WMO/CAS/WWW

FIFTH INTERNATIONAL WORKSHOP ON TROPICAL CYCLONES


Topic 4.3: Tropical cyclone formation
Large-scale control and mesoscale influences; Intraseasonal


Rapporteur: J. L. Evans
Department of Meteorology
The Pennsylvania State University
503 Walker Building
University Park, PA 16802, USA

Email: evans@essc.psu.edu
Fax: 1-814-865-3663

Working Group: Bill Frank, Miloud Bassafi, Klaus Fraedrich, Lloyd Shapiro, Tetsuo Nakazawa


4.3.1: Introduction

Six features of the large-scale tropics were identified by Gray (1968) as necessary conditions for tropical cyclogenesis: sufficient ocean thermal energy [SST>26_C to a depth of 60m], enhanced mid-troposphere relative humidity, conditionally unstable atmosphere, enhanced lower troposphere relative vorticity, weak vertical shear of the horizontal winds at the genesis site and displacement by at least 5_ latitude away from the equator. The first three – thermodynamic – parameters have been identified as seasonal indicators of genesis potential, while the latter– dynamical – parameters are a measure of the daily likelihood of genesis (McBride and Zehr 1981), although the presence of all of these environmental features is not sufficient to ensure tropical cyclogenesis. Thus, the necessary, but not sufficient, criteria for tropical cyclogenesis may be summarized as: the ability to support deep convection in the presence of a low-level vorticity maximum, which reduces the local Rossby radius of deformation and focuses the convective heating locally (Ooyama 1982). This broad description of tropical cyclone development forms the basis of the CISK [Conditional Instability of the Second Kind] (Ooyama 1963; Charney and Eliassen 1964), WISHE [Wind Induced Surface Heat Exchange] (Emanuel 1986; 1995a), and “free ride” CISK (Fraedrich and McBride 1989) theoretical models.

Over the past decade, the role of the Madden Julian Oscillation [MJO] in modulating the intraseasonal variability of convection in the global tropics has been well documented and its effects on the likelihood tropical cyclogenesis have been elucidated. Interactions across scales and the mesoscale structure of the evolving tropical convective system have been documented and genesis theories that include these interactions are being developed. The effects of equatorial waves in organizing convection and energy on the tropical storm scale have been explored, as has the role of convection in spinning up a tropical storm from a weak incipient vortex. Finally, the implications of this work on potential changes in formation in other climate regimes have been elucidated.

While improved observations and field campaigns allow for better resolution of the mesoscale details of formation, the question remains as to whether this is essentially a deterministic or a stochastic process (e.g., Simpson et al. 1997; Ooyama 1982). Is it possible to identify an individual incipient disturbance that will develop into a tropical storm or are larger scale environmental forcings the crucial factor for development? While we do not seek to answer this question here, we explore the range of large-scale and mesoscale influences on tropical cyclogenesis documented in recent years.

4.3.2: Intraseasonal Variations in Formation

McBride and Zehr (1981) identified the convective potential of a region as being associated with the seasonal signal of tropical cyclogenesis. More recently, the Madden Julian Oscillation [MJO; 30-60 day mode] (Madden and Julian 1972, 1994) has been linked to periodic organization of tropical convection within the storm season. Hendon and Liebmann (1994) and Evans et al. (2000) use different approaches to relate progression of the MJO through the tropical Indian and western Pacific basins to increasing organization of convection in those regions. At the onset of the MJO, convective activity is disorganized and relatively short lived. As the MJO approaches, the convective systems become increasingly more organized (more symmetric and longer mean lifetimes). By the peak of the MJO, tropical cyclogenesis is likely. Evans and Jaskiewicz (2001) analyze the evolution of western Pacific convection organization and subsequent tropical cyclogenesis through three MJO cycles during the TOGA COARE field phase and note an increase in strength of the axisymmetric Hadley circulation coincident with the increasing level of convective organization and activity as each MJO cycle progresses through the tropical western Pacific. An MJO cycle in eastern Pacific tropical cyclogenesis is also observed (Maloney and Hartmann 2000; Molinari and Vollaro 2000). Modulation of the likelihood of tropical cyclogenesis by the MJO is in good agreement with the three week active/ three week inactive periods of tropical cyclogenesis proposed earlier by Gray (1979).

Liebmann et al. (1994) suggest that the increase in likelihood of genesis for western Pacific storms is not due to the MJO acting as an inherent control on genesis, but rather that the increased convective activity overall provides many more seed systems from which a tropical storm could develop. The fraction of systems actually developing from convective cluster to tropical storm is unchanged between “favorable” [convection active] and “unfavorable” [convection suppressed] phases of the MJO.

Molinari et al. (1997) and Maloney and Hartmann (2000) characterized the effects of the MJO on tropical cyclogenesis in the eastern North Pacific. Provided these synoptic-scale easterly waves remain coherent in their passage from Africa to the Caribbean and into the eastern North Pacific, their likelihood of resulting in tropical cyclogenesis does not depend on their initial strength, but on phasing with the convectively active cycle of the MJO in regions of dynamic instability over the warm tropical waters. Eastern Pacific cyclogenesis is suppressed in the “unfavorable” phases of the MJO (Molinari and Vollaro 2000). As the MJO moved through the region, locations of genesis progressed eastward and northward and moved with the deep convection in the ITCZ (Molinari and Vollaro 2000).

As discussed below (Section 4.3.3), regions of preferred wave growth (and tropical cyclogenesis) are characterized by a sign reversal of the meridional gradient of large-scale potential vorticity [PV] field, which satisfies the Charney-Stern condition for dynamical instability (Charney and Stern 1962). Dickinson and Molinari (2000) attribute the maintenance of this PV gradient reversal to the in situ deep convection in the eastern Pacific. While this would reinforce the MJO signal for eastern Pacific genesis, the causality of this link is not clear (Section 4.3.3).
4.3.3: Large Scale Controls on Tropical Cyclone Formation

Each component of Gray’s genesis parameter plays a physical role in preconditioning the environment for genesis and so modulates storm genesis frequency. DeMaria et al. (2001) have developed an operational genesis parameter that incorporates 5-day means of vertical wind shear, vertical instability and mid-level moisture variables. They find that these large-scale parameters must all be favorable for genesis to occur, but this set does not contain a sufficiency criterion for genesis. Thus the search for the sufficiency condition and for the mechanisms for generating the incipient vortex continues.

Schubert et al. (1991) present a PV model for the Hadley cell and argue that the continuous PV source associated with convection in the ITCZ will act to destabilize and, through a combined barotropic-baroclinic instability, to break down the ITCZ periodically. Based on this model, they argue that Africa should not be a unique region for easterly wave formation and suggest that similar disturbances could also be generated in the Australian monsoon. Dickinson and Molinari (2000) perform a climatology of sign reversals of meridional PV gradients over Africa and Australia. They demonstrate that the source of the sign reversal in the PV in each region is the deep convection associated with the active monsoon. Dickinson and Molinari (2000) identify a 2–6 day peak in the variance of the African monsoon [outgoing longwave radiation (OLR) and low-level winds] consistent with easterly wave activity, but find no corresponding peak in the Australian data [either for 2-6 or 2-10 days]. They suggest that the lack of easterly wave activity could result from either the confined region of PV reversal [3000 km in Australia, cf 5000 km in Africa], the lack of topography in Australia [and thus a weaker easterly jet], or the stability of the Australian flow pattern.

Lander (1990) documents the role of easterly waves in tropical storm formation in the western North Pacific and Australian regions for the special case of twin tropical storms that straddle the equator at formation. He suggests that these storms may form in association with equatorial Rossby waves. Dickinson and Molinari (2002) diagnose the evolution of a mixed Rossby-gravity wave packet with slow eastward group velocity (with westward phase speed of individual waves) as a faster eastward MJO cycle traverses the tropical central and western Pacific. For their boreal summer case, convective heating straddles the equator east of 150°E, but Northern Hemisphere convection dominates west of this longitude. This longitudinal asymmetry of convection relates to the evolution of the mixed Rossby-gravity waves. In the eastern part of the domain, the waves remain equatorially trapped and symmetric. West of 150°E, the waves undergo a transition as the waves appear to track along the dynamic equator associated with PV development in the western Pacific ITCZ, and move away from the equator. A region of active convection forms on the northwest side of a near-equatorial wave trough (in the region of maximum low-level convergence for a mixed Rossby-gravity wave). This convective region moves away from the equator and eventually develops into a tropical storm. The ultimately formation of a tropical depression is attributed to enhanced Ekman pumping (Ekman 1905) in the evolving depression (consistent with CISK or free ride development described above).

Kuo et al. (2001) examined the evolution of wave activity over the western North Pacific in association with the background confluent flow, scale contraction of short Rossby waves, and nonlinear dynamics. In their barotropic model, nonlinear energy/enstrophy accumulation in the confluence zone results from continuous wave forcing upstream. The energy accumulation is maximized for zonal wavelengths near 2000 km. In a monsoon-like mean flow, the process provides for upstream dispersion from a downstream northwestward-propagating disturbance, which contributes to a new vortex formation. The overall process may contribute to a sequence of vortex formations in the confluent eastern end of the western North Pacific monsoon trough.

Ferreira and Schubert (1997) explore the barotropic mode of instability of the convecting ITCZ suggested by Schubert et al. (1991). Two modes of instability are identified: either the trough breaks down into a group of several cyclones, or it mixes and regains symmetry as one large system. The shape of the initial PV strip used to characterize the ITCZ, as well as the presence of another cyclone, affected which mode of breakdown occurred. The presence of an additional cyclonic vortex in the ITCZ sped its breakdown.

Briegel and Frank (1997) map the genesis locations of western North Pacific tropical storms. They partition the large-scale tropical environment into monsoon trough and ITCZ zones that are separated by a confluence zone. They find that 74% of tropical cyclogenesis events occur in the monsoon trough region, with 62% of these forming in the confluence zone. They note that genesis in this region may be separated by as little as 5-7 days, and attribute repeated genesis in this region to preconditioning of the low-level flow (increased convergence of tropical moisture) by the maturing storm as it moves away. However, this is also a zone of wave energy accumulation and such repeated genesis could also be attributed to dispersion of wave packets (Holland 1995; Molinari et al. 1997).

Briegel and Frank’s (1997) analysis of the relative frequency of tropical cyclogenesis in the monsoon trough (28%) and confluence zone (46%) is somewhat at odds with the results of Ritchie and Holland (1999), who find ~40% of storm genesis in the monsoon trough, with only ~30% of genesis events located in the confluence zone. The short time periods used in both studies (two and three years respectively) is likely the source of much of this discrepancy and highlights the effect of interannual variability on tropical cyclogenesis in this region. Ritchie and Holland (1999) suggest that genesis in the monsoon trough is due breakdown of the monsoon as proposed by Ferreira and Schubert (1997) and Schubert et al. (1991). In the confluence zone, genesis may be initiated through easterly waves that propagate into the region, possibly interacting with a TUTT (tropical upper-tropospheric trough). Both the wave’s presence and the possible interaction with the TUTT
1 will locally enhance low-level convergence and thus moist convection, and provide a favorable environment for tropical cyclogenesis. Cyclogenesis from waves embedded in the easterlies was documented for 12 (15%) of their cases, with formation due to wave energy dispersion (Zehnder 1991; Holland 1995) attributed as the genesis mechanism in 12% of cases. Formation in the monsoon gyre was rare (6%), due to the infrequent occurrence of this synoptic environment. However, once the gyre was present, six cases of genesis occurred in a three-week period (Ritchie and Holland 1999). Increasing organization of mesoscale convective systems [MCS] in the 72 h prior to genesis is observed within the monsoon trough and confluence zone developments which supports the hypothesis of mutual interaction between a mesoscale vortex and large-scale flow (Ferreira and Schubert 1997; Simpson et al. 1997; Holland 1995; Zehr 1992). The MCS development is far less apparent for the “pure” easterly wave developments.

4.3.4: Mesoscale Influences on Tropical Cyclogenesis

Mesoscale influences on tropical storm formation include effects contributing to (or limiting) the development of the incipient vortex or disturbance and the feedbacks between this disturbance and convection. These processes are considered here in turn.

a) Development of mesoscale incipient vortices

Monsoon depressions (e.g., Harr et al. 1996), African easterly waves (e.g., Zehnder 1991), and subtropical cyclones (Hebert and Poteat 1975) have long been understood to provide source disturbances from which, under appropriate thermodynamic conditions, a tropical storm could develop. While the tropical eastern Pacific was widely recognized as being a highly active region of tropical cyclogenesis, until the most recent decade there were few explanations for the source of the many incipient disturbances needed to support this genesis. Refraction of equatorial Rossby waves (Zehnder 1991) or amplification of the waves in the eastern Pacific ITCZ (Dickinson and Molinari 2002) are also advanced as sources of tropical cyclogenesis in this region. Zehnder (1991) examines the evolution of an n=2 Rossby wave and n=4 mixed Rossby-gravity wave advected in easterly flow across a mountain range representative of the Sierra Madres of Mexico. Interaction of the Rossby wave with the northwest-southeast oriented mountain (centered on 15°N) leads to the development of a lee trough and a secondary vorticity maximum that propagates downstream with period about 5 days (about half the wave period). Both regions are potential locations for tropical cyclogenesis. The faster mixed Rossby-gravity mode is centered on the equator and so interacts less with the mountain range to the north. Even still, a lee vorticity maximum develops on the southern part of the mountain ridge as this wave traverses the region. Zehnder and Gall (1991) document two case studies that provide observational support for this as a tropical storm formation mechanism.

Molinari et al. (1997) demonstrate that both the Caribbean and tropical eastern Pacific have regions in which there is a reversal in meridional gradient of the base state PV, which satisfies the Charney Stern condition for inertial instability (Charney and Stern 1962) and so identifies preferred regions for wave growth. Indeed, Kiladis and Wheeler (1995) observe that the eastern Pacific is a source region for equatorial Rossby waves. Molinari et al. (1997) note that African easterly waves tend to weaken and lose their convection in the central Atlantic and propose that the inertially unstable region in the Caribbean reinvigorates these waves, providing sufficient energy for them to complete their trek into the eastern Pacific. They examine the 1991 season and identify tropical storm development downstream of these inertially unstable zones consistent with the continued progression of equatorial waves across the Western Hemisphere tropics. Dickinson and Molinari (2002) propose that preferential wave growth occurs in the dynamically unstable and convectively active eastern Pacific ITCZ. They argue that the reversal in the mean meridional PV gradient here is maintained by convectively-forced PV generation. In contrast, Tomas and Webster (1997) attribute the presence of the inertial instability to cross-equatorial advection of absolute vorticity driven by high cross-equatorial pressure gradients in the ITCZ regions (due to land/sea or SST gradients), and see the convection as a consequence of this dynamic environment. Viewing the convection as a consequence of the dynamic instability, rather than as its driver, is consistent with the lag of the convective heating compared to the peak in the PV gradient reversal observed by Molinari et al. (1997).

Simpson et al. (1997) track the evolution of Tropical Cyclone Oliver (1993) from a cluster of weak mesovortices, through merger of these systems into two MCS. Subsequently, one MCS developed an eye structure and sheared the other system, which became the tropical storm rainband. As these processes were occurring, the large-scale monsoon trough was strengthening, and thus reducing the Rossby deformation radius of the environment and concentrating the effects of the convective heating locally. Such genesis from accumulation of mesoscale vorticity had not been observed previously. With the ongoing improvements in observation capability in the tropics, largely from remotely sensed or operated data platforms, it is becoming possible to survey the tropics and to determine the frequency of such genesis events. Studies of vortex intensification and in the presence of an asymmetric PV anomaly for barotropic (Möller and Montgomery 1999; Enagonio and Montgomery 2001) and balanced baroclinic (Möller and Montgomery 2000) vortices provide theoretical support for this genesis mechanism.

Bosart and Bartlo (1991) and Davis and Bosart (2001, 2002) explore the internal and environmental effects on the evolution of Hurricane Diana (1984) from an initially subtropical system. They identify three phases of internal development: a subtropical vortex initially spins up at the southern extreme of a quasi-stationary front; the transformation from cold to warm core occurs as remotely [convectively] generated PV anomalies are forced on the poleward side of the vortex due to interactions with an upper-level trough-ridge couplet – these PV anomalies are mixed into the vortex center; immediately after transition to warm core, the storm intensity is sustained as the convective structure [spiral bands] becomes organized; at last, the storm intensifies as a tropical system, and finally develops a clear eye.

Pfeffer and Challa (1992) and Challa et al. (1998) have explored the contribution to tropical cyclogenesis of eddy heat and momentum fluxes due to interactions with the storm environment – such as the upper trough-ridge couplet in the Diana case. They find that the eddy fluxes induce a surface radial inflow that enhances oceanic surface fluxes and moisture transports, and then intensifies a weak incipient vortex. In the absence of these fluxes, symmetric intensification mechanisms (such as CISK or WISHE) are insufficient to drive intensification of the same incipient vortex (Pfeffer and Challa 1992). Challa et al. (1998) demonstrate that the secondary (in-up-out) circulation due to fluxes of eddy angular momentum at upper levels is sufficient to drive this surface moisture convergence. The results of Tuleya (1991) suggest that the structure of the incipient disturbance is also likely important in determining whether or not it will develop.

No matter how the incipient vortex develops, it must remain coherent until it is in a convectively favorable environment for tropical cyclogenesis to occur. Thus, the response of the mesoscale incipient vortex to environmental forcing will also determine its likelihood of genesis. Jones (1995) demonstrates that a “dry” vortex in even weak sustained vertical wind shear will tilt in the vertical and inevitably decay with time. However, Schecter et al. (2002) develop a quasi-geostrophic theory for the response of a dry vortex to a brief period of vertical wind shear. They argue that the shear-induced tilt of the vortex core can be rectified through damping of the first internal mode asymmetries induced by the tilt and feedback of wave energy to the symmetric vortex. As foreshadowed by Montgomery and Kallenbach (1997), the vorticity axisymmetrization is sensitive to the radial gradient of storm vorticity. While Schecter et al. (2002) recognize that the quasi-geostrophic approximation is not ideal for a tropical cyclone-like vortex, they argue that this theory should logically extend to the asymmetric balance [AB] framework for tropical cyclones developed by Shapiro and Montgomery (1993). Indeed, Möller and Jones (1998) demonstrate the applicability of the AB theory for a tropical cyclone-like vortex in vertical wind shear, qualitatively reproducing the key elements of the primitive equation results of Jones (1995).

As discussed in Section 4.3.3, other mechanisms for tropical cyclogenesis involving mesoscale interactions include in situ development from monsoon trough convection (Zehr 1992), development of MCS (e.g., Ritchie and Holland 1999), monsoon trough breakdown (Ferreira and Schubert 1997) and interaction of a weak low-level disturbance with an upper low (e.g., Bosart and Bartlo 1991; Montgomery and Farrell 1993). The role of convection in both preconditioning the formation environment and in transforming the seed vortex to the familiar warm core tropical storm structure is considered next.

b) Convection feedbacks on genesis

Bister (2001) has documented the effects of formation latitude and the spatial distribution of convection on the early evolution of a tropical storm. Emanuel (1995a) observes that downdrafts forming in eyewall convection in a sub-saturated environment will cool and dry the local boundary layer, and thus suppress convection locally. He argues that high mid-tropospheric relative humidity is needed to suppress the negative effects of these downdrafts, and allow continued core convection in the developing storm. Should the eyewall convection be cut off, convection at large radius outside the radius of maximum winds – “peripheral” convection – may become established. Using the axisymmetric model of Bister and Emanuel (1997), Bister (2001) demonstrates that the presence of an outer ring of active convection spins up the tangential wind locally, and reduces the radial pressure gradient and so weakens the low-level inflow to the core. The secondary circulation of the storm now operates through this outer convection, which leads to subsidence in the core, weakens the eyewall convection and so suppresses mid-troposphere moisture buildup in the core. Further, this peripheral ring of convection modulates the location of subsequent convection – perhaps its greatest negative impact on eyewall convection (Bister 2001).

Corbosiero and Molinari (2002) performed an analysis of the distribution of lightning data to track active core (r < 100 km from the center) and outer band (100 km < r < 300 km) convection in 35 Atlantic storms within 400 km of the coast or over land. They stratify according to vertical wind shear [r  500km, 850-200 hPa] and find a slight preference for downshear left convection in the storm core and a strong signal for downshear right convection in the outer band. This is consistent with the development of deep divergent circulations in the evolving storms that act to oppose the vertical wind shear and minimize the tilt of the storm center. While these results applied across tropical depressions through to hurricanes, they were less consistent for weaker storms, which indicates that tropical depressions were less able to respond to vertical wind shear than stronger storm systems.

A number of authors (e.g., Kimball and Evans 2002; Frank and Ritchie 2001; Elsberry and Jeffries 1996) have demonstrated that strong vertical wind shear [>10-12.5 m s-1 over 850-200 hPa] acts to weaken a developing tropical system through suppression of its core convection. However, even a weak storm can act to organize itself against the shear, so that it may persist and even intensify in a region of strong vertical wind shear (Elsberry and Jeffries 1996; Simpson et al. 1998). Further, an adjacent upper trough or TUTT may act to reinforce the vertical motion and convection preferentially to oppose the action of the shear (Kimball and Evans 2002; Hanley et al. 2001).

The mechanisms responsible for tropical storm formation from mesoscale convective systems have only recently been elucidated. In this process, a deep warm cored cyclone must be created from a shallow, mid-level (lower tropospheric cold cored) vortex (e.g., Simpson et al. 1997, 1998; Bister and Emanuel 1997). To mitigate against the cooling effects of evaporation in saturated downdrafts, the free troposphere in the region of the developing eye should approach saturation (Montgomery and Farrell 1993; Emanuel 1995a,b, 1997; Simpson et al. 1998; Lighthill 1998). If the air remains sub-saturated these cool downdrafts will stabilize the boundary layer under the eyewall, so that convection will develop far from the storm center, and cut off the inflow of high _e air needed to feed the eyewall convection (Bister 2001). The moistening also reduces the local Rossby radius, such that the upper low has sufficient dynamical scale to drive growth of the low-level cyclone. This need for moistening in the region of the incipient mid-level vortex provides one clue to the large-scale condition of low vertical wind shear observed by Gray and others: minimal vortex-relative flow will allow for local moistening to be most efficient, promote more rapid development of the incipient disturbance and minimize the chance that negative environmental influences will destroy the system (Bister and Emanuel 1997; Simpson et al. 1998). The mid-level vortex must build down toward the surface to generate the surface inflow and increased surface fluxes needed for subsequent development (Emanuel 1995b, 1997; Fraedrich and McBride 1989). Bister and Emanuel (1997) propose that this will occur through buoyantly-driven vortex-relative flow advecting high PV air downward along isentropic surfaces. Once the surface winds respond to the downward building vortex, surface fluxes can increase the _e of air inflowing towards the nearly saturated core of the developing system, and provide source air for the “hot towers” of the eyewall (Simpson et al. 1998). The system can now develop independent of its environment and genesis may be said to have occurred.

4.3.5: Tropical Cyclogenesis and Climate Change

Gray (1968, 1979) identified six large-scale features of the tropical environment necessary conditions for tropical cyclogenesis (see Section 4.3.1). Through scaled combination of these environmental variables, he developed a “seasonal genesis parameter” [SGP] that related the magnitudes of these parameters to the frequency of tropical cyclogenesis and demonstrated that the SGP reproduced the seasonal and regional variability of tropical cyclone formation globally. Ryan et al. (1992) used the SGP to diagnose the realism of the control simulation of a global climate model [GCM] and of a climate change simulation in which carbon dioxide levels were enhanced compared to present day conditions. While there were some deficiencies in the control simulation, the regions of tropical cyclogenesis were consistent with observations and the peak activity in the western North Pacific was clearly implied. Application of this SGP diagnostic to the climate change simulation resulted in the suggestion of a doubling of tropical cyclone numbers in this warmer climate regime. However, this frequency tendency was dominated by changes in the thermal parameters of the SGP, and specifically the SST excess over 26°C. McBride and Zehr (1981) identified the thermodynamic parameters as indicators of the seasonality of genesis, rather than of frequency. Ryan et al. (1992) note that, if the thermal parameters are used as a seasonal indicator, and the dynamic components are used to indicate genesis frequency, no increase in tropical storm numbers is implied. Royer et al. (1998) acknowledge the difficulties with using the thermal parameters as direct indicators of genesis frequency. They apply only the dynamic components of the SGP in analyzing their climate change simulation, and restrict their analysis to zones of active tropical convection. Watterson et al. (1995) apply the SGP to ECMWF reanalysis fields to study the ability of the SGP to reproduce observed interannual variability. While the SGP once again reproduced the long-term mean patterns of tropical cyclogenesis, the SGP analyses of interannual variability were problematic. Based on recent theories of tropical cyclogenesis (Sections 4.3.2 and 4.3.3), it is likely that the inability of the SGP to capture mesoscale influences on genesis limits its applicability on finer timescales. Since the SGP was derived as a tool for isolating the mean tropical cyclogenesis environment in the current climate, such restrictions are not surprising. Coupled atmosphere-ocean GCMs [AOGCMs] are now beginning to resolve the larger scales relevant to mesoscale storm evolution (e.g., Meehl et al. 2000a). For example, Bengtsson et al. (1996) investigate changes in tropical storm frequency in the ECHAM3 GCM and conclude that tropical storm numbers are likely to decrease in a warmer climate.

The agreement between the dynamical SGP fields diagnosed from GCM simulations and the observed global distribution of tropical storm activity suggests that GCM simulate the large-scale environment for tropical cyclogenesis reasonably well, at least in the mean. These skillful control simulations are necessary to give confidence in GCM climate change scenarios (e.g., Easterling et al. 2000, Meehl et al. 2000b). Chan and Evans (2002) explore the realism of the large-scale environment for tropical cyclogenesis in the western North Pacific. They characterize the structure of the monsoon trough system in the western north Pacific and East Asia in terms of an oceanic monsoon trough, with a confluence point and an easterly ITCZ regime (Briegel and Frank 1997). Chan and Evans (2002) explore the interannual variability for ensembles of AMIP
2 simulations using the NCAR Community Climate Model (CCM3.6) and the NASA-GISS (version SI2000) GCMs during the boreal summers of 1989-1993. They compare the results of these GCM simulations with ECMWF Reanalyses and rainfall data. While the basic structure of the ITCZ system is present in each of the eleven ensemble members studied, systematic errors in location existed for each model. Further, the member-to-member variability for each ensemble had the same magnitude as the interannual variability from neutral to warm event. Thus, evaluation of simulations of even the summertime mean [June, July, August] monsoon trough regime in the western North Pacific reveals that improvements are needed to reproduce key characteristics of the monsoon trough relevant to tropical cyclogenesis.

In an effort to evaluate the likely effect of changes in the thermodynamic background environment on tropical cyclogenesis, Dutton et al. (2000) explored the variation of deep tropical convection with SST for a 134-year transient
3, coupled AOGCM simulation using the NCAR CSM1 model. Comparison of OLR and SST in the tropics in the control climate with the results of Graham and Barnett (1987) reveals that CSM1 captures the observed increase in cloud top height [decrease of minimum OLR] above a threshold SST of approximately 25°C [compared to 27°C in Graham and Barnett 1987]. As CO2 concentrations increase, the simulated threshold temperature for tropical convection progressively increases to approximately 26°C and 26.7° C at 2_CO2 and 3.4_CO2. This coupled GCM response to increasing CO2 concentrations implies that the expansion of the 26°C isotherm will not yield an expansion of the regions of deep tropical convection. Indeed, the convective rainfall spatial pattern does not differ systematically between years in this transient simulation. Further, while the average precipitation rate [40°N-40°S] increases slightly as the transient simulation progresses, the areal coverage of regions with extreme precipitation rates remains approximately constant. There is some suggestion that the spatial distribution of convection is more El Niño like (see also Krishnamurti et al. 1998).

Goldenberg et al. (2001) demonstrate the importance of variation in both SST and vertical wind shear on tropical storm frequency. As argued by Emanuel (1997, 1999) and Holland (1997), the increased SST could potentially lead to a higher upper limit on both potential and realized storm intensity. This increase in maximum potential intensity was demonstrated by Knutson et al. (1998) for a series of incipient storms placed in environments representative of the control and 2_CO2 climate regimes. Finally, the statistical scaling of Emanuel (2000) and the observational study of Evans (1993) both demonstrate that the vast majority of tropical storms never attain their peak potential intensity.

4.3.6: Research and Forecast Challenges

Anecdotal evidence for skill in operational model simulations of tropical cyclogenesis is beginning to emerge. However, quantitative measures of this genesis skill are still required to evaluate model reliability. Advances in theoretical understanding and observational analysis of tropical cyclogenesis suggest new diagnostics for genesis potential applicable to analysis of the operational models. Translation of these theories into models applicable to the forecast process is a challenge straddling both the research and forecast communities.

4.3.7: Summary and Recommendations

In the last decade, a much broader range of systems have been recognized as potential candidates for the incipient vortex leading to tropical depression and the mechanisms for genesis have been further elucidated. Environmental impacts on tropical storm development have been explored more comprehensively, including positive and negative influences of the dynamical structure of the environment in both moist and dry approximations. Internal system dynamics relevant to potential vorticity mixing and convective distribution/feedbacks on storm evolution have been better understood. Through extension of fundamental understanding of the necessary conditions for tropical cyclogenesis to analyses of GCM simulations, inferences on the effect of climate change on tropical storm frequency can now be drawn.

4.3.8: Bibliography


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1 Ritchie and Holland (1999) do not see any effect of the TUTT observed in their composite on the tropical storm formation; they speculate that this is due to the coarse resolution of their gridded analyses. However, Ferreira and Schubert (1999) propose that tropical cyclones may be responsible for the formation of TUTT cells and that this formation is preconditioned by the large-scale environmental shear. Thus, the TUTT could be a consequence, rather than a driver, of the tropical cyclogenesis.

2 Atmospheric Modeling Intercomparison Project

3 CO2 concentrations increase 1 percent per year throughout the simulation