5. Case Studies
As previously mentioned in Section 2.4, we used 0.5° CFSR data to create maps
highlighting prominent synoptic and mesoscale features and important dynamical
processes during the 3–4 Jan 1999 (Type G) and 11–12 Dec 2008 (Type BC) ice storms.
In order to describe the local thermodynamic environment and the thermodynamic
evolution of air parcels entering each ice storm area, we also generated sounding profiles
from radiosonde data and calculated backward trajectories with the NOAA HYSPLIT
model. The results for each case are discussed separately below.
5.1 3–4 Jan 1999 Ice Storm
5.1.1 Overview
The 3–4 Jan 1999 ice storm impacted 11 CWAs and 149 counties between West
Virginia and Maine (Fig. 5.1). Local storm reports indicate that icing amounts generally
ranged from 0.25 in (0.64 cm) to 0.50 in (1.27 cm), but some locations received nearly
1.00 in (2.54 cm). According to Storm Data, significant ice accumulations caused more
than 75,000 customers to lose power in the northeastern U.S. (NCDC 1999). Tree
damage and downed power lines were most prevalent throughout northern Virginia,
western Pennsylvania, northeastern Pennsylvania, and east-central New York. The
heaviest freezing rain fell across the Allegheny Plateau, as well as along and just east of
the central and northern Appalachians (excluding central Pennsylvania, where most
precipitation fell as snow and sleet). The spatial characteristics of the ice storm region
suggest that topography played a crucial role during this event.
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5.1.2 Synoptic Evolution
Figures 5.2–5.9 illustrate the evolution of the synoptic-scale circulation and
associated thermal boundaries and moisture transport between 1200 UTC 2 Jan 1999 and
0000 UTC 4 Jan 1999. At 1200 UTC 2 Jan, an inverted pressure trough in the sea level
pressure (SLP) field is located over the Mississippi Valley, near the right entrance region
of a 300-hPa jet streak (Fig. 5.2). Downstream of the pressure trough, a thickness ridge
coincident with a region of southeasterly geostrophic flow and implied low-level warm
advection overspreads the Ohio Valley. Across eastern New England, we see the
remnants of a synoptic-scale thickness trough corresponding to a prior cold-air intrusion
from southern Canada. A second thickness trough, independent of the synoptic-scale
thickness trough, is positioned just east of a narrow pressure ridge between northern
Georgia and central Virginia. The juxtaposition of the pressure trough with high pressure
over the interior Northeast results in a strong southwest-to-northeast pressure gradient
across the Appalachians, and thereby establishes a favorable synoptic environment for
“hybrid” cold-air damming (Bailey et al. 2003). During the initial stages of the 3–4 Jan
1999 ice storm, southerly flow in the 850–700-hPa layer allows for low- to midlevel
transport of very warm, moist air between the Gulf of Mexico and the Tennessee and
Ohio Valleys (Fig. 5.3). Consequently, a large zonal moisture gradient develops across
the Appalachians, with PW values greater than 24 mm over western Kentucky and
Tennessee and PW values less than 16 mm over North Carolina and Virginia. Low PW
and equivalent potential temperature (θe) values east of the Appalachians are
representative of a continental polar (cP) mass, which becomes reinforced near the
surface by cold-air damming.
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As time progresses, a mature surface cyclone evolves from the existing pressure
trough and gradually intensifies as it tracks west of the Appalachians. At 0000 UTC 3
Jan, the low center is located beneath a narrow, meridionally oriented jet streak between
the Gulf Coast and northern Michigan (Fig. 5.4). The associated thickness ridge has
rapidly pushed northeastward along the western slopes of the Appalachians, but persistent
cold-air damming deters the eastward progression of the warm front, and the thickness
trough remains locked to the eastern slopes between western North Carolina and northern
Virginia. Consequently, the pressure ridge strengthens, and an inverted pressure trough
forms along the East Coast due to the eastward displacement of the maritime tropical
(mT) air mass emanating from the Gulf of Mexico. This pressure trough is coincident
with a region of enhanced southerly 850–700-hPa winds, which allow warm, moist air to
overspread the southeastern U.S. and eventually reach the mid-Atlantic and New England
states (Fig. 5.5). Meanwhile, strong southerly low- to midlevel flow continues to support
poleward moisture transport and warm advection west of the Appalachians.
By 1200 UTC 3 Jan 1999, the primary low has entered the Great Lakes region and
deepened to 996 hPa (Fig. 5.6). The dissipation of the thermal trough and associated
pressure ridge and the passage of a frontal trough suggest that cold-air damming has
ceased entirely along the central and southern Appalachians. As suggested by the
meridionally elongated surface trough, secondary cyclogenesis is also occurring over the
mid-Atlantic region. Rauber et al. (2001b) noted that secondary coastal cyclogenesis may
occur along the existing warm frontal boundary during Type G events, and in fact,
several Type G events in our climatology featured secondary cyclogenesis southeast of
the parent cyclone. Meanwhile, the thickness ridge migrates northeastward and thickness
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values exceeding 540 dam overspread upstate New York and New England. At 1200
UTC 3 Jan 1999, the moisture plume now extends northward from the western Atlantic,
with PW values approaching 40 mm and θe values exceeding 320 K near the Delmarva
Peninsula (Fig. 5.7). Strong low- to midlevel southerly flow, coupled with large
meridional thickness and PW gradients, results in pronounced warm advection and
moisture transport, and helps establish a favorable thermodynamic profile for freezing
rain across the northeastern U.S. One exception is east of the northern Appalachians
(New Hampshire and Maine), where relatively low PW and θe values indicate that the cP
air mass has not yet been modified. This region is characterized by a weak thickness
trough and pressure ridge, which suggests that cold-air damming may also play a role
during ice storms in eastern New England.
Eventually, the secondary cyclone becomes separated from the primary cyclone
and intensifies as it moves up the New England coast (Fig. 5.8). At 0000 UTC 4 Jan
1999, a synoptic-scale thickness trough approaches the East Coast and low-level cold
advection dominates west of the secondary cyclone. Meanwhile, low-level warm
advection via southeasterly geostrophic flow is implied across Maine (thickness values
now exceed 540 dam throughout the entire state). As the secondary cyclone tracks
northeastward, the moisture plume begins to exit New England (note the sharp zonal
gradient in the PW field). However, the 850–700-hPa winds remain southerly across
Maine, allowing PW and θe values to approach 24 mm and 310 K, respectively, along the
coast (Fig. 5.9).
5.1.3 Mesoscale Dynamics
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Besides the synoptic-scale circulation, thermal boundaries, and moisture
transport, it is important to consider the role of mesoscale processes, such as terrain–flow
interactions and frontogenesis, in establishing favorable meteorological conditions for
prolonged freezing rain. Figures 5.10–5.13 illustrate the evolution of the SLP, 10-m
wind, and surface temperature fields between 0600 UTC 3 Jan 1999 and 1800 UTC 3 Jan
1999. At 0600 UTC 3 Jan 1999, the juxtaposition of the Great Lakes cyclone and
associated frontal trough with the anticyclone over northern Maine creates a strong
pressure gradient over the northeastern U.S. (Fig. 5.10). Due to land–ocean frictional
differences and flow blocking by the southwest-to-northeast oriented Appalachians, the
surface winds become cross-isobaric east of the Appalachians and turn east-northeasterly
toward lower pressure. Thus, although the synoptic-scale pressure pattern is not
especially favorable for cold-air damming, cold surface air continues to be reinforced
equatorward along the eastern slopes of the Appalachians, and the pressure ridge persists
between northern Virginia and eastern Massachusetts. Figure 5.11 highlights the strong
baroclinicity often observed in the vicinity of the Appalachians during Type G events. By
0600 UTC 3 Jan 1999, above-freezing surface air has spread northward across Ohio
within the warm sector of the Great Lakes cyclone. Relatively mild air also builds near
the Delmarva Peninsula as a result of low-level warm advection east of the cold-air
damming region. Meanwhile, the cold-air damming region is characterized by a narrow
thermal trough (T < 0°C) extending from central Pennsylvania to North Carolina. The
juxtaposition of very cold air across eastern New York and New England with relatively
mild air over the western Atlantic also creates a strong meridional temperature gradient
along the southern New England coast.
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During the following 12 h, the secondary cyclone separates from the parent
cyclone and intensifies along the mid-Atlantic coast (Fig. 5.12). As the low center tracks
northeastward and the zonal pressure gradient amplifies, the surface winds become more
southerly, thereby suggesting an increase in low-level warm advection across New
England. By 1800 UTC 3 Jan 1999, both the pressure ridge and the thermal trough along
eastern slopes of the central Appalachians have completely dissipated (Figs. 5.12 and
5.13). Subfreezing air still remains across northeastern New York and northern New
England, but in most locations warm advection via east-southeasterly onshore flow has
resulted in significant warming. The only exception is over interior Maine, where easterly
winds reinforce cold air west of the eastward retreating anticyclone and maintain a slight
pressure ridge along the eastern slopes of the northern Appalachians. Relative to Fig.
5.11, the baroclinicity has weakened along the mid-Atlantic and southern New England
coasts, but the surface temperature gradient has actually strengthened in southern Maine.
This increased baroclinicity likely results from differences in thermal advection implied
by southerly flow near the coast (warm advection) and easterly flow across the interior
(neutral temperature advection). The fact that the strongest baroclinicity is located inland
(as opposed to immediately at the coast) highlights the importance of surface topography
in trapping cold air along the eastern slopes of the northern Appalachians.
5.1.4 Local Thermodynamic Environment
Figures 5.14 and 5.15 illustrate the vertical distribution of temperature, dewpoint,
and wind over Pittsburgh, PA (0000 UTC 3 Jan 1999), and Albany, NY (1200 UTC 3 Jan
1999), respectively. As Fig. 5.1 demonstrates, the spatial distribution of freezing rain
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during the 3–4 Jan 1999 ice storm was unique in that significant ice accumulations were
reported both east and west of the highest elevations. The Pittsburgh, PA, sounding
reveals a well-defined surface-based subfreezing layer beneath a pronounced melting
layer. A saturated inversion layer lies between 925 hPa and 850 hPa, with temperatures
increasing from –4.5°C to 7.8°C. Although the Pittsburgh, PA, profile is notably drier
and more conducive to evaporational cooling above 850 hPa, the magnitude and depth of
the warm layer likely ensures that any hydrometeors reaching the inversion layer will
melt completely. Also note that low-level wind shear (10-kt southeasterly winds at 970
hPa; 65-kt south-southwesterly winds at 850 hPa) is indicative of warm advection and a
low-level jet (i.e., warm conveyor belt). Despite the presence of a surface-based
subfreezing layer, a saturated inversion, and a warm nose aloft, the vertical distribution of
temperature, dewpoint, and wind is somewhat different in the Albany, NY, sounding.
Compared to the Pittsburgh, PA, sounding, the area between the 0°C isotherm and
temperature profile is larger for the subfreezing layer, but considerably smaller for the
elevated melting layer. However, nearly saturated conditions below 700 hPa likely inhibit
evaporational cooling and thus favor the complete melting of hydrometeors. Regarding
vertical wind shear, the directional shear between 1000 hPa (southeasterly winds) and
850 hPa (southerly winds) is insignificant, but the strong speed shear (roughly 70 kt) is
indicative of a low-level jet and warm conveyor belt.
5.1.5 Parcel Trajectories
Figures 5.16 and 5.17 illustrate the source regions and trajectories for air parcels
500 m, 1500 m, and 3000 m above ground level (AGL) at the two sounding locations. At
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Pittsburgh, PA (0600 UTC 3 Jan 1999), air parcels arriving within the subfreezing layer
(500 m AGL) and near the top of the inversion (1500 m AGL) originated in the boundary
layer over south-central Canada (Fig. 5.16). These parcels slowly tracked southeastward
following the cold-air outbreak from Canada, then turned anticyclonically along the midAtlantic coast and ascended slightly as they approached and crossed the Appalachians.
The air parcel arriving above the inversion (3000 m AGL) descended from roughly 2000
m AGL over the Midwest to less than 500 m AGL over the western Atlantic. While
completing an anticyclonic loop off the Southeast coast, this parcel likely underwent
significant diabatic heating in the boundary layer (potential temperature [θ] increased
from ~280 K to ~293 K along the trajectory). During the last 12 h, this parcel rapidly
ascended within the warm sector of the primary cyclone. At Albany, NY (1800 UTC 3
Jan 1999), air parcels arriving within the subfreezing layer and near the top of the
inversion originated over northeastern Canada, at roughly 1000 m AGL and 2500 m
AGL, respectively (Fig. 5.17). These parcels turned cyclonically and gradually ascended
while passing over Hudson Bay during the first 48 h, then descended into the boundary
layer as they tracked across the Great Lakes and Northeast. Upon reaching the
northwestern Atlantic, these parcels looped back toward the East Coast and rose slightly
as they entered the ice storm region. The air parcel arriving above the inversion
descended rapidly from about 6000 m AGL over southwestern Canada to roughly 1000 m
AGL across the western Atlantic, where it remained quasi-stationary for 48–60 h. During
the last 24 h, this parcel likely experienced diabatic heating (θ increased from ~292 K to
~303 K along the trajectory) as it turned northward and ascended near the East Coast.
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5.2 11–12 Dec 2008 Ice Storm
5.2.1 Overview
The 11–12 Dec 2008 ice storm affected eight CWAs and 74 counties in a westsouthwest-to-east-northeast band from central Pennsylvania to eastern Maine (Fig. 5.18).
This event was the most destructive ice storm to impact New York and New England
since the historic 5–9 Jan 1998 ice storm. As Fig. 5.19 illustrates, the 11–12 Dec 2008 ice
storm was associated with copious amounts of precipitation. During the 24-h period
ending at 1200 UTC 12 Dec, quantitative precipitation estimates totaled 1–3 in (25.4–
76.2 mm) liquid equivalent throughout the ice storm area, with locally higher amounts
over elevated terrain. Storm Data also revealed widespread reports of ice accretion
between 0.50 in (1.27 cm) and 1.00 in (2.54 cm). Severe ice accumulations caused
considerable tree damage and downed numerous power lines and utility poles,
particularly across eastern New York, central New England, and southern Maine.
Combined power outages in the Albany, NY (ALY), Boston, MA (BOX), and Gray, ME
(GYX), CWAs exceeded 1.1 million, and thousands of additional power outages occurred
in the other five CWAs (NCDC 2008). The orientation of the ice storm region suggests
that the heaviest freezing rain fell on the poleward side of a zonally elongated frontal
boundary (i.e., warm front).
5.2.2 Synoptic Evolution
Figures 5.20–5.27 illustrate the evolution of the synoptic environment between
0000 UTC 11 Dec 2008 and 1200 UTC 12 Dec 2008. Initially, we see a low-level
baroclinic zone and an upper-level jet maximum extending from the Ohio Valley to the
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Canadian Maritimes (Fig. 5.20). The strong meridional thickness gradient is reinforced
by a thickness trough and an associated surface anticyclone across southeastern Canada.
1000–500-hPa thickness values exceed 540 dam throughout much of the northeastern
U.S., indicating sufficient warm air aloft for mixed precipitation. Meanwhile, a weak area
of low pressure is located near the Gulf Coast, south of the equatorward jet entrance
region. Dominant high pressure over the western Atlantic implies southerly return flow
along the Southeast coast. At 0000 UTC 11 Dec 2008, a broad swath of warm, moist air
(PW > 32 mm and θe > 320 K) extends from the western Atlantic basin northeastward to
Nova Scotia (Fig. 5.21). This large region of moisture-rich air is characterized by two
distinct ribbons of high PW: one originating over the Gulf of Mexico, and the other near
the western periphery of a synoptic-scale anticyclonic circulation over the western
Atlantic. Predominantly southwesterly flow within the 850–700-hPa layer allows for
effective moisture transport along and east of the Appalachians.
As time progress, the surface cyclone gradually deepens and moves across the
southeastern U.S. At 1200 UTC 11 Dec 2008, the low center is not only positioned south
of the equatorward entrance region of the primary jet, but is also located near the left exit
region of a second jet maximum over the Gulf of Mexico (Fig. 5.22). This coupled jet
configuration provides a favorable environment for surface cyclone intensification. The
evolution of the surface cyclone and associated regions of baroclinicity closely resembles
the Shapiro–Keyser model of cyclogenesis discussed by Schultz et al. (1998); note the Tbone structure of the thickness fields and zonally elongated warm frontal boundary.
Further north, the juxtaposition of the warm front along the mid-Atlantic coast with the
thickness trough over eastern Quebec results in the intensification of the meridional
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thickness gradient and corresponding northern jet maximum. As Fig. 5.23 reveals, the
typical Shaprio–Keyser T-bone structure is also evident in the PW field. Maximum PW
(> 48 mm) and θe (> 330 K) values occur within the south-southwesterly 850–700-hPa
flow ahead of the cold front (i.e., warm conveyor belt), while relatively high PW values
(> 36 mm) also extend northeastward along the warm front. The formation of a separate
θe maximum within the anticyclonic circulation off the Southeast coast suggests that the
Atlantic Ocean will serve as an important moisture source for this event. Large PW and
θe gradients across the Appalachians signify the juxtaposition of a cP air mass to the
north and west with a mT air mass to the south and east.
During the next 12 h, the surface cyclone continues to deepen and remain south of
the equatorward jet entrance region as it tracks east of the southern Appalachians (Fig.
5.24). The associated surface warm front becomes better-defined and gradually migrates
toward the southern New England coast. Although the surface anticyclone over
southeastern Canada begins to retreat eastward, two distinct pressure ridges develop
across the St. Lawrence Valley and eastern New England. As in Figs. 5.20 (24 h prior)
and 5.22 (12 h prior), the ice storm region is characterized by abundant warm air aloft
(thickness values between 540 and 552 dam). Given the proximity of the ice storm region
to the surface warm front, the interior Northeast lies in a favorable location for
precipitation associated with frontal overrunning. By 0000 UTC 12 Dec 2008, the
moisture plume has become much narrower and more organized along the East Coast
(Fig. 5.25). Strong low- to midlevel south-southwesterly flow aids in the poleward surge
of very warm, moist air (PW > 48 mm and θe > 330 K) near the Southeast coast. Southsouthwesterly moisture transport also intensifies over the mid-Atlantic and southern New
83
England states, allowing the mT air mass to overspread the region above the shallow cold
surface layer.
At 1200 UTC 12 Dec 2008, the surface cyclone is crossing Long Island and
continues to intensify as it rapidly moves northeastward (Fig. 5.26). The location of the
upper-level jet entrance region not only places the surface cyclone in a favorable position
for deepening, but also facilitates QG forcing for ascent throughout New England.
Meanwhile, the moisture plume continues to migrate northeastward, with PW values
approaching 40 mm and θe values exceeding 320 K over the Gulf of Maine (Fig. 5.27).
Deep south-southwesterly flow implies vigorous low- to midlevel moisture transport and
warm advection across eastern New England.
5.2.3 Mesoscale Dynamics
Figures 5.28–5.31 illustrate the evolution of the SLP, 10-m wind, and surface
temperature fields between 1800 UTC 11 Dec 2008 and 0600 UTC 12 Dec 2008. At 1800
UTC 11 Dec 2008, the southwest-to-northeast pressure gradient induces pressure-driven
channeling (e.g., Whiteman and Doran 1993) down the southwest-to-northeast oriented
St. Lawrence Valley and north–south oriented Champlain and Hudson Valleys (Fig.
5.28). Given the predominantly meridional background temperature gradient (Fig. 5.29),
enhanced northerly or northeasterly surface winds imply cold advection and help
reinforce shallow cold air at the surface. In response to the orographically forced
northeasterly flow, a mesoscale pressure ridge and thermal trough develop throughout the
St. Lawrence Valley. Slight pressure riding also occurs within the region of northeasterly
flow across eastern New England. Further south, we see the surface warm front denoted
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by a pressure trough and associated baroclinic zone extending northeastward from the
Delmarva Peninsula. This frontal boundary is characterized by strong confluence in the
10-m wind field, with predominantly northeasterly flow on the poleward side and
southerly flow on the equatorward side.
Between 1800 UTC 11 Dec 2008 and 0600 UTC 12 Dec 2008, the surface low
deepens and tracks northeastward, causing the pressure gradient to increase across New
York and New England (Fig. 5.30). Pressure-driven channeling continues throughout the
St. Lawrence, Champlain, and Hudson Valleys, and northeasterly winds persist across
eastern New England. Meanwhile, the surface warm front gradually approaches Long
Island and southern New England. An increase in the pressure gradient and cyclonic
curvature of the isobars along the frontal boundary reinforces the confluence in the 10-m
wind field. As a result of this persistent confluence, the surface temperature gradient has
strengthened considerably along the mid-Atlantic and southern New England coast by
0600 UTC Dec 2008 (Fig. 5.31). Moreover, given the proximity of the frontal boundary
to the coast, land–ocean contrasts in both temperature and friction likely accelerate the
surface frontogenesis (e.g., Szeto et al. 1999). Also note that the 0°C isotherm, which
approximately denotes the southern extent of the ice storm region, has remained nearly
stationary between northern Pennsylvania and central New England since 1800 UTC 11
Dec 2008.
5.2.4 Local Thermodynamic Environment
Figures 5.32 and 5.33 illustrate the vertical distribution of temperature, dewpoint,
and wind over Albany, NY (0000 UTC 12 Dec 2008), and Gray, ME (0600 UTC 12 Dec
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2008), respectively. Both profiles are representative of the classical freezing rain
sounding, with a surface-based subfreezing layer, a low-level inversion, and an elevated
melting layer. Furthermore, both profiles are saturated with respect to water below 700
hPa and saturated with respect to ice above 700 hPa. Significant directional wind shear
within the 1000–850-hPa layer also implies low-level warm advection. Despite these
fundamental similarities, the Gray, ME, sounding exhibits a shallower subfreezing layer,
a deeper and more well-defined warm layer, and a greater low- to midlevel mixing ratio
than the Albany, NY, sounding. Additionally, the low-level wind shear and implied warm
advection are stronger in the Gray, ME, sounding (northerly winds near the surface
veering to southerly at 850 hPa) than in the Albany, NY, sounding (weak northeasterly
winds near the surface veering to weak southerlies at 850 hPa).
5.2.5 Parcel Trajectories
Figures 5.34 and 5.35 illustrate the source regions for air parcels arriving 500 m,
1500 m, and 3000 m AGL near the two sounding locations. At Albany, NY (0000 UTC
12 Dec 2008), the air parcel within the surface-based subfreezing layer (500 m AGL)
descended from roughly 3000 m AGL over interior Canada (cP air mass) to the boundary
layer across New England, where it turned anticyclonically around the surface high
pressure center (Fig. 5.34). The air parcel above the inversion (3000 m AGL) had
extensive contact with the western Atlantic (mT air mass) before undergoing rapid ascent
as it approached the ice storm region. Throughout the 120-h period, this parcel likely
experienced substantial diabatic heating (θ increased from ~280 K to ~302 K along the
trajectory). At Gray, ME (0600 UTC 12 Dec 2008), the air parcel arriving within the
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subfreezing layer descended from 1500 m AGL near Hudson Bay before circulating
around the anticyclone north of New England (Fig. 5.35). The evolution of the air parcel
arriving above the inversion layer at Gray, ME, was rather complex. During the first 24 h,
this parcel descended from roughly 7000 m AGL over interior Canada to the boundary
layer over the western Atlantic. This parcel subsequently entered the synoptic-scale
anticyclonic circulation associated with the subtropical high, likely experienced
considerable diabatic heating (θ increased from ~285 K to ~304 K along the trajectory),
and remained in the boundary layer until rapidly ascending to 3000 m AGL during the
last 6 h. Backward trajectories from Albany, NY, and Gray, ME, clearly illustrate that the
11–12 Dec 2008 ice storm was characterized by the juxtaposition of two air masses with
very distinct thermodynamic properties (a cP air mass below a mT air mass). Moreover,
parcels arriving above the low-level inversion at both locations were likely tied to the
low- to midlevel moisture transport observed throughout the ice storm region.
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Figure 5.1: Map highlighting the 149 counties impacted by the 3–4 Jan 1999 ice storm.
Figure 5.2: 300-hPa wind speed (shaded, every 5 m s−1), 1000–500-hPa thickness
(dashed, every 6 dam), and mean sea level pressure (black, every 4 hPa) at 1200 UTC 2
Jan 1999.
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Figure 5.3: Precipitable water (shaded, every 4 mm), 850–700-hPa layer-averaged
equivalent potential temperature (white, every 10 K), and 850–700-hPa layer-averaged
moisture transport (arrows, expressed as the product of wind velocity and mixing ratio,
multiplied by 100) at 1200 UTC 2 Jan 1999.
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Figure 5.4: As in Fig. 5.2, except at 0000 UTC 3 Jan 1999.
Figure 5.5: As in Fig. 5.3, except at 0000 UTC 3 Jan 1999.
90
Figure 5.6: As in Fig. 5.2, except at 1200 UTC 3 Jan 1999.
Figure 5.7: As in Fig. 5.3, except at 1200 UTC 3 Jan 1999.
91
Figure 5.8: As in Fig. 5.2, except at 0000 UTC 4 Jan 1999.
Figure 5.9: As in Fig. 5.3, except at 0000 UTC 4 Jan 1999.
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Figure 5.10: Mean sea level pressure (black, every 2 hPa) and 10-m streamlines (blue) at
0600 UTC 3 Jan 1999. Surface elevation is shaded every 200 m.
93
Figure 5.11: 2-m temperature (red, every 2°C) at 0600 UTC 3 Jan 1999. Surface
elevation is shaded every 200 m.
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Figure 5.12: As in Fig. 5.10, except at 1800 UTC 3 Jan 1999.
Figure 5.13: As in Fig. 5.11, except at 1800 UTC 3 Jan 1999.
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090103/0000 72520 PIT (Pittsburgh, PA)
Figure 5.14: Skew-T with temperature (red, °C), dewpoint (blue, °C), and wind (barbs,
kt). Solid black, orange, and green lines denote constant values of temperature, potential
temperature, and equivalent potential temperature, respectively. Dashed green lines
denote constant values of mixing ratio. The 0°C isotherm is highlighted in purple. Based
on rawinsonde data from the 0000 UTC 3 Jan 1999 Pittsburgh, PA (PIT), sounding.
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090103/1200 72518 ALB (Albany, NY)
Figure 5.15: As in Fig. 5.14, except based on rawinsonde data from the 1200 UTC 3 Jan
1999 Albany, NY (ALB), sounding.
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Figure 5.16: 120-h backward trajectories ending at 0600 UTC 3 Jan 1999 over
Pittsburgh, PA (PIT), which is denoted by the black star. Red, blue, and green lines
represent air parcels arriving 500 m, 1500 m, and 3000 m above ground level (AGL),
respectively. The top time series indicates the altitude of each parcel (m AGL), while the
bottom time series indicates the ambient potential temperature (K), at 6-h intervals.
Calculated from NCEP/NCAR reanalysis data.
98
Figure 5.17: As in Fig. 5.16, except for backward trajectories ending at 1800 UTC 3 Jan
1999 over Albany, NY (ALB).
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Figure 5.18: Map highlighting the 74 counties impacted by the 11–12 Dec 2008 ice
storm.
Figure 5.19: Quantitative precipitation estimates valid for the 24-h period ending at 1200
UTC 12 Dec 2008. From the National Mosaic & Multi-sensor QPE.
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Figure 5.20: As in Fig. 5.2, except at 0000 UTC 11 Dec 2008.
Figure 5.21: As in Fig. 5.3, except at 0000 UTC 11 Dec 2008.
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Figure 5.22: As in Fig. 5.2, except at 1200 UTC 11 Dec 2008.
Figure 5.23: As in Fig. 5.3, except at 1200 UTC 11 Dec 2008.
102
Figure 5.24: As in Fig. 5.2, except at 0000 UTC 12 Dec 2008.
Figure 5.25: As in Fig. 5.3, except at 0000 UTC 12 Dec 2008.
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Figure 5.26: As in Fig. 5.2, except at 1200 UTC 12 Dec 2008.
Figure 5.27: As in Fig. 5.3, except at 1200 UTC 12 Dec 2008.
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Figure 5.28: As in Fig. 5.10, except at 1800 UTC 11 Dec 2008.
Figure 5.29: As in Fig. 5.11, except at 1800 UTC 11 Dec 2008.
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Figure 5.30: As in Fig. 5.10, except at 0600 UTC 12 Dec 2008.
Figure 5.31: As in Fig. 5.11, except at 0600 UTC 12 Dec 2008.
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081212/0000 72518 ALB (Albany, NY)
Figure 5.32: As in Fig. 5.14, except based on rawinsonde data from the 0000 UTC 12
Dec 2008 Albany, NY (ALB), sounding.
107
081212/0600 74389 GYX (Gray, ME)
Figure 5.33: As in Fig. 5.14, except based on rawinsonde data from the 0600 UTC 12
Dec 2008 Gray, ME (GYX), sounding.
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Figure 5.34: As in Fig. 5.16, except for backward trajectories ending at 0000 UTC 12
Dec 2008 over Albany, NY (ALB), and calculated from the NCEP Global Data
Assimilation System (GDAS) model.
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Figure 5.35: As in Fig. 5.16, except for backward trajectories ending at 0600 UTC 12
Dec 2008 over Gray, ME (GYX), and calculated from the NCEP Global Data
Assimilation System (GDAS) model.
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