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Article Type:
Research Article

Tropical Transition of Tropical Storm Kirogi (2012) over the Western North Pacific: Synoptic Analysis and Mesoscale Simulation

Wataru Yanase aMeteorological Research Institute, Japan Meteorological Agency, Tsukuba, Japan

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https://orcid.org/0000-0002-0352-5297
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Udai Shimada aMeteorological Research Institute, Japan Meteorological Agency, Tsukuba, Japan

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Naoko Kitabatake bMeteorological College, Japan Meteorological Agency, Kashiwa, Japan

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Eigo Tochimoto cNational Research Institute for Earth Science and Disaster Resilience, Tsukuba, Japan

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Online Publication:
13 Sep 2023
Print Publication:
01 Oct 2023
DOI:
https://doi.org/10.1175/MWR-D-22-0190.1
Page(s):
2549–2572
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Abstract

Tropical transition (TT) is a cyclogenesis process in which a baroclinic disturbance is transformed into a tropical cyclone. Many studies have analyzed TT events over the North Atlantic. This study assesses TT processes from a possible subtropical cyclone to Tropical Storm Kirogi at a relatively high latitude over the western North Pacific in an environment of enhanced baroclinicity in August 2012. Analyses based on satellite observations, the JRA-55 reanalysis, and a simulation with 2.5-km horizontal grid spacing demonstrate three stages during the TT: the baroclinic, intermediate, and convective stages. Over the baroclinic stage, Kirogi had an asymmetric comma-shaped cloud pattern with convection in the northern and eastern parts of the cyclone. This convection is attributed to quasigeostrophic forcing and frontogenesis associated with advection of warm and moist air. Vorticity locally generated by this convection was advected to the cyclone center by cyclone-relative northerly flow. Kirogi also had a shallow warm-core structure due to the interaction with an upper-level cold trough extending from the midlatitudes. In the intermediate stage, the warm and moist air in the lower troposphere and the cold trough in the upper troposphere wrapped around Kirogi. In the convective stage, Kirogi attained characteristics of a typical tropical cyclone with convection concentrated near the cyclone center and a deep warm-core structure. These results demonstrate that baroclinic processes can directly trigger formation of a tropical storm at relatively high latitudes over the western North Pacific in a similar manner to that over the North Atlantic.

Significance Statement

Tropical cyclogenesis is an important process for early identification of tropical cyclone hazards. Tropical transition is a tropical cyclogenesis process that is triggered by a subtropical or extratropical disturbance. It is unique to relatively high latitudes and has social importance particularly for midlatitude countries. There have been fewer studies on tropical transition over the western North Pacific than over the North Atlantic. This study demonstrates the dynamics of a distinct tropical transition event that led to the formation of Tropical Storm Kirogi (2012) at a relatively high latitude over the western North Pacific.

© 2023 American Meteorological Society. This published article is licensed under the terms of the default AMS reuse license. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Wataru Yanase, wtryanase@gmail.com

Abstract

Tropical transition (TT) is a cyclogenesis process in which a baroclinic disturbance is transformed into a tropical cyclone. Many studies have analyzed TT events over the North Atlantic. This study assesses TT processes from a possible subtropical cyclone to Tropical Storm Kirogi at a relatively high latitude over the western North Pacific in an environment of enhanced baroclinicity in August 2012. Analyses based on satellite observations, the JRA-55 reanalysis, and a simulation with 2.5-km horizontal grid spacing demonstrate three stages during the TT: the baroclinic, intermediate, and convective stages. Over the baroclinic stage, Kirogi had an asymmetric comma-shaped cloud pattern with convection in the northern and eastern parts of the cyclone. This convection is attributed to quasigeostrophic forcing and frontogenesis associated with advection of warm and moist air. Vorticity locally generated by this convection was advected to the cyclone center by cyclone-relative northerly flow. Kirogi also had a shallow warm-core structure due to the interaction with an upper-level cold trough extending from the midlatitudes. In the intermediate stage, the warm and moist air in the lower troposphere and the cold trough in the upper troposphere wrapped around Kirogi. In the convective stage, Kirogi attained characteristics of a typical tropical cyclone with convection concentrated near the cyclone center and a deep warm-core structure. These results demonstrate that baroclinic processes can directly trigger formation of a tropical storm at relatively high latitudes over the western North Pacific in a similar manner to that over the North Atlantic.

Significance Statement

Tropical cyclogenesis is an important process for early identification of tropical cyclone hazards. Tropical transition is a tropical cyclogenesis process that is triggered by a subtropical or extratropical disturbance. It is unique to relatively high latitudes and has social importance particularly for midlatitude countries. There have been fewer studies on tropical transition over the western North Pacific than over the North Atlantic. This study demonstrates the dynamics of a distinct tropical transition event that led to the formation of Tropical Storm Kirogi (2012) at a relatively high latitude over the western North Pacific.

© 2023 American Meteorological Society. This published article is licensed under the terms of the default AMS reuse license. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Wataru Yanase, wtryanase@gmail.com
Keywords: Cyclogenesis/cyclolysis; Subtropical cyclones; Tropical cyclones; Nonhydrostatic models

1. Introduction

Tropical cyclones (TCs) generally form in tropical environments with high sea surface temperature (SST), low stability favorable for cumulus convection, and weak vertical shear (Gray 1975; Emanuel and Nolan 2004). At relatively high latitudes, however, some types of TCs form in different environments. For example, McTaggart-Cowan et al. (2015) focused on TC genesis over relatively low SST oceans (<26.5°C), and concluded that genesis was boosted by reduced static stability due to a cold tropopause. According to a global climatology analyzed by McTaggart-Cowan et al. (2013), 30% of TCs form in baroclinic conditions including lower-tropospheric baroclinicity, upper-tropospheric disturbances, or both. In particular, tropical transition (TT) is an important process in which a baroclinic disturbance is transformed into a warm-core TC (Davis and Bosart 2004). TT events often occur over the North Atlantic and have been investigated intensively (Davis and Bosart 2003; Hulme and Martin 2009a,b; Cordeira and Bosart 2011; Mauk and Hobgood 2012; Bentley et al. 2016, 2017). Empirical knowledge or statistical prediction models based on pure tropical dynamics may not be applicable to such baroclinically induced TCs. Wood et al. (2023, manuscript submitted to Trop. Cyclone Res. Rev.) reported that the National Hurricane Center’s predictions of TT events were less reliable than those of the average tropical cyclogenesis events over the North Atlantic.

One of the baroclinic disturbances that can lead to a TT process is a subtropical cyclone (SC). According to the Glossary of Meteorology (American Meteorological Society 2022), SCs occur in regions of weak to moderate horizontal temperature gradients and extract available potential energy, as do extratropical cyclones, but they also extract some or most of their energy from convective redistribution of heat acquired from the sea, as do tropical cyclones. An SC has a shallow warm-core structure composed of a warm core in the lower troposphere like a TC and a cold core in the upper troposphere like an extratropical cyclone (da Rocha et al. 2018). The climatology of SCs over the North Atlantic was explored by Evans and Guishard (2009) and Guishard et al. (2009) based on the cyclone phase space analysis (Hart 2003) which objectively classifies a cyclone with a shallow warm-core structure as an SC. They showed that the combination of warm SST and atmospheric baroclinicity determines the regions and seasons that are favorable for SC formation. Their composite analysis also indicated that SC formation is associated with the intrusion of upper-tropospheric disturbances into regions of relatively warm SST.

The TT process from an SC to a TC often shows characteristics partly similar to those in the late stage of an extratropical cyclone: e.g., a warm seclusion in the lower troposphere and a wrap-up of a PV streamer in the upper troposphere (Hulme and Martin 2009b; Cordeira and Bosart 2011; Maier-Gerber et al. 2019). Bentley et al. (2017) proposed three types of SCs over the North Atlantic based on the influence of upper-tropospheric disturbances: 1) the cutoff low category formed in reduced stability owing to cutoff lows south of anticyclonic wave breaking events; 2) the meridional trough category formed in dynamically forced ascent east of meridionally elongated troughs; and 3) the zonal trough category formed under the influence of relatively high SST and westward extending troughs south of anticyclonic wave breaking events.

Global climatological studies have demonstrated that SCs (or hybrid cyclones) and TT events also occur over basins other than the North Atlantic (McTaggart-Cowan et al. 2013; Yanase et al. 2014). There are several studies on SCs and TT over the South Atlantic (McTaggart-Cowan et al. 2006; Evans and Braun 2012; Gozzo et al. 2014) and the eastern and central North Pacific (Morrison and Businger 2001; Bentley and Metz 2016). As TCs form most frequently over the western North Pacific, it is worth examining SCs and TT events in this region. Sadler (1976, 1978) reported TC cases that developed under the influence of westward-moving upper-tropospheric disturbances called tropical upper-tropospheric trough (TUTT) cells. A few statistical studies also demonstrated that TCs sometimes form under the influence of upper-tropospheric disturbances over the western North Pacific (Fudeyasu and Yoshida 2019; Takemura and Mukougawa 2021). A composite analysis by Hamaguchi and Takayabu (2021) indicated that TUTTs forced quasigeostrophic ascending motion, moistened the mid- to upper troposphere, enhanced convection, and generated tropical depression–type disturbances between 10° and 20°N. Chang et al. (2019) is the first case study that explicitly discussed TT over the western North Pacific based on the cyclone phase space; pre-Typhoon Peipah (2007) was a deep cold core cyclone accompanied by quasigeostrophic ascending motion, transformed to a deep warm core cyclone at latitudes between 10° and 20°N, and two days later became a tropical storm. Although these studies focused on upper-tropospheric disturbances, few studies have analyzed baroclinic processes during TT that directly trigger the formation of tropical storms at relatively high latitudes over the western North Pacific.

Tropical Storm Kirogi over the western North Pacific had an asymmetric cloud pattern at relatively high latitudes (25°–30°N) implying baroclinic dynamics around the time of tropical storm formation on 6 August 2012. The purpose of this study is to assess the formation processes of Kirogi based on standard analyses of TT and SCs and to discuss whether they resemble the TT processes over other basins. A better understanding of differences in dynamics and environments between typical TC formation and TT over the western North Pacific will be useful for operational cyclone analysis in this region and statistical prediction of TC development based on dynamical and environmental predictors such as the Statistical Hurricane Intensity Prediction Scheme (SHIPS; DeMaria and Kaplan 1999; Shimada et al. 2018). This will also help to develop a global picture of TT. Section 2 of this paper describes the dataset used for the analyses and the design of the numerical simulation. Section 3 presents the TT characteristics of Kirogi based on observations and reanalysis. Section 4 analyzes the TT dynamics of Kirogi based on a numerical simulation. Section 5 discusses similarity to TT events over other basins, roles of vertical shear, and predictability. Section 6 summarizes our main conclusions.

2. Methodology

a. Observations and analysis datasets

The best track product from the Japan Meteorological Agency (JMA) was used for the location and minimum sea level pressure of Kirogi (JMA 2022). Because the JMA does not use the SC as a category of cyclones in their operational analysis, some TCs in their best track data may have had characteristics of an SC.

MTSAT-2 geostationary satellite observations with a resolution of 4 km were used for the analyses of cloud patterns and horizontal winds. The atmospheric motion vector is a horizontal wind product that is computed on 0.5° longitude–latitude grids by tracking the observed radiance patterns (Oyama 2010). The pressure levels of the horizontal winds were estimated from the observed radiance and atmospheric variables of the JMA global atmospheric model. We utilized a water vapor band that corresponds to the pressure levels between 200 and 400 hPa.

The JRA-55 reanalysis data (Kobayashi et al. 2015) with 1.25° horizontal grid spacing were used for examining large-scale conditions around Kirogi. The data were produced using a JMA global atmospheric model with TL319 spectral horizontal resolution and 60 vertical levels with four-dimensional variational data assimilation (4D-Var). The Lanczos filter with 30-day cutoff period and 120-day data window was used to extract subseasonal-scale environmental fields by removing synoptic-scale disturbances. The 40-yr mean fields between 1979 and 2018 were used as climatological fields to discuss the characteristics of the environmental fields in 2012.

b. Numerical simulation

To analyze the detailed structure and dynamics of Kirogi, we conducted a numerical simulation using the JMA Nonhydrostatic Model (Saito et al. 2006). The cloud processes were simulated by a two-moment bulk-type microphysics scheme that calculated mixing ratios for cloud water, cloud ice, rain, snow, and graupel as well as number concentrations for ice phases (Lin et al. 1983; Murakami 1990) along with a convective parameterization scheme (Kain and Fritsch 1990). The planetary boundary layer was parameterized by the level-2.5 closure of the Mellor–Yamada–Nakanishi–Niino turbulence scheme (Nakanishi and Niino 2004) with the surface-layer scheme proposed by Beljaars (1995). The model also calculated long- and short-wave radiation processes and ground temperature as described in JMA (2013).

The horizontal grid spacing was 2.5 km for a domain of 4000 km × 4000 km centered at 30°N, 160°E based on the Lambert conformal conical projection (Fig. 1). The vertical grid spacing for 48 layers with the model top at 21 801 m above mean sea level (MSL) increased linearly from 40 m at the lowest level to 868 m at the highest level. The initial and boundary atmospheric conditions and SST were obtained by interpolating the JRA-55 reanalysis data. Time integration was conducted at time steps of 10 s for 144 h starting at 0000 UTC 5 August 2012. As the TT process of Kirogi was not easy to simulate (see section 5c), we determined the experimental design so that the intensity and structure of Kirogi in the simulation resemble those in the observation and analysis.

Fig. 1.
Fig. 1.

Tracks of Kirogi in the simulation (black) and the JMA best track analysis (gray). Circles indicate 6-hourly locations of Kirogi with larger circles for 0000 UTC, where the digits denote 4, 6, 8, and 10 Aug 2012. Colors represent SST (°C) at 0000 UTC 5 Aug in ocean areas and ground elevation (m) in land areas in the model domain.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

We mainly analyzed a 10-km grid dataset that was obtained by averaging 4 × 4 grid points of the original 2.5-km grid dataset, as the original dataset was too noisy to analyze meso-β- or larger-scale structures of Kirogi. In some analyses described later, the 10-km grid dataset was further smoothed using a 400-km binomial filter, which calculates a weighted mean with 41-point binomial coefficients in the horizontal directions.

c. Analysis of a cyclone and atmospheric dynamics

The center of the cyclone was determined carefully particularly in the cyclogenesis stage when mesoscale convective systems disturbed the sea level pressure (SLP) field of a weak cyclone. Detecting the SLP minimum was found to be ineffective for the cyclogenesis stage. Therefore, we utilized the pressure centroid method proposed by Nguyen et al. (2014). The method iteratively calculates the centroid of the SLP anomaly. First, the location of minimum SLP is used as the first guess of the cyclone center, and the SLP centroid is calculated within a circle of radius R around the first guess:
x¯=∑r=0RxiPi′∑r=0RPi′ and y¯=∑r=0RyiPi′∑r=0RPi′,
Pi′=Penv−Pi,
where x is longitude, y is latitude, P′ is the SLP deficit, the index i denotes the grid points in the circle, and an overbar denotes the cyclone center. The SLP deficit is the difference between environmental SLP (Penv) and local SLP (Pi). We set R at 100 km and Penv at 1010 hPa. The resulting centroid is used as the new guess of the cyclone center, and (1) was repeated until the center location converged. This method provided a smoother track than the simple SLP-minimum detection method for the cyclogenesis stage, whereas the two methods resulted in almost identical tracks for later stages when the cyclone was intense.

The cyclone phase space assesses synoptic-scale cyclone structures objectively (Hart 2003; Evans and Guishard 2009). The thermal wind indices, −VTL and −VTU, are defined for the lower (900–600 hPa) and upper (600–300 hPa) troposphere, respectively. Positive and negative indices indicate warm- and cold-core structures of the cyclone, respectively. A TC has a deep warm-core structure (−VTL>0 and −VTU>0), and an SC has a shallow warm-core structure (−VTL>0 and −VTU<0).

Environmental vertical shear was defined as the difference in horizontal wind between 850 and 200 hPa averaged in an annulus bounded by circles with the radii of 200 and 800 km from the cyclone center (Kaplan and DeMaria 2003). For the JRA-55 reanalysis, the cyclone center was taken from the best track record.

Based on the quasigeostrophic approximation, the Q vector was defined as
Q=(−Rdp∂Vg∂x⋅∇hT,−Rdp∂Vg∂y⋅∇hT),
where Rd is the gas constant for dry air, p is pressure, Vg is the geostrophic wind vector, ∇h is the horizontal gradient operator, and T is temperature. Convergence of the Q vector indicates quasigeostrophic forcing for ascent (Hoskins et al. 1978).
The vorticity equation was examined in Cartesian coordinates relative to the cyclone center:
∂ζ∂t=−Vr⋅∇hη−w∂ζ∂z−η∇h⋅Vr−k^⋅(∇hw×∂Vr∂z)+ρ−2k^⋅(∇hρ×∇hp),
where ζ is the vertical component of relative vorticity, Vr is the cyclone-relative horizontal wind vector, η is the vertical component of absolute vorticity, w is the vertical wind, ρ is the density, and k^ is the unit vector in the vertical direction. From left to right, the terms on the right-hand side are the cyclone-relative horizontal advection, vertical advection, stretching, tilting, and solenoidal terms.
The time tendency of the circulation about the cyclone center was derived from the integration of (4) over a closed circular area (Davis and Galarneau 2009; Galarneau and Weisman 2020):
∂C∂t=−η¯δ˜A−∮η′Vr′⋅n^dl+∮w(k^×∂Vr∂z)⋅n^dl+∮ρ−1∇hp⋅n^dl,
where C is the circulation, δ˜ is the mean divergence of the closed horizontal area A, an overbar denotes the average value around the perimeter of the area, a prime denotes a perturbation from the average, and n^ is the unit vector locally outward normal to the perimeter. From left to right, the terms on the right-hand side are the stretching, eddy flux, tilting, and solenoidal terms. We referred to the derivation of the equation in Davis and Galarneau (2009); note that the solenoidal term disappears in their equation owing to the pressure coordinate. The right-hand side is not exactly equivalent to the left-hand side in the analysis, because short-time fluctuations were not resolved by 1-h interval of model output, and also because some variables including turbulent mixing were not available from the model output.

3. Observations and reanalysis

The locations and minimum pressures of Kirogi are shown in Figs. 1 and 2, respectively, based on the JMA best track record. Kirogi was first identified as a tropical depression (winds < 34 kt; 1 kt ≈ 0.51 m s−1) at 24.0°N, 161.8°E at 0600 UTC 3 August 2012, and showed limited development and northward motion over the next three days. It became a tropical storm (winds ≥ 34 kt) at 27.7°N, 162.6°E at 1200 UTC 6 August (referred to as genesis time). Kirogi’s minimum pressure decreased from 6 to 7 August and remained steady on 8 August (Fig. 2), while Kirogi moved northward and then northwestward (Fig. 1). From 9 to 10 August, Kirogi moved into the midlatitude environment with lower SST and its minimum pressure increased. It finally became an extratropical cyclone at 0600 UTC 10 August.

Fig. 2.
Fig. 2.

Time evolution of minimum sea level pressure of Kirogi in the best track analysis (thick gray) and the main simulation (thick black). Three sensitivity experiments, EXP1 (thin solid blue), EXP2 (thin dashed black), and EXP3 (thin dashed blue), are discussed in section 5c. The labels “B,” “I,” and “C” indicate the baroclinic stage (6 Aug), the intermediate stage (7 Aug), and the convective stage (8 Aug), respectively, determined in the last paragraph in section 3.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

The cloud patterns of Kirogi from 6 to 8 August are shown in Fig. 3. From 0000 to 1200 UTC 6 August, Kirogi had an asymmetric comma-shaped cloud pattern with a wide cloud area in the eastern and northern parts of the cyclone (Figs. 3a,b), implying a baroclinic structure (Cotton et al. 2011) around the time of genesis. The atmospheric motion vector indicated southwesterly flow in the upper troposphere around the wide cloud area. From 1200 UTC 6 August to 1200 UTC 7 August (Figs. 3b–d), the western end of the comma-shaped cloud started to wrap up and organized into a convective vortex separate from the wide cloud area. After 0000 UTC 8 August, the convective vortex was almost isolated from the dissipating wide cloud area to the east, and became an independent system like a typical TC (Figs. 3e,f). Thus, the cloud patterns indicate a remarkable structural transition from 6 to 8 August.

Fig. 3.
Fig. 3.

MTSAT-2 geostationary satellite observation of Kirogi. Infrared imagery (blue–white shading) along with atmospheric motion vectors derived from a water vapor band (yellow barbs). (a) 0000 UTC 6 Aug, (b) 1200 UTC 6 Aug, (c) 0000 UTC 7 Aug, (d) 1200 UTC 7 Aug, (e) 0000 UTC 8 Aug, and (f) 1200 UTC 8 Aug 2012. All panels show the same domain (20°–40°N, 150°–173°E) with grid intervals of 5°.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

Environmental vertical shear is generally considered to be detrimental to the genesis and intensification of TCs (Gray 1975; Frank and Ritchie 2001). Figure 4a shows the time evolution of the environmental vertical shear around Kirogi based on the JRA-55 reanalysis. The genesis of Kirogi on 6 August occurred during a temporary increase in the vertical shear exceeding 10 m s−1 (thick black curve), which is statistically less favorable for TC genesis (Nolan and McGauley 2012). The temporary increase in the vertical shear on 6 August is mainly attributed to the change in the upper-level winds (thin dashed curves) rather than the lower-level winds (thin dotted curves), indicating synoptic-scale modulation in the upper troposphere.

Fig. 4.
Fig. 4.

Time evolution of vertical shear and horizontal winds (m s−1) averaged in an annulus bounded by circles with the radii of 200 and 800 km from the cyclone center in (a) the JRA-55 reanalysis and (b) the simulation. The thick black curve indicates the magnitude of the vertical shear between 850 and 200 hPa. The red and blue curves are zonal and meridional components, respectively, of the vertical shear (thick solid), wind at 850 hPa (thin dotted), and wind at 200 hPa (thin dashed).

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

The upper-level potential vorticity (PV) is an important factor for interpreting TT. Figure 5 shows daily PV fields on the 350-K isentropic surface based on the JRA-55 reanalysis from 3 to 8 August. From 3 to 4 August, when pre-Kirogi was almost stationary and classified as a tropical depression in the best track record, a high PV disturbance was located to the northwest of pre-Kirogi (Figs. 5a,b). This disturbance originated from a high PV area in the subtropical central/eastern North Pacific that was associated with a TUTT. The northern side of the TUTT was characterized by low PV, implying anticyclonic wave breaking (Bentley et al. 2017). Although these characteristics resemble a TT process in the cutoff low category (Bentley et al. 2017), the cyclone remained a tropical depression during this period. From 5 to 6 August, another high PV disturbance approached Kirogi from the north (Figs. 5c,d), which extended from a high PV area in the midlatitudes to the western side of Kirogi. From 7 to 8 August, the disturbance further elongated and wrapped around Kirogi cyclonically and finally dissipated (Figs. 5e,f). These characteristics resemble a TT process in the meridional trough category (Bentley et al. 2017) and appear to have directly affected the genesis of Tropical Storm Kirogi on 6 August.

Fig. 5.
Fig. 5.

Potential vorticity on the 350-K isentropic surface (PVU; colors) and the 26.5°C SST isotherm (green curve) from the JRA-55 reanalysis at 1200 UTC (a) 3 Aug, (b) 4 Aug, (c) 5 Aug, (d) 6 Aug, (e) 7 Aug, and (f) 8 Aug 2012. Small and large crosses indicate the location of Kirogi at the tropical depression and tropical storm stages, respectively.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

Synoptic-scale temperature and wind fields and the Q vectors are presented in Fig. 6. On 6 August, the environment in the subtropics was characterized by a moderate southwest–northeast gradient of temperature throughout the troposphere, which was separated from the main baroclinic zone at the midlatitudes (Figs. 6a,d). In the upper troposphere, a cold anomaly accompanied the local maximum of the high PV southwest of Kirogi (Figs. 5d, 6a), which is a structure typical of an upper cold disturbance. This disturbance was characterized by northerly flow on the western side and southwesterly on the eastern side, which was consistent with the atmospheric motion vectors based on satellite observations (Fig. 3b); note that the upper-level trough was identified as a cloud-free region in the IR imagery. The cyclonic circulation associated with Kirogi and the upper cold disturbance in the moderate baroclinic environment induced convergence of the Q vector on its eastern side (Fig. 6g) with the maximum at 650 hPa (not shown), indicating quasigeostrophic forcing for ascent. On 7 August, the upper cold disturbance moved to the southeastern side of Kirogi (Fig. 6b). By this time, Kirogi was enclosed by upper-level cyclonic circulation, and was less affected by the strong wind associated with the upper cold disturbance (see also Figs. 3d, 4a). In addition, Kirogi became isolated from the tropical warm air to the southeast in the lower troposphere (Fig. 6e), and was also separated from the convergence of the Q vector (Fig. 6h), implying decrease in baroclinic influences on cyclone development. On 8 August, Kirogi had a warm anomaly throughout the troposphere (Figs. 6c,f), whereas the convergence of the Q vector was weak (Fig. 6i).

Fig. 6.
Fig. 6.

Synoptic-scale conditions from the JRA-55 reanalysis at 1200 UTC (a),(d),(g) 6 Aug; (b),(e),(h) 7 Aug; and (c),(f),(i) 8 Aug 2012. (top) Temperature (°C; colors) and horizontal wind vectors at 300 hPa. (middle) Temperature (°C; colors) and horizontal wind vectors at 850 hPa. (bottom) The Q vectors (kg−1 m2 s−1), convergence of the Q vectors (10−18 kg−1 m s−1; colors), and geopotential height (contour interval 50 m) at 700 hPa. The length scales of the vectors are shown in bottom-right corner in the right panels. Green curves are the 26.5°C SST isotherm. Crosses indicate the location of Kirogi.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

In addition to synoptic-scale processes, it is worth exploring subseasonal-scale environmental fields and their difference from the climatology (Fig. 7), because TT events are not frequent over the western North Pacific. On the 350-K isentropic surface, a high PV area extending from the subtropical central-eastern North Pacific and a low PV area on its north side were also evident on this time scale (Fig. 7a), indicating an enhancement of the TUTT structure compared to the climatological field (Fig. 7b). Geopotential heights at 850 hPa indicate that a ridge over the North Pacific intensified to the north of Kirogi and weakened to the south (Figs. 7c,d). This characteristic is similar to the pressure pattern favorable for tropical storm formation near upper-level disturbances over the North Atlantic (Galarneau et al. 2015). The temperature in the subtropics around the central Pacific was lower than the climatology at 300 hPa (not shown) and 850 hPa (Fig. 7f). This cold anomaly enhanced the horizontal gradient of temperature around Kirogi (Figs. 7e,f). The precipitable water changed little around Kirogi, but its meridional gradient was enhanced (Figs. 7g,h). The SST was lower than the climatology around Kirogi (not shown). The anomalous structure of the atmosphere around Kirogi was conductive to baroclinic cyclogenesis.

Fig. 7.
Fig. 7.

Subseasonal-scale conditions based on the 30-day-filtered JRA-55 reanalysis at 1200 UTC 6 Aug in (left) 2012 (colors and white contours) and (right) its anomaly (red and blue shading) from the 40-yr climatological fields (black contours). (a),(d) Potential vorticity at 350-K isentropic surface (PVU). (c),(d) Geopotential height at 850 hPa (m). (e),(f) Temperature at 850 hPa (°C). (g),(h) Precipitable water (kg m−2). The contour intervals are the same between the 2012 fields and the climatological fields. Crosses indicate the location of Kirogi.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

In summary, Kirogi’s genesis occurred in an enhanced baroclinic environment and showed a comma-shaped cloud pattern consistent with the quasigeostrophic forcing for ascent and an interaction with an upper-level cold disturbance extending from midlatitudes. These characteristics implying baroclinic processes on 6 August appear different from the tropical characteristics on 8 August. This result motivated us to further assess multiscale processes during the transition period from 6 to 8 August based on a numerical simulation. Hereafter, for convenience, we roughly separate the transition period into a baroclinic stage on 6 August, an intermediate stage on 7 August, and a convective stage on 8 August (Fig. 2) considering the above-mentioned characteristics.

4. Numerical simulation

a. Validation of basic characteristics

We first validated the simulation results by comparison with observations. The model reproduced the northward and northwestward motion of Kirogi during the transition period from 6 to 8 August (Fig. 1). The minimum pressure in the simulation was also similar to the best track record from 6 to 8 August (Fig. 2), although the intensity was overestimated for the first 24-h integration period on 5 August. The increase in environmental vertical shear around the time of genesis on 6 August was also reproduced with a slight overestimation compared with the value based on the reanalysis (Fig. 4). From 9 to 10 August, the cyclone moved northward more slowly in the simulation than in the observation. As a result, Kirogi in the simulation was located at lower latitude with higher SST (Fig. 1) and weaker vertical shear (Fig. 4) than in the observations, presumably resulting in the slower increase in minimum pressure. We focus on the transition period from 6 to 8 August, in which the model reproduced the location, intensity, and vertical shear realistically.

To compare with the MTSAT-2 observation (Fig. 3), simulated infrared imagery and horizontal wind vectors at ∼10 km MSL (∼300 hPa) are presented in Fig. 8. The model reproduced the following characteristics of the observed cloud pattern. In the baroclinic stage, the simulated Kirogi had a comma-shaped pattern, and the wide cloud area in the northern and eastern parts of the cyclone was accompanied by southwesterly flow (Figs. 8a,b). In the intermediate stage, the western end of the comma-shaped cloud started to wrap up and organized into a convective vortex separate from the wide cloud area (Figs. 8c,d). In the convective stage, the convective vortex was almost isolated from the dissipating wide cloud area to the east, and became an independent system (Figs. 8e,f).

Fig. 8.
Fig. 8.

Simulated infrared brightness temperature (°C; colors) and horizontal wind vectors at ∼10 km MSL (yellow barbs). (a) 0000 UTC 6 Aug, (b) 1200 UTC 6 Aug, (c) 0000 UTC 7 Aug, (d) 1200 UTC 7 Aug, (e) 0000 UTC 8 Aug, and (f) 1200 UTC 8 Aug. Black contours represent SLP every 5 hPa. Crosses denote cyclone centers.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

b. Cyclone-scale characteristics

An overview of structural change of Kirogi is given in the cyclone phase space diagram (Hart 2003) in Fig. 9. In the baroclinic stage on 6 August, Kirogi had a shallow warm-core structure (i.e., a warm-core structure in the lower troposphere and a cold-core structure in the upper troposphere), which resembled a characteristic SC over the North Atlantic (Evans and Guishard 2009). In the intermediate stage on 7 August, the structure in the upper troposphere changed from a cold core to a warm core. In the convective stage on 8 August, Kirogi finally developed a deep warm-core structure. Thus, the cyclone phase space indicated that Kirogi underwent a TT process from an SC to a TC. It should be noted that the cyclone phase space based on the JRA-55 reanalysis also indicated a shallow warm-core structure in the baroclinic stage (not shown); however, it failed to identify a deep warm-core structure in the convective stage presumably because the analysis was affected by an upper-tropospheric high pressure system near Kirogi and by insufficient representation of a weak TC in the reanalysis.

Fig. 9.
Fig. 9.

The cyclone phase space for Kirogi in the simulation. Horizontal and vertical axes denote the thermal wind indices in the lower and upper troposphere, respectively. Small and large circles indicate 6-hourly and daily indices, respectively, where the digits denote 0000 UTC from 5 to 11 Aug. The top-right, bottom-right, and bottom-left quadrants correspond to a deep warm-core, a shallow warm-core, and a deep cold-core cyclones, respectively.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

To diagnose favorable conditions for TT processes, Figs. 10a–c presents the coupling index which is defined by the difference between potential temperature (PT) at the dynamic tropopause (2-PVU potential vorticity surface; 1 PVU = 10−6 K kg−1 m2 s−1) and equivalent potential temperature (EPT) at 850 hPa (McTaggart-Cowan et al. 2015; Bentley et al. 2017); low value of coupling index (<22.5°C) indicates reduced bulk tropospheric stability which are considered favorable for TT processes. The data were smoothed by a 400-km binomial filter. Whereas Kirogi was under the influence of a synoptic-scale reduced stability area throughout the transition period, the relationship between convection and the reduced stability area changed with the stages. In the baroclinic stage (Fig. 10a), convection in the wide cloud area was organized in the eastern part of the reduced stability area. In the intermediate and convective stages (Figs. 10b,c), however, convection near the cyclone center developed around the center of the reduced stability area. This change may imply a difference in the organization mechanisms of convection between the stages.

Fig. 10.
Fig. 10.

Coupling index analysis of the simulation. (a)–(c) Coupling index (K; colors) and vertical shear vectors between the dynamic tropopause and 850 hPa. (d)–(f) PT (K; colors) and horizontal wind vectors at the dynamic tropopause. (g)–(i) EPT (K; colors) and horizontal wind vectors at 850 hPa. (left) 1200 UTC 6 Aug, (center) 1200 UTC 7 Aug, and (right) 1200 UTC 8 Aug. The length scales of the vectors are shown in bottom-right corner in the right panels. Thick contours represent ascent of 0.1, 0.2, and 0.5 m s−1 at ∼5 km MSL. Thin contours represent SLP every 5 hPa. Crosses denote cyclone centers.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

The contribution of low PT at the dynamic tropopause to coupling index is examined in Figs. 10d–f. Note that low PT at the dynamic tropopause corresponds to high PV on an isentropic surface. In the baroclinic stage (Fig. 10d), a low PT disturbance extended southward from the midlatitudes on the west side of the cyclone center, which corresponded to the high PV streamer observed in the JRA-55 reanalysis (Fig. 5d). The convection was organized on the eastern side of this disturbance. In the intermediate stage (Fig. 10e), the low PT disturbance wrapped around the cyclone center cyclonically (see also Fig. 5e). The convection near the cyclone center was located below the northern tip of this cyclonically wrapped low-PT area, implying the influence of reduced stability due to the upper cold air. In the convective stage (Fig. 10f), however, the low PT disturbance began to dissipate near the cyclone center (see also Fig. 5f) where convection was strong. The dissipation of an upper cold disturbance in a convective region during TC genesis was also presented in a case study on Hurricane Michael (2000) by Davis and Bosart (2003) and in an idealized experiment by Montgomery and Farrell (1993).

The contribution of high EPT at 850 hPa to coupling index is also examined in Figs. 10g–i. In the baroclinic stage (Fig. 10g), the highest EPT area was found in the eastern half of the cyclone owing to the northward advection of high EPT air from the low latitudes. The convection in the wide cloud area was organized around the northeastern periphery of this high EPT area. In the intermediate and convective stages (Figs. 10h,i), a high EPT area around the cyclone center became isolated from the high EPT area to the southeast, indicating a pattern similar to the warm core seclusion of an extratropical cyclone. This isolated high EPT area contributed to the low coupling index around the cyclone center, particularly in the convective stage when the low PT at the dynamic tropopause began to dissipate as mentioned above (Figs. 10c,f,i). The convection near the cyclone center was organized in this high EPT area.

As the convection organized on the periphery of the area of high EPT in the lower troposphere in the baroclinic stage, we also examined the scalar frontogenesis function (Schultz and Doswell 1999; Yanase et al. 2022); we analyzed the PT field at 500 m MSL (Figs. 11a–c) because it was difficult to interpret the EPT field at 850 hPa owing to the modification by convection. The data were smoothed by a 400-km binomial filter. In the baroclinic stage, the convection was organized in the frontogenesis region around the steep gradient of PT on the northern and eastern side of the cyclone center (Fig. 11a; see also Fig. 10g). This frontogenesis was partly attributed to the deformation term in the frontogenesis function (not shown). The frontogenesis was also affected by positive diabatic and divergence terms and by negative tilting term, implying feedbacks from the vertical motion in the direct circulation. In the intermediate and convective stages, this frontogenesis area separated from the cyclone center and dissipated (Figs. 11b,c).

Fig. 11.
Fig. 11.

Lower-tropospheric thermal characteristics for Kirogi in the simulation. (a)–(c) PT (K; colors), frontogenesis (thick solid contours every 1 × 10−9 K m−1 s−1), frontolysis (thick dashed contours every 1 × 10−9 K m−1 s−1 with gray shading), and horizontal wind vectors at ∼500 m MSL. (d)–(f) Total surface heat flux (W m−2; colors) and horizontal wind vectors at 20 m MSL. (left) 1200 UTC 6 Aug, (center) 1200 UTC 7 Aug, and (right) 1200 UTC 8 Aug. The length scales of the vectors are shown in bottom-right corner in the right panels. Thin contours represent SLP every 5 hPa. Crosses denote the cyclone centers.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

As the heat flux from the ocean surface plays an important role in the development of a TC, the model output of the total heat (sensible heat plus latent heat) flux during the transition period are presented in Figs. 11d–f. In the baroclinic stage, large heat flux occurred over a relatively broad area north and northeast of the cyclone center where EPT was relatively low and horizontal wind speed was large (Figs. 10g,11d). This distribution tended to reduce the EPT gradient indicating negative feedback on the baroclinicity in the lower troposphere. Note that the large wind speed north and northeast of the cyclone center was also observed by the Oceansat scatterometer at 1230 UTC 6 August (not shown). In the intermediate stage, the cyclone center was separated from this large heat flux area and was accompanied by small heat flux (Fig. 11e). In the convective stage, however, the large heat flux was concentrated near the cyclone center (Fig. 11f), giving positive feedback over the high EPT area (Fig. 10i).

The synoptic-scale disturbances in the upper and lower troposphere also affected the evolution of vertical shear. The simulated vertical shear that temporarily exceeded 10 m s−1 in the baroclinic stage had a northwesterly component consisting of an upper-level westerly flow and a lower-level southeasterly flow (Fig. 4b), which is also in good agreement with the vertical shear based on the JRA-55 reanalysis (Fig. 4a). Figures 10a–c presents the horizontal distributions of vertical shear of the horizontal winds between the dynamic tropopause (Figs. 10d–f) and 850-hPa pressure level (Figs. 10g–i); although the analyzed upper levels differ between Fig. 4 (200 hPa) and Fig. 10 (dynamic tropopause), this difference does not qualitatively affect the following characteristics. In the baroclinic stage, at the upper level, the zonal flow was strong westerly owing to the approach of the cold disturbance, whereas the meridional flow was small due to offsetting of southerly and northerly components on the east and west side of the cold disturbance, respectively (Fig. 10d). At the lower level, a southeasterly flow was intense between the cyclone and a subtropical ridge to the northeast (Fig. 10g). By the intermediate stage, the upper-level westerly flow weakened as the upper-level cold disturbance wrapped around the cyclone center and began to dissipate, which resulted in the decrease in vertical shear (Figs. 10b,e). The changes in the upper-level winds were also observed in the atmospheric motion vector analysis (Fig. 3). In summary, the temporary increase in the vertical shear in the baroclinic stage is attributed to the approach of upper-level cold disturbance and its dissipation. The influence of the vertical shear on cyclone dynamics is discussed in section 5b.

c. Mesoscale vortex dynamics

As the thermal structure changed during the transition period, we also compared the vortex dynamics of the baroclinic and convective stages. Figure 12 shows relative vorticity along with two major terms in the vorticity equation at 500 m MSL around the cyclone center; the terms were calculated using a 400-km binomial filter. In the baroclinic stage, the vorticity was high in the northern and eastern parts of the cyclone, where active convection stretched vorticity (Figs. 12a,d). The vorticity was also high near the cyclone center, although convection was inactive. In this area, the vorticity was advected from the north by the northerly flow relative to the cyclone motion (Fig. 12g); i.e., the origin of the high vorticity at the cyclone center was the stretched vorticity in the convective area in the northern and eastern parts. The tilting effect and vertical advection were relatively small at this level (not shown). In the intermediate stage, however, the vorticity became concentrated symmetrically around the cyclone center, and was enhanced directly by the local convection through the stretching effect (Figs. 12b,e), but the vorticity was reduced by the horizontal advection of weak vorticity from the outer region (Fig. 12h). These processes were enhanced during the convective stage (Figs. 12c,f,i), which is consistent with the dynamics of a typical TC.

Fig. 12.
Fig. 12.

Vortex dynamics in the simulation. (a)–(c) Relative vorticity at ∼500 m MSL (10−4 s−1; red shading), ascent at ∼5 km MSL (thick contours at 0.1, 0.2, and 0.5 m s−1), and SLP (thin contours every 5 hPa). (d)–(f) Stretching term (10−8 s−2; colors) and 1.0 × 10−4 s−1 relative vorticity (thick contours) at ∼500 m MSL. (g)–(i) Horizontal advection term (10−8 s−2; colors), storm-relative horizontal wind vectors [the length scale is shown in bottom-right corner in (i)], and 1.0 × 10−4 s−1 relative vorticity (thick contours) at ∼500 m MSL. (left) 1200 UTC 6 Aug, (center) 1200 UTC 7 Aug, and (right) 1200 UTC 8 Aug. Cyclone centers are located at the centers of the panels (crosses).

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

The baroclinic stage was characterized by rapid northward shift of the cyclone center (Fig. 13). At 0000 UTC 6 August, a local minimum of SLP developed northeast of the cyclone center owing to strong convection (Fig. 13a; see also Fig. 8a). By 0600 UTC, the cyclone center had shifted 1.5° northward toward this local minimum (Fig. 13b). This process continued until around 1200 UTC (Fig. 13c). Thus, the strong convection north of the cyclone center appears to have played an important role in shifting the cyclone center a large distance northward. Note that this process in the Earth-relative framework corresponds to the vorticity advection from the convective region to the cyclone center by the northerly flow in the cyclone-relative framework (Fig. 12g).

Fig. 13.
Fig. 13.

Relative vorticity at ∼500 m MSL (red shading) and SLP (blue contours every 1 hPa) in the simulation. (a) 0000, (b) 0600, and (c) 1200 UTC 6 Aug. Black curves indicate the tracks of Kirogi. Small and large black circles indicate 1- and 6-hourly locations of Kirogi, respectively, where the digits denote 0000 UTC from 6 to 8 Aug. White circles correspond to the time for each panel. Black arrows point to the local minima of SLP.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

d. Azimuthally averaged characteristics

We examined the azimuthally averaged structure of Kirogi because it is fundamental information on TCs. To examine the evolution of the radial structures of the cyclone, Fig. 14 shows radius (r)–time Hovmöller diagrams of vertically integrated condensed water (the sum of cloud water, cloud ice, rain, snow, and graupel) as a proxy for convection, azimuthal wind at 500 m MSL, and total surface heat flux. In the baroclinic stage, intense convection and wind occurred outside the radius of 250 km at 0000 UTC 6 August (Figs. 14a,b), and then shifted inward to the radius between 100 and 200 km by about 1200 UTC 6 August; note that the intense convection and wind mainly occurred in the northern and eastern parts of the cyclone (see Figs. 8b, 11a). In the intermediate stage, convection and wind were weak at 0000 UTC 7 August, but then started to intensify at the radius between 50 and 150 km around 1200 UTC 7 August. In the convective stage, convection and wind were enhanced, particularly around the radius of 100 km, which is consistent with structures of typical TCs. The evolution of surface heat flux was similar to that of azimuthal wind (Figs. 14b,c). Thus, the cyclone size determined by the radius of peak azimuthally averaged condensed water and wind was large during the baroclinic stage when the cyclone was asymmetric, and then became small during the convective stage when the cyclone became more symmetric.

Fig. 14.
Fig. 14.

Radius–time Hovmöller diagram of azimuthally averaged properties about the cyclone center in the simulation. (a) Vertically integrated condensed water (kg m−2). (b) Azimuthal wind at ∼500 m MSL (m s−1). (c) Total surface heat flux (W m−2). The double-headed arrows with labels “B,” “I,” and “C” indicate the baroclinic, intermediate, and convective stages, respectively.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

The evolution of the vertical structure in the inner core (r < 100 km) is presented in time–height (z) Hovmöller diagrams (Fig. 15). The variables changed in a different manner in the lower and upper troposphere (below and above about 6 km). In the baroclinic stage, the low-level vorticity increased and deep ascent was enhanced temporarily around 1200 UTC 6 August (Figs. 15a,b). At this time, the cyclone center approached the convection in the northern part (Figs. 8, 12, and 13). Later in this stage, the ascent dissipated and the low-level vorticity decreased slightly. In contrast, the upper-level vorticity increased and the upper-level PT decreased (Figs. 15a,c), indicating the intrusion of the upper-level cold disturbance toward the cyclone center (Fig. 10d). As a result, Kirogi had a low-level warm-core structure and an upper-level cold-core structure in this stage, which is consistent with the cyclone phase space (Fig. 9). In the intermediate stage, the low-level vorticity developed slowly in the presence of moderate ascent (Figs. 15a,b), whereas the upper-level cold disturbance dissipated gradually due to moderate diabatic heating (Figs. 15c,d). In the convective stage, the low-level vorticity was enhanced owing to the deep convection, and a deep warm-core structure developed due to large diabatic heating.

Fig. 15.
Fig. 15.

Time–height Hovmöller diagram of properties averaged within a 100-km radius about the cyclone center. (a) Relative vorticity (10−4 s−1). (b) Vertical wind (10−2 m s−1). (c) PT anomaly (K) from the outer region (an average between the 900- and 1000-km radius from the cyclone center). (d) Diabatic heating associated with cloud microphysics and cumulus parameterization (K h−1). (e)–(i) Terms in the circulation budget at a 100-km radius (10−8 s−2): (e) time tendency; (f) the sum of stretching, eddy flux, tilting, and solenoidal terms; (g) stretching term; (h) eddy flux term; and (i) tilting term. The double-headed arrows with labels “B,” “I,” and “C” indicate the baroclinic, intermediate, and convective stages, respectively.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

The budget of the circulation around the circle with radius of 100 km from the cyclone center was examined (Figs. 15e–i) to support the vorticity equation analysis (Fig. 12). The terms were divided by the area so that they were identical to the contribution of the time tendency of the area averaged vorticity based on Stokes’s theorem. Important characteristics of the circulation tendency (Fig. 15e) were in good agreement with the total (Fig. 15f) of stretching (Fig. 15g), eddy flux (Fig. 15h), tilting (Fig. 15i), and solenoidal terms (not shown because of small magnitude compared to the other terms) except for a discrepancy below 1 km MSL. In particular, the dominant characteristic was the positive tendency of the lower-tropospheric circulation around 1200 UTC 6 August corresponding to increase in vorticity (see also Fig. 15a). This is attributed to the eddy flux effect (Fig. 15h), which corresponds to the horizontal advection of vorticity from the convective region in the northern and eastern parts into the cyclone center owing to the cyclone-relative northerly flow (see also Fig. 12g). Whereas the tendency became relatively small and fluctuating in the intermediate and convective stages, the tendency in the lower troposphere was positive on average, indicating the intensification of Kirogi. This positive tendency is attributed to the stretching and tilting terms (Figs. 15g,i). The stretching effect was caused by the ascent near the cyclone center (see also Figs. 12b,c,e,f and 15b), which was particularly large below 1 km MSL. The tilting term was negative below 1 km MSL and positive above (Fig. 15i), indicating upward transport of the large circulation from low levels (Davis and Galarneau 2009). The discrepancy in the circulation budget below 1 km MSL is considered to result from the effects of surface friction and turbulent mixing which were not estimated due to the lack of model output. When convection was enhanced (e.g., around 0000 UTC 8 August), the stretching effect increased in the lower troposphere, and the tilting effect increased in the upper troposphere. The large tilting effect in the upper troposphere was associated with a convective updraft near the radius of 100 km which increased and decreased vorticity inside and outside the radius, respectively, in the vorticity budget analysis (not shown). This characteristic is in good agreement with the schematic of tilting term illustrated in Fig. 1 in Davis and Galarneau (2009). This positive tilting term was offset by the negative stretching and eddy flux terms associated with outflow.

In summary, the vortex structure and dynamics also differed between the baroclinic and convective stages; the latter was related to typical TC dynamics more closely, although the cyclone was located at relatively high latitudes.

5. Discussion

a. Characteristics of SC and TT

We discuss whether Kirogi can be classified as an SC and a TT event around the time of its genesis by comparing our analysis of Kirogi with those in previous studies of SCs and TT particularly over the North Atlantic.

Kirogi’s genesis in the baroclinic stage occurred under the influence of an upper cold disturbance (Figs. 5, 6, 10), which is a typical synoptic-scale condition for SCs and TT (Davis and Bosart 2003; Evans and Guishard 2009; Hulme and Martin 2009a,b; McTaggart-Cowan et al. 2013; Bentley and Metz 2016; Bentley et al. 2016, 2017). The resultant shallow warm-core structure of Kirogi seen in the cyclone phase space (Fig. 9) can be classified as an SC according to Evans and Guishard (2009). The shallow warm-core structure was also identified in the PT anomaly field averaged over a 100-km radius about the cyclone center (Fig. 15c).

The upper cold disturbance extended southward from the midlatitudes and approached Kirogi from the northwest (Fig. 5). This characteristic resembles the meridional trough category in Bentley et al. (2017). Their composite analysis for this category demonstrated that troughs first induce convection on their eastern side owing to quasigeostrophic forcing for ascent, and then support SC formation owing to reduced coupling index through the vertical alignment of the troughs with surface cyclones. This evolution is consistent with our Q vector and coupling index analyses from the baroclinic to intermediate stages (Figs. 6 and 10).

The structural evolution of Kirogi in a moderate baroclinic environment from the baroclinic to intermediate stages partly resembled the occlusion process of a mature extratropical cyclone (Figs. 3, 8). In the baroclinic stage, convection in the northern and eastern parts of the cyclone was attributed to frontogenesis in the lower troposphere (Fig. 11a) embedded within a larger region of quasigeostrophic forcing for ascent (Fig. 6g). High values of vorticity produced by the convection in this frontogenesis area were advected horizontally into the cyclone center (Figs. 12g, 15h); this is a similar process to that of an SC case over the North Atlantic (Hulme and Martin 2009b) and an occlusion (or warm seclusion) of an extratropical cyclone (Galarneau and Weisman 2020). The cyclone center was also attracted to the convective area to the north, and shifted northward rapidly (Fig. 13), which was also observed in the SC case by Hulme and Martin (2009b).

In the intermediate stage, the warm and moist air near the cyclone center becomes separated from the tropical air in the lower troposphere (Figs. 6e, 10h, 11b), while the upper cold disturbance wrapped around Kirogi (Figs. 5e, 6b, 10e). These characteristics were also observed in the TT events over the North Atlantic (Hulme and Martin 2009a,b), and resembled an occluded thermal structure in the lower troposphere and a treble clef PV structure in the upper troposphere of an occluded extratropical cyclone (Posselt and Martin 2004). The analysis of EPT surfaces also indicated a structure that was partly similar to a trough of warm air aloft (trowal; Posselt and Martin 2004) in an occluded extratropical cyclone (not shown), although it was much more difficult to determine the EPT surface at the lower latitudes.

The vertical shear temporarily exceeded 10 m s−1 around the genesis time in the baroclinic stage (Fig. 4), which was attributed to the approach of an upper cold disturbance (Fig. 10). In the intermediate stage, the vertical shear was reduced because the upper cold disturbance wrapped around the cyclone center and dissipated. This reduction in the vertical shear associated with the structural change in an upper cold disturbance was similar to the TT process of Hurricane Michael (2000) analyzed in Davis and Bosart (2003). In addition, the evolution of Kirogi’s cloud pattern around the baroclinic stage (Figs. 3a–c, 8a–c) appears to be the shift from down-shear convection to up-shear convection in a relative sense; the latter can reduce upper-level PV in large area around the cyclone center due to divergent flow and efficiently eliminate the vertical shear in the area (Davis and Bosart 2004).

The environment during the TT process of Kirogi was characterized by a combination of moderate horizontal temperature gradient and SST exceeding 26.5°C in the subtropics (Fig. 6), which is consistent with the climatological study of SCs over the North Atlantic (Guishard et al. 2009). Idealized numerical simulations also support the idea that such environments are responsible for SC-like structures and TT-like evolution (Yanase and Niino 2018, 2019). August in 2012 was characterized by steep horizontal gradient of temperature compared to the climatology in the subtropics around the central Pacific owing to a cold anomaly on its north side (Figs. 7e,f). This enhanced baroclinicity appears to be favorable for forcing ascending motion through the Q-vector convergence and frontogenesis during the cyclogenesis of Kirogi. The northeast–southwest temperature gradient around Kirogi was different from the more meridional temperature gradient associated with TT events over the North Atlantic analyzed in Hulme and Martin (2009a); as a result, the horizontal pattern of Kirogi appears to be rotated somewhat clockwise from those of the North Atlantic cases. The cold anomaly in the troposphere north of Kirogi may be partly attributed to the enhanced TUTT pattern (Fig. 7b), because a positive anomaly of upper-level PV is accompanied by a negative temperature anomaly beneath it. It is worth noting that the geneses of Typhoon Damrey (28 July) and Typhoon Haikui (3 August) in 2012 also occurred near the enhanced TUTT. Another remarkable characteristic was an enhanced ridge to the north of the genesis location of Kirogi (Fig. 7d). According to a composite analysis of TCs near upper-level PV streamers over the North Atlantic (Galarneau et al. 2015), intense high pressure systems to the north of the TCs can support the development of the TCs in the following manner: 1) enhanced easterly flow between the TCs and the high pressure systems increases surface heat flux and westward moisture transport; 2) these moist processes facilitate an upshear propagation of convection, which can help to reduce vortex tilt and to aid TC development (Rappin and Nolan 2012). The numerical simulation of Kirogi consistently shows that the surface heat flux was largest on the north side of the cyclone (Fig. 11d) and the convection shifted to the relatively upshear side of the cyclone center (Figs. 8a–c) around the baroclinic stage. As mentioned earlier, upshear convection can help to reduce vertical shear around the cyclone center (Davis and Bosart 2004).

b. Vertical shear and vortex reformation in the baroclinic stage

The temporary increase in vertical shear during the baroclinic stage appears to have influenced the cyclone formation, although strong shear is generally considered to be unfavorable for TC development. To examine the sensitivity of the vertical shear to its definition, we averaged vertical shear in annuli of different sizes using the JRA-55 reanalysis for 1200 UTC 6 August (Fig. 16a). We refer to an annulus with the inner radius of ri (0–500) km and the outer radius of ro (200–1600) km as a ri–ro-km annulus. Note that the analysis in Fig. 4 corresponds to the 200–800-km annulus. When the inner radius was fixed at 0 km, the vertical shear was strong northwesterly for large outer radii (the 0–800-km, 0–1200-km, and 0–1600-km annuli). On the other hand, the vertical shear was weak southwesterly for small outer radius (the 0–200-km annulus), which is strongly affected by the local flow to the east of the upper-level cold trough associated with a PV streamer (Figs. 10a,d). The 0–200-km annulus appears to be too small to interpret an environment for the asymmetric convection with horizontal scale exceeding 200 km (Fig. 10a). When the outer radius was fixed at 800 km, the vertical shear was less sensitive to the inner radius (the 0–800-, 200–800-, and 500–800-km annuli) because the area inside the inner radius was relatively small. These sensitivities should be taken into account, when our result is compared with other case studies. We chose the 200–800-km annulus because it was used in several studies (Kaplan and DeMaria 2003; Fischer et al. 2017). The sensitivity of vertical shear to annulus size for the simulation (Fig. 16b) was similar to that for the JRA-55 reanalysis, supporting the credibility of the simulation.

Fig. 16.
Fig. 16.

Vertical wind shear (m s−1) around the cyclone center averaged in annuli with different radii at 1200 UTC 6 Aug for (a) the JRA-55 reanalysis and (b) the simulation. In (b), the plot of the 0–1200-km annulus is underneath that for 500–800-km annulus, because the two values are similar to each other.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

Statistical studies showed that some tropical cyclogeneses occur even within strong vertical shear exceeding 10 m s−1 (Bracken and Bosart 2000) including TCs under the influence of upper-level disturbances (Fischer et al. 2017). Jones (1995) discussed two different roles of vertical shear (her Fig. 4): 1) the vertical shear tilts a vortex, forcing ascending motion with adiabatic cooling on the downshear side of the vortex; then the axis of the tilted vortex rotates cyclonically due to the interaction between upper and lower vortices, forcing ascending motion 90° to the right of the tilt direction due to the motion along an isentropic surface (vortex tilt mechanism); 2) the cyclonic circulation of a vortex within the horizontal temperature gradient balanced with the vertical shear advects warm air causing ascending motion on the downshear side of the vortex (warm-air advection mechanism). Idealized experiments in Yanase and Niino (2019) demonstrated that the warm-air advection mechanism is characterized by a large horizontal scale, a warm anomaly on the downshear side, and a upshear tilt of a vortex, whereas the vortex tilt mechanism is characterized by a small horizontal scale, a cold anomaly on the downshear side, and a downshear tilt of a vortex. Diabatic processes also affect the location and intensity of convection (Rappin and Nolan 2012; Rogers et al. 2020).

Nguyen and Molinari (2015) demonstrated downshear reformation of Tropical Storm Gabrielle (2001) using a nonhydrostatic model with 1-km horizontal grid spacing. The environmental shear of about 10 m s−1 tilted the vortex of Gabrielle downshear left, which enhanced asymmetric convection and produced a mesovortex. This mesovortex finally developed into a dominant vortex of Gabrielle. This process in Gabrielle was similar to the reformation of the cyclone center near the convection downshear-left of the original vortex and the resultant rapid northward shift of Kirogi in the baroclinic stage (Fig. 13). On the other hand, there were also several differences between the two cases; the following discussion is based on the vertical shear averaged in the 0–500-km annulus (west-northwesterly shear; see Fig. 16b) for consistency with Nguyen and Molinari (2015). 1) The tilt of the vortex in Kirogi was more complicated than that in Gabrielle. Figure 17 shows the centroids of pressures at levels between 0 and 10 km MSL; to mitigate the influence of mesoscale convection on pressure fields, the centroids were calculated within a circle of radius R (100 km) as described in section 2c except that Penv is set at the maximum pressure within a circle of radius R at each level. The centroids at 5 and 10 km MSL were located upshear-left (north-northwest) and upshear-right (west-southwest) of the centroid at 0 km MSL, respectively. Apparently, the centroids above 8 km MSL were strongly affected by the upper-level cold disturbance (see also Fig. 10d). 2) The asymmetric distribution of convection in Kirogi (Fig. 10a) occurs on a larger scale than that in Gabrielle (Fig. 7 in Nguyen and Molinari 2015). 3) The temperature field was also different between Kirogi and Gabrielle. Figure 18 shows the radial–vertical structure for shear-relative quadrants of Kirogi at 1200 UTC 6 August. In the lower troposphere, the largest PT anomaly occurred in the downshear-right quadrant because of the horizontal temperature gradient of the environmental fields (see also Fig. 11a). The deep ascent in the downshear-left quadrant was associated with a positive PT anomaly below 2 km MSL due to warm-air advection (see also Fig. 11a) along with diabatic heating within the ascent (not shown). In contrast, the ascent in the downshear-left quadrant of Gabrielle was accompanied by a negative PT anomaly in the lower troposphere (Figs. 6 and 8 in Nguyen and Molinari 2015), which was attributed to the adiabatic cooling associated with the vortex tilt. Thus, the structures of Kirogi during the vortex reformation appear to be affected by the baroclinicity and the upper-level disturbance.

Fig. 17.
Fig. 17.

Pressure (contours every 1 hPa) at 0 km (red), ∼5 km (green), and ∼10 km (blue) MSL and ascent at ∼5 km MSL (light, medium, and dark gray shadings for 0.1, 0.2, and 0.5 m s−1, respectively). Circles connected by black lines denote the centroids of pressure every 1 km from 0 to 10 km MSL, with large circles for 0 km (red), 5 km (green), and 10 km (blue) MSL. The arrow at the top right denotes the shear vector averaged in the 0–500-km annulus.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

Fig. 18.
Fig. 18.

Radius–height distribution of azimuthally averaged properties in each shear-relative quadrant at 1200 UTC 6 Aug. Potential temperature anomaly from the average of the four quadrants (K; colors) and vertical wind (black contours every 0.1 m s−1; solid and dashed contours denote positive and negative values, respectively). (a) Upshear-left quadrant, (b) downshear-left quadrant, (c) upshear-right quadrant, and (d) downshear-right quadrant.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

c. Predictability

It is worth describing the predictability of the TT process of Kirogi. Figure 2 presents some examples of sensitivity experiments. When a cumulus scheme was not used (EXP1), a cyclone hardly developed. However, this does not simply indicate that a cumulus scheme was necessary for a simulation with 2.5-km horizontal grid spacing, because the results also depended on initial conditions. EXP2 and EXP3 are simulations with and without the cumulus scheme, respectively, initialized at 1200 UTC 5 August 2012 (12 h later than the main simulation). Both experiments reproduced the development of Kirogi to some extent, and show little difference until 9 August. Therefore, the simulation of Kirogi’s TT process was very sensitive to experimental design. This result is consistent with the recent report on phase transitions (Wood et al. 2023, manuscript submitted to Trop. Cyclone Res. Rev.) in that operational predictions of TT events were less reliable than those of the average tropical cyclogenesis events.

The distribution of vorticity and ascent in the sensitivity experiments are presented in Fig. 19. At the baroclinic stage, convection was more or less organized on the north and east side of the cyclone in all the experiments (Figs. 19a–c; see also Fig. 12a). These results confirm that baroclinic processes including quasigeostrophic forcing for ascent and frontogenesis were responsible for the large-scale organization of convection. However, convection in EXP1 was organized less tightly than the other experiments. This was followed by less concentrated convection and vorticity near the cyclone center at the convective stage in EXP1 than in the other experiments (Figs. 19d–f; see also Fig. 12c). These results imply that the baroclinically organized convection was responsible for the subsequent formation of the TC.

Fig. 19.
Fig. 19.

Relative vorticity at ∼500 m MSL (10−4 s−1; red shading) and ascent at ∼5 km MSL (thick contours at 0.1, 0.2, and 0.5 m s−1) in the sensitivity experiments at (a)–(c) 1200 UTC 6 Aug and (d)–(f) 1200 UTC 8 Aug. (left) EXP1, (center) EXP2, and (right) EXP3. Thin contours represent SLP every 5 hPa. Cyclone centers are located at the centers of the panels (crosses).

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

6. Summary

We analyzed the structural transition during and after the cyclogenesis of Tropical Storm Kirogi in August 2012 based on observations, a reanalysis dataset, and a numerical simulation. We approximately separated the transition period into a baroclinic stage on 6 August, an intermediate stage on 7 August, and a convective stage on 8 August as illustrated in Fig. 20.

Fig. 20.
Fig. 20.

Schematic illustration of the TT process of Kirogi at (a) the baroclinic stage, (b) the intermediate stage, and (c) the convective stage. Orange and cyan areas denote a low-level warm moist air and an upper-level cold trough, respectively, with arrows indicating the horizontal flow at each level. Gray scalloped areas denote clouds. Green curved bands in (a) and (b) indicate frontogenesis. A purple arrow in (a) indicates the vertical shear vector. “L” indicates the center of the cyclone at the lower troposphere.

Citation: Monthly Weather Review 151, 10; 10.1175/MWR-D-22-0190.1

In the baroclinic stage including the genesis time (Fig. 20a), Kirogi was characterized by a comma-shaped cloud pattern with a wide cloud area north and east of the cyclone center. The convection was organized in the frontogenesis area and produced large relative vorticity by stretching. This vorticity was advected horizontally toward the cyclone center in the cyclone-relative framework, whereas it shifted the cyclone center northward in the Earth-relative framework. In the upper troposphere, a high PV streamer or cold trough extended southward from the midlatitudes to the western side of Kirogi. The vertical shear was increased temporarily under the influence of this upper-level disturbance. The cyclone had a shallow warm-core structure like SCs in previous studies.

In the intermediate stage (Fig. 20b), the western end of the comma cloud began to wrap up, and the warm and moist air near the cyclone center became separated from the tropical airmass to the southeast. In the upper troposphere, the high PV disturbance wrapped around Kirogi. The occluded thermal ridge in the lower troposphere and a treble clef PV distribution in the upper troposphere were similar to those in occluded extratropical cyclones. The decrease in vertical shear by this time was associated with the weakening of the upper-level high PV, which appears to be promoted by the divergent flow caused by the upshear convection.

In the convective stage (Fig. 20c), the warm and moist air was isolated from the tropical airmass in the lower troposphere, and the upper cold disturbance almost dissipated. As a result, Kirogi had a deep warm-core structure. Deep convection was organized near the cyclone center, and directly enhanced the vorticity at the center. The surface heat flux was also concentrated near the cyclone center.

Taken together, the structural transition of Kirogi resembled TT events from an SC to a TC in the North Atlantic. Compared with the TT of Typhoon Peipah over the western North Pacific (Chang et al. 2019), the TT of Kirogi occurred at relatively high latitudes, showed a shallow warm-core structure like an SC, and was associated with a strong baroclinic process that directly triggered the formation of a tropical storm. It should be also noted that the baroclinic process during Kirogi’s genesis occurred in an enhanced baroclinic environment in the subtropics over the central North Pacific in August 2012.

As the TT processes are case-specific, more case studies are necessary to understand SCs and TT over the western North Pacific and to construct their global perspective. We also expect that a comprehensive understanding of SCs and TT will provide a more complete picture of a wide spectrum of cyclones (Hart 2003; Yanase and Niino 2019; Wood et al. 2023, manuscript submitted to Trop. Cyclone Res. Rev.) by serving as a bridge between diabatically driven cyclones like TCs and baroclinically driven cyclones like extratropical cyclones.

Acknowledgments.

We acknowledge helpful comments by Drs. Masahiro Hayashi, Akiyoshi Wada, Kazuto Takemura, Takeshi Horinouchi, Ryuichi Kawamura, Sachie Kanada, Ron McTaggart-Cowan, Jonathan E. Martin, and Ian Morrison, and thorough reviews by three anonymous reviewers and Dr. Clark Evans. The experiments were conducted using Fujitsu Server PRIMERGY CX2550 in the Meteorological Research Institute. This research was supported by JSPS KAKENHI Grants JP19H01973, JP19H05696, JP19H00705, and JP21H01164, and by the Cooperative Program (140, 2022) of Atmosphere and Ocean Research Institute, The University of Tokyo.

Data availability statement.

The JMA best track product is available at https://www.jma.go.jp/jma/jma-eng/jma-center/rsmc-hp-pub-eg/besttrack.html. The JRA-55 reanalysis is available at https://jra.kishou.go.jp/JRA-55/index_en.html. The atmospheric motion vectors and the JMA nonhydrostatic model are made available under a contact with the JMA, because these are basically operational products collected and developed by JMA. The output data from the numerical simulation is too large to be publicly archived with available resources, and are available upon request to the corresponding author.

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