC intensification occurs. Rapid intensi-fication occurred in both years when the southwesterlymonsoon from the South China Sea and the easterly wind


on the southern edge of a subtropical high were conflu-ent. However, moisture conditions were different betweenthe two years. In the 2000 season, rapid intensificationoccurred under relatively dry conditions, while intensifi-cation occurred under relatively moist conditions in the2004 season. Since a southwesterly monsoon was trappedwest of the Philippines in the South China Sea during the2000 season, the center of a monsoonal trough was lo-cated in that place. In addition, easterly winds broughtdry air into the area where the rapid intensification oc-curred. Therefore, the area was relatively dry, althoughthe TCHP around the area was relatively high. This means

that high TCHP is not always correlated with high lower-tropospheric relative humidity.

In contrast, the monsoonal trough shifted eastwardin the 2004 season. Under the prevailing atmosphericenvironmental conditions, the rapid intensification areawas covered with moist air, although the TCHP east ofthe Philippines was relatively low. Easterly winds on thesouthern edge of the subtropical high meandered east ofthe area where rapid intensification occurred. The typi-cal wave-like perturbations in the easterly flow over thecentral Pacific, with a wavelength of 2000 km, meridi-onal wave number n = 3, and a period of 5 to 10 days

(a)

(b)

(c)

Fig. 6. Tropical cyclone heat potential (kJ cm–2) in the North Pacific from a composite analysis of cases of rapid intensification(central pressure falling more than 10 hPa per six-hour) (a) from 1998 to 2004, (b) in 2000, and (c) in 2004. The compositeanalysis was made using the daily TCHP at the time of rapid intensification. “Typhoon” symbols are shown at locations ofrapid intensification.


(Reed and Recker, 1971) slowed down upon approachingthe monsoonal westerly flow. At the same time,perturbations moving eastward in the monsoonal wester-lies tended to become trapped in the confluent zone(Holland, 1995). This meandering led to the enhancementof tropical cyclogenesis around the rapid intensificationzone under high TCHP, although further external forcingwas not required for the intensification beyond its owninteraction with the ocean (Briegel and Frank, 1997).

Therefore, the lower-tropospheric relative humidity maynot be a necessary condition, or the reanalysis system withits coarse horizontal resolution may not express the hu-midified condition around a TC. According to Chan andLiu (2004), the interannual variation in monsoonal troughstrength is mostly contributed by the ENSO event ratherthan by the local SST. This implies that a higher localTCHP alone may not significantly impact lower-tropo-spheric relative humidity.

(a)

(b)

Fig. 7. Winds, relative humidity at a height of 850 hPa, and isobaric height of 850 hPa, all from a composite analysis of cases ofrapid intensification using NCEP-DOE AMIPII R-2 data (a) in 2000 and (b) in 2004. “Typhoon” symbols are shown at loca-tions of rapid intensification. The composite analysis was made using a six-hour NCEP R-2 reanalysis at the time of rapidintensification.


4.3 Vertical shear between levels of 200 hPa and 850hPaFigure 8 presents composite maps showing the winds

at a height of 200 hPa and the second power of the verti-cal wind shear between 200 hPa and 850 hPa in the 2000(Fig. 8(a)) and 2004 (Fig. 8(b)) seasons. Figure 8 revealsthat TC rapid intensification occurs under weak verticalwind shear in a divergent field at a level of 200 hPa. Theweak vertical wind shear spread more widely in the 2000season than in the 2004 season. In the 2004 season theamplitude of vertical wind shear was relatively high inthe mid-latitudes, along the southern edge of the subtropi-cal high, and in the South China Sea. Therefore, the areawith weak vertical wind shear in the 2004 season wasmuch smaller than in the 2000 season, leading to a nar-rower divergent field at a level of 200 hPa in the 2004season. The horizontal distribution of the vertical windshear and the divergent field may influence not only thelocations of rapid TC intensification but also the TC tracksfrom 2004, especially the extraordinary number of land-

falls in Japan.Weak vertical wind shear is one of the crucial condi-

tions for TC formation (Gray, 1979). It seems that weakvertical wind shear is a favorable condition for TC inten-sification, too. In other words, strong vertical shear maypossibly have inhibited intensification, even when TCspassed over warm sea water. Rapid TC intensificationinvariably occurred under weak vertical wind shear with-out being inhibited by the unfavorable environment.

5. Impact of Tropical Cyclone Heat Potential on In-tensification of Typhoons Chaba (2004) andSongda (2004)This is a case study investigating the impact of TCHP

on TC intensification. Chaba and Songda became tropi-cal storms over the sea around the Marshall Islands. Theymoved along a similar track, making landfall in Japanwhile passing through a similar course (Fig. 9). Theirdurations, about 11 days for each typhoon, were relativelylong compared to the mean duration of typhoons, which

(a)

(b)

Fig. 8. As Fig. 7 except for winds at the height of 200 hPa and vertical shear between the levels of 200 hPa and 850 hPa.


was about six days in 2004 (Wada, 2005a). Chaba reachedits peak strength with a central pressure of 910 hPa overthe sea west of the Mariana Islands from 23 to 25 August(Fig. 10(a)). In contrast, Songda reached its peak strengthwith a central pressure of 925 hPa over the ocean north-west of Saipan from 31 August to 2 September, and againsouthwest of Okinawa from 4 to 5 September (Fig. 10(b)).The tendencies of intensification for the two typhoonswere different.

During the development and mature stages of Chaba,the SST had been maintained above approximately 29°C.In contrast, the TCHP derived from the reanalysis byMOVE gradually decreased from approximately 120kJ cm–2 to 50 kJ cm–2 during the period. The potentialintensities of Chaba and Songda derived from the em-pirical relation shown in Fig. 4(c) are 948.7 hPa and953.13 hPa. These PIs are weaker than the MCP in thebest-track data. There might also be errors in the best-track data. However, TCHP is not solely responsible forthe rapid intensification of the two typhoons. The differ-ence between the potential intensities of Chaba and

Fig. 9. Tracks of typhoons Chaba (2004) and Songda (2004).“Typhoon” symbols indicate the central position of Chaba(2004). Filled typhoon symbols indicate the central posi-tion of Songda (2004).

Fig. 10. Time series of daily SST, three-day mean SST, TCHP, and central pressure in the cases of Typhoons (a) Chaba (2004) and(b) Songda (2004). Upper panel indicates time series of the SST and TCHP, while lower panel indicates that of central pres-sure.


Songda is only about 5 hPa, while that between theiranalyzed MCPs is 15 hPa. This implies that the TC inten-sity of Songda was subject to influence by air-sea inter-action.

During the decay stage of Chaba, the central pres-sure rose sharply when the typhoon passed through a low-TCHP area, even though the SST had been maintainedabove 29°C. This suggests that the tendency of Chaba’scentral pressure during its decay stage was not correlatedwith the SST but rather was related to TCHP. Figure 11portrays the relationship between SST and central pres-sure and between TCHP and central pressure during thedecay stages of Chaba and Songda. No correlation be-tween SST and central pressure was found (Fig. 11(a)),while the correlation coefficient between TCHP and cen-tral pressure was about 0.47, which is relatively high com-pared to that between SST and central pressure (Fig.11(b)). Both of Songda’s correlation coefficients arehigher than those for Chaba. This is consistent with thesuggestion that the TC intensity of Songda is subject toinfluence by air-sea interaction. In the case of Songda,the correlation coefficient between TCHP and central

pressure in Fig. 11(c) (about 0.96) is higher than that be-tween SST and central pressure in Fig. 11(d) (about 0.76).This reveals that TCHP is more strongly correlated withcentral pressure than SST during the decay stage.

In the case of Songda, SST had been maintainedabove 28°C during the development and mature stages.In contrast, the TCHP varied with high amplitudes. Inparticular, during the rapid intensification stage ofSongda, the TCHP around the area was above 120kJ cm–2. A peak in TCHP was also found during the ma-ture stage of Songda. In contrast, the variations in TCHPwere not always correlated with the underlying SST whereSongda passed by. The intensity of Songda or its intensi-fication was influenced by the variation of TCHP. There-fore, local TCHP is better correlated with the intensityand intensification than the local SST.

6. Relationship between Warm Core Eddies and theIntensification of Typhoons Chaba (2004) andSongda (2004)Oceanic reanalysis data taken by MOVE for the du-

ration of Chaba and Songda was used to explore the rela-

Fig. 11. (a) Relationship between SST and central pressure during the decay stage of typhoon Chaba (2004). (b) As (a) except ofthe relationship between TCHP and central pressure. (c) As (a) except for typhoon Songda (2004). (d) Relationship betweenTCHP and central pressure during the decay stage of typhoon Songda (2004).


tionship between local TCHP variations and the tenden-cies of central pressure in the cases of Chaba and Songda.The five-day mean TCHP field for 21 to 25 August re-veals that Chaba developed over the ocean with highTCHP, above 80 kJ cm–2 (Fig. 12(a)). When Chabareached its peak strength, the typhoon movednorthwestward where the TCHP became relatively low.During its sustenance of MCP at 910 hPa, Chaba passedover two warm core eddies (WCEs). From 26 to 31 Au-gust, Chaba weakened in intensity (Fig. 12(b)) as it movednorthwestward toward Japan. The typhoon made landfallwith a central pressure of 950 hPa although the TCHPwas low where the typhoon passed by. The strong typhoonbrought on a disaster caused by torrential rains, destruc-tive winds, and storm surges.

Songda had moved west-northwestward from 26 to31 August after the passage of Chaba (Fig. 12(b)). TheTCHP became low where Chaba passed by. This decreasein TCHP was probably caused by multiple effects ofupwelling and turbulent mixing. These effects are thesame as those of sea surface cooling (SSC) (Wada, 2005b).Since the TCHP also became low when Songda passedby, it reached its peak strength of MCP of 925 hPa, 15hPa weaker than that of Chaba. Nevertheless, Songdasustained a central pressure of around 925 to 930 hPa from1 to 5 September (Fig. 12(c)). During this period, Songdapassed over two WCEs.

One of the WCEs was located in around 135°E, 21°N(hereafter W1). W1 had been already dominant before thearrival of Chaba. During the passage of Chaba, the TCHPof W1 became lower than before the passage of the ty-phoon. During the passage of Songda, however, the TCHPincreased. The other WCE was located around 128°E,27°N (hereafter W2) along the Kuroshio stream. ThisWCE appeared only from 1 to 5 September when Chabapassed by to the right of WCE W2. Upstream from W2,where there was a source of the Kuroshio stream, anotherWCE was found southeast of Taiwan (hereafter W3). ThisWCE had been continuously salient, even during the pas-sage of Chaba and Songda. WCE W3 may help WCE W2develop or maintain its amplitude through the Kuroshiostream.

Unlike warm eddies in the Gulf of Mexico, those inthe western North Pacific can be separated or propagatedfar from the place of formation. Thus, each evolution ofWCEs W1, W2, and W3 is examined from 1998 to 2004using the reanalysis data taken by MOVE. The areas ofthese WCEs are defined as follows: W1 is located be-tween 130°E and 135°E and between 20°N and 22.5°N,W2 is located between 127.5°E and 130°E and between26°N and 29.5°N, and W3 is located between 121°E and125°E and between 19°N and 22.5°N.

Figure 13 depicts the evolution of TCHP around theWCEs, accompanied by seasonal variations. The evolu-

Fig. 12. Distribution of mean TCHP in the western North Pa-cific by MOVE reanalysis data. (a) 21 to 25 August. (b) 26to 31 August. (c) 1 to 5 September 2004. These panels in-clude the typhoon names and their central pressures. W1,W2 and W3 indicate warm eddies. The domain of W1 cov-ers from 130° to 135°E and 20° to 22.5°N, that of W2 from127.5° to 130°E and 26° to 29.5°N, and that of W3 from121° to 125°E and 19° to 22.5°N. Open “typhoon” symbolsdenote the locations of Chaba (2004). Filled ones denotethose of Songda (2004).


tion of TCHP has high frequencies, especially in W1 andW3. Each TCHP value for W1 (Fig. 13(a)) and W2 (Fig.13(b)) in the 2003 season is the highest, while the valuefor W3 (Fig. 13(c)) in the 2004 season is the highest inany year. The TCHP value for W2 in the 2004 season isrelatively low, partly because moderately strong typhoonsfrequently passed WCE W2 during the 2004 season. Incontrast, the TCHP value for W3 in the 2004 season was

(a)

Chaba

Songda

(b)

Fig. 13. Time series of the TCHP around (a) W1, (b) W2, and(c) W3 from 1998 to 2004. The domains of W1, W2, andW3 are as in Fig. 12.

Fig. 14. (a) SST and ocean current after 300-hour integration,at 1200 UTC 5 September 2004. (b) TMI and AMSRE fu-sion daily SST on 5 September 2004. “Typhoon” symbolswith lines indicate the central positions and tracks of Chaba(2004) and Songda (2004).

extraordinarily high compared to the value for 1998 to2003. The formation mechanism of the extraordinaryWCE W3 in the 2004 season is beyond the scope of thisstudy. However, it may be related to the extraordinarynumber of typhoon landfalls in Japan with only a smallincrease in central pressure from the peak.


7. Effect of the Baroclinic Ocean Response to Ty-phoon Chaba (2004) on Warm Core Eddies andon the Intensity of Typhoon Songda (2004)After the passage of Chaba, the TCHPs of W1 and

W2 (Fig. 12(c)) evidently increased. Each increase inTCHP probably influenced the sustenance of TC inten-sity of the subsequent typhoon Songda. However, theprocess responsible for the TCHP increase has never beenexamined. Our interest was to explore the process ofTCHP increase in the WCEs. In order to investigate theeffect of the baroclinic ocean response to Chaba on theWCEs, especially W1 and W2, numerical simulationswere performed by MRI.COM.

Two specifications for atmospheric forcing were pre-pared for the numerical simulations in the present study.One was determined from the NCEP R-2 wind stress andwind velocity by linear interpolation; the other replacedthe above mentioned wind stress and wind velocity withvalues obtained from JMA/GANAL by linear interpola-tion with an artificial Rankin vortex. The Rankin vortexwas determined from the best-track central position, adistance of the radius of 50 knots wind velocity, and maxi-mum sustained wind velocity. Hereafter, this experimentwill be referred to as JMAR and the other as NCEP. Pro-duction of JMAR atmospheric forcing from the JMA/GANAL and best-track data is described in detail in Wada(2002). The time interval of atmospheric forcing in JMARor NCEP is 10 minutes, corresponding to the time step ofthe numerical simulation from MRI.COM. The initial time

of each numerical simulation is 2400 UTC 23 August2004. Two numerical integrations were performed be-tween the initial time and 2400 UTC 4 September 2004.

Figure 14(a) illustrates the horizontal distribution ofSST simulated by MRI.COM for 1200 UTC 5 September2004 in JMAR. The simulation enables one to reproduceSSC after the passages of Chaba and Songda. Around137°E, 22°N, the SSC caused by the passage of Songdaoverlapped with that caused by the passage of Chaba.Songda sustained its peak intensity where the typhoonpassed over the relatively low SST compared to the SSTbefore the passage of Chaba. Figure 14(b) illustrates thehorizontal distribution of TMI and AMSER fusion dailySST on 5 September 2004. SSC was produced around thetrack of Chaba, especially to the right of the running di-rection. This represents the ocean response to a fast mov-ing storm (Wada, 2005b). SSC was dominant around137°E, 22°N. These results are consistent with results ofJMAR seen in Fig. 14(a).

Figure 15 presents a vertical section of sea tempera-

Fig. 15. Vertical section of sea temperature in JMAR experi-ment, differences between JMAR and NCEP R-2, and dif-ference in ocean current along 22°N between them. “C”indicates cold water and “W” indicates warm water in theJMAR experiment. Chaba (2004) passed by near 137°E,22°N where upwelling is dominant.

Fig. 16. Time series of differences in tropical cyclone heatpotential (kJ cm–2) and depth of 26°C isotherm (m) aroundthe center position of (a) Chaba (2004) and (b) Songda(2004) between JMAR and NCEP R-2. Solid line representstropical cyclone heat potential, and broken line representsdepth of the 26°C isotherm.


ture, the difference in sea temperature between JMAR andNCEP, and the temperature of the ocean current betweenthem along the 22°N line at 2400 UTC 31 August 2004.Around 137°E, SSC by upwelling is dominant. The peakof sea temperature cooling from the conditions just be-fore the passage of Chaba was greater than 6°C at a deptharound 40 m, accompanied by cyclonic circulation. Incontrast, the northwestward or westward near-inertial

current induced by the passage of Chaba produced a con-vergence near the surface outside the sea temperaturecooling region around the typhoon center. Around theconvergence region, the sea temperature in JMAR washigher than that in NCEP around a depth of 80 m. En-trainment induced by wind forcing on the left side of thetrack may have contributed to the increase in near-sur-face sea temperature through downward transportation ofwarm near-surface water into the ocean interior, sincevertical turbulent mixing on the left side of the track wastoo small to entrain cold water by upwelling behind thetyphoon.

Figures 14 and 15 provide no evidence of the TCHPincrease around each typhoon track because the increasein SST is negligibly small compared to the decrease ofSST due to the passage of the typhoons. In order to clarifythe TCHP increase, TCHP was estimated by the formulaof Leipper and Volgenau (1972), using the results of nu-merical simulations in JMAR and NCEP. Figure 16 de-picts a time series of differences in TCHP and the depthof the 26°C isotherm (Z26) between JMAR and NCEP,around the center of Chaba and Songda. For Chaba, theTCHP in JMAR is smaller than that in NCEP, althoughZ26 is deeper in JMAR than it is in NCEP (Fig. 16(a)). Incontrast, the TCHP in JMAR is larger than that in NCEPin the case of Songda, with a relatively deep Z26 (Fig.16(b)).

It is worth noting that the difference between NCEPand JMAR without the Rankin vortex is negligible (notshown). This suggests that the Rankin vortex plays animportant role in producing TCHP variations around thetracks. Figure 17 indicates each evolution of TCHP andZ26 around the WCEs: W1 (Fig. 17(a)), W2 (Fig. 17(b)),and W3 (Fig. 17(c)). The TCHP and Z26 in JMAR washigh around W1 (Fig. 17(a)) and W3 (Fig. 17(c)), partlybecause Typhoon Aere (2004), which was generated be-tween Chaba and Songda, had passed by near the area. Inparticular, a near-inertial variation is found in Fig. 17(a).In fact, the Rankin vortex of Aere is artificially includedin the JMAR atmospheric forcing. In contrast, TCHP andZ26 in NCEP were high around WCE W2 (Fig. 17(b))until 29 August. After that, both TCHP and Z26 increasedgradually in JMAR, accompanied by inertial oscillations.The inertial oscillations found in W2 were probablycaused by near-inertial currents induced by the passageof Songda.

8. Discussion

8.1 Relationship between accumulated tropical cycloneheat potential and tropical cyclone intensityWe have demonstrated in the previous sections that

the potential intensity derived from ATCHP was morereliable than that derived from ASST or the duration of a

Fig. 17. Time series of TCHP and depth of the 26°C isothermaround (a) W1, (b) W2, and (c) W3 from 24 August to 5September in 2004. “NCEP” and “JMAR” represent kindsof atmospheric forcing. Domains of W1, W2, and W3 areas in Fig. 12.


TC. This was also reported in Wada (2006), using anonhydrostatic model coupled with a slab ocean model.The SSC caused by the passage of a TC, accompanied bymixed layer deepening by entrainment, leads to a decreasein TCHP around the TC center. This is because the TCHPvariations are accompanied by variations in mixed-layerthickness as well as in the sea temperature above 26°C.In contrast, differences in central pressure between thecoupled and non-coupled experiments reported in Wada(2006) are correlated not with TCHP but with ATCHP.This suggests that increases in central pressure simulatedby the atmosphere-ocean coupled model are closely re-lated not to the SSC around the TC center but to the re-duction of ATCHP where TCs have been experienced.

In an observational study, Cione and Uhlhorn (2003)reported that the energy available to the TC was an orderof magnitude greater than the energy extracted by the TC.A numerical study by Chan et al. (2001) found that theoriginal peak intensity of a TC did not weaken after leav-ing a warm core eddy, even though the TC attained itspeak at the center of the warm core eddy. These results,along with those of the present study, reveal that a TCcarries a kind of reservoir of heat from the ocean whichis closely related to its intensity. This reservoir is not re-lated to the TCHP at a particular place but rather to ATCHPduring the overall duration. Local SSC caused by the pas-sage of a TC plays a role in inhibiting any ATCHP in-crease.

ATCHP can influence TC intensity through transpor-tation of water-vapor flux (or latent heat flux) from theocean to the atmosphere by eddy diffusion in the plan-etary boundary layer. The transportation of water-vaporflux is generally estimated under the assumption that theatmosphere at the sea surface is saturated, in equilibriumwith the SST. The exchange coefficients of the momen-tum and enthalpy fluxes are calculated in the surface-boundary scheme of the atmospheric model. Accordingto Emanuel (1995) and Schade and Emanuel (1999), thesecoefficients, especially for moisture flux, play an essen-tial role in determining TC intensity.

In order for the moisture flux to affect TC intensity,the moisture supplied to the atmospheric boundary layerneeds to be transported to the warm core of a TC throughsecondary circulation around the eyewall. During trans-portation to the warm inner core by neutral moist ascent,the moisture flux continues to be transported from theocean to the atmosphere, even when the TC becomesweak. Therefore, a TC can sustain its intensity by thecontinuous supply of moisture flux. The above-mentionedprocess occurs in a cumulative way so that the ATCHPbut not the TCHP is important to TC intensity. The proc-ess is also linked to the effects of deep-layer inflow onintensification described by Ooyama (1982).

The ATCHP becomes higher when the duration is

longer. As for the duration of TCs, the eastward shift ofTC-genesis locations leads to an elongation of the TCtrack over warm SST regions in El Niño years (Camargoand Sobel, 2005). The study of Camargo and Sobel (2005)shows that ENSO indices are positively correlated withaccumulated cyclone energy (ACE) in the western NorthPacific. In contrast, since 1995, the ACE indices for allbut two Atlantic hurricane seasons have been above nor-mal; the exceptions are the El Niño years of 1997 and2002 (Trenberth, 2005). The concept of ACE is similar tothat of ATCHP from the standpoint of the accumulationof energy (kinetic or thermal). Therefore, the concept ofaccumulation may be more important in discussing TCintensity and intensification than the concept of localiza-tion.

In order to qualitatively evaluate the effect of ATCHPincreases due to the passage of TCs on their central pres-sures, we estimated the ATCHP increase in 2004 aroundWCEs W1 and W2. The difference in ATCHP betweenNCEP and JMAR was 0.6 MJ day–1 for W1, and 0.22MJ day–1 for W2, corresponding to intensifications of ap-proximately 12 hPa day–1 and 4.5 hPa day–1 determinedin the case of JMAR using the relation shown in Fig. 4(b).In fact, other atmospheric environmental conditions, suchas vertical shear or an upper trough, may affect the inten-sification of TCs, as suggested in the previous section. Inany case, the ATCHP contributes to sustaining TC inten-sity.

A comparison between 2000 and 2004 reveals thatrapid intensification areas are confined to places wherethe divergence in the upper troposphere is relatively small,although the TCHP was relatively high on average aroundthe basin in both years. This shows that rapid TC intensi-fication is controlled by the atmospheric environment ora related TC track rather than the average TCHP distribu-tion. Moreover, the pattern in the upper troposphere maybe related to the tracks typically making landfall in Ja-pan, while the average TCHP distribution may be relatedto TC intensity through the accumulation of TCHP along-side TC tracks.

8.2 Formation mechanism of warm eddies in the west-ern North PacificThe ATCHP increase caused by the ocean response

to a TC is explored here. Lin et al. (2005) discussed theimpact of warm core eddies on the intensity of TyphoonMaemi (2003). However, they did not mention the im-pact of TCs on variations in WCEs, nor the impact ofthose WCEs on subsequent TCs passing through the areaafter the passage of a previous TC. Bender and Ginis(2000) reported that the intensity of a subsequent hurri-cane did not intensify but rather was suppressed due tothe passage over the SSC area of a previous hurricane.According to the observational research of Pudov et al.


(1978) and Fedorov et al. (1979), the sea temperaturearound the typhoon center was reduced by upwelling,while areas outside the reduction area were warming, es-pecially around a depth of 50 m. A schematic diagramfrom Ginis (1995) indicates that a warming region for-ward and outside of the radius of deformation of a TCwas brought about not only by entrainment but also by asurrounding convergence caused by divergent flow in-duced by cyclonic wind stress near the TC center. Thiscauses downward transportation of warm sea water fromthe sea surface to the ocean interior. On the left side ofthe TC track, an anticyclonical circulation is formed orenhanced by near-inertial oscillation. The anticyclonicalcirculation is favorable for sustaining the amplitude ofWCEs. It is worth noting that the sea temperature warm-ing is not remarkable at the sea surface but is significantaround a depth of 50 m, corresponding to the mixed layerbase. This is one reason why TCHP increases significantlyin areas surrounding the TC.

Unlike the TCHP around WCE W1, the TCHP aroundWCE W2 increases in NCEP. The following processesare believed to influence TCHP variations around WCEW2: the baroclinic response to Typhoon Aere (2004), thedevelopment of WCE W3, and advection from WCE W3toward WCE W2. Initially, SSC caused by the passage ofAere led to a decrease in TCHP around WCE W2 inJMAR. Subsequently, Typhoon Songda (2004) ap-proached WCE W2. Near-inertial currents caused by thepassage of Songda led to an increase in TCHP ahead ofthe direction of TC translation. In addition, warm wateris transported from WCE W3 through the Kuroshio. Theincrease in TCHP has become significant since 1 Sep-tember 2004.

Songda passed by to the left of the track of TyphoonChaba (2004). This is similar to another case involvingtyphoons in October and November 2004: TyphoonsMa-on (2004), Tokage (2004), and Nock-ten (2004), seenin Fig. 18. Ma-on experienced rapid intensification around130°E, 23°N. After the passage of Ma-on, Tokage passedby on the left side of Ma-on’s track. Tokage was weakerthan Ma-on. However, Tokage sustained its intensity whilemoving westward. Moreover, Nock-ten passed by to theleft of Tokage’s track, sustaining its peak intensity. Thistrack pattern was notably found around 120 to 135°E, 16to 23°N, denoted by a cloud symbol in Fig. 18. SSC wasfrequently found here in 2004 due to the frequent pas-sage of typhoons (Wada, 2005a).

According to Riehl (1972), TCs are not at their peakintensity when they begin to recurve with a relatively slowspeed of TC translation. Around the recurvature area, sig-nificant SSC results from the slow translation. The SSCcaused by the moving TC appears along the track as anarrow band of cold water in the western North Pacific.

In the particular case of a TC moving northwestward,

the TCHP can increase south (decrease north) of the trackdue to transportation of warm water by near-inertial cur-rents (by upwelling and vertical mixing). Previous stud-ies of the ocean response to storms have focused on theSST cooling rate, which was dominant in a slow transla-tion (Price, 1981; Wada, 2002). In the case of Chaba,however, the typhoon re-developed during a slow trans-lation when the typhoon passed over warm eddy W1. Af-ter the passage of Chaba, SSC occurred alongside thetrack. The pattern of SSC was similar to previousmodeling studies (e.g. Bender and Ginis, 2000). It is worthnoting that it is difficult to define a threshold betweenslow and fast translation in the present study. This diffi-culty is because the best-track TC position has been re-corded every six hours, except around Japan, where rapidTC intensification rarely occurs. Moreover, the calcula-tion-period of ATCHP does not involve the mature stagewhen the SST cooling rate is the greatest with a slow trans-lation. If we were to separate all the cases into two cat-egories; for example, fast-moving TCs and slow-movingones, the ATCHP-TC intensity relation would be modi-fied. However, the other question arises in the aforemen-tioned discussion: whether or not a TC can continue tointensify with a slow translation? This is beyond the scopeof this study.

Nock-ten

Nock-ten

Tokage

TokageMa-on

Ma-on

Fig. 18. Lines depicting the tracks of typhoons Ma-on (2004),Tokage (2004), and Nock-ten (2004). “Typhoon” symbolsindicate central position of each typhoon. Size of the ty-phoon symbol represents intensity of each typhoon. Cen-tral positions with central pressures above 960 hPa areshown grey, while those with central pressures below 960hPa are shown black. “Cloud” region represents a suste-nance region for typhoon intensity.


The combination of WCE and SSC leads to a sharpgradient SST alongside the track. According to the nu-merical study of Chan et al. (2001), the intensificationprocess over a sharp SST gradient is similar to that overa warm core eddy. The present study demonstrates thatthe formation of a sharp SST gradient is equivalent to theformation of anticyclonical circulation on the left side ofthe TC track, which is consistent with the result of Chanet al. (2001).

However, local SSC is not responsible for the subse-quent movement of TCs toward warm regions in the lowlatitudes. Previous studies have reported that the behaviorof TCs is hardly affected by SSC in general; so manyfactors could influence the behavior that it would be dif-ficult to isolate the effect of SST. From the standpoint ofthe effect of SST distribution of TC translation, Changand Madala (1980) reported in their idealized numericalexperiments that the behavior of TCs was influenced bySST distribution. In their result, TCs tended to move intoregions of warmer SST, translating the mean flow overthe ocean downwind, with SST gradients perpendicularto the mean ambient flow vector. In the present study, theSST gradient had been produced after the passage of aprevious TC (Chaba or Ma-on in this case). Futhermore,the subsequent TC (Songda or Tokage in this case) movedinto a high-TCHP region (not shown). Even though thereport of Chang and Madala (1980) is for a case that isnot terribly realistic, it may be necessary to reexaminethe relationship between subsequent TC translation andSST or TCHP distribution produced as a result of the pas-sage of a previous TC.

These TC track patterns have not been typical in re-cent decades. Interestingly, the above mentioned TC trackpattern has occurred only nine times in the western NorthPacific since 1981, as seen in Table 2. It is worth notingthat this track pattern has increased in recent years. Table2 also indicates that this pattern often appears in El Niño

or El Niño-like seasons. The baroclinic ocean responseto TC intensity may not always have a positive impact.Indeed, the baroclinic ocean response in the case of a“cross” track pattern is one of the factors responsible forintensification or sustenance of peak TC intensity in ElNiño and El Niño-like seasons.

9. ConclusionWe employed tropical cyclone heat potential (TCHP)

in the western North Pacific from 1998 to 2004 as esti-mated from oceanic reanalysis data to explore the rela-tionship between TCHP and tropical cyclone (TC) inten-sity, as well as that between the TRMM/TMI three-daymean sea surface temperature (SST) and TC intensity. Aregressive empirical function derived from the TCHP wasmore reliable for determining the maximum potential in-tensity of TCs than empirical functions derived from SST.The ATCHP, an accumulation of TCHP around a TC centerevery six hours from the formation to the mature stage, isbetter correlated with the intensity in the mature stagethan is the ASST, an accumulation of SST around thecenter every six hours, along with the duration.

Rapid intensification of a TC was found to occurwhen TCHP was relatively high in the western North Pa-cific from 1998 to 2004. Rapid TC intensification wasfrequently observed in 2000 and 2004. Atmospheric andoceanic environments, including the relative humidity andwind in the lower troposphere, and the vertical shear be-tween levels of 200 hPa and 850 hPa, occurring between2000 and 2004, all are different. Rapid intensification isconfined to areas where the divergence in the upper tropo-sphere is relatively small. The pattern in the upper tropo-sphere may be related to the tracks typically making land-fall in Japan. Rapid intensification is controlled both byTCHP on a basin scale and atmospheric environment.

In the cases of Typhoons Chaba (2004) and Songda(2004), the intensity and rapid intensification were re-

Table 2. List of typhoons having minimum central pressure within a week. The second typhoon moved to the right side of the firsttyphoon, then later moved to the left side.

Year Month El First typhoon MCP of the first typhoon(hPa)

Second typhoon MCP of the second typhoon(hPa)

2004 Aug.−Sep. Chaba 910 Songda 925

October Tokage 940 Nockten 9452002 Jun.−Jul. E Chataan 930 Halong 945

Aug.−Sep. E Rusa 950 Sinlaku 950

1998 October Zeb 900 Babs 9401993 December E Lola 955 Manny 9551992 November E Gay 900 Hunt 9401987 August E Betty 890 Cary 9601981 November Hazen 955 Irma 905


lated not to SST variations but to variations in TCHP.However, the potential intensity derived from ATCHP inthe cases of Chaba and Songda was weaker than the best-track TC intensity. At SSTs above 29°C, the tendenciesof TC intensity agree well with those of TCHP. The TCHPvariations were brought about by the baroclinic oceanresponse to Chaba. This implies that the atmospheric en-vironment may also contribute to rapid intensification ir-respective of TCHP. The SSC due to the passage of Chabaled to a suppression of its intensity, while near-inertialcurrents produce a surrounding convergence followed bydownward transportation of warm water into the oceaninterior.

Recently, Oey et al. (2006) reported Loop Currentwarming by Hurricane Wilma (2005)-induced conver-gences in the northwestern Caribbean Sea. However, nei-ther the process of oceanic variations in warm eddies norits influences on TC intensity in the western North Pa-cific have ever been investigated. The present study sug-gests that TC intensity tends to be sustained, not sup-pressed, due to the increase in TCHP when a TC moveson the left side of the previous TC track within a week.This process may be seen in Songda, Typhoons Tokage(2004) and Nock-ten (2004).

These TC track patterns have not always been foundin recent decades. Interestingly, only the nine patternshave occurred since 1981 in the western North Pacific.The track pattern frequently appears in El Niño or El Niño-like seasons, particularly since 2000. In the case of the“cross” track pattern, the baroclinic ocean response is oneof the factors responsible for the intensification and suste-nance of TC intensity, but only in El Niño and El Niño-like seasons, through the enhancement of TCHP result-ing from the baroclinic ocean response. In the presentstudy, we discovered the significance of ATCHP for TCintensity as well as one of the processes causing TCHP toincrease before a TC passed by. However, we do not knowhow or why “cross” tracks occur only in El Niño and ElNiño-like seasons. Thus, this has now become an issuefor further investigation.

AcknowledgementsThe authors are grateful to Dr. Hideyuki Nakano for

providing an ocean general circulation model and usefulcomments about the model. The authors express their sin-cere thanks to three anonymous reviewers for their ben-eficial comments. The numerical experiments were per-formed on the supercomputer (NEC SX-6) at the Mete-orological Research Institute. Both the General MappingTools (GMT) and the Grid Analysis and Display System(GrADS) were used to draw pictures. Part of the workwas supported for the first author by funds from a coop-erative program (No. 126) provided by the Ocean Re-search Institute, the University of Tokyo.

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