economic loss after Hurricane Andrew in Florida and Louisiana in August 1992 was in the order of. U.S.$ 30 billion. In 1970 more than 200,000 persons.
Introduction
Tropical Cyclones Andreas H. Fink, Peter Speth Institute for Geophysics and Meteorology, University of Cologne, Kerpener Strasse 13, D-50923 Cologne, Germany
Correspondence to: A.H. Fink
482
Tropical cyclones are by far the most devastating of all natural disasters, in terms of both loss of human life and economic damage. It is estimated that the economic loss after Hurricane Andrew in Florida and Louisiana in August 1992 was in the order of U.S.$ 30 billion. In 1970 more than 200,000 persons were killed in Bangladesh after a storm surge caused by the Bhola cyclone, named after a small town located in the delta of the river Ganges. From the scientific point of view, very few atmospheric processes are as poorly understood as tropical cyclogenesis, i.e., the complex and interrelated dynamic and thermodynamic processes that lead to the formation and intensification of a tropical cyclone. This lack of knowledge is partly related to the fact that tropical cyclones spend most of their life time over remote tropical ocean areas where in situ observations are scarce or even lacking. Nonetheless, measurements from aircraft reconnaissance and advances in computer, radar, satellite, and other remote sensing technology have lead to a constant increase in our knowledge, especially with respect to the structure of an archetypal tropical cyclone and the environmental conditions that favor the formation of tropical storms. This contribution summarizes selected aspects of tropical cyclones. After introducing some basic definitions, we present a number of simple energetic arguments regarding that which controls the upper bound on intensity of tropical storms. We then describe briefly the basic circulation and geometry of a storm. The following section reviews the necessary preconditions and the initial disturbances that are able to initiate tropical cyclones, followed by a climatology of tropical cyclone frequency and occurrence for the tropical oceans. Finally, we address the important issue of understanding and detecting the Naturwissenschaften 85, 482–493 (1998)
Q Springer-Verlag 1998
long-term variability and trends in tropical cyclone activity. Here we present the current state of knowledge with respect to changes in tropical storm frequency, intensity, and geographical distribution in a greenhouse-warmed world.
Basic Definitions “Tropical cyclone” is the generic term for a lowpressure system over tropical or subtropical waters with organized convection and a definite cyclonic surface wind circulation (counter-clockwise in the Northern Hemisphere and clockwise in the Southern Hemisphere). In contrast to middle-latitude lowpressure systems, which derive their energy from strong horizontal temperature gradients associated with cold or warm fronts, tropical cyclones are driven by the latent heat released by the condensation of immense amounts of water vapor within their convective rainbands. For reasons discussed below, tropical cyclones are characterized by a large temperature excess in higher levels compared to the surrounding atmosphere. Therefore they are called “warm-core” systems. In contrast, middle-latitude lows are “cold-core” features. Another striking difference pertains to the level of maximum winds; in tropical storms the winds are strongest from the surface up to 2–3 km (usually referred to as the lower troposphere), while extratropical cyclones have a distinct wind maximum at about a height of 9 km (usually called the upper troposphere). Tropical cyclones are classified according to their maximum sustained wind speeds at the surface. Unfortunately, maximum sustained winds are defined as 1-min averages in the Atlantic and East Pacific and 10-min averages elsewhere, which complicates comparison of the statistics of one basin with those of another. Tropical lows with surface winds between 10 and 17 m/s are called “tropical depressions.” Once the maximum sustained winds exceed 17 m/s, the system is assigned a namein all basins except the northern Indian Ocean, and the system is called a “tropical storm.” If the wind speed further increases beyond 33 m/s, the cyclone is called a “hurricane” in the Atlantic and East Pacific and a “typhoon” in the western North Pacific west of 1807E. Consequently, the terms “hurricane” and “typhoon” are regionally specific names for strong tropical cyclones. These systems are called (severe) tropical cyclones or storms in the remaining tropical ocean basins.
Simple Energetic Considerations As defined above, tropical cyclones are warm-core, circular vortices, maintained by the transfer of latent and, much less important, sensible heat from the ocean to the atmosphere. The energy supply from the sea is immense, although limited by the physics of sea-air heat exchange. The latent heat is converted to sensible heat in the deep convective clouds, and this sensible heat is continuously converted into mechanical energy, i.e., the kinetic energy of the circulating air masses. In a steady-state tropical storm there is a balance between the supply of heat energy from the ocean and the loss of energy through surface frictional forces and radiative cooling. As a consequence, a tropical cyclone can be considered as a Carnot heat engine that converts heat into work [1]. The “fuel” of that heat engine is the atmospheric water vapor. A famous theorem in thermodynamics states that the maximum heat energy that can be converted into mechanical energy within one cycle of a Carnot engine, i.e., its thermodynamic efficiency, h, is given by:
hp
TinPTout Tin
(1)
where Tin is the temperature at which heat is added to the engine, and Tout is the temperature at which it is removed. For a tropical cyclone, heat is added at the sea surface and removed through outgoing longwave radiation at the level where the air flows out of the top of the storm. Tin and Tout are therefore the sea surface temperature (typically at about 287 C) and the temperature at the outflow level (around –757 C); therefore h is roughly 1/3. How much energy can there be in a hurricane? This depends on how much sea water can be evaporated, which in turn depends on the wind speed and the saturation deficit of an air parcel just above the sea. Neglecting radiational cooling, an expression for the maximum heat input is: Emax p Lq* (1PH)
(2)
where L is the latent heat of evaporation, and H is the relative humidity, which varies between 0 and 1, and q* is the saturation specific humidity which increases exponentially with temperature. Multiplying Eq. 2 by Eq. 1 expresses the maximum mechanical energy available to the engine (Eph Emax), that is: Ep
TinPTout Lq* (1PH) Tin
(3)
Let us now neglect radiation and assume a steadystate hurricane in which latent heat gain is balanced 483
by frictional loss. Turbulent friction occurs mainly in the very turbulent layer adjacent to the sea. The rate at which latent heat flows into the engine is: CwVE
(4)
where Cw is the heat transfer coefficient for water vapor and V is a representative surface wind speed. On the other hand, the rate of dissipation of mechanical energy is given by: CdV 3
(5)
where Cd is the drag coefficient. Setting Eq. 4 equal to Eq. 5 provides an expression for the maximum wind speed: Vmax p
" CC
w
E
(6)
d
For the lack of better knowledge, it is widely assumed in climate models that the exchange coefficients of heat, water vapor, and momentum are equal in magnitude; a typical value of the ratio Cw/ Cd is therefore 1. Given a sea-surface temperature of 297 C, a cloud-top temperature of –757 C, and a typical surface relative humidity of 80%, the maximum wind speed according to Eq. 6 would be roughly 66 m/s. Indeed, tropical cyclones with surface winds of this magnitude are almost exclusively observed over such high ocean temperatures. It should be pointed out that the maximum wind speed obtained from Eq. 6 depends critically on the value of the relative humidity, H, and that a tropical cyclone rarely attains its maximum potential intensity estimated from such simple thermodynamic models. This is due mainly to counteracting dynamic processes, such as the vertical variations in the magnitude of the wind and the upwelling of colder ocean water to the surface by the strong surface winds. Keep in mind that radiation was also neglected in Eq. 2 for Emax.
Structure of a Mature Tropical Cyclone The circulation within a mature tropical cyclone, that is, a storm which forms an eye, is rather complex. A detailed discussion is beyond the scope of this contribution. However, there are two major circulation types, an axisymmetric primary circulation and an asymmetric secondary circulation. The primary circulation consists of the tangential or swirling winds around the axis of rotation. This is strongest in the vortex core and above the frictional layer at heights of 2–3 km. In a steady-state balanced situ484
ation (in the absence of friction), the so-called gradient wind balance, the pressure gradient force acting to accelerate the rotating parcel inward, equals the combined Coriolis and centrifugal forces, trying to accelerate the parcel into the opposite direction. In such a hypothetical steady state the air would rotate continuously around the eye of the cyclone. However, given a constant pressure gradient force, the surface friction slows air parcels in the frictional layer, which causes a decrease in the Coriolis and centrifugal forces and thus a deflection of the parcel toward the low-pressure center. Eventually it spirals into the core of the cyclone. Measurements have verified that the gradient wind balance is a good approximation of the swirling winds in most parts of the vortex. As a consequence, many features associated with the swirling winds can persist for tens of rotation periods, each lasting about 1 h close to the center [2]. The surface wind component directed inward constitutes the lower branch of the asymmetric secondary circulation that maintains the primary circulation against friction and radiative cooling by supplying thermal energy and angular momentum. At the surface, the secondary circulation causes environmental air to converge in the eyewall region, that is, the region beneath the circular convective rainband which surrounds the eye of the cyclone (see Fig. 1a). If a surface air parcel undergoes a 60 hPa pressure drop on its way into the eyewall region, it should cool by about 57 C through adiabatic expansion. However, in reality it maintains a nearly constant temperature by heat gained from the ocean. In addition, it acquires water vapor from the ocean, such that the relative humidity in the parcel increases from about 75% to around 85%. As a consequence, the parcel reaches the eyewall much more buoyant than the environmental air surrounding the cyclone. If it is displaced vertically above a certain level, called the level of free convection (LFC), the parcel ascends freely to heights of up to 14 km or more (see Fig. 1a), simply because it remains warmer than the surrounding air due to the persistent latent heat release. During its ascent the total energy content (i.e., kinetic plus potential plus latent plus sensible) heat content of the parcel is approximately conserved, although in the process there are large conversions from latent to sensible heat. The vertical wind speed in the eyewall is on the order of 5 m/s, much less than the 20 m/s which occurs in thunderstorm complexes over the midwestern United States. The huge overall energy release in a tropical cyclone is the result of the size of the area covered by deep organized convection. At upper levels (12–15 km, see
Fig. 1. a) Schematic presentation of the diabatically driven secondary circulation of a mature tropical cyclone. The eyewall convection is separated into an inner and outer eyewall, which is clearly discernible from the radar reflectivity pattern in b). Note the spiraling “feeder rain bands” and the embedded convective cells (black areas). (From [14])
Fig. 1a) the air flows anticyclonically out of the vortex due to a high-pressure region aloft and the stable stratification which inhibits further ascent. The pressure gradient force accelerates an air parcel radially outward. Under the action of the Coriolis force it curves more and more anticyclonically. The described diabatically driven secondary circulation is strongest near the surface and predominates over the primary circulation away from the vortex core. The two circulations add to the total flow in such a way that the maximum low-level surface winds are observed below the inner eyewall. As mentioned above, the eyewall is a concentric ring of strong convection surrounding the eye of the tropical cyclone. The eye is a region of calm winds at the axis of rotation, no precipitation, and sometimes even clear sky. The eye has a typical diameter of 30–60 km and appears when the storm approaches
hurricane or typhoon intensity. Within the eye, air is forced to descend to levels 1–3 km above the surface (Fig. 1a). Adiabatic compression of the sinking air is the cause of the great temperature excess that typically maximizes to values up to 10–157 C just above the moist 2- to 3-km-deep surface layer. The question now arises: What mechanisms give rise to the formation of the eye? Surprisingly, the dynamics involved are not yet understood in detail. While Gray [3] ascribes the eye to a dynamically forced centrifuging of mass out of the eye into the eyewall due to an unbalanced flow, Willoughby [2] attributes the formation of the eye to the moist convection in the eyewall. Sustaining the eyewall convection requires a strong and permanent moisture supply. However, if the storm achieves hurricane force winds, moist air parcels advected from the environment can no longer effectively reach the center since the air blows more and more tangentially along lines of constant pressure. Hence, the maximum of updrafts and rainfall must move out of the center of the vortex where the ascending motion is replaced by convectively forced sinking motion. Finally, more mass is removed from the eye via the low-level outflow into the evolving convective ring than is replaced by the upper-level inflow that originates from the top of the eyewall convection (Fig. 1a). This causes the pressure to drop further within the eye. It should be pointed out here that in a recent publication Emanuel [4] gives some strong arguments that convection in the eyewall cannot make the eye warmer than the eyewall. From theoretical and modeling considerations, he infers that the mechanical spin-up of the eye’s rotating winds that occurs through lateral stresses at the eye-eyewall boundary requires additional forced subsidence warming at the center of the eye to satisfy the thermal wind balance. (It is assumed that strong radial turbulent momentum diffusion occurs when the radial gradient of the wind speed at the eye-eyewall boundary exceeds a certain value. As a consequence the tangential wind speed increases at the rim of the eye.) From the arguments above it is clear that fluctuations in the eyewall convection should give rise to corresponding pressure and surface wind fluctuations. Such an oscillation of the central pressure and maximum winds is in fact observed in intense tropical cyclones by a sequence of contracting convective rings. From the radar reflectivity pattern in Fig. 1b it can be seen that much of the precipitation within the cyclone occurs in the spiraling rain bands, often called the “feeder bands.” They eventually lead to the formation of the concentric outer eyewall, which can also be seen in Fig. 1b. As this outer ring ap485
proaches the inner eyewall, its associated subsidence (Fig. 1a) suppresses the convection in the inner eyewall. This causes a rapid increase in the central pressure of the storm. In contrast, as the outer eyewall replaces the inner eyewall, the cyclone rapidly intensifies. The described strong pressure oscillations, which have a variable period between 1 h and 1 day, are a special problem in forecasting the landfall intensity of tropical cyclones.
Tropical Cyclone Formation As discussed above, the development and maintenance of tropical cyclones depends critically on the moisture supply from the surface. From observations it is known that warm ocean waters of at least 26.57 C throughout a sufficient depth of about 60 m are a necessary condition for tropical cyclogenesis. Due to the enormous heat capacity of water, evaporational cooling of the ocean surface is almost imperceptible. However, strong vertical mixing induced by the strong surface winds would considerably lower the temperature in the uppermost layer of the ocean unless the mixing occurs entirely within a warm water body of more than 26.57 C. Conversely, the low heat capacity and limited moisture availability of a land surface prevents tropical cyclone formation over land and causes a landfalling cyclone to dissipate rapidly. It should be pointed out here that the lack of energy supply, but not the increased surface friction, is the primary process that kills tropical cyclones over land. Another commonly known prerequisite of tropical cyclone formation is the existence of a non-negligible Coriolis force. Near the equator the Coriolis force is negligible, and air parcels can reach the center of a low-pressure system in a straight-line motion that is perpendicular to the lines of constant pressure. To force a parcel on a spiraling motion into a nearequatorial low-pressure system would require a preexisting curvature in the track of the air parcel. In this case the centrifugal acceleration would counteract the pressure gradient force. However, in the real world, pressure perturbations due to convection remain very small near the equator, and almost all cyclones form poleward of 57 latitude. North and south of this latitude the deflection of air parcels to the right in the Northern Hemisphere, and to the left in the Southern Hemisphere, is sufficient to create a rotating vortex in which a low-level air parcel must travel along a spiraling trajectory before it reaches the low-pressure center. Next, a certain level of thunderstorm activity is needed that allows the heat stored in the ocean to 486
be liberated for tropical cyclone development. These conditions are closely bound to the location of the Intertropical Convergence Zone (ITCZ). Deep convection occurs repeatedly within the lowpressure zone of the ITCZ due to the low-level convergence of moisture-laden winds underneath a conditionally unstable atmosphere. The ITCZ migrates north and south, more or less following the sun’s solstice, which causes a distinct seasonality in tropical cyclone occurrences in all basins. The fact that the sea surface temperature falls below or exceeds the threshold value of 26.57 C in the course of the year is secondary. Finally, moist middle-tropospheric levels reduce the evaporation of thunderclouds that occurs when dry ambient air is entrained into the cloud; they are hence a further factor facilitating tropical cyclone formation within the ITCZ. In addition to the three fundamental preconditions discussed above, a fourth factor is of capital importance; low vertical wind shear, i.e., a slow change in the magnitude of the wind with height. Weak shear in the vertical (less than about 10 m/s difference in the surface and upper-level winds at about 12.5 km) allows the heat released by condensation to concentrate in a vertical column. High wind shear not only effectively inhibits tropical cyclogenesis but also destroys a mature tropical cyclone by interfering with the organization of deep convection around the cyclone center. Even if all four environmental factors favor cyclogenesis, tropical cyclones do not form spontaneously. A pre-existing, near-surface disturbance with a sizeable spin and weakly organized thunderstorm activity due to convergent surface winds is required. In the Tropics two types of disturbances are wellknown: (a) the so-called “easterly waves,” which occur in the convergence zone of the easterly trades, and (b) disturbances associated with a “monsoon trough.” Similar to the ITCZ, the atmospheric conditions prevailing in the monsoon trough (the term “trough” denotes a low-pressure area) are conducive to thunderstorm activity. The monsoon trough is associated with the monsoon shear line that separates converging low-level equatorial westerlies and the easterly trades poleward of the monsoon shear line. If a north-south oriented paddle were thrown into this streamflow in the Northern (Southern) Hemisphere, it would start to rotate counterclockwise (clockwise) around an axis centered at the monsoonal shear line. This weak cyclonic shear and the location of the monsoon troughs well poleward of 57 latitude are among the environmental factors that favor tropical cyclone formation. In fact, about 80% of the tropical cyclones form from disturbances within the monsoon trough and/or the ITCZ [5].
The only major exception is the North Atlantic basin. Due to the strong north-south temperature gradient between the hot Sahara Desert and the cooler Congo Basin and Gulf of Guinea region, a strong low-level easterly jet develops over Central and West Africa between May and October. Instabilities in the jet sometimes grow to African easterly waves which frequently travel across the Atlantic without further development. However, some of them organize into Atlantic hurricanes, especially in July to September. In fact, tropical waves in the Atlantic trades are not only the most important tropical cyclogenesis mechanism in the North Atlantic basin but are also the triggering mechanism of the most intense Atlantic hurricanes. For example, more than 90% of intense hurricanes that made landfall along the eastern coast of the United States since the advent of satellite monitoring in 1967 originated from African easterly waves [6]. It has even been observed that African easterly waves travel across the Atlantic and the Caribbean Sea and develop into hurricanes over the East Pacific after crossing Central America. Globally about 15% of the tropical cyclones originate from disturbances in the easterly trade wind flow. It should be mentioned that tropical cyclones sometimes form at old, stagnant, subtropical frontal systems. This is an important cyclogenetic mechanism for both early-season (June–July) and late-season (October–November) Atlantic hurricanes, especially over the warm waters of the Gulf of Mexico and the Gulf Stream off the eastern coast of the United States. Additionally, upper-level cold vortices which sometimes invade the Tropics from middle latitudes can develop into tropical cyclones. In both cases a middle-latitude type low-pressure area with an upper-level cold core and considerable wind shear transforms into a true warm core tropical system due to the latent heat released by the thunderstorm activity in its center. The warm Gulf Stream and the Kuroshio Current east of Japan are regions in which such “semitropical” systems can form at a surprisingly high latitude of up to 307N [5].
Although the above preconditions for tropical cyclone formation prevail over vast tropical ocean areas for weeks or even months, an annual average of only about 87 tropical disturbances intensify to become named tropical storm strength (Table 1). Since the cyclones form over data-free or data-sparse oceanic areas, and since reconnaissance aircraft are not launched before a disturbance reaches a certain degree of organization, the factors why some disturbances intensify to tropical storm strength while others decay are not well understood. It is generally accepted that a certain amount of pre-existing uppertropospheric outflow, i.e., divergent winds, is conducive for intensification. Tropical cyclone forecasters therefore closely check whether upper-tropospheric winds are favorable. The problem is that small errors in the analyzed total wind field can cause large errors in the divergence of the wind field. As mentioned above, another factor favoring tropical cyclone formation and intensification is the existence of moist middle-tropospheric levels. A pre-existing mesoscale (100–200 km) convective system may have left such a moist layer at middle to upper tropospheric levels, but a subsequent blow-up of convection must occur underneath the pre-moistened upper troposphere. Such upper-level moisture anomalies are also difficult to analyze over the data-poor tropical oceans. Taking these difficulties into consideration, it can easily be understood that predicting tropical cyclone formation is a very complicated task. Cyclone forecasters therefore rank the probability of tropical cyclone formation in a region of persistent thunderstorm activity with a discernible low-level cyclonic circulation as “poor,” “fair,” or “good.” However, sometimes a tropical cyclone formation alert must be canceled, or a cloud cluster formation ranked as “poor” develops into a tropical cyclone within a day. An extreme example of an explosive intensification is Super Typhoon Forrest in September 1983; Forrest’s central pressure fell from 976 to 876 hPa in just under 24 h.
Table 1. Averaged annual total numbers of named tropical cyclones (i.e., peak surface winds of at least 17 m/s) and their standard deviations (std) for all tropical basins and for the 30-year period 1968 through to 1997. In the rightmost column the global average over all basins is displayed
Mean Std
Western North Pacific, South China Sea
East Pacific (~1807W)
Atlantic Caribbean Sea, Gulf of Mexico
North Indian Ocean
Australian Region, Southwest Pacific ( 1 1007E)
Southwest Indian Ocean (~1007E)
global
26.9 4.3
16.6 4.6
9.3 3.6
5.5 2.2
16.2 3.9
12.1 3.2
86.7 7.9
487
Observed Tropical Cyclone Occurrence Figure 2 plots the tracks of all tropical cyclones which in 1988–1990 reached at least named tropical cyclone strength (i.e., winds exceeding 17 m/s). For a better graphical representation, the Tropics are divided into the Eastern (Fig. 2a) and Western Hemispheres (Fig. 2b). It can be seen from Fig. 2 that about twice as many tropical cyclones occur in the Eastern as in the Western Hemisphere. Moreover, two-thirds are observed in the Northern Hemisphere, due mainly to the absence of cyclones over the Southeast Pacific and South Atlantic. As a consequence of the intense cold ocean currents at the eastern rims of both basins in the Southern Hemisphere the most southerly positions of the respective ITCZs are located at the equator or even in the Northern Hemisphere. Over the South Atlantic, strong vertical wind shear also unfavors tropical cyclone formation. Note that the average number of named tropical cyclones given below for each individual basin refer to the 30-year base period 1968–1997, and are reproduced in Table 1. Figure 2 also demonstrates that the most active cyclone basin is the western North Pacific including the South China Sea. During the period 1968–1997 almost 27 disturbances reached named tropical storm intensity in an average year. Roughly two-thirds of these became typhoons. Although the typhoon season peaks in August, the western North Pacific is the only region where tropical cyclogenesis has been observed in all months of the year. This basin is also known for its very large and intense storms. About four typhoons each year reach Super Typhoon intensity (i.e., winds exceeding 67 m/s), and Super Typhoon Tip was measured to have a central pressure of 870 hPa on 12 October 1979. Factors that promote the frequent genesis of typhoons include the very large and deep warm water body with temperatures of 29–307 C in the northern summer, an active monsoon trough, and persistent upper-level divergence. Typhoons in this region frequently make landfall at the coasts of Vietnam, mainland China, Taiwan, Korea, and Japan and cause much damage and many fatalities each year. With nine tropical system per year (average for 1968–1997), the Philippines is the country that is most frequently affected by tropical storms in the Tropics. The second most active basin is the eastern North Pacific, with an average of 16.6 named cyclones during the hurricane season from the middle of May to the end of November. Most of the cyclones dissipate over water north of 207N and between 120 and 1407W (Fig. 2b) because they encounter cold waters and high wind shear. However, an average of one 488
Fig. 2. The tracks of all named tropical cyclones (i.e., maximum sustained winds exceeding 17 m/s) for the 3-year period 1988–1990
storm per year propagates as far as the central Pacific. In August–September 1994 Hurricane John traveled all the way from the eastern to the western Pacific, and with a life span of 31 days John is the longest lasting tropical cyclone on record. Especially during the late hurricane season eastern Pacific hurricanes can recurve and make landfall at the Mexican Coast including Baja California. This happened in October 1997, as Hurricane Pauline severely damaged the town of Acapulco. Since the 1997 hurricane season also saw Hurricane Linda, the strongest East Pacific cyclone on record, with sustained winds of 160 knots and an estimated minimum pressure of
900 hPa, the simultaneous 1997–1998 El Nin˜o event has widely been held responsible for these weather extremes. However, as discussed below, a clearcut effect of El Nin˜o on the tropical cyclone activity in that basin is not supported by current scientific knowledge. In the Southern Hemisphere an annual average of about 16 named cyclones occur in the Australian/ Southwest Pacific Ocean region and about 12 in the Southwest Indian Ocean west of 1007E. In both of these areas the initial disturbances are commonly located in the monsoon troughs. In the Southwest Indian Ocean 85% of cyclones occur between December and March, with a peak in January. Interestingly, the Southwest Indian Ocean is the only basin other than the western North Pacific where tropical cyclones occur in winter. Cyclones form rather regularly in July and August in the Southern Hemisphere central Indian Ocean. In the Australian/ Southwest Pacific region the tropical cyclone season extends from late October through May. In the Southwest Pacific region the frequency distribution has a single peak in February, while a double peak occurs in the tropical seas north of Australia in January and March, with a distinct minimum in February. The reason for the double peak is simple: the ITCZ is located mostly north of Australia in January and March, while it moves to its southernmost position in February and is then located over the continent of Australia. As a consequence, the potential for tropical cyclone development is reduced over the adjacent tropical seas. In the North Atlantic, including the Caribbean Sea and the Gulf of Mexico, an average of nine named cyclones are observed between June and November. In an average year five reach hurricane strength. As discussed above, the most intense hurricanes form from easterly waves, south of 207N and just east of the Cape Verde Islands, mainly during the peak hurricane season in August and September. The strongest Atlantic hurricane on record was Hurricane Gilbert in September 1988, with an estimated central pressure of 888 hPa. The least active cyclone basin is the North Indian Ocean. Only five or six cyclones form each year in the Bay of Bengal and the Arabian Sea, and only two of these reach hurricane force winds. However, five to six times as many tropical cyclones form over the Bay of Bengal than in the Arabian Sea [7]. The cyclone frequency has a bimodal distribution, with peaks in April–May and October–November. Between June and September strong vertical shear inhibits tropical cyclone development over this area. Although only 7% of the global tropical cyclones form in that basin, they are the most disastrous in
the world. The shallow waters of the Bay of Bengal, the low flat coastal terrain and the funneling shape of the coastline cause severe storm surges that flood the delta of the Ganges River. In 1970 more than 200,000 persons were killed in Bangladesh, and in 1991 the death toll was more than 100,000. In some years, such as 1988 and 1997, tropical storms in the South China Sea cross the Malay Peninsula into the Bay of Bengal.
Year-to-Year Variation in Tropical Cyclone Activity Interestingly, the global average of 87 tropical cyclones per year is rather stable for a weather system that is regarded as a relatively rare phenomenon in any one basin. Its annual average variation is only 9% (Table 1 and Fig. 3d), with the extreme variations being –17% and c18%. In contrast, interannual variations in tropical cyclone activity can be quite remarkable in a given basin. This statement especially holds for the North Atlantic, where the average annual variation is nearly 40% (Table 1). Thanks to the pioneering work by Prof. W. Gray from the Colorado State University and others, factors that control tropical storm activity in the North Atlantic are reasonably well understood. Since 1984 Professor Gray and his collaborators have issued Atlantic tropical cyclone seasonal forecasts as early as December of the preceding year. The forecasts are updated in early June at the beginning of the hurricane season, and in early August at the start of the most active period (for the interested reader, Professor Gray’s forecasts are available in the Internet at http://tropical.atmos.colostate.edu/forecasts/ index.html). Gray’s forecasts are based on a multivariate linear regression analysis [8]. Major factors that enter his formula and, that are known to affect Atlantic hurricane formation are the following. First is the state of the El Nin˜o–Southern Oscillation (ENSO) in the Pacific Ocean; during ENSO warm (or El Nin˜o) years increased upper-level westerlies over the tropical Atlantic lead to an increased vertical wind shear. The prolonged 1991–1994 El Nin˜o resulted in the least active period on record, with a total of only 15 hurricanes. The substantial decrease in the number of named Atlantic tropical storms during El Nin˜o periods is evident from the filled bars in Fig. 3a, especially from the 1970s onward. Second is the phase of the Quasi-Biennial Oscillation (QBO). Stratospheric winds at and above the top of tropical thunder489
storms change direction from west to east roughly every 12–15 months. If stratospheric winds blow from the east, Atlantic hurricane activity is reduced, presumably due to increased upper-tropospheric to lower-stratospheric wind shear. A third factor is African West Sahel rainfall. A wet Sahelian rainy seasons seems to be associated with a higher chance of low-latitude “Cape Verde” type hurricanes, i.e., storms that form from African easterly waves. Gray’s predictions have demonstrated considerable accuracy. However, a major problem is the need for accurate long-lead forecasts of the most influential forcing mechanism, the state of the ENSO. For example, in 1997 Gray overestimated the Atlantic hurricane activity, partly because he forecast a weak El Nin˜o for 1997–1998, whereas it has in fact turned out to be one of the strongest on record. In contrast, the effect of El Nin˜o on the number of named East Pacific storms is not very clear (Fig. 3b). Nonetheless, it is generally believed that El Nin˜o exerts an effect on the basin-wide tropical cyclone activity. In the West Pacific cyclones originate farther east and closer to the equator in both hemispheres [9]. In the Southern Hemisphere fewer cyclones form over the Coral Sea and threaten the eastern Australian coast, while in the Northern Hemisphere more cyclones are observed to move from the East to the Central Pacific. Therefore the chance of the Hawaiian Islands being struck by a cyclone increases. In September 1992, an El Nin˜o-year, Hurricane Iniki caused substantial damage on the Hawaiian island of Kauai. In conclusion, some interannual variability is related to ENSO, but with the exception of the Atlantic less is known about additional causes of the year-to-year variations in other basins, and how they are interrelated.
Multidecadal Variability and Trends in Tropical Cyclone Activity Despite the very remarkable increase in the globally averaged surface temperature in the past two decades, which has also affected the land and oceanic areas in the Tropics, no increase or decrease in the
Fig. 3. Total numbers of named tropical cyclones for: a) the Atlantic basin (1944–1997), b) the eastern North Pacific (1949–1997), c) the western North Pacific (1945–1997), and d) the sum over all tropical basins (1968–1997). Filled bars, El Nin˜o-years; right abscissa, longterm average number 490
total number of named storms is discernible from Fig. 3d for the period 1968–1997. A reliable global number cannot be obtained for earlier years due to the lack of global data coverage in the pre-satellite era. Some authors have argued that the number of tropical cyclones in the world’s most active basin, the western North Pacific, has increased in recent years [10]. Figure 3c apparently confirms this; however, the 1960s saw a similar highly active period, and the increase may reflect only a decadal variation. The presence of island observing stations and busy commercial shipping routes and the advent of reconnaissance aircraft observations since the middle 1940s have resulted in rather reliable tropical cyclone records for the western North Pacific and the North Atlantic. In contrast, the increase in the corresponding statistics on the East Pacific is artificial (Fig. 3b). Due to its remoteness, the cyclone activity in this basin was long underestimated before the advent of satellite observations at the end of the 1960. In conclusion, tropical cyclone statistics do not support a significant trend in the number of named tropical cyclones, with the possible exception of the North Indian Ocean where data seem to support a downward trend over the past two decades (not shown). However, a distinct decadal variability (i.e., temporal variations on the order of 10–50 years) in the number and preferred tracks of intense Atlantic hurricanes is evident in the historical record. For example, in the more active 47-year period 1921–1967, 15 intense hurricanes with sustained wind speeds of more than 50 m/s made landfall in Florida, while there were 12 for the remaining eastern coast of the United States. In the less active periods, 1900–1920 and 1968–1993, together 47 years, the corresponding numbers were 4 and 5, respectively [6]. While the variability in the early part of this century can partly be explained by a shift in hurricane tracks from the Gulf coast in the 1900s toward the Florida and midAtlantic coast in the 1920s [11], the considerable decline in the number of intense landfalling hurricanes at these coastal strips in recent decades agrees with a similar trend in the Caribbean Sea and in the Gulf of Mexico. This is clearly visible from Fig. 4, which displays the tracks of intense hurricanes for the 24year periods 1944–1967 (upper panel), and 1968–1991 (lower panel). It should be pointed out here that the satellite detection of a few weaker storms in the remote areas of the Atlantic in recent years may have prevented the detection of a similar trend in the total number of named storms. Despite the fact that ENSO warm events have been more prominent in the past two decades, it cannot explain the above decadal varia-
tions. Finding the causes of this decadal variation has become an important scientific challenge, given the prospect of a return of more active hurricane seasons in the Caribbean Sea and at the Atlantic and Gulf coasts of the United States. These areas face an increasing risk, since the population and property development have immensely grown, while the hurricane preparedness and awareness of local inhabitants have declined. The question arises as to the potential causes of the decadal-scale Atlantic hurricane variability. Recent research, especially by Prof. Gray and his colleagues (see, for example [6]), suggests that the decadal changes in the ocean temperature in the tropical and subtropical Atlantic are responsible for the changes in the distribution and intensity of hurricanes in that region. A concurrent cooling (warming) of the North (South) Atlantic was observed in the 1960s which is thought to be the major cause of a weaker West African monsoon. The decadal West African monsoon signal is reflected in the well-known prolonged Sahelian drought which began in the early 1970s and still is continuing. The timing of the decadal changes in the Atlantic hurricane activity and the West African rainfall suggests that these climatic fluctuations are interrelated. Pasch et al. [12] give evidence that the total number of African easterly waves that leave West Africa, and that are known to be the most important incipient disturbances for tropical cyclogenesis in the Atlantic, is not related to the year-to-year changes in Atlantic hurricane activity. However, it appears that the changes in the Atlantic sea-surface temperatures and in the low-level monsoonal flow off the coast of West Africa, which are at least partly responsible for the decadal Sahelian drought, created less favorable conditions for cyclone formation in the East Atlantic. For the quiet season, 1994, Pasch et al. noted that cooler ocean temperatures, higher stability, and vertical wind shear inhibited a better organization of African wave disturbances over the eastern Atlantic; the waves became poorly organized, or even dissipated. Gray [6] has speculated that sea-surface temperature anomalies in the Atlantic will weaken and take on the opposite sign in coming years due to an increased heat advection in the Gulf Stream. This would increase the potential of strong landfalling hurricanes along the eastern United States coast. In the past three decades more than 90% of these systems originated from African easterly waves, a factor that would become stronger and more frequent in Gray’s scenario. However, it remains speculative whether the most active 3-year period on record, 1995–1997, and the concurrent warming of the 491
North Atlantic, signals a decadal upturn in the Atlantic hurricane activity.
Tropical Cyclones and Global Climate Change
Fig. 4. The tracks of all intense (i.e., maximum 1-min surface winds exceeding 50 m/s) Atlantic hurricanes for the 24-year periods: a) 1944–1967 and b) 1968–1991
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Coupled ocean-atmosphere general circulation models (OAGCMs) predict an increase in the oceanic area enclosed by the 26.57 C sea-surface temperature isotherm as a result of increasing greenhouse gas concentrations. It is a fallacy that the increase in tropical ocean temperatures automatically results in an increasing number of more intense tropical cyclones. The world’s leading scientists in the field of tropical cyclone research recently reviewed the current state of knowledge with respect to possible developments next century [9]. As noted above, tropical cyclogenesis depends on a variety of environmental factors, of which ocean temperature is merely one. While the spatial resolution of present OAGCMs is too low to simulate mature tropical cyclones, they can give a crude prediction regarding the environmental factors that favor tropical cyclogenesis. Many simulations show intensified tropical atmospheric east-west and north-south overturning circulation cells, known as the Walker and Hadley circulation cells, which is associated with an upper-tropospheric warming within the ITCZ as a consequence of increased latent heat release over warmed oceans. While the increased overturning in the Walker and Hadley cells causes stronger vertical wind shear, the increased dry static stability due greater warming of the upper than of the lower troposphere tends to counteract the decrease in moist static stability at lower levels. The latter is caused by the increased low-level water vapor content. Moreover, Holland [13] showed with a thermodynamic tropical cyclone model that the vertical temperature and moisture profiles predicted by an OAGCM causes the threshold ocean temperature for tropical cyclone formation to increase to about 287 C, thus leaving the area susceptible to tropical cyclogenesis at about the same size as presently. Moreover, the maximum potential intensity of tropical cyclones underwent an only modest increase of 10–20%. Summarizing all the above results, the current state of knowledge does not support any major changes in the area or the global location of tropical cyclone genesis in a greenhouse-warmed world. However, increased ocean temperatures at higher latitudes of the West Atlantic and West Pacific Oceans could, given a favorable low vertical wind
shear, sustain single tropical cyclones at unprecedented high latitudes and/or change the character of the extratropical transition of the tropical cyclone.
Outlook We have discussed several unresolved questions with respect to the dynamic and thermodynamic interplay that leads to a formation and intensification of a tropical cyclone. The prospects for future progress are somewhat mixed. On the one hand, increasing computer power will allow scientists to run more sophisticated hurricane models, for example, high-resolution coupled ocean-atmosphere models capable of simulating the large-scale environmental flow, and which can resolve a mature hurricane or even single clouds. This will undoubtedly contribute to a more thorough understanding of hurricane dynamics. On the other hand, model simulations must be validated against observations. At present, the extent of classical observational networks (surface observations and weather balloons) is shrinking due to budgetary reductions, and remote sensing data do not always provide a suitable substitute for in situ observations. In 1987 the very valuable reconnaissance aircraft flights into West Pacific typhoons were canceled by the United States government. Moreover, data collected in thousands of these flights since 1945 have been lost. This lead to a further: there are remarkably few scientists working on the problem of understanding and predicting tropical cyclones, compared to those studying other natural disasters of similar social significance, such as earthquakes [1]. It is indisputable that the potential for loss of life and destruction of property caused by tropical cyclones has dramatically increased in recent years as a result of coastal development and urbanization in cyclone-prone areas. Society must weigh the cost of maintaining an observational network and funding basic scientific research against the losses in human lives and economic damage.
1. Emanuel KA (1997) Climate variations and hurricane activity: some theoretical issues. In: Diaz FH, Pulwarty RS (eds) Hurricanes. Springer, Berlin Heidelberg New York 2. Willoughby HE (1996) Mature structure and evolution. In: Elsberry RL (ed) Global perspectives on tropical cyclones. World Meteorological Organization, Geneva 3. Gray WM (1991) Comments on “gradient balance in tropical cyclones”. J Atmos Sci 48 : 1201 4. Emanuel KA (1997) Some aspects of hurricane inner-core dynamics and energetics. J Atmos Sci 54 : 1014 5. Frank WM (1985) Tropical cyclone formation. In Elsberry RL (ed) A global view of tropical cyclones. Office of Naval Research, Arlington 6. Gray WM, Sheaffer JD, Landsea CW (1997) Climate trends associated with multidecadal variability of Atlantic hurricane activity. In: Diaz Diaz FH, Pulwarty RS (eds) Hurricanes. Springer, Berlin Heidelberg New York 7. McBride JL (1996) Tropical cyclone formation. In: Elsberry RL (ed) Global perspectives on tropical cyclones. World Meteorological Organization, Geneva 8. Landsea CW, Gray WM, Mielke PW, Berry KJ (1994) Seasonal forecasting of Atlantic hurricane activity. Weather 49 : 273 9. Henderson-Sellers A, Zhang H, Berz G, Emanuel K, Gray W, Landsea C, Holland G, Lighthill J, Shieh S-L, Webster P, McGuffie K (1998) Tropical cyclones and global climate change: A postIPCC assessment. Bull Am Meteor Soc 79 : 19 10. Chan JCL, Shi J (1996) Long-term trends and interannual variability in tropical cyclone activity over the western North Pacific. Geophys Res Lett 23 : 2765 11. Diaz FH, Pulwarty RS (1997) Decadal climate variability, Atlantic hurricanes, and societal impacts: an overview. In: Diaz FH, Pulwarty RS (eds) Hurricanes. Springer, Berlin Heidelberg New York 12. Willoughby HE (1988) “Atlantic tropical systems of 1994 and 1995: A comparison of a quiet season to a near-record-breaking one”. Austr Meteorol Mag 36 : 183 13. Pasch R, Avila JLA, Jiing J-G (1998) The maximum potential intensity of tropical cyclones. Mon Weather Rev 126 : 1106 14. Holland G (1997) The dynamics of the tropical cyclone core. J Atmos Sci 54 : 2519
Acknowledgements. The authors thank Ms. Peggy Allard for her assistance in preparing the figures. The authors also thank two anonymous reviewers for their suggestions.
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