1. Introduction
1.1 General Purpose
The purpose of this thesis is to document the spatial structure and
temporal evolution of heavy rainstorms that sometimes develop well poleward of
landfalling or near-coastal-tracking Atlantic basin tropical cyclones (TCs) (refer to
the appendix for a list of acronyms used in this thesis). A statistical and
composite climatology of these predecessor rain events (PREs) was established
using a database of 47 cases occurring downstream of 21 TCs between 1998
and 2006. The PREs from three TCs—two from within the climatological period
and one from before—were chosen for detailed synoptic-scale and mesoscale
study based on their capacity to illustrate significant findings from the climatology
as well as physical mechanisms crucial to their formation. A null case also was
selected so the environment downstream of the TC could be contrasted with
those cases that spawned PREs. The aim of this type of approach is to continue
the mission of the warm-season project of the Collaborative Science,
Technology, and Applied Research (CSTAR) program under which the research
was conducted—namely, to improve the prediction of heavy precipitation events
over the northeastern U.S. through integration of academic research into useful
operational frameworks. In order to provide an appropriate background for this
particular project, the remainder of this chapter will document previous research
on the climatological and physical aspects of warm-season extreme rain events
(EREs), extratropical transition (ET), and the precipitation distribution occurring
both directly with, and well downstream of, landfalling and transitioning TCs.
1
1.2 Motivation and Overview
It has been documented recently that inland freshwater flooding is the
main weather-related cause of death when a TC makes landfall in the U.S.
(Rappaport 2000). Complicating this issue is the observation that numerous
recent landfalling and near-coastal-tracking TCs have been preceded by
coherent areas of heavy rainfall, some of which can produce their own flash
floods (M. Jurewicz and D. Vallee 2006, personal communication). Area forecast
discussions (AFDs) from the National Weather Service (NWS) weather forecast
office (WFO) in Taunton, MA, prior to the onset of PREs ahead of Katrina (2005)
and Ophelia (2005) suggest that these events pose a critical forecast challenge.
Timing and placement of such unexpected heavy rain can have disastrous
effects, as occurred when the PRE ahead of Frances (2004) deluged the New
York City subway system during rush hour on the morning of 8 September (Luo
et al. 2004). It is hypothesized that this lack of anticipation—and thus of
preparation—arises from a number of factors, including 1) the inability of current
numerical guidance to accurately depict these organized areas of heavy rain, 2)
the possibility that subtle features not easily detected may help focus heavy
rainfall in the moist tropical air mass ahead of a TC, and 3) the lack of attention
an operational forecaster may give toward predicting antecedent precipitation
when the rainfall evolution associated with the TC itself is perceived to be the
main concern. Consequently, a primary motivating factor of this research is to
increase awareness in the meteorological community of the potential for distinct
areas of heavy rainfall to occur far from the centers of TCs.
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Significant progress has been made in our understanding of the ET
process, but the operational forecaster is still faced with a number of challenges
when trying to predict the sensible weather that may result when a TC
approaches the midlatitudes. Jones et al. (2003) explains that these difficulties
arise mainly because one set of numerical prediction models is currently used in
the midlatitudes, while another set is used in the tropics. Neither of these sets
successfully captures the interaction of the two regimes with any consistency.
Recent CSTAR-related research, however, has successfully improved the
forecast of precipitation accompanying landfalling and transitioning TCs (D.
Vallee 2006, personal communication) through the creation of conceptual models
(DeLuca 2004; Srock 2005). These works, among others (Bosart and Carr 1978,
hereafter BC78; Jones et al. 2003), have noted that heavy rain can form when
the large-scale environmental flow allows tropical moisture to stream far ahead of
the TC itself. Only BC78, in detailing the excessive rain centered on Wellsville,
NY, far ahead of Agnes (1972), have provided a thorough investigation of this
type of event in the refereed literature to date (see section 1.4.3).
After consulting with a number of current operational meteorologists, the
author believes that many are aware of a small number of PREs, particularly if
they experienced them directly. However, no known comprehensive study of the
overall properties of PREs, nor of the mechanisms that work to bring them about,
has been conducted. This gap in understanding can have serious implications
both for areas that are directly affected by a TC and those that are not. The
current study will attempt to bridge this gap by providing forecasters with new
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climatological perspectives of PREs and detailed analyses of several cases. With
a reported increase in heavy precipitation events in the last century over the U.S.
(Karl and Knight 1998), and the suggestion that Atlantic TC activity will remain
high for the foreseeable future due to a favorable combination of warm sea
surface temperatures and reduced vertical wind shear (Goldenberg et al. 2001),
an extensive study of PREs needs to be undertaken now so forecasters can
better anticipate their occurrence.
1.3 Significant Warm-Season EREs
1.3.1 Overview
Since PREs are separate from the main rain shields associated with TCs,
a relevant background literature review should begin with a general discussion of
significant warm-season EREs. These events are associated commonly with
flash floods, which produce the most fatalities of any element of convective
storms (Doswell et al. 1996, hereafter D96; NOAA 2006) and cause more
property damage than all other weather-related phenomena in an average year
(Fritsch et al. 1998). A significant portion of these statistics likely can be
attributed to the lag in forecasting and warning of flash floods in comparison to
tornadoes (D96). Relatively poor predictability of warm-season in comparison to
cool-season precipitation (Funk 1991; Fritsch et al. 1998) arises from the
predominance of forcing for ascent by convective processes (Olson et al. 1995),
which often takes place near mesoscale convergence zones (Heideman and
Fritsch 1988; Koch and Ray 1996), as opposed to synoptic-scale forcing, which
4
is better portrayed by current numerical models. Flash floods are particularly
difficult to forecast due to the complex interaction of meteorological and
hydrological factors (D96). The occurrence and inadequate prediction and
warning of such events in turn has marked societal and economic impacts on
industries that rely heavily on accurate forecasts of significant precipitation, such
as government, business, agriculture, entertainment, and transportation (Fritsch
et al. 1998).
1.3.2 Climatologies
Heavy precipitation and flash flood climatologies have been constructed
for decades on both the national (La Rue and Younkin 1963; Maddox et al. 1979,
hereafter M79; Karl and Knight 1998; Brooks and Stensrud 2000; Schumacher
and Johnson 2005; Schumacher and Johnson 2006) and regional (Winkler 1988;
Houze et al. 1990; Giordano and Fritsch 1991; Bradley and Smith 1994; LaPenta
et al. 1995; Konrad 1997; Junker et al. 1999; Moore et al. 2003) levels. Recently,
Schumacher and Johnson (2006) classified an event as extreme if the total
rainfall received by a particular location exceeded its 50-yr recurrence amount,
thus limiting the study only to those events considered extreme for a given
region. Using this scheme, Schumacher and Johnson (2006) found that EREs
were most common within a southwest–northeast band from the southern Plains
through the Ohio Valley and Great Lakes states during the period 1999–2003
(see their Fig. 4). ERE occurrence in the southern part of this band peaked in late
spring with the predominance of mesoscale convective systems (MCSs),
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although there was a secondary maximum in early fall when significant synopticscale forcing was more likely (M79; Houze et al. 1990; Bradley and Smith 1994;
Schumacher and Johnson 2006). The vast majority of EREs in the north,
however, were due to MCSs developing in weakly forced environments (Bradley
and Smith 1994) during midsummer, which corresponds well with the overall
national peak in ERE occurrence (M79; Giordano and Fritsch 1991; Brooks and
Stensrud 2000).
The East Coast saw significantly fewer EREs during the period, partially
because the 50-yr recurrence amount for the East Coast is significantly higher
than that of areas farther inland (Fig. 1.1), indicating that heavy rainfall is more
common there overall. In particular, Fig. 1.1 shows that localized maxima in
recurrence amount are located along the Southeast U.S. coastlines bordering the
Atlantic Ocean and Gulf of Mexico, as well as the eastern slopes of the southern
and central Appalachians. EREs most commonly occur in this portion of the U.S.
in September due to the direct impacts of TCs (La Rue and Younkin 1963;
LaPenta et al. 1995; Brooks and Stensrud 2000) or their interactions with
synoptic-scale disturbances (LaPenta et al. 1995). Despite comprising less than
ten percent of all EREs east of the Rockies, these tropical-related events
produce the largest rainfall totals overall (Schumacher and Johnson 2006).
1.3.3 Meteorological Ingredients
1.3.3.1 General Considerations
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Numerous studies have discussed the utility of ingredients-based
methodologies to augment the output from numerical models when trying to
predict summertime EREs (e.g., Funk 1991; D96). D96 state that the largest
precipitation totals occur where the rainfall rate is highest for the longest amount
of time, thereby reducing this forecast problem to two main variables. Difficulties
arise, however, when trying to forecast the different factors that affect rainfall rate
and duration. For example, D96 state that high rainfall rates are generally
produced by a combination of rapid upward vertical motion, significant moisture
through a deep layer, and high precipitation efficiency. Long duration at any one
location is favored by slow system movement and upstream cell development
The remainder of this section will document previous work on the
meteorological processes that cause the aforementioned basic ingredients to
combine in midlatitudes, thus creating an environment favorable for EREs in the
warm season. The same factors combine to produce heavy precipitation in
tropical settings, but the manner in which they are brought together may differ
substantially from typical midlatitude situations (D96).
1.3.3.2 Favorable Synoptic-Scale Patterns
The high rainfall rates and slow movement contributing to flash floods east
of the Rockies described by D96 are usually associated with the elements of at
least one of three characteristic large-scale patterns defined by M79. The first is
called a synoptic flash flood pattern, in which convection breaks out in the warm,
moist air mass ahead of a quasi-stationary surface front often oriented from
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south-southwest to north-northeast (Fig. 1.2a). The slowing down or stalling of a
front with this orientation happens with some regularity along the Appalachians
(O’Handley and Bosart 1996), especially when the winds aloft have a significant
component parallel to the front (M79) and the mountain range (Schumacher et al.
1996), allowing convection to organize into a linear structure along the front (Dial
and Racy 2004).
Fig. 1.2b shows that the area at risk for the heaviest rainfall with the
synoptic setup is at the nose of a low-level jet streak as described by M79 and
Junker et al. (1999), but other studies (e.g., Elsner et al. 1989; Moore et al. 2003)
have found the heaviest rainfall near the left-exit region of the low-level jet streak.
Significant precipitation often results when either of these sectors of a low-level
jet streak becomes positioned ahead of a slow-moving midlevel trough (Fig.
1.2c). Sometimes the exit region of a low-level jet streak can couple with the
entrance region of an upper-level jet streak (Uccellini and Johnson 1979). For
example, Junker et al. (1999) noted the importance of the low-level jet streak in
transporting moisture toward a region of MCS development in the Midwest, but
also found that midwestern MCSs are collocated with the equatorward-entrance
region of an upper-level jet streak about 60% of the time. The presence of the
equatorward entrance region of an upper-level jet streak may allow heavy rain to
cover a more extensive area than it would if favorable jet dynamics were not
present (Kane et al. 1987).
Frontal flash flood events (M79) occur in the cool, moist air mass poleward
of a zonally oriented quasi-stationary surface boundary, sometimes near a weak
8
mesolow that can act to enhance low-level convergence (Fig. 1.3a). The
orientation of the low-level jet streak perpendicular to the surface front creates an
overrunning situation, where warm, moist air is transported above the cooler air
at the surface (Fig. 1.3b). The greatest chance for heavy rainfall occurs
downstream of a weak 500 hPa short-wave trough and along or just upstream of
a ridge axis extending northwest of a midlevel anticyclone (Fig. 1.3c). D96
hypothesize that the margins of such a ridge favor excessive precipitation
because of the tendency for sharp temperature and moisture boundaries to
collect there.
M79 term their third type of event a mesohigh flash flood, in which
convection prior to the ERE produces a nearly stationary outflow boundary. New
convection develops on the cool side of the outflow boundary southwest of the
mesohigh, in an area of significant low-level moisture that is often located ahead
of a slow-moving surface front (Fig. 1.4a). Heavy rain is favored further if the exit
region of the low-level jet streak becomes juxtaposed with the outflow boundary
(Fig. 1.4b) and a weak midlevel short-wave trough located upstream (Fig. 1.4c).
The orientation of the greatest flash flood threat near the large-scale ridge axis in
Fig. 1.4c is similar to the midlevel frontal pattern depicted in Fig. 1.3c. With the
risk of heaviest rain on the warm side of the synoptic-scale front, weak
subsidence on the periphery of the ridge may inhibit deep convection long
enough for moisture to accumulate there and for low-level diabatic heating to
increase the midlevel lapse rate and to decrease the vertical stability (D96).
9
Heavy rainfall may occur over one area in distinct episodes with the
mesohigh type of setup in M79, increasing the potential for flash flooding. The
initial rain serves to saturate the underlying soil, allowing for significant runoff of
the subsequent heavy rainfall. Multiple episodes of heavy rain over one area can
hinder the utility of NWS flash-flood guidance and make for an especially
challenging meteorological and hydrological forecast, especially if severe
weather is also present (Schwartz et al. 1990).
Winkler (1988) identified an additional large-scale pattern conducive to
flash flooding termed Type IV (Fig. 1.5) when she examined events over
Minnesota. In this case, a narrow moisture ridge at the surface intersects a
midlevel ridge axis with weak flow aloft, leading to a spatially confined, slowmoving ERE of short duration. With the upper-level trough axis still well to the
west, the focus for precipitation comes from a weak surface boundary in the
moist air mass.
1.3.3.3 Synoptic-Scale and Mesoscale Forcing for Ascent
Synoptic-scale forcing for ascent is usually diagnosed by some form of the
quasigeostrophic (QG) omega equation. Using the traditional form (e.g.,
Bluestein 1992, section 5.6; Holton 2004, section 6.4), previous studies have
noted that summertime EREs often are associated with low-level warm-air
advection (e.g., Junker et al. 1999), indicated by veering of the wind with height,
and/or weak differential cyclonic vorticity advection (e.g., BC78; Junker et al.
1999). Nevertheless, ascent brought about by synoptic-scale processes is
10
usually too slow to raise a parcel of air to its level of free convection (LFC) or
produce observed ERE rainfall rates (Doswell 1987). Many EREs therefore occur
when mesoscale forcing for ascent dominates over or augments synoptic-scale
forcing for ascent (Doswell 1987; Fritsch et al. 1998; Junker et al. 1999). D96
suggest that the background lift provided by synoptic-scale QG forcing mostly
aids mesocale processes by moistening and destabilizing the atmosphere near
where EREs occur.
Low-level convergence can provide the lift needed to bring air parcels to
their LFC and result in rapid ascent, especially in the warm season (Heideman
and Fritsch 1988). Summertime convergence zones along which heavy
precipitation can form span many different scales and types. Only half of all
summertime rainfall in the U.S. occurs in proximity to the deep synoptic-scale
fronts present in the M79 synoptic and frontal flash flood composites (Heideman
and Fritsch 1988). Excluding rainfall from TCs, other, more subtle, triggering
mechanisms can include thunderstorm outflow boundaries (M79), sea-breeze
fronts (Heideman and Fritsch 1988; Koch and Ray 1996; Fovell 2005), and
coastal fronts (Bosart et al. 1972; Appel et al. 2005), as well as prefrontal
troughs, airmass boundaries, and heat troughs (Koch and Ray 1996). It is often
challenging to detect such triggering mechanisms using traditional surface
mesoanalyses; radar and satellite imagery have proven to be more effective
(Koch and Ray 1996).
If low-level confluence associated with one of the types of boundaries
mentioned above acts on a preexisting thermal gradient, a region of
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frontogenesis will be produced, and the thermal gradient may intensify. The
frontogenesis produces a direct thermal circulation, which acts to enhance
vertical motion near the front (Moore et al. 2003). Frontogenesis also can arise
due to evaporational cooling when precipitation falls and/or to the juxtaposition of
cloudy and clear air, leading to differential radiative heating (Langmaid and
Riordan 1998). The diabatic cooling caused by falling precipitation and reduced
radiative heating also can work to strengthen cold-air damming along the
Appalachian Mountains (Fritsch et al. 1992).
1.3.3.4 Forcing for Ascent by Physiographic Processes
Orographic processes can be the only source of lift necessary in areas of
large terrain gradients (Barros and Kuligowski 1998), such as the EREs that led
to the Big Thompson and Rapid City (Maddox et al. 1978) and Fort Collins
(Petersen et al. 1999) flash floods. These EREs occurred in upslope flow regions
in the Rocky Mountains, in which weak differential cyclonic vorticity advection
caused by subtle short waves rotating around a mid- to upper-level trough
upstream supplied a background environment of weak synoptic-scale forcing. A
“bentback ridge” (Maddox et al. 1978) just to the east contributed to deep
southeasterly flow against the mountains. This setup is consistent with the M79
frontal and mesohigh composites, but orographic lift was the main mechanism
responsible for raising parcels in a moist, conditionally unstable air mass below a
marked temperature inversion to their LFC. A further result of the deep
southeasterlies in these cases was to anchor precipitation cells along the
12
foothills, which caused long periods of heavy rain and exacerbated the severity of
the flooding.
Heavy rain events with significant orographic upward motion also can take
place in proximity to the Appalachians, as occurred in the central part of the
mountain chain in January 1996 (Barros and Kuligowski 1998) and in western
Virginia in June 1995 (Pontrelli et al. 1999). However, if the orographic uplift is
not sufficient to bring parcels to their LFC, Barros and Kuligowski (1998) noted
that the elevated terrain can still produce or enhance precipitation via the
transport of warm, moist air over a cold pool that collects against the topographic
barrier (indicative of cold-air damming), or through the seeder–feeder mechanism
(see also, e.g., Bell and Bosart 1988).
Another feature dependent on regional physiography that can act to
produce or enhance precipitation is the coastal front (Bosart et al. 1972). Srock
(2005, section 1.5.2) thoroughly reviews previous research on this topic, so only
a brief summary will be included here. Coastal fronts arise due to land–sea
temperature contrasts, orography, and the shape of the coastline, and are
observed most frequently in winter along the coasts of New England, North
Carolina, and Texas (Bosart et al. 1972; Bosart 1975). They also can arise due to
frictional convergence near the coastline produced by gradients in surface
roughness between the land and the sea (Roeloffzen 1986). Coastal fronts tend
to act like warm fronts, such that overrunning enhances precipitation on their cold
sides (Bosart et al. 1972; Marks and Austin 1979; Srock 2005).
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More recent studies have examined the occurrence of warm-season
coastal fronts (e.g., DeLuca 2004; Appel et al. 2005; Srock 2005). The results
indicate that the mechanisms by which they form are similar to the cool season,
but the temperature contrasts along warm-season coastal fronts are weaker than
their wintertime counterparts. Nevertheless, coastal fronts still can act to enhance
precipitation significantly, especially if they are induced ahead of polewardmoving TCs (DeLuca 2004; Srock 2005). In researching summertime coastal
fronts in North Carolina, Appel et al. (2005) found that most are diurnal in nature,
remain just offshore, and are infrequently associated with cold-air damming.
They can develop in response to a mesoscale atmospheric circulation on the
western edge of the Gulf Stream, and are usually observed when an anticyclone
resides nearby or just east of the North Carolina coast.
1.3.3.5 Moisture, Instability, and Shear
Significant moisture through a deep tropospheric layer is necessary to
sustain moist convection capable of producing EREs and flash floods (e.g., M79).
To illustrate this point, M79 calculated that the mean precipitable water for each
of their composites was greater than 37 mm. More recently, Konrad (1997)
concluded that a ridge axis in the 700 hPa mixing ratio pattern was the most
common ingredient associated with heavy rainfall over the southeastern U.S.,
even though the magnitude of the atmospheric moisture content showed no
direct relationship with observed rainfall amounts. Furthermore, Funk (2003)
explained that the source of the moisture must be considered in addition to its
14
magnitude. If the moisture is transported from a maritime tropical air mass
characterized by cumulus clouds containing many small drops and a high density
and size distribution, higher rainfall totals than expected may occur.
D96 noted several reasons why deep moisture was so important to heavy
rain development, including that it 1) results in rainfall rates directly proportional
to the precipitable water of the column through increased condensation, 2)
enhances precipitation efficiency through the reduction of dry-air entrainment,
which promotes evaporation, and 3) creates a relatively low LFC, which allows
buoyant processes to increase updraft strength if the parcel can be lifted to that
level.
Convection can be either surface-based or elevated in nature. Surfacebased convection typically occurs in the warm sector ahead of a frontal
boundary, as in the M79 synoptic and mesohigh flash flood composites. Positive
convective available potential energy (CAPE) and a negative lifted index on an
atmospheric sounding indicate the potential for surface-based convection. On a
plan-view map, warm-sector convection often develops near a low-level θe ridge
axis (Funk 1991), which can be used as a proxy for CAPE (Schwartz et al. 1990).
It also has been shown that the vertical distribution of positive area in a long,
narrow manner reduces parcel vertical velocity and acceleration (Blanchard
1998). The resultant increase in parcel residence time within the warm cloud
layer promotes greater precipitation efficiency due to the collision–coalescence
process (Beard and Ochs 1993).
15
Moore et al. (2003) conducted an extensive study of the features that
bring about elevated convection, which are summarized in Fig. 1.6. This type of
activity occurs on the cool side of a front as warm, moist air is transported over
the surface boundary by a low-level jet streak (Colman 1990a,b; Moore et al.
2003; Schumacher and Johnson 2005), making the environment favorable for
elevated convection strikingly similar to the M79 frontal flash flood composite.
Deep moisture and the equatorward entrance region of an upper-level jet streak
also are common features, as previously discussed in this subsection and in
subsection 1.3.3.2, respectively. Potential instability [indicated by equivalent
potential temperature (θe) decreasing with height] and CAPE can exist in
elevated layers above the frontal inversion (Moore et al. 2003), with low-level
frontogenesis (Colman 1990b) and/or warm-air advection (Moore et al. 2003)
typically providing the lifting mechanism necessary to release the elevated
instability. It can be inferred from Henry (1988) that the Showalter index and K
index are better-suited for evaluating elevated instability that may promote
flooding rains than surface-based CAPE and the lifted index. The Showalter
index and K index for the composites in M79 were less than −2 and greater than
35, respectively, indicating at least an 80% chance for thunderstorm
development utilizing the forecasting rule of thumb proposed by Henry (1988).
Diagnosis of the potential for heavy convective precipitation often begins
with an evaluation of moisture content and instability (D96), but surface moisture
flux convergence (MFC) also can be an important aid in predicting heavy rainfall
location and extent up to three hours in advance (Banacos and Schultz 2005).
16
MFC has a long history of use in ERE case studies (e.g., BC78; Elsner et al.
1989) because it can be helpful for identifying mesoscale boundaries, such as
those described in section 1.3.3.3, when they can be resolved by gridded surface
analyses. However, use of MFC assumes that the dominant lifting mechanisms
are not located significantly above the boundary layer and that CAPE and
convective inhibition are favorable for deep, moist convection (Banacos and
Schultz 2005). Junker et al. (1999) discussed that MFC also is evident in the exit
region of an 850 hPa jet streak that transports moisture toward an area of heavy
rainfall, but it is significant spatial extent and duration of MFC that distinguishes
large EREs from smaller ones with lesser rainfall rates.
Another ingredient that is important to assess when predicting warmseason EREs is the vertical wind shear, which can diagnose the anticipated
organization of the convection and the system motion vector. EREs often are
associated with weak-to-moderate speed shear, which tends to organize the
convection into linear structures (Houze et al. 1990; D96; Schumacher and
Johnson 2005) and allow individual convective elements to move slowly. If, in
addition to slow system movement, the most favorable ingredients for cell
redevelopment discussed previously are located upstream from the active
convection, training echoes will be observable on radar, and the duration of
heavy rainfall at any one location will be prolonged (D96). Flash flooding can
occur in precipitating systems that form in strongly sheared environments, but
either the rainfall rates must be exceptionally high or significant upstream
17
redevelopment must take place for an extended period of time to counterbalance
the faster movement of individual cells.
Significant veering of the wind with height has been positively correlated
with rainfall amounts occurring with EREs in the southeastern U.S. (Konrad
1997). The deep (but not necessarily strong) warm-air advection this veering
implies is most common in frontal flash flood events because of the orientation of
the low-level jet streak with respect to the midlevel ridge axis (see Figs. 1.3b,c).
Synoptic flash flood events are associated more commonly with veering winds
only in the lowest layers of the troposphere and unidirectional speed shear aloft
(see Figs. 1.2b,c).
1.4 Precipitation Distribution Associated with Landfalling and Transitioning TCs
1.4.1 Background on ET
TCs have profound effects on many regions of the world, as they bring
high wind, heavy rain, and storm surge (Jones et al. 2003), as well as sporadic
tornadoes (Edwards and Pietrycha 2006) with them, particularly to coastal and
maritime locations. When TCs make landfall, they weaken as the land surface
cuts off the surface heat and moisture fluxes that drive them over the warm
ocean. They may undergo ET, regardless of whether or not they make landfall,
as they encounter the midlatitude westerlies, but the degree of transition and
redevelopment as an extratropical cyclone varies on a case-by-case basis.
Summaries on the current understanding of ET are presented in Jones et
al. (2003), DeLuca (2004), and Srock (2005), so only a brief discussion will be
18
given here. The two-step process of ET illustrated by Fig. 1.7 (Jones et al. 2003)
first outlines the environmental changes encountered by a TC when it
approaches the midlatitudes. These may include, but are not necessarily limited
to, stronger temperature and moisture gradients, increased vertical shear, lower
SSTs, and a larger Coriolis parameter. The TC responds through structural
changes, most notably the development of significant asymmetries in the
temperature, wind, and moisture fields. The overall intensity of the transitioning
TC decreases, but its areal extent expands. The TC then may interact with an
upper-level trough or low-level baroclinic zone if either is in sufficient proximity, or
simply decay in the Tropics without completing ET if neither of these midlatitude
features is present. When the TC becomes an extratropical cyclone, the ET
process takes an average of 36 h (Evans and Hart 2003).
Potential reintensification as an extratropical cyclone is the subject of a
third stage in the ET model proposed by Klein et al. (2000), although
reintensification is not necessary to have extreme impacts on land- and oceanbased interests (Jones et al. 2003). Redevelopment is often observed when a
midlatitude cyclonic circulation is located northwest of the TC (Harr and Elsberry
2000; Harr et al. 2000) at an average distance of 1700 km (Foley and Hanstrum
1994), but can occur more gradually with less intensification if the midlatitude
system is northeast of the TC (Harr and Elsberry 2000; Harr et al. 2000). The
strength of the subsequent extratropical cyclone generally depends on the
degree of interaction between the TC and upper-level dynamical and low-level
thermodynamic processes (Klein et al. 2002).
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1.4.2 Precipitation Directly Associated with a TC
1.4.2.1 Large-Scale Considerations
Forecasting rainfall amount, location, and duration is one of the most
complicated tasks associated with a landfalling or transitioning TC (Jones et al.
2003).
A successful quantitative precipitation forecast requires accurate
prediction of TC track, intensity, structural changes, and interactions with
midlatitude circulations, which often are represented inaccurately by numerical
models. Unreliable predictions can result in excessive rainfall amounts occurring
in unexpected locations, causing significant loss of life and property damage.
Heavy rainfall is usually symmetric about the TC track in mature storms
(Frank 1977; Marks 1985; Rodgers et al. 1994), although asymmetries may
occur depending on the environmental shear in the vicinity of the storm, as well
as the TC motion (Chen et al. 2006). A left-of-track shift in the rainfall distribution
commonly occurs at higher latitudes (Atallah et al. 2007) as the TC undergoes
ET and interacts with a midlatitude circulation (Jones et al. 2003). The relevant
features involved in this process are illustrated schematically in Fig. 1.8a by
Atallah et al. (2007) using the potential vorticity framework (Hoskins et al. 1985).
The approach of a strong, positively tilted midlatitude trough enhances the lowlevel thermal gradient north and west of the TC, gradually altering the
mechanisms forcing ascent from symmetric diabatic heating to the differential
vorticity and temperature advections featured in QG theory. Additionally,
downstream ridging induced by diabatic heating can enhance the upper-level
divergence occurring within the equatorward entrance region of a southwesterly
20
upper-level jet streak, leading to more favorable jet dynamics and an increased
likelihood of intense precipitation, sometimes well downstream of the TC center
(Atallah and Bosart 2003; Jones et al. 2003; Atallah et al. 2007). Atallah et al.
(2007) explained that the effects of diabatic heating can be inferred qualitatively
by the nonconservation of potential temperature on the dynamic tropopause.
A right-of-track precipitation distribution occurs more readily at lower
latitudes than at higher latitudes as the TC interacts with a well-defined east–
west oriented ridge that has likely been enhanced by diabatic heating and upperlevel outflow from the TC (Fig. 1.8b; Atallah et al. 2007). In this case, the
upstream trough is much broader and weaker, and the forcing for vertical motion
through vorticity advection by the thermal wind shifts to the northeast of the TC.
Most storms that exhibit this shift make landfall along the Gulf Coast and then
track along the Appalachians, which suggests qualitatively that upslope flow
along the eastern slopes will promote a right-of-track rainfall pattern. This type of
rainfall pattern tends to occur with TCs of relatively small spatial scale, though
the precipitation distribution often shifts from right-of-track to left-of-track as the
TCs continue moving poleward (Atallah et al. 2007).
1.4.2.2 Small-Scale Considerations
The mesoscale and physiographic features discussed as potential foci for
heavy warm-season rainfall in sections 1.3.3.3 and 1.3.3.4, respectively, also can
act to enhance precipitation directly within or out ahead of a TC (DeLuca 2004;
Srock 2005). Specifically, Srock (2005) discusses how coastal fronts and
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topographic barriers can modulate the precipitation distribution, especially in the
southeastern U.S. Mesoscale coastal fronts played key roles in augmenting
rainfall rates with the approach of Agnes (1972) (Bosart and Dean 1991), Marco
(1990) (Srock 2005), and Bob (1991) (DeLuca 2004). Floyd (1999) is an example
of how an intense coastal front associated with a troposphere-deep baroclinic
zone can act in concert with the large-scale factors discussed in section 1.4.2.1
to produce an extensive shield of rain spanning a considerable distance from the
TC center (Atallah and Bosart 2003).
Orographic contributions to rainfall have been documented in connection
with numerous Atlantic basin storms impinging upon the Appalachians, such as
Jerry (1995) (Srock 2005), and smaller-scale mountain ranges of the
northeastern U.S. (DeLuca 2004). Sinclair (1993) documented Cyclone Bola
(1988), which deposited copious amounts of rainfall in upslope regions of New
Zealand. In general, any time a TC circulation causes a moist airstream to
intersect a topographic barrier, regardless of the distance from the TC center,
there is the potential for enhanced rainfall.
1.4.3 Heavy Precipitation Well Downstream of a TC
Jones et al. (2003) included a short section in their review paper about the
potential for areas of heavy rain separate from the main TC circulation to develop
well ahead of the center of the TC, even if it never makes landfall. Noting that this
rain is often not anticipated because of its distance from the TC center, they
described the increased risk of flooding that can occur if the heavy rainfall directly
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associated with the TC falls over the same region as the antecedent
precipitation.
Awareness that heavy rainfall can break out well ahead of a TC is not
new, although there has been relatively little work reported in the refereed
scientific literature on this type of event. Peripheral awareness of this possibility
dates back to at least 1955, when Mook (1955) described a long line of
convergence that extended well northeast of a weakening Diane (1955). This
line, which DeLuca (2004) suggested may have been a weak coastal front,
spawned distinct areas of heavy rain in Pennsylvania and Connecticut that were
partially enhanced by orographic lift, and another near Boston that was
associated with a weak cyclonic circulation that developed along the
convergence line.
The most in-depth published analysis of a PRE was on the episode
centered near Wellsville, NY, far ahead of Agnes (1972) (BC78). The rainfall
distribution in Fig. 1.9a shows that the region in New York occurred as a
separate entity from the heavy precipitation directly associated with Agnes. BC78
found that the inland flooding occurred 1) within a midlevel confluence zone
between the tropical and midlatitude airstreams (Fig. 1.9b), 2) in an environment
characterized by weak differential cyclonic vorticity advection, and 3) as a narrow
tongue of deep tropical moisture was transported well poleward of the TC
through a low-level ridge line. The heavy rain in New York eventually became
embedded within a surface pressure trough as a small-scale cyclonic circulation
began to develop just north of the PRE.
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Subsequent studies of Agnes (DiMego and Bosart 1982a,b; Bosart and
Dean 1991) continued to document heavy rainfall in advance of the TC in
addition to the Wellsville, NY, heavy rains. They found that the highest
precipitation rates were concentrated on the windward slopes of the
Appalachians in the presence of orographic lift and moisture, as well as along
and just west of a weak coastal front extending well to the north of Agnes. The
space–time cross section from Bosart and Dean (1991) shown in Fig. 1.10
demonstrates that the passage of the coastal front was marked by a significant
wind shift from easterly to north-northwesterly, an increase in wind speed, and a
temperature drop of approximately 3°C, coincident with the onset of the most
intense rainfall rates.
Vigorous ascent likely resulted from a weak,
frontogenetical direct thermal circulation in the presence of a moist tropical air
mass characterized by a moist–neutral stability profile.
An upper-level jet streak combined with low-level convergence and
overrunning along a TC-induced θe gradient to produce significant rainfall in
advance of Marco (1990) (Srock 2005). A coastal front in eastern North Carolina
ahead of Floyd (1999) helped focus heavy rainfall there as the storm approached
(Atallah and Bosart 2003; Colle 2003). The tropical disturbance that eventually
would become Leslie (2000) interacted with a stalled frontal boundary in southern
Florida to produce extremely heavy rains (Franklin et al. 2001). In studying the
accuracy of the Z–R relationships used by the NWS with the Weather
Surveillance Radar-1988 Doppler (WSR-88D), Ulbrich and Lee (2002) found that
two periods of heavy precipitation preceded the arrival of rainfall directly
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associated with Helene (2000) in upstate South Carolina: one associated with the
initial passage of a cold front (Fig. 1.11a), and another poleward of the front as it
became quasi-stationary and oriented east–west (Fig. 1.11b).
Finally, Lin et al. (2001) have found that heavy precipitation with a
significant orographic component preceded the arrival of a tropical depression
(TD) in 1959 and Tropical Storm Rachel (1999) in Taiwan, as well as another TD
in Japan in 1974. They concluded that the role of the TCs was to enhance the
moist low-level jet streak impinging upon the steep mountain slopes, which aided
in the release of conditional instability. The proximity of a slow-moving synopticscale system, such as a quasi-stationary front, helped slow the movement of the
convective systems and prolong the rain.
1.5 Goals and Organization of the Thesis
This thesis aspires to marry the previous research presented here on
EREs and the precipitation distribution associated with TCs to show how the
approach of a TC at great distances can help bring about the ingredients
necessary for the formation of PREs. An extensive climatology and several
detailed case studies will be used to provide forecasters with a practical
framework for better predicting the occurrence of these PREs, with an eye
toward continuing the recent trend of successful incorporation of CSTAR
research into NWS operations.
The thesis will be organized as follows. Chapter 2 will provide an overview
of the data sources, demonstrate the PRE identification process, and explain the
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various methods used in constructing the climatology, the three case studies,
and the null case study. Chapter 3 will present the first known statistical and
composite climatology of PREs, with an emphasis on their overall properties and
the main synoptic-scale features associated with them. Chapter 4 will document
the governing synoptic-scale and mesoscale dynamics and thermodynamics
associated with the three selected case studies and one null case. Chapter 5 will
relate the findings of the case studies to the literature review and suggest a
forecast strategy for PREs by comparing the different case studies among
themselves and with the null case. Technology transfer of this research into NWS
operations also will be discussed in this chapter. Chapter 6 will conclude with a
synthesis of this research and suggestions for future work.
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