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Meso-␤-scale Pressure Dips Associated with Typhoons
HIRONORI FUDEYASU* AND SATOSHI IIZUKANational Research Institute for Earth Science and Disaster Prevention, Tsukuba, JapanTAIICHI HAYASHI
Disaster Prevention Research Institute, Kyoto University, Kyoto, Japan(Manuscript received 25 May 2005, in final form 19 June 2006)
ABSTRACT
Using a mesoscale model, the formation process of a pressure dip, which was characterized by a rapid
decrease and subsequent increase in surface pressure lasting less than an hour, was investigated. A simulated pressure dip accompanied by Typhoon Zeb was closely related to warm potential temperature anomalies in the lower troposphere. As the typhoon moved into the midlatitude westerlies, the inflow of a dry air
mass into the moist region of the typhoon occurred to the west of the typhoon center. Then, a downdraft
developed due to evaporation and sublimation. Below the melting levels where there were fewer hydrometeors, however, evaporation was insufficient to offset the adiabatic warming. As a result, warm potential
temperature anomalies were created in the lower level, resulting in the formation of a pressure dip. The
features of a pressure dip associated with other typhoons observed over Japan were also examined. The
features were summarized as follows: 1) the pressure dip was observed only on the western side of thetyphoon center at a distance of 50–300 km from the center, 2) the pressure dip was accompanied by a sudden
cessation of rainfall, and 3) all typhoons with a pressure dip were found during boreal autumn whenlarge-scale environmental conditions were characterized by westerlies with a trough and a dry air mass to
the west of Japan at upper levels and fronts at lower levels. The formation process of the simulated pressure
dip in other typhoons was found to be similar to that of Typhoon Zeb. The present study suggests thatpressure dip is an inherent feature of the asymmetric structure of a typhoon undergoing transition to a
extratropical cyclone.

1. Introduction
Fujita (1952) first described the phenomenon of apressure dip that is occasionally observed in association
with typhoons crossing the Japanese islands or moving
along the southern coast of Japan (e.g., Nakajima et al.
1980; Fujii 1992; Maeda 1994; Fudeyasu and Tsukamoto 2000). The occurrence of a pressure dip is generally only recognized by the analysis of barographic
data. Figure 1 shows the time series data of surface
wind direction and speed, temperature, dewpoint temperature, and pressure recorded at station K (see Fig. 3)* Current affiliation: Japan Agency for Marine-Earth Scienceand Technology (JAMSTEC), Yokosuka, Japan.

Corresponding author address: H. Fudeyasu, Japan Agency for
Marine-Earth Science and Technology (JAMSTEC), 2-15 Natsushima, Yokosuka, Kanagawa 237-0061, Japan.
E-mail: fudeyasuh@jamstec.go.jp
DOI: 10.1175/MWR3337.1© 2007 American Meteorological Society

MWR3337

during the passage of Typhoon Zeb. A pressure dip,
consisting of a rapid decrease and subsequent increasein surface pressure, is distinct from the gradual decrease in surface pressure that is recorded during thepassage of the typhoon center. Previous studies haveshown that pressure dips are meso-␤-scale phenomena
with a bandlike structure and typical dimensions of
100–200 km by 10–50 km. The maximum observed amplitude of pressure dips is 7–9 hPa, and the duration
recorded at a single station is always less than 1 h.
Pressure dips are accompanied by strong winds, occasionally causing severe damage at various locations
over the Japanese islands. For example, it was reported
that extensive damage resulting from instantaneouswind speeds of 54.3 m s⫺1 occurred during a pressure
dip associated with Typhoon Mireille (Maeda 1994;
Fujita 1992).
The formation mechanism of pressure dips is poorly
understood. Matsumoto and Okamura (1985) used
Doppler radar data and surface observations to study a

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FIG. 1. The record of the surface pressure, temperature, dewpoint temperature, and wind speed and direction as observed atstation K from 1100 to 1800 UTC 17 Oct 1998. See Fig. 3 for thelocation of station K.

pressure dip observed over Japan during the passage ofTyphoon Gay and concluded that the pressure dip wasdue to the passage of an internal gravity wave excitedby the convection of the typhoon. Tsujimura (1993)
concluded rather that the pressure dip associated with
Gay was due to a solitary internal gravity wave. On theother hand, Inoue et al. (1999) argued that the pressure
dip observed within Typhoon Zeb was a low pressure
region that formed behind the gust front of a gravity
current caused by an intense rainband. Thus, it is apparent that the formation mechanism of pressure dips isa controversial issue.
Another controversial issue is the genesis location ofpressure dips with respect to the center of the accompanying typhoons. Fujita (1952) described pressure dipson both sides of the movement direction of Typhoon
Della at a distance of 50–600 km from the typhoon
center. Nakajima et al. (1980) reported pressure dips
both ahead of and behind the advancing Typhoon
Owen at a distance of 50–100 km from its center. In thiscase, the pressure dips ahead of the typhoon were short
lived and of lesser amplitude than those behind thetyphoon. The pressure dips associated with Typhoon
Gay and Typhoon Zeb were observed to the left andrear of the northeast-moving typhoons (Matsumoto

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and Okamura 1985; Inoue et al. 1999; Fudeyasu andTsukamoto 2000). The pressure dips that accompanied
Typhoon Mireille developed on the rear of the typhoon
at a distance of about 200 km from its center (Fujita
1992; Fujii 1992; Maeda 1994).
In this study, we first examined the formation mechanism of pressure dips using a high-resolution numericalmodel. We focused on the pressure dips associated with
Typhoon Zeb because its track was the most suitablefor investigating the structure and occurrence of pressure dips because of the availability of comprehensive
observational data. The model successfully reproduced
the major features of the pressure dip associated with
Zeb. We then attempted to clarify the formation
mechanism of pressure dips. Another aim of this study
was to examine the conditions necessary for the occurrence of pressure dips using the available observational
data.
This paper is organized into the following parts. Section 2 describes the observational data and the model
used in the study while section 3 compares the simulated pressure dips associated with Zeb with observational data. Additionally, the internal structure of apressure dip is described. Section 4 examines the formation mechanism of pressure dips. Section 5 presents
the features found in observed and simulated pressure
dips associated with other typhoons. Section 6 discusses
the conditions needed for pressure dip formation while
section 7 summarizes the current study.

2. Data and modela. Data
To identify past occurrences of pressure dips, we analyzed the temporal variations of surface pressure as recorded by the barographs at meteorological stations of
the Japan Meteorological Agency (JMA) during theperiod from 1980 to 1998. The stations are located at
approximately 50-km intervals throughout the Japanese islands. We used the following quantitative definitions to distinguish a pressure dip from other features, such as local changes in pressure.
1) The maximum amplitude of the dip in surface pressure observed at each station had to be greater than
1 hPa and had to last for less than 1 h.
2) The pressure dip, as defined in the first definition
above, had to be associated with an individual typhoon, and had to be recorded by at least 10 meteorological stations.
3) The movement of the pressure dip, as detected byconditions detailed in the first definition above, had
to be recognized in tandem with the movement of
the accompanying typhoon.

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These quantitative definitions are based on the characteristics of a pressure dip described by Fujita (1952),
who stated that a pressure dip is a mesoscale atmospheric phenomenon with a bandlike structure and thatthe propagation of the pressure dip must be recognized
by reference to barographs. The second definition detailed above is required in order to depict the spatial
structure of pressure dips. Therefore, in this study, weonly considered those pressure dips with significant amplitude and long persistence that were detected over
the Japanese islands.
Tracks of typhoons were obtained from the besttrack archives of the Regional Specialized Meteorological Centers (RSMC) Tokyo-Typhoon Center. The
dataset consisted of the names, positions, maximum
surface pressures, and maximum wind speeds of typhoons, generally recorded at 6-h intervals. Estimatesof the hourly positions of typhoon centers were determined by linear interpolation based on the 6-h data.
Hourly rainfall patterns were estimated using the Radar-Automated Meteorological Data Acquisition System (AMeDAS) data provided by the JMA. The RadarAMeDAS data comprise an hourly averaged rainfall
analysis created from a composite of observations fromoperational precipitation radars and the AMeDAS data
(Obayashi 1991). The Radar-AMeDAS data during theperiod from April 1988 to March 2001 have a horizontal
resolution of 5 km.
Details of the large-scale atmospheric conditions
around the typhoons with pressure dips were reconstructed from the reanalysis data provided by the National Centers for Environmental Prediction–National
Center for Atmospheric Research (NCEP–NCAR).
The data had been collected every 6 h between 1980
and 1998 with a horizontal resolution of 2.5° (Kalnay etal. 1996). Equivalent black body temperature (TBB)
from the Japanese Geostationary Meteorological Satellite spanning the period of 1980–98 was also used.

b. ModelA numerical simulation was performed using thefifth-generation Pennsylvania State University–NCAR
Mesoscale Model (MM5), version 3.5. The MM5 is anonhydrostatic model developed as a community mesoscale model (Dudhia 1993; Grell et al. 1995). The
model incorporates multiple nested grids and includes
both explicit moist physical processes and cumulus parameterization schemes. This model solves the nonlinear, primitive equations using Cartesian coordinates in
the horizontal direction and terrain-following sigma coordinates in the vertical direction. Pressures at thesigma levels were determined from a reference statethat was estimated using the hydrostatic equation for

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FIG. 2. Model domains of the numerical simulations and terrain.
Regions higher than 400 m are shaded. Track (line) and 3-hourly
locations (closed circles) of the simulated typhoon derived fromthe D2 simulation during the period from 0000 UTC 17 Oct to0000 UTC 18 Oct 1998.

any given sea level pressure and temperature based ona standard lapse rate.
In this study, the model contained 30 sigma levels offine resolution near the surface, while the top of themodel was at 50 hPa. The simulation was conductedusing the two domains shown in Fig. 2. The outer domain (D1) had 271 ⫻ 271 grid points, centered at 33°N,
133°E, with a horizontal resolution of 15 km. The innerdomain (D2) had 460 ⫻ 460 points with a horizontal
resolution of 5 km. The outer domain was designed to
produce the initial and boundary conditions for D2,
while D2 was designed to explicitly resolve the finestructures of typhoons and associated pressure dips. We
also conducted a simulation using a horizontal resolution of 3 km, but the results were essentially the same as
those from the model with a horizontal resolution of
5 km.
A simple ice scheme (Dudhia 1989) was used for the
cloud microphysics. In this scheme, cloud water andrainwater below 0°C are treated as ice and snow, respectively. The convective scheme of Grell (1993) wasused as a cumulus parameterization only in D1. The
model also included the planetary boundary layer(Hong and Pan 1996), a five-layer soil model, and a
cloud–radiation interaction scheme (Dudhia 1989,
1993; Grell et al. 1995).

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TABLE 1. Summary of initial time and integrated time for simulations.
Domain 0

Domain 1

Domain 2

TyphoonInitial time

Integrated
time (h)Initial time

Integrated
time (h)Initial time

Integrated
time (h)Gay
Wayne
HattieMireille
Oliwa
Vicki

0000 UTC 21 Oct 1981
1800 UTC 17 Oct 1989
0000 UTC 6 Oct 1990
0600 UTC 26 Sep 1991
1200 UTC 14 Sep 1997
0000 UTC 21 Sep 1998

48
72
72
48
72
48

1800 UTC 21 Oct 1981
1200 UTC 18 Oct 1989
1800 UTC 6 Oct 1990
0000 UTC 27 Sep 1991
0300 UTC 15 Sep 1997
1500 UTC 21 Sep 1998

30
54
54
30
57
33

0000 UTC 22 Oct 1981
0000 UTC 19 Oct 1989
0000 UTC 7 Oct 1990
0000 UTC 27 Sep 1991
0600 UTC 15 Sep 1997
1500 UTC 21 Sep 1998

24
48
48
30
54
33

The initial and lateral boundary conditions for D1 inthe simulation of Typhoon Zeb were derived from the
Global Asian Meteorological Experiment (GAME) reanalysis dataset produced by the Meteorological Research Institute–JMA (see details online at http://gainhub.mri-jma.go.jp/GAME_reanal.html). The dataset
includes the 3D analyzed atmospheric fields over the
Asian and Pacific regions (80°N–30°S, 30°E–180°), covering the period from April to October 1998. The data
have a horizontal resolution of 0.5° for both longitude
and latitude, with 17 vertical levels and a 6-h time interval. Sea surface temperature is represented by theskin temperature of the NCEP–NCAR reanalysis
dataset (Kalnay et al. 1996).
The simulation of D1 was performed for 42 h, starting at 1800 UTC 16 October 1998, after the model was
integrated for 18 h during an initial spinup at 0000 UTC
16 October. During the initial spinup, continuous dynamical assimilation was performed by adding forcing
functions to the governing model equations to gradually “nudge” the model state toward the reanalysis data
for wind-only components (Grell et al. 1995). The relaxation boundary conditions were adopted as the lateral boundary conditions (Anthes et al. 1987) in which
the model-predicted values were relaxed to those estimated from the large-scale analysis data. The D2 simulation began at 0000 UTC 17 October 1998 and was
performed for 36 h using the initial and boundary conditions interpolated from the results of the D1 simulation. At the start of the D2 simulation, the center of
Typhoon Zeb was located to the southwest of Kyushu,
Japan (Fig. 2).
We also used the NCEP–NCAR reanalysis dataset as
the initial and lateral boundary conditions in order tosimulate other typhoons. This dataset was also used to
examine the sensitivity of the conditions used for simulating Zeb. The horizontal resolution of the datasets iscoarse in comparison with the GAME reanalysis
dataset, and consequently another outer domain (D0)was incorporated into the model in order to nest thesimulation of D1. The mesh of D0 had 120 ⫻ 120 points

and a horizontal resolution of 45 km. For the coarse
domain D0, the Kuo scheme (Grell et al. 1995) wasused as the cumulus convection parameterization, but
all other configurations of the model are the same as
those used for the D1 simulation. The simulation of D0
was performed using the continuous dynamical assimilation with the typhoon bogus schemes incorporated inMM5. The initial and integration times of all simulations conducted in this study are summarized in Table
1. In the following analysis, we only use the resultsof D2.

3. The pressure dip associated with Typhoon Zeb
a. Features of the observed pressure dip
Typhoon Zeb was upgraded to a tropical storm(maximum 1-min sustained wind speeds in excess of
17.2 m s⫺1 or 35 kt) from a tropical depression on 11
October 1998 when it was situated at 10°N, 140°E (Fig.
3). The typhoon turned toward the northwest over thevicinity of the Philippine Sea and increased in intensity
to that of a super typhoon (maximum 1-min sustained
wind speed in excess of 70 m s⫺1 or 135 kt). The typhoon made landfall in Kyushu at 0700 UTC 17 October, with a central pressure of 975 hPa and a maximum
wind speed of 25 m s⫺1. The typhoon then passed over
Shikoku and Honshu, the main island of the Japaneseislands, at 1400 UTC 17 October. After reaching the
Sea of Japan, Zeb moved north-northeast and eventually weakened. It was downgraded to an extratropical
cyclone when it was located near Hokkaido at 0000
UTC 18 October.
During the passage of Zeb through the Japanese islands, a considerable decrease and subsequent increasein surface pressure was recorded at several meteorological stations. Figure 4 shows the time series data ofsurface pressure recorded by barographs at these meteorological stations. In addition to a gradual change insurface pressure related to the passage of Zeb, a sudden
3–8-hPa drop in surface pressure was also discernible.
The pressure dips appeared a few hours after the pas-

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FIG. 3. Hourly locations of pressure dip axis (solid lines) for Typhoon Zeb (black circles).
Open circles indicate stations where the pressure dip was observed, whereas crosses indicate
stations where the pressure dip was not observed.

sage of the typhoon center for all stations except station
J, where a pressure dip was recorded prior to the arrivalof the typhoon center. It should be noted that the pressure dip existed only to the northwest of Zeb as it
tracked to the northeast (Fig. 3). There is no sign of apressure dip at stations I and L located to the southeastof the typhoon. A pressure dip was also not observedon the northern part of the Kyushu and Chugoku districts during the period from 0800 to 1200 UTC, indicating that the extent of the pressure dip axis was defined by the isopleths of minimum pressure within thepressure dip. Although most meteorological stations recorded the pressure dips after the passage of the typhoon center, a number of stations far from the typhoon center recorded a pressure dip in advance of thetyphoon center during the period from 1300 to 1500
UTC. Station A, located to the northwest of the typhoon track (Fig. 3), recorded no pressure dip. We interpreted this as an indication that the pressure dip axis
did not extend to lower latitudes.
To depict the mesoscale structure of the pressure
dips, we conducted a time-to-space conversion (TSC)
analysis (e.g., Koch and O’Handley 1997). Figure 5 presents the spatial structure of the surface pressure derived from the TSC analysis of surface pressure record-

ings at 28 meteorological stations within a distance of
200 km from the typhoon center between 1400 and 1500
UTC 17 October (see Fig. 3). The surface pressure had
an asymmetric distribution. A low pressure region of
100-km length and 20–30-km width extended toward
the north-northwest from a point 50 km west of thetyphoon center. An area of high pressure to the east of
the low-pressure region represented the rise in surfacepressure recorded in the time series data immediately
prior to the onset of the pressure dip (Fig. 4).
The arrival of the pressure dip was accompanied by asudden change in surface winds (Fig. 1), although thesewind changes were not observed at all meteorologicalstations because of local effects, such as mountainous
terrain. Wind speeds weakened after the passage of thetyphoon center, but strong winds were observed immediately prior to the arrival of the pressure dip. Station K
was located to the left-hand side of the moving direction of Zeb and would therefore be expected to record
an anticlockwise change in wind direction with the passing of the typhoon. The wind direction, however, recorded a clockwise change for the 30 min prior to thearrival of the pressure dip.
The approach of the pressure dip was also accompanied by rapid changes in surface temperature and dew-

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FIG. 4. Barographic charts recorded at meteorological stations during the passage of Typhoon Zeb from 16 to 17 Oct 1998. Station locations are indicated in Fig. 3. The detection ofa pressure dip at each station is indicated by a dashed circle.

point temperature, which fell 3°C one hour before thepressure dips appeared (Fig. 1). These trends were only
observed at stations near the track of the typhoon.
Temperature changes are related to surface fronts associated with Zeb, rather than the pressure dip itself(see Fig. 22). In fact, Fujita (1952) documented no significant change in surface temperature associated witha pressure dip. However, a few hours before a pressure
dip appeared, a sudden decrease in surface temperatureand dewpoint temperature was recorded at some meteorological stations (see also Fig. 1 of Fujita 1952, Fig.
4 of Matsumoto and Okamura 1985, Fig. 3 of Fujii 1992,
and Fig. 2 of Maeda 1994). As stated above, thesechanges in temperature appear to be related to fronts
associated with the typhoons rather than the pressure
dip. This topic will be discussed later in more detail.
Figure 6 shows the hourly rainfall amount derived
from the Radar-AMeDAS data at 0700, 0900, 1300, and1500 UTC 17 October 1998. Intense rainfall precedes
the arrival of the pressure dip axis. The rainfall was
weak at the center of the pressure dip and stopped withthe passing of the pressure dip axis. The AMeDAS
data, which contain a record of surface wind, temperature, and rainfall averaged over 10-min intervals since
1995, show that rainfall changes occurred abruptly overa period of less than 30 min (not shown).

b. Validation of simulated pressure dip
In this section, we validate the features of the Typhoon Zeb simulation. The track of Zeb simulated in
D2 with the finest resolution, determined from the areaof minimum sea level pressure of the simulated typhoon, is shown in Fig. 2. The simulated typhoon trackis generally in good agreement with the observed typhoon track (Fig. 3). The central pressure of the simulated typhoon are about 980 hPa on 17 October when
the typhoon passed over Kyushu, Shikoku, and Honshu
before reaching the Sea of Japan. The recorded pressures within the center of Zeb at this time were in therange 975–980 hPa. The track of the simulated typhoon,
however, is delayed by an hour in comparison with theactual movement of Zeb. This is because the initial

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FIG. 5. Sea level pressure field for Typhoon Zeb during 1400–1500 UTC derived by a TSC
analysis. Contour interval is 2 hPa. The cross represents the position of the typhoon center.

position of D2 shifted westward because of the insufficient simulation of D1, which had a coarse resolution.
This discrepancy may be related to a bias in the boundary conditions or in the assimilation used in the present
model. Although the arrival time of the simulated typhoon differs from actual observations, the track andintensity of Typhoon Zeb are well reproduced by themodel.
Next, we compared the simulated pressure dips withthose observed within Zeb. Figure 7 presents the timeseries data of the simulated sea level pressure at 3-minintervals at each station (see Fig. 3). Portions of thetime series data at stations A and B are not shown
because the typhoon center had passed over these stations prior to the initial time of the D2 simulation at
0000 UTC 17 October (Fig. 2). Of note, the locations ofstations H, I, J, K, L, M, and N have been shifted 1°
westward from their actual locations to accommodate
differences between the simulation and observational
data for the period after Zeb made landfall over Shikoku.
The time series data of simulated sea level pressurerecord a diplike decrease in sea level pressure associated with the passage of the typhoon center, although
their amplitudes are generally smaller than the observed amplitudes (Fig. 4). The peak of the pressure dip

was recorded at station C at 0900 UTC, station D at
0955 UTC, station E after 1050 UTC, station F at 1115
UTC, and station G at 1355 UTC. The pressure dips
generated within the simulation occur about an hour
after the observed pressure dips were recorded at eachstation. This time discrepancy occurs because the trackof the simulated typhoon was 1 h late compared withthe actual track of Zeb (Figs. 2 and 3). A pressure dipof 4.5 hPa was recorded at station H at 1325 UTC, andpressure dips were also recorded at station J at 1445
UTC and station K at 1515 UTC. These times are comparable to the timing of the actual pressure dips observed at each station. As in the observational data,
there is no sign of a pressure dip at stations I and L. In
contrast to the observational data, the simulated pressure dip was not recorded at stations M and N, although
the reason for this is not clear.
Figure 8 shows the time sequence data of the simulated sea level pressure from 0500 to 1800 UTC 17 October 1998 together with the simulated hourly rainfall.
At 1300 UTC when the typhoon center was locatedover western Shikoku (Fig. 8d), a low pressure region
distinct from the typhoon center extended toward thenorth from an area west of the typhoon center. This low
pressure region corresponded to the pressure dip axis,

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FIG. 6. Hourly rainfall region of more than 1 mm h⫺1 derived from the Radar-AMeDAS data and the pressure dip
axis at 0700, 0900, 1300, and 1500 UTC 17 Oct 1998. The solid lines denote the location of the pressure dip axis.

as defined by isopleths of the minimum pressure within
the pressure dip. The passage of the pressure dip axis
can be traced by a sudden decrease and then increase in
the simulated sea level pressure recorded at several stations (Fig. 7). The pressure dip has a length of 200 km
and a width of 10–30 km, which is consistent with thedimensions derived from observational data (Fig. 5). It
should also be noted that as in the actual situation (Fig.
6), rainfall precedes the arrival of the simulated pressure dip axis. The simulated rainfall is weak at the center of the pressure dip and ceases with the passage ofthe pressure dip axis, which is consistent with the rainfall features observed (Fig. 6).
The simulated pressure dip was not recorded before
0500 UTC (Fig. 8a) but developed in the northwest
quadrant of the typhoon after 0800 UTC (Fig. 8b) when
the typhoon center was located over the sea to thesouthwest of Kyushu. This indicates that the formationof the pressure dip is not related to the topography of
the Japanese islands (Fig. 2). The pressure dip axis thenmoved northeastward with the typhoon and extended
to an area west of the typhoon center (Fig. 8c). The
pressure dip was most clearly discernible at 1300 UTC(Fig. 8d). As the typhoon moved farther northeast, thepressure dip broke down (Figs. 8e,f) and was not recorded at stations M and N (Fig. 7).
Although the amplitude of the simulated pressure
dips was weaker than that of the observed dips, thefeatures of the simulated pressure dips were sufficiently
similar to those of the observed pressure dips that we
could confidently use the simulation data to investigate
the formation mechanism of pressure dips.

c. Structure of the simulated pressure dip
Figure 9 shows the vertical sections of the potential
temperature anomaly at each pressure level across station H at 1300 UTC (see Fig. 8d) together with the sea
level pressure. The potential temperature anomaly wascalculated as the deviation from the average potential
temperature of the entire D2 area. The upper tropo-

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FUDEYASU ET AL.

FIG. 7. Simulated time series data of the sea level pressure for meteorological stations during
the passage of a typhoon from 16 to 17 Oct 1998. Station locations are indicated in Fig. 3. The
detection of a pressure dip at each station is indicated by a dashed circle.

sphere contains warm anomalies associated with thewarm core of the typhoon that result in a broad lowering in sea level pressure. In the troposphere below 700
hPa, warm potential temperature anomalies existed
east of 132°E because of a northward warm advection
associated with cyclonic circulation, while the anomalies were relatively cool to the west of 132°E. The lowlevel front at 132°E was consistent with the changes insurface temperature observed prior to the pressure dip,which were related to fronts associated with the typhoons rather than the pressure dip. At 131.5°E in thelower troposphere, within the colder region of the lowlevel front at 132°E, there were remarkable warm potential temperature anomalies from 700 to 900 hPa that
corresponded to a dip in sea level pressure.
Here we use the warm potential temperature anomalies, assuming a hydrostatic balance, as a tool to estimate the magnitude of the pressure dip. The hypsometric equation is written as follows:⌬z ⫽ z2 ⫺ z1 ⫽

冉冊

RT
p1
ln
g
p2

共z1 ⬍ z2 , p1 ⬎ p2兲,

共1兲

where p is the pressure (Pa), z is the height (m), g is theacceleration due to gravity (m s⫺2), R is the universal
ideal gas constant, and T is the mean temperature (K)of the layer of thickness ⌬z. The anomalous temperature ⌬T is measured as the deviation from the mean
temperature T and determines the pressure p1 ⫹ ⌬p as
follows:
ln共 p1 ⫹ ⌬p兲 ⫽TT ⫹ ⌬Tln

冉冊p1⫹ ln p2.
p2

共2兲

Under the conditions described in Fig. 9, T ⫽ 310 K,
p1 ⫽ 900 hPa, p2 ⫽ 700 hPa, and ⌬T ⫽ 5.0 K, whichmakes ⌬p ⫽ ⫺3.8 hPa. This estimate of the pressure
decrease within pressure dips is comparable to thesimulation value of about 4.5 hPa, suggesting that thewarm potential temperature anomalies in the lower troposphere are the major cause for the formation of pressure dips.
Figure 10 shows the vertical sections of vertical velocity, relative humidity with winds in the plane, cloudwater/ice mixing ratio, and rainwater/snow mixing ra-

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FIG. 8. Simulated sea level pressure fields and hourly rainfall within D2 at (a) 0500, (b) 0800, (c) 1100, (d)1300, (e) 1500, and (f) 1800 UTC 17 Oct 1998. Contour interval is 2 hPa. The detection of a pressure dip isindicated by a dashed circle. Regions of rainfall of 10–20 mm h⫺1 are heavily shaded while regions withgreater than 20 mm h⫺1 are lightly shaded.

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FIG. 9. Cross section of (a) sea level pressure and (b) potential
temperature anomalies at 1300 UTC 17 Oct 1998. Location ofcross section is indicated by the line A–B in Fig. 8d, at the latitudeof 34.1°N. Contour interval is 1 K. Regions with anomalies of 3–5
K are heavily shaded while those greater than 5 K are lightly
shaded. The detection of warm potential temperature anomalies
related to a pressure dip is indicated by a dashed circle.

tio. Near the typhoon center, there is a relatively moist
region accompanied by strong ascending flow throughout the entire troposphere (Figs. 10a,b). Over the pressure dip region immediately west of the ascendingflows, there is a locally strong descending flow from 300
hPa down to the surface, which is accompanied by awavelike vertical circulation to the west. In association
with the strong descending flow, there is the relatively
dry region below 600 hPa in the same narrow region as
the warm potential temperature anomalies. The dry region is connected with a dry region in the upper andmiddle troposphere and is accompanied by the westerly
winds.
At higher altitude, ice was advected outward fromthe typhoon center, forming the overhanging anvil overthe pressure dip region (Fig. 10c). This ice water in theanvil aggregated and transformed into snow as it approached the 0°C level (Fig. 10d). It should be notedthat in the present simulations, cloud water and rainwater below 0°C are treated as cloud ice and cloudsnow in the cloud microphysics (Dudhia 1989). Therefore, when cloud snow fell through the 0°C level, it
melted and fell to the ground as rain. The change in fallvelocities from snow to rain caused the concentrationsof water content to be larger above the melting level

1235

than below the melting level. The pressure dip was
found below the overhanging anvil. These images imply
that the pressure dip is linked to warm potential temperature anomalies in the lower troposphere, accompanied by the intrusion of an upper dry air mass caused bya locally strong descending flow.
The descending flow area is found along the confluence zone resulting from the interaction between thetyphoon’s circulation and the environmental westerly
winds when the typhoon enters the zone of midlatitude
westerlies (Fig. 11). The inflow of a dry air mass by the
midlatitude westerlies into the moist area of the typhoon induces a frontal zone of equivalent potential
temperature in the middle troposphere. At the beginning of the simulation, the confluence zone accompanied by descending airflow existed far to the north of
the typhoon (Figs. 11a,b). As the typhoon moved
northeastward, the confluence and descending airflow
intensified at the rear of the typhoon (Figs. 11c,d). As
the typhoon continued to move still farther northeastward, the confluence to the rear of the typhoon rapidly
became weaker and the descending airflow disappeared(Figs. 11e,f). Therefore, the occurrence of the simulated
descending flow area seems to be related to the evolution of the interaction between the typhoon’s circulation and midlatitude westerlies.

4. Formation mechanisms of the simulatedpressure dip
a. Heat budget
We will now consider how the warm potential temperature anomalies in the lower troposphere form. The
process of anomaly formation can be investigated bydescribing the anomalies in terms of a thermodynamic
equation. The time tendency of potential temperatureis estimated by the difference in the potential temperature sampled at a time interval of 3 min. The diabatic
term, including radiation, diffusion, and net heating related to phase changes of water, is defined as the residual derived from both the horizontal and verticaladvection terms and the time tendency.
Figure 12 displays the time series data of the potential temperature anomalies, which are calculated as thedeviation of temperature at a point from the averagetemperature over the D2 area at sea level pressure.
Data were calculated at each pressure level over station
H (Fig. 3), where the largest dip in pressure was detected (Fig. 7). During the period from 0700 to 1500
UTC, sea level pressure decreased in association withthe passage of the typhoon (Fig. 12a). This decrease insurface pressure was related to the warm core in theupper troposphere (Fig. 12b). The decrease due to the1236

MONTHLY WEATHER REVIEWVOLUME 135

FIG. 10. Same cross-sectional profiles as in Fig. 9 but for (a) vertical velocity, (b) relative humidity and the zonal
and vertical components of the winds (m s⫺1 arrows, scale at bottom left), (c) cloud water/ice mixing ratio, and (d)rainwater/snow mixing ratio. Contour interval in (a) is 0.5 m s⫺1. Dark shading represents areas with verticalvelocity less than 0.0 m s⫺1 while light shading represents areas of vertical velocity less than ⫺0.5 m s⫺1. Contour
interval in (b) is 10% and the regions greater than 80% are shaded. Contour interval in (c) is 2.0 ⫻ 10⫺5 kg kg⫺1
and in (d) is 1.0 ⫻ 10⫺3 kg kg⫺1. The solid lines indicate the level of 0°C.

pressure dip appeared at station H from 1245 to 1325
UTC, coinciding with the development of warm potential temperature anomalies in the lower troposphere.
The thermodynamic balance showed that the warmingin the lower level from 1315 to 1330 UTC was caused
mainly by vertical advection (Fig. 13). Although thecooling due to the diabatic term contributed partly tocounteracting the warming, the adiabatic warming of
the descending air in the lower level was not balanced
by diabatic cooling. The cooling due to the horizontal
advection term accompanied by the cyclonic circulation
was relatively small in the lower level. This indicatesthat the adiabatic descent of air is responsible for theformation of warm potential temperature anomalies in

the lower troposphere. However, at a level above 600
hPa, warming due to vertical advection was balanced by
the diabatic term. Warming at 400 hPa resulted fromthe outflow of warmer air from the typhoon center,
while cooling at 550 hPa seemed to be related to the
inflow of environmental air with a lower potential temperature caused by the westerlies.

b. Trajectory analysis
To further clarify the formation of the warm potential temperature anomalies in the lower troposphere,
we conducted a backward trajectory analysis using 3D
winds sampled at 3-min intervals. Initially, 80 parcels
were uniformly defined in the 700-hPa level over

APRIL 2007

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FIG. 11. Winds (m s⫺1 arrows, scale at the bottom left) at 500 hPa for (a) 0500, (c) 1300, and (e) 1800 UTC
17 Oct 1998. (b), (d), (f) Same as in (a), (c), (e) but for equivalent potential temperature and verticalvelocity. Regions larger than 0.2 m s⫺1 are lightly shaded while regions smaller than ⫺0.2 m s⫺1 are heavily
shaded. Contour interval of equivalent potential temperature is 10 K. Solid broken lines indicate theconfluence zone between the typhoon’s circulation and the midlatitude westerly winds while crosses represent the position of the typhoon center.

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FIG. 12. Time series data of (a) sea level pressure and (b)anomalies of potential temperature at station H for the period
0600–1600 UTC 17 Oct 1998. Contour interval is 1 K. Dark shading indicates areas with an anomaly of 3–5 K while light shading
indicates areas with anomalies greater than 5 K.

the surrounding area of station H (34.0°–34.7°N,
131.5°–132.0°E) at 1300 UTC. This region includes boththe pressure dip axis and surrounding areas. Figure 14
shows the trajectories of the parcels over 12 h between
0100 and 1300 UTC. The origins of the parcels that
resided over the pressure dip can be traced back in timeto the eastern coast of China (Fig. 14a), while the parcels in the surrounding area originated from the eastnortheast and, to a lesser degree, from the south (Fig.
14b).
Figure 15 presents the time series data of several
averaged quantities of parcels that originated from thewest (western parcels), which were initially released at
the 700-hPa level over the surrounding area of station
H. Also shown in the figure are the time series data ofparcels that originated from the east-northeast (easternparcels), which were initially situated slightly east of
station H. In the colder region over the northern coastof Honshu, the eastern parcels remained below 3000 m
until 1000 UTC (Figs. 14b and 15a). At 1000 UTC, theeastern parcels were brought into the typhoon centerand ascended to 3000 m. The potential temperature andequivalent potential temperature increased with the
moist-adiabatic lapse rate (Fig. 15b).
The western parcels could be traced back to the upper level around 10 000 m (above 300 hPa) at 0100 UTC

VOLUME 135

FIG. 13. Vertical profiles of (a) time tendency of potential temperature, (b) horizontal advection, (c) vertical advection, and (d)diabatic effect at station H for the period 1315–1330 UTC 17 Oct
1998.

(Fig. 15a). Up until 1030 UTC, strong upper-level westerly winds propelled the western parcels to the eastnortheast (Fig. 14a) on the gently descending slope of
the isentropic surface within the upper troposphere.
The western parcels conserved potential temperatureand equivalent potential temperature (Fig. 15b) andsubsequently had low humidity (Fig. 15d). The initialmotion of the western parcels was thus predominantly
adiabatic. Once the western parcels had overtaken thenortheastward-moving typhoon between 1030 and 1300
UTC, they moved rapidly downward from 8000 to 3000
m within the intense descending flow (Figs. 15a,c). During this period, the water mixing ratio of the westernparcels increases along with relative humidity (Fig.
15d), indicating that water was supplied from the moist
region of the typhoon. Although the mixing ratio of thewestern parcels increased, the equivalent potential temperature remained steady as the potential temperature
decreased from 1030 to 1200 UTC (Fig. 15b). From1245 to 1300 UTC, when the pressure dips began toappear in the vicinity of station H, the changes in potential temperature, humidity, and water mixing ratio
were small, suggesting that evaporative cooling was insufficient to offset adiabatic warming because there
were fewer hydrometeors below the melting level.

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1239

FIG. 14. Trajectories of 80 parcels over 12 h between 0100 and 1300 UTC 17 Oct 1998. The origins of theparcels are at the 700-hPa level over the (a) pressure dip axis and (b) surrounding areas.

Thus, the western parcels were warmer by 4 K than theeastern parcels (Fig. 15b).
The above trajectory analysis suggests that downdrafts play an important role in the formation of pressure dips. As a typhoon moves into the midlatitude
westerlies, dry air associated with the westerlies enters
the typhoon, which contains saturated air. The interaction between the typhoon’s circulation and the westerly
winds forms equivalent potential temperature fronts at
the midlevel, while an overhanging anvil is evident at
the upper level. Next, dry air entering just below theanvil is cooled by evaporation and sublimation (Figs.
15b,d), causing it negative buoyancy (Fig. 15f), hence
producing downdrafts (Fig. 15c). Thus, the formation ofa mesoscale downdraft results from the cooling of
evaporation and sublimation related to dry air masses
carried by the upper-level westerly winds toward the
moist typhoon region. The downdraft peaked at 1215
UTC (Fig. 15c), when the parcels reached the melting
level (Figs. 10c,d). These parcels underwent accelerated
evaporation and melting until they passed through themelting level. After passing the melting level, evaporative cooling became insufficient to offset adiabatic
warming because there was an insufficient amount ofwater in the lower level (Figs. 15e,f). As a result, warm
potential temperature anomalies appeared in the lower
level, causing a pressure dip. The horizontal scale ofpressure dips was determined by the width of the regionof mesoscale downdrafts that occurred around themelting level. Warming in the lower levels gave theparcels positive buoyancy (Fig. 15f), resulting in the

deceleration of downdrafts (Fig. 15c). It should be
noted that the downdraft impinging on the lower levels
excited gravity waves on the rear side (Figs. 10a,b), but
the induced surface pressure change was much weaker
than that caused directly by the downdrafts (Fig. 9).

5. Pressure dips associated with other typhoons
a. Features of observed pressure dips
In the previous sections, we described the characteristics of pressure dips observed during the passage ofZeb. Here we describe the general characteristics ofpressure dips. During the period from 1980 to 1998, 89
tropical cyclones of at least tropical storm intensity
passed across Japan or moved along the southern coastof Japan. Seven of these typhoons were accompanied
by pressure dips. A greater proportion of the typhoons
may have been accompanied by pressure dips, but suchdips in typhoons (whose tracks are at sea, far from theJapanese islands) could not be detected. Furthermore,
our definitions excluded pressure dips with a small amplitude and short persistence. The characteristics ofeach pressure dip and typhoon are summarized inTable 2, while the typhoon tracks and locations of thepressure dip axis defined by the isopleths of minimumpressure within the pressure dip are shown in Fig. 16.
All of these typhoons had considerable intensity withcentral pressures below 980 hPa at the time they approached the Japanese islands (Table 2). They moved
quickly toward the northeast or north-northeast during
their passage over the Japanese islands, weakened rap-

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MONTHLY WEATHER REVIEWVOLUME 135

FIG. 15. Time series data of several averaged parameters of the western (solid line) and eastern (dashed line)
parcels over 12 h between 0100 and 1300 UTC 17 Oct 1998: (a) height, (b) potential temperatures and equivalentpotential temperatures, (c) vertical velocity, (d) water vapor mixing ratio and relative humidity, (e) rainwater/snow
mixing ratio, and (f) the anomalies of potential temperature.

idly, and eventually transformed into extratropical cyclones.
Typhoon Gay approached near Tokyo on 22 October
1981 without making landfall (Fig. 16a). At that time,
the pressure dip axis extending to the north and south
was recorded along the southern coast of Honshu. Ty-

phoon Wayne moved along the southern coast of theJapanese islands from 19 to 20 September 1989 (Fig.
16b). The pressure dip axis related to the movement of
the typhoon was observed in Kyushu and along thesouthern coast of Shikoku. Typhoon Hattie accelerated
northeastward along the coast of the Japanese islands

TABLE 2. Summary of the features of the detected seven typhoons with pressure dips.

TyphoonLandfalldate

Central pressure
at landfall (hPa)

Movement speed
at 32°N (km h⫺1)

No. of stations
that observed dip

Max amplitudeof dip (hPa)Avg amplitudeof dip (hPa)

Length of dip
axis (km)

Gay
Wayne
HattieMireille
Oliwa
Vicki
Zeb

22 Oct 1981
19 Sep 1989
8 Oct 1990
27 Sep 1991
16 Sep 1997
22 Sep 1998
17 Oct 1998

965
975
975
940
965
960
980

40
25
25
45
35
50
20

21
14
30
18
15
13
28

9
7
5
6
5
4
8

4.2
2.4
2.5
2.7
2.1
1.9
3.9

150
100
100
150
100
150
150

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FIG. 16. Progressive locations of the pressure dip axis (solid lines) and accompanying typhoons (black circles) for
Typhoons (a) Gay, (b) Wayne, (c) Hattie, (d) Mireille, (e) Oliwa, and (f) Vicki. Positions of typhoon centers, tracked
at 3-h intervals, are denoted by closed circles. Hourly and 3-hourly locations of the pressure dip axis are indicated bythin and thick lines, respectively.

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MONTHLY WEATHER REVIEWVOLUME 135

FIG. 17. Amplitude of the pressure drop recorded at different
stations.

from 7 to 9 October 1990 (Fig. 16c). The pressure dip
axis extending from the southeast to the northwest occurred along the southern coast of the Japanese islands.
Typhoon Mireille passed over the western coast of theJapanese islands in late September 1991 (Fig. 16d). On
this occasion, the pressure dip axis extended from eastsoutheast to west-northwest and was observed in thenorthern part of Kyushu and the western part of the
Chugoku district. Typhoon Oliwa made landfall on thesouthern part of Kyushu in mid-September 1997 (Fig.
16e). The pressure dip axis associated with Oliwa extended from south to north, appearing from southern
Kyushu to the western Chugoku district. Typhoon
Vicki traveled over the Kinki district in late September
1998 (Fig. 16f) and was accompanied by an eastsoutheast to west-northwest oriented pressure dip axis.
Figure 17 presents the amplitude of pressure dips
recorded at various meteorological stations. The recorded amplitudes ranged from 1 to 9 hPa, with a typical value of 2 hPa, while the observed pressure dips
were generally of 20–40-min duration (not shown).
These values are similar to those reported in previous
studies (e.g., Fujita 1992). There is no apparent relationship between the amplitude of the pressure dipsand the speed and intensity of the associated typhoons(not shown).
Figure 18 shows the geographical locations of thepressure dip axes with respect to the centers of thetyphoons. All the pressure dip axes were located to thewest of the typhoon centers, with the majority within
the northwestern quadrant. The distance between thepressure dip axes and the accompanying typhoon centers ranged from 50–300 km. All subsequent references
to location relative to the typhoon center are described
relative to the direction of movement of the typhoon.
Fujita (1952) reported that pressure dips appeared on

FIG. 18. Locations of the pressure dip axis relative to each
typhoon center based on geographic coordinates.

both the left- and right-hand sides of Typhoon Della,
but that the pressure dip on the left-hand side of thetyphoon had an amplitude of 7 hPa, while those on theright-hand side had a lesser amplitude. Nakajima et al.
(1980) documented pressure dips on both the left- andright-hand sides of Typhoon Owen, but the dip on theright-hand side of the typhoon lasted only several hours
and was weaker than the dip to the left of the typhoon.
In all other documented cases, pressure dips were observed only to the left of the typhoon centers. Becasue
all typhoons with pressure dips moved northeastward,
we can also conclude that the occurrence of pressure
dips with a significant amplitude and duration are confined to the left and the rear of the direction of thetyphoon’s movement.
Figure 19 shows the rainfall patterns when the dip
associated with each typhoon has its maximum peak.
As described earlier in the case of Typhoon Zeb (Fig.
6), there is a remarkable change in rainfall intensity
across the pressure dip axis. The intense rainfall region
occurs in advance of the pressure dip axis, while little
rainfall is recorded following the passing of the pressure
dip axis. Thus, the cessation of rainfall around pressure
dips is a common feature of this phenomenon.

b. Simulation of pressure dips
To confirm the process of pressure dip formation described for Typhoon Zeb, we examined the structure of
the simulated pressure dips associated with other typhoons. Figure 20 shows the simulated sea level pressure in D2 for five typhoons (Wayne, Hattie, Mireille,

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1243

FIG. 19. Hourly rainfall region greater than 1 mm h⫺1
derived from the Radar-AMeDAS data and the pressure
dip axis at (a) 0600 UTC 19 Sep 1989 (Typhoon Wayne), (b)1900 UTC 7 Oct 1990 (Typhoon Hattie), (c) 1000 UTC 27
Sep 1991 (Typhoon Mireille), (d) 1100 UTC 16 Sep 1997
(Typhoon Oliwa), and (e) 0600 UTC 21 Sep 1998 (Typhoon
Vicki).

Oliwa, and Vicki) at the time when the pressure dips
were most developed. Warm potential temperature
anomalies exist in the lower troposphere over the pressure dip axis (Fig. 21). Although the amplitudes of thesimulated pressure dips are somewhat different fromthe observed values, the simulated pressure dip wasclosely related to deficiencies in the simulated warm
potential temperature anomalies in the lower tropo-

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MONTHLY WEATHER REVIEWVOLUME 135

FIG. 20. Simulated sea level pressure
fields within domain 2 at (a) 1500 UTC
19 Sep 1989 (Typhoon Wayne), (b) 1100
UTC 7 Oct 1990 (Typhoon Hattie), (c)0900 UTC 27 Sep 1991 (Typhoon
Mireille), (d) 2000 UTC 15 Sep 1997 (Typhoon Oliwa), and (e) 1100 UTC 22 Sep
1998 (Typhoon Vicki). Contour interval
is 2 hPa except in (b) and (e) it is 1 hPa.
The detection of a pressure dip axis isindicated by a dashed circle.

sphere. These results strongly support the conclusionthat the pressure dips associated with the other typhoons have a comparable structure and formative process to that described in previous sections for the caseof Typhoon Zeb. The pressure dips observed within
Typhoon Gay, however, were not successfully simulated. The pressure dip recorded within Typhoon Gay
may have been generated by a different mechanism,
such as an internal gravity wave excited by the convection of the typhoon (Matsumoto and Okamura 1985).
However, the exact reason for the inability to properly
simulate the pressure in Typhoon Gay remains unclear.

6. Discussion
The present study suggests that the formation mechanism of a pressure dip is similar to that of a wake low(e.g., Johnson and Hamilton 1988; Johnson 2001). A
wake low generally appears near or just behind thesquall lines, in association with warming caused by unsaturated descent. The descending flow plays a role in
the formation of the warming and is connected to rear
inflow into the squall line (Johnson and Hamilton

1988). Past studies propose several hypotheses regarding the causes of the strong descending flow (Miller andBetts 1977; Johnson and Hamilton 1988; Zhang and
Gao 1989; Schmidt and Cotton 1990; Gallus 1996). In
addition, it is suggested that the environmental flow inwhich the squall line occurs determines the intensity of
the rear inflow into the squall line at the middle andupper troposphere.
In the case of pressure dips, the midlatitude westerlies corresponding to rear inflows are extremely strong(see Fig. 23), while many wake lows have weaker rear
inflow into the squall line at the middle and upper troposphere (Houze 1993). The midlevel confluence between the typhoon’s circulation and the midlatitude
westerlies is linked to the formation of a steep upwardsloping westward front at midlevels where an intense
downdraft develops. A pressure dip therefore appears
to be caused by the effect of the intense midlatitude
westerlies on typhoons.
Previous studies (Uccellini and Koch 1987; Ramamurthy et al. 1993; Koch and Siedlarz 1999) have investigated the large pressure fluctuations associatedAPRIL 2007

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1245

FIG. 21. Cross sections of sea levelpressure and potential temperature
anomalies at the (a) longitude of
130.0°E, (b) longitude of 129.3°E, (c)longitude of 129.0°E, (d) latitude of
32.8°N, and (e) longitude of 136.2°E, indicated by line A–B in Fig. 20 at thesame time frame as Fig. 20. The contour
interval is 1 K. The regions with anomalies of 3–4 K are heavily shaded while
regions with anomalies greater than 4 K
are lightly shaded. The detection ofwarm potential temperature anomalies
related to a pressure dip is indicated bya dashed circle.

with extratropical cyclones, which are somewhat similar
to the characteristics of pressure dips. They suggested
that the fluctuations were excited by internal gravity
waves through a geotropic adjustment process andshearing instability. Although the above process may
be important for the generation of some pressure dips,for example, those recorded within Typhoon Gay (Matsumoto and Okamura 1985), the simulated results inthe present study indicate that downdrafts play a fundamental role in the formation of pressure dips.
Ritchie and Elsberry (2001) used a mesoscale nu-

merical model to study the transition from tropical cyclones to extratropical cyclones. Under ideal baroclinic
conditions, a tropical cyclone undergoing transformation develops a low-level warm anomaly to the westsouthwest of the cyclone, which has similar characteristics to the pressure dip. The results of this study and of
Ritchie and Elsberry’s (2001) study demonstrate that atropical cyclone undergoing transformation within midlatitude, large-scale, atmospheric conditions generates
a diplike structure in surface pressure. Thus, one could
conclude that a pressure dip is a unique phenomenon

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MONTHLY WEATHER REVIEWVOLUME 135

FIG. 22. Winds (m s⫺1 arrows, scale at bottom) and equivalent potential temperature at 850 hPa for seven different typhoons as they made landfall over the Japaneseislands. The contour interval is 5 K and the crosses represent the position of thetyphoon centers.

intrinsic to the transition of a tropical cyclone, although
not every transforming typhoon is accompanied by thedevelopment of a pressure dip.
Figure 22 shows the potential temperature and windat 850 hPa, while Fig. 23 shows the TBB and wind at 500
hPa at the time when each of the typhoons listed inTable 2 made landfall in the Japanese islands. In allcases, there were steep thermal gradients in the lowlevels to the north and west of the typhoon centers. As
mentioned in section 3, sudden changes in surface temperature and dewpoint temperature that preceded thearrival of the pressure dips (Fig. 1) are related to the

influence of the surface fronts. At the mid- and upper
troposphere, there is a cloudy region known as thedelta rain shield (Shimazu 1998), which extends northeast from the typhoon centers (Fig. 23). The midlatitude westerlies are present in the Japanese islands, accompanied by a trough to the west of Japan. Such largescale atmospheric flow is favorable for the northeast
movement of typhoons across the Japanese islands andfor the detection of pressure dips at land-based meteorological stations. These features are the characteristics
of a typhoon moving into the midlatitude areas ofstrong baroclinicity.

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1247

FIG. 23. Winds (m s⫺1 arrows, scale at bottom) and equivalent blackbody temperature (TBB) for seven different typhoons as they made landfall over the Japaneseislands. The contour interval is 10 K. Regions with a TBB of 240–260 K are heavily
shaded while regions with a TBB smaller than 240 K are lightly shaded. The crosses
represent the position of the typhoon centers.

It is interesting to note that the pressure dips were
only observed during September and October (Table2), when the atmospheric conditions described above
develop most commonly. Pressure dips reported in previous studies were also associated with typhoons thatcrossed Japan during September and October (e.g., Nakajima et al. 1980). One notable exception is Typhoon
Della, which made landfall in June 1949 (Fujita 1952).
Although the conditions described above also occur
during June and early July, typhoons rarely make landfall over Japan. Typhoons most frequently cross Japanin late July and August, but pressure dips are not ob-

served during the summer because the westerly winds
retreat to the north. Thus, it would appear that pressure
dips tend to occur in September and October when atyphoon enters the zone of midlatitude westerlies. The
atmospheric conditions described above might be favorable for the transformation of the typhoon structure
into an extratropical storm (Muramatsu 1982; Kitabatake 2002). The typhoons that had pressure dips wereof considerable intensity at the time they approached
the Japanese islands. These typhoons then rapidly weakened and eventually transformed into extratropical cyclones during their passage over the Japanese islands.

1248

MONTHLY WEATHER REVIEWThe area of downdraft overlapped with the confluence zone generated by the interaction between thetyphoon’s circulation and the environmental westerly
winds as the typhoon entered the zone of midlatitude
westerlies (Fig. 11). Although the same formation
mechanism was found to the north and south of thetyphoon center in our simulations, pressure dips werediscernible only over that part of the downdraft region
on the western side of the typhoon center (Figs. 8 and
18). The intensity of confluence between the typhoon’s
circulation and the environmental westerly winds affected the relative amount of dry air entering the moist
region at mid- and upper level, which in turn contributed to the difference in the intensity of the downdraft.
The downdraft to the north was not intense enough toform the pressure dip (Fig. 11). On the other hand, theamplitude of the pressure dips was related to the potential temperature anomalies, which were the deviations from environmental potential temperature. The
environmental low-level potential temperature to thesouth, the warm region of the surface front (Fig. 22),
was relatively warm. Therefore, warm potential temperature anomalies related to a pressure dip in lower
levels tended to develop only on the western side of thetyphoon center and to the north of the surface front
where there was advection of cold air (see Fig. 24). This
explains why pressure dips appeared only in limitedareas although the downdrafts occurred over a relatively elongated area. Furthermore, the northeastward
movement of the typhoon and the locational difference
between the midlevel and surface fronts (see Fig. 24a)
caused the frontal passage prior to the pressure dip to
be unobserved at stations far from the track of the typhoon, as described in section 3a.
After analyzing U.S. hurricanes and Japanese typhoons, Fujita (1992) noted that there are two different
types of pressure dips: (i) cold-sector dips, which formin the wake of a rainband in the northwestern sector,and (ii) warm-sector dips, which form in any sector of
the storm where there is free cold-air inflow. The
former type is similar to the pressure dips consideredhere while the latter type seems to correspond to pressure dips that we excluded in the present study. However, because of the limitations of the available data,
the characteristics of the pressure dips associated withhurricanes are not clear. Such an examination could beconsidered for future study.

7. Summary
We simulated the pressure dips associated with Typhoon Zeb using a mesoscale numerical model. The
model successfully reproduced the major features of

VOLUME 135

FIG. 24. Schematic of the (a) circulation around the pressure dip
and (b) cross section through the pressure dip indicated by theline A–B in (a). Arrows in (a) indicate streamlines at the midlevel
while those in (b) indicate the trajectories of the upper dry air.
The broken line in (a) is the confluence zone and equivalentpotential temperature fronts resulting from the interaction between the typhoon’s circulation and the environmental westerly
winds at the midlevel. The frontal symbols indicate fronts, thebroken line in (b) is the melting level, and the letter L refers thesurface low pressure region (i.e., pressure dip). The cross represents the position of the typhoon center. The hatched regions
indicate the region of warm potential temperature anomaly related to the pressure dip. Note that horizontal scales differ in thetwo schematics.

APRIL 2007

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FUDEYASU ET AL.

pressure dips and thus permitted us to investigate theformation mechanism of the pressure dips.
The formation mechanism of the pressure dips is illustrated schematically in Fig. 24. The simulated pressure dips were closely linked to warm potential temperature anomalies in the lower troposphere, which areaccompanied by downdrafts throughout the entire troposphere. In association with the anomalies, relative
humidity and rainwater decreased while an overhanging anvil above the melting level developed. As thetyphoon moved into the midlatitude westerlies, theconfluence zone and equivalent potential temperature
fronts resulting from the interaction between the typhoon’s circulation and the environmental westerly
winds appeared at midlevel and the inflow of a dry air
mass occurred (Fig. 24a). As dry air carried by the midlatitude westerlies entered the moist region below theanvil, the cooling of evaporation and sublimation
caused negative buoyancy, which resulted in the formation of downdrafts (Fig. 24b). The parcels underwent
accelerated evaporation until passing the melting level.
Below the melting levels, where there were fewer hydrometeors, evaporative cooling was insufficient to offset adiabatic warming. As a result, the warm potential
temperature anomalies occurred in the lower levels, resulting in the formation of pressure dips. The warm
potential temperature anomalies related to a pressure
dip in lower levels tended to develop only on the western side of the typhoon center and to the north of thesurface front. The horizontal scale of pressure dips wasdetermined by the width of the region of mesoscale
downdrafts around the melting level. Similar processes
were found to operate during the simulation of pressure
dips associated with other typhoons.
In the present study, we have also described the features of pressure dips observed during the period from1980 to 1998 using barographic data from different meteorological stations. During this period, pressure dipsof significant amplitude and duration were recorded in
7 out of 89 typhoons with at least tropical cyclone intensity. The observed pressure dips occurred only on
the western side of the typhoon centers, at a distance of
50–300 km from the typhoon center, and were accompanied by a sudden cessation of rainfall. All of thesetyphoons moved rapidly toward the northeast or northnortheast during their passage over the Japanese islands before weakening rapidly and eventually transforming into extratropical cyclones. The pressure dips
were observed under large-scale environmental conditions characterized by westerly winds associated with a
midlatitude trough with a dry air mass to the west ofJapan at upper levels and fronts at lower levels. The
pressure dips were detected in typhoons that occurred

during the boreal autumn when the above large-scale
environmental conditions appear around Japan. No
pressure dips were observed in typhoons during theNorthern Hemisphere’s summer when most typhoons
approach Japan. We therefore would suggest that apressure dip is an inherent feature of the asymmetricstructure of a typhoon undergoing extratropical transition.
The proportion of typhoons associated with pressure
dips may be higher than the detected numbers because
we were only able to detect pressure dips in typhoons
that passed directly over the Japanese islands. It may bepossible to find more pressure dips in typhoons that arelocated solely over the Pacific Ocean from highresolution satellite data. Such data would increase ourability to understand the characteristics of meso-␤-scaleatmospheric phenomena associated with tropical cyclones, including pressure dips, and lead to a better
understanding of tropical cyclones.
Acknowledgments. This paper is part of H. Fudeyasu’s Ph.D. thesis at Kyoto University. Thanks are extended to the staff and students at the Disaster Prevention Research Institute, Kyoto University, for their help
and encouragement. The authors also expresses hearty
thanks to Prof. O. Tsukamoto for his thoughtful comments and for providing observational data. Many
thanks also to Drs. T. Fujii, Y. Tsujimura, H. Okamura,
N. Kitabatake, T. Satomura, H. Ishikawa, T. Itano, andJ. Yoshino for fruitful discussions, comments, and criticisms. The authors are grateful for the efforts of many
scientists at both PSU and NCAR to maintain MM5
and they acknowledge the Meteorological Research Institute and Japan Meteorological Agency for supplying
meteorological data of past tropical cyclones. This work
is partly supported by the Research Project for Suitable
Coexistence of Human, Nature, and the Earth under
the Japanese Ministry of Education, Sports, Culture,
Science, and Technology.
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