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BRACKEN AND BOSART
353
The Role of Synoptic-Scale Flow during Tropical Cyclogenesis over the
North Atlantic Ocean
W. EDWARD BRACKEN
AND
LANCE F. BOSART
Department of Earth and Atmospheric Sciences, University at Albany, State University of New York, Albany, New York
(Manuscript received 13 February 1998, in final form 22 January 1999)
ABSTRACT
The synoptic-scale flow during tropical cyclogenesis and cyclolysis over the North Atlantic Ocean is investigated using compositing methods. Genesis and lysis are defined using the National Hurricane Center (NHC,
now known as the Tropical Prediction Center) best-track data. Genesis (lysis) occurs when NHC first (last)
identifies and tracks a tropical depression in the final best track dataset. Storm-centered composites are created
with the Analysis of the Tropical Oceanic Lower Level (ATOLL; ;900 hPa) and 200-hPa winds for June–
November produced by NHC for the years 1975–93. Results show that significant regional differences exist in
200-hPa flow during genesis across the Atlantic basin. Composites of genesis in the western part of the basin
show a 200-hPa trough (ridge) located to the west (east) of the ATOLL disturbance. In the eastern half of the
basin composites of genesis show a sprawling 200-hPa ridge centered northeast of the ATOLL disturbance. The
major axis of this elliptically shaped 200-hPa anticyclone extends zonally slightly poleward of the ATOLL level
disturbance. Another composite of relatively rare genesis events that are associated with the equatorward end
of frontal boundaries show that they generally occur in the equatorward entrance region of a jet streak in
conjunction with an ATOLL cyclonic vorticity maximum in a region where vertical shear is minimized.
An approximation of the Sutcliffe–Trenberth form of the quasigeostrophic omega equation is used to estimate
the forcing for vertical motion in the vicinity of developing tropical cyclones. Forcing for ascent is found in all
three genesis composites and is accompanied by a nonzero minimum in vertical shear directly above the ATOLL
cyclonic vorticity maximum. Vertical shear over developing depressions is found to be near 10 m s 21 , suggestive
that weak shear is necessary during tropical cyclogenesis to help force synoptic-scale ascent. Composites of
tropical cyclone lysis show much weaker ATOLL cyclonic vorticity when compared to the genesis composites.
The magnitude of the vertical shear and the forcing for ascent above the lysis ATOLL disturbance are stronger
and weaker, respectively, than in the genesis composites. These differences arise due to the presence of a jetstreak and a longer half-wavelength between the trough and ridge axes in the lysis 200-hPa flow composite.
The genesis flow patterns are decomposed by crudely removing the signature of the developing cyclone and
its associated convection. Two separate and very different flow patterns commonly observed during genesis over
the eastern and western Atlantic Ocean are found to be very similar once the flows are decomposed. Both flows
are characterized by strong deformation at low levels and at 200 hPa with an upper-level jet exit region near
the developing depression.
1. Introduction
Animations of satellite imagery illustrate the ubiquity
of organized cloud clusters (horizontal scale typically
;400–600 km) in the Tropics. Observations of these
cloud clusters (Simpson and Riehl 1981) have shown
that only 1%–2% develop into tropical cyclones. Fundamental atmospheric observations in tropical cyclone
breeding grounds are available from ships, satellites,
aircraft, and land-based stations, but they are too limited
in both space and time to adequately sample the cloud
clusters. Consequently, one of the most difficult tasks
Corresponding author address: Dr. W. Edward Bracken, Department of Earth & Atmospheric Sciences, The University at Albany/
SUNY, 1400 Washington Ave., Room ES-234, Albany, NY 12222.
E-mail: [email protected]
q 2000 American Meteorological Society
for today’s operational tropical meteorologist is forecasting which cloud cluster will be the seedling that
later becomes a tropical cyclone. How can one distinguish between a cloud cluster that will develop and one
that will not develop? What processes influence the transition from cloud cluster to tropical cyclone? This paper
will try to address these questions by focusing upon the
role that synoptic-scale systems play in North Atlantic
tropical cyclogenesis.
The process of tropical cyclogenesis involves a complex interaction between different scales of atmospheric
motion. In general, the Tropics are characterized by a
decrease in entropy of saturated air with height and
consequently are conditionally unstable to undilute parcel ascent. Ooyama (1964) and Charney and Eliassen
(1964) recognized that the circulations of a larger-scale
tropical vortex and embedded cumulus clouds could op-
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MONTHLY WEATHER REVIEW
erate together in a conditionally unstable atmosphere in
such a way as to allow the system to amplify in time.
This idea of system amplification caused by a cooperation between the larger-scale vortex and the cumulative
effects of cumulus clouds was termed conditional instability of the second kind (CISK). Emanuel (1986,
1989, 1991) and Rotunno and Emanuel (1987) take a
different view. They hypothesize that tropical cyclones
are maintained against dissipation by moist convection
whose energy source is supplied entirely by self-induced
anomalous fluxes of moist enthalpy from the sea surface
with almost no contribution from preexisting conditional instability. This type of instability has been
termed wind-induced surface heat exchange (WISHE).
Theoretically, genesis is a continuous series of events,
or more accurately, a process that culminates in the
formation of a self-sustaining vortex. Very generally,
genesis can be thought of as a process that leads to the
formation of a tropospheric deep vortex and enables that
vortex to enter a state of self-development or self-intensification. This intensification of the vortex can be
measured by an increase in interrelated quantities like
vorticity and wind speed or a decrease in central pressure. Some theories hypothesize that this state of selfdevelopment begins when these quantities reach a finite
value (i.e., CISK- or WISHE-like processes begin to
become important). In reality, however, the vortex may
never become fully self-sustaining and the onset of this
instability is not instantaneous. In general, the instability
gradually becomes more important as a warm core forms
within the vortex. It is this internal instability in conjunction with external environmental influences such as
the sea surface temperature (SST) distribution and environmental dynamics and thermodynamics that determine the ultimate intensity of the cyclone. Events leading up to the formation of a persistent, tropospheric deep
vortex will be considered the genesis process. Additionally, before one can claim an understanding of tropical cyclogenesis one must also gain an understanding
of the processes that inhibit genesis and those that cause
tropical cyclone demise (cyclolysis).
Previous research on tropical cyclogenesis (i.e., Gray
1968, 1988) has revealed several necessary conditions
for tropical cyclogenesis. The four most widely agreed
upon are the following: 1) the presence of a synopticscale low-level cyclonic vorticity maximum stronger
than the background planetary vorticity, 2) average oceanic mixed layer temperatures at least 278C, 3) background planetary vorticity greater than 0.8–1.3 3 1025
s21 (the value of the Coriolis parameter 38–58 off the
equator), and 4) vertical shear of the horizontal wind
over the 850–200-hPa layer generally ,10–15 m s21
[Zehr (1991, 1992) states that shear ,12.5 m s21 is
conducive for genesis]. In addition, some studies (e.g.,
Colon and Nightingale 1963; Fett 1966; Gray 1988)
have stressed a fifth condition that involves some external influence acting upon the disturbance to promote
genesis. External influences may include 1) an upper-
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tropospheric anticyclone (or ridge) (e.g., Riehl 1948,
1950; Erickson 1963), 2) an upper-tropospheric cyclone
(or trough) (e.g., Koteswaram and George 1957; Ramage 1959; Erikson 1967; Sadler 1967, 1975, 1976,
1978; Yanai 1968; McBride and Keenan 1982; Davidson
et al. 1990; Bosart and Bartlo 1991; Reilly 1992; Montgomery and Farrell 1993; Molinari et al. 1995; Briegel
and Frank 1997), 3) a lower-tropospheric wind surge
(e.g., Morgan 1965; Fujita et al. 1969; Love 1985a,b;
Lee et al. 1989; Zehr 1991; Briegel and Frank 1997),
and 4) an upper-tropospheric jet streak (e.g., Bracken
and Bosart 1997).
The initial cyclonic vorticity maximum that becomes
the tropical cyclone can be very broadly placed into
several categories: 1) a disturbance associated with a
monsoon trough or the intertropical convergence zone
(Riehl 1954, 1979), 2) a disturbance embedded in easterly trade wind flow such as the easterly wave (EW)
(Carlson 1969a,b; Burpee 1972, 1974, 1975; Reed et al.
1977; Saha et al. 1981; Thorncroft 1995; Thorncroft
and Hoskins 1994a,b), 3) stagnant subtropical frontal
zones that originate in midlatitudes (Frank 1987), 4) old
midlatitude mesoscale convective systems (MCSs; Bosart and Sanders 1981), and 5) upper-level cutoff lows
that penetrate to lower levels (Avila and Rappaport
1996). All of the work presented here will focus on
tropical cyclogenesis in the North Atlantic Ocean basin.
Almost all of the cyclones in the North Atlantic region
have their seeds in the form of African EWs (Riehl 1954;
Erickson 1963; Simpson et al. 1968; Frank 1970), old
frontal boundaries that originated in midlatitudes (e.g.,
Bosart and Bartlo 1991), old midlatitude MCSs (e.g.,
Bosart and Sanders 1981), or the low-level reflection of
an upper-level cutoff low (Avila and Rappaport 1996).
This paper will examine composite synoptic-scale
flows using a gridded dataset in the vicinity of developing tropical depressions. More specifically, this paper
will identify and examine the synoptic-scale flows most
commonly observed during North Atlantic basin tropical cyclogenesis, and explore the fundamental dynamics that are implied by those flows. Composites are created for genesis events in very limited areas across the
Atlantic basin and will be storm-centered. This will allow an examination and comparison of genesis events
occurring in climatologically different background
flows. Section 2 will discuss the dataset used to create
the composites. Section 3 will discuss the procedure
used to create the composites. In section 4 genesis and
lysis composite flows and composite diagnostics will be
presented. Section 5 and 6 will present the results of
the compositing and the conclusions, respectively.
2. Data
Gridded wind analyses for the North Atlantic Ocean
created by the National Hurricane Center (NHC, now
known as the Tropical Prediction Center) and archived
at the Hurricane Research Division of the National Oce-
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BRACKEN AND BOSART
anic and Atmospheric Administration/Atlantic Oceanographic and Meteorological Laboratories (NOAA/
AOML/HRD) for the years 1975–93 will be used here
to examine the synoptic-scale flow during genesis.
Twice-daily analyses of the meridional and zonal wind
components were available at two levels: Analysis of
the Tropical Oceanic Lower Layer (ATOLL ;900 hPa)
and 200 hPa (Shapiro 1986; Goldenberg and Shapiro
1996). This ATOLL–200-hPa dataset covers a region
from about 1248 to 58W and 458N to the equator. The
grid has 35 and 80 points in the meridional and zonal
direction, respectively, giving the dataset a resolution
of about 1.58 3 1.58.
An advantage of the ATOLL–200-hPa dataset is that
the analyses were created using the same methods over
a 20-yr period. National Meteorological Center (NMC;
now the National Centers for Environmental Prediction)
analyses at 850 and 200 hPa were used as a first guess
for the ATOLL–200-hPa analyses. Rawinsondes, pibal
winds, fixed and moving ship surface observations, and
manually derived satellite winds were used to improve
upon the NMC 850-hPa analysis, yielding the ATOLL
analysis. Aircraft reports, 200-hPa rawinsonde observations and manually derived satellite winds were also
used to improve upon the NMC 200-hPa analysis, yielding the final 200-hPa analysis. A more detailed discussion of the dataset and methods used to create the analyses can be found in Wise and Simpson (1971). A
discussion on the caveats of using this ATOLL–200hPa dataset is given in section 5a of this paper.
The position of tropical depressions are determined
by using the NHC best-track data (Jarvinen et al. 1984).
A second dataset identical in format to the NHC besttrack data was used to identify cases of tropical cyclolysis during the period 1967–87 (hereafter called the
lysis best-track dataset). This dataset was created at
NHC during an effort to create a CLIPER-type (climatology and persistence; Neumann 1972) forecasting
model for tropical depressions (C. J. Neumann 1998,
personal communication). Unfortunately, the effort to
create this new forecasting model was halted in 1988
when the individual compiling the data passed away.
Therefore, it might be expected that the data in the lysis
best-track dataset would be of the same quality as the
original best-tract dataset. This lysis best-track dataset
can be obtained by contacting the first author of this
paper.
The NHC best-track dataset is not completely objective and some caveats on using it will be discussed here.
The dataset was created partially by using the subjective
interpretations of hurricane forecasters and satellite analysts. Hence the starting time of a depression given by
NHC might be in error. The foremost reason for this
possible error is the lack of observations near developing depressions. This may result in an incorrect classification of system strength. It is hoped that these types
of errors will be random and the compositing procedure
will average them out. It should be expected, however,
355
FIG. 1. Formation points of tropical depressions over the North
Atlantic Ocean that later went on to become tropical storms (marked
by triangles) for 1975–93. Points of cyclolysis of tropical depressions
(between 308 and 708W and equatorward of 258N) that never became
tropical storms (marked by squares) for 1975–87.
that all hurricane forecasters and satellite analysts will
wait until persistent deep moist convection is present in
association with a cyclonic surface circulation before
declaring a system a depression. One might then consider the conditions presented in the composites as those
favorable for persistent deep moist convection in association with a cyclonic surface circulation and weak
shear, conditions highly favorable for tropical depression formation.
3. Compositing procedure
For the purpose of compositing, tropical cyclogenesis
(cyclolysis) is defined as the time a depression first (last)
appears in the NHC best-track data or the lysis besttrack data. Storm-centered composites for an area
;3000 km 3 3000 km were created for cyclones during
1975–93. Data for each storm is used only once in the
composites and only at the time of genesis or lysis. The
genesis of several storms in the composite occurred near
the edge of the data domain. For these storms using a
composite area of ;3000 km 3 3000 km results in
regions of missing data near the edge of the composite
domain. This problem was addressed by not computing
a composite at those grid points where data for a particular storm were missing. As a result, some composite
analyses will have regions of missing data along the
southern and eastern edges.
A map of genesis across the basin (Fig. 1) reveals
two large regions where genesis is favored: 1) east of
the southeast United States and over the Gulf of Mexico,
and 2) in the 108–208N band between Africa and South
America. An examination of all genesis events in each
of these large regions above (158 genesis events in total)
shows that the synoptic-scale flow differs considerably
between the two regions (not shown). Smaller representative subregions were selected from the larger regions to examine the differences in flow during genesis
(Fig. 1). Choice of genesis subregion locations was primarily based upon the expectation that an increase in
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frequency and coverage of observations in those regions
would yield a more reliable analysis. The first genesis
subregion (24 storms) is the Bahamas region (208–308N,
608–808W; composite storm position of 26.48N,
69.68W). The second genesis subregion (31 storms) is
the Cape Verde region (58–208N, 158–408W; composite
storm position of 12.88N, 28.18W). A third composite
(nine storms; composite storm position 18.08N, 85.68W)
of genesis events occurring late in the season near midlatitude frontal boundaries in the Gulf of Mexico and
Caribbean Sea will also be presented. A fourth composite (18 storms; marked by squares in Fig. 1; composite storm position 15.88N, 53.88W) consists of overwater cyclolysis events in the Atlantic basin equatorward of 258N and between longitudes 308 and 708W
(total number of overwater lysis events over the entire
Atlantic basin was 58). Cyclolysis occurring poleward
of 258N will be excluded to minimize any biases created
by cold SSTs and strong vertical shear characteristic of
midlatitudes and to eliminate the possibility of sampling
a tropical cyclone that is becoming extratropical.
In a series of papers, McBride (1981a,b) and McBride
and Zehr (1981) presented an observational analysis of
tropical cyclogenesis in the Atlantic and Pacific basin.
Twelve composite datasets were created from rawinsonde data comparing developing and nondeveloping
cloud clusters. Composites given in McBride and Zehr
(1981) showed that developing cloud clusters in the
western Atlantic were accompanied by a trough–ridge
couplet in the 200-hPa streamline analyses (their Fig.
11). They concluded that relatively strong upward vertical motion on the order of 100 hPa day21 (;2 cm s21 )
within a circle 48 latitude in radius centered on the cluster is needed for it to develop.
Synoptic-scale vertical velocities can be qualitatively
inferred for the composites shown in this paper using
diagnostics based upon quasigeostrophic (QG) theory.
In the QG omega equation the vertical velocity is forced
by two competing and overlapping processes: 1) differential vorticity advection, and 2) the Laplacian of
thermal advection. Although an alternative vertical motion method based on Q vectors (Hoskins et al. 1978)
avoids the cancellation problem of the traditional method, both methods would be impractical here since only
wind data at ;900 and 200 hPa is available. Instead,
the sense of the synoptic-scale forcing will be approximated for each composite by taking advantage of the
Trenberth (1978) QG methodology whereby vertical
motion is forced through vorticity advection by the thermal wind. Sutcliffe (1947) was the first to demonstrate
the importance of vorticity advection by the thermal
wind for the development of cyclones and anticyclones.
A highly simplified version of the Sutcliffe–Trenberth
omega equation will be used here to help locate midtropospheric regions where forcing for synoptic-scale
ascent and descent may be present at the level of QG
theory. The method is based on
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v } 2V t · =zavg ,
(1)
where v [ vertical velocity (dp/dt), z [ relative vorticity, V t [ V 200 2 V 900 , =zavg [ =(z 900 1 z 200 )/2, and
} [ ‘‘is proportional to.’’ The reader is cautioned that
this assessment of the sense of vorticity advection by
the thermal wind given in (1) is highly approximated.
Using a value of 1024 m s22 (approximately an acceleration of 10 m s21 over 24 h) for the inertial acceleration and a value of 10 m s21 for the wind velocity, the
magnitude of the Rossby number (Ro) at 158 and 108
lat is ;0.25 and 0.40, respectively. These Ro values
suggest that to first approximation QG theory can be
used qualitatively to help understand the dynamics of
the composite flows.
4. Compositing results
a. Bahamas and Cape Verde genesis cases
At the ATOLL level the Bahamas and Cape Verde
composite flow streamlines and isotachs (Figs. 2a,c)
show flow near the center of the disturbance taking on
the shape of an ‘‘inverted V,’’ similar to patterns given
in Frank and Johnson (1969) and Frank (1969). In both
regions the axis of the inverted V is slightly tilted southwest–northeast with an axis of maximum winds on the
poleward side of the inverted V. Based upon the characteristics in Fig. 2, it seems probable that most of the
ATOLL disturbances in the Bahamas and Cape Verde
composites are EWs, likely of African origin. The composite ATOLL relative vorticity shows a maximum in
cyclonic vorticity associated with both genesis cases
(Figs. 3a,c). Although both composite flows at the
ATOLL level are very similar, the composite 200-hPa
flows are very different (Figs. 2b,d). Contrasting the
200-hPa streamlines in the two genesis composites
shows that, in general, genesis in the Bahamas region
is associated with a highly amplified trough–ridge pattern and genesis in the Cape Verde region is associated
with a zonally oriented elliptically shaped anticyclone.
In the Bahamas region, composite 200-hPa isotachs
and relative vorticity show several interesting features.
Strong westerlies are observed along the poleward edge
of the domain (Fig. 2b) with weaker flow on each side
of the trough axis. A 200-hPa relative vorticity dipole
straddles the ATOLL vorticity maximum with cyclonic
(anticyclonic) vorticity upstream (downstream) (Fig.
3b). At 200-hPa the cyclonic vorticity axis does not
extend very far poleward and is confined to the equatorward side of the strong westerlies, possibly indicating
that this composite trough includes a number of lowlatitude trough fracture events (Dean and Bosart 1996).
The bulk of the cyclonic vorticity maximum is confined
to the southwest quadrant of the composite domain and
is located downstream of a positively tilted ridge axis.
Finally, a narrow band of cyclonic vorticity extends east
and then northeast from the equatorward edge of the
vorticity maximum. A larger-scale composite (not
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357
FIG. 2. Composite flow streamlines (thin solid), isotachs (every 1 m s 21 ; thick solid), and
divergence (every 0.2 3 1025 s21 in all panels except every 0.1 3 1025 s21 in e; greater than 0.2
and less than 20.2 shaded with positive values surrounded by thin solid lines and negative values
surrounded by thin dashed lines) at the ATOLL level for (a) Bahamas subregion, (c) Cape Verde
subregion, and (e) lysis subregionp; and the 200-hPa level for (b) Bahamas subregion, (d) Cape
Verde subregion, and (f ) lysis subregion. Thin dotted lines cross at the center of the composite
depression. Composite domain approximately 3000 km 3 3000 km. Blank areas in analyses are
regions where data from a storm were missing (see section 3 for further explanation).
shown) illustrates that this narrow band of cyclonic vorticity is actually the tail end of the time-mean tropical
upper-tropospheric trough (Sadler 1976) over the Atlantic Ocean.
Composite 200-hPa relative vorticity and isotachs in
the Cape Verde composite also show several key features. Relatively strong westerly flow is observed on the
poleward edge of the domain and strong easterly flow
is seen over the equatorward half of the domain (Fig.
2d). As a result of this flow configuration the ridge axis
at 200 hPa extends almost zonally northeast of the
ATOLL vorticity maximum and then arcs southwestward ahead of the low-level disturbance over the western half of the domain. Therefore, for genesis cases in
the Cape Verde region the ATOLL vorticity maxima is
observed to be squeezed between the poleward edge of
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FIG. 3. Composite flow streamlines (thin solid) and relative vorticity (every 0.3 3 1025 s21 ;
with cyclonic vorticity thick solid and anticyclonic vorticity thick dashed) at the ATOLL level for
(a) Bahamas subregion, (c) Cape Verde subregion, and (e) lysis subregion; and the 200-hPa level
for (b) Bahamas subregion, (d) Cape Verde subregion, and (f ) lysis subregion. Thin dotted lines
cross at the center of the composite depression. Composite domain approximately 3000 km 3
3000 km.
a relatively strong easterly jet and the ridge axis at 200
hPa. In addition, relatively strong westerlies are observed at 200 hPa poleward of the ridge axis, especially
northeast of the ATOLL vorticity maxima.
Vertical motions on the synoptic scale are too small
to be measured directly, but since divergence is available at the ATOLL and 200-hPa levels (Fig. 2), the
average vertical velocity in the ATOLL–200-hPa layer
may be computed using the kinematic method (O’Brien
1970). It will be assumed here that vertical velocity at
the ground and 100 hPa are zero and divergence varies
linearly between levels. Calculations of the kinematic
vertical velocity for both genesis composites (not
shown) indicate the upward vertical velocity over the
developing cyclone center averages 1–2 cm s21 or 70–
120 hPa day21 , numbers quantitatively very similar to
the results given in McBride and Zehr (1981). It should
be noted, however, that small uncertainties in the hor-
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BRACKEN AND BOSART
359
FIG. 4. Composite Sutcliffe–Trenberth forcing (contours every 5 3 10211 s22 ; with positive
forcing shaded and surrounded by thin solid lines and negative forcing by thin dashed lines) for
(a) Bahamas subregion, (c) Cape Verde subregion, and (e) lysis subregion; and the ATOLL–200hPa vertical shear of the horizontal wind (alternating solid and dashed lines at 1 ms 21 intervals,
and shaded greater than 16 m s21 ) for (b) Bahamas subregion, (d) Cape Verde subregion, and (f )
lysis subregion. Thin dotted lines cross at the center of the composite depression. Composite
domain approximately 3000 km 3 3000 km.
izontal wind field can create large errors in vertical velocities computed using the kinematic method. Therefore, the quantitative estimates of vertical velocities given here should be viewed with caution. However, the
divergence patterns displayed in both genesis composites (ATOLL convergence and 200-hPa divergence) certainly are consistent with upward vertical motion above
the ATOLL vorticity maximum.
Other perhaps more desirable techniques of comput-
ing vertical velocities can be developed using QG theory. As discussed earlier, the Sutcliffe–Trenberth method of inferring vertical motion is used here to locate
regions of forcing for upward vertical motion. Regions
of inferred upward (downward) vertical motion will coincide with regions of positive (negative) forcing or
cyclonic vorticity advection (anticyclonic vorticity advection) by the thermal wind. The Bahamas composite
(Fig. 4a) shows a very well-defined maximum (mini-
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mum) in forcing, and therefore a maximum in inferred
upward (downward) motion, directly over and slightly
northeast (west) of the ATOLL vorticity maximum. A
second maximum in forcing for ascent appears in the
extreme northwest corner of the domain. Two minima
straddle the first maximum in forcing, indicating inferred downward motion in those areas.
The use of QG theory permits a diagnosis of what is
physically causing (or forcing) upward vertical motion
over the developing disturbance; the kinematic method
does not allow such a physical interpretation. The Sutcliffe–Trenberth forcing shows that it is the high-amplitude 200-hPa trough–ridge couplet in the Bahamas
genesis composite that creates a situation favorable for
cyclonic vorticity advection by the thermal wind and
implied upward vertical motion over the ATOLL vorticity maximum. That upward vertical motion in turn
likely creates an environment conducive for persistent,
organized, deep moist convection near the center of the
ATOLL vorticity maximum.
Sutcliffe–Trenberth forcing computed for the Cape
Verde composite also shows inferred upward vertical
motion (Fig. 4c), although the forcing here is weaker
than in the Bahamas composite. Note that a given
amount of synoptic-scale forcing will yield a stronger
response (stronger upward or downward vertical velocities) in a warm air mass with small static stability (as
is found in the Tropics) when compared to a cool air
mass with larger static stability [as is found in the midlatitudes (e.g. Doswell 1987)]. In the Cape Verde composite, the area of inferred upward motion is found near
the ATOLL vorticity maximum. The upward vertical
motion over the developing disturbance is being forced
primarily through the advection of smaller values of
anticyclonic vorticity toward regions of larger values of
anticyclonic vorticity, not through the advection of cyclonic values of vorticity toward regions of anticyclonic
vorticity as was found in the Bahamas case. What is
important dynamically for forcing upward vertical motion is that vorticity advection over the disturbance is
in the cyclonic sense, the sign of the vorticity does not
matter dynamically. Therefore, in the Cape Verde case
it is the strong anticyclone and the cyclonic vorticity
advection along its equatorward flank at 200 hPa that
creates an environment favorable for upward vertical
motion.
The results from the Sutcliffe–Trenberth method are
consistent with qualitative expectations from the QG
omega equation. In that equation vertical motion is
forced by differential vorticity advection and the Laplacian of temperature advection. If one assumes that
temperature advection is negligible in the Tropics, then
upward vertical motion occurs where vorticity advection
increases with height. A diagnosis of vorticity advection
(not shown, but may be qualitatively determined with
Figs. 3a–d) indicates that at the ATOLL level in the
Bahamas composite vorticity advection is cyclonic (anticyclonic) west (east) of the depression center. At 200
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hPa the Bahamas composite shows cyclonic vorticity
advection over and east of the depression position;
therefore, in the Bahamas composite, vorticity advection
increases with height over and to the east of the depression center. Similarly, in the Cape Verde composite,
vorticity advection at the ATOLL level is cyclonic (anticyclonic) west and southwest (south and southeast) and
nearly neutral north of the depression center. At 200 hPa
cyclonic vorticity advection is concentrated in a band
oriented northeast–southwest located over and slightly
northwest of the depression center. Therefore, in the
Cape Verde composite, vorticity advection increases
with height over and north of the depression center. This
qualitative visual inspection of vorticity advection patterns in the Bahamas and Cape Verde region support
the existence of upward motion over the developing
depression.
Upward vertical motion over a low-level vorticity
maximum, by itself, does not bring about tropical cyclogenesis. Weak vertical wind shear is also a necessary,
but still insufficient, condition for tropical cyclogenesis.
Theoretically, shear will influence how mass is circulated or ventilated through a vertical column in the cloud
cluster (Gray 1968). If cloud-cluster ventilation is small,
then enthalpy and moisture in a vertical column can
increase and hydrostatically result in reduced surface
pressure. If the ventilation is large, then mass circulation
through the column is too fast to allow enthalpy and
moisture to concentrate and sustained surface pressure
falls cannot occur. An alternate explanation for why
strong vertical shear decreases the intensity of tropical
cyclones was presented by DeMaria (1996). That paper
hypothesized that midlevel warming, associated with a
shear-induced tilted maximum in potential vorticity, stabilizes the lower troposphere and reduces convective
activity near cyclone centers.
The vertical shear in these composites is computed
over the ATOLL–200-hPa layer. In the Bahamas composite (Fig. 4b) the vertical shear is relatively large over
the poleward and equatorward edges of the domain. An
oblong region of weak shear is observed near the center
and immediately south and west of the developing disturbance. The Cape Verde composite (Fig. 4d) also
shows stronger shear both poleward and equatorward
with a region of weak shear oriented east-northeast to
west-southwest across the center of the developing disturbance. Therefore, both genesis composites do show
weak shear, a situation favorable for enthalpy and moisture concentration near the center of the developing disturbance.
Vertical shear in the ATOLL–200-hPa layer at the
time of genesis was calculated over the grid point nearest the depression center and the eight grid points that
surround it. The values from those nine grid points were
then averaged to come up with a value of shear directly
over the depression. Those calculations show that shear
is weak but nonzero for both the Bahamas and Cape
Verde composites during genesis. In the Bahamas com-
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BRACKEN AND BOSART
posite the mean shear magnitude is 11.9 m s21 and the
mean shear direction is southwesterly (from 2158). In
the Cape Verde composite the mean shear magnitude is
9.1 m s21 and the mean shear direction is south-southeasterly (from 1608). These results will be discussed in
more detail in section 5b of this paper (see also Fig.
11).
b. Lysis cases
The disturbances in the lysis composite can be viewed
as a special set of lysis cases. These disturbances are
tropical depressions that obviously must have formed
in an environment that was once conducive to genesis
but has become unfavorable for the continued existence
of the depression (or conducive for cyclolysis). These
depressions in the lysis best-track dataset never went on
to become tropical storms. It must be remembered that
the disturbances in the lysis composite are undergoing
cyclolysis and therefore the composite flow pattern is
an environment that will not allow for the continued
existence of the tropical cyclone. Hence, rather than
considering the lysis cases as representative of nondevelopers it is more accurate to interpret them to be developers that were cut off in the ‘‘prime of their life.’’
Ideally one would want to contrast a composite of nondevelopers against the genesis composite, but that is not
possible with these datasets.
At the ATOLL level the disturbance for the lysis composite weakly resembles an EW defined by a feeble
inverted trough and vorticity maximum (Figs. 2e, 3e).
At 200 hPa (Figs. 2f, 3f) a trough axis is located west
of the ATOLL vorticity maximum with a very prominent jet streak in the southwest flow east of the trough
axis that extends from near the center of the disturbance
northeastward. A vorticity maximum is observed in the
base of the trough at 200 hPa. An area of strong anticyclonic vorticity dominates the east and northeast part
of the domain and is part of a dipole of vorticity that
straddles the jet streak axis. It is interesting that the flow
in the extreme eastern half of the domain is qualitatively
very similar to the Cape Verde genesis composite at this
level. Since these westward-moving lysis cases were at
some time in their past genesis cases (note that these
cases are not included in the Cape Verde composite) it
is perhaps not surprising that the flow to the east resembles the genesis composite.
Sutcliffe–Trenberth forcing of inferred vertical motion does show weak upward vertical velocities over
and near the ATOLL vorticity maximum in the lysis
composite. Composites of divergence (Figs. 2e,f) indicate ATOLL convergence and 200-hPa divergence
and therefore upward motion over the disturbance center. Cyclonic vorticity advection at 200 hPa (not shown)
and Sutcliffe–Trenberth diagnostics (Fig. 4e) both indicate the presence of weak forcing for upward vertical
motion. ATOLL–200-hPa shear (Fig. 4f) shows weak
shear equatorward and strong shear poleward, resulting
361
in a very strong gradient of shear across the ATOLL
vorticity maximum. In addition, the areal coverage of
strong shear is much larger and closer to the ATOLL
vorticity maximum, especially to the north and west. It
is hypothesized that strong shear near the disturbance
center associated with the jet streak on the east side of
the trough axis and weak forcing for ascent over and
near the center of a very weak vortex all act to destroy
the developing warm core and cause cyclolysis. It is
hypothesized that the quintessential difference between
the Cape Verde genesis composite and the lysis composite is the presence of strong shear in the lysis composite near the center of the ATOLL disturbance.
c. Anomalies of composites from climatology
Anomalies from climatology for the Bahamas, Cape
Verde, and lysis composites were created to show how
the genesis and lysis composites differ from the mean
conditions over the basin. Anomaly fields are defined
as departures of the composite fields from a weighted
climatology created using the daily ATOLL–200-hPa
analyses from 1975 to 1994. Shown in Fig. 5 are the
composite analyses minus the climatological analyses.
At the ATOLL level all three composites show an anomalous cyclonic vortex of 2.7 3 1025 s21 , 2.1 3 1025
s21 , and 0.9 3 1025 s21 in the Bahamas, Cape Verde,
and lysis composites, respectively (Figs. 5a–c). In the
Bahamas composite (Fig. 5a) the subtropical ridge is
anomalously strong on the poleward side of the anomalous cyclonic vortex. In the Cape Verde composite (Fig.
5b) anomalous southerly flow is observed on the southeastern side of the cyclonic vortex, possibly indicative
of a cross-equatorial flow into the disturbance from the
Southern Hemisphere. The lysis anomalies (Fig. 5c)
show only an anomalous cyclonic vortex with an anomalously strong anticyclone east-northeast of that vortex
and a very weakly anomalous cyclonic vortex north of
the main vortex.
The 200-hPa anomalies in the Bahamas region (Fig.
5d) show that during genesis the westerly component
of the wind is anomalously weak and the flow is anomalously cyclonic (anticyclonic) west (east) of the
ATOLL vortex. This anomalous vorticity dipole is the
source of the observed 200-hPa trough–ridge couplet
seen in Fig. 3b. In the Cape Verde composite (Fig. 5e)
the 200-hPa ridge axis on the poleward and westward
side of the developing vortex is seen as anomalously
strong. In addition, the flow poleward and westward of
the vortex has an anomalously strong southerly component. Therefore, during Cape Verde genesis events
the 200-hPa ridge poleward and west of the lower-tropospheric disturbance is anomalously strong. In the lysis
composite (Fig. 5f) the 200-hPa flow pattern shows an
anomalously strong trough west of the lower tropospheric vortex and a stronger ridge east of the vortex.
Therefore, during these lysis events the flow at 200 hPa
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VOLUME 128
FIG. 5. Composite anomalies (composite minus weighted climatology) for the Bahamas and Cape
Verde genesis cases and the lysis cases. (a), (b), (c) ATOLL anomalous streamlines (thick lines)
and relative vorticity (every 0.3 3 1025 s21 ; with cyclonic vorticity thin solid and anticyclonic
vorticity thin dashed) for the Bahamas, Cape Verde, and lysis composites, respectively. (d), (e),
(f ) Same as in (a), (b), and (c) except at 200 hPa. (g), (h), (i) Anomalous Sutcliffe–Trenberth
forcing (thick solid contours every 5 3 10211 s22 with only positive values contoured) and 200hPa–ATOLL shear anomalies (contoured every 1 m s21 ; less than 5 m s21 of shear shown by thin
lines with positive values solid and negative values dashed; shear greater than 6 m s 21 shaded
every 1 m s21 according to the color bar). Composite domain approximately 3000 km 3 3000 km.
was anomalously strong from the south and southwest
over and northwest of the low-level vortex.
Anomalies of the vertical shear of the horizontal wind
and the Sutcliffe–Trenberth forcing appear in Figs. 5g–i.
In the Bahamas region (Fig. 5g) the Sutcliffe–Trenberth
forcing is anomalously strong over and northeast of the
low-level vortex. Shear in the Bahamas composite is
anomalously weak well poleward and southeast of the
low-level vortex center and anomalously strong both
east and west of the vortex (Fig. 5g). The Sutcliffe–
Trenberth forcing in the Cape Verde composite (Fig. 5h)
is anomalously strong over and poleward of the lowlevel vortex. Shear in the Cape Verde composite (Fig.
5h) is anomalously strong over almost the entire analysis
domain with a relative minimum east and south of the
center and a relative maximum poleward of the lowlevel vortex. In the lysis composite (Fig. 5i) Sutcliffe–
Trenberth forcing is again anomalously strong over and
northeast of the low-level vortex. Shear in that composite (Fig. 5i) is anomalously strong over the entire
domain and is anomalously highest northeast of the center. This observed tendency for genesis to occur away
from those regions where shear is anomalously low and
lysis to occur where shear is anomalously high possibly
indicates that shear must be carefully balanced between
values too low to force upward motion and values too
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BRACKEN AND BOSART
high to allow moisture and enthalpy accumulation in a
deep vertical column near the center of the depression.
d. Composite flow decompositions
In an effort to determine commonalties and differences between genesis and lysis flows the composites
will be decomposed and examined here. Flows will be
decomposed by removing the approximate signature of
the cyclone/easterly wave and its associated convection.
This is done by computing and subtracting out a symmetric, convergent cyclonic vortex at the ATOLL level
and a symmetric, divergent anticyclonic vortex at the
200-hPa level centered on the ATOLL level disturbance.
The result of that subtraction (hereafter called the background flow) reveals similarities in the background flow
between the two genesis composites and differences between the genesis and lysis composite background flows
at the ATOLL and 200-hPa level.
An initial comparison of the undecomposed Bahamas
and Cape Verde genesis composites reveals similarities
only at the ATOLL level and large differences at the
200-hPa level. However, many similarities at both the
ATOLL and 200-hPa level are revealed once the flow
is decomposed. At the ATOLL level the background
flow is characterized by strong deformation with the
axis of dilatation in the Bahamas region oriented southwest–northeast northwest of the center of the composite
(Fig. 6a) and east–west over the center of the composite
in the Cape Verde region (Fig. 6c). This flow configuration favors a concentration of moisture and thermal
gradients near the depression in the genesis composites.
At the 200-hPa level the background flows are nearly
identical and again characterized by strong deformation.
Both 200-hPa background flows (Figs. 6b,d) show two
regions of cyclonic flow north and south of the center
connected to each other by an axis of cyclonic vorticity
that passes just to the west of the center and two regions
of anticyclonic vorticity east and west of the center. This
results in the placement of the depression in a general
southeasterly flow on the equatorward side of a zonally
oriented ridge axis, which then allows for the existence
of vorticity advection in the cyclonic sense and forces
upward motion over the developing depression. Therefore, the genesis composites exhibit some fundamental
similarities once the signature of the developing cyclone/easterly wave and its associated convection are
removed.
A comparison of the decomposed genesis and lysis
flows reveal differences, the most striking of which is
seen at the 200-hPa level (Figs. 6b,d,f). At 200 hPa the
lysis depression is located directly beneath a zonally
oriented ridge axis and is not equatorward of a ridge
axis as was seen in both genesis composites (Fig. 6f).
This may indicate that cyclonic vorticity advection at
200 hPa over the depression is precluded, resulting in
an absence of forced upward motion. Another significant
difference between the two composites is the presence
363
of a well-defined 200-hPa trough and its associated
southwesterly jet streak in the vicinity of the lysis depression. This trough and its associated jet streak may
act to produce strong vertical shear of the horizontal
wind in the vicinity of the depression, precluding development.
e. Late season genesis in the western Caribbean and
southern Gulf of Mexico
As was mentioned previously, a small portion (3%–
5%) of tropical cyclones form in subtropical regions
along stagnant frontal zones, which originate in midlatitudes associated with more baroclinic environments.
This section will briefly describe a composite of nine
tropical cyclogenesis events in the western Caribbean
and southern Gulf of Mexico (158–258N, 808–958W)
subjectively picked from the September–November period of 1977–91. Satellite images and synoptic surface
maps were inspected. Those storms that appeared to be
associated with old midlatitude frontal boundaries were
included in the composite (storms included in the composite can be determined by using Fig. 10). Composites
were created using methods exactly the same as those
used to create the Bahamas and Cape Verde composites.
An examination of satellite imagery during the nine
genesis events (not shown) indicated that all events occurred in close proximity to the equatorward end of an
old midlatitude frontal boundary. The possible presence
of a frontal boundary indicates that jet streak dynamics
may be playing an important role during these tropical
cyclogenesis events. Therefore, these events will be presented and discussed separately from the Bahamas and
Cape Verde composites.
Composite fields for these cases of ‘‘frontal’’ tropical
cyclogenesis are given in Fig. 7. The composite ATOLL
streamlines (Fig. 7a) show a pattern different from the
Bahamas and Cape Verde composites (Figs. 2a,c). The
inverted V pattern is not observed in the streamlines of
the frontal composite, possibly indicating that the developing disturbances in these cases may not originate
from EWs. Instead, the streamlines indicate a more symmetric vortex at the ATOLL level with stronger flow on
its poleward side. At 200 hPa (Fig. 7b) an anticyclone
is centered just east of the ATOLL-level cyclonic vorticity maximum (Fig. 7c). Parcels that flow poleward
on the west side of the 200-hPa anticyclone should experience strong acceleration into a confluent jet-entrance
region located between a weak trough and the poleward
side of the anticyclone (Fig. 7b). Since the ageostrophic
wind blows perpendicular and to the left of the parcel
acceleration (in the Northern Hemisphere) the acceleration of parcels into the jet should result in a thermally
indirect transverse circulation across the jet axis at 200
hPa. Divergence (convergence) on the right (left) side
of the 200-hPa jet-entrance region (Fig. 7b) supports
this conjecture. This divergence above ATOLL level
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VOLUME 128
FIG. 6. Background flow computed by decomposing the total composite flow as described in
text (full barbs 5 m s21 and half-barbs 2.5 m s21 ) for (a) Bahamas subregion, (c) Cape Verde
subregion, and (e) lysis subregion at ATOLL level; and for (b) Bahamas subregion, (d) Cape
Verde subregion, and (f ) lysis subregion at 200 hPa. Composite domain approximately 3000 km
3 3000 km.
convergence (Fig. 7a) kinematically indicates upward
vertical motion over the low-level vorticity maximum.
The ATOLL and 200-hPa vorticity fields (Figs. 7c,d)
for the frontal composite do show significant differences
from the Bahamas and Cape Verde composites. At the
ATOLL level (Fig. 7c) the cyclonic vorticity maximum
is stronger than in the Bahamas and Cape Verde composites. A band of cyclonic vorticity extends east-northeast from the ATOLL vorticity maximum, possibly indicating the presence of a composite frontal boundary.
At 200 hPa (Fig. 7d) a vorticity dipole straddles the jetentrance region poleward of the anticyclone. This 200hPa anticyclone is characterized by two major ridge
axes. As in the Cape Verde composite a zonal ridge axis
extends westward from the anticyclone center. The presence of a jet poleward of the anticyclone helps create
a second ridge axis that extends poleward almost meridionally from the anticyclone center.
Sutcliffe–Trenberth forcing diagnostics (Fig. 7e)
show three regions of inferred upward motion over the
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365
FIG. 7. Frontal genesis composite; flow streamlines (thin solid), isotachs [every 1 m s 21 in (a)
and every 2 m s21 in (b) thick solid], and divergence [every 0.2 3 1025 s21 in (a) and every 0.3
3 1025 s21 in (b); greater than 0.2 and less than 20.2 shaded in (a) and greater than 20.3 and
less than 20.3 shaded in (b) with positive values surrounded by thin solid lines and negative values
surrounded by thin dashed lines] for (a) the ATOLL level and (b) the 200-hPa level. Composite
flow streamlines (thin solid) and relative vorticity (every 0.5 3 1025 s21 ; with cyclonic vorticity
thick solid and anticyclonic vorticity thick dashed) for (c) the ATOLL level and (d) the 200-hPa
level. Composite Sutcliffe–Trenberth forcing (every 5 3 10211 s22 ; with positive forcing shaded
and surrounded by thin solid lines and negative forcing in thin dashed lines) in (e). ATOLL–200hPa vertical shear of the horizontal wind (thick solid every 2 m s 21 and shaded greater than 16 m
s21 ) in (f ). Thin dotted lines cross at the center of the composite depression. Composite domain
approximately 3000 km 3 3000 km.
composite domain. The most prominent area of inferred
upward motion extends northeast–southwest over the
ATOLL cyclonic vorticity maximum. Two other regions
of implied ascent are observed over the northwest and
equatorward edge of the domain. As was noted in the
other genesis composites the area of inferred upward
motion over and near the ATOLL cyclonic vorticity
maximum is forced by cyclonic vorticity advection by
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MONTHLY WEATHER REVIEW
the thermal wind. But in this composite (and in the Cape
Verde composite), the area of inferred upward motion
is located upstream of a 200-hPa ridge axis. This frontal
composite is different from the Cape Verde composite
in that there are two prominent 200-hPa ridge axes that
play a role in creating inferred upward motion. One axis
extends westward from the anticyclone and the other
extends poleward almost meridionally from the anticyclone center. Cyclonic vorticity advection upstream
of the two axes combine to create one coherent area of
inferred upward motion with the strongest forcing over
and poleward of the ATOLL cyclonic vorticity maximum.
As might be expected, shear is very strong over the
poleward half of the composite domain in the vicinity
of the jet (Fig. 7f). But over the equatorward half of
the domain shear is small, thereby creating a strong
gradient of shear across the developing ATOLL cyclonic
vorticity maximum. This results in shear values near the
developing disturbance larger than those observed in
the lysis composite. However, in this genesis composite
the ATOLL cyclonic vorticity and forcing for ascent are
much stronger than in the lysis composite. This result
suggests that a particular shear value alone does not
determine if tropical cyclogenesis occurs.
5. Discussion
The Bahamas and Cape Verde genesis composites
presented here possess characteristics that are very similar to those seen in published papers on tropical cyclogenesis. A case study in the Australian–South Pacific
region of cyclones Irma and Jason by Davidson et al.
(1990) and work by McBride and Keenan (1982) suggests that tropical cyclogenesis can occur in association
with an equatorward-moving upper-tropospheric trough.
In the western North Pacific region upper-tropospheric
troughs and ridges in close proximity to developing
tropical cyclones have been associated with genesis
(Sadler 1976, 1978). Figures 1b,c in Sadler (1976) possess features that are very similar to those seen in the
Bahamas and Cape Verde genesis composites. Both Sadler (1976) figures show essentially the same dynamical
signature. In his figures the developing system is located
southwest of an upper-level anticyclone and equatorward of a zonally oriented ridge axis, a flow pattern
very similar to the Cape Verde genesis composite. It is
also interesting to note the presence of an upper-level
trough in the Sadler (1976) figures, as was seen in the
Bahamas composite. Sadler (1976) suggests that the
flow pattern given in those figures leads to enhanced
outflow and divergence in the cyclone outflow layer. As
was shown in this paper, those flow patterns create regions of upper-tropospheric cyclonic vorticity advection
that are favorable for ascent and therefore create an
environment that will support persistent, deep moist
convection.
Published observations of cyclogenesis in the western
VOLUME 128
North Atlantic basin show how often genesis is accompanied by upper-tropospheric troughs and ridges. Erickson (1963) gives a description of observations during
the development of Tropical Cyclone Debbie (1961)
near Cape Verde and notes the presence of an upperlevel anticyclone during the genesis process. Colon and
Nightingale (1963) noted that out of 40 cases of cyclone
development over the North Atlantic, 28 occurred west
of an upper-level anticyclone or east of a trough in poleward flow. The development of Tropical Cyclone Alma
(1962) is discussed in Yanai (1968). Alma is shown to
form as a result of the interaction between an African
easterly wave (AEW) and an upper-tropospheric trough
over the Caribbean Sea region. In a case study by Bosart
and Bartlo (1991), the genesis of Tropical Cyclone Diana is shown to occur as the result of the interaction of
an upper-tropospheric trough and an old frontal boundary off the southeast United States. In a case study of
the high-latitude development of Tropical Cyclone Dorothy (1966), Erikson (1967) emphasizes the important
role of upper-tropospheric vorticity advection associated
with an upper-tropospheric trough above a well-defined
low-level disturbance.
a. Validity of composites
An obvious question that must be answered is, Do
the individual genesis cases that make up the composite
resemble the composite analysis? This question can be
approached quantitatively using the two-sided Student’s
t-test or by a qualitative subjective inspection of the
individual cases in the composites. Both methods will
be used here. The two-sided Student’s t-test is used to
inquire whether the mean of a sample (the composite)
is significantly different than the mean of the population
of all genesis events (Panofsky and Brier 1968). Calculations of the statistical significance of composite vorticity, wind speed, vertical shear, and Sutcliffe–Trenberth forcing indicate that all features in the Bahamas,
Cape Verde, and lysis composites are at least significant
to the 95% level (not shown). The only exception to
the last statement is the 200-hPa cyclonic vorticity maximum in the lysis composite (note that the 200-hPa jet
streak in that composite is significant to the 95% level).
Before a subjective investigation of the individual
composite members begins, several issues concerning
the characteristics of the datasets should be noted. The
ATOLL–200-hPa analyses are created objectively and
as a result will contain errors. Analysis errors may arise
for several reasons: 1) measuring error, 2) analysis
scheme smoothing, 3) a lack of observations, and 4) the
tendency for the analyses to resemble the 12-h forecast
in data void regions. Errors in the best-track dataset may
arise due to the highly subjective nature of the decision
process at NHC and the lack of observations near developing cyclones. NHC’s methods of identifying the
beginning or end of a tropical cyclone’s life cycle are
not objective. An exact determination of the time and
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BRACKEN AND BOSART
location of tropical cyclogenesis is extremely difficult,
if not impossible. Undoubtedly there will be some differences in opinion among the individual forecasters at
NHC as to exactly where and when a tropical cyclone
has developed or decayed. It is hoped that the criteria
to declare a cloud cluster a tropical depression will be
made primarily based upon observations and be consistent among NHC forecasters. It is also expected that
the errors in both datasets will be random and that by
creating composites these random errors will be removed from the analyses and only features that play a
significant role in the tropical cyclogenesis process will
remain. These errors may occasionally be observed in
the individual members that make up the composite.
Therefore, one should not expect each member to look
exactly like the composite. One should expect most
composite members to be similar to the composite with
variations in the position and intensity of significant
features due to the errors noted above. An additional
source of difference is natural variability. No two cases
of genesis look exactly alike. If there is a common dynamical signature it should be observed in the composite.
A subjective examination of the 200-hPa flow for
each member in the Cape Verde composite (Fig. 8)
shows that 87% (27 of 31; note that storms with missing
data were not tabulated) of the members are similar to
the composite analysis (Figs. 2, 3). These 27 cases are
marked by the presence of an anticyclone or ridge axis
to the poleward or east and easterly flow over or equatorward of the ATOLL-level vorticity maxima. A small
shift in the location/time of genesis or the strength of
the synoptic-scale features would result in excellent
agreement between these 27 cases and the composite.
The four remaining cases (Frances 1976; Dennis 1981;
Irene 1981; Dennis 1987) are accompanied by flows that
are not very similar to the composite but are not completely contrary to the composites either. Subjective inspection of the 200-hPa flow for each member in the
Bahamas composite (Fig. 9) show close similarities between 83% (20 of 24) of the cases and the composite
analysis. All 20 cases possess a trough–ridge couplet in
the flow streamlines that would very closely match the
overall composite analysis if the strength of some of the
synoptic-scale features were altered or if a small shift
in the location/time of genesis occurred. The four cases
that do not match the composite (Blanche 1975; Caroline 1975; Barry 1983; Ana 1991) do not possess a
significant trough–ridge couplet at 200 hPa. Even the
nine frontal cases (Fig. 10) generally match the composite analyses. All frontal cases are characterized by
the presence of a 200-hPa trough and ATOLL disturbance (not shown) located near the equatorward jet entrance region of a 200-hPa westerly jet streak located
poleward of the developing storm on the downstream
side of the trough. Henri (1979) is the only storm that
does not develop in the equatorward jet-entrance region
367
of a westerly jet streak; instead Henri develops in a
poleward jet-entrance region of an easterly jet.
A final check on the representativeness of the composites can be made by using composites for other independent regions. Since the area covered by the Bahamas and Cape Verde subregions are relatively small,
additional composites can be created for regions adjacent to the Bahamas and Cape Verde subregions with
an independent set of depressions. If the composite analyses in the Bahamas and Cape Verde subregions match
the composite analyses in adjacent regions, then one
might conclude that the composites are representative
of conditions during genesis. Or, in statistical terms, if
two independent subsamples of a given population both
yield the same mean (or composite), then the means are
likely representative of the mean of the total population.
Composites analyses in regions adjacent to the Bahamas
and Cape Verde regions were calculated using a total
of 60 separate genesis events (not shown). A comparison
of the Bahamas and Cape Verde composites, with composite analyses in those adjacent regions, shows that
they do very closely match each other. This indicates
that the Bahamas and Cape Verde composite flows are
representative of conditions commonly observed in
those regions during tropical cyclogenesis.
b. Role of synoptic-scale flows
A primary requirement for tropical cyclogenesis is
the existence of persistent, organized, deep moist convection. Sustained convection is the essential ingredient;
without a mechanism to produce it, other quantities like
vertical shear of the horizontal wind, low-level vorticity,
and SST are irrelevant. Existence of organized deep
moist convection in association with a lower-tropospheric cyclonic vorticity maximum, small vertical
shear, and SST .278C should be favorable for tropical
cyclogenesis. The role of the synoptic-scale flow in creating this favorable genesis environment is now discussed.
Meteorologists who forecast weather for midlatitude
locations quickly realize that precipitation is commonly
observed on the downstream side of an upper-tropospheric trough axis. The physical reasons for why this
occurs are found within QG theory. If the upper-level
flow is assumed to be in geostrophic balance and temperature advection is neglected, then for an idealized
sinusoidal 200-hPa flow pattern, upward vertical motion
on the downstream side of a 200-hPa trough axis and
the upstream side of a 200-hPa ridge axis must be primarily forced through an upward increase in cyclonic
vorticity advection. A fundamental difference between
the Cape Verde and Bahamas composites is the existence
and orientation of the trough and ridge axes. In the
Bahamas composite both the trough and ridge axis are
meridionally oriented so that forcing for upward motion
exists between the trough and ridge axes. In the Cape
Verde composite the trough axis is absent so that upward
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FIG. 8. 200-hPa flow streamlines (thick solid) and ATOLL-level cyclonic vorticity (greater than 1.0 3 1025 s21 shaded) for all cases included in Cape Verde subregion composite.
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FIG. 9. 200-hPa flow streamlines (thick solid) and ATOLL-level cyclonic vorticity (greater than 1.0 3 1025 s21 shaded) for all cases included in Bahamas subregion composite.
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FIG. 10. 200-hPa flow streamlines (thick solid) and isotachs (greater than 15 m s 21 shaded every 5 m s21 ) for all cases included in frontal tropical cyclogenesis composite.
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motion is forced only equatorward of a zonally oriented
ridge axis. However, both composites are accompanied
by forced upward motion created by the same mechanism, an upward increase in cyclonic vorticity advection. Likewise, it can be shown that in the frontal genesis
composites that upward QG vertical motion over the
ATOLL cyclonic vorticity maximum lies beneath the
200-hPa equatorward jet-entrance region and that this
ascent is forced by the same mechanism as in the Bahamas and Cape Verde composites, that is, an upward
increase in cyclonic vorticity advection. This inferred
ascent pattern is consistent with that found in the vicinity of jet streaks (e.g., Bluestein 1993, 397–407).
The coexistence of SST .278C, a synoptic-scale flow
pattern favorable for convection, and a lower-tropospheric cyclonic vorticity maximum (i.e., an EW or an
old midlatitude frontal boundary) is not enough for tropical cyclogenesis. It is believed that weak vertical wind
shear is required so that enthalpy and moisture can accumulate and increase in magnitude in a vertical column
near the depression center (Gray 1968). If ventilation
is minimized then enthalpy and moisture in the vertical
column can increase and hydrostatically result in lower
surface pressure. Alternatively, DeMaria (1996) hypothesized that strong shear inhibits convection near the
cyclone center by stabilizing the lower troposphere.
Nevertheless, shear will not directly play a role in the
initiation of deep moist convection; it primarily determines the structure of convection after it is initiated. As
the vertical shear increases, the structure of convection
will move from ordinary cells (or single cells) to multicells (e.g., Bluestein 1993, 487–492). Ordinary-cell
type thunderstorms are characterized by small ventilation and therefore are conducive to the accumulation of
enthalpy and moisture in a vertical column.
The Bahamas and Cape Verde genesis composite flow
analyses do show weaker shear over the ATOLL-level
disturbance. Weak shear in the Cape Verde composite
is perhaps not surprising given the existence of a strong
anticyclone aloft, but weak shear in the Bahamas composite associated with an upper-level trough might be
unexpected. Weak shear in the Bahamas composite is
the result of the superposition of generally weak southerly flow at both 200 hPa and the ATOLL level. In
general, upper-level troughs (especially those at midlatitudes) are associated with jet streaks and relatively
stronger flow. The Bahamas composite trough is not
accompanied by very strong flow. This trough is not
typical of troughs commonly seen embedded in the
westerlies at midlatitudes. Wind speeds at 200 hPa are
strongest well poleward of the trough and weakest along
the trough axis and at the center of the anticyclone. No
strong jet streaks are observed upstream or downstream
of the trough axis.
The favorable trough–ridge pattern in the Bahamas
composite should be contrasted with the unfavorable
trough–ridge pattern in the lysis composite. The most
significant differences between the two composites are
371
FIG. 11. Histogram of ATOLL–200 hPa (;900 hPa) layer-average
shear magnitude (m s21 ) at the time of genesis for all depressions
during the time period 1975–93 (139 storms). Shear is computed by
subtracting the ATOLL wind from the 200-hPa wind. The calculation
is made for the grid point nearest the depression and the eight grid
points that surround it. The shear value used is the average of those
nine grid points.
the longer half-wavelength between the trough and ridge
axis and the existence of a strong jet streak in the southwesterly flow between the trough and ridge in the lysis
composite at 200 hPa. The consequences of these differences are weaker cyclonic vorticity advection and
forcing for ascent and stronger shear over the ATOLL
vorticity maximum in the lysis composite.
Most of the previous discussion has concentrated on
the detrimental effects of shear. Therefore, it may seem
paradoxical to hypothesize that some weak shear is necessary for tropical cyclogenesis, but that idea will be
investigated here. The magnitude of vertical shear during genesis has been the subject of controversy. In the
past it has been suggested that zero or near-zero vertical
shear must be present during genesis (e.g., McBride and
Zehr 1981). An examination of vertical shear for all
genesis events available in the ATOLL–200-hPa dataset
from 1975 to 1993 over the Atlantic basin (a total of
139 storms) is shown in Fig. 11. Shear is computed by
subtracting the ATOLL wind from the 200-hPa wind.
The calculation is made for the grid point nearest the
depression and the eight grid points that surround it.
The average value at those nine grid points for all 139
storms are shown in Fig. 11. The same calculations were
made using the grid point nearest the depression and
the grid point northwest, northeast, southwest, and
southeast of it (5 points in total). The results from the
5-point calculation were nearly identical to the 9-point
calculation. The results show that during genesis the
372
MONTHLY WEATHER REVIEW
200-hPa–ATOLL layer vertical shear magnitude is near
10 m s21 (the mean is 10.5 m s21 , the median is 10.0
m s21 , and the standard deviation is 4.8 m s21 ).
The results presented here do not show a tendency
for the vertical shear to be near zero during genesis. It
is interesting to note that while it is commonly believed
that shear inhibits the genesis process, the existence of
some weak shear is required to force synoptic-scale ascent. This suggests that shear magnitude during genesis
must be carefully balanced between values small enough
to prevent strong ventilation and values large enough
to be dynamically significant (i.e., allow for advective
processes, e.g., forcing for ascent, to exist). Figure 5
can also be used to give some credibility to this hypothesis.
In the Bahamas genesis composite the shear is anomalously weak north and southeast of the low-level vortex; the rest of the analysis domain is covered by anomalously strong shear (Fig. 5g). East and west of the
center the shear attains its maximum anomalous value.
Directly over and near the center, shear is only anomalously high by 0–5 m s21 . The area of shear that is
weakly anomalously positive in the Bahamas composite
is collocated with an area of anomalously high Sutcliffe–Trenberth forcing for ascent (Fig. 5g). The same
general characteristics found in the Bahamas composite
can be seen in the Cape Verde composite (Fig. 5h). Shear
over and near the developing low-level vortex is only
anomalously high by 0–5 m s21 in a region where the
Sutcliffe–Trenberth forcing for ascent is anomalously
high. In the lysis composite the Sutcliffe–Trenberth
forcing for ascent is again anomalously high (Fig. 5i),
but in that composite the shear over and near the depression is greater than 5 m s21 above normal.
To summarize, the composites show that shear during
genesis is a relative minimum over and near the developing depression. However, genesis does not tend to
occur in regions where shear is nonexistent (zero) or
even anomalously weak. Genesis tends to occur where
the shear is weak (when compared to the environment
surrounding the storm) but just slightly (0–5 m s21 )
anomalously positive. One possible explanation of this
is that in order for there to be forced synoptic-scale
ascent there must be some shear. On the other hand,
there cannot be so much shear that mass circulation
through the column is too fast to allow enthalpy and
moisture concentration. The case of too much shear
seems to be what is observed in the lysis composite.
Therefore, it may be concluded that shear magnitude
during genesis must be carefully balanced between values small enough to prevent strong ventilation and values large enough to create forcing for ascent. However,
from these results the exact nature of the most favorable
shear profile for genesis is unclear. Perhaps a concentration of shear over a given layer of the troposphere
might be more favorable than shear of similar magnitude
over other layers in the troposphere.
The identification of regions favorable for tropical
VOLUME 128
cyclogenesis on any given day is not simple. Locating
regions of weak shear by identifying upper-level anticyclones or regions of easterly flow aloft is not adequate
enough to determine if conditions are favorable for genesis. One must locate regions where the synoptic-scale
flow will favor persistent, organized deep moist convection. An examination of satellite images shows that
over the tropical Atlantic Ocean MCSs commonly do
not last more than one or two diurnal heating cycles
(Laing and Fritsch 1997). The initiation of an MCS
depends upon the existence of three things: 1) low-level
moisture, 2) instability, and 3) lift (Doswell 1987). The
supply of the first ingredient, low-level moisture, is certainly adequate for deep moist convection in the Tropics.
For the second ingredient, instability, an assumption
commonly made is that the air mass above the tropical
oceans is almost always conditionally unstable and supplies an abundant and widespread energy source for
deep moist convection. This assumption is valid only
for undilute vertical displacements of air parcels that do
not include the effects of water loading. The idea that
convective available potential energy (CAPE) in the
Tropics is large has recently been challenged by Betts
(1982) and Xu and Emanuel (1989). They showed that
soundings from the deep Tropics are nearly neutral to
reversible adiabatic ascent when adiabatic condensate
is included in the definition of buoyancy. Therefore, the
assumption of widespread regions of large CAPE over
the tropical oceans may need to be replaced by widespread near neutrality with only small regions where
CAPE is large. The final ingredient in the initiation of
an MCS, lift, is a mechanism that is forced by processes
acting on many different scales of motion (Doswell
1987). Lift is required to bring a parcel to its level of
free convection. Lift can be generated on several different scales: storm scale, mesoscale, synoptic scale, and
planetary scale. Planetary-scale upward motion occurs
only very slowly over areas much larger than an MCS
and will likely not play the primary role in the initiation
of convection and tropical cyclogenesis. Planetary-scale
features such as the Madden–Julian Oscillation (Madden
and Julian 1994) may act to enhance convection over
large areas of the Tropics by creating a larger-scale environment that is favorable for ascent forced by processes on smaller scales. In contrast, storm-scale lift is
very localized and cannot initiate upward motion over
an area large enough to have an impact on the scale of
an MCS.
The initiation of deep moist convection, assuming
instability and moisture exist, is then mostly controlled
by lift on the mesoscale and synoptic scale. Both scales
interact to produce the lift that may initiate an MCS.
Lift and subsidence on different scales can oppose one
another and result in weak vertical velocities. But sometimes lift on different scales can cooperate to produce
strong upward motion. It is theorized that during tropical
cyclogenesis the mesoscale and synoptic scale both cooperate to create an environment favorable for persis-
FEBRUARY 2000
BRACKEN AND BOSART
tent, organized deep moist convection. The role of the
synoptic scale in producing ascent in the Tropics and
subtropics is believed to be twofold: 1) produce a large
area of tropospheric-deep lift over the low-level cyclonic vorticity maximum, thereby creating a favorable
environment for strong upward motion and thunderstorm initiation when mesoscale lift is also present; and
2) destabilize the air and/or remove any capping trade
wind inversions through that ascent if the air is neutrally
buoyant or stable over large areas. Given favorable synoptic-scale conditions for ascent accompanied by a lower-tropospheric cyclonic vorticity maximum, weak vertical shear, and SST .278C, then the situation might be
considered favorable for tropical cyclogenesis. It has
been shown here that the conditions in the frontal, Bahamas, and Cape Verde genesis composites meet all
these requirements (SSTs are not shown in composites,
but are assumed .278C since genesis occurred and the
climatological SSTs are .278C) and therefore illustrate
the most common example of conditions favorable for
tropical cyclogenesis over the North Atlantic basin.
One of the primary results presented here appears to
conflict with results presented in Zehr (1992). That paper concludes that upper-level wind patterns do not take
an active role as a forcing function on tropical cyclogenesis over the western North Pacific Ocean. In contrast, the upper-level synoptic-scale wind pattern appears to have a strong control over where and when
tropical cyclogenesis occurs over the Atlantic Ocean in
part through the creation of synoptic-scale ascent and
persistent organized, deep moist convection. Perhaps
these contrasting results indicate that tropical cyclogenesis is fundamentally very different over the different
ocean basins and/or is a consequence of differences in
computational methods.
Finally, the most important question is then, How
does the contemporaneous existence of synoptic-scale
ascent, weak shear, low-level cyclonic vorticity, and
SST .278C ultimately result in a warm core vortex?
That process is highly dependent upon interactions between different scales of motion and will not be addressed here, but will be addressed in a later paper. The
synoptic scale plays several different roles in tropical
cyclogenesis; it can 1) provide a low-level cyclonic vorticity maximum; 2) create an environment with weak
vertical shear of the horizontal wind; 3) create an environment favorable for persistent, organized deep moist
convection; and 4) perhaps moisten mid- and upper levels through that convection thereby diminishing the negative impacts of evaporatively cooled downdrafts. The
details of the interactions between these synoptic-scale
processes and the mesoscale flow will ultimately determine where, when, and how a warm core vortex forms.
6. Summary
The synoptic-scale flow during tropical cyclogenesis
in the North Atlantic basin has been examined using
373
storm-centered composites. Composites are created using only wind data at the ATOLL (;900 hPa) and 200hPa level. The results suggest that two different largescale upper-tropospheric flow patterns are most commonly observed during genesis over the basin. One flow
pattern is characterized by an upper-tropospheric
trough-ridge couplet and is most commonly observed
in the Bahamas region. The low-level cyclonic vorticity
maximum in the Bahamas composite is located beneath
the poleward flow east (west) of the upper-level trough
(ridge) (Fig. 2). The second flow pattern is commonly
observed in the Cape Verde region and characterized by
an upper-tropospheric ridge axis poleward of the lowlevel cyclonic vorticity maximum (Fig. 2). A third less
commonly observed flow pattern is observed in association with the equatorward end of old, stagnant midlatitude frontal boundaries. This third pattern is characterized by a 200-hPa jet streak entrance region on the
downstream side of a trough (Fig. 10).
Although the two most commonly observed genesis
patterns are very different at the 200-hPa level, they do
have many similarities when the flow field is decomposed. Removal of the signature of the developing cyclone/easterly wave and its associated convection from
the composite flow field leaves a flow pattern characterized by strong deformation (Fig. 6). Both composites
show a jet exit region at 200 hPa in the vicinity of the
developing depression with an axis of dilatation oriented
northeast–southwest to the northwest of the developing
disturbance. At the ATOLL level the decomposed flow
field again shows strong deformation with an axis of
dilatation northwest and over the depression in the Bahamas and Cape Verde composites, respectively.
The configuration of the undecomposed full flow
fields presented in all three genesis composites are favorable for tropical cyclogenesis for several reasons.
First, they provide a low-level cyclonic vorticity maximum. Second, the flow patterns force synoptic-scale
ascent through cyclonic vorticity advection by the thermal wind above the low-level cyclonic vorticity maximum. It is hypothesized that this inferred synoptic-scale
upward motion then creates an environment favorable
for the creation of persistent, deep moist convection
through several different processes: 1) destabilization of
the lower troposphere; 2) removal of capping trade wind
inversions that may preclude deep, moist convection; 3)
generation of low-level cyclonic vorticity by horizontal
convergence associated with Ekman pumping; and 4)
creation of a favorable environment for mesoscale lift.
Third, the flow pattern is configured in such a way that
the vertical shear of the horizontal wind is minimized
over and near the low-level cyclonic vorticity maximum
but yet large enough to force synoptic-scale ascent. Last,
the presence of the persistent deep, moist convection
might then act to moisten mid- and upper levels of the
troposphere and thereby diminish the stabilizing effects
of evaporatively cooled downdrafts.
Composites of overwater tropical cyclolysis events
374
MONTHLY WEATHER REVIEW
illustrate that not all 200-hPa troughs are favorable for
tropical cyclogenesis. Some 200-hPa troughs are associated with weakening depressions. Tropical cyclolysis
events are generally associated with 1) small forcing for
synoptic-scale ascent resulting from weak cyclonic vorticity advection by the thermal wind, 2) a weak ATOLL
cyclonic vorticity maximum, and 3) strong shear associated with a jet streak on the downstream side of the
200-hPa trough axis.
Another factor that has not been addressed here but
could be playing a role in tropical cyclogenesis is the
intrinsic potential of the incipient disturbance. Some
lower-tropospheric cyclonic vorticity maxima might
have a structure that would make them more favorable
for genesis. A modeling study by Kwon and Mac (1990)
suggests that some AEWs are more favorably structured
thermodynamically for growth. Unfortunately, the dataset used here contains no thermodynamic information
and wind data is available at only two levels. Mesoscale
details about the vertical structure of the incipient disturbance are not obtainable. Those limitations aside, the
data do show the presence of synoptic-scale forcing for
ascent during genesis, indicating that some weak shear
is necessary to permit the existence of advective processes that can force ascent within the developing depression.
The results presented here suggest a strategy that forecasters may use to help identify areas favorable for tropical cyclogenesis. Locating regions of weak shear by
identifying upper-level anticyclones or regions of easterly flow aloft is too general of a strategy and will not
sharply focus attention on those regions that are most
likely for genesis. Without a mechanism to produce persistent, organized deep moist convection, other factors
like vertical shear of the horizontal wind, low-level vorticity, and SST are irrelevant. Once a synoptic-scale flow
pattern favorable for ascent and persistent, organized
deep moist convection is established, then these other
factors become important. For most cases, identifying
regions where vertical shear, low-level vorticity, and
synoptic-scale vertical motion are in general agreement
with the values found in the Bahamas, Cape Verde, and
frontal composites in conjunction with SST .278C
should significantly reduce the size of the area where
tropical cyclogenesis is possible.
Acknowledgments. The research presented here represents part of the first author’s (WEB) Ph.D. thesis from
SUNY/Albany. Some of the work presented here was
completed while WEB visited the Hurricane Research
Division (HRD) of NOAA/ERL/AOML and worked
there as part of the U.S. Department of Commerce Student Career Experience Program. A very special thanks
goes to Bob Burpee of HRD who was instrumental in
arranging cooperation between HRD and SUNY/Albany. Without his foresight to realize how productive
this cooperation would be and his willingness to take a
chance on a graduate student, this research may have
VOLUME 128
never have been undertaken. Peter Black, Frank Marks,
and Hugh Willoughby all made contributions to this
paper through conversations on the topic of tropical
cyclogenesis. Sam Houston was kind enough to help
prepare some figures in this paper. Thanks also to Bob
Kohler who assisted in the retrieval of the ATOLL–200hPa data from the HRD archives. Howard Friedman also
deserves a special thanks for all his administrative work
in arranging WEB’s stay at HRD. Celeste Iovinella has
also been a huge help in preparing this paper. Comments
from the three anonymous reviewers greatly improved
this paper also. This research was supported by NSF
Grants ATM9612485 and ATM9413012.
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