**2. Life cycle of tropical cyclones**

This section describes the general understanding about life cycle of TCs. The description includes a brief overview about their genesis, structural evolution, propagation and dissipation.

#### **2.1. Cyclogenesis**

western pacific, south-western and south-eastern Indian Ocean and North Indian Ocean (NIO) region. The cyclonic storms are often known as hurricanes and typhoons in the Atlantic and northwest Pacific, whereas TCs in other Ocean basins. The average frequency of occurrence, season and intensity of TCs vary from basin to basin. NIO basin shows bi-modal TC season with maximum frequency during post-monsoon period (October–December) and are comparatively stronger than pre-monsoon ones. Though the size of the TCs is relatively smaller and their intensity is comparatively less over NIO basin as compared to the other global basins, this region is quite important in view of the densely populated rim countries with poor socioeconomic conditions. Hence, loss of life and property is quite significant in this region. Based on the intensity, TCs formed over NIO basin can be classified into (i) depression if the associated 10-m maximum sustained wind (MSW) is in between 17 and 33 kt; (ii) cyclonic storm (CS) if MSW is in between 34 and 47 kt; (iii) severe cyclonic storm (SCS) if it has MSW of 48–63 kt; (iv) very severe cyclonic storm (VSCS) if the MSW is within the range 64–90 kt, and extremely severe cyclonic storm (ESCS) if the MSW is in range of 91–119 kt; and (v) super cyclonic storm (SuCS) if it has MSW of 120 kt or more (http://www.rsmcnewdelhi.imd.gov.in). This classification may differ from those over other global basins including that of the widely used Saffir-Simpson hurricane wind scale or SSHWS (http:// www.nhc.noaa.gov/aboutsshws.php). Both types of classifications, that is, the earlier one from India Meteorological Department (IMD) and the SSHWS consider the tropical low-pressure system as a depression if MSW is <34 kt. The consideration of CS and SCS lies in the tropical storm category (MSW lies in the range 34–63 kt) of SSHWS. The VSCS category over NIO basin is similarto that of the category 1 hurricane (MSW lies in the range 64–82 kt) type. The category 3 (MSW lies in between 96 and 112 kt) and category 4 (MSW is within the range 113-136 kt) major hurricanes are comparable to ESCS over NIO basin. The IMD classification categorizes MSW above 120 kt as super cyclonic storm (SuCS), whereas SSHWS considers the desired wind speed above 137 kt for category 5 hurricane. However, the basic structure of NIO TCs

194 Recent Developments in Tropical Cyclone Dynamics, Prediction, and Detection

Having a prolonged coast line, about 96 districts (lying within 100 km from the coast) of India are vulnerable to the occurrence of TCs with varying intensity [2]. Out of these 96 districts, ∼59% are at least highly vulnerable. The number of CS and SCS with a core of MSW (between 34 and 63 kt) crossing different countries of the NIO region is found to be 504 during 1891– 2015 (derived from the available IMD data at http://www.rmcchennaieatlas.tn.nic.in). Out of these, about 328 (>65%) crossed the Indian coasts, whereas 127 (>25%) have crossed the east coast of India between Gopalpur and Kolkata. In general, the proneness to TCs is quite high for the coastal districts of West Bengal, Odisha, Andhra Pradesh, and Tamil Nadu [2, 3]. In view of these, TCs over NIO basin can be considered as quite lethal and expensive natural disaster as they bring widespread destruction in these regions. The consequent loss of human life and properties impacts the economy of a country. Therefore, it is important to forecast the evolution of TCs by using numerical models as the frequency of such storms is increasing in several basins of the world in the present warming period [4]. Therefore, it is attempted to put forward the recent developments in understanding the related meteorological characteristics, their predictability, climatological aspects and the gaps identified in the area of TC research.

is similar to hurricanes and typhoons.

The TC research has evolved over several decades and researchers use observations as well as numerical models for this purpose. For example, some observational studies [5–7] discuss about TC formation and evolution. The pioneering works by Gray [8, 9] have shown that the formation of TC at any location depends on six factors: (i) appropriate Coriolis parameter '*f*' that is practically effective 5° away from the equator in both the hemispheres, (ii) low-level positive relative vorticity (*ζ*r), that is, existence of initial disturbance, (iii) low tropospheric vertical wind shear (*S*z), (iv) ocean thermal energy (*E*) signified by sea surface temperature (SST), that is, SST should be ≥26.5°C within vertical extent of 60 m, (v) atmospheric instability measured in terms of difference in equivalent potential temperature (*θ*e) between the surface and 500 mb or Δ*θ*e and (vi) mid-tropospheric relative humidity 'RH'. Some of these aspects are discussed explicitly over the years using observations and numerical models [7, 10]. The first three parameters produce a dynamic potential (*fζ*r/*S*z), while the remaining three parameters yield a thermal potential (*E* Δ*θ*e RH). And, the product of dynamic and thermal potentials provides the seasonal genesis frequency.

Cyclogenesis do not occur spontaneously even if all the environmental conditions are met. Further, only about 10% of all cyclonic disturbances intensify into TCs. These low-pressure systems gradually form from a pre-existing (or initial) disturbance that consists of wind vortex and organized clouds. Thus, the necessary conditions for tropical cyclogenesis must be supported by the deep convection in the presence of a low-level absolute vorticity maximum and the initial convection must survive for sufficient time. The survival ability of the initial convection depends on '*ζ*r', '*f* ' atmospheric stability (defined by Brunt Vaisala frequency '*N*') and depth of the system (*H*). This ability is defined by the Rossby radius of deformation '*L*R' typically for a large tropical cyclonic system (www.meted.ucar.edu):

$$L\_{\aleph} = \frac{NH}{\zeta\_r + f} \tag{1}$$

The average life expectancy of a TC is about 1 week, whereas it is found that few cyclones remain active for more than 4 weeks (exact time frame may change from basin to basin) as seen in case of a hurricane, provided the system must be able to stay over the warm tropical waters. In most of the TCs, the Coriolis and centripetal forces oppose the pressure gradient force [11]. In the lowest kilometres near the surface, the frictional force destroys the gradient balance and consequently, air spirals inward towards the storm centres. The primary circulation (horizontal axisymmetric) during tropical cyclogenesis gains latent heat through the process of evaporation and exchange of sensible heat with the underlying ocean as it spirals towards the storm centre [12]. Consequently, it gains large angular momentum and kinetic energy because of the acceleration towards the low-pressure centre. The evaporation of sea spray provides the necessary moisture supply. Because of the high velocity demanded by the quasi-conservation of angular momentum, the air may not penetrate beyond some small radius. To conserve the angular momentum, the air spirals upward in the eyewall forming intense ring of cumulus cloud and a calm eye at the centre and brings in the latent heat it acquired during the upward motion in the boundary layer to the free atmosphere. Due to the cooling of this rising air, latent heat releases into the atmosphere to add more energy to the storm. Across the top of the boundary layer, the turbulent eddies generated by the mechanical mixing due to the prevailing strong winds cause a significant downward flux of sensible heat from the free atmosphere through subsidence (**Figure 1**).

**Figure 1.** Vertical cross section of a mature cyclonic storm and associated basic characteristics (adopted from http:// www.hko.gov.hk/informtc/nature.htm).

As the convective updrafts in the eyewall ascends to the tropopause, the latent heat is converted to sensible heat through condensation in order to provide the much-needed buoyancy for lifting air from the surface to tropopause level. After reaching at the upper level, the air turns outwards and eventually spread out at high altitudes, where it forms anticyclonic circulation and eventually the cool air above the eye begins to sink into the central core (**Figure 1**). Thus, the storm can be termed as a quasi-steady thermodynamic heat engine that is primarily driven by latent heat release. This heat engine runs between a warm heat reservoir as sea that is at ∼300°K and a cold reservoir located at 15–18 km up in the troposphere having a temperature ∼200°K. A baroclinic structure is maintained by the latent heat release in the warm core, which is continuously converted to kinetic energy that is responsible to drive the TC.

Apart from the basic factors discussed so far, there is a significant role of Madden-Julian oscillation (MJO) and El Niño in the frequency of occurrence of TCs (see [13]). In certain scenarios, equatorial Rossby waves (ER), mixed Rossby-gravity waves (MRG), Kelvin waves and easterly waves also influence the tropical cyclogenesis [14]. However, equatorial Kelvin waves do not appear to play a major role in tropical cyclogenesis. Several hurricanes in the North Atlantic form from African waves, MJO has a significant role in tropical cyclogenesis in North Pacific and the formation of cyclonic storms in the northwest Pacific is associated with MRG waves [13, 15]. These waves enhance the local conditions for the genesis of TCs by increasing upward motion, convection and the low-level vorticity by altering the local vertical shear pattern. The larger-scale waves, such as the MJO and ER, can also alter the mean zonal wind in large spatial and temporal scales in order to influence the mean flow.

The active phase of MJO is generally found over Indian Ocean, the maritime continent, and western pacific [13], which seem to play a major role in regulating the frequency of occurrence (usually increases) and formation of TCs in these regions. MJO increases the westerly wind which blows from west to east and its active phase through the region increases convective activity. During El Niño events, the atmospheric response to SST anomalies (SSTA) in the equatorial Pacific perturbs the Walker circulation [16]. The most common form of genesis occurs when they interact with the Asian monsoon. However, such type of interaction is still not studied well though few studies emphasized the role of El Niño or El Niño Southern Oscillation (ENSO) in TC formation over the Bay of Bengal (BOB) region indicating a decrease in the number of TCs [17].

#### **2.2. Structural evolution**

tion and exchange of sensible heat with the underlying ocean as it spirals towards the storm centre [12]. Consequently, it gains large angular momentum and kinetic energy because of the acceleration towards the low-pressure centre. The evaporation of sea spray provides the necessary moisture supply. Because of the high velocity demanded by the quasi-conservation of angular momentum, the air may not penetrate beyond some small radius. To conserve the angular momentum, the air spirals upward in the eyewall forming intense ring of cumulus cloud and a calm eye at the centre and brings in the latent heat it acquired during the upward motion in the boundary layer to the free atmosphere. Due to the cooling of this rising air, latent heat releases into the atmosphere to add more energy to the storm. Across the top of the boundary layer, the turbulent eddies generated by the mechanical mixing due to the prevailing strong winds cause a significant downward flux of sensible heat from the free atmosphere

196 Recent Developments in Tropical Cyclone Dynamics, Prediction, and Detection

**Figure 1.** Vertical cross section of a mature cyclonic storm and associated basic characteristics (adopted from http://

As the convective updrafts in the eyewall ascends to the tropopause, the latent heat is converted to sensible heat through condensation in order to provide the much-needed buoyancy for lifting air from the surface to tropopause level. After reaching at the upper level, the air turns outwards and eventually spread out at high altitudes, where it forms anticyclonic circulation and eventually the cool air above the eye begins to sink into the central core (**Figure 1**). Thus, the storm can be termed as a quasi-steady thermodynamic heat engine that is primarily driven by latent heat release. This heat engine runs between a warm heat reservoir as sea that is at ∼300°K and a cold reservoir located at 15–18 km up in the troposphere having a temperature ∼200°K. A baroclinic structure is maintained by the latent heat release in the warm core, which

Apart from the basic factors discussed so far, there is a significant role of Madden-Julian oscillation (MJO) and El Niño in the frequency of occurrence of TCs (see [13]). In certain scenarios, equatorial Rossby waves (ER), mixed Rossby-gravity waves (MRG), Kelvin waves

is continuously converted to kinetic energy that is responsible to drive the TC.

through subsidence (**Figure 1**).

www.hko.gov.hk/informtc/nature.htm).

The general structure of the TC can be understood through the visualization of vertical crosssection of a mature TC as depicted in **Figure 1**, which consists of eye, eyewall and rainbands. The centre of the structure signifies the low-pressure cyclone eye, where a strong downward flow occurs indifferent to the immediate neighbouring updrafts. However, subsidence is also visible alongside the updrafts in the neighbourhood eyewall region away from the eye. The appearance of eye, its growth, intensification of eyewall and disappearance of eye are described in this section.

The life cycle of the TC is shown in **Figure 2(a)**, where the inner and outer cores of a TC are considered besides its intensification in order to depict the strengthening and weakening. **Figure 2(b)** depicts the different stages of TC life period including genesis, development, mature stage and dissipation by considering the evolution of TC Phailin (2013) in BOB region as an example to the illustration shown in **Figure 2(a)**.

In the intensification period (or phase 1), the momentum from outer core towards the inner one helps in strengthening the 700 hPa wind field and subsequently it helps in the eye wall cloud formation [18]. Prior to the appearance of eye, the intensification process is quite slow (at a rate ∼8 hPa/day). The increase in maximum wind field is ∼5 ms−1day−1 and the outer core strengthens at a rate of ∼2 ms−1day−1. Gradually, when the eye appears, the central pressure is about 987 hPa and the rate of intensification increases by ∼250 times, at a rate of about 20 hPa/ day. The rapidly deepening cyclone (at a rate ∼42 hPa/day) supports an earlier eye formation. During the filling phase, the central pressure starts rising by drawing momentum through the outer core and strengthening the outer core's wing.

**Figure 2.** (a) Conceptual rendering from the main events in the life cycle of a tropical cyclone [18] and (b) different stages of tropical cyclone Phailin formed over Bay of Bengal [80].

The phase 2 is usually marked by the strengthening of outer core wind (similar to the stage (c) of TC Phailin shown in **Figure 2(b)**), whereas the inner core wind diminishes. During this phase, the eye expands and the filling of the inner core continues through the inflowing air towards the cyclone centre. During the filling phase, the inertial stability of the outer core is twice as large as that of the deepening stage making the outer core rigid for the inflowing air. Gradually the expansion of the eye ceases and the central core fills. It is important to note that the longer a cyclone spends in phase 2, stronger the outer radius will be and the radius of the damaging winds expand as long as the eye exists.

During phase 3, the outer wind starts weakening (e.g. stage (d) of TC Phailin shown in **Figure 2(b)**) with the disappearance of eye. Once the eye is vanished, the inflow of angular momentum ceases that was responsible for the strengthening of the outer core and from where the decay of outer radius low level wind field begins. These characteristics are valid for the cyclones which do not suffer landfall because the landfall would erode the wind field irrespective of the appearance of eye.

Though the basic principles of structural evolution may hold good for the TCs occurring in NIO basin (refer to **Figure 2(b)** for different stages of the TC Phailin), the formation of a distinguished eye structure may always not be feasible. A distinguished eye may be seen in case of a very severe cyclonic storm in this basin during phase 2. However, an explicit analysis in this direction is not available in literature for the NIO basin even though few studies like [19] computed the radius of maximum wind seen in case of TC intensification.

#### **2.3. Propagation**

**Figure 2.** (a) Conceptual rendering from the main events in the life cycle of a tropical cyclone [18] and (b) different

stages of tropical cyclone Phailin formed over Bay of Bengal [80].

198 Recent Developments in Tropical Cyclone Dynamics, Prediction, and Detection

TCs generally originate in tropics and thereafter, travel westward [20, 21] or turn poleward and recurve towards eastward direction [21, 22] or suffers extratropical transition over land or water [23] before dissipation. If a time scale of 1–3 weeks is considered, then the evolution of Rossby wave train significantly influences the track of a TC. Across the subtropical regions, under the influence of synoptic scale ridging, the TCs tend to move more westerly, and under the influence of synoptic scale trough, TCs tend to recurve into the mid-latitude [24]. On a seasonal scale, it is seen that over the Indian Ocean, the advancement of monsoon has a considerable impact on TCs' growth and their track [25].

In principle, TCs move under the influence of its surrounding environment. When the easterlies are added with the wind at certain level from the storm, the resulting effect forces the system to move in a westward direction [26]. Since the winds are not constant with height, it complicates the movement. The 'β effect' or 'β drift' pushes the cyclone towards the northwest direction in the northern hemisphere. It superimposes a weak northwest ward (southwest ward) steering current upon the TC in the northern (southern) hemisphere.

Apart from the factors mentioned earlier, the wind shear around anti-cyclonic flow at the top of the TCs also impacts their movement and can influence the track as much as the 'β drift'. There is a more complex phenomenon which influences the motion of a cyclone, known as 'Fujiwhara effect' [27]. Fujiwhara interaction describes the mutual rotation of two vortices about a common centre [28]. This centre typically refers to the mass weighted centroid of the two vortices, if they are of equal strength. In the presence of the β effect, the two vortices rotate around each other relative to the centre of rotation. This centre of rotation is not fixed and, instead, moves northwest ward in response to the 'β effect'. 'Fujiwhara effect' is noticed over other basins of the world including Atlantic, but is not applicable for TCs formed over NIO.

#### **2.4. Dissipation**

The most common way of dissipation of a TC is its landfall. When the storm moves over land, it deprives itself from warm water and the available moisture over ocean. Consequently, it is deprived from the energy source and the warm core with thunderstorms near the centre turns into a remnant low-pressure area due to quick loss of energy. Weakening can also occur if it encounters a vertical wind shear that causes the heat engine and convection shift away from the centre. The rate of power dissipation of TCs can be computed [29] as

$$E\_D = C\_D \rho \nu^3$$

where *E*D is the rate of energy dissipation per unit time per unit horizontal surface area, *v* defines the wind speed, *ρ* is for air mass density, and '*C*D' is the drag coefficient that depends upon the surface irregularities. Since the power dissipation in TCs is proportional to the cube of its wind velocity, the severity can be computed as the cumulative sum of the cube of the wind velocity over time according to the above equation.

#### **3. Role of ocean in genesis and intensification**

There are two sources which are capable of changing the TC intensity, one is internal variability and other one is environmental interaction. One important aspect of later source is the interaction between the ocean and the storm system. Usually TC is regarded as the most forceful case in air-sea interaction studies where energy from the warm ocean waters is delivered via surface heat flux [30]. The ocean response is quite sensitive to the surface drag coefficient. Emanuel [31] used a simple numerical model to establish the progress of hurricane intensity. Their findings advocate that in most cases, the intensity depends on three factors, viz. initial intensity of cyclone, thermodynamic state of atmosphere through which the cyclone propagates and finally the heat exchange with the upper layer of the ocean underlying the core of the cyclone. Rapid intensification of TC is noticed when it passes over the deep upper ocean mixed layer and that upper ocean thermal structure plays a significant role in the intensification process [32–34]. Sutyrin [35] performed simulations with a coupled model of the oceanic and atmospheric boundary layers and concluded that the interaction is strong enough to change the supply of heat and moisture fluxes from the ocean into the atmosphere significantly within few hours of the formation of the storm and consequently, influence the TC intensity.

The intensity of TC increases with increase in SST and upper ocean heat content [36]. The positive feedback occurs when genesis and intensification happens. During this phase, the evaporation from the ocean surface stimulates surface wind that subsequently increases the moisture supply and consequently increasing the latent heat that is further utilized to drive the circulation. As a negative feedback, the decrease in SST results in the decrease in total heat flux (sum of latent heat and sensible heat), resulting in decrease in intensity of the storm. Besides these interactions, some of the mechanical energy supplied by the TC is dispersed laterally and vertically by the internal inertia-gravity waves with time [37].

On the other hand, the intensification of TC depends not only on SST but also on subsurface ocean thermal structure also considered as an important predictor for the TC intensification (e.g. see [38–40]). In the changing climate scenario, SST plays a bigger role during pre-monsoon season as compared to the post-monsoon period for governing TC activity over NIO region [41]. In contrast, the same may not be valid for other basins including North Atlantic Ocean, where an increasing trend in correlation between SST and TC power dissipative index is observed [42]. The influence of the changing climate on the TC genesis and intensification in the NIO region may therefore not be limited to the analysis relating SST only.
