Heavy Precipitation Associated with Southern ... - AMS Journals

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Heavy Precipitation Associated with Southern Appalachian Cold-Air Damming and Carolina Coastal Frontogenesis in Advance of Weak Landfalling Tropical Storm Marco (1990) ALAN F. SROCK AND LANCE F. BOSART University at Albany, State University of New York, Albany, New York (Manuscript received 30 September 2008, in final form 22 January 2009) ABSTRACT An analysis is presented of Tropical Storm Marco (1990), a storm that dropped copious amounts of rain over the southeast United States. Marco was noteworthy because of its role in the formation and evolution of two distinct episodes of cold-air damming and coastal frontogenesis over Georgia and the Carolinas. These mesoscale features led to greater than 300 mm of precipitation in 2 days over the near-coastal southeast United States; much of the rain occurred while Marco was over 400 km away. This case is further complicated by two other nearby tropical cyclones, which affected Marco’s track and the overall rainfall distribution. Synoptic and mesoscale analyses of the development of the coastal front and cold-air damming episodes show that the location of Marco helped to orient low-level winds toward the Appalachians. As rain developed inland, a pocket of relatively cool air, the ‘‘cool pool,’’ formed near the mountain slopes and was partially blocked by the higher terrain. Low-level analyses show that the coastal front on the oceanward edge of the cool pool became a focusing mechanism for ascent and precipitation, as moist, tropical air advected inland by Marco was forced upward at the density gradient. The results indicate that a weak tropical cyclone can directly effectuate intense precipitation distant from the storm center, both by causing moist tropical flow toward land and by inducing mesoscale features that focus the precipitation and lead to heavy rainfall and flooding.

1. Introduction a. Purpose Landfalling and near-landfalling tropical cyclones (TCs) pose a challenging forecast problem because of the potential devastation to people and property from storm surges, strong winds, and heavy precipitation (e.g., Sheets 1990; Rappaport 2000). Although storm surges and high winds tend to be weaker when distant from the storm, regions of heavy precipitation associated with a TC can occur hundreds of kilometers away from the storm center. In the case of landfalling and near-landfalling TCs, topography-related mesoscale features can combine with enhanced forcing for ascent and moisture from synoptic features and the TC to cause extremely damaging, heavy rainfall over land. Rappaport (2000) states that inland precipitation and associated freshwater Corresponding author address: Alan F. Srock, Dept. of Atmospheric and Environmental Sciences, University at Albany/SUNY, Albany, NY 12222. E-mail: [email protected] DOI: 10.1175/2009MWR2819.1 Ó 2009 American Meteorological Society

flooding from landfalling TCs are often the most dangerous threat because of the loss of life and the destruction of property. Further study into the effect of surface winds, wind shifts, and land/ocean airmass contrasts on the precipitation distribution of TCs was suggested as a way to improve understanding and forecast skill by the Fifth Prospectus Development Team of the U.S. Weather Research Program (PDT-5; Marks et al. 1998). Nearshore surface features can modify the final TC precipitation distribution through mesoscale effects such as orographically forced ascent (upslope), coastal frontogenesis, and cold-air damming (CAD). This paper will examine the formation and enhancement of coastal fronts (CFs) and CAD events induced by the presence of a TC in proper position with respect to the local terrain. The period of 9–13 October 1990 was chosen for study because two distinct episodes of coastal frontogenesis and CAD caused exceptionally heavy precipitation near the southeast U.S. coast; these mesoscale features were enhanced by Tropical Storm Marco in the eastern Gulf of Mexico. Marco was the primary TC affecting the

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FIG. 1. Topography (shaded; m), significant station location, cross-section line (dashed), and storm positions every 6 h (see legend for symbol reference) from 0000 UTC 9 Oct 1990 to 1200 UTC 13 Oct 1990. The TC position at 1200 UTC on a given date is shown by an open symbol next to the date.

Southeast’s rainfall during the period, but the remnants of Hurricane Klaus and the rapid approach of Hurricane Lili from the east near the end of the period greatly affected the final rainfall pattern (Fig. 1 shows the tracks of the TCs). Marco and the remnants of Klaus were responsible for over 500 mm of precipitation near the Georgia–South Carolina border, $57 million in damage (in 1990 dollars), and seven deaths (all due to inland flooding) from the heavy rainfall (Mayfield and Lawrence 1991). The rainfall distribution with this case is uncommon compared to other landfalling TC events, since much of the heaviest rainfall in coastal Georgia and the Carolinas occurred while Marco was . 400 km away. Figure 2 shows the National Centers for Environmental Prediction/ Hydrological Prediction Center (NCEP/HPC) archived 24-h accumulated precipitation ending at 1200 UTC 11 October and 1200 UTC 12 October, along with Marco’s position at the end of each period (Figs. 2a,b, respectively). The heaviest rainfall in both periods is located between the Appalachian Mountains and the coast; however, the highest rainfall totals in the 24-h period ending at 1200 UTC 11 October appear over nearcoastal South Carolina, 12 h before Marco makes landfall in the Florida Panhandle. In this paper, we will show that the presence and location of Marco were crucial for the development of two CAD episodes east

of the southern Appalachians, each coincident with a shallow, intense CF over near-coastal Georgia and South Carolina.

b. Previous work Topographic effects can significantly enhance precipitation in any rainfall-producing system. Orographic precipitation enhancement without a CF or CAD is generally related to upslope effects (e.g., Passarelli and Boehme 1983; Barros and Kuligowski 1998). Intentionally selecting cases without CFs, Passarelli and Boehme (1983) found that regions of upslope precipitation received 20%–60% more rainfall than nearby flat or downslope terrain. Focusing solely on precipitation directly attributable to a landfalling TC, Haggard et al. (1973) looked at 70 yr of TCs that made landfall in the United States and subsequently traversed the Appalachian Mountains (defined as the TC center passing over elevation greater than 300 m). The authors found TCs that crossed high topography usually had precipitation maxima in the areas of sharpest elevation increase. Cold-air damming (e.g., Richwein 1980; Forbes et al. 1987; Bell and Bosart 1988) refers to the process where cold air is slowed by a topographic barrier and moves

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FIG. 2. The 24-h accumulated precipitation (contoured every 25 mm starting at 50 mm; contours end at 150 mm) ending at (a) 1200 UTC 11 Oct 1990 and (b) 1200 UTC 12 Oct 1990. The location of Marco at the end of the time period is denoted by the M. Adapted from archived image at NCEP/HPC (courtesy of W. Junker).

preferentially in the direction of the pressure gradient force. This occurs most often on the eastern side of approximately north–south-oriented mountains with a high pressure system to the north, as cold air moving toward the eastern slopes has insufficient kinetic energy to go over the barrier and is then forced to decelerate. The deceleration weakens the effect of the Coriolis force, so the cold air will preferentially move away from the higher pressure to the north, creating an equatorward bulge of high pressure and low temperature (e.g., Bailey et al. 2003, their Fig. 1). Forbes et al. (1987) showed that the presence of the cold pool was important for the location and type of precipitation. As warmer, moister air from the Atlantic Ocean was advected toward the mountains, it was lifted up and over the density gradient at the edge of the cold pool, instead of at the mountainside topography gradient as would be expected with terrain-driven ascent alone.

Bell and Bosart (1988) found that CAD east of the Appalachians most frequently occurs in late fall and early winter, when the temperature difference is greatest between the warm ocean and cold land. Their study also found a low-level wind maximum oriented parallel to the mountains in the cold air, which helped to push the cold dome equatorward. Fritsch et al. (1992) suggested that the cold pool could be maintained by the blocking of incoming solar radiation due to cloud cover over the cold dome and evaporative cooling due to precipitation falling through the cold dome; however, Brennan et al. (2003) found evaporative cooling in the cold air only helped to maintain the temperature deficit while the cold pool was unsaturated. To further categorize CAD events, Bailey et al. (2003) objectively classified members of a CAD climatology. Their primary division between CAD types depended on formation and maintenance method—whether the polar

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cold air was advected into the region by equatorward synoptic forcing, cooled in situ due to heavy rain and evaporative cooling, or a combination of both. Most of their cases showed a clear signal of high pressure to the north or northeast of the mountains, but, in the few cases contained in their last category (unclassifiable), a primary feature leading to CAD development was often a low pressure to the south. Coastal fronts (e.g., Bosart et al. 1972; Bosart 1975; Nielsen 1989) are thermal and density airmass boundaries that form primarily as a result of land–sea temperature contrasts. Usually forming approximately 12 h before the passage of a coastal low pressure system, CF locations and intensities are affected by ‘‘orography, coastal configuration, land–sea temperature contrast and friction’’ (Bosart et al. 1972). Coastal fronts are usually shallow features (e.g., Nielsen and Neilley 1990), and CAD can assist in the strengthening of the CF gradient if the cold pool extends close enough to the coastline. However, Nielsen and Neilley (1990) point out that CAD and CF events can and often do occur without a significant presence of the other. Riordan (1990) examined a CF that formed off the coast of the Carolinas over the sea surface temperature gradient at the edge of the Gulf Stream. As this CF moved inland, the temperature gradient intensified as the CF interacted with an alreadyexisting CAD event against the Appalachians, which led to especially heavy precipitation in the region. Coastal fronts are also often found directly ahead of a TCs track. Bosart and Carr (1978) and Bosart and Dean (1991) studied the heavy precipitation ahead of Hurricane Agnes (1972), which occurred up to one day ahead of Agnes on a line stretching from Virginia to New York. The authors noted a surface convergence line coincident with a strong temperature and dewpoint gradient, driven primarily by land–sea thermal contrasts. This CF was directly attributable to the tropical cyclone, forming near the shore and continuing ahead of the TC as Agnes moved north. Atallah and Bosart (2003) and Colle (2003) examined aspects of the precipitation distribution of Hurricane Floyd (1999) through synoptic and modeling analyses; these studies found that precipitation ahead of Floyd’s track was generally enhanced along the CF from approximately 12 h before through the time of storm passage. Cote (2007) discussed predecessor rain events (PREs), defined as rainfall . 100 mm in 24 h ahead of a TC. A PRE’s primary moisture source must be advected poleward by the TC, but the lifting mechanism and precipitation are not directly attributable to the TC. His work showed that the average distance from TC to PRE is 935 km, with a ;36-h time lag before TC passage. Only some of the cases in Cote (2007) involved a CF, but the pres-

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ence of a CF along with the poleward moisture transport was shown to focus precipitation at and inland of the CF. Precipitation ahead of a landfalling and recurving TC track can also be greatly modified by extratropical transition (ET; e.g., Bosart and Carr 1978; Bosart and Lackmann 1995; Klein et al. 2000; Jones et al. 2003; Evans and Prater-Mayes 2004). The ET process occurs as a TC morphs from a warm-core to cold-core cyclone and gains extratropical cyclone characteristics (e.g., wind field asymmetries and fronts). An upstream trough and downstream ridge are necessary components for ET, although the relative strength of the TC vorticity and the upstream vorticity in the trough can vary. During ET, precipitation tends to shift toward left of track, as enhanced ascent ahead of the upstream trough shifts the region of strongest precipitation toward the approaching trough (Atallah et al. 2007). A good synopsis of the current understanding of ET is presented in Jones et al. (2003, see especially their Fig. 11). Binary interaction (e.g., Fujiwhara 1921, 1923; Brand 1970) also plays a significant role in Marco’s evolution, given the proximity of Hurricanes Klaus and Lili. Dritschel and Waugh (1992) and Prieto et al. (2003) studied the effect of distance, storm radii, and relative vorticity strength on different forms of binary interaction. These studies have shown that cyclonic vorticity centers tend to orbit each other cyclonically; however, the type of interaction can vary greatly (see especially Prieto et al. 2003, their Fig. 2). The paper will be organized as follows. Section 2 will discuss the data used for the study. Section 3 will present a synoptic overview of Marco and the surrounding environment, while section 4 will look at the mesoscale effects (i.e., CAD and CF) that led to the enhanced rainfall. Section 5 will discuss these results and offer conclusions and suggestions for future work.

2. Data sources The primary sources for gridded data are the 40-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40; Uppala et al. 2005), with 1.1258 horizontal grid spacing, 22 vertical levels (pressure coordinates), and 6-h temporal resolution, and the NCEP North American Regional Reanalysis (NARR; Mesinger et al. 2006), which has ;32-km horizontal grid spacing and 29 vertical levels, with 3-h temporal resolution. Both reanalyses were compared with observations to ensure the validity of any conclusions drawn using the data. Considering the relatively coarse grid spacing of the ERA-40, the exact positioning and values of near-surface fields and derived quantities like vertical motion should be treated with some

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skepticism; however, relative maxima and minima still provide useful information. The ERA-40 proved superior for locating the tropical cyclones and for large-scale flow features, but comparison of the reanalyses with observed soundings and surface data suggests that the NARR better represented the state of the atmosphere on the mesoscale in the CF and CAD regions. Surface data were collected from three sources at NCAR: the NCEP Automated Data Processing Global Surface Observations (ADP), the U.S. Air Force Data Surface Airways, version 3 (DATSAV3), and the International Comprehensive Ocean–Atmosphere Dataset (ICOADS); duplicate stations were removed from the file. The National Hurricane Center (NHC) best-track dataset [i.e., the Atlantic basin hurricane database (HURDAT)] was used for tropical cyclone positions while the TCs were contained in the dataset; note that Lili was classified as a subtropical storm until 0000 UTC 11 October. Also, Klaus was not tracked in HURDAT after 1200 UTC 9 October, so its position for the next 24 h was subjectively determined using low-level vorticity from the NARR and ERA-40 and nearby surface observations. After 1200 UTC 10 October, Klaus’s location was indeterminable using available data. Other data sources include archived radiosonde data from the National Climatic Data Center (NCDC) and visible satellite images from the University at Albany archive. The General Meteorological Package (GEMPAK; Koch et al. 1983) was used for all data archival and most of the plotting routines. To facilitate calculation of differences and gradients using surface data, the irregularly positioned surface observations were interpolated using a Barnes analysis (e.g., Barnes 1964) in GEMPAK to a 0.38 3 0.38 horizontal grid. The interpolation was only performed on the land- and buoy-based observations (ADP and DATSAV3), since the ship data from the ICOADS were too sparse and erratic to include at the grid spacing desired over land for objective analysis. Thus, surface data figures will plot station observations and subjectively analyzed contours using both land and marine observations, but objectively analyzed features will be shaded using only the ADP and DATSAV3 data.

3. Synoptic (large scale) setup a. Case overview Marco officially reached tropical storm strength over northern Cuba at 1200 UTC 9 October. At that time, the tightest mean sea level pressure (MSLP) and 10002500hPa thickness gradients were located just ahead of the midlatitude trough over the western Great Plains (Fig. 3a). A surface high, which dominated the central Atlantic at

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previous times (not shown), had moved off to the northeast as the precursor to Lili began to travel westward under the ridge. By 1200 UTC 10 October, the primary midlatitude surface cyclone was over Indiana and Ohio, associated with the thermal trough over the Midwest, whereas the sharpest 1000–500-hPa thickness gradient was located west of the Appalachians, stretching southwestward to the Mississippi Delta (Fig. 3b). A weak CAD signature was present in the central Carolinas, as shown by a MSLP maximum, which bulged southwestward parallel to the mountain axis. At the same time, an area of 700-hPa ascent ,24.0 3 1023 hPa s21 was over near-coastal Florida, Georgia, and South Carolina. Marco had moved over the Florida Keys, and a tightening of the pressure gradient to its east and northeast over the previous 24 h suggested an even stronger feed of tropical air. Precipitation was heavy in parts of central Georgia and South Carolina during this period, as shown by the accumulated rainfall at Columbia (CAE), South Carolina, Augusta (AGS), Georgia, and Athens (AHN), Georgia (Fig. 4). By 1200 UTC 11 October, the dominant 700-hPa ascent maximum was located ahead of the synoptic-scale thermal trough stretching from North Carolina to Ontario (Fig. 3c). The offshore shift of the 700-hPa ascent maximum near South Carolina suggests that precipitation in the most heavily flooded areas (over nearcoastal Georgia and South Carolina) waned during this period. Radar summaries (not shown) and accumulated precipitation (Fig. 4) at the time confirm that the strongest echoes and rainfall had moved toward the coast and offshore, and would remain there over the next 12 h. At 1200 UTC 11 October, Marco was near its peak intensity just off the western coast of Florida, barely grazing the coast for its first landfall (Fig. 1; e.g., Mayfield and Lawrence 1991). Marco made a second landfall at TD strength near 0000 UTC 12 October, and had reached central Georgia by 1200 UTC 12 October (Fig. 1). Rainfall had returned to central Georgia and South Carolina by 1200 UTC 12 October (see Fig. 4), as shown by two distinct 700-hPa ascent maxima within 300 km of Marco: one just off the northern Georgia coast and the other over higher terrain in northern Georgia (Fig. 3d). Marco weakened further after this time, and the associated ascent and precipitation induced by the TC decreased rapidly as well.

b. Large-scale environment and key features At 1200 UTC 9 October, a southwesterly jet .75 m s21 on the dynamic tropopause (DT; e.g., Hoskins et al. 1985; Nielsen-Gammon 2001) stretched from Texas to Hudson Bay on the downstream side of the associated DT potentially cold (cyclonic vorticity) anomaly (Fig. 5a).

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FIG. 3. ERA-40 MSLP (contoured, hPa), 1000–500-hPa thickness (dashed; dam), and 700-hPa vertical motion (shaded; 31023 hPa s21) at 1200 UTC for (a) 9 Oct, (b) 10 Oct, (c) 11 Oct, and (d) 12 Oct 1990. The positions of Klaus, Lili, and Marco are indicated by using the symbol convention in Fig. 1.

The equatorward entrance region of this jet was located over a cold front with the same orientation (not shown), stretching southwestward from the low-level cyclone over Lake Erie (Fig. 5a). Farther downstream, potential temperatures on the DT near Marco and Klaus were . 360 K, and nearby wind shear between the DT and 850-hPa levels was anticyclonic (not shown), suggesting a warm-core structure through the depth of the troposphere, consistent with 1000–500-hPa thickness values .576 dam (Fig. 3a). The poleward advection of this warm air would have likely aided in building the ridge that was developing over the western Atlantic. The ERA-40 analyzed stronger low-level vorticity to the west of Klaus’s estimated position at 1200 UTC 10 October (Fig. 5b); this may be related to Klaus advecting too far west in the model representation, vorticity elongation stretching northward from

Marco, or a combination of both. Marco and the remnants of Klaus moved primarily northward over the next 24 h (Fig. 1), although Klaus shows a much more pronounced westward path. Klaus’s westward shift was likely partially due to a binary interaction; Marco was far stronger than Klaus at this time, and the TCs were within 1000 km of each other, suggesting Klaus was undergoing some form of straining out while moving cyclonically around Marco (e.g., Prieto et al. 2003). A visible satellite image from 1831 UTC 10 October shows the strong tropical moisture feed east of Florida that curved cyclonically around Marco’s northern and eastern flank and brought warm, moist air to the coastal Carolina and Georgia region (Fig. 6a). Backward trajectories calculated from the ERA-40 at approximately the same time show the source region of the air at 850 hPa over coastal South Carolina and Georgia was located

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FIG. 4. The 3-h accumulated precipitation (mm) ending at the period listed for CAE (black bars), AGS (gray bars), and AHN (white bars) for periods ending 1200 UTC 10 Oct–1800 UTC 12 Oct 1990. Data are from the NCDC hourly precipitation dataset.

over the tropical Atlantic between 950 and 1000 hPa 48 h earlier (Fig. 7a). These trajectories approximately follow the track of Klaus (Fig. 1), and are consistent with a stream of air with precipitable water (PW) values .55 mm that had reached Georgia and South Carolina at 1800 UTC 10 October. For comparison, the mean precipitable water derived from soundings at Charleston, South Carolina for the month of October (from 1948 to 2005) is approximately 33 mm, whereas the plus-two standard deviation-level for the same period is ;50 mm (see http://www.crh.noaa.gov/unr/include/pw.php?sid5chs); this shows the anomalously high moisture available for precipitation, especially during the early stages of this event. Extrapolation of Klaus’s track indicates that the moist air from Klaus would have reached the southeast U.S. coast around 1800 UTC 10 October, suggesting that the especially high PW values near the coast are partially due to the remnant moist air of the decaying TC. Backward trajectories from 0000 UTC 11 October show a continuation of the tropical feed, although the flow at the time was more southerly (Fig. 7b). Precipitation between 1200 UTC 10 October and 1200 UTC 11 October was exceptionally heavy (greater than 150 mm) over coastal Georgia and South Carolina (Fig. 2a); thus, much of the decrease in precipitable water between 1800 UTC 10 October and 1200 UTC 11 October would likely be explained by heavy rainfall over the southeast United States. By 1200 UTC 11 October, the primary low-level vorticity maximum associated with the midlatitude trough had moved north of the Great Lakes, but the DT trough still stretched equatorward to the Gulf of Mexico (Fig. 5c). The low-level vorticity maximum stretching from North

Carolina to Pennsylvania formed from a combination of vorticity advection over the Appalachians (not shown) and lee troughing on the eastern slopes. Visible satellite from 1331 UTC 11 October shows the cloud shield above Marco and coastal Georgia and South Carolina, as well as the continued moisture feed from the tropics (Fig. 6b). Backward trajectories from 1200 UTC 11 October confirm the near-surface, tropical source with PW values .45 mm advecting around the east side of Marco toward the Georgia coast (Fig. 7c). However, PW values at the Georgia coast are nearly 10 mm less at 1200 UTC 11 October than at 1800 UTC 10 October; since both low-level trajectories show a south to southeast (tropical) source region, the difference in PW near the coast was most likely due to the added moisture associated with Klaus at the former time. At 1200 UTC 12 October, Marco was located southeast of the upper-level jet associated with the synoptic trough to the west and the associated positive PV maximum (Fig. 5d). Jones et al. (2003) stated that a TC just downstream of an upper-level jet and positive PV maximum is in a favorable position for ET, but Marco did not redevelop or accelerate poleward over the next 24 h. Instead, Marco slowed its northward motion and turned to the east. Marco’s northward deceleration was partially a result of the upper-level trough lifting out and not picking up the TC, whereas its eastward shift was likely related to another binary interaction with Lili, which was moving rapidly westward (Fig. 1). Both TCs had a northward component of motion after this time, but the weakening Marco moved southeastward relative to (or cyclonically around) Lili. Therefore,

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FIG. 5. ERA-40 dynamic tropopause potential temperature (shaded; K), winds (m s21; full barb is 5 m s21 and pennant is 25 m s21), and 925–850-hPa mean relative vorticity (contoured every 0.5 3 1024 s21 starting at 0.5 3 1024 s21) at 1200 UTC for (a) 9 Oct, (b) 10 Oct, (c) 11 Oct, and (d) 12 Oct 1990. The positions of the TCs are as in Fig. 3.

instead of potentially undergoing ET and rapidly lifting out of the region with the trough to the north, Marco’s poleward progression slowed, which kept the storm over South Carolina. Visible satellite imagery at 1331 UTC 12 October showed good vertical extent in the cloudiness just ahead of Marco (Fig. 6c). The southwestern sharp back edge of convection near the Georgia–South Carolina border was primarily due to the advection of low-level cold air around Marco from west of the Appalachians (discussed further in section 4). Backward trajectories from 1200 UTC 12 October again show a tropical, near-surface source region for the air over coastal South Carolina and Georgia, with PW .45 mm (Fig. 7d). Again, Marco’s circulation sustained the tropical moisture feed along the storm’s eastern side, which led to the continued heavy rainfall.

Most of the synoptic features discussed in this section suggest a favorable environment for precipitation over Georgia and the Carolinas, but these features alone cannot fully account for the extremely heavy and localized rainfall with this case. The next section will focus on the low-level mesoscale features that led to the local extreme enhancement of precipitation.

4. Mesoscale analysis a. Surface observations To locate the CAD/CF events, surface observations and meteograms will be used to show the formation and development of the integral mesoscale features. Although a front should be technically based solely on a thermal and/or density gradient (e.g., Sanders 1999), a dynamical definition may also assist in locating the

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FIG. 6. Visible satellite images from University at Albany archive for (a) 1831 UTC 10 Oct, (b) 1331 UTC 11 Oct, and (c) 1331 UTC 12 Oct 1990. The position of Marco is given by M at 1800 UTC 10 Oct, 1200 UTC 11 Oct, and 1200 UTC 12 Oct 1990, respectively.

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FIG. 7. ERA-40 PW (shaded above 25 mm), 850-hPa winds (m s21, where full barb is 5 m s21 and pennant is 25 m s21), and 48-h back trajectories from 850-hPa starting at (a) 1800 UTC 10 Oct, (b) 0000 UTC 11 Oct, (c) 1200 UTC 11 Oct, and (d) 1200 UTC 12 Oct 1990. Trajectory shadings (blue, purple, gray, and black) represent the 850–900-, 900–950-, 950–1000-, and .1000-hPa layers, respectively. The TC positions are as in Fig. 3.

boundary. For the thermal definition of a front, surface virtual potential temperature (uy) is used, since it eliminates effects from varying elevation and is directly related to density. For the dynamical definition, we will use the horizontal velocity gradient tensor, which accounts for convergence, deformation, and vorticity (VTEN; e.g., Stonitsch and Markowski 2007). The velocity gradient tensor will be large when horizontal convergence, confluence, or vorticity is large in the area; since all of these can be significant in frontal zones, regions of large VTEN suggest that the dynamical aspects of fronts may be present in the area. A region with a tight uy gradient and large VTEN collocated would suggest the potential for a convectively active frontal area. As the first significant CF/CAD event began around 1200 UTC 10 October, a region of high pressure and low

uy began to bulge southwestward to the east of the Appalachians, with a strong uy gradient at the oceanward edge of the potentially cold pool (Figs. 8a,b). Winds blew almost directly onshore at Georgia and South Carolina coastal stations, whereas inland winds were oriented roughly parallel to the mountains. A .10 K (100 km)21 uy gradient was located to the west of the Appalachians, associated with a cold front trailing southward from the synoptic cyclone to the north (see Figs. 3b and 5b), but remained west of the mountains throughout both CF/CAD events. A cyclonic wind shift was collocated with the ;8 K (100 km)21 coastal uy gradient in coastal North and South Carolina, while another cyclonic wind shift was located near the ;4 K (100 km)21 uy gradient in southeast Georgia and northeast Florida. Although the intensities of both wind shifts had not

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FIG. 8. Observed surface analysis of (a) MSLP (contoured; hPa), temperature (dashed; 8C), and standard station observations at 1200 UTC 10 Oct; and (b) virtual potential temperature (plotted and contoured; K), velocity gradient tensor (objectively shaded every 2 3 1025 s21 starting at 4 3 1025 s21), and winds (m s21, where full barb is 5 m s21 and pennant is 25 m s21) at 1200 UTC 10 Oct 1990. The TC positions are as in Fig. 3.

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reached the minimum threshold for VTEN shading, the combination of the uy gradient and wind shift suggest a CF-type boundary was forming in both locations. By 1800 UTC 10 October, the temperature at Charleston (CHS), South Carolina has increased 28C in 6 h, whereas the temperature at CAE remained nearly constant over the same period (Fig. 9a; see Fig. 1 for station locations). Sea level isobars from central North Carolina to northeastern Georgia bulged to the southwest, highlighting the relatively high pressure air partially blocked by the mountains; winds in this region were oriented nearly parallel to the mountain axis, as expected with CAD. Along the coast, the uy gradient in southeastern North Carolina had weakened over the previous 6 h (Fig. 9b), suggesting that differential diabatic heating due to the lack of cloud cover over land had weakened the northern end of the boundary (Fig. 6a). However, over Georgia and South Carolina, cloudiness and precipitation limited the inland heating and maintained a stronger uy gradient (e.g., Fritsch et al. 1992). In southern Georgia and northeastern Florida, rainfall and cloudiness over land helped to maintain the cold pool and tighten the uy gradient, while the collocated VTEN increased to .6 3 1025 s21 over the same 6-h period. At 0000 UTC 11 October, the temperature and dewpoint at CHS continued to rise during the day, whereas both decreased slightly at CAE (Fig. 10a). The combination of higher pressure to the northeast and Marco to the south-southwest helped to maintain geostrophic flow toward the Appalachians, and heavy rain continued inland in the Carolinas. The overlap of the uy gradient and the VTEN maximum was especially apparent at this time over near-coastal Georgia and South Carolina, attesting to the strong thermal and dynamical structure of this CF. Since extrapolation of Klaus’s track placed the decayed TC inland over the Carolinas by 0000 UTC 11 October, the exceptionally moist air associated with the former TC (PW . 55 mm, 6 h earlier; Fig. 7a) was forced up and over the CF, leading to enhanced rainfall. Soon after this period, both the uy gradient and VTEN weakened in Georgia and the Carolinas, and neither reinvigorated significantly until ;0000 UTC 12 October (not shown). At the peak of the second CF/CAD event (0600 UTC 12 October), dynamical forcing took on a larger role, as shown by the large swath of VTEN . 4 3 1025 s21 ahead of Marco (Fig. 11). During this second CF/CAD event, the CF boundary was located farther inland, likely because it was advected closer to the mountains by the cyclonic flow around Marco. Winds inland paralleled the mountains and were oriented nearly perpendicular to isobars, suggesting that ageostrophic flow was dominant and that CAD was an important feature. This second

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CAD/CF event was shorter lived and more transitory, as the both the tightest uy gradient and highest VTEN values stayed just ahead of Marco and remained oriented nearly parallel to the mountains. Previous CF research has emphasized a strong CF event and intense precipitation 12–24 h ahead of the storm passage along the storm track (e.g., Bosart and Dean 1991; Atallah and Bosart 2003; Colle 2003); the development of the second CF event just ahead of Marco shows similar characteristics. Cooler and drier continental air was beginning to advect toward Marco’s southern flank from west of the Appalachians around 1200 UTC 12 October (Fig. 12), limiting the primary convection to the eastern and northern sides of the weakening storm as Marco continued to move northeast (Fig. 6c). Cooling from clearing over eastern North Carolina helped to tighten the uy gradient at the coast, but daytime heating weakened this gradient over the next 12 h (not shown). The uy gradient and VTEN ahead of Marco also weakened as Marco dissipated after this time. To further examine the temporal evolution of the two intense CF/CAD periods, meteograms were constructed for CHS, CAE, AHN, and AGS (Fig. 13). At CHS (Fig. 13a), winds were generally from the southeast throughout both CF events (approximately 1200 UTC 10 October–0600 UTC 11 October and 0000–1200 UTC 12 October). Winds at CHS had a northerly component from 1800 UTC 11 October to 0000 UTC 12 October; during this period, radar summaries position the primary convection offshore, and only CHS (near the coast) reported significant rainfall accumulation (not shown). Inland at AHN, rainfall was generally light through the first event and only strong for a few hours of the second CF event (Figs. 4 and 13b). During both intense CF events, winds at AHN had a northeasterly component, suggesting that AHN was in the cold dome for both CAD events. The temperature and dewpoint remained nearly constant at AHN and stayed colder than CHS throughout the period, suggestive of a persistent surface boundary between these two stations. Two stations in the center of the cold pool highlight the two intense periods of heavy precipitation. At CAE (Fig. 13c), moderate and heavy rain began just after 1200 UTC 10 October and continued though 0700 UTC 11 October. Light, intermittent precipitation followed until 0000 UTC 12 October; then, after a few hours of moderate rain, died off again until Marco moved nearly overhead around 1500 UTC 12 October. Observations were similar at AGS (Fig. 13d), where precipitation started just after 1200 UTC 10 October, weakened by 0600 UTC 11 October and was intermittent until 0000 UTC 12 October, when moderate rain and thunderstorms were reported (cf. the heavy precipitation recorded between

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FIG. 9. As in Fig. 8, but for 1800 UTC 10 Oct 1990.

0600 and 1200 UTC 12 October in Fig. 4). During both CF/CAD events, winds at CAE and AGS were from the northeast, whereas temperatures averaged ;38C colder than CHS but only slightly warmer than AHN, illustrat-

ing that both CAE and AGS were within the cold pool. The two distinct periods of significant precipitation were reported at both stations in the center of the cold pool while the CF boundaries were most intense over land.

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b. Vertical structure To determine the vertical structure of the CAD/CF events, a high-resolution, three-dimensional dataset was

needed to measure the depth of the cold air. The NARR was chosen to supplement sounding and surface observations for examination of this case, as it had the best resolution and the most accurate low-level winds of

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the available datasets. Comparison of a plan view of 975-hPa frontogenesis from the NARR (Fig. 14a) with frontogenesis calculated from surface observations (Fig. 14b) at 0000 UTC 11 October shows the two pri-

mary regions of frontogenesis roughly collocated (the synoptic cold front and CF). The NARR uy gradient just southeast of the Appalachians was nearly 8 K (100 km)21 stronger than observed, primarily because of the

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especially low uy pool of air in central Georgia. This thermal analysis discrepancy explains the higher frontogenesis values at the coast in the NARR representation. Except for the temperature in the cold pool, the

general structure of the thermal and height (pressure) fields are sufficiently similar to suggest that qualitative analyses can be performed using NARR low-level fields at 0000 UTC 11 October. Comparisons of other times

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FIG. 13. Meteograms of surface observations from 0000 UTC 10 Oct to 1800 UTC 12 Oct 1990 of temperature (8C), dewpoint (8C), MSLP (hPa), winds (m s21, where full barb is 5 m s21 and pennant is 25 m s21), and present weather for (a) CHS, (b) AHN, (c) CAE, and (d) AGS.

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FIG. 14. (a) NARR 975-hPa geopotential heights (contoured; m), virtual potential temperature (dashed; K), horizontal frontogenesis [shaded beginning at 2 K (3 h)21 (100 km)21], and winds (m s21, where full barb is 5 m s21 and pennant is 25 m s21) at 0000 UTC 11 Oct 1990. (b) Surface MSLP (contoured; hPa), virtual potential temperature (dashed; K), horizontal frontogenesis [shaded beginning at 2 K (3 h)21 (100 km)21] and winds (m s21, where full barb is 5 m s21 and pennant is 25 m s21) at 0000 UTC 11 Oct 1990.

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FIG. 15. Comparison of NARR (solid blue) and observed (dashed red) soundings for (a) CHS and (b) AHN at 0000 UTC 11 Oct 1990. Winds are plotted using same convention as in Fig. 14.

and locations shows that the NARR was too cool in the cold dome in both events, which led to slightly stronger winds parallel to the mountains as the along-mountain pressure gradient intensified (not shown). NARR-derived soundings are compared with observed soundings at 0000 UTC 11 October to show the validity of NARR vertical profiles of temperature, dewpoint, and wind. At CHS (Fig. 15a), the thermodynamic and wind fields show good qualitative and quantitative agreement, though the near-surface air in the NARR is slightly cooler and drier than in the observations. The NARR and observations at AHN show good qualitative agreement above 550 hPa but have important differences at low levels (Fig. 15b). The primary discrepancy is the temperature and dewpoint below 900 hPa; the NARR temperature in the lowest 100 hPa is ;58C colder than observations. Above the cold dome, the NARR and observed winds generally agree within ;5 m s21 and ;208 in orientation throughout the depth of the troposphere, excepting a significant westerly shift between 550 and 750 hPa in the observed data, suggestive of warm-air advection in that layer. Comparison of other soundings with the NARR at other times show that the cold dome is ;58C too cool during both CAD events (not shown). However, except for the thermal structure of the cold dome, the NARR seems to

represent the vertical structure of the atmosphere sufficiently well throughout this case. Cross sections from the NARR showing the low-level structure of the two CAD/CF events are shown in Fig. 16 (see Fig. 1 for cross-section line). At 1200 UTC 10 October, when the first CF/CAD episode began, the near-surface wind shift boundary was coincident with the eastern edge of the cold pool and an area of frontogenesis .5 K (100 km)21 (3 h)21 (Fig. 16a; cf. Fig. 8b). This near-surface frontogenesis caused enhanced forcing for ascent at the boundary, which combined with background isentropic lift to induce heavy precipitation at and just inland of the leading edge of the cold dome. This temperature and wind structure continued through 1800 UTC 10 October and 0000 UTC 11 October (Figs. 16b,c). Though the NARR cold pool became too cold during this event (as noted earlier), the area above the cold dome between the mountains and the coast was still dominated by ascent, with an ascent maximum directly above the surface horizontal uy gradient and frontogenesis maximum. Thus, one would have expected the strongest precipitation at and just inland of the CF (e.g., CAE, Figs. 4 and 13c), whereas stations farther inland would have accumulated far less rain during this period (e.g., AHN; Figs. 4 and 13b). The top of the cold pool was approximately the height of the

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FIG. 16. NARR cross sections of virtual potential temperature (contoured; K), vertical motion (dashed; 31023 hPa s21, only contouring ascent), horizontal winds (m s21, where full barb is 5 m s21 and pennant is 25 m s21), and horizontal frontogenesis [shaded beginning at 2.5 K (3 h)21 (100 km)21] at (a) 1200 UTC 10 Oct, (b) 1800 UTC 10 Oct, (c) 0000 UTC 11 Oct, and (d) 0600 UTC 12 Oct 1990. The cross-section line is shown in Fig. 1; the approximate location of the coastline is delineated by the triangle below the plot.

higher terrain to the northwest, so flow retardation due to CAD was still occurring. The low-level air is likely too stable in the NARR representation, since the lapse rate at the top of the cold pool is not as strong in the observed thermal vertical structure; however, given the thermal structure indicated in the observed soundings, some flow blocking would be expected. Winds continued to blow parallel to the Appalachians in the cold dome to the southeast of the ridge axis, whereas southerly and southeasterly flow dominated both the offshore flow and the flow above the cold dome. A cross section through the peak of the second CF event shows a similar structure at 0600 UTC 12 October (Fig. 16d). The core of the frontogenesis was closer to the mountains in this case, as the CF was advected farther inland by the easterly flow around Marco. The cross section at this time sliced through the core of the maximum NARR ascent, which was farther inland and above the strongest low-level frontogenesis. The second CF event is more transitory as the low-level density

boundary and ascent maximum move north-northeastward ahead of Marco (not shown).

5. Discussion and conclusions A case study of the development of CAD and associated CFs induced and enhanced by nearby Tropical Storm Marco has been presented. In this case, the presence, location, and track of Marco, along with Marco’s interaction with both Hurricanes Klaus and Lili, were necessary ingredients for heavy precipitation in near-coastal Georgia and the Carolinas. By helping to induce and enhance two distinct CAD/CF events, relatively weak TCs were able to cause significant damage through copious amounts of rainfall and resultant flooding. Cold-air damming was important because it shifted the primary core of ascent toward the CF at the oceanward edge of the cold dome. The development and maintenance of these CAD events are noteworthy because the pressure gradient against the southern Appalachians was

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primarily caused by a low to the south of the mountains. Although there was a fairly strong high over the central Atlantic at the beginning of the first CAD/CF event (around 1200 UTC 10 October), its location alone was not favorable for the ideal pressure gradient against the southern Appalachians. As Marco developed and moved into the eastern Gulf of Mexico, the pressure gradient tightened and became more orthogonal to the spine of the Appalachians. Without Marco reorienting and intensifying the pressure gradient, the development of this CAD event would have likely been far weaker. Branick et al. (1988) discussed a mesoscale convective system which traveled eastward out of New Mexico and Texas toward the Georgia and Florida coasts. They noted that CAD formed in the Carolinas due to pressure falls induced by the passage of a squall line through Georgia. Marco caused a similar development of CAD (predominantly induced by lower pressures to the south), but was even more damaging to the Carolinas because of the tropical moisture advection toward and over the cold dome. Traditionally, CAD has been considered a cold-season phenomenon, where a high pressure system to the north led to geostrophic winds oriented toward the mountains. With a dominant northern high, the cold dome often develops through the advection of cold air as well as in situ cooling (e.g., Bailey et al. 2003). However, especially in the first CAD/CF event in this study, the poleward source of cold air (through advection) is a minor factor, since much of the flow directed at the mountains has tropical or near-tropical origins. Cooling in central South Carolina and near-coastal Georgia in the clear air overnight preceding 1200 UTC 10 October likely combined with latent cooling from nearby precipitation to create the initial pocket of cooler air, which was then maintained by the continuous rain and lack of solar heating due to cloud cover during the first heavy rainfall episode. This conclusion can be verified by examining the development of the cold dome; the difference between uy inside and outside of the cold pool is small before precipitation begins (not shown), but magnifies with time because the onshore, oceanic flow warms during the day (see Figs. 8b, 9b, and 10b). CAD development without a dominant poleward high would be expected far less frequently than with a cold poleward high, since it requires both a strong low equatorward of the mountains and mechanisms to diabatically cool and then maintain the cool air in the blocked region (e.g., radiative cooling, evaporative cooling from precipitation, cloud cover). For this reason, a cold dome in a CAD event primarily caused by a low to the south would likely be better termed a ‘‘cool pool,’’ since it is relatively cold in comparison with its surroundings, but significantly warmer than the cold pool in a classic wintertime CAD event.

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The heaviest precipitation in this case fell primarily within 200 km of the coast (Figs. 2a,b). As the cold dome developed and the boundary at the coastal edge evolved into a CF, moist tropical onshore flow was forced up and over the cold dome. The CF, just inland from the coast, acted as a focusing mechanism for ascent during each event. Mesoscale analyses of uy gradient and VTEN highlighted both the thermodynamic and dynamic features of the CF. The heaviest precipitation (Fig. 2) and intense flooding noted over near-coastal Georgia and South Carolina (Mayfield and Lawrence 1991) were collocated with the strongest ascent and most favorable VTEN and uy gradient. The NARR reasonably replicated both CAD/CF episodes, and its analysis showed the expected strong ascent at and just inland of the CF (Figs. 16a–d). Unlike the results from Haggard et al. (1973), which found the heaviest TC rainfall in regions of flow ascending over higher terrain, the heaviest precipitation with Marco was shifted away from the mountains by the CAD/CF episodes. Thus, the position, longevity, and intensity of the two CAD/CF events were important for the overall precipitation location and distribution. The moist remnants from Hurricane Klaus also added to the heavy rainfall in the first CAD/CF event. The PW difference between the ERA-40 at 1800 UTC 10 October (Fig. 7a) and 1200 UTC 11 October (Fig. 7c) shows a decrease in PW of ;10 mm in the core of the tropical moisture plume over the southeast United States. As the remnants of Klaus moved over coastal Georgia and South Carolina, the especially moist air would have been forced up and over the leading edge of the CF boundary, creating additional rain and helping to maintain the cold dome. If Klaus was not present, heavy rainfall would probably still have been recorded in Georgia and the Carolinas, but Klaus’s additional moisture most likely amplified the final precipitation totals. Previous research has discussed the heavy precipitation associated with landfalling TCs; however, many of those studies have focused on the heaviest rainfall occurring just ahead and within 12–24 h of storm passage (e.g., Bosart and Dean 1991; Atallah and Bosart 2003; Colle 2003). The results of these studies show a similar evolution and rainfall distribution to the second CAD/ CF event studied herein. Bosart and Carr (1978) and Cote (2007) examined cases of enhanced precipitation well downstream of the TC. In both studies, poleward moisture transport on the eastern side of the TC was the primary source of rain. Cote (2007), looking at storms affecting the East Coast, stated that the average PRE forms nearly 1000 km away and 36 h ahead of TC passage. The first CAD/CF event with Marco could be considered a form of PRE because the additional moisture over 400 km and 36 h away from Marco was

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FIG. 17. Conceptual model of a cold-air damming and coastal front event induced by a nearby tropical cyclone. Important features are denoted on the diagram.

advected there primarily by the TC. However, this case differs from the traditional PRE because the augmented precipitation was not induced by midlatitude, synopticscale features, but was rather focused to a specific region by CAD and coastal frontogenesis, mesoscale phenomena directly induced by the nearby TC. These results highlight the need to consider weak TCs as possible heavy rain producers, even at a distance. Marco (and the remnant moisture from Klaus) inundated near-coastal Georgia and South Carolina with precipitation. Figure 17 shows a conceptual model that includes some of the key features in this event. A tropical cyclone west of Florida suggests geostrophic flow oriented toward the Appalachian Mountains, with moist tropical air advected around the TC’s eastern side toward the southeast U.S. coast. If a thermal boundary or CF already exists at the coastline (possibly formed by differential diabatic heating or land/sea temperature contrasts), the tropical air would be forced to rise at the boundary and would lead to heavy precipitation at and just inland of the CF. Precipitation which fell on the inland side of the CF would evaporatively cool the air until saturation, and would help to decrease the temperature in the cool pool. The more dense, colder air would experience flow blocking, which would in turn enhance CAD and cause a push of cold air equatorward

on the leeward side of the mountains. Although this type of CAD/CF development is uncommon, the potential for significant precipitation well ahead of a nearby TC should be considered if flow blocking is possible against nearby topography. Acknowledgments. Funding for this work was provided by NSF Grants ATM-0304254 and ATM-0553017, as well as an AMS 21st Century Campaign Graduate Fellowship. The authors thank Dr. Anantha Aiyyer and Jason Cordeira for their assistance in developing the backward trajectory code, Celeste Iovinella for assorted technical help with the manuscript preparation, and Nicholas Metz for many useful discussions throughout the research process. The authors also thank two reviewers for their comments on the manuscript. Much of the data were provided by the National Climatic Data Center, the ECMWF, and the Data Support Section of the Computational and Information Systems Laboratory at NCAR. NCAR is supported by grants from the National Science Foundation. REFERENCES Atallah, E. H., and L. F. Bosart, 2003: The extratropical transition and precipitation distribution of Hurricane Floyd (1999). Mon. Wea. Rev., 131, 1063–1081.

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