VOLUME 138

MONTHLY WEATHER REVIEWDECEMBER 2010

Multiscale Interactions in the Life Cycle of a Tropical Cyclone Simulated in a Global
Cloud-System-Resolving Model. Part I: Large-Scale and Storm-Scale Evolutions*
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 24 July 2010)
ABSTRACTThe Nonhydrostatic Icosahedral Atmospheric Model (NICAM), a global cloud-system-resolving model,
successfully simulated the life cycle of Tropical Storm Isobel that formed over the Timor Sea in the australsummer of 2006. The multiscale interactions in the life cycle of the simulated storm were analyzed in thisstudy. The large-scale aspects that affected Isobel’s life cycle are documented in this paper and the corresponding mesoscale processes are documented in a companion paper.
The life cycle of Isobel was largely controlled by a Madden–Julian oscillation (MJO) event and the associated westerly wind burst (WWB). The MJO was found to have both positive and negative effects on thetropical cyclone intensity depending on the location of the storm relative to the WWB center associated with
the MJO. The large-scale low-level convergence and high convective available potential energy (CAPE)
downwind of the WWB center provided a favorable region to the cyclogenesis and intensification, whereas
the strong large-scale stretching deformation field upwind of the WWB center may weaken the storm by
exciting wavenumber-2 asymmetries in the eyewall and leading to the eyewall breakdown.
Five stages are identified for the life cycle of the simulated Isobel: the initial eddy, intensifying, temporary
weakening, reintensifying, and decaying stages. The initial eddy stage was featured by small-scale/mesoscaleconvective cyclonic vortices developed in the zonally elongated rainband organized in the preconditioned environment characterized by the WWB over the Java Sea associated with the onset of an MJO event over the East
Indian Ocean. As the MJO propagated eastward and the cyclonic eddies moved southward into an environment
with weak vertical shear and strong low-level cyclonic vorticity, a typical tropical cyclone structure developed over
the Java Sea, namely the genesis of Isobel. Isobel experienced an eyewall breakdown and a temporary weakening
when it was located upwind of the WWB center as the MJO propagated southeastward and reintensified as its
eyewall reformed as a result of the axisymmetrization of an inward spiraling outer rainband that originally formed
downwind of the WWB center. Finally Isobel decayed as it approached the northwest coast of Australia.

* School of Ocean and Earth Science Technology Publication Number 7984 and International Pacific Research Center Publication
Number 713.
1
Current affiliation: Faculty of Education and Human Sciences, Yokohama National University, Yokohama, Japan.

Corresponding author address: Dr. H. Fudeyasu, Yokohama National University, 79-1 Tokiwadai, Hodogaya-ku, Yokohama, Kanagawa
240-8501, Japan.
E-mail: fude@ynu.ac.jp
DOI: 10.1175/2010MWR3474.1
Ó 2010 American Meteorological Society

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1. Introduction
To simulate and understand the life cycle of a tropical
cyclone (TC), including genesis, intensification, structure
and intensity changes, and decay, has long been a captivating subject at the frontier of science and remains
challenging because of the complex multiscale interactions involved. The life cycle of a TC consists of a succession of stages starting from the establishment of a favorable
environment to TC genesis. During the genesis stage,
sustained convective activities, which may stem from avariety of processes in a favorable environment, develop
into a surface mesoscale or synoptic vortex. Although themesoscale processes in the genesis stage have been theleast understood aspect of the life cycle of a TC, the climatological large-scale conditions favorable for TC genesis have been well known since Gray (1968, 1979). These
include a sea surface temperature (SST) above about
26.58C with a deep ocean mixed layer, a cyclonic lowlevel vorticity anomaly, weak and preferably easterly
vertical shear, and a region of organized deep convection
with a moistened mid-to-lower troposphere.
Previous studies have paid special attention to the roleof the Madden–Julian oscillation (MJO; Madden andJulian 1971, 1972, 1994; Zhang 2005) in the organization ofdeep convection as a precursor to TC genesis (Liebmann
et al. 1994; Maloney and Hartmann 2000a,b; Hall et al.
2001; Bessafi and Wheeler 2006; Barrett and Lance 2009).
In addition to the MJO, various tropical/equatorial waves,such as the easterly waves, equatorial Rossby waves,
mixed Rossby–gravity waves, can also modulate the largescale vertical vorticity and vertical shear, thus significantly
affecting TC genesis and life cycle over different ocean
basins (Landsea 1993; Briegel and Frank 1997; Ritchieand Holland 1999; Thorncroft and Hodges 2001; Molinari
and Vollaro 2000; Maloney and Hartmann 2001; Hallet al. 2001; Dickinson and Molinari 2002; Bessafi and
Wheeler 2006; Frank and Roundy 2006).
The genesis stage is generally followed by an intensification stage. During this stage, the typical structureof a TC, such as the eye, eyewall, and spiral rainbands,
would emerge. After its formation, the structure and intensity of a TC may be subject to changes due to bothinternal processes and external forcing (Wang and Wu
2004). Previous studies have shown that strong vertical
shear may interrupt the intensification or weaken a TC by
ventilating the warm core or triggering asymmetries ineyewall convection (Simpson and Riehl 1958; Gray 1968;
McBride and Zehr 1981; Merrill 1988; Wang and Holland
1996; Bender 1997; Frank and Ritchie 1999, 2001; Wanget al. 2004; Black et al. 2002; Heymsfield et al. 2006; Zenget al. 2007, 2008). Several internal dynamical processes canlead to an interruption of TC intensification or a weakening,

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such as the eyewall breakdown due to the interaction
between spiral rainbands and eyewall convection (Wang
2002a,b) and the concentric eyewall cycle (Willoughby
et al. 1982). The last stage of the life cycle of a TC is its
decay under the effects of an unfavorable environment
or when it moves over cold water or makes landfall, or
its transition to become an extratropical cyclone.
Significant progress has been made in modeling TCs
with high-resolution atmospheric models since Liu et al.
(1997). The modeling studies have greatly improved ourunderstanding of different stages of the life cycle of a TC.
However, few studies have succeeded in simulating thewhole life cycle of a TC because most high-resolution
models are run in a limited area for several days and thus
generally could not capture the life cycle of a TC that isalways embedded in a large-scale flow with time scales
longer than a week. Although two-way nesting with amoving mesh following a TC can reduce this discrepancy,
the best way to avoid the problem completely is to run a
global cloud-resolving model. The ever-increasing capability of powerful computers in recent years makes it possible to simulate the whole life cycle of a TC by running a
global nonhydrostatic model at resolutions high enough toresolve cloud systems.
The first global cloud-system-resolving model, thatis, the Nonhydrostatic Icosahedral Atmospheric Model(NICAM), with horizontal grid spacing of a few kilometers, simultaneously models cloud clusters, organized mesoscale convective systems, and the associated large-scale
circulation, such as the MJO (Nasuno et al. 2007; Satohet al. 2008). Miura et al. (2007) reported a successful simulation of a real MJO event in the austral summer 2006 in
NICAM with realistic settings (hereinafter the MJO experiment). The MJO experiment reproduced the largescale organized cloud systems, including the multiscale
cloud clusters, in the convective phase of the MJO. As afollow-up report of the simulated MJO by Miura et al.
(2007), Fudeyasu et al. (2008) presented the successful
simulation of the complete life cycles of two real TCs,
Bondo and Isobel, which formed over the Indian Ocean
in the MJO experiment. Initialized with the atmosphericconditions two weeks before the genesis of Isobel, themodel captured reasonably well not only the timing andmotion but also the mesoscale structure and evolution of
Isobel.
The main purpose of this study is to conduct a detailed
analysis on the multiscale interactions involved in the life
cycle of the simulated Isobel based on the output of
the MJO experiment. This paper (Part I) focuses on thelarge-scale aspects of the genesis and subsequent stormscale evolution. In a companion paper (Fudeyasu et al.
2010, hereinafter Part II), we will examine the involved
mesoscale processes. The rest of the paper is organized as

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follows. Section 2 briefly introduces the MJO experimentand the analysis methods. The simulated MJO and its
associated westerly wind burst (WWB) and Tropical
Storm Isobel are compared with those observed in section
3. The life cycle of the simulated Isobel is described in
section 4. The evolution of the large-scale environmental
conditions and the storm-scale structure change of thesimulated Isobel are analyzed in section 5. The influenceof the MJO on the life cycle of the simulated Isobel isfurther discussed in section 6. The major results are summarized in the last section.

2. The NICAM MJO experiment and analysis
methods
NICAM has been developed at the Frontier ResearchCenter for Global Change (FRCGC), at the Japan Agency
for Marine Earth Science and Technology (JAMSTEC),
and the Center for Climate System Research (CCSR),
at the University of Tokyo, on the world’s fastest-evercomputer—the Earth Simulator in Japan. The model
adopts the icosahedral horizontal grid system, which covers
the globe quasi-homogeneously. The model physics include the explicit cloud microphysics scheme of Grabowski(1998) for moist processes, a simple bucket model forland surface processes, and the level-2 turbulence closure
scheme of Nakanishi and Niino (2006) for subgrid-scalevertical mixing [see Tomita and Satoh (2004) and Satohet al. (2008) for more details].
The MJO experiment used in this study was a 32-day
simulation with realistic settings (the realistic initial conditions and terrain) and a horizontal resolution of about
7 km and 40 unevenly distributed vertical levels with themodel top at about 38 km in a terrain-following height coordinate (Miura et al. 2007; Nasuno et al. 2009). The model
was initialized at 0000 UTC 15 December 2006 with initial atmospheric variables and soil moisture from the
National Centers for Environmental Prediction (NCEP)
analysis, which has a horizontal resolution of 1.08 by 1.08.
The model provided outputs of 6-hourly snapshot threedimensional variables and 1.5-hourly averaged twodimensional variables.
The SST was spatially and temporally interpolated
from the weekly Reynolds SST with a horizontal resolution of 1.08 by 1.08 (Reynolds and Smith 1994). Miuraet al. (2008) performed two sensitivity experiments using the same model setting with the NICAM MJO experiment, but with a horizontal resolution of about
14 km. The time evolution of the simulated MJO event
did not differ significantly between the simulations withthe observed time-varying SST and the fixed climatological SST. The results implied that the simulated MJO
was inherently an internal atmospheric mode and the4287

observed SST was not critical to the eastward propagation of the MJO at least in the model for this case.
To facilitate our analysis of evolution of the large-scale
circulation, we removed the TC signal from the model
output using the smoothing algorithm developed by
Kurihara et al. (1990). At first, the NICAM output wasarea averaged onto 1.08 by 1.08 grids. The disturbance field
was filtered from the original field with successive applications of a simple smoothing operator. As a result, thebasic fields consisting of only the large-scale circulation
without either TC circulation or small-/mesoscale features
were well retained, while waves having wavelengths lessthan 98 longitude–latitude were entirely removed fromthe original field and those with longer wavelengths werepartially damped.
Definition of the TC center is critical to the calculationof the symmetric and asymmetric structures of the TC.
Several methods were used to determine the storm centerin this study. After 1800 UTC 28 December 2006 when
Isobel had a clear signal in the sea level pressure (SLP)
field, Isobel’s center (xc, yc) was defined as the masscenter (Wang 2007) given belowðð( pr ps )x dx dy( pr ps )y dx dy, yc 5 ð, (1)xc 5 ð
( pr ps ) dx dy( pr ps ) dx dywhere ps is the SLP and pr is a reference pressure (here1010 hPa). The integration domain is a circular area
whose radius is approximately 210 km from the surface
vortex center. Several iterations are required to get a converged solution because the integration radius depends on
the storm center. Before 1800 UTC 28 December, thecirculation center provides promising results. First, thewind field at each level was decomposed into basic fieldand disturbance field. The circulation center was definedas the center of the axisymmetric circulation along which
the azimuthal-mean wind at 850 hPa reaches the maximum at a given time. Since the results using the circulation
center at 850 hPa after 1800 UTC 28 December 2006 were
essentially the same as that obtained based on the masscenter, the 850-hPa circulation center was used to determine Isobel’s center in the initial eddy stage (see section 4).

3. Verification of the simulationa. The MJO and the associated WWB
Figure 1 shows the Hovmöller diagrams of anomalies
in the observed blackbody temperature (TBB) and thesimulated outgoing longwave radiation (OLR) averaged
between 158N and 158S. The observed TBB was derived
from the merged IR channels from five geostationary
satellites of the Climate Prediction Center (CPC-IR).

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FIG. 1. Hovmöller diagrams of anomalies of (a) CPC-IR and (b) NICAM-OLR averaged between 158N and 158S.
The anomalies are derived from the time average during 15 Dec 2006–15 Jan 2007. The solid line with filled circles
represents Isobel’s tracks from (a) JTWC and (b) NICAM.

The simulated OLR was area averaged onto a 0.25 grid
to match the resolution of the satellite products. The
evolution of the MJO event, characterized by an eastward propagation of the 5000–10 000-km convective envelope at approximately 5 m s21 over 608E–1808, waswell reproduced in the model. However, compared withobservations, the convective envelope in the simulation
shifted 108–208 longitude eastward over the East Indian
Ocean and the Maritime Continent. Miura et al. (2007),
Masunaga et al. (2008), Nasuno et al. (2009), and Liu et al.
(2009) verified the evolution of this NICAM-simulated
MJO event against various observations. They foundthat NICAM realistically simulated the three-dimensional
structure of the MJO but the simulated MJO grew fasterand had the amplitude larger than that observed.
Since the WWB associated with the MJO event isclosely related to Isobel’s life cycle, we briefly discuss the
behavior of the WWB in both observation and simulation
here. Although there are various definitions for the
WWB (e.g., Murakami and Sumathipala 1989; Harrison
and Vecchi 1997; Seiki and Takayabu 2007), we simply
defined the WWB as the westerlies at 850 hPa greaterthan 1 m s21 in the basic field associated with the MJO.
To compare the evolution of the observed and simulated
WWB quantitatively, we introduced the WWB front andWWB center. The former was defined by the easternboundary (easternmost point) of the WWB, while the latter was the location of the maximum westerly in the WWB.
Figure 2 shows the evolution of the WWB in the NCEP
analysis and the NICAM simulation. In NICAM, the
WWB front and center started to appear over the East
Indian Ocean on 24 December and moved eastward toward the western Pacific (Figs. 2b,d). The real WWB also
started to appear on 24 December but the location of the
WWB front and center were 108–208 longitude west of
those simulated (Figs. 2a,c). As a result, the WWB in thesimulation arrived at the Maritime Continent about twodays earlier. After the passage of the WWB front over
the Java Sea, both the observed and simulated MJO continued to move eastward but the simulated front shifted
108–158 latitude southward away from the equator. The
intensity of the simulated WWB showed an evolution
similar to that observed but about 30% stronger (Fig. 2e).
This is consistent with the analysis by Liu et al. (2009), who

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FIG. 2. Temporal evolution of (a),(b) the WWB front (solid line) and front point (closed circle) and (c),(d) the
WWB center at 0000 UTC derived from (a),(c) the NCEP analysis and (b),(d) the NICAM simulation. Solid lines
represent (a),(c) the JTWC and (b),(d) the simulated Isobel’s tracks. The dotted line in (b) represents the track
relative to the WWB front point. (e) Temporal evolution of the WWB intensity derived from the NCEP analysis(solid) and the NICAM simulation (dot) and (f) the location of Isobel relative to the WWB center. Tracks andlocations at 0000 UTC (solid line and closed circle) from the simulated Isobel are compared with those from the
JTWC best track (cross). Numbers mean the dates of the month.

showed that the simulated MJO event grew faster over the
Indian Ocean and the Maritime Continent and had an
intensity about 30% stronger than the observed. We willshow in section 3b that biases in the simulated MJO were
responsible for the discrepancies in the life cycle of thesimulated Isobel.

b. An overview of the observed and simulated Isobel
The genesis of Isobel was associated with the onset of
the MJO over the East Indian Ocean in both simulationand observation. The observed Isobel was located to thesoutheast of the WWB center based on the Joint Typhoon Warning Center (JTWC) best-track data (Fig. 2f).
Similar to the genesis of the observed Isobel, the simulated Isobel formed in the southeast quadrant of the
WWB about 700 km from the WWB center. According
to the best-track data from both JTWC and the Australian

Bureau of Meteorology (BoM), the actual Isobel was firstdetected over the Timor Sea south of the Java Island on
31 December 2006 (Fig. 3a). The track of the simulated
Isobel was very close to that observed from 31 December.
The central SLP of the observed Isobel remained steadyaround 1000 hPa for 2 days after its formation according to the JTWC and BoM best-track data (Fig. 3b) andthen deepened and reached the minimum of 994 hPa at
1800 UTC 2 January by JTWC and 980 hPa at 1200 UTC
2 January by BoM. [Postanalysis of Isobel by BoM indicated that the circulation did not reach tropical cyclone
intensity (see online at http://www.bom.gov.au/weather/wa/cyclone/about/seasonsummary200607.shtml). On theother hand, the JTWC final report (available online at http://
metocph.nmci.navy.mil/jtwc/atcr/2007atcr/2007atcr.pdf)
indicated that Isobel reached the tropical storm intensity.
Our focus in this study is to understand the multiscale

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FIG. 3. (a) Tracks and locations at 0000 UTC (line and closed circle) from the simulated Isobel, compared withthose from the BoM best-track (line and cross) and the JTWC best-track (line and plus). Numbers mean the dates ofthe month. (b) Temporal evolution of the central SLP (hPa) from the NICAM simulation (solid) and BoM and JTWC
best tracks (dashed). (c) Temporal evolution of the simulated maximum azimuthal-mean surface wind speed (m s21).

interactions involved in the simulated Isobel. Therefore,
the difference in the storm intensity from the observation
is not the key.] The central SLP of the simulated Isobel
was about 10 and 35 hPa deeper than that of JTWC and
BoM, respectively. This discrepancy could be attributed tothe simulated too strong WWB in the model. Nevertheless,the model reproduced the stagnant nature of the movement during the first 2 days and the subsequent intensification and weakening of the observed Isobel (Figs. 3a,b).
Figure 4 shows the hourly rainfall rate derived from theTropical Rainfall Measuring Mission (TRMM) Microwave
Imager (TMI) and the Quick Scatterometer (QuikSCAT)
sea surface winds (Freilich and Dunbar 1999; Chelton andFreilich 2005). Prior to the first detection of Isobel by JTWC
and BoM, there were regions with organized convection
associated with the MJO and its WWB around the Java
Island on 30 December 2006. The zonally elongated rainband labeled A in Fig. 4 was organized over the Java Sea at
1200 UTC 30 December 2006. The rainband A moved
southeastward and entered the Timor Sea on 1 January
2007 and evolved into cloud clusters. Another rainfall region, labeled B in Fig. 4, was organized over the Timor Sea
south of the Java Island on 31 December. BoM and JTWC
first detected the center of Isobel near the cloud cluster B on
31 December. On 1 January 2007, the cluster B weakened
over the Timor Sea. The detected center of Isobel was between clusters A and B, as seen from the circulation centerin the surface wind field on 1 January 2007. The two clusters
merged to the north of the circulation center over the Timor
Sea at 0000 UTC 2 January and subsequently became thespiral rainband and the partial eyewall of Isobel.
A similar rainband formed over the Java Sea in thesimulation was about two days too early (Fig. 5) compared with that in the observations (Fig. 4). This wasprimarily due to the too early growth of the WWB andthe associated convection over the Maritime Continent in
the simulation (Figs. 1 and 2). Nevertheless, although thedetails are somewhat different, the overall rainfall features of the simulated Isobel are comparable to thoseobserved (Figs. 4 and 5), including the asymmetric rainfall
distribution in the inner-core region. The spiral rainband
southwest of actual Isobel’s center at 0000 UTC 2 January(Fig. 4) also appeared in the simulation between 1800 UTC
1 January and 0000 UTC 2 January (Fig. 5). As Isobel
was approaching the northwest coast of Australia after
0000 UTC 3 January, intense rainfall occurred over the land
south of Isobel’s center (Fig. 4). This asymmetric rainfall
pattern between 1200 UTC 3 January and 0000 UTC
4 January was also reproduced by the model (Fig. 5).
Note that although the early formation stage of thesimulated Isobel was not compared well with that in the
BoM or JTWC best-track data, satellite observations
showed the linkage between the cloud/rainfall patternorganized in the MJO over the Java Sea and the genesis of

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FIG. 4. Rainfall rate (mm h21, shaded) derived from the TRMM TMI and surface wind (m s21) measured by NASA’s Quick Scatterometer. Light (dark) shading represents areas with rainfall rate larger than 0.5 mm h21 (5 mm h21). The crosses and pluses from31 Dec 2006 to 3 Jan 2007 represent the cyclone center derived from the BoM and JTWC best track, respectively.

Isobel. This linkage was well simulated in NICAM as will
be discussed in the next sections. Therefore, the MJO experiment captured reasonably well the life cycle of Isobel
as already shown in Fudeyasu et al. (2008). This allows us
to further understand the multiscale interactions in Isobel’slife cycle based on the NICAM simulation in this study.

4. The life cycle of the simulated Isobel
To facilitate the discussion, we divided the life cycle ofthe simulated Isobel into five stages (Table 1). We refer tothe period from 1800 UTC 27 December to 1800 UTC

28 December 2006, prior to Isobel being detectable in
SLP field, as the initial eddy stage. At 1800 UTC 27 December, convective cyclonic eddies appeared in a zonallyoriented rainband over the Java Sea (see Part II). These
eddies moved southeastward over the Java Sea and madelandfall on one of the Lesser Sunda Islands at 1800 UTC
28 December. The surface relative cyclonic vorticity averaged within a radius of 50 km from the initial eddy’s
center first increased and then decreased slightly (Fig. 6a)as the eddies approached the Lesser Sunda Islands(Fig. 3a). The maximum azimuthal-mean surface windspeed increased (Fig. 3c) and the radius of maximum wind

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FIG. 5. Rainfall rate (mm h21, shaded) and surface (10-m height) winds (m s21) from the NICAM simulation. The crosses
represent Isobel’s center.

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TABLE 1. List of different stages in the simulated Isobel’s life cycle and characteristics of the storm evolution, rainfall pattern, largescale setting, and the dominant effects: minimum surface pressure (SP), mid–upper troposphere (UL), vertical shear (VS), TC-scalevorticity (TC), low-level horizontal cyclonic shear (CS), steering flow (layer mean flow) (SF), inner core convection (eyewall) (EW), lowlevel horizontal convergence (CO), low-level stretching deformation (SD), large-scale velocity (CV), low-level zonal divergence (ZD),
and wavenumber (WN).
Life stage(abbreviation)

Characteristics
PeriodPreconditioning (PC) 0000 UTC 27 Dec–1800
UTC 27 Dec 2006
Initial eddy (IE)
1800 UTC 27 Dec–1800
UTC 28 Dec 2006
Intensifying (IN)
1800 UTC 28 Dec–0600
UTC 31 Dec 2006
Temporary
0600 UTC 31 Dec 2006–1800
weakening (TW)
UTC 1 Jan 2007
Early TW
0600 UTC 31 Dec 2006–0600
UTC 1 Jan 2007
Late TW
0600–1800 UTC 1 Jan 2007
Reintensifying (RI)
1800 UTC 1 Jan–1800
UTC 2 Jan 2007
Decaying (DE)
1800 UTC 2 Jan–0000
UTC 5 Jan 2007

TC evolution
No detection

RainfallIE genesis
Contraction
No SP signal
TC developing EW formation
SP decreasing
TC weakening EW breakdown
SP increasing
WN-1 asymmetryWN-2 asymmetry
TC developing EW reformation
SP minimum
TC weakening EW breakdown
SP increasing(RMW) decreased (Fig. 6b). Rainfall averaged within aradius of about 50 km increased whereas that averaged
within a radius of 100 km changed little and that averaged
between radii of 100 and 200 km decreased (Fig. 6c),indicating a contraction of the rainfall region (Fig. 5).
The initial eddy stage was followed by a period during
which Isobel’s central SLP decreased to the first minimumof about 975 hPa at 0000 UTC 31 December (Fig. 3b).
The period from 1800 UTC 28 December to 0600 UTC
31 December is thus referred to as the intensifying stage(Table 1). During this period, both the maximum surface
wind speed and the surface cyclonic relative vorticity averaged within a radius of 50 km increased rapidly (Figs. 3cand 6a), while the RMW remained between 30 and 40 km(Fig. 6b). The rainfall increased in the outer region (Fig. 6c).
A distinct feature at this stage is the increase in the radiusof the azimuthal-mean 15 m s21 surface wind (Fig. 6b),indicating the development of the storm-scale circulation. Isobel showed little movement over the Timor Sea
during the intensifying stage (Fig. 3a).
The intensification of the simulated Isobel was interrupted at around 0600 UTC 31 December and the storm
weakened until 1800 UTC 1 January 2007 with the central
SLP increased by 15 to 990 hPa (Fig. 3b). This period isreferred to as the temporary weakening stage (Table 1).
The maximum surface wind speed decreased by about
15 m s21 (Fig. 3c) and the increase in the surface cyclonic
relative vorticity slowed down considerably (Fig. 6a).
Rainfall showed a significant decrease, in particular, withina radius of 50 km (Fig. 6c), indicating a weakening of
the inner-core convection. The simulated Isobel showed

Large-scale setting

Rainband formation MJO onset

Dominant effects

Easterly trade winds
Weak UL winds

Strong CO/CV/VSHigh CAPE
Strong CO/CV
Strong VS/SF
Strong CS/CV
Weak VS/SF

Strong UL easterly

Strong VS

Enhanced WWB

Shift of WWB center Positive SD/ZD
Enhanced WWB
Change in VS
Strong SF
Strong UL northerly Strong VS/SF
Landfall effectan unusual track with a cyclonic rotation over the Timor
Sea during the late temporary weakening stage (Fig. 3a).
The simulated Isobel started to reintensify from1800 UTC 1 January, referred to as the reintensifying stage(Table 1). Its central SLP decreased and reached its lifetime minimum value of about 960 hPa at 0600 UTC
2 January (Fig. 3b). The maximum azimuthal-mean surface
wind speed increased by about 25 m s21 (Fig. 3c) and thesurface relative vorticity averaged within a radius of 50 km
reached the lifetime minimum value of about 24 31024 s21 (Fig. 6a). By the end of the reintensifying stage,
Isobel headed southeastward toward the northwest coastof Australia and started to weaken (Figs. 3 and 6), namely,
the decaying stage of the simulated Isobel. Both the central
SLP and the surface relative vorticity increased rapidly(Figs. 3b and 6a). The maximum surface wind speed decreased with an increase in the RMW (Figs. 3c and 6b).
Isobel made landfall at 1200 UTC 4 January with its central
SLP of about 990 hPa (Fig. 3a). Therefore, the MJO experiment simulated the week-long life cycle of Isobel comparable to the typical life cycle of a TC (Wang and Wu 2004).

5. Evolution of the large-scale circulation andstorm-scale structure change
a. Preconditioning of the genesis and intensificationof Isobel
Figure 7 shows the large-scale basic winds at 850, 500,
and 200 hPa from 1800 UTC 27 December 2006 to
1200 UTC 2 January 2007. On 27–28 December, the WWB
developed over the East Indian Ocean and the Java Sea

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FIG. 6. Temporal evolution of (a) relative vorticity(1024 s21) averaged within a radius of 50 km from Isobel’s center, (b) the radii of the maximum azimuthalmean surface wind (km, solid) and 15 m s21 surface
wind (km, dashed), and (c) 12-hourly mean rainfall rate
(mm h21) averaged between radii of 0–50 (filled box),
50–100 (gray box), and 100–200 km (open box).

associated with the onset of the MJO (Fig. 7a). The lowlevel WWB extended east-southeastward during 29–30
December, while the easterly trade winds extended westward around the latitude of 158S south of Isobel (Fig. 7b).
The upper troposphere was dominated by easterlies inthe equatorial region and westerlies to the south in the
midlatitude.
The evolution of several large-scale parameters derived
from the basic wind field and averaged within a radius of
500 km from Isobel’s center is shown in Fig. 8. Note that
before the initial eddy stage Isobel’s center was fixed at thecenter of the initial eddies over the Java Sea at 1800 UTC
27 December (also in Figs. 9–13). The low-level cyclonicshear and convergence of zonal wind (du/dy and du/dx)
increased over the Java Sea from 27 December, providinga favorable large-scale condition for the organization ofsmall-/mesoscale convective vortices (Figs. 8a,b). In thefollowing days, the low-level large-scale cyclonic shear
around Isobel was enhanced by the WWB and easterly
trade winds over the Timor Sea (Fig. 7b). The convective
available potential energy (CAPE) following Isobel’s
center was high (.500 J kg21) over the Maritime Continent before 27 December, while it decreased afterward asa result of enhanced convection associated with the WWB
over the Java Sea during the initial eddy stage (Fig. 8d).
For a comparison, Fig. 8 also shows the averaging following the MJO during 24–28 December 2006 (labeled
MJO in Fig. 8). The centers of the average were determined based on the position of Isobel relative to theeastward-propagating WWB front, as shown in Fig. 2b.
Namely, the average regions were located in the Indian
Ocean during 24–25 December, while in the Java Sea
during 26–28 December. The average region relative to
the MJO front was always accompanied with strong
horizontal cyclonic shear and convergence of zonal wind
(du/dx and du/dy) and strong vertical shear (Figs. 8a,c),
whereas convergence/shear of meridional wind (dy/dx
and dy/dy) were negligible (not shown). The CAPE increased with time as the MJO moved toward the Maritime Continent (Fig. 8d). Therefore, the preconditioned
large-scale environment set high CAPE around the Maritime Continent, and low-level cyclonic vorticity and convergence over the Java Sea, leading to active convection
and the genesis of Isobel.
Figure 9 shows the time–latitude (time–longitude) crosssection of zonal winds of the basic flow zonally (meridionally) averaged within 500 km from the longitude (latitude) of Isobel’s center at 850 hPa. The WWB associatedwith the MJO appeared in the equatorial region on
26 December 2006 and extended southward gradually, reaching the Java Sea at around 68S at 0000 UTC
27 December. The initial eddies formed southeast of the
maximum low-level westerlies in the WWB at 1800 UTC
27 December (Figs. 2f and 9). The westerlies continued toextend southeastward and reached the Lesser Sunda Islands by 29 December. Isobel formed and developed in thewesterly slightly north of the cyclonic shear line (0 m s21 of
averaged zonal winds) in the lower troposphere.

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FIG. 7. The horizontal wind (m s21) and wind velocity (m s21, contour) of the basic fields of (left) 850-, (middle) 500-, and (right)
200-hPa 24-hourly mean started at (a) 1800 UTC 27 Dec, (b) 1200 UTC 29 Dec, (c) 0600 UTC 31 Dec 2006, and (d) 1200 UTC 1 Jan 2007.
The contour interval is 5 m s21. The cross represents the storm center at the starting time of the averaged period.

Figure 10 shows the time–height cross section of thebasic wind field averaged within a radius of 500 km fromthe storm center with vectors showing the vertical shear,
Isobel’s motion, and deep/shallow layer mean flow. From
26 December, the easterly vertical shear increased rapidly(Figs. 8c and 10), which resulted from the developmentof the low-level westerlies associated with the WWB. As
the initial eddies moved southward away from the region of the equatorial upper-level easterlies, the vertical shear decreased during both the late initial eddy
stage and the early intensifying stage (Figs. 7b, 8c, and 10).
The development of Isobel benefited from the enhanced

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FIG. 8. Temporal evolution of (a) zonal divergence (1026 s21, du/dx, solid), zonal shear (1026 s21, du/dy, solid),
and meridional divergence (1026 s21, dy/dy, solid) at 850 hPa; (b) relative vorticity (1026 s21, solid), divergence(1026 s21, solid), and horizontal stretching deformation (1026 s21, solid) at 850 hPa; (c) vertical shear (m s21, solid)
between 200 and 850 hPa; and (d) CAPE (J kg21, solid) in the basic fields averaged within a radius of 500 km from
Isobel’s center. The location of Isobel’s center is fixed at the center of the initial eddies at 1800 UTC 27 Dec during
0000 UTC 24 Dec–1800 UTC 27 Dec. Results for the average region moved following the WWB front with Isobel’s
center relative to the WWB front point at 0000 UTC 28 Dec as the reference shown in Fig. 2b are displayed as adashed line (labeled MJO).

low-level large-scale cyclonic vorticity and weak vertical
shear over the Timor Sea in the intensifying stage. The
mid–upper-level winds were relatively weak around Isobel(Figs. 7b and 10), consistent with the weak steering flow
and little movement of Isobel (Fig. 3a).
The vertical profiles of relative vorticity averaged withina radius of 301 km from Isobel’s center and potential
temperature anomaly averaged within a radius of 49 km
are shown in Fig. 11. The cyclonic vorticity started todevelop below about 4 km at 1800 UTC 27 December,a signal of the genesis of initial eddies. After their geneses, the initial eddies moved eastward until 0600 UTC
28 December and then southward until 0000 UTC
29 December (Fig. 3a). At the end of the initial eddy
stage, the cyclonic vorticity developed upward into theupper troposphere (Fig. 11a). The warm anomaly formedin the upper troposphere in the inner core with the maximum temperature anomaly at about 9–12 km (Fig. 11b).
This was accompanied by a decrease in Isobel’s central SLP(Fig. 3b), indicating that Isobel developed a TC structure.
The azimuthal-mean rainfall rate shows a contractionof the rainfall region in the initial eddy stage (Fig. 12),which indicates an intensification of the initial eddies.

This is consistent with the findings of Zehr (1992), who
noted the inward contraction of convection during the TC
genesis from observations. On 30 December in the lateintensifying stage, the azimuthal-mean heavy rainfall occurred between radii of 20 and 40 km with little rainfall
within 20 km, a sign of the formation of an eye and eyewall. By 1800 UTC 30 December, a rainband developed
about 50 km to the west of Isobel’s center and propagatedradially outward (Fig. 5).
By the end of the initial eddy stage, the circulation
center tilted westward with height, and in the late initialeddy and early intensifying stages the upper-level centers coincided with the mid- and low-level centers (notshown), indicating that Isobel developed a vertically aligned
deep circulation. Figure 13a shows the ‘‘axisymmetricity’’defined as
0BBaxisymmetricity(%) 5B1@

ð r5301km
0ð r53011
EKE drCC
C3 100,
A
KE dr

0

(2)

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FIG. 9. (a) The time–latitude cross section of the 850-hPa zonal wind (m s21) of the basic field zonally averaged
within 500 km from the longitude of Isobel’s center and (b) the time–longitude cross section of the 850-hPa zonal
wind of the basic field meridionally averaged within 500 km from the latitude of Isobel’s center. The contour interval
is 2 m s21. Light shading represents westerlies and dark shading represents areas of westerlies stronger than 4 m s21.
The solid line with closed circles represents Isobel’s track.

where KE is the azimuthal-mean kinetic energy, EKE
is the azimuthal-mean eddy kinetic energy (Wang 2007,2008), and r is the radius from the storm center. The
axisymmetricity was quite small in the early initial eddy
stage while it increased in the late initial and intensifying
stages. The development of symmetric structure with vertically aligned deep circulation of Isobel was consistent
with the favorable environment with a weak vertical shear.

b. Unfavorable large-scale circulation and theinterruption of intensification
The cyclonic vorticity of Isobel decreased in the mid–
upper troposphere during the temporary weakening stage,
which was accompanied by a weakening of the warm corein the upper troposphere and an increase in the central
SLP (Figs. 3b and 11). Convection in the inner core wassuppressed whereas convection outside the inner core only
weakened slightly during the temporary weakening stage(Fig. 12). These structural changes were consistent with
the negative effect of the large-scale circulation, which
was characterized by strong vertical shear in the early
temporary weakening stage and then a strong horizontal stretching deformation field in the late temporary
weakening stage. From 31 December 2006 to 1 January
2007, upper-level easterlies strengthened around Isobel(Fig. 7c), leading to an increase in easterly vertical shear(Figs. 8c and 10). Corresponding to the strong easterly
shear, the upper-level circulation centers of Isobel shifted
westward from the mid–lower-level centers (not shown)with deep convection enhanced downshear (Fig. 5), consistent with previous studies (Wang and Holland 1996;
Frank and Ritchie 1999).

The axisymmetricity decreased in the temporary weakening stage (Fig. 13a). The asymmetry was dominated by
the wavenumber-1 component in the mid–upper troposphere in the late intensifying and early temporary weakening stages (Figs. 13b,c), which can be explained by the
influence of the strong vertical shear. From 1200 UTC
1 January, Isobel moved to the region with weak uppertropospheric winds between strong easterlies to the northand midlatitude westerlies to the south (Fig. 7d). As a result, the vertical shear decreased (Figs. 8c and 10) and thewestward vertical tilt of the storm center was reduced in
the late temporary weakening stage (not shown).
Isobel remained in the westerlies in the lower troposphere during the intensifying and the temporarily weakening stages (Fig. 9). The WWB moved southeastward
and the WWB center shifted to the east of Isobel (Figs. 2f,
7d, and 9b), leading to the low-level divergence of zonal
wind (du/dx) around Isobel (Fig. 8a). In contrast, its meridional convergence (dy/dy) between the southerlies eastof the subtropical high developed over the East Indian
Ocean and the northerlies of the southeastwardly extended WWB increased around Isobel. These mid–lowertropospheric zonal divergence and meridional convergencecontribute to an increase in the large-scale stretching deformation field (Fig. 8b). The axisymmetricity remained
small in the late temporary weakening stage. Although thetotal asymmetry was still dominated by a wavenumber-1
asymmetry (Figs. 13b,c), a wavenumber-2 asymmetry appeared in the mid–lower troposphere in the late temporary
weakening stage (Fig. 13d) when the large-scale stretching
deformation field was strong. During 1200–1800 UTC
1 January, the rainfall rate in Isobel showed two spiral rainbands (Fig. 5), indicating the development of4298

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FIG. 10. Time–height cross section of horizontal winds (m s21, barb) of the basic flow averaged within a radius of 500 km from Isobel’s center. The contour interval is 2 m s21 of zonal
wind. Light shading represents westerlies and dark shading represents zonal wind stronger than
4 m s21. Full wind barb represents 10 m s21. The vectors above show the vertical shear vector,
Isobel’s motion, and the deep-layer (850–300 hPa) and shallow-layer (850–700 hPa) mean
winds derived from the basic wind field within a radius of 500 km.

a wavenumber-2 asymmetry in Isobel in the strong
stretching deformation field.

c. Changes in large-scale circulation andreintensification of Isobel
During 1–2 January 2007, the WWB extended southeastward and strengthened in the lower troposphere,causing the appearance of northwesterlies in the midtroposphere over the Java Sea and the eastern Timor
Sea (Fig. 7d). The midlevel northwesterlies resulted ina fast southeastward movement of Isobel (Figs. 3a and 10).
The vertical shear over Isobel changed from easterly shear
to north-northeasterly shear accordingly (Fig. 10). The
cyclonic vorticity of Isobel developed again throughout thetroposphere at the beginning of the reintensifying stage
with the warm core in the upper troposphere strengthened(Fig. 11) and Isobel reintensified (Figs. 3b,c).
The reintensification of Isobel was associated with theeyewall reformation through the axisymmetrization of
an inward-propagating outer spiral rainband (Figs. 5 and13a). At 1500 UTC 1 January, the eyewall convection
showed some discrete cloud features. A large outer spiralrainband was initiated to the southeast of Isobel’s centerat around 158–168S, downwind of the strengthened WWB
center. The rainband was enhanced by an intense confluence zone between the strong northwesterlies of the
WWB and southeasterly trade winds at 1800 UTC 1 January (Figs. 5 and 7d). By 2100 UTC 1 January, the spiralrainband continuously enhanced and spiraled cyclonically inward. During 2100 UTC 1 January and 0600 UTC
2 January, as Isobel’s center quickly moved southeastward,
the inner end of the spiral rainband wrapped into the coreregion and organized into a new partial eyewall (Fig. 5). In
addition, accompanied with the cyclonic inward propagation of the outer rainband was a cyclonic rotation of
Isobel’s track during 1500 UTC 1 January–0300 UTC
2 January (Fig. 3a). Isobel started to decay from 0600
UTC 2 January mainly as a result of the land effect as it
was approaching the northwest coast of Australia (Fig. 3a).

6. Discussion on the influence of the MJO on
Isobel’s life cycle
Isobel’s life cycle was closely related to the evolutionof the MJO and the associated WWB in the simulation

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FIG. 11. Time–height cross section of (a) absolute vertical vorticity (1025 s21) averaged within a radius of 301 km
from Isobel’s center and (b) potential temperature anomaly (K) averaged within a radius of 49 km. The contour
interval in (a) is 2 3 1025 s21, and light (dark) shading represents areas with vertical vorticity less than 24 3 1025 s21(28 3 1025 s21). The contour interval in (b) is 2 K, and light (dark) shading represents areas with an anomaly greaterthan 4 K (8 K). The warm anomaly was defined as the difference from potential temperature averaged in 308S–08 and
808–1408E at 0000 UTC 23 Dec 2006.

as discussed in the last section. Hall et al. (2001) examined statistically the relationship between the TC activity and the MJO in the Australian region and foundthat the MJO strongly modulates the climatologicalpattern of TC geneses in the region, especially pronounced to the northwest of Australia, where Isobel
formed. Previous studies also showed that the WWB
associated with an MJO event could produce twin cyclones straddling the equator as a result of the Rossby
wave response to the heating associated with the active
convection phase of the MJO (Nitta et al. 1992; Sui andLau 1992, Lin and Johnson 1996). The cyclonic gyrenorth of the equator, twin of Isobel, did not appear inthis case. This might be due to the fact that convection
associated with the MJO occurred around the Maritime
Continent, mainly south of the equator. As a result, the
WWB extended southeastward over the Timor Sea withconvection suppressed in the Northern Hemisphere(Fig. 2) as predicted by the simple model of Gill (1980).
To examine how the large-scale circulation associatedwith the WWB as an external forcing contributed to thelife cycle of Isobel, a composite analysis was executed. To
eliminate the effects of the Maritime Continent, the signal of Isobel, and the southward movement of the WWB
away from the equator, the composite analysis based on
the basic wind fields was done for the period of 24–
25 December 2006 when the WWB developed over theeastern Indian Ocean. Figure 14 shows the composite

wind field with respect to the WWB center. The horizontal wind field is dominated by two large-scale zonallyelongated cyclonic gyres: one to the north and one to thesouth of the WWB center (Figs. 14a,b). The large-scale
cyclonic vorticity accompanied with the WWB with high
CAPE provided favorable conditions for the initiation anddevelopment of small-scale/mesoscale convective vortices

FIG. 12. Time–radius cross section of the azimuthal-mean rainfall
rate (mm h21) based on 1.5-h model output. Areas with 10–
20 mm h21 are lightly shaded, those of 20–30 mm h21 are heavily
shaded, and those higher than 30 mm h21 are filled.

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FIG. 13. (a) Time–height cross section of the axisymmetricity (%) defined in Eq. (2). Time–height cross sectionof (b) the total, (c) wavenumber-1, and (d) wavenumber-2 asymmetric kinetic energies (m2 s22) averaged withina radius of 301 km. The contour interval in (a) is 20% and light (dark) shading represents areas larger than
20% (80%). The contour interval in (b) is 20 m2 s22 and light (dark) shading represents areas larger than 20 m2 s22(80 m2 s22). Contour intervals are 10 m2 s22 in (c) and 5 m2 s22 in (d) and light (dark) shading represents areas
larger than 10 m2 s22 (20 m2 s22).

over the Java Sea and eventually led to the genesis of
Isobel. Thus, the WWB associated with the MJO played
an important role in the genesis and intensification of
Isobel. Meanwhile, the WWB had a negative effect on thegenesis of Isobel. During the early initial eddy stage, thevertical shear was strong over the initial eddies due tothe low-level WWB and upper-level equatorial easterlies
over the Java Sea (Figs. 7a, 8c, and 10), which is unfavorable for cyclogenesis (Gray 1968). The vertical shear
decreased as the initial eddies moved southward awayfrom the equator. Eddies then organized and developed
into Tropical Storm Isobel over the Timor Sea.

The location of Isobel relative to the WWB center isimportant to the structure and intensity changes of thesimulated Isobel. During the initial eddy stage and intensifying stage, Isobel was located southeast of the WWB
center (Fig. 2f), where a strong large-scale convergenceexists in the lower troposphere (Figs. 8b and 14c), favorable
for the development of deep convection and the genesis of
Isobel. In contrast, as the WWB moved southeastward,
Isobel experienced a temporary weakening when Isobel
was located southwest of the WWB center (Fig. 2f), wherea strong stretching deformation and a strong large-scale
divergence of zonal wind were dominant (Figs. 8a,b andDECEMBER 2010

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FIG. 14. Composite (a) horizontal winds (m s21), (b) relative vorticity (s21), (c) divergence (s21), and (d) stretching
deformation (s21) based on the basic wind fields at 850 hPa with respect to the WWB center during 24–25 Dec 2006.
The contour interval in (a) is 2 m s21 and light (dark) shading represents areas larger than 2 m s21 (6 m s21). The
contour interval in (b) is 2 3 1026 s21 and light (dark) shading represents areas larger/lower than 66 3 1026 s21 (610 31026 s21). The contour interval in (c) is 2 3 1027 s21 and light (dark) shading represents areas larger (smaller) than 2 31027 s21 (22 3 1027 s21). The contour interval in (d) is 1 3 1026 s21 and light (dark) shading represents areas larger(smaller) than 2 3 1026 s21 (22 3 1026 s21). Solid (dashed) contours denote positive (negative) values. Track of Isobel
relative to the WWB center and locations at 0000 UTC are shown by solid line and closed circles (as in Fig. 2f).

14d). The strong stretching deformation was responsible
for the development of two spiral rainbands (Fig. 5) and awavenumber-2 asymmetric dynamical structure (Fig. 13d)in the late temporary weakening stage. This differs fromthe dominant wavenumber-1 asymmetry developed in thestrong vertical shear in the early weakening stage.
Davis et al. (2008) analyzed the tropical cyclogenesis
over the eastern Pacific in 2005 and 2006 and found thatthe horizontal stretching deformation in the backgroundflow at 900 hPa differed significantly between the developing and nondeveloping systems. Developing systems
were generally located in the zero or negative stretching
deformation field while nondeveloping systems mostly
occurred in the positive stretching deformation field. The
structures of nondeveloping systems were generally zonally elongated in the strong stretching deformation field,
which was very similar to the wavenumber-2 asymmetricstructure developed in Isobel in the late temporary

weakening stage in our simulation. Although Davis et al.
(2008) did not analyze TCs at the mature stage, the strong
horizontal stretching deformation in the environmental
flow may have a negative effect on TC intensity by triggering convective asymmetries in the inner core. Therefore,
depending on the location of Isobel relative to the WWB
center, the MJO played both positive and negative effects
on the genesis and intensity of the simulated Isobel.

7. Conclusions
The global cloud-system-resolving model NICAM successfully simulated the major features of Isobel’s life cycle
in an MJO experiment. This allowed us to investigate theeffect of the MJO on the evolution of the simulated Isobel, as summarized in Table 1. With high CAPE aroundthe Maritime Continent, the WWB over the Java Sea
associated with the MJO onset over the East Indian

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MONTHLY WEATHER REVIEW

Ocean enhanced the large-scale cyclonic vorticity andconvergence in the lower troposphere, favorable for active convection, leading to the development of a zonallyelongated rainband (the preconditioning stage of Isobel).
Convective eddies initiated and organized in the rainband in the preconditioned environment (the initial eddy
stage). As the WWB extended eastward, the initial eddies
moved southeastward. After the initial eddies arrived at
the favorable environment with weak vertical shear andhorizontal cyclonic shear enhanced between the WWB
and easterly trade winds over the Timor Sea, Isobel developed the typical TC structure (the intensifying stage).
With the increase in vertical shear, a wavenumber-1 asymmetric structure developed in Isobel’s eyewall, leading toa breakdown of the eyewall and a temporary weakening of
the storm (the early temporary weakening stage). As Isobel
shifted to the west of the WWB center, the wavenumber-2
asymmetric structure in the inner core of Isobel was excitedin a strong environmental stretching deformation field in
the mid–lower troposphere (the late temporary weakening stage). As Isobel quickly moved southeastward, it
experienced an eyewall reformation process through theaxisymmetrization of an inward-propagating outer rainband that formed downwind of the WWB center (thereintensifying stage). Isobel finally weakened as a resultof the land effect as it approached the northwest coast ofAustralia (the decaying stage).
The life cycle of the simulated Isobel therefore waslargely controlled by a large-scale environmental condition characterized by the MJO onset and the associated
WWB, the easterly trade winds, and changes in vertical
shear and the steering flow. The WWB associated with
the MJO, on one hand, contributed positively to the intensification of Isobel as Isobel was located southeast of
the WWB center where large-scale convergence occurred,and on the other hand, affected negatively on Isobel’s
development as Isobel was located southwest of the WWB
center where the environmental stretching deformation
was strong. Therefore, the MJO had both positive andnegative effects on the intensity of the simulated Isobel,
depending on the relative location of Isobel and the WWB
center associated with the MJO.
Although the NICAM MJO experiment provides hightemporal and spatial resolution dataset for a detailed
study of the multiscale interactions involved in the life
cycle of the simulated Isobel, the results reported here
should not be taken to imply that there is skill at forecasting TC behavior at such a lead time. We noticed thatthe simulated Isobel was stronger than the observed
owing to the stronger MJO in the simulation. Nevertheless, to understand the multiscale interactions involved in
the simulated Isobel can help understand the role of the
MJO in the life cycle of a TC. Therefore, the difference in

VOLUME 138

the simulated storm intensity from that observed should
not be a big problem for our purpose in this study. This
paper covers only the evolution of the large-scale environmental flow and the storm-scale evolution of Isobel.
The multiscale interactions among the large-scale, stormscale, and mesoscale processes in the simulated Isobel
will be discussed in Part II.
Acknowledgments. The authors are grateful to Dr. Ping
Liu for his helpful comments on the NICAM-simulated
MJO verification. This study was supported in part by NSF
Grant ATM-0754039 and in part by JAMSTEC, NASA,and NOAA through their sponsorships of the International Pacific Research Center (IPRC) in the School of
Ocean and Earth Science and Technology (SOEST) at the
University of Hawaii. Part of this study was also supportedby the JST/CREST in Japan. The Earth Simulator wasused for the NICAM simulation.

REFERENCES
Barrett, B. S., and M. L. Lance, 2009: Links between tropical cyclone activity and Madden–Julian oscillation phase in the
North Atlantic and Northeast Pacific basins. Mon. Wea. Rev.,
137, 727–744.
Bender, M. A., 1997: The effect of relative flow on the asymmetricstructure in the interior of hurricanes. J. Atmos. Sci., 54, 703–724.
Bessafi, M., and M. C. Wheeler, 2006: Modulation of south Indian
Ocean tropical cyclones by the Madden–Julian oscillation andconvectively coupled equatorial waves. Mon. Wea. Rev., 134,
638–656.
Black, M. L., J. F. Gamache, F. D. Marks, C. E. Samsury, and
H. E. Willoughby, 2002: Eastern Pacific Hurricanes Jimena of
1991 and Olivia of 1994: The effect of vertical shear on
structure and intensity. Mon. Wea. Rev., 130, 2291–2312.
Briegel, L. M., and W. M. Frank, 1997: Large-scale influences on
tropical cyclogenesis in the western North Pacific. Mon. Wea.
Rev., 125, 1397–1413.
Chelton, D. B., and M. H. Freilich, 2005: Scatterometer-based assessment of 10-m wind analyses from the operational ECMWF
and NCEP numerical weather prediction models. Mon. Wea.
Rev., 133, 409–429.
Davis, C., C. Snyder, and A. C. Didlake, 2008: A vortex-based
perspective of eastern Pacific tropical cyclone formation. Mon.
Wea. Rev., 136, 2461–2477.
Dickinson, M., and J. Molinari, 2002: Mixed Rossby–gravity waves
and western Pacific tropical cyclogenesis. Part I: Synoptic
evolution. J. Atmos. Sci., 59, 2183–2196.
Frank, W. M., and E. A. Ritchie, 1999: Effects of environmental flow
upon tropical cyclone structure. Mon. Wea. Rev., 127, 2044–2061.
——, and ——, 2001: Effects of vertical wind shear on the intensity
and structure of numerically simulated hurricanes. Mon. Wea.
Rev., 129, 2249–2269.
——, and P. E. Roundy, 2006: The role of tropical waves in tropical
cyclogenesis. Mon. Wea. Rev., 134, 2397–2417.
Freilich, M. H., and R. S. Dunbar, 1999: The accuracy of the
NSCAT 1 vector winds: Comparisons with National Data
Buoy Center buoys. J. Geophys. Res., 104, 11 231–11 246.

DECEMBER 2010

FUDEYASU ET AL.

Fudeyasu, H., Y. Wang, M. Satoh, T. Nasuno, H. Miura, andW. 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 ina global cloud-system-resolving model. Part II: System-scale
and mesoscale processes. Mon. Wea. Rev., 138, 4305–4327.
Gill, A. E., 1980: Some simple solutions for heat-induced tropicalcirculation. Quart. J. Roy. Meteor. Soc., 106, 447–462.
Grabowski, W. W., 1998: Toward cloud resolving modeling oflarge-scale tropical circulations: A simple cloud microphysics
parameterization. J. Atmos. Sci., 55, 3283–3298.
Gray, W. M., 1968: Global view of the origin of tropical disturbances and storms. Mon. Wea. Rev., 96, 669–700.
——, 1979: Hurricanes: Their formation, structure and likelyrole in the general circulation. Meteorology over the Tropical Oceans, D. B. Shaw, Ed., Royal Meteorological Society,155–218.
Hall, J. D., A. J. Matthews, and D. J. Karoly, 2001: The modulationof tropical cyclone activity in the Australian region by the
Madden–Julian oscillation. Mon. Wea. Rev., 129, 2970–2982.
Harrison, D. E., and G. A. Vecchi, 1997: Westerly wind events inthe tropical Pacific, 1986–95. J. Climate, 10, 3131–3156.
Heymsfield, G. M., J. Halverson, E. Ritchie, J. Simpson, J. Molinari,and L. Tian, 2006: Structure of highly sheared Tropical Storm
Chantal during CAMEX-4. J. Atmos. Sci., 63, 268–287.
Kurihara, Y., M. A. Bender, R. E. Tuleya, and R. J. Ross, 1990:
Prediction experiments of Hurricane Gloria (1985) using
a multiply nested movable mesh model. Mon. Wea. Rev., 118,
2185–2198.
Landsea, C. W., 1993: A climatology of intense (or major) Atlantic
hurricanes. Mon. Wea. Rev., 121, 1703–1713.
Liebmann, B., H. H. Hendon, and J. D. Glick, 1994: The relationship between tropical cyclones of the western Pacific andIndian Oceans and the Madden–Julian oscillation. J. Meteor.
Soc. Japan, 72, 401–412.
Lin, X., and R. H. Johnson, 1996: Heating, moistening, and rainfall
over the western Pacific warm pool during TOGA COARE.
J. Atmos. Sci., 53, 3367–3383.
Liu, P., and Coauthors, 2009: An MJO simulated by the NICAM at
14- and 7-km resolution. Mon. Wea. Rev., 137, 3254–3268.
Liu, Y. B., D. L. Zhang, and M. K. Yau, 1997: Multiscale numerical
study of Hurricane Andrew (1992). Part I: Explicit simulationand verification. Mon. Wea. Rev., 125, 3073–3093.
Madden, R. A., and P. R. Julian, 1971: Detection of a 40–50-day
oscillation in the zonal wind in the tropical Pacific. J. Atmos.
Sci., 28, 702–708.
——, and ——, 1972: Detection of global-scale circulation cells in
the tropics with a 40–50-day period. J. Atmos. Sci., 29, 1109–
1123.
——, and ——, 1994: Observations of the 40–50-day tropicaloscillation—A review. Mon. Wea. Rev., 122, 814–837.
Maloney, E. D., and D. L. Hartmann, 2000a: Modulation of eastern
North Pacific hurricanes by the Madden–Julian oscillation.
J. Climate, 13, 1451–1460.
——, and ——, 2000b: Modulation of hurricane activity in the Gulf
of Mexico by the Madden–Julian oscillation. Science, 287,
2002–2004.
——, and ——, 2001: The Madden–Julian oscillation, barotropic
dynamics, and North Pacific tropical cyclone formation. Part I:
Observations. J. Atmos. Sci., 58, 2545–2558.

4303

Masunaga, H., M. Satoh, and H. Miura, 2008: A joint satellite andglobal cloud-resolving model analysis of a Madden–Julian
Oscillation event: Model diagnosis. J. Geophys. Res., 113,
D17210, doi:10.1029/2008JD009986.
McBride, J. L., and R. Zehr, 1981: Observational analysis of tropical cyclone formation. Part II: Comparison of nondeveloping
versus developing systems. J. Atmos. Sci., 38, 1132–1151.
Merrill, R. T., 1988: Environmental influences on hurricane intensification. J. Atmos. Sci., 45, 1678–1687.
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.
——, ——, ——, ——, and ——, 2008: Response of coarseresolution models only partly cloudy. Science, 320, 613.
Molinari, J., and D. Vollaro, 2000: Planetary- and synoptic-scale
influences on eastern Pacific tropical cyclogenesis. Mon. Wea.
Rev., 128, 3296–3307.
Murakami, T., and W. L. Sumathipala, 1989: Westerly bursts during the 1982/83 ENSO. J. Climate, 2, 71–85.
Nakanishi, M., and H. Niino, 2006: An improved Mellor–Yamadalevel-3 model: Its numerical stability and application to a regional prediction of advection fog. Bound.-Layer Meteor., 119,
397–407.
Nasuno, T., H. Tomita, S. Iga, H. Miura, and M. Satoh, 2007:
Multiscale organization of convection simulated with explicit
cloud processes on an aquaplanet. J. Atmos. Sci., 64, 1902–
1921.
——, H. Miura, M. Satoh, A. T. Noda, and K. Oouchi, 2009: Multiscale organization of convection in a global numerical simulation of the December 2006 MJO event using explicit moist
processes. J. Meteor. Soc. Japan, 87, 335–345.
Nitta, T., T. Mizuno, and K. Takahashi, 1992: Multiscale convective systems during the initial phase of the 1986/87 El Niño.
J. Meteor. Soc. Japan, 70, 447–466.
Reynolds, R. W., and T. M. Smith, 1994: Improved global sea
surface temperature analyses using optimum interpolation.
J. Climate, 7, 929–948.
Ritchie, E. A., and G. J. Holland, 1999: Large-scale patterns associated with tropical cyclogenesis in the western Pacific. Mon.
Wea. Rev., 127, 2027–2043.
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.
Seiki, A., and Y. N. Takayabu, 2007: Westerly wind bursts and theirrelationship with intraseasonal variations and ENSO. Part I:
Statistics. Mon. Wea. Rev., 135, 3325–3345.
Simpson, R. H., and H. Riehl, 1958: Mid-tropospheric ventilation
as a constraint on hurricane development and maintenance.
Proc. 10th Conf. on Hurricanes, Miami Beach, FL, Amer.
Meteor. Soc., D4.1–D4.10.
Sui, C. H., and K. M. Lau, 1992: Multiscale phenomena in thetropical atmosphere over the western Pacific. Mon. Wea. Rev.,
120, 407–430.
Thorncroft, C., and K. Hodges, 2001: African easterly wave variability and its relationship to Atlantic tropical cyclone activity.
J. Climate, 14, 1166–1179.
Tomita, H., and M. Satoh, 2004: A new dynamical framework ofnonhydrostatic global model using the icosahedral grid. Fluid
Dyn. Res., 34, 357–400.
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.

4304

MONTHLY WEATHER REVIEW——, 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.
——, 2007: A multiply nested, movable mesh, fully compressible, nonhydrostatic tropical cyclone model-TCM4:
Model description and development of asymmetries without explicit asymmetric forcing. Meteor. Atmos. Phys., 97,
93–116.
——, 2008: Structure and formation of an annular hurricane simulated in a fully compressible, nonhydrostatic model—TCM4.
J. Atmos. Sci., 65, 1505–1527.
——, and G. J. Holland, 1996: Tropical cyclone motion and evolution in vertical shear. J. Atmos. Sci., 53, 3313–3332.
——, and C. C. Wu, 2004: Current understanding of tropical cyclone structure and intensity changes—A review. Meteor.
Atmos. Phys., 87, 257–278.

VOLUME 138——, M. T. Montgomery, and B. Wang, 2004: How much vertical
shear can a tropical cyclone resist? Bull. Amer. Meteor. Soc.,
85, 661–662.
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.
Zehr, R., 1992: Tropical cyclogenesis in the western North Pacific.
NOAA Tech. Rep. NESDIS 61, 181 pp.
Zeng, Z.-H., Y. Wang, and C.-C. Wu, 2007: Environmental dynamical control of tropical cyclone intensity—An observational study. Mon. Wea. Rev., 135, 38–59.
——, L.-S. Chen, and Y. Wang, 2008: An observational study ofenvironmental dynamical control of tropical cyclone intensity
in the Atlantic. Mon. Wea. Rev., 136, 3307–3322.
Zhang, C., 2005: Madden-Julian Oscillation. Rev. Geophys., 43,
RG2003, doi:10.1029/2004RG000158.