A Case Study of the Formation of an Eastern Pacific
Tropical Cyclone
by
Terence Kung
B.S. Atmospheric and Oceanic Sciences (1997)
University of Wisconsin at Madison
Submitted to the Department of Earth, Atmospheric and Planetary Sciences in Partial
Fulfillment of the Requirements for the Degree of
MASTER OF SCIENCE IN METEOROLOGY
at the
MASSACHUSETTES INSTITUTE OF TECHNOLOGY
June 1999
© 1999 Massachusetts Institute of Technology. All Rights Reserved.
Author ......
........ .........
...... ...
Department of Earth, Atmospheric and Planetary Sciences
May 6, 1999
Certified by ..............
...
Kerry A. Emanuel
Professor of Meteorology
Thesis Supervisor
Accepted by........................................
Ronald G. Prinn
I
Department Head
MASSACHUSETTS INSTITUTE
OF TECHNOLOGY
JUN 0 1 1999
LIBRARIES
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A Case Study of the Formation of an Eastern Pacific Tropical Cyclone
by
Terence Kung
Submitted to the Department of Earth, Atmospheric and Planetary
Sciences on May 7, 1999 in partial fulfillment of the requirements
for the Degree of Master of Science in Meteorology
Abstract
A case study is performed to investigate the nature of tropical cyclogenesis in the
eastern Pacific Ocean. Focus is given to the formation and development of the initial
circulation which eventually intensified into Hurricane Fefa. Using satellite imagery, the
author studies the development of convective activity in the genesis region. Gridded
reanalysis data are used to document the synoptic-scale flow, with emphasis on tracing the
easterly wave which is associated with the formation of Fefa. The data show that the
easterly wave propagated across the Caribbean Sea and the Central American mountains,
and the initial circulation developed while the wave had moved into the eastern Pacific.
The wave is found to have moved through an unstable basic state while it was in the
Caribbean, which is favorble for its growth and maintenance. Two phenomena are
observed prior to the formation of the low-level circulation. These include an easterly jet
in the eastern Pacific that may have been associated with the blocking effect of the Central
American mountains, and a southerly wind surge into the monsoon trough region.
In addition, aircraft observations collected during Tropical Experiment in Mexico
are used to study the evolution of the mesoscale system. Initially, a circulation in the
middle troposphere with a cold core in the boundary layer, and a shear line located to the
west were found. One day later, a low-level warm core vortex had developed, and it was
displaced from the mid-level vortex. It is suggested that the low-level vortex formed from
the spin-up of the monsoon trough, independent of the mid-level vortex.
Thesis Supervisor: Kerry A. Emanuel
Title: Professor of Meteorology
Acknowledgements
I would like to thank professor Kerry Emanuel for stimulating my interest in the
problem of tropical cyclogenesis, and suggesting this topic to me. I thank him for his
guidance, constructive criticisms along the way, and reviewing several drafts of my thesis.
I would also like to thank professor Alan Plumb for providing me financial support, and
acting as my academic advisor. I value the friendships with fellow graduate students in the
department. I appreciate their warmth and concerns with me.
I thank Lodovica Illari for supplying me tapes of the NCEP/NCAR reanalysis data.
I also thank Luis Farfan of the University of Arizona for sending me the satellite images,
and helping me out in some displaying problems.
I thank all brothers and sisters in the Boston Chinese Evangelical Church for their
love and spiritual support. Most of all, I thank my parents for their love, support and
understanding.
The research was supported under NSF grant ATM-9528471.
Table of Contents
1 1.1 Introduction
1.2 Overview of the Tropical Experiment in Mexico (TEXMEX) and Hurricane
Fefa
1.3 Review of previous work on tropical cyclogenesis
11
2 Data and analysis methods
2.1 Doppler radar data
2.2 In situ data
2.3 Satellite data
2.4 NCEP/NCAR reanalysis data
23
26
29
30
30
3 Observations of the synoptic-scale circulation
31
4 Aircraft data analysis results
4.1 Flight 1P
Flight 2E
Flight 3P
4.2 Comparison to the genesis of Guillermo
51
51
58
62
70
5 Summary and suggestions for Future Work
73
6 Bibliography
76
13
14
_~II~~_ -I.__XII~ULII.
List of Figures
Figure 2.1: Tracks of aircraft-estimated vortex centers of TEXMEX cases that developed
into hurricanes (from D. Raymond)
24
Figure 3.1: Winds and relative vorticity from the NCEP/NCAR reanalysis at 700 mb
32
Figure 3.2: Winds and relative vorticity from the NCEP/NCAR reanalysis at 1000 mb
37
Figure 3.3: Infrared images from GOES
41
Figure 3.4: Absolute vorticity from the NCEP/NCAR reanalysis at 700 mb
45
Figure 4.1: Observations in pre-Fefa MCS during flight IP
52
Figure 4.2: Observations in pre-Fefa MCS during flight IP (cont'd)
57
Figure 4.3: Observations in pre-Fefa MCS during flight 2E
59
Figure 4.4: Observations in Tropical Storm Fefa MCS during flight 3P
63
Chapter 1
1.1 Introduction
Tropical cyclones are fascinating natural phenomena for many meteorologists. One reason
is that over the years these storms have caused enormous loss of human lives and significant
economical impact. The second reason is that they are still not fully understood and pose great
challenges to the scientific community. An outstanding problem in current tropical cyclone research
is their formation, or the genesis problem. It is not only a subject of scientific interest, but also of
practical interest. It is because many mariners and coastal or island communites are quite
vulnerable to sudden formation of tropical cyclones. If the time and location of their formation can
be well predicted, more advanced warnings can be given to these communities and thereby human
and economic loss can be minimized.
The necessary environmental conditions for tropical cyclones to form have been well
established. They include large values of low-level relative vorticity, a coriolis parameter greater
0
than some finite value, weak vertical wind shear, and sea surface temperatures exceeding 26 C (eg.
Gray 1968). However, the sufficient conditions for their formation are less clear, as evidenced by
the fact that only a small fraction of the disturbances observed in the Tropics grow into tropical
storms even when the necessary conditions are met. One of the difficulties is that there are limited
data over the oceans where tropical cyclones form. A detailed observation of the formation process
is needed to further our understanding of this problem.
The subtropical eastern Pacific Ocean is the most prolific region in the world for the
formation of tropical cyclones. The tropical cyclones in this region are often associated with
African easterly waves (Avila and Pasch 1992). While these easterly waves have often been
observed to trigger tropical cyclones, the real mechanism has not yet been well understood. It is not
dl^ l*ri~~.r.ar--irt*Y6-
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clear whether they transform directly into tropical cyclones, or they merely provide a favorable
background environment for pre-existing tropical disturbances to develop into tropical cyclones.
The purpose of this study is to provide a detailed documentation of the formation of
Tropical Cyclone Fefa in the eastern Pacific. Fefa formed during the Tropical Experiment in
Mexico (TEXMEX), an intensive research program on tropical cyclone formation in the eastern
North Pacific. Aircraft observations provide detailed mesoscale data during the cyclone's
formation. Attention is given to the thermodynamics of the genesis process. In addition,
NCEP/NCAR reanalysis, received on 2.50 x 2.50 grids, are used to study the characteristics of the
large-scale environment. Particular attention is given to the evolution of the easterly wave
associated with the formation of Fefa.
In section 1.2, an overview of TEXMEX and hurricane Fefa is given. The next section
contains a thorough review of previous studies of tropical cyclogenesis.
Chapter 2 gives a description of the data sets and methods of analysis. Chapter 3 describes the
observations of the large-scale environment. Chapter 4 contains an analysis of the aircraft data. A
summary of the overall findings of this study and suggestions for future work are given in the final
chapter.
1.2 Overview of the Tropical Experiment in Mexico (TEXMEX) and Hurricane Fefa
One of the outstanding problems in tropical cyclone research is why so many disturbances
in the tropics fail to develop into tropical cyclones. Rotunno and Emanuel (1987, hereinafter RE)
suggested that convective downdrafts in the core of the incipient disturbances bring air of low
equivalent potential temperature (0e) into the boundary layer, suppressing further convection and
thereby preventing their subsequent development. For the disturbances to further develop, the
negative effect of the downdrafts has to be overcome. In principle, this can be accomplished by an
increase of the equivalent potential temperature in the middle troposphere, an increase of relative
humidity so that evaporation of rain is suppressed, and/or an increase of wind speed in the
boundary layer so that the sea surfaces fluxes can keep replenishing the Ge in that layer. The
simulations of RE and Emanuel (1989), in which a warm core vortex was used in the initial state,
suggested that an increase of Oe in the middle troposphere is necessary for tropical cyclogenesis.
The main goal of TEXMEX was to test a hypothesis stated in the TEXMEX Operations Plan
(Emanuel 1991): The elevation of ee in the middle troposphere just above a near surface vorticity
maximum is a necessary and perhaps sufficient condition for tropical cyclogenesis. It was assumed
that the elevation of ee is accomplished by deep convection bringing high 8e to the middle
troposphere, as occurred in the models initialized by warm core vortices. In the case study of the
genesis of hurricane Fefa, analysis showed that there was a moderate increase of ee at 3 km
altitude, and the pre-Fefa system evolved quickly into a tropical storm. The TEXMEX hypothesis
is not rejected, at least, by this case study. But in another TEXMEX case study ie. the genesis of
hurricane Guillermo (Bister and Emanuel 1997, hereinafter BE), observations suggested that the
increase of Oe in the middle troposphere is not a necessary condition for tropical cyclogenesis.
During the first flight into the pre-Fefa system, the surface circulation is already strong,
and it is displaced well to the west of the 700 mb circulation (Raymond et al., 1998). The fact that
no flights surveyed the earliest state of the system suggests that this case study may not be the most
ideal one for testing the TEXMEX hypothesis. But the study is interesting in its own right, partly
because the formation of Fefa looks so different from the formation of Guillermo. This suggests
that different mechanisms might have taken place, or played the major roles, in the genesis process
in the two cases. One unresolved question posed in BE is whether the downward extension of the
cold core vortex is a necessary condition for the development of the warm core vortex. The case
study of Fefa serves to further address this question. Moreover, the focus of BE was on the
mesoscale aspects. In this study, emphasis is also put on the large-scale environment. The premise
of this study is that by combining detailed analyses of the large-scale data and the mesoscale data,
one can obtain a more complete picture of the mechanisms involved in tropical cyclogenesis.
1.3 Review of previous work on tropical cyclogenesis
Two major theories have been proposed to explain the instability mechanism responsible
for the growth of tropical depression. The first one is referred to as Conditional Instability of the
Second Kind (CISK), proposed by Charney and Eliassen (1964). This mechanism states that
tropical depressions intensify by utilizing the Convective Available Potential Energy (CAPE)
through a cooperative feedback between cumulus clouds and the large-scale flow. In the region
where there is low level cyclonic vorticity, frictional convergence at the surface forces vertical
motion. The rising parcels are warmed relative to the environment by latent heat release. This
warming is balanced by an in-up-out seconary circulation that provides convergence in the lower
troposphere, which increases the low level vorticity through vortex stretching. The increased
vorticity then leads to increased frictional convergence at the surface and thereby more vertical
motion and latent heat release. Thus a positive feedback loop is established. The CISK mechanism
does not require large-scale saturation, but it does require that the atmosphere be conditionally
unstable. The existence of CAPE in the atmosphere associated with conditional instability has been
challenged (eg. Xu and Emanuel 1989). They showed that soundings from the deep tropics are
nearly neutral to reversible ascent in the lower and middle troposphere. Another important
drawback of CISK is that its closure is valid only after an incipient vortex has attained sufficient
strength and organization. It does not explain how the incipient vortex forms. As noted by Ooyama
(1982), this theory was never intended for inquiry into the question of tropical cyclogenesis.
More recently, an alternative theory to CISK, known as air-sea interaction theory or WindInduced Surface Heat Exchange (WISHE) has been proposed (Emanuel 1986; RE 1987; Emanuel
1989). The basic idea of WISHE is that tropical cyclones intensify and maintain themselves
entirely by self-induced anomalous fluxes of moist enthalpy from the sea surface, which are
primarily determined by the magnitude of the surface winds. An advantage of WISHE is that it
does not require any contribution from preexisting CAPE in the troposphere.
However, the
WISHE mechanism requires that an initial vortex be sufficiently strong in order for a cyclone to
develop into hurricane intensity in a reasonable time. The theory does not account for the processes
by which the finite amplitude rotary system (or 'starter vortex') itself forms.
As suggested above, it is important to know how the 'starter vortex' forms in order to
solve the problem of tropical cyclogenesis. One proposed source of these initial vortices is given by
the propagation of synoptic-scale easterly waves. These waves generally originate over western
Africa, move across the Atlantic, and eventually reach the Caribbean Sea (eg., Burpee 1972; Reed
et al., 1988). They typically have a cyclonic vortex at midddle levels with a weak cold core below
the vortex and a warm core above, as required by thermal wind balance. The waves often manifest
themselves as vorticity anomalies elongated into the form of a trough aligned in the northeastsouthwest direction (eg., Shapiro 1986). When the waves move across Central America and enter
the eastern Pacific, they apparently gain energy (Nitta and Takayabu 1985). The reason for this is
r-- -------rr-r~---~
still unknown. Observations from satellite imagery suggest a link between the passage of easterly
waves over the Caribbean and the subsequent formation of tropical cyclones in the eastern Pacific.
Avila and Pasch (1992) were able to associate each eastern Pacific tropical cyclone with a
synoptic-scale easterly wave from Africa for the 1991 hurricane season. However, it is not clear
from observations what physical mechanisms may result in the reduction of horizontal scale of the
waves. Besides, it is known that a large fraction of the easterly waves fail to transform into
hurricanes (Avila 1991). While Shapiro (1986) showed that strong African waves can maintain
their structure while propagating across the Atlantic Ocean and into the eastern Pacific, there have
been suggestions that eastern Pacific easterly waves might develop due to local or upstream
instabilities, instead of being the same waves that formed in Africa (Molinari et al., 1997).
From a theoretical perspective, Shapiro (1977) proposed a storm genesis criterion based on
the nonlinear dynamics of the easterly wave. Attempts have been made to simulate the
transformation of an easterly wave into a tropical storm (Kurihara and Tuleya 1981; Kurihara and
Kawase 1985). They showed that diabatic heating due to the condensation of water vapor is
essential for the development of the wave into a tropical storm. They also showed that for
westward wave propagation, an easterly shear in the basic flow is favorable for wave growth
because it helps to advect warmed upper air in the same sense as the phase propagation of the
wave. Pytharoulis (1999) used the United Kingdom Meteorological Office (UKMO) Unified model
to show that the transformation of an initially cold core African easterly wave into a warm core
tropical cyclone took place after strong convective bursts. He emphasized the importance of the
role of regions with enhanced latent heat fluxes along the track of these waves for strong
convective events. He further showed that evolution over warm sea surface temperatures (SST) has
enhanced likelihood for development.
-iarC~-~
Some researchers have emphasized the role of the orography of Central America on
eastern Pacific tropical cyclogenesis. One proposed mechanism is that an easterly wave incident on
the Central American mountains is modified by the topography (Zehnder 1991).
That author
showed that there is a southward deflection of the incident flow by the anticyclonic circulation over
the mountains and a reduction of planetary vorticity. To conserve potential vorticity, cyclonic
vorticity must occur in the lee of the mountains. The area of vorticity has a structure that is
consistent with the circulations that intensify in the WISHE theory. Another mechanism related to
the topographic influence of the Central American mountains that does not require a synoptic-scale
easterly wave to propagate across the mountains was proposed by Mozer and Zehnder (1996).
They showed that when there is a stably stratified easterly barotropic flow over idealized
topography representative of the Sierra Madre, a barotropically unstable jet results from the
blocking of the low level flow by the topography. Low level vorticity maxima are continuously
produced downstream of the mountain, resulting in a low level easterly wave train in the region that
corresponds to the eastern Pacific. The authors suggested that surface fluxes of heat and deep
convection in the Pacific would allow for a further intensification of the vorticity that developed in
their study.
One different approach to the problem of tropical cyclogenesis places emphasis on external
forcing. As early as in the 50's, the importance of eddy angular momentum fluxes associated with
upper-level wave asymmetries in the environment of mature tropical cyclones was recognized
(Pfeffer 1958). Observational studies have demonstrated an initiation of the cyclogenesis process
through interaction with surrounding upper-tropospheric synoptic systems, particularly upper-level
troughs (eg., Riehl 1948; McBride and Keenan 1982). Sadler (1976) showed that tropical cyclone
formation in the western North Pacific often occurs in association with intense cyclonic cells
embedded in the Tropical Upper Tropospheric Trough (TUTT). Observations during the
~^_I__^
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Australian Monsoon Experiment (AMEX) document the formation of tropical cyclones Irma and
Jason (Davidson et al., 1990). Their work suggest that both cyclones formed in association with an
equatorward-moving TUTT. But while Sadler hypothesized that the upper-level troughs played a
kinematic role in facilitating the development of the tropical cyclone's characteristic upper-level
outflow channel, other researchers (eg., Holland 1983) demonstrated that the role of the adjacent
upper-level troughs in developing tropical cyclones is to bring about large imports of relative
angular momentum by eddy fluxes at outer radii in the upper troposphere. Challa and Pfeffer
(1990) further substantiated this idea by using a three-dimensional primative equation model to
illustrate rapid hurricane development in cases of strong eddy momentum flux convergence toward
the storm center. Using a moist nonlinear balanced model, Montgomery and Farrell (1993) showed
that an upper tropospheric potential vorticity (PV) anomaly approaching a weak surface cyclone
may result in the spin-up of the surface cyclone. Molinari et al. (1995) stressed the importance of
scale reduction of the upper-tropospheric PV anomaly by the upper-tropospheric anticyclone
associated with the tropical cyclone prior to their superposition. They noted that this helps to
reduce the magnitude and duration of vertical wind shear associated with the upper-tropospheric
PV anomaly which would hinder the tropical cyclone's development. In our study, there was no
evidence in the NCEP reanalysis data of any independent upper-tropospheric positive PV
anomalies in the vicinity of the region where hurricane Fefa formed, though, to be sure, the eastern
Pacific is nearly devoid of data.
One important source of initial cyclonic circulations over the eastern Pacific region is
Mesoscale Convective Systems (MCSs). The MCSs form over inland Mexico or over the Central
American mountains. They propagate westward under the influence of the tropical easterlies, and
eventually move over the warm waters of eastern Pacific, where they may evolve into hurricanes.
MCSs typically have a horizontal scale of hundreds of kilometers, and a lifetime of hours. In
satellite imagery, MCSs can be identified as the large areas of cloud with temperatures lower than 70 0 C. Studies show that some MCSs develop an inertially stable, warm core vortex in a trailing
stratiform region referred to as Mesoscale Convective Vortex (MCV) (Brandes 1990; Chen and
Frank 1993). The MCVs have a horizontal scale of approximately 100 to 200 km, and the vorticity
maximum is in the middle troposphere. Bosart and Sanders (1981) tracked a mid-latitude MCS in
an observational study. When it was still over the continent, it developed a mid-level vortex but
there was no surface circulation. But later when the system moved over the ocean, the system
developed into a tropical cyclone. Velasco and Fritsch (1987) found evidence from satellite
imagery that MCSs lead to tropical cyclogenesis. They suggested that when these MCSs propagate
into a large-scale environment favorable for tropical cyclogenesis, the generated MCV may play a
catalytic and crucial role in initiating tropical storm development. The question of how MCSs form
is still unanswered, and is beyond the scope of this study.
Theories have been proposed to explain the formation and maintenance of MCVs. Latent
heat release due to convection warms the stratiform-cloud region of an MCS and evaporation of
rain cools the region below the cloud, sharpening the potential temperature gradient near the cloud
base. Mesoscale convergence and stretching that develops just above the cloud base (Houze 1977)
produces a positive PV anomaly. Raymond and Jiang (1990) (hereinafter RJ) proposed that the
interaction of the PV anomalies with a sheared environment can produce enough low-level lifting to
feed the observed convection. Chen and Frank (1993) studied the formation of the mesovortex by a
three-dimensional hydrostatic mesoscale model. They suggested that within a saturated stratiform
rain region, the local Rossby radius is reduced to a scale below the scale of the stratiform rain
region of the MCS due to the reduction of vertical stability in the saturated region. As a result,
gravity waves cannot be sustained. This facilitates the conversion of latent heating to warming and
creates balanced mesovortices within the stratiform region.
I-l-----L1-^-~--rur o--l--^-rr~------rCL
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In an observational study, Fritsch et al. (1994) (FMK, hereinafter) documented a case in
which a warm core mesoscale cyclone forms in conjunction with a mid-latitude MCS, and through
repeated cycles of deep convection, it strengths and grows. In their case, the PV and potential
temperature patterns were basically the same as those postulated by RJ. FMK found that in most
of the redevelopment cycles, each outbreak of new convection occurred well within the vortex
circulation, usually near its center. Note that in FMK, the mesovortex developed in an atmosphere
characterized by a low-level southwesterly jet of high ee air overlain by a deep, warm, midlevel
layer of slightly weaker wind speeds and very weak vertical wind shear. The low-level jet plays the
role of feeding new convective developments. Though many studies have documented cases where
MCSs developed into tropical depressions, a fundamental problem is that in MCS events the
midlevel vortices are attended by surface mesohighs with anticyclonic circulations. The persistence
of high surface pressure presents great difficulties in arguing that the midlevel vortices can be
instrumental in initiating tropical cyclogenesis. For a tropical depression to form, the surface
mesohigh must be replaced by a mesolow with cyclonic circulation. One way to achieve this is by a
downward migration of the midlevel vortex. Once the vortex migrates downward to the surface and
becomes in contact with the oceanic heat and moisture source, the air-sea interaction instability can
get started and the surface cold core may transform into warm core. But the question still remains
as to what physical mechanism is responsible for the downward migration of the midlevel vortex.
Rogers (1997) proposed that a possible mechanism is increasing the strength of the PV anomaly by
initiating convection within the existing anomaly. Strengthening the PV anomaly will increase the
penetration depth (Hoskins et al., 1985). It is conceivable that after several convective
redevelopments, the cyclonic circulation will eventually reach the surface. By a three-dimensional
mesoscale model, Rogers attempted to simulate the MCS event in FMK, and found that the
cyclonic circulation penetrated approximately 100 mb closer to the surface. As in FMK, the
~_____--_----__--rirri-Lu;rur
m~
convection is initiated by the advection of high Oe air within a strong low-level jet ascending over
the cold pool left from the previous convective cycle.
Some researchers have emphasized the importance of vortex interaction in tropical
cyclogenesis. Ritchie et al. (1993) hypothesized that certain configurations of the midlevel PV
anomalies, or certain configurations of the large-scale flow would lead to a merger of the
mesovortices, resulting in a large rotation center of enhanced PV, and thereby an increased
penetration depth. In a study of the formation of Typhoon Irving during the Tropical Cyclone
Motion (TCM-92) experiment, Ritchie and Holland (1997) noted a cooperative interaction between
environmental and mesoscale dynamics. They found that besides the interactions between midlevel
vortices which lead to a combined vortex of greater depth, there is an interaction between midlevel
vortices and the low-level circulation which produces a development downward of the midlevel
vorticity, and strengthens the surface vortex. Simpson et al. (1997) investigated this scaleinteraction in other cases of tropical cyclone formation, and emphasized the importance of the
spinup of the monsoon trough in the background, which provides enhanced conditions for the
downward development of the midlevel vortices to the surface.
There is yet another theory for the downward migration of the midlevel vortex. BE
hypothesized that evaporation of rain below the stratiform precipitation region increases the
relative humidity in the lower troposphere and leads to a downdraft that advects the vortex
downward. Using an axisymmetric model, they showed that if the precipitation lasts long enough,
the vortex will be able to reach the boundary layer. As a result, convection redevelops and this
leads to a further increase of vorticity below the maximum heating and the formation of a warm
core near the surface.
In summary, we now have a modest understanding of the mechanism by which a tropical
storm intensifies. However, no well-accepted, closed theory for the formation of the initial
~~
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circulation yet exists. Plenty studies have shown that easterly waves are intimately related to
tropical cyclogenesis. But the mechanism responsible for the transformation of these waves into
tropical cyclones is not completely clear. For genesis in the eastern Pacific region, it is apparent
that the Central American mountains play a role. But questions still remain as to what their exact
role is and how important it is compared to other factors. The influence of upper-level PV
anomalies is not well understood. It is not known how important in practice they are for tropical
cyclogenesis, and in what exact way they initiate the formation process. The problem of how
MCSs evolve into tropical cyclones is also not completely solved. Various mechanisms have been
proposed for the downward migration of the mid-level vortex to the surface. But it is not known
which mechanism is dominant in practice. Furthermore, there still might be other mechanisms not
yet discovered. Our understanding of tropical cyclogenesis will remain incomplete until all the
aforementioned problems are solved.
Chapter 2
Data and analysis methods
TEXMEX was designed to observe intensively the process of tropical cyclogenesis, with
the principal aim of testing the aforementioned hypothesis about how genesis takes place. This
hypothesis was tested by making measurements inside developing and nondeveloping cloud
clusters, using the WP-3D aircraft operated by the National Oceanic and Atmospheric
Administration's (NOAA) Office of Aircraft Operations and the National Center for Atmospheric
Research (NCAR) Lockheed Electra. Both aircraft were equipped to make in situ measurements of
standard meteorological variables, and the WP-3D had the additional capability of deploying
Omega dropwindsondes (ODW) and making detailed Doppler radar measurements. The aircraft
measurement systems are described in detail in the TEXMEX Operations Plan (Emanuel 1991).
The eastern tropical North Pacific region was selected for the field program, as it has the
highest frequency of genesis per unit area of any region worldwide (Elsberry et al., 1987) and the
main genesis region is only a few hundred kilometers south of Acapulco, Mexico, which has an
airport well suited to research flight operations. The field phase of the experiment began on 1 July
and ended on 10 August 1991. During this project there were six intensive operation periods
(IOPs) that surveyed one short-lived convective system, one nondeveloping mesoscale cloud
cluster, and four cloud clusters that ultimately developed into hurricanes. The tracks of those
systems that developed into hurricanes are shown in Fig. 2.1. Detailed description of the aircraft
flight operations are provided in the TEXMEX Data Summary (Renno et al., 1992).
5.
-110
.
-105
,
.
-100
-95
-90
-85
Figure 2.1: Tracks of aircraft-estimated vortex centers of TEXMEX cases that developed into
hurricanes (from D. Raymond). Circle shows where each disturbance was declared a tropical storm
by National Hurricane Center. E is for Enrique, F for Fefa, G for Guillermo and H for Hilda.
As the principal working hypothesis concerned thermodynamic transformations of the
lower and middle troposphere, most flight operations were conducted near the 700-mb level and in
the subcloud layer. Most of the NOAA WP-3D flights at 700 mb deployed ODWs, and the tail
Doppler radar on the WP-3D operated almost uninterrupted through all of the flight operations. In
order to maximize the temporal continuity of observations of evolving cloud clusters, while obeying
operational contraints, the aircraft flew alternating missions at approximately 14-h intervals. Most
flight missions lasted 7-9 hours, of which 1-3 hours were used in transit to the target area.
A case study of the genesis of hurricane Guillermo has been performed (BE). In this work,
the MCS that developed into hurricane Fefa was chosen as the case study, because it has the best
documented data apart from the Guillermo case. It has relatively good data coverage. The early
state of the system was captured during the first flight. The system also has an interesting low level
cold core structure like the pre-Guillermo system. This study serves as an additional case study to
test the hypothesis made in the BE study, and also allows us to do a comparison between the two
cases.
The MCS that developed into hurricane Fefa on 31 July 1991 was the target object of
IOP4 from 28 July 1991 to 29 July 1991. Three flights were flown during IOP4. The first and the
third flights were flown with the WP-3D. These flights are labeled IP and 3P, respectively. The
second flight was flown with the Electra, and is labeled 2E. The flights are summarized in table
2.1. Each flight consisted of several flight legs at 3 km (700 mb), and several more at 300 m.
Flight
Target
Date
Time(UTC) at 700mb
Time (UTC) at 300 m
IP
pre-Fefa
07/28/91
21.47 - 01.26 F
01.33 - 03.15 F
2E
pre-Fefa
07/29/91
11.46- 14.05
14.07- 15.02
3P
TS Fefa
07/29/91
01.35 - 03.27 F
03.35 - 04 40 F
Table 2.1 Summary of flights into (pre-)Fefa. F is for following day, TS is for tropical storm
In the stratiform precipitation areas, measurements made by the the ODWs often showed
100% relative humidity, indicating wetting of the instruments. Owing to the wetting, the ODW data
are not used in the data analysis. Geostationary Operational Environmental Satellite (GOES)
imagery was obtained from L. Farfan of the University of Arizona.
-r~lr------I-~
-I~
2.1 Doppler radar data
The characteristics of the WP-3D Doppler radar are given in Table 2.2. The unambiguous
range and velocity are dictated by the pulse repetition frequency and the wavelength. The number
of samples per each radar grid volume depends on the distance of the grid volume from the aircraft.
There are at least 32 samples per each radar grid volume. The number of samples per grid volume
increases with distance from the aircraft, being 128 for distances larger than 38.4 km.
Radar characteristic
Value
Frequency
9.315 GHz
Wavelength
3.22 cm
Pulse length
0.55 ps
Pulse repetition frequency
1600 per sec
Beam width - along track
1.350
Beam width - across track
1.900
Unambiguous velocity
12.88 m/s
Unambiguous range
93.75 km
Table 2.2 The characteristics of the NOAA WP-3D aircraft's Doppler radar
At least two beams from different angles are needed for the determination of the three
dimensional wind. In TEXMEX the WP-3D Doppler radar was used in the Fore-Aft Scanning
Technique (FAST) mode. In this mode, the antenna tilt angle, defined forward or aft from the
perpendicular to the ground track, is changed between each rotation of the antenna about the
aircraft's longitudinal axis.
When FAST is used, Doppler velocities are contaminated by the component of the velocity
of the aircraft in the direction of the antenna. Thus, wind measurement is prone to errors in the
aircraft's ground velocity. On the other hand, the FAST mode is practical when it is not possible to
plan the flight pattern in advance. This is because the two components of the wind can be retrieved
from one flight leg.
The Doppler data were mapped to a 3 x 3 x 0.5 km grid (0.5 km in the vertical direction)
by averaging the data in the horizontal direction and interpolating in the vertical direction. The
components of the aircraft's ground velocity and precipitation particle fallspeed in the direction of
the antenna were subtracted from the radial velocities. The terminal fallspeed was estimated using
empirical fallspeed - radar reflectivity relations (eg. see Marks and Houze 1987). The velocities
were then unfolded using Bargen and Brown's method (1980). An independent measure of the wind
speed is needed to unfold the velocities. The in situ aircraft measurement of wind was used for this
purpose. Manual editing of the data followed automatic unfolding. Possible wrong unfolding was
corrected for by adding a suitable number of unambiguous velocities (see Table 2.2) to any
suspicious values. If the velocity still seemed unrealistic, the data were deleted. Most of the data
deleted were from 0.5, 1.0, and 1.5 km altitudes. Data from these altitudes appear to have been
compromised by sea clutter. Suspicious looking winds at higher altitudes were relatively rare, and
probably owing to second or multiple trippers, or sidelobes. Finally, the three- dimensional wind
was calculated from the two radial components in the following way (the computer program was
written by John Gamache of Hurricane Research Division/ NOAA): First, the horizontal wind
components were calculated assuming that the vertical component is zero. Then, the horizontal
components were used to calculate the first guess of the divergence field. Starting from the lowest
level, the following procedure was repeated for each level until a prescribed accuracy was attained
or until 50 iterations had been done. The vertical wind component for the first iteration was
I__I__(1_Ylllli~_lIIILIL
II~--
calculated from the anelastic continuity equation using vertical wind velocities from the level
below, and the first guess estimate of divergence at the current level. Then, the three wind
components were adjusted using the Least Squares method. The solution using this method usually
converged to the desired accuracy in less than 50 iterations.
The data from separate flight legs had to be merged in a suitable way for the analysis of
the whole MCS. A correction was made for the movement of the MCS by using a translation
velocity, estimated by tracking the vortex center from one flight to the next, to move each data
point to an appropriate position at some reference time. If more than one datapoint was moved to
the same gridvolume, an average was calculated. The data were mapped to a 5 x 5 x 0.5 km grid,
0.5 km being the vertical resolution.
Sea clutter can affect the data far from the aircraft, and when the antenna is pointing
downward. When the aircraft flies at 3 km altitude and the antenna is pointed horizontally, the
mainlobe of the beam will touch the sea surface at 90 km range. The problem
of sea clutter
contamination becomes worse when the antenna is pointed downward. For example, when data are
collected from an altitude of 1.5 km above the sea surface and the flight altitude is still 3 km, the
mainlobe touches the sea surface as close as 45 km from the aircraft. The Doppler velocities in the
case studied were noisy below 2 km altitude. Where data from below 2 km was used in the
analysis, care was taken that any suspicious looking winds were deleted. In addition to sea clutter
problem, errors in the in the measured ground speed of the aircraft and errors in the estimated
precipitation fallspeed can affect the wind data.
Airborne radars have additional error sources compared to ground based radars. The error
in the antenna pointing angles relative to the aircraft are smaller than 0.50, and can be accounted
for. The antenna position with respect to the ground is calculated using information of the aircraft
altitude, given by the Inertial Navigation System. The aircraft altitude angles have errors less than
0.50. The location and the velocity of the aircraft are retrieved by integrating the accelerations
given by the Inertial Navigation System. The ground velocity of the aircraft is subtracted from the
Doppler velocities to eliminate the velocity component that is owing to the movement of the
Doppler radar itself. Therefore, errors in the ground velocity will introduce errors in Doppler
winds.
The error in the measured ground speed is about 2-3 m/s. The error introduced to Doppler
velocities owing to an error in the ground speed is constant with height. Therefore, vertical
differences of wind are not affected. Assuming an error of 1 m/s in the terminal fallspeed (see Atlas
et al., 1973), the associated error in the horizontal wind speed can be estimated. The errors are less
than 1 m/s for horizontal distances of more than 4 km from the flight track, assuming that no data
farther than 4 km above or below the aircraft is used in the analysis. Moreover, for straight flight
tracks this error is perpendicular to the flight track and shows up as spurious divergence or
convergence.
2.2 In situ data
Intercomparison of instruments onboard NOAA WP-3D and the NCAR Electra was made
using data from two sets of intercomparison flights in the beginning and at the end of the field
experiment. The differences of the temperature and the dewpoint temperature measures by the two
aircrafts were less than 0.3 K during both flights, and they were accounted for in the data analysis,
by interpolating the differences in time and subtracting them from the data of the other aircraft.
Data were excluded from the analysis if the magnitude of the vertical velocity exceeded 1 m/s, in
order to minimize the effect of active convective updrafts and downdrafts on the analyzed fields.
Data were also excluded if the measured dewpoint temperature exceeded the measured
temperature. However, no data were excluded from the 300 m analyses. Using the same method as
with the Doppler data, the data were renavigated to the appropriate locations at a given time, and
80 s (10 km) averages were calculated. These averages were then analyzed by hand.
2.3 Satellite data
The Geostationary Operational Environmental Satellite (GOES-7) provided us with
infrared and visible images during the TEXMEX project. The imagery has a horizontal resolution
of 8 km and covers Central and North America from 70 to 550N and from 600 to 140'W. In this
study we only use the infrared images. Hourly image data for the period of the TEXMEX project
are interpolated to 0.10 x 0.10 area elements and converted to infrared brightness temperatures.
The satellite imagery is used to track the disturbance that developed into hurricane Fefa. Particular
attention is given to the location of convective regions.
2.4 NCEP/NCAR Reanalysis data
This recently released dataset is described in detail by Kalnay et al. (1996). In this study,
we only used the wind fields. All the data have a temporal spacing of 6 hours (0000, 0600, 1200
and 1800 UTC) and are put on a 2.50 lat x 2.50 long grid. The dataset has 17 mandatory levels
ranging from 1000 to 100 hPa.
Chapter 3
Observations of the synoptic-scale circulation
The tropical easterly wave that developed into Fefa was well observed. The wave first
emerged from the northwest coast of Africa on 17 July. It then propagated across the Atlantic
Ocean and the Caribbean Sea in the next eight days without development, and finally emerged over
the eastern Pacific Ocean on 25 July, as indicated by the Balboa, Mexico, rawinsonde data
(Rappaport and Mayfield 1992).
The location and propagation of the wave in the NCEP/NCAR re-analysis is shown in
Fig. 3.1, which shows the wind vectors and isopleths of the relative vorticity at 700 mb. At 00
UTC 20 July (Fig. 3. la) the wave axis, which is defined by the location of the line of wind shift or
vorticity maximum, was near 380W, with the axis tilted in a northeast-southwest direction. The
0
latitudinal extent of the wave to the south seems to reach 10 N or even further south. The relative
vorticity contours show that there is a maximum of cyclonic vorticity located around the axis and it
is flanked by regions of anticyclonic vorticity to the east and west. The wave had a similar position
and structure at 500 mb and 850 mb (not shown), but it is better defined at the 700 mb level.
The wave propagated steadily westward at a speed of about 70 longitude per day. At 00
UTC 22 July the axis was near 50 0W (Fig. 3. lb). On the follwing day the wave propagation seems
to accelerate and the wave axis reached 64°W at 00 UTC 23 July. At this time the relative vorticity
maximum becomes less obvious (Fig. 3. Ic). The relative vorticity maximum at 850 mb disappears
and the winds in the eastern Caribbean at that level were predominantly easterlies (not shown). At
0
00 UTC 24 July the wave axis past south of Haiti and reached 75 W. The wave became better
defined again as seen in the relative vorticity isopleths (Fig. 3.1d). The wave continued to
propagate westward and 24 hours later the axis was near 80'W, very close to the mountains of
er~~L-l-i^l-lp~
^cu*y^-iur*~- ~arr.
20W
20
b)
30N
27N -
/
+
24Ne-
-2e
21N
1V8N
r
5
12N-
-5e
e 0
1.5e-05
-
9N
11W
5er0 Z
' --
N
.0
eu
15N
6N
3N
05
2e-055
.
...
7%0
%
1e0
5e5e06
5e-06
-
1e-05
- 5e-06
1e-05
e 5--
100W
90W
80W
70W
60W
50W
40W
30W
2
20
Figure 3.1: Winds and relative vorticity from the NCEP/NCAR re-analysis at 700 mb and 0000
UTC for (a) 20 July, (b) 22 July, (c) 23 July, (d) 24 July, (e) 25 July, (f) 26 July, (g) 27 July, (h)
28 July. The reference vector is in m/s.
30N
27N
24N
21 N
18N
15N
12N
90W
-100W
lW
70W
80W
60W
30W
40W
50W
20W
20
d)
30N-
1.5e05--5
100W
V
-2e-05
.,
e ---ee0 -2 5-.-5
,5
..
I-
219N
6N
-e 18N -
u
e
-le-05
5
1
EeW
- 5
2,e,,
2
27N
0 e-
-e-065
~
90W
.5e-05
70W
80W
60W
50W
40W
20
Figure 3.1 (cont'd)
Figure 3.1 (cont'd)
le-
41
30W
201
-^- L.
.-.--~--l~d~--isXI(sII~
I
I*W1_4-.^-.--
20W
20
f)
30N
27N
--
'\\
\-
24N
21N
18N
\ --..,-. _,- ,le:,
- ----------------
15N
12N9N6N-
/
"
Se-05
1 -
1e-05,
e--
•
3N_.
EQ-
12(3W 115W
110W
105W
100W
95W
.
90W
85W
80W
75W
70W
20
Figure 3. 1 (cont'd)
___
_Ilpn
10
h)
30N
27N 24N
21N 18N15N
12N
9N
6N
3N EQ
120W 115V
50W
10
Figure 3.1 (cont'd)
__l~~Xli___
-ICXII.~~F-L---.~I_~I^*IYIIIC-~I~-
--
i~-
Y--.------~-I-~I~Y~~II~LIUI~Y~-~I~ipll
L. _-XI_1~~
1
Central America (Fig. 3.le). The relative vorticity isopleths show a maximum more to the
southeast, near Panama. Fig. 3.1d also shows that on 24 July there is wavelike structure located at
50N, 990 W. The feature seems to retrograde eastward and reach 90'W at 00 UTC 26 July. It is
uncertain whether that feature might be an artifact of the analysis, and whether it plays any
significant role in the genesis of Fefa. At the same time, the easterly wave was making its passage
across Central America, emerging into the eastern Pacific Ocean. The region of maximum relative
vorticity remains over Central America on 26 and 27 July, while the wave axis entered the eastern
Pacific. At 00 UTC 28 July, the wave axis reached 104 0W. A relative vorticity maximum appeared
over the eastern Pacific near 103 0 W which is distinguished from the one over Central America
(Fig. 3.1h). An examination of the 6 hourly fields reveals that the relative vorticity maximum
seems to have developed in situ and emerged from very low latitudes. This maximum remained
relatively weak until 29 July, when it became more prominent and was located at the genesis region
of Fefa.
During the period in which the wave was propagating across Central America, there is no
eveidence of any organized low level circulation over the eastern Pacific. Fig. 3.2a shows the winds
and isopleths of relative vorticity in the NCEP/NCAR re-analysis at 1000 mb at 00 UTC 26 July.
Vorticity over the eastern Pacific was generally near zero or slightly negative. There is an area of
cyclonic vorticity in the southern Caribbean near Panama. This vorticity maximum is a
climatological feature of that region (eg. Nitta and Takayabu 1985), and it persisted throughout the
period of our interest (Fig. 3.2 a-d).
One day later, at 00 UTC 27 July, the easterly wave was already in the eastern Pacific, as
seen in the 700 mb analysis. But at 1000 mb there still is no evidence of closed circulation over the
Pacific (Fig. 3.2b). We can see the ITCZ located along 90N, extending from the Central America
to about 100°W as an axis of deformation. The signature of the ITCZ is also evident at 925 mb
24N
2130N
1 e-05
18N 15N -
<--
,1
le
e
2e-05
j
3N-
e 0
..
120W
115W
110W
105W
100W
95W
90W
85W
80W
75W
70W
65W
60W
10
b)
30N
27N
-
24N
1N
. . .
'
W
12W
UTC
18N)
for
(a) 26 July,
e5e-0590W-9e-05
---W
7e-05" e1005
W
85W
75W
27 July, (c) 28 July, (d) 29 July.
-0 5
N -e
EQ
12)W
85W-8--
115W
110W
105W
100W
95W
90W
80W
70W
65W
60W
10
Figure 3.2: Winds and relative vorticity from the NCEP/NCAR re-analysis at 1000 mb and 0000
55W
~-~- -r;-i-~-r-
-n~-~~---sl--~..
Irxl-i~------------)I i-llrrriua
--l.i--
27N
S'---2e-05
24N.
21N
l
e-05"
7/
18N
15N
.
. . ..
zg
'
_
_
.1--
e-05
9N -
6N
\
3N EQ
120W
115W
110W
-- +A&0
100W
105W
95W
90W
85W
80W
75W
70W
65W
60W
55W
50W
10
d)
27N
V,
N
S-......
24N -e
le-05
--1e-05
21N-
0
18N
12N -
-
-le-05
1e
9N-
5
e50
1-ee.05
6N 3N\
EQ
120W
115W
110W
105W
100W
95W
90W
85W
80W
75W
70W
65W
10
Figure 3.2 (cont'd)
60W
55W
50W
(not shown). The winds were northeasterlies to the east of the mountains of Central America, and
the wind speed reached over 10 m/s.
The first evidence of a region of cyclonic vorticity at low levels in the eastern Pacific
occurred at 00 UTC 28 July (Fig. 3.2c). A distinct area of positive relative vorticity was located at
90N, 960W. The maximum of vorticity had a magnitude in excess of 2 x 105' s '. During the
formation of this vorticity maximum, the 700 mb wave axis has already reached 104 0W. An
important contribution to the formation of the area of cyclonic vorticity is provided by the
development of strong easterly winds to the south of the Gulf of Tehuantepec, which were absent a
day earlier. The maximum speed of the easterly winds reached 8 m/s. This easterly jet was also
observed at the 925 mb level (not shown), and was present since 12 UTC 27 July. An explanation
of the formation of this easterly jet was provided by Mozer and Zehnder (1996). According to their
argument, when the winds to the east of the mountains of Central America are from an easterly
direction, they will be incident on the mountains. This results in a blocking of the flow at low levels
and a diversion of the flow to the south. They then argue that by the conservation of the Ertel
potential vorticity this blocking causes an increase in magnitude of the easterly winds in the lee of
the mountains ie. over the eastern Pacific. In a study of the genesis of Hurricane Guillermo, Farfan
and Zehnder (1997) also observed the prescence of strong easterly winds south of Mexico and west
of Central America, preceding the formation of the storm. But one interesting difference is that in
their case the 700 mb wave was propagating northwestward and was still located in the western
Caribbean when the strong easterlies in the eastern Pacific formed, while in this case the wave has
entered the eastern Pacific. It is also interesting to note that approximately one day prior to the
formation of the strong easterly winds south of Mexico, the speed of the winds incident on the
central American mountains increased, which is most notable at the 925 mb level.
LL
-IULIULe~lBsI*~I~-3L16.
-.al~-LICV"
*~
*^r.rC~lllllrPPYLP-UYI~Y"IIBLIC-YI-
The area of maximum vorticity had not increased significantly in its horizontal extent by
00 UTC 29 July (Fig. 3.2d). Note that the southerly winds to the south of the region has increased
in magnitude, indicating the occurrence of a wind surge. In the region of positive relative vorticity,
the easterly winds were replaced by southerly winds, and a cyclonic curvature of the flow can
clearly be seen. The wind surge is also evident at the 925 mb level (not shown), occurring at about
the same time. It is suggested that the formation and growth of the low level cyclonic circulation
can be associated with to the development of the easterly jet to the north followed by a southerly
wind surge later.
The development of the mesoscale system was investigated using satellite imagery. At 00
UTC on 25 July (Fig. 3.3a) , there was an absence of convection in a large region south of the Gulf
of Tehuantepec. Any convective activity was confined to the west or to latitudes south of 100N,
and was rather scattered. At 01 UTC on 26 July (Fig. 3.3b), more convective activity is seen in the
eastern Pacific, though it was still disorganized. The area of convection extended from about
110 0W to the the coast of Central America, and was still confined to south of 100 N. Note that
there is a train of clouds oriented in a NNE-SSW direction in the Caribbean Sea north of
Honduras. This is the time when the wave axis was crossing Central America. 12 hours later (Fig.
3.3c), convective activity in the eastern Pacific decreased somewhat, and the train of clouds in the
Caribbean had lost much of its signature. At 12 UTC on 27 July (Fig. 3.3d), convective activity in
the eastern Pacific rejuvenated. There were several regions of moderately intense convection, some
of which were located to the north of 100N. On the next day (Fig. 3.3e), convection became more
organized, and was greater in areal extent. A concentrated area of convection is first seen at 12
UTC on 29 July (Fig. 3.3f). It was located near 120N, 109 0W, and corresponds quite well with the
location of the circulation center derived from aircraft observations. At that time the system had
developed into a tropical depression. When Fefa reached tropical storm strength at 01 UTC on
b)
Figure 3.3: Infrared images from GOES at (a) 0000 UTC 25 July, (b) 0100 UTC 26 July, (c)
1200 UTC 26 July, (d) 1200 UTC 27 July, (e) 0000 UTC 28 July, (f) 1200 UTC 29 July, (g) 0100
UTC 30 July.
d)
Figure 3.3 (cont'd)
f)
Figure 3.3 (cont'd)
Figure 3.3 (cont'd)
30 July (Fig. 3.3g), the convection associated with the circulation center became more intense and
its area increased. Note that there is another small area of convection to the northwest of the major
convection area. That small area of convection might be associated with the low-level circulation
center, which was displaced to the northwest at that time.
Burpee (1972) noted the existence of a sign reversal of the meridional gradient of potential
vorticity (qy) over Africa, which satisfies the Charney-Stem necessary condition for instability of
internal jets. Molinari et al. (1997) extended the geographical location under observations of this
kind of PV structure and discussed the relationship between the fluctuation of the Caribbean and
eastern Pacific sign reversals of the Ertel potential vorticity and eastern Pacific tropical
cyclogenesis. In a later study, Molinari et al. (1999) show that in the period preceding the genesis
30N
27N
24N
21N
18N
15N
12N
9N
6N
b)
30N
24N
15N
21N18N15N 12N- .
05e055e-05
4e
6e-05
I
" 3e-05
...
N
6N
-
1e-05
2e-Uzu3e-z05
e2e-05
e-05
3N -
EQ
120W
C
4e
-e-05
110W
100W
90W
80W
70W
60W
50W
40W
30W
Figure 3.4: Absolute vorticity from NCEP/NCAR re-analysis at 700 mb and 0000 UTC for (a) 20
July, (b) 21 July, (c) 22 July, (d) 23 July, (e) 24 July, (f) 25 July.
30N
27N 24N
21N
3e-05
18N
15N 12N - 4e-05
9N
.
e-.,0
6N 3N
EQ
120W
110W
d)
30N
27N
0
24N
21N -<4e-05
18N -3e05
15N -5e12N
9N
6N
3N EQ
110W
100W
Figure 3.4 (cont'd)
f)
30N
27N
/
21N
18N
12N
9N
6N
3N
EQ
.
3
o
Figure 3.4 (cont'd)
of Hurricane Hernan of 1996, the easterly wave moved through a background state which had a
strong sign reversal in the meridional gradient of the absolute vorticity, which is favorable for the
maintenance and growth of the wave. It is found that such a favorable background state also
existed before the genesis of Fefa. Fig. 3.4 a-f shows the isopleths of the absolute vorticity at
700 mb. There is clear evidence of a sign reversal in the western Caribbean Sea as early as 20
July, when the easterly wave was still in the eastern Atlantic. The strength of the sign reversal
increased somewhat over the next three days, while the easterly wave was propagating into the
Caribbean Sea. On 24 and 25 July, when the wave was propagating across the western Caribbean,
the strength of the sign reversal decreased. In the eastern Pacific, on the other hand, there is no
clear evidence of a sign reversal. In conclusion, these analyses suggest that the easterly wave
moved through an unstable basic state in the Caribbean Sea.
49
50
Chapter 4
The target object of IOP4 of TEXMEX was the MCS that eventually developed into
Hurricane Fefa. This IOP was conducted from 28 July 1991 to 29 July 1991. Altogether three
flights were flown. The first and the third flights were flown with the NOAA WP-3D, and were
labeled 1P and 3P, respectively. The second flight was flown with the NCAR Electra, and was
labeled 2E. Flights 1P and 2E consisted of several flight legs at both 3 km and 300 m altitudes,
while flight 3P consisted of flight legs at 3 kmn and 450 m altitudes. In the following, a summary of
the observations made during these flights is presented.
4.1 Flight 1P
This is the first flight into the pre-Fefa system. Fig. 4. l1a shows the Doppler radar wind at
3 km altitude. A mesoscale circulation pattern can be seen. Fig. 4. lb shows the in situ observations
of the wind at the same level. Combining these two figures, a vorticity center is estimated to be
located near 10.3' N, 105.90 W. Fig. 4.1c shows the Doppler radar wind at 1 km altitude. No
closed circulation was found at this altitude. There was a WSW-ENE oriented shear line between
100 and 110 N and 1070 and 1060 W, which was to the west of the 700 hPa (3 km) vorticity center.
The winds south of this shear line are generally SW'ly, and are stronger than the winds at higher
levels. The in situ observations of the wind at 300 m (Fig. 4. ld), which were collected at a later
time, only captured the SW portion of the 700 hPa vorticity center, showing SW'ly winds of 15 20 m/s. The flight pattern also captured the shear line to the west of the 700 hPa vorticity center,
which had moved northward by about 80 latitude since the 3 kman altitude flight. Assuming that the
vortex is in balance with the thermal field, the vertical shear of the wind between two levels gives a
proxy for the temperature anomalies in the corresponding layer.
-,07
- 16
-105
b)
12
Figure 4.1: Observations in pre-Fefa MCS during flight 1P. (a) Doppler wind field at 3 kmn, (b)in
situ winds at 3 km. Long barb is 5 m/s.
c)
-=L
-7
j
\yJJJJ
\~
J JJ
-
j
-/JJ
J J
j
J =
d)
Figure 4.1 (cont'd): (c) Doppler wind field at 1 kIan, (d) in situ winds at 300 m.
1
1 2
4
S
7
8
9
10 11 12 1 1
i
15
8 17 t8 1920 21 22
3 24 29 2
27 2
336
so
80
339
9
,g
II
iMo
10--3
c
167.
---- ------
aGI
Figure 4.1 (cont'd): (e) change of wind from 1 to 3 kan, (f) time series of virtual potential
temperature on leg 2, (g) relative humidity at 3 kmn in percents, (h) ee at 3 klan. Temperature in
Kelvins.
21 30 31
4
i
00
-1o
2
3
4
S
2
12
3
3
4
4
5
5
1
8
7
8
5
8
6
7
7
8
3
I
9
10 11
12 13
14 110
17
1
21
22 23
24 2
28
10
15
20
145 1145
314+
312S5
1
13J14
1I
10 11 12 13 14
I
I
iI
11
19 16 "17 t8 12 20
I
21 22 23
24 25 20
Figure 4.1 (cont'd): (j) time series of Oe on leg 4, (k) time series of in situ winds on leg 4, (1)time
series of virtual potential temperature on leg 4, (m) map showing leg 2 and leg 4.
25
If the wind shear is cyclonic (anticyclonic), then the vortex has a cold (warm) core. Fig. 4. le shows
the vertical difference between the wind at 3 km and 1 km. There is a region of cyclonic shear near
10.80 N, 105.80 W, suggesting a cold core located to the north, and not very far away from the
vortex. Fig. 4. If shows the time series of the virtual potential temperature recorded on leg 2 of the
3 km flight pattern, as indicated in Fig. 4. l1b. The conspicuous trough in the curve corresponds to
the cold core. However, there is also a region of anticyclonic shear near 10.80 N, 105.60 W,
suggesting a warm core located to the NE of the vorticity center. The relative humidity at 3 km
altitude is quite high in the region of the vortex core (Fig. 4. 1g). It varies between 80% and 100%.
The relative humidity in the environment is only about 70%. The analysis of the equivalent
potential temperature (0e) at 3 km (Fig. 4.1.h) shows a maximum near the vortex core, collocated
with the region of highest relative humidity. Fig. 4.1 .j shows the time series of Oe recorded in leg 4
of the 3 km flight pattern. This leg runs directly through the vortex center, indicated by the shift of
wind (Fig. 4.1k). There is an obvious peak of ee (342.5K) at the vortex center. Note that the ee in
the environment of the MCS is only about 335K, which is about 6K lower. The analysis of the
virtual potential temperature roughly agrees with the Fig. 4.1e. It does not show a distinct cold
anomaly in the vortex core. This is further illustrated in the time series curve of the virtual
potential temperature in leg 4 (Fig. 41).
The analyses of the ee and virtual potential temperature in the boundary layer (1 km
altitude) are shown in Fig. 4.2a and Fig. 4.2b, respectively. Although the flight pattern missed the
region of the 700 hPa vortex center, the flight leg closest to the vortex center does show that both
variables have negative anomalies near that region. So it suggests that in the region of the 700 hPa
vortex, a cold core exists at low levels. In the region of the trough or the shear line, however, both
variables have relatively high values, though the winds in that region are no stronger than those
c)
d)
14ti
-
302
30
~c~Sol ~
ro~l-----
44.
42
it0
-
301
300.
301
2"2.
26.5
041
-
2
3
4
S
7
0
O10 11 12 13 14 i5 to 17 1
t 20 2 2
223
24 25 2
27 21
30
.
-
.
---
230
205
1
-
--1 2
4
3
4
6
7
9 0
10 11 12 11 IA4
1 1617 IS
I0 20 21 22 23 243
Figure 4.2: Observations in pre-Fefa MCS during flight 1P. (a) Oe at 300 m, (b) virtual potential
temperature at 300 m, (c) time series of e on leg 9, (d)time series of virtual potential temperature
on leg 9.
25 27 28 20 30
02
-/-
0
Figure 4.2 (cont'd): (e) map showing leg 9.
near the vortex core. This is clearly illustrated in the time series of these two variables in leg 10
(Fig. 4.2c, 4.2d). Both variables increase substantially in the region of the trough.
Flight 2E
Whereas in the first flight no distinct vortex can be seen at 3 km altitude, during flight 2E
a vortex center became obvious. Fig. 4.3a indicates a vortex center located at 12.20 N, 109.30 W.
The height of the 700 hPa surface shows a minimum at the vortex center (Fig. 4.3b). Compared to
the first flight, the altitude of the 700 hPa height minimum has decreased by about 20 m. At 3 kan
altitude, the relative humidity generally varies between 80% and 90% near the vortex center (Fig.
4.3c). But note that to the north of the vortex center, there is a small region where the relative
humidity is rather low (<80%), indicating the presence of some dry air. This is illustrated in the
time series curve of the relative humidity in leg 4 of the flight pattern (Fig. 4.3d). Note the drop in
~0
I0
C
3j3
7 .....
12
-
+°
,;-
1,3~a
\
4------p-
i
1
/
3/60
.44
-4
C
I,
.1low,I
1 0
Figure 4.3: Observations in pre-Fefa MCS during flight 2E. (a) in situ winds at 3 km, (b) altitude
of 700 hPa surface (m), (c) relative humidity at 3 km in percents.
I
1
--
2
-
3
4
5
-
6
7
8
9
10
1
11
14
15
18
17
18
16
f)
10-
f
I/
S
Sr
-15
-10
-
0
5
10
I5
20
25
30
----
---------------
Figure 4.3 (cont'd): (d) time series of relative humidity on leg 4, (e) time series of in situ winds on
leg 4, (f) virtual potential temperature at 3 kin.
60
313
302 j
X 314
t
301
...........
300
.........
312
1 2 3 4 5 6 7
8
21 22 23 2 25 20 27 25 29 30 31 32 33
1
4
2
9 10 11 12 13 14 15 15 17 1o 1
1
3
4
5
8
7
3
9
2
3
14
17
18
Figure 4.3 (cont'd): (g) time series of virtual potential temperature on leg 2, (h) time series of
virtual potential temperature on leg 6, (j) map showing leg 2, leg 4 in the 3 km altitude flight
pattern and leg 6 in the 300 m altitude flight pattern.
61
20
1
TO 11 12 13 14 15 It 17 15 19 l 2071
22
2
2223
24 2
22
relative humidity just to the north of the vortex center, which is indicated by the point of nearly
calm wind (Fig. 4.3e).
During this flight, we also notice that a lower tropospheric warm core has gradually
developed inside the cold core. This is manifest in the reversal of the temperature gradient (Fig.
4.3f). This reversal if gradient is further illustrated in the time series plot of the virtual potential
temperature in leg 2 (Fig. 4.3g). In the boundary layer, due to the less than optimal flight pattern,
we cannot get detailed measurements of the virtual potential temperature around the vortex center.
But there still seems to be a reversal of temperature gradient in the southern side of the 700 hPa
vortex center, as indicated in the time series plot for this variable in leg 6 (Fig. 4.3h). Thus it may
be concluded that there is a small warm core inside a cold core at both levels. The analysis of Ge at
3 km altitude shows an increase of about 2K near the vortex center, whereas in the boundary layer,
the Ge has increased substantially (by - 5K). Note that there are two regions of low level Oe
maxima. One is located to the east of the 700 hPa vortex center, where the recorded wind speed is
largest (25 m/s). Another is located about 60 km to the SW of the vortex center, and the winds
there are only moderately strong (15 m/s). In the boundary layer, no closed circulation was found.
But since the flight leg is oriented straight (in a SW-NE direction) and passed to the SE of the 700
hPa vortex, it was probably not sufficient to resolve a closed circulation.
Flight 3P
During this flight the system has already intensified into a tropical storm named Fefa.
Doppler radar observations at 2 km altitude show an obvious circulation wind pattern (Fig. 4.4a).
The 700 hPa vortex center is located near 12.70 N, 1110 W. The maximum wind at 1 km altitude
(not shown) was about 30 m/s. The altitude of the 700 hPa height minimum decreases further by
about 15 m from flight 2E. The analyses of Ge and virtual potential temperature at 3 km alititude
13
P-1
(1
b)
'-7
7-
Figure 4.4: Observations in Tropical Storm Fefa during flight 3P. (a) Doppler wind field at 2 km,
(b) in situ winds at 3 lan.
I~~~ --~-~~-^--~-- ---
-~-1.~1...~.1.~-
338
3*f
33%
3939
\338
333.
3 336
3f
33f
3/F
317
3/f
314
j 17
3/jI___________
3/3
'/
3/6
_____
3/1
3,'
3/5
3/
-1I9
Figure 4.4 (cont'd): (c) Oe at 3 km, (d) virtual potential temperature at 3 km.
__YIIII---^~^~_ -~IW~YCi
~jJ L~~__~_J~~
e)
f)
31
-
330
31*
2
1
4
3
7
5
6
9
10
11
12
13
14
15
1I
17
1
1
2
3
5
4
8
7
$
10
0
1i
12
11
14 15
17
16
18
h)
g)
31-
l
10-
4
1
?\\\
-15o
-10
---
1
-
-1
4C 314
-5
5
1
2
3
4
s
0
7
6
9
10
11
12
13
14
15
I
time series of virtual potential temperature
Figure 4.4 (cont'd): (e) time series of ee on leg 2, (f)
on leg 2, (g) time series of in situ winds on leg 2, (h) time series of virtual potential temperature on
leg 9.
17
11
..~-^_.i~_~--i-~~.
20-
/
23--
--
-5'
-10
1
10
25
0
----
I
30f
11
-
i-
' ~
30
--
Figure 4.4 (cont'd): (j)time series of in situ winds on leg 9, (k) in situ winds at 450 m, (1)virtual
potential temperature at 450 m, (m) map showing leg 2 and leg 9.
(Fig. 4.4 c and d respectively) do not show a maximum at the circulation center, indicated by the in
situ winds (Fig. 4.4b). This is further shown in the time series plots for these two variables in leg 2
(Fig. 4.4e, 4.4f). This leg passed very close to the circulation center, which is indicated by the shift
of wind from southerly to northerly (Fig. 4.4g). The largest values of these two variables are
instead found about 60 km northwest of the center. Near the 700 hPa circulation center, the virtual
potential temperature is relatively low. Fig. 4.4h shows the time series of the virtual potential
temperature on leg 9. The circulation center is indicated by a shift of wind from southerly to eastsoutheasterly (Fig. 4.4j). The maximum value of the virtual potential tern erature was recorded
well after the aircraft had passed the vortex center. The time series of ee has a similar trend (not
shown).
When we look at the winds measured during the 450 m altitude flight pattern (Fig. 4.4k),
we find that the circulation center is located near 13.30 N, 111.70 W. Therefore, the vortex in the
boundary layer is not vertically aligned with the 700 hPa vortex. It is displaced to the northwest,
which might indicate a tilting of the vortex. Since the 700 hPa vortex was detected only 45 minutes
earlier, the translation of the system between the flight at these two levels cannot account for the
vortex displacement. It is interesting to look at the vertical wind shear to verify whether the shear
might have caused a tilting of the vortex. Vertical displacements of the vortex of tropical cyclones
in the Western Pacific have been observed using reconnaissance aircraft data (Huntley and Diercks
1981). They showed that during their development stage many tropical cyclones have a significant
tilt in their vertical axes, and that the vertical tilt was highly correlated with the intensity and
direction of vertical wind shear. In this case, the magnitude and direction of the 700-1000 mb wind
shear is estimated using the NCEP/NCAR re-analysis data. At 00 UTC 30 July, when the vertical
displacement of the vortex was observed by the P-3 aircraft, the 700-1000 mb wind shear in the
region of the tropical storm was found to be east-southeasterly, and has a magnitude of about
7 m/s. Since the 700 mb circulation center was located to the southeast of the surface circulation
center, there seems to be no correlation between the direction of the wind shear and the direction of
the tilt. In fact, their directions are nearly opposite. This observation suggests that the vortex at the
boundary layer is unlikely to be a direct result of a downward extension of the mid level vortex. It
is conjectured that the surface circulation spun up rather independently at the beginning, as a result
of the processes mentioned in the previous section. At a later stage, the mid-level vortex may have
interacted with the low-level vortex, resulting in a more cohesive vortex and further intensification
of the storm.
The displacement of the low level vortex might explain the large values of ee and virtual
potential temperature found to the northwest of the 700 hPa center, because convection may be
enhanced near the low-level vortex center. In fact, the largest values of these two variables are
found near the 450 m circulation center. The relatively low Oe at the 700 hPa center is probably
due to the penetration of dry air into the core (Raymond et al., 1998). In the boundary layer, the
virtual potential temperature has increased by about 2 degrees K since the previous flight (Fig.
4.41).
In order to get a rough estimate of the time evolution of thermodynamic variables in the
region of the vortex, averages of the variables were calculated in a box. The box was generally
chosen to be centered on the vortex, whose position was estimated by the wind field. For flight 1P
at 3 km altitude, the box was chosen to be a square one of roughly 140 km width. At 300 m, it was
chosen to be a rectangle in order to capture the values along the shear line. For flight 2E, the box at
both levels was a square one of about 130 km width. Since the 300 m flight pattern could not
resolve any vortex, the vortex center was assumed to be at the same location as the 700 hPa vortex.
For flight 3P at 3 km altitude, the box was 120 km in the meridional direction and 130 km in the
zonal direction. At 450 m, it was a square one of about 110 km width. For this flight, both boxes
were centered on the vortex at that level. Note that the time difference between each pair of
consecutive flights is 14 hours. The boxes for which the averages were calculated are shown in Fig.
4. lb, 4. Id, 4.3a, 4.4b and 4.4k. The results are shown in the table below.
Flight 1P
Flight 2E
Flight 3P
RH,3 km
91
86
90
RH,300 m
94
91
90
Qe, 3 km
340
342
340
ee, 300 m
342
345
347
ev, 3 km
315
316
315
ev,450 m
301
302
304
Table 4.1: Averages of in situ data in a box around the vortex from the two
flight altitudes for flights IP, 2E and 3P. Relative humidity in percent,
temperature variables in Kelvins
Between flights IP and 2E, the relative humidity at both levels has decreased. This might
indicate some intrusion of dry air into the system. The ee at 3 km increases slightly by 2 K, while
in the boundary layer it increases by 5 K. Note that between these flights there is a transition from
a low level cold core stage to the development of a small warm core inside the cold core. The
virtual potential temperature changes little at both levels between these flights.
Between flight 2E and 3P, the system has intensified into a tropical storm. Note that the
relative humidity has reincreased at 3 km, but stays rather constant in the boundary layer. Both the
Oe and virtual potential temperature in the boundary layer increase by 2-3 K. But at 3 km altitude,
the values of these two variables dropped slightly. Again, the displacement of the low level vortex
away from the 700 hPa vortex might account for these results, as mentioned earlier.
4.2 Comparison to the genesis of Guillermo
Guillermo developed less than a week after the formation of Tropical Storm Fefa. The preGuillermo MCS developed while an easterly wave was propagating into the Gulf of Mexico from
the Caribbean Sea. Using flight 1P for Fefa and flight 2P for Guillermo as the starting point for
comparison (The first flight, lE, for Guillermo, was conducted well to the west of the system and
therefore cannot capture it well), the pre-Fefa system took about 28 hours to become a tropical
storm, while the pre-Guillermo system took about 42 hours. Thus Fefa seems superficially to be a
fast developer compared to Guillermo. But note that during the first flight, the low-level circulation
associated with the pre-Fefa system was already strong. During flight 2P for Guillermo, a distinct
vortex center was already observed at 3 km altitude. A closed circulation was also found in the
boundary layer. Whereas for Fefa, there was not an obvious vortex center at 3 km altitude during
flight 1P. Moreover, no closed circulation was found in the boundary layer. Instead, a shear line
was observed. The winds observed to the south of the shear line were in fact stronger than the
maximum winds associated with the pre-Guillermo system's 975 hPa circulation. In terms of the
thermodynamic variables, the relative humidity was very high (-90 %) and the ee has a positive
anomaly at the vorticity center at 3 km altitude for flight IP of pre-Fefa and 2P of pre-Guillermo.
At this level, a negative virtual potential temperature anomaly was found in the core of the preGuillermo system, but that is not found in the pre-Fefa system. In other words, we cannot see a
pronounced cold core at 3 km altitude in the first flight into Fefa.
In the boundary layer, both flights show a negative anomaly in Oe and virtual potential
temperature in the region of the 700 hPa vorticity center, suggesting that this level was occupied by
a cold core in both cases. The pre-Guillermo system has changed little from flight 2P to 3E. We
now try to compare the observations from flight 2E for Fefa with those from flight 4P for
Guillermo. There was an obvious difference in the evolution of the relative humidity. For the preGuillermo system, the relative humidity increased at both levels (by 2% at 3 km altitude and by
7% at 300 m altitude) in flight 4P, indicating a significant moistening of the boundary layer. But
for the pre-Fefa system, it decreased by 3-5% at both levels in flight 2E, indicating some intrusion
of dry air. In both cases, high Oe values were observed near the vortex center at 3 km altitude.
Note that for the pre-Fefa system, there has been a 2 degrees K elevation in the ee at this level
since the first flight. In the boundary layer, the ee increased by about 3 K in both cases. One
similarity of these two systems is that the analysis of virtual potential temperature shows a
reversal of gradient outside the center of the core, indicating that a small warm core has developed
inside a cold core.
We now compare the observations from flight 3P for Fefa and those from flight 5E for
Guillermo, since both systems have just reached tropical storm intensity in these flights. An
important difference is that there is a displacement of the low level vortex to the northwest of the
700 hPa vortex for Fefa. Perhaps as a result, the highest values of ee and virtual potential
temperature at 3 km altitude were located to the northwest of the 700 hPa vortex center. This is in
contrast to Guillermo, where the highest values of these variables were found at the 700 hPa vortex
center. Besides, a reversal of the gradient of the virtual potential temperature can still be seen in
Guillermo, while it can no longer be seen in Fefa. During flight 6P for Guillermo, when it has
already become a hurricane, the relative humidity at 3 km altitude near the vortex core decreases
by about 4%. But for Fefa, it re-increases by about 4% in the last flight. The relative humidity in
the boundary layer, on the other hand, has remained almost constant in both cases. Compared to
the previous flight, the Oe and the virtual potential temperature in the boundary layer have
increased by about 2 degress K in both cases while the tangential wind speed increases. This
suggests that the intensification owing to the feedback mechanism between the wind and the
surface fluxes may have been in operation.
Chapter 5
Summary and Suggestions for Future Work
Observations of the synoptic-scale circulation in chapter 3 suggest that the formation of
Hurricane Fefa was closely associated with the passage of an easterly wave, which was traced by
the portions of PV anomalies in the reanalysis data. The wave propagated westward across the
Caribbean Sea and the Central American mountains, and finally emerged into the eastern Pacific.
The initial circulation at 700 mb formed in the trough region of the wave. Analyses show that there
was a strong sign reversal of the meridional gradient of the absolute vorticity in the Caribbean
while the easterly wave was propagating through that region. It is suggested that this unstable basic
state led to the maintenance and growth of the wave, thus providing a favorable background
condition for the genesis of Fefa.
Reanalysis data suggest that there are at least two phenomena that occurred prior to the
formation of the initial circulation in the boundary layer. The first of these is the development of an
easterly jet to the south of the gulf of Tehuantepec. This easterly jet may have been associated with
the blocking and deflection of the flow around the south end of the Central American mountains, as
suggested by earlier studies (Mozer and Zehnder 1996). The second phenomenon is a surge of
southerly winds into the region of the ITCZ. It is uncertain whether the wind surge led to an
increase in the relative vorticity in the monsoon trough, or that it simply advected the trough further
north. It is suggested that both the easterly jet and the southerly wind surge might have played a
role in the formation of the low-level circulation, though it is uncertain how important a role they
have played.
Thermodynamic measurements from aircrafts presented in chapter 4 show that during the
early stage, i.e. the first flight into the pre-Fefa system, there was a positive anomaly of relative
humidity and ee in the vicinity of the vortex at 3 km altitude. But the analysis of virtual potential
temperature does not show a distinct negative anomaly in the vortex core at this level, though a
cold core exists in the boundary layer. This is in contrast to the pre-Guillermo system (investigated
in another TEXMEX case study), whose cold core is pronounced in both levels. Another feature
which differentiates the Fefa case from the Guillermo case is that for the former case, there was a
shear line located to the west of the 700 mb vortex at low levels. Observations from subsequent
flights indicate an intrusion of dry air into the tropical depression, resulting in a decrease of Oe at
the 700 mb level. On flight 3P, it is observed that there is a displacement of the 700 mb vortex
from the boundary layer vortex. An analysis of the vertical wind shear does not suggest that the
shear had caused the displacement or tilting of the vortex. It is conjectured that the surface
circulation formed from the spin-up of the low-level trough, rather than from a downward
extension of the mid-level vortex. An interaction between the mid-level vortex and the low-level
vortex might have resulted in a more cohesive vortex and further intensification of the storm. Thus
this study suggests that the downward extension of the mid-level vortex into the boundary layer is
not a necessary condition for the development of a warm core vortex.
One limitation of this work is that we cannot make generalizations for all cases of tropical
cyclogenesis from one case study. From the single case study here, there seem to be several
elements favorable for cyclogensis. These include the movement of an easterly wave through an
unstable basic state; flow across the Central American mountains which results in an easterly jet in
the eastern Pacific; the pre-existing monsoon trough; and a wind surge of uncertain origin. There is
a possibilty that some elements are essential to the cyclogenesis process while some others are
incidental or only play a subsidiary role. One future direction is to do additional case studies. By
looking at many cases of cyclogenesis one might be able to identify an element that is dominant
over the others. Moreover, one aspect of this work that merits further study is the origin and nature
of the wind surge. Another question that remained unanswered is whether a wave incident upon the
mountains is necessary to create a flow configuration that leads to the formation of an easterly jet
in the Pacific. An alternative direction of research is to use a hierarchy of numerical models to
isolate the individual contributions from each element. The ultimate goal is to determine the
necessary and sufficient condition for tropical cyclogenesis in the eastern Pacific. The results may
also be of interest to other basins where tropical cyclones form, such as the western Pacific and the
Australian regions, as they have some characteristics in common with the eastern Pacific.
__JYY__WII_______1I_~_,I_....
._^__
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