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- Author or Editor: Lynn K. Shay x
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Abstract
The interaction between a tropical cyclone (TC) and the underlying ocean is investigated using an atmosphere–ocean coupled model. The atmospheric model is developed from the Pennsylvania State University (Penn State)–National Center for Atmospheric Research (NCAR) mesoscale model version 4 MM4 and the ocean model consists of a mixed layer and an inactive stagnant layer beneath. Coupling between the atmosphere and the ocean models is achieved through wind stress and surface heat and moisture fluxes that depend on the sea surface temperature (SST). In the absence of a background flow, the atmospheric component consists of only a predefined vortex with an initial central pressure and the radius of the 15 m s−1 wind. The basic control experiments demonstrate that the coupled model can simulate the development of a TC and its interaction with the ocean.
Changes in TC intensity are sensitive to those of SST and the response is almost instantaneous. An SST of ∼27°C is found to be the threshold for TC development. In addition, the initial depth of the ocean mixed layer has an appreciable effect on TC intensity, which also depends on the movement of the TC. Furthermore, the vertical structure of ocean (vertical temperature gradient in the stagnant layer and temperature differential between the two layers) plays a significant role in modulating TC intensity.
In the presence of a warm core eddy (WCE), a TC intensifies before its center reaches the edge of the WCE. Although the TC attains maximum intensity at the center of the WCE, it does not weaken to its original intensity after leaving the WCE. During the entire passage of the TC, the SST at the center of the WCE decreases by about only 1°C, and the WCE generally maintains its original characteristics. However, two cold pools are observed around its periphery. A similar intensification process occurs when a TC moves over a sharp SST gradient and a locally deep ocean mixed layer. These results are explained by the interaction between the ocean and the TC circulation as well as the change in the total surface heat flux.
Abstract
The interaction between a tropical cyclone (TC) and the underlying ocean is investigated using an atmosphere–ocean coupled model. The atmospheric model is developed from the Pennsylvania State University (Penn State)–National Center for Atmospheric Research (NCAR) mesoscale model version 4 MM4 and the ocean model consists of a mixed layer and an inactive stagnant layer beneath. Coupling between the atmosphere and the ocean models is achieved through wind stress and surface heat and moisture fluxes that depend on the sea surface temperature (SST). In the absence of a background flow, the atmospheric component consists of only a predefined vortex with an initial central pressure and the radius of the 15 m s−1 wind. The basic control experiments demonstrate that the coupled model can simulate the development of a TC and its interaction with the ocean.
Changes in TC intensity are sensitive to those of SST and the response is almost instantaneous. An SST of ∼27°C is found to be the threshold for TC development. In addition, the initial depth of the ocean mixed layer has an appreciable effect on TC intensity, which also depends on the movement of the TC. Furthermore, the vertical structure of ocean (vertical temperature gradient in the stagnant layer and temperature differential between the two layers) plays a significant role in modulating TC intensity.
In the presence of a warm core eddy (WCE), a TC intensifies before its center reaches the edge of the WCE. Although the TC attains maximum intensity at the center of the WCE, it does not weaken to its original intensity after leaving the WCE. During the entire passage of the TC, the SST at the center of the WCE decreases by about only 1°C, and the WCE generally maintains its original characteristics. However, two cold pools are observed around its periphery. A similar intensification process occurs when a TC moves over a sharp SST gradient and a locally deep ocean mixed layer. These results are explained by the interaction between the ocean and the TC circulation as well as the change in the total surface heat flux.
Abstract
The thermodynamic effect of downdrafts on the boundary layer and nearby updrafts are explored in idealized simulations of category-3 and category-5 tropical cyclones (TCs) (Ideal3 and Ideal5). In Ideal5, downdrafts underneath the eyewall pose no negative thermodynamic influence because of eye–eyewall mixing below 2-km altitude. Additionally, a layer of higher θ e between 1- and 2-km altitude associated with low-level outflow that extends 40 km outward from the eyewall region creates a “thermodynamic shield” that prevents negative effects from downdrafts. In Ideal3, parcel trajectories from downdrafts directly underneath the eyewall reveal that low-θ e air initially moves radially inward allowing for some recovery in the eye, but still enters eyewall updrafts with a mean θ e deficit of 5.2 K. Parcels originating in low-level downdrafts often stay below 400 m for over an hour and increase their θ e by 10–14 K, showing that air–sea enthalpy fluxes cause sufficient energetic recovery. The most thermodynamically unfavorable downdrafts occur ~5 km radially outward from an updraft and transport low-θ e midtropospheric air toward the inflow layer. Here, the low-θ e air entrains into the updraft in less than 5 min with a mean θ e deficit of 8.2 K. In general, θ e recovery is a function of minimum parcel altitude such that downdrafts with the most negative influence are those entrained into the top of the inflow layer. With both simulated TCs exposed to environmental vertical wind shear, this study underscores that storm structure and individual downdraft characteristics must be considered when discussing paradigms for TC intensity evolution.
Abstract
The thermodynamic effect of downdrafts on the boundary layer and nearby updrafts are explored in idealized simulations of category-3 and category-5 tropical cyclones (TCs) (Ideal3 and Ideal5). In Ideal5, downdrafts underneath the eyewall pose no negative thermodynamic influence because of eye–eyewall mixing below 2-km altitude. Additionally, a layer of higher θ e between 1- and 2-km altitude associated with low-level outflow that extends 40 km outward from the eyewall region creates a “thermodynamic shield” that prevents negative effects from downdrafts. In Ideal3, parcel trajectories from downdrafts directly underneath the eyewall reveal that low-θ e air initially moves radially inward allowing for some recovery in the eye, but still enters eyewall updrafts with a mean θ e deficit of 5.2 K. Parcels originating in low-level downdrafts often stay below 400 m for over an hour and increase their θ e by 10–14 K, showing that air–sea enthalpy fluxes cause sufficient energetic recovery. The most thermodynamically unfavorable downdrafts occur ~5 km radially outward from an updraft and transport low-θ e midtropospheric air toward the inflow layer. Here, the low-θ e air entrains into the updraft in less than 5 min with a mean θ e deficit of 8.2 K. In general, θ e recovery is a function of minimum parcel altitude such that downdrafts with the most negative influence are those entrained into the top of the inflow layer. With both simulated TCs exposed to environmental vertical wind shear, this study underscores that storm structure and individual downdraft characteristics must be considered when discussing paradigms for TC intensity evolution.
