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Multiscale Interactions in the Life Cycle of a Tropical Cyclone Simulated in a Global
Cloud-System-Resolving Model. Part II: System-Scale and Mesoscale Processes*
HIRONORI FUDEYASU1 AND YUQING WANG
International Pacific Research Center, School of Ocean and Earth Science and Technology, University of Hawaii
at Manoa, Honolulu, Hawaii

MASAKI SATOH
Atmosphere and Ocean Research Institute, University of Tokyo, Kashiwa, and Japan Agency for Marine Earth Scienceand Technology, Yokohama, JapanTOMOE NASUNO
Japan Agency for Marine Earth Science and Technology, Yokohama, JapanHIROAKI MIURA AND WATARU YANASE
Atmosphere and Ocean Research Institute, University of Tokyo, Kashiwa, Japan(Manuscript received 22 April 2010, in final form 28 July 2010)
ABSTRACTThe life cycle of Tropical Storm Isobel was simulated reasonably well in the Nonhydrostatic Icosahedral Atmospheric Model (NICAM), a global cloud-system-resolving model. The evolution of the large-scale circulationand the storm-scale structure change was discussed in Part I. Both the mesoscale and system-scale processes in
the life cycle of the simulated Isobel are documented in this paper. In the preconditioned favorable environment
over the Java Sea, mesoscale convective vortices (model MCVs) developed in the mesoscale convective systems(MCSs) and convective towers with cyclonic potential vorticity (PV) anomalies throughout the troposphere[model vortical hot towers (VHTs)] appeared in the model MCVs. Multiple model VHTs strengthened cyclonic
PV in the interior of the model MCV and led to the formation of an upright monolithic PV core at the centerof the concentric MCV (primary vortex enhancement). As the monolithic PV core with a warm core developed
near the circulation center, the intensification and the increase in horizontal size of the cyclonic PV were enhanced through the system-scale intensification (SSI) process (the secondary vortex enhancement), leading to thegenesis of Isobel over the Timor Sea. The SSI process can be well explained by the balanced dynamics.
After its genesis, the subsequent evolution of the simulated Isobel was controlled by both the external influence and the internal dynamics. Under the unfavorable environmental conditions, the development ofasymmetric structure reduced the axisymmetric diabatic heating in the inner core and the SSI process became
ineffective and the storm weakened. Later on, as the eyewall reformed as a result of the axisymmetrization of an
inward-propagating outer spiral rainband, the SSI process became effective again, leading to the reintensificationof Isobel. Therefore, the large-scale environmental flow provided the precondition for the genesis of Isobel andthe triggering mechanism for subsequent storm-scale structure change as discussed in Part I. The system-scale
and mesoscale processes, such as the evolution of MCVs and merging VHTs, were responsible for the genesis,
while the eyewall processes were critical to the storm intensity change through the SSI process.

* School of Ocean and Earth Science Technology Publication Number 7985 and International Pacific Research Center Publication
Number 714.
1
Current affiliation: Faculty of Education and Human Sciences, Yokohama National University, Yokohama, Japan.

Corresponding author address: Dr. Hironori Fudeyasu, Yokohama National University, 79-1 Tokiwadai, Hodogaya-ku, Yokohama,
Kanagawa 240-8501, Japan.
E-mail: fude@ynu.ac.jp
DOI: 10.1175/2010MWR3475.1
Ó 2010 American Meteorological Society

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1. Introduction
The first global cloud-system-resolving model, thatis, the Nonhydrostatic Icosahedral Atmospheric Model(NICAM; Tomita and Satoh 2004; Satoh et al. 2008)
with realistic settings and a horizontal grid spacing of
7 km simulated reasonably well the life cycle of Tropical Storm (TS) Isobel formed over the Timor Sea in
the austral summer of 2006 (Fudeyasu et al. 2008). In
Fudeyasu et al. (2010, hereafter Part I), we focusedon the evolution of the large-scale circulation and thestorm-scale structure change at different stages in the life
cycle of the simulated Isobel. We showed that the westerly wind burst (WWB) over the Java Sea associated withthe onset of an Madden–Julian oscillation (MJO) event
over the East Indian Ocean played a critical role in theorganization of small-/mesoscale convective cyclonic vortices and the genesis of Isobel. The temporary weakening
was found to be related to the eyewall breakdown triggered by strong vertical shear and a strong environmental
stretching deformation field, whereas the reintensification was associated with the eyewall reformation as a result of the axisymmetrization of an inward-propagatingouter spiral rainband. In this paper, we will document
the system-scale and mesoscale processes in the life cycle of the simulated Isobel with the focus on processes
related to the genesis and intensity change of the simulated Isobel.
The genesis remains the least understood stage in thelife cycle of a tropical cyclone (TC) because of the complex scale interactions involved and lack of in situ fineobservations over the open ocean where the TC forms.
In a preconditioned large-scale environment, the genesis
generally involves the mesoscale convective organization, including vortical hot towers (VHTs) and mesoscaleconvective vortices (MCVs) embedded in mesoscale convective systems (MCSs; e.g., Ritchie and Holland 1997;
Simpson et al. 1997; Bister and Emanuel 1997; Hendricks
et al. 2004; Montgomery et al. 2006). MCVs generally
form in the stratiform precipitation region at midtroposphere and have horizontal scales in the order of 10–
100 km (Bister and Emanuel 1997; Ritchie and Holland
1997) and VHTs are deep convective cells with potential
vorticity (PV) anomalies throughout the depth of thetroposphere on scales of 1–20 km embedded in a MCV(Hendricks et al. 2004; Montgomery et al. 2006). Only
when these multiscale systems work cooperatively cana cyclogenesis emerge.
Top down and bottom up are two hypotheses for TC
genesis. In the top-down hypothesis, the geneses of
MCVs occur within the large-scale stratiform precipitation region in the midtroposphere. The merging of
midtropospheric MCVs increases the horizontal scale

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and thus the downward penetration of anomalous midlevel cyclonic PV, leading to the cyclogenesis at thesurface (Ritchie and Holland 1997; Simpson et al. 1997).
The evaporative cooling-induced subsidence from stratiform clouds may also bring the midlevel PV anomaly
downward, leading to the genesis of a surface cyclonicvortex (Bister and Emanuel 1997). The bottom-up hypothesis describes the upward development of lowertropospheric cyclonic vortices via convective processes.
In their high-resolution cloud-resolving model simulations, Hendricks et al. (2004) and Montgomery et al.
(2006) found that VHTs play an essential role in thegenesis of a TC. The merging of multiple VHTs increases the cyclonic vorticity with a low-level vorticity
maximum and collective heating can spin up the stormscale circulation from the lower troposphere upward.
Tory et al. (2006a) explained that top-down and bottomup hypotheses may operate at two different phases in
the TC genesis: top down first during the preconditioning phase and then bottom up in the rapid intensification
stage. Earlier studies showed that the MCVs provide thenecessary condition for deep convection and initial concentration of cyclonic vorticity (e.g., Zhang and Fritsch
1987; Bartels and Maddox 1991; Zhang and Bao 1996a,b).
Thus, the preexisting MCVs may provide favorable environment for the genesis of VHTs and accelerate themerging process of neighboring VHTs. In a recent study,
Fang and Zhang (2010, 2011) simulated the initial development of Hurricane Dolly (2008) in a cloud-resolving
model. They found that the genesis of Hurricane Dolly
was essentially a bottom-up process.
The genesis is generally followed by a rapid intensification during which the storm becomes self-sustaining
and capable of intensifying through the system-scale
process, such as the wind-induced surface heat exchange(WISHE) mechanism (Ooyama 1969; Rotunno and
Emanuel 1987) and the system-scale intensification (SSI)
process (Shapiro and Willoughby 1982; Tory et al.
2006a,b, 2007). The SSI process describes the kinematics
of TC intensification in which the secondary circulationin response to heat and momentum sources intensifies
the slowly evolving primary circulation by bringing highangular momentum toward the inner-core region. This
process can be explained well by the balanced dynamics,
namely the Sawyer–Eliassen equation (e.g., Eliassen 1951;
Shapiro and Willoughby 1982; Schubert and Hack 1982;
Holland and Merrill 1984).
The evolution of the inner core and eyewall structure
is responsible for the intensity change of a TC (Wang
2002a,b; Wang and Wu 2004). The eyewall forms in an
intensifying TC as a convective ring, which coincides
with the strongest updrafts and the maximum tangential
winds. Although the mechanism for suppression of eyewall

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convection is complex (Frank and Ritchie 2001), strong
vertical shear of the environmental flow often has anegative effect on eyewall convection and leads to theweakening of a TC. Interaction between the eyewall andstrong spiral rainbands can lead to eyewall breakdown
and partial eyewall replacement accompanied by an intensity fluctuation of the TC (Wang 2002b). In somespecial cases, a new convective ring, or concentric eyewall, can form outside the primary eyewall. In such acase, as the outer convective ring propagates inward andamplifies; the inner, primary eyewall weakens and is
eventually replaced by the concentric eyewall (Willoughby
et al. 1982). The eyewall replacement is usually accompanied by a weakening and then a reintensification of
the TC, producing a large fluctuation in TC intensity.
Therefore, the eyewall process including its responseto any external forcing is crucial to the intensity change
of a TC.
The objectives of this study are 1) to describe themesoscale and system-scale processes in the life cycleof the simulated Isobel in NICAM, 2) to elucidate thephysical mechanisms that are responsible for the genesis
and intensity change of the simulated Isobel throughazimuthal-mean tangential wind budget analysis and thebalanced dynamics based on the Sawyer–Eliassen equation, and 3) to synthesize the multiscale interactions involved in the life cycle of the simulated Isobel based on
results from both Part I and this part. The rest of thepaper is organized as follows. The methodology for diagnostic and budget analyses is described in section 2.
Section 3 examines the mesoscale and system-scale processes leading to the genesis of Isobel. The dynamics
leading to the subsequent intensity change in Isobel are
discussed in section 4. The multiscale interactions in thelife cycle of Isobel are synthesized in section 5. The main
findings are summarized in the last section.

2. Methodology
Data used in the analyses of this study were the sameas those used in Part I, namely from the NICAM simulation with realistic settings at 7-km resolution for amonth in the austral summer of 2006, referred to as the
MJO experiment (Miura et al. 2007). Details of the simulated tropical storms in the MJO experiment can befound in Fudeyasu et al. (2008; Part I). In this study, theoutputs of the MJO experiment, including the snapshotsof three-dimensional data at 6-h intervals and 1.5-hourly
averaged two-dimensional data at the surface, were usedin the analysis.
The azimuthal-mean tangential wind budget was conducted by casting the governing equations into cylindricalcoordinates following the moving storm with the storm

motion vector subtracted from the total wind field. The
tendency equation for the azimuthal-mean tangential
wind can be written as
›y t
5
›t

yr h

y9r h9

w

›y t
›z

w9

›y9t
1 F sg ,›z(1)where z is height, t is time, y t is tangential wind, y r isradial wind, w is vertical velocity, h 5 f 1 z is absolute
vertical vorticity, z is relative vertical vorticity, f is theplanetary vorticity, and F sg is the azimuthal-mean tangential wind tendency due to vertical mixing and surface friction. The overbar denotes the azimuthal mean
at a given height, while the prime denotes the deviation (eddy) from the azimuthal mean. The five terms on
the right-hand side of Eq. (1) are radial advection of
the azimuthal-mean absolute angular momentum by
the azimuthal-mean radial wind (mean radial advection), eddy radial flux, vertical advection of azimuthalmean tangential wind by the azimuthal-mean verticalmotion (mean vertical advection), eddy vertical flux,
and surface friction and subgrid-scale vertical mixing.
The last term is treated as the residual in Eq. (1) in ourbudget.
To quantify the dynamical processes, we examinedthe evolution of Ertel’s PV defined as
PV 5h  $u,r(2)

where h is the three-dimensional absolute vorticity vector,
u is potential temperature, $u is the three-dimensional
gradient of potential temperature, and r is air density. Here PV is measured in PVU (1 PVU [ 1 31026 m2 s21 K kg21).
To investigate the SSI process, the Sawyer–Eliassen
equation in pseudoheight Z coordinates is used. The
equation can be found in Hendricks et al. (2004),
Montgomery et al. (2006), and Fudeyasu and Wang(2011). The forcing terms include momentum forcing
and diabatic heating:F 5 y9r z9
Q5

g
u0


y9r

w9

›y9
1 F sg ,
›Z

›u9
›r

w9

momentum forcing,›u9
1 u_ ,›Z

diabatic heating,

(3)

(4)

where terms in the momentum forcing represent theazimuthal-mean eddy radial/vertical flux of the asymmetric momentum and azimuthal-mean subgrid-scalevertical turbulent mixing, including surface friction. In
the diabatic heating, g is the acceleration due to gravity(9.81 m s22), u0 is the reference potential temperature

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MONTHLY WEATHER REVIEW(300 K), the first two terms represent the azimuthal-mean
eddy radial/vertical flux of the asymmetric potential
temperature, and u_ is the azimuthal mean diabatic heating rate. The diabatic heating rate u_ for each model grid
was calculated by u_ 5 ›u/›t 1 u  $u, where u is thethree-dimensional wind vector. In our calculations, if
the absolute vorticity h 5 f 0 1 z , 0.5f 0 , where f0 is
the planetary vorticity at the cyclone center, we set
h 5 0.5 f 0 to allow the discriminant (Fudeyasu and Wang2011) to satisfy the elliptical condition. This does notaffect the global solution to the Sawyer–Eliassen equation as we tested. The boundary conditions for the
Sawyer–Eliassen equation are the same as those ofHendricks et al. (2004) and Montgomery et al. (2006).
The computational domain is 0–15 km of pseudoheight inthe vertical with 500-m grid interval and 0–504 km in theradial direction with 7-km grid spacing. Variables fromthe Sawyer–Eliassen calculation were then converted
back to the physical height coordinate from the pseudoheight coordinate.

3. Genesis processes
a. Evolution of mesoscale vortices
Figure 1 shows the evolution of PV averaged in a layer
between 1 and 8 km above the sea surface and horizontal wind fields at 3-km height. On 25 December 2006,
there were little or no areas with cyclonic PV less than
21 PVU in the mid–lower troposphere over the JavaSea (note that cyclonic PV is negative in Southern
Hemisphere). During the period from 0000 to 1800 UTC
27 December 2006—that is, the preconditioning stage ofthe simulated Isobel (see Part I)—the WWB occurred
over the Java Sea associated with the onset of an MJO
event over the East Indian Ocean. The low-level largescale cyclonic vorticity and convergence together withhigh convective available potential energy (CAPE) over
the Maritime Continent triggered the development of aneast–west elongated rainband over the Java Sea (Figs. 5
and 8 in Part I). From 27 December, the cyclonic PV
anomalies with horizontal scales of 50–200 km started
to develop in the rainband (Fig. 1). The genesis anddevelopment of the cyclonic PV anomalies were associated with the MCSs embedded in the rainband.
Figure 2 shows the vertical structure of PV, relative
humidity, vertical velocity, and diabatic heating rate
zonally averaged between 1148 and 1178E at 0000 UTC
28 December 2006 over the Java Sea. The strong updraftwith high relative humidity throughout the depth of thetroposphere occurred around 4.58–5.58S (Figs. 2b,c) near
the cyclonic shear line of the low-level wind field (notshown). The diabatic heating rate with its maximum at

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a height of about 8–10 km was associated with deep
convection, whereas that around 4–6-km height near
the melting level was mainly associated with stratiform clouds (Fig. 2d). The cyclonic PV anomalies at the
2–5-km heights were indications of the midtroposphericvortex development related to diabatic heating in thestratiform clouds (Fig. 2a), which is consistent withthe finding of Raymond and Jiang (1990). We refer tothe midtropospheric vortex with horizontal scale morethan 50 km as model MCV. During the initial eddy stage
from 1800 UTC 27 December to 1800 UTC 28 December,the model MCVs moved southeastward along with therainband toward the Lesser Sunda Islands (Fig. 1). These
MCVs merged and eventually developed into a concentric MCV with the intensified cyclonic PV anomaly nearits center, the precursor of the genesis of Isobel.
Figure 3 shows the time–radius cross section of theazimuthal-mean PV averaged between 3–6-km heights
and the time–height cross sections of PV averaged
within a radius of 301 km following Isobel’s center. Notethat before the initial eddy stage the center was fixed as
the center of Isobel at 1800 UTC 27 December. The
cyclonic PV anomaly appeared first in the mid–lower
troposphere between 2- and 5-km heights in the preconditioning stage (Fig. 3b). The organization of MCVs
around Isobel’s center was accompanied by cyclonic PV
less than 21 PVU with a horizontal scale over 30 km in
radius at 1800 UTC 27 December (Fig. 3a). The radiusof cyclonic PV increased to 90 km in the early initialeddy stage, but decreased to about 60 km as the cyclonic
PV intensified to 25 PVU at the center of the initialeddies in the late initial eddy stage, indicating the contraction of the MCV.
Figure 4 shows a plane view of PV averaged in themid–lower troposphere between 1- and 6-km heights,
the wind fields at 3-km height, and the mid- and lowerlevel circulation centers (see Part I). By the end of theinitial eddy stage, the circulation center tilted toward thenorthwest and west with height, which is consistent withthe strong easterly vertical shear (Figs. 8c and 10 in Part I).
At 0000 UTC 28 December, a few strong cyclonic PV
anomalies with central values less than 24 PVU andradii less than 40 km were embedded in the MCVs. The
centers of these small-scale vortices did not coincidewith the circulation centers at that time. Figure 5 shows
the vertical structure of PV anomalies along the segment
A–C labeled at 0000 UTC 28 December in Fig. 4. The
PV anomalies marked A and B in Fig. 4 had high cyclonic PV values less than 29 and 26 PVU, respectively,
up to 11-km height, whereas the PV anomaly C had highcyclonic PV in the lower troposphere (Fig. 5a). The PV
anomalies A and B were accompanied by strong updrafts (Fig. 5c). Strong updraft southeast of the PV

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FIG. 1. PVU (shaded) averaged in the layer between 1- and 8-km heights above the sea surface and horizontal winds(vector) at 3-km height. Regions with cyclonic vorticity less than 21 PVU are shaded. The focused views on the innercore PV are shown at 0600 and 1200 UTC 31 Dec 2006, 1800 UTC 1 Jan 2007, and 0000 and 0600 UTC 2 Jan 2007.

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FIG. 2. Vertical–meridional cross sections of (a) PVU, (b) relative humidity (%), (c) vertical velocity (m s21), and(d) diabatic heating rate (1023 K s21) zonally averaged between 1148 and 1178E at 0000 UTC 28 Dec 2006. In (a), thecontour interval is 0.3 PVU, and regions less than 20.3 PVU (20.9 PVU) are lightly (heavily) shaded. In (b), thecontour interval is 10%, and the regions greater than 70% (90%) are lightly (heavily) shaded. In (c), the contour
interval is 0.1 m s21 and regions greater than 0.1 m s21 (less than 0 m s21) are heavily (lightly) shaded. In (d), thecontour interval is 1.0 3 1023 K s21 and regions greater than 0.1 3 1023 K s21 (2.0 3 1023 K s21) are lightly(heavily) shaded.

anomaly C was associated with the PV anomaly D (Fig. 4),
which tilted northward with height. Warm anomaliesof about 4 K (defined as the temperature deviationfrom the reference temperatures averaged at the sameheight in a region of 308S–08, 808–1408E at 0000 UTC
23 December 2006) occurred around 10-km height, an
indication of active convection with each cyclonic PV
anomaly (Fig. 5e). The warm anomalies, however, were
too weak to have signals in the sea level pressure (SLP)
field (Fig. 5g).
These deep cyclonic PV anomalies with a horizontal
scale of about 40 km are associated with deep convection, very similar to those associated with VHTs as discussed in Hendricks et al. (2004) and Montgomery et al.
(2006). However, the horizontal scale of the vortices inthe simulation is larger than VHTs derived from observations (Reasor et al. 2005; Sippel et al. 2006), mainly
because of the relatively coarse model resolution. In ourfollowing discussions, we refer to the deep, cyclonic PV

anomalies embedded in model MCVs as model VHTs orconvective PV anomalies because they exhibited vortex
interactions very similar to the observed VHTs in theinitial eddy stage of the simulated Isobel.
The multiple model VHTs were identifiable in themodel MCVs until 0600 UTC 28 December, while their
number decreased by the end of the initial eddy stage(Fig. 4). A monolithic PV core with peak cyclonic PV
less than 214 PVU formed near the center of the concentric MCV at 1800 UTC 28 December. Because thelifetime of convective clouds is generally less than 1–2 h,each VHT could not be tracked with the 6-hourly output of the MJO experiment. However, surface relative
vorticity and rainfall derived from 1.5-hourly output
showed the merging of VHTs in the organized convective rainband (not shown). The merging of multiple
cyclonic PV anomalies formed a monolithic PV core,
which is consistent with the finding of Tory et al. (2006b,
2007).

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FIG. 3. (a) Time evolution of the azimuthal-mean PVU as a function of radius averaged in a layer between 3- and
6-km heights with contour interval of 1 PVU and regions less than 21 PVU (26 PVU) lightly (heavily) shaded.
(b) Time–height cross section of PVU averaged within a radius of 301 km from Isobel’s center with contour intervalof 0.1 PVU and regions less than 20.2 PVU (20.8 PVU) lightly (heavily) shaded.

Figure 5b shows the vertical structure of the monolithic PV core with a peak cyclonic PV of 219 PVU. The
monolithic PV core was accompanied by strong updrafts
over 8 m s21 (Fig. 5d). Although the peak cyclonic
PV anomaly of the monolithic PV core at 1800 UTC
28 December was about 2–3 times that of previous
VHTs at 0000 UTC 28 December, the horizontal scale ofthe monolithic PV core remained small and was similar
to those of previous VHTs. The warm anomalies
intensified in the upper troposphere with the maximum
temperature anomaly of over 8 K and expanded radiallyoutward (Fig. 5f), resulting in low pressure at the surface (Fig. 5h). Note that the center of the monolithic PV
core now coincided with the circulation center (Fig. 4)
and a vertically aligned, deep cyclonic circulation developed. The system thus possessed the typical structure ofa TC, namely the genesis of Tropical Storm Isobel.

b. Storm-scale process
To evaluate the physical mechanism responsible forthe genesis of the simulated Isobel in the initial eddy
stage, we conducted the budget and diagnostic analyses
for two times: 0000 UTC 28 December (the early initial eddy stage) when multiple VHTs were active, and1800 UTC 28 December (the late initial eddy stage) after
the formation of monolithic PV core. Figures 6 and 7
show the azimuthal-mean radial and tangential winds,vertical velocity, the time tendency in tangential winds,
and terms that contribute to the azimuthal-mean tangential wind budget in Eq. (1) for the two times,

respectively. The storm centers used in the calculation forsymmetric/asymmetric components were defined as the
850-hPa circulation center at 0000 UTC 28 December andthe mass center (see Part I) at 1800 UTC 28 December(also in Figs. 8–12).
In the early initial eddy stage, the weak azimuthalmean inflow occurred in the midtroposphere with thepeak at 4-km height between radii of 150–300 km andweak outflow occurred in a relatively deep layer in theupper troposphere (Fig. 6a). Weak cyclonic tangential
winds extending to large radii appeared below 7-km
height (Fig. 6c) and weak updrafts occurred between
radii of 30–50 km in the mid–upper troposphere (Fig. 6b).
In the late initial eddy stage, strong updrafts occurred
throughout the depth of the troposphere within a radiusof about 30 km from the center (Fig. 7b). Tangential
winds more than 8 m s21 occurred in the lower troposphere extending radially outward with the maximum inthe boundary layer between the radii of 15 and 50 km(Fig. 7c). There were two inflow layers: a strong one in
the boundary layer below 2 km and a weak one in the
midtroposphere with a weak outflow in between thatpeaked at around 4-km height (Fig. 7a).
The azimuthal-mean tangential wind increased withtime throughout the depth of the troposphere withina radius of about 250 km with the maximum increase in
the mid–lower and upper troposphere between the radiiof 20 and 100 km in the early initial eddy stage (Fig. 6d).
The spinup of the azimuthal-mean tangential wind resulted from the mean radial/vertical advections of absolute angular momentum in the mid–lower troposphere and4312

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FIG. 4. As in Fig. 1, but for PVU (shaded) averaged in the layer between 1- and 6-km heights with different shading
scales shown at the bottom of the lower-right. Cross, circle, and square represent the circulation centers at 850, 700,
and 500 hPa, respectively.

the eddy fluxes in both the upper and lower troposphere(Figs. 6e,f). In the late initial eddy stage, areas with the
spinup of tangential wind expanded to large radii (Fig. 7d).
The maximum increase occurred from the surface up to
12-km height in the inner-core region and was mainly
contributed to by the mean advections (Fig. 7e). Contribution by the eddy process to the azimuthal-mean tangential wind budget was generally small in the late initialeddy stage (Fig. 7f).

Since the mean advection associated with the azimuthalmean radial and vertical winds (secondary circulation) is
responsible for the spinup of the azimuthal-mean primary circulation, it is our intent to examine what contributed to the azimuthal-mean secondary circulation
during the initial eddy stage of the simulated Isobel.
Figures 8a,b show the radius–vertical cross section of
the azimuthal-mean diabatic heating rate (the dominant
forcing term in the Sawyer–Eliassen equation, as will beDECEMBER 2010

FUDEYASU ET AL.

FIG. 5. Vertical cross sections of (a),(b) PVU; (c),(d) vertical velocity (m s21); (e),(f) warm anomalies (K); and(g),(h) sea level pressure (hPa) at (left) 0000 and (right) 1800 UTC 28 Dec 2006. Locations of the cross sections aremarked as line segments in Fig. 4. The contour interval is 1 PVU and regions less than 21 PVU (25 PVU) are lightly(heavily) shaded in (a),(b). The contour interval is 0.5 m s21 and regions more than 0.5 m s21 (2.0 m s21) are lightly(heavily) shaded in (c),(d). The contour interval is 1 K and the regions more than 4 K (8 K) are lightly (heavily)
shaded in (e),(f). The warm anomaly was calculated as the difference of potential temperature from the reference
potential temperature averaged in 308S–08 and 808–1408E at 0000 UTC 23 Dec 2006.

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FIG. 6. Azimuthal-mean (a) radial wind (m s21), (b) vertical velocity (m s21), (c) tangential wind (m s21), (d) thetime tendency in tangential wind (m s22) estimated based on term on left-hand side of Eq. (1), (e) mean radial andvertical advections, and (f) eddy radial and vertical fluxes estimated based on terms on right-hand side of Eq. (1) at
0000 UTC 28 Dec 2006. The contours are at 0, 61, 62, 65, and 610 m s21 and regions greater than 1 m s21 (lessthan 1 m s21) are heavily (lightly) shaded in (a). The contour interval is 0.2 m s21 and regions greater than
0.2 m s21 (less than 20.05 m s21) are heavily (lightly) shaded in (b). The contour interval is 2 m s21 and regions
greater than 2 m s21 (10 m s21) are lightly (heavily) shaded in (c). The contour interval is 1 3 1024 m s22 andregions greater than 0 m s22 (2 3 1024 m s22) are lightly (heavily) shaded in (d)–(f). Solid contours denotepositive values and dashed contours, negative values.

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FIG. 7. As in Fig. 6, but at 1800 UTC 28 Dec 2006.

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FIG. 8. Azimuthal-mean diabatic heating rate (K s21) at (a) 0000 UTC 28 Dec, (b) 1800 UTC 28 Dec, (c) 1800 UTC
30 Dec, (d) 1200 UTC 31 Dec 2006, (e) 0600 UTC 2 Jan, and (f) 0000 UTC 3 Jan 2007. The contour interval is 1 31024 K s21. Regions greater than 4 3 1024 K s21 are filled and regions greater than 2 3 1024 K s21 are heavily
shaded, while regions greater than 1 3 1024 K s21 are lightly shaded. Dashed horizontal lines represent the height of
6 km used in the sensitivity calculations in (e).

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discussed later) in the early and late initial eddy stages,
respectively. The azimuthal-mean diabatic heating rate
was negligible in the early initial eddy stage (Fig. 8a), butbecame large within a radius of 30 km in the late initialeddy stage (Fig. 8b). The latter was associated withconvective heating in the strong updraft in the monolithic PV core (Fig. 5). The secondary circulation diagnosed from the Sawyer–Eliassen equation for the early
initial eddy stage was quite different from the model
simulation (not shown), indicating that the azimuthalmean secondary circulation at the early initial eddy stagecannot be explained by the balanced dynamics. On theother hand, the azimuthal-mean secondary circulation
for the late initial eddy stage was reproduced reasonably
well using the Sawyer–Eliassen equation (Figs. 7a,b and9a,b), although considerable differences exist in the region outside about 200–300-km radii and in the upper
troposphere. This is consistent with Hendricks et al.
(2004) and Montgomery et al. (2006), who also showedthat the storm-scale secondary circulation in the genesisstage was nearly in both gradient wind and thermal windbalances and could be well explained by the balanced
dynamics using the Sawyer–Eliassen equation.
Since the Sawyer–Eliassen equation is linear, its solution is additive. This allows us to understand which
forcing process is responsible for the secondary circulation at the late initial eddy stage. We thus performed
two calculations: in one calculation only the azimuthalmean diabatic heating forcing term in the last term on
the right-hand side of Eq. (4) (see Fig. 8) was included
in solving the Sawyer–Eliassen equation and in the othercalculation only the eddy terms, namely the first andthe second terms in Eqs. (3) and (4), were included. The
results showed that the secondary circulation at the lateinitial eddy stage was mostly driven by the azimuthalmean diabatic heating (not shown). That is, as a monolithic PV core with warm core developed at the circulation
center, the secondary circulation was enhanced by convective heating in the monolithic PV core, contributing
to the intensification of the primary circulation.

4. Intensity change
After Isobel passed over the Lesser Sunda Islands
at 1800 UTC 28 December, the concentric MCV withthe monolithic PV core developed into Tropical Storm
Isobel over the Timor Sea, which then intensified from1800 UTC 28 December to 0600 UTC 31 December(Fig. 1). The azimuthal-mean cyclonic PV averaged between 3- and 6-km heights increased up to 220 PVU in
the inner core of the simulated Isobel (Fig. 3a). The
cyclonic PV less than 21 PVU extended radially outward up to 150 km from the storm center, indicating an

4317

increase in the horizontal scale of the simulated Isobel.
The outward expansion of high cyclonic PV was closely
related to the formation of an outer spiral rainband(Figs. 5 and 12 in Part I). By 1800 UTC 30 December thesimulated Isobel showed the typical structure of a mature TC with the maximum tangential wind near thesurface under the eyewall (Fig. 10c), updrafts in theeyewall and outer spiral rainbands, subsidence in the eye(Fig. 10b), inflow in the mid–lower troposphere andoutflow both in the core just above the inflow boundary
layer and in the upper troposphere (Fig. 10a). Convective heating occurred through the depth of the troposphere between radii of 20 and 40 km in the eyewall andin the mid–upper troposphere in outer rainbands between radii of 100 and 250 km (Fig. 8c).
The intensification of the primary circulation occurred
through the depth of the troposphere but most strongly
in the inner-core region in the intensifying stage (Fig. 10d).
The inner-core intensification was mainly attributed tothe mean advections in the lower and upper troposphereand to eddy fluxes in the midtroposphere (Figs. 10e,f). As
the case in the later initial eddy stage, now the azimuthalmean radial and vertical winds were captured well by
the solution of the Sawyer–Eliassen equation althoughsome noise existed in the upper-tropospheric outflow(Figs. 9c,d). Therefore, the intensification of the primarycirculation in the inner-core region can be well explainedby the balanced dynamics, namely the SSI process.
In the temporally weakening stage from 0600 UTC
31 December 2006 to 1800 UTC 1 January 2007, theintensification of the simulated Isobel was interrupted by
the influences of strong vertical shear first and a strong
environmental stretching deformation in the mid–lower
troposphere later as discussed in Part I. At 0600 UTC
31 December by the end of the intensifying stage, asymmetric PV monopole structure in the inner-core region appeared with a peak of 214 PVU (Fig. 1). Isobel
experienced an abrupt structure change at 1200 UTC
31 December with the central PV monopole split into
two areas: one to the west and one to the east of Isobel’s
center. The asymmetric structure of the inner-core PV
remained until 1200 UTC 1 January 2007. The decay ofthe symmetric structure in Isobel during the early temporally weakening stage was accompanied with the westward tilt of the vertical vorticity center due to the effectof strong easterly vertical shear as discussed in Part I,
which is similar to the PV pattern in the case of TC Chris
discussed in Tory et al. (2006b).
It is interesting to see that the weakening of the primary circulation started from the upper troposphere in
the core region (Figs. 3b and 11c), consistent with thefinding of Frank and Ritchie (2001), who found that
their simulated storm weakened from the top down in4318

MONTHLY WEATHER REVIEWVOLUME 138

FIG. 9. Azimuthal-mean (a),(c),(e) radial winds (m s21) and (b),(d),(f) vertical velocity (m s21) derived from the
Sawyer–Eliassen equation at (a),(b) 1800 UTC 28 Dec 2006 in the same domain as in Fig. 7, at (c),(d) 1800 UTC
30 Dec 2006 in the same domain as in Fig. 10, and at (e),(f) 0600 UTC 2 Jan 2007 in the same domain as in Fig. 12.
Contours are at 0, 61, 62, 65, and 610 m s21 and regions greater than 1 m s21 (less than 1 m s21) are heavily(lightly) shaded in (a), (c), and (e). The contour interval is 0.2 m s21 and regions greater than 0.2 m s21 (less than
20.05 m s21) are heavily (lightly) shaded in (b),(d), and (f). Solid contours denote positive values; dashed contours,
negative values.

DECEMBER 2010

FUDEYASU ET AL.

4319

FIG. 10. As in Fig. 6, but at 1800 UTC 30 Dec 2006.

a vertical shear environment. The radial wind was exhibited by a deep outflow layer up to a radius of about
350 km between 2- and 4-km heights with a deep inflow
layer between 5- and 12-km heights immediately below

the upper-tropospheric outflow (Fig. 11a). The mid–
upper-tropospheric inflow implied a convergence into
the eye and eyewall region, consistent with a subsidence
in the inner core below about a 10-km height (Fig. 11b).

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MONTHLY WEATHER REVIEWVOLUME 138

FIG. 11. As in Fig. 6, but at 1200 UTC 31 Dec 2006.

Such a secondary circulation led to a spindown of
the tangential wind in the mid–lower troposphere anda spinup of the tangential wind in the mid–uppertropospheric inflow layer mainly outside the inner-core
region (Figs. 11d,e). Eddies played a role in spinning

down the primary circulation except for in the lower
troposphere and in the core region (Fig. 11f). As a result,the azimuthal-mean tangential wind decreased in a deep
layer below 12-km height within a radius of about 100 km
except for a small area in the eye between 1- and 8-km

DECEMBER 2010

FUDEYASU ET AL.

heights (Fig. 11d), an indication of inward mixing byboth the azimuthal-mean inflow and eddy transport (Figs.
11e,f). Consistent with the inflow immediately below
the upper-tropospheric outflow, there was a spinup oftangential wind outside the core region between 7- and
12-km heights (Fig. 11d). Note that the azimuthal-meandiabatic heating rate in the inner core of the storm wasalmost dismissed in the weakening stage (Fig. 8d). The
SSI process became ineffective and the storm weakened.
Later on, the simulated Isobel experienced an eyewall
reformation as discussed in Part I. This led to a reintensification of the storm from 1800 UTC 1 January to1800 UTC 2 January 2007. At 1800 UTC 1 January, ahigh cyclonic PV band was initiated outside the eyewall(Fig. 1) in an outer rainband south of Isobel’s center(Fig. 5 in Part I). The cyclonic PV band associated withthe rainband displayed a hooklike structure extending
southeastward from the core of Isobel (Fig. 1). The PVband then spiraled cyclonically inward, merged into theinner-core PV from the north by 0000 UTC 2 January,and enhanced the cyclonic PV in the inner core. This was
accompanied with a reformation of the quasi-symmetric
PV monopole structure with a peak cyclonic PV lessthan 224 PVU by 0600 UTC 2 January. This mesoscaleintensification process through the inward transport of
the rainband-induced cyclonic PV anomalies supports
the hypothesis of May and Holland (1999). They estimated diabatic heating and PV generation in TC rainbands based on the observed vertical motion profiles
and found high levels of PV generation in the stratiform
rain regions. They suggested that such PV generation
represents a significant PV source for the system as awhole as it is transported into the core and thus has
implications for TC.
After the recovery of the quasi-symmetric structure of
the inner-core PV, the simulated Isobel reintensified
and reached its lifetime maximum intensity by 0600 UTC
2 January. Again the storm gained the typical structureof a mature TC (Fig. 12), similar to the one during theintensifying stage discussed earlier (Fig. 10), but thestorm now developed a deeper cyclonic circulation(Fig. 12c) and a larger horizontal size with a wider
eyewall updraft (Fig. 12b). Accordingly, the azimuthalmean diabatic heating rate in the eyewall occurred
between radii of 30 and 60 km and in spiral rainbands
around a radius of 100 km (Fig. 8e). The azimuthalmean tangential wind tendency showed a large spinup
in the eyewall in the mid–lower troposphere and immediately outside the eyewall in the mid–upper troposphere while a spindown in the eye region in the upper
troposphere occurred (Fig. 12d), an indication of thedevelopment of an outward tilting of the radius of maximum wind (Fig. 12c).

4321

The spinup of the primary circulation was mainly
contributed by the mean transport of absolute angularmomentum in the mid–lower troposphere in the innercore region and in the eyewall (Fig. 12e) and by the eddy
fluxes in the midtroposphere between radii of 50 and
150 km (Fig. 12f). The mean transport also contributed
to the spinup of the tangential wind outer core with amaximum contribution between 7- and 10-km heights(Fig. 12e). Note that the eddy process contributed mainly
to the tangential wind budget negatively (Fig. 12f) andthus played a role in weakening the storm intensity,
consistent with the idealized simulation results of Yanget al. (2007). The Sawyer–Eliassen equation reproduced
the secondary circulation well in the reintensifying stage(Figs. 9e,f), indicating that the reintensification of thesimulated Isobel after the reformation of the eyewall
and the recovery of the eyewall convection can be explained by the balanced dynamics, namely the SSI process again.
The inner-core cyclonic PV decreased after 1800 UTC
2 January as Isobel approached the northwest coast ofAustralia and continued to decay (Figs. 1 and 3). The
azimuthal-mean diabatic heating rate became negligible
in the inner-core region (Fig. 8f) and the weakening of
the tangential winds started from the inner-core region(not shown). These changes overall are very similar to
those in the temporary weakening stage.

5. A synthesis of the multiple-scale interaction
in Isobel’s life cycle
This study (i.e., Part I and Part II) documented boththe external influence of the large-scale environmental
flow on the structure and intensity changes and the involved mesoscale and system-scale processes in the life
cycle of the simulated Isobel, as synthesized schematically in Figs. 13 and 14. The WWB associated with theonset of an MJO event over the Indian Ocean provided
a favorable large-scale environmental condition for thegenesis of Tropical Storm Isobel over the Java Sea(Fig. 13). This preconditioned environment was featured
by strong large-scale cyclonic vorticity and convergencein the mid–lower troposphere and high CAPE. Convection was initiated and organized in a zonally elongatedrainband with embedded MCSs near the cyclonic shear
line in the preconditioned large-scale environment.
Some model MCVs with vorticity maxima in the
midtroposphere developed in association with diabatic
heating in stratiform clouds (Fig. 14a), consistent withprevious finding of Raymond and Jiang (1990). Convective towers with cyclonic PV anomalies throughout
the depth of the troposphere were initiated and developed in the MCVs (Fig. 14b). This featured the initial

4322

MONTHLY WEATHER REVIEWVOLUME 138

FIG. 12. As in Fig. 6, but at 0600 UTC 2 Jan 2007.

eddy stage of the simulated Isobel. Although it is difficult for us to examine the formation mechanisms ofthe convective vortices with a 7-km model resolution,
the dipole PV structure with positive and negative PV

anomalies around the convective updraft in VHTs embedded in the merging MCVs (Fig. 5a) implies the importance of tilting of horizontal vorticity to the formationof VHTs (Montgomery et al. 2006). The MCVs provided

DECEMBER 2010

FUDEYASU ET AL.

4323

FIG. 13. Schematic diagram showing the vorticity evolution and the associated environmental
flow on the (left) synoptic scale, (middle) system scale, and (right) mesoscale in the inner PV
core. Regions with high cyclonic vorticity of 210 PVU are filled and regions with cyclonicvorticity of 24 PVU are heavily shaded, while regions with weak cyclonic vorticity of 21 PVU
are lightly shaded. Solid arrows indicate the lower-level flow, while broken arrows show theupper-level flow.

necessary conditions for the formation of VHTs (Zhang
and Fritsch 1987; Bartels and Maddox 1991; Zhang and
Bao 1996a,b).
Multiple model VHTs merged and strengthened cyclonic PV anomaly in the interior of the merging MCV
and led to the formation of an upright monolithic PV
core at the center of concentric MCV (primary vortex
enhancement; Fig. 14c), as discussed in Hendricks et al.

(2004), Montgomery et al. (2006), and Tory et al. (2006b).
As the monolithic PV core with a warm core developed
near the circulation center, the cyclonic PV was enhanced by the SSI process driven by diabatic heating in
the inner-core region (secondary vortex enhancement;
Fig. 14d). This was accompanied by the intensificationand the increase in horizontal size of the cyclonic circulation, leading to the formation of Tropical Storm Isobel.

4324

MONTHLY WEATHER REVIEWVOLUME 138

FIG. 14. Schematic diagram of the vertical cross section showing the genesis processes of (a) large-scale preconditioning, (b) meso/small-scale convective vortices evolution, (c) primary vortex enhancement, and (d) secondary
vortex enhancement. Regions with significant diabatic heating are shaded, and cloud regions are schematicallydrawn. In (a)–(c), opened arrows indicate the MCV, while solid arrows show the convective vortex (VHT). In (d),opened arrows show the system-scale primary circulation, while solid arrows indicate the system-scale secondary
circulation in response of the formed warm core. Note that the Northern Hemisphere motion is used for cyclonicrotation in the figure while Tropical Storm Isobel is a Southern Hemispheric tropical cyclone.

Although the azimuthal-mean secondary circulation
contributed largely to the spinup of the primary circulation in the formation stage, eddy process also played
a positive role in the upper and lower troposphere during the early genesis stage (Fig. 6f), consistent with theresults of Hendricks et al. (2004).
In the intensifying stage, Isobel moved into the Timor
Sea and continued to intensify under the favorable largescale environment with weak vertical shear and strong
low-level cyclonic shear (Fig. 13). The high cyclonic PV
core intensified and became axisymmetric, and the eyewall formed with enhanced convective heating. The secondary circulation in response to diabatic heating in
the eyewall accelerated the azimuthal-mean primarycirculation through inward transport of absolute angularmomentum. This intensification process (SSI) can be explained well by the balanced dynamics based on the

Sawyer–Eliassen equation with the eddy processes playing some secondary roles.
As the simulated Isobel moved southward and the
MJO propagated eastward, Isobel was exposed to theregion first with a strong vertical shear and later with astrong stretching deformation field upwind of the WWB
in the MJO. This change in large-scale environmental
flow forced convective asymmetries in the inner core of
Isobel, leading to a breakdown of the eyewall and atemporary weakening of the storm (the temporary weakening stage). The weakening of the primary circulation
occurred mainly in the inner core and started from the
top down, consistent with previous results from idealized simulations in Frank and Ritchie (2001). This spindown process was related to the strengthening and inward
penetration of an inflow immediately below the uppertropospheric outflow, the azimuthal-mean subsidence inDECEMBER 2010

FUDEYASU ET AL.

the inner-core region, and a deep outflow layer in the mid–lower troposphere above the inflow boundary layer. As
convection in the eyewall was suppressed, the azimuthalmean diabatic heating rate became negligible in the innercore region, and the SSI process thus became ineffective
and the storm weakened.
In the late temporally weakening stage, as the storm
moved southeastward, the vertical shear changed fromeasterly shear to north-northeasterly shear over the inner core, leading to the development and intensificationof a strong outer spiral rainband to the south-southeastof the storm center, namely downwind of the WWB
center. The outer rainband propagated cyclonically inward and axisymmetrized, leading to the reformationof the eyewall and the reintensification of the simulated
Isobel. Therefore, the eyewall reformation as a result of
the axisymmetrization of the inward-propagating outer
spiral rainbands was an important process to the reintensification of the simulated Isobel. The eyewall reformation
was accompanied by recovery of a quasi-axisymmetricinner-core structure with enhanced azimuthal-mean diabatic heating in the eyewall, allowing the SSI process tooperate effectively and the reintensification of the simulated Isobel. Note that although the axisymmetrization
contributed to the reintensification, the latter might also
be speeded up the axisymmetrization process of the inward-propagating spiral rainband and the reformationof the eyewall. As the simulated Isobel moved farther
southeastward and approached the northwest coast ofAustralia, it weakened because of the continued decreasein moisture supply from the underlying ocean surface
and eventually it dissipated after it made landfall overnorthwest Australia.

6. Conclusions
The life cycle of Tropical Storm Isobel was reasonably
simulated in the global cloud-system-resolving model,
NICAM. In this study, we have investigated both thelarge-scale and storm-scale evolutions (Part I) and thesystem-scale and mesoscale processes (Part II) in the life
cycle of the simulated Isobel. As we have synthesized in
the last section, the life cycle of the simulated Isobel involved multiscale interactions from cloud-scale convective
activities to large-scale circulation related to the onset of
MJO and the associated WWB. We show that the largescale environmental flow provided the preconditioned
environment for the genesis of the simulated Isobel andtriggered the subsequent structure change of the storm.
The mesoscale and the system-scale processes, such as theevolution of MCVs and merging VHTs were responsible
for the genesis, while the eyewall processes were critical to
the storm intensity change through the SSI process.

4325

In the special case of Isobel, we have found that the
WWB associated with the onset of a MJO event played
dual roles in the life cycle of the simulated Isobel. On
one hand, the WWB contributed positively to the intensification of Isobel as Isobel was located southeast of
the WWB center where large-scale convergence occurred with high CAPE. On the other hand it contributed to the development of the asymmetric structure in
the inner core and thus negatively to Isobel’s intensification as Isobel was located southwest of the WWB
center where the environmental stretching deformation
was strong. Therefore, depending on the relative location of Isobel and the WWB center associated with the
MJO, the MJO had both positive and negative effects on
the intensity of the simulated Isobel. We thus suggestthat caution needs to be given to the relative location
between the WWB and the TC when the possible effectof the MJO on a TC is discussed.
In a simulation study of real TCs, Tory et al. (2006b)
considered two processes for the genesis: the primary vortex enhancement through the convective vortex enhancement and the formation of PV core in convective updrafts(Hendricks et al. 2004; Montgomery et al. 2006) and thesecondary vortex enhancement through the SSI process.
Our results also demonstrate that both mesoscale andsystem-scale processes are critical to the genesis of thesimulated Isobel in a preconditioned large-scale environment. We also show that the SSI process operates
effectively only when the storm intensifies with intense
convection in the inner-core region. The SSI process becomes ineffective when there is little azimuthal-mean diabatic heating in the inner core. This suggests thatspecial attention should be paid to the eyewall process inorder to understand and predict the structure and intensity changes of a TC.
Note that our results are based on NICAM MJO experiment at 7-km resolution, which could only marginally capture the important mesoscale processes in thelife cycle of Isobel. This resolution is not high enough to
examine the life cycle of individual VHTs and the statistical behavior of VHTs. Although now NICAM canbe run at 3.5-km resolution for monthly long simulations, it is still not high enough to study convection indetails. Therefore, studies with regional fine-resolution
cloud-resolving models are still important for our understanding of the life cycle of TCs. Nevertheless, ourcase study for Isobel based on the NICAM simulation
demonstrates the potential use of nearly cloud-resolving
global models to study the life cycle of TCs.
In addition, as we pointed out already in Part I, although the NICAM MJO experiment provided us an
opportunity to study the multiscale interactions involvedin the life cycle of the simulated Isobel, the results do not4326

MONTHLY WEATHER REVIEW

imply skill at monthly-long forecasting of TC behaviors.
The timing and structure and intensity change werelargely controlled by the large-scale environmental settings associated with the onset of an MJO event. Realistic simulation and prediction of the MJO are stillchallenging to most regional and global models (Sperberand Waliser 2008; Kim et al. 2009). Nevertheless, global
cloud-system-resolving models have recently shown consistent improvements in the simulation of the MJO(Miura et al. 2007; Liu et al. 2009). Therefore, the extended realistic simulation and skillful prediction of TCs
by numerical models are expected to be improved alongwith the continuous improvements in the ability of global
models in simulating and predicting the tropical variability of convective activities at intraseasonal timescales.
Acknowledgments. The authors are grateful to theanonymous reviewers for their helpful comments. This
study was supported in part by NSF Grant ATM0754039 and in part by JAMSTEC, NASA, and NOAA
through their sponsorships of the International PacificResearch Center (IPRC) in the School of Ocean andEarth Science and Technology (SOEST) at the University of Hawaii. Part of this study was also supported by
the JST/CREST in Japan. The Earth Simulator wasused for the NICAM simulations.

REFERENCES
Bartels, D. L., and R. A. Maddox, 1991: Midlevel cyclonic vortices
generated by mesoscale convective systems. Mon. Wea. Rev.,
119, 104–118.
Bister, M., and K. A. Emanuel, 1997: The genesis of Hurricane
Guillermo: TEXMEX analyses and a modeling study. Mon.
Wea. Rev., 125, 2662–2682.
Eliassen, A., 1951: Slow thermally or frictionally controlled meridional circulation in a circular vortex. Astrophys. Norv., 5,
19–60.
Fang, J., and F. Zhang, 2010: Initial development and genesis of
Hurricane Dolly (2008). J. Atmos. Sci., 67, 655–672.
——, and ——, 2011: Evolution of multiscale vortices in the development of Hurricane Dolly (2008). J. Atmos. Sci., 68, in press.
Frank, W. M., and E. A. Ritchie, 2001: Effects of vertical wind
shear on the intensity and structure of numerically simulated
hurricanes. Mon. Wea. Rev., 129, 2249–2269.
Fudeyasu, H., and Y. Wang, 2011: Contribution by balanced dynamics to the intensification of a tropical cyclone simulated in
TCM4: Outer core spinup process. J. Atmos. Sci., in press.
——, ——, M. Satoh, T. Nasuno, H. Miura, and W. Yanase, 2008:
Global cloud-system-resolving model NICAM successfully simulated the lifecycles of two real tropical cyclones. Geophys. Res.
Lett., 35, L22808, doi:10.1029/2008GL036003.
——, ——, ——, ——, ——, and ——, 2010: Multiscale interactions in the life cycle of a tropical cyclone simulated in a
global cloud-system-resolving model. Part I: Large-scale andstorm-scale evolutions. Mon. Wea. Rev., 138, 4285–4304.

VOLUME 138

Hendricks, E. A., M. T. Montgomery, and C. A. Davis, 2004: The
role of ‘‘vortical’’ hot towers in the formation of Tropical
Cyclone Diana (1984). J. Atmos. Sci., 61, 1209–1232.
Holland, G. J., and R. T. Merrill, 1984: On the dynamics of tropical
cyclone structural changes. Quart. J. Roy. Meteor. Soc., 110,
723–745.
Kim, D., and Coauthors, 2009: Application of MJO simulation
diagnostics to climate models. J. Climate, 22, 6413–6436.
Liu, P., and Coauthors, 2009: An MJO simulated by the NICAM at
14- and 7-km resolutions. Mon. Wea. Rev., 137, 3254–3268.
May, P. T., and G. J. Holland, 1999: The role of potential vorticity
generation in tropical cyclone rainbands. J. Atmos. Sci., 56,
1224–1228.
Miura, H., M. Satoh, T. Nasuno, A. T. Noda, and K. Oouchi, 2007:
A Madden–Julian Oscillation event realistically simulated bya global cloud-resolving model. Science, 318, 1763–1765.
Montgomery, M. T., M. E. Nicholls, T. A. Cram, and A. B. Saunders,
2006: A vortical hot tower route to tropical cyclogenesis.
J. Atmos. Sci., 63, 355–386.
Ooyama, K., 1969: Numerical simulation of the life cycle of tropical
cyclones. J. Atmos. Sci., 26, 3–40.
Raymond, D. J., and H. Jiang, 1990: A theory for long-lived mesoscale convective systems. J. Atmos. Sci., 47, 3067–3077.
Reasor, P. D., M. T. Montgomery, and L. F. Bosart, 2005: Mesoscale observations of the genesis of Hurricane Dolly (1996).
J. Atmos. Sci., 62, 3151–3171.
Ritchie, E. A., and G. J. Holland, 1997: Scale interactions during theformation of Typhoon Irving. Mon. Wea. Rev., 125, 1377–1396.
Rotunno, R., and K. A. Emanuel, 1987: An air–sea interaction
theory for tropical cyclones. Part II: Evolutionary study using
a nonhydrostatic axisymmetrical numerical model. J. Atmos.
Sci., 44, 542–561.
Satoh, M., T. Matsuno, H. Tomita, H. Miura, T. Nasuno, and S. Iga,
2008: Nonhydrostatic icosahedral atmospheric model (NICAM)
for global cloud resolving simulations. J. Comput. Phys., 227,
3486–3514.
Schubert, W. H., and J. J. Hack, 1982: Inertial stability and tropical
cyclone development. J. Atmos. Sci., 39, 1687–1697.
Shapiro, L. J., and H. E. Willoughby, 1982: The response of balanced hurricanes to local sources of heat and momentum.
J. Atmos. Sci., 39, 378–394.
Simpson, J., E. Ritchie, G. J. Holland, J. Halverson, and S. Stewart,
1997: Mesoscale interactions in tropical cyclone genesis. Mon.
Wea. Rev., 125, 2643–2661.
Sippel, J. A., J. W. Nielsen-Gammon, and S. E. Allen, 2006: The
multiple-vortex nature of tropical cyclogenesis. Mon. Wea.
Rev., 134, 1796–1814.
Sperber, K. R., and D. E. Waliser, 2008: New approaches to understanding, simulating, and forecasting the Madden–Julian
oscillation. Bull. Amer. Meteor. Soc., 89, 1917–1920.
Tomita, H., and M. Satoh, 2004: A new dynamical framework ofnonhydrostatic global model using the icosahedral grid. Fluid
Dyn. Res., 34, 357–400.
Tory, K. J., M. T. Montgomery, and N. E. Davidson, 2006a: Prediction and diagnosis of tropical cyclone formation in an NWP
system. Part I: The critical role of vortex enhancement in deep
convection. J. Atmos. Sci., 63, 3077–3090.
——, ——, ——, and J. D. Kepert, 2006b: Prediction and diagnosisof tropical cyclone formation in an NWP system. Part II: A
diagnosis of Tropical Cyclone Chris formation. J. Atmos. Sci.,
63, 3091–3113.
——, N. E. Davidson, and M. T. Montgomery, 2007: Prediction anddiagnosis of tropical cyclone formation in an NWP system.

DECEMBER 2010

FUDEYASU ET AL.

Part III: Diagnosis of developing and nondeveloping storms.
J. Atmos. Sci., 64, 3195–3213.
Wang, Y., 2002a: Vortex Rossby waves in a numerically simulated
tropical cyclone. Part I: Overall structure, potential vorticity,and kinetic energy budgets. J. Atmos. Sci., 59, 1213–1238.
——, 2002b: Vortex Rossby waves in a numerically simulated
tropical cyclone. Part II: The role in tropical cyclone structure
and intensity changes. J. Atmos. Sci., 59, 1239–1262.
——, and C. C. Wu, 2004: Current understanding of tropical cyclone structure and intensity changes—A review. Meteor.
Atmos. Phys., 87, 257–278.
Willoughby, H. E., J. A. Clos, and M. G. Shoreibah, 1982: Concentric eye walls, secondary wind maxima, and the evolutionof the hurricane vortex. J. Atmos. Sci., 39, 395–411.

4327

Yang, B., Y. Wang, and B. Wang, 2007: The effect of internallygenerated inner core asymmetric structure on tropical cyclone
intensity. J. Atmos. Sci., 64, 1165–1188.
Zhang, D. L., and J. M. Fritsch, 1987: Numerical simulation of themeso-beta scale structure and evolution of the 1977 Johnstown
flood. Part II: Inertially stable warm-core vortex and the mesoscale convective complex. J. Atmos. Sci., 44, 2593–2612.
——, and N. Bao, 1996a: Oceanic cyclogenesis as induced bya mesoscale convective system moving offshore. Part I: A 90-h
real-data simulation. Mon. Wea. Rev., 124, 1449–1469.
——, and ——, 1996b: Oceanic cyclogenesis as induced by a mesoscale convective system moving offshore. Part II: Genesis
and thermodynamic transformation. Mon. Wea. Rev., 124,
2206–2225.