4334 MONTHLY WEATHER REVIEW VOLUME 136

The Effects of Complex Terrain on Severe Landfalling Tropical Cyclone Larry (2006) over Northeast

HAMISH A. RAMSAY School of Meteorology, and Cooperative Institute for Mesoscale Meteorological Studies, University of Oklahoma, Norman, Oklahoma

LANCE M. LESLIE School of Meteorology, University of Oklahoma, Norman, Oklahoma

(Manuscript received 19 October 2007, in final form 1 April 2008)

ABSTRACT

The interaction between complex terrain and a landfalling tropical cyclone (TC) over northeastern Australia is investigated using the fifth-generation Pennsylvania State University–National Center for At- mospheric Research (PSU–NCAR) Mesoscale Model (MM5). Severe TC Larry (in March 2006) made landfall over an area of steep coastal orography and caused extensive damage. The damage pattern sug- gested that the mountainous terrain had a large influence on the TC wind field, with highly variable damage across relatively small distances. The major aims in this study were to reproduce the observed features of TC Larry, including track, intensity, speed of movement, size, decay rate, and the three-dimensional wind field using realistic high-resolution terrain data and a nested grid with a horizontal spacing of 1 km for the finest domain (referred to as CTRL), and to assess how the above parameters change when the terrain height is set to zero (NOTOPOG). The TC track for CTRL, including the timing and location of landfall, was in close agreement with observation, with the model eye overlapping the location of the observed eye at landfall. Setting the terrain height to zero resulted in a more southerly track and a more intense storm at landfall. The orography in CTRL had a large impact on the TC’s 3D wind field, particularly in the boundary layer where locally very high wind speeds, up to 68 m sϪ1, coincided with topographic slopes and ridges. The orography also affected precipitation, with localized maxima in elevated regions matching observed rainfall rates. In contrast, the precipitation pattern for the NOTOPOG TC was more symmetric and rainfall totals decreased rapidly with distance from the storm’s center. Parameterized maximum surface wind gusts were located beneath strong boundary layer jets. Finally, small-scale banding features were evident in the surface wind field over land for the NOTOPOG TC, owing to the interaction between the TC boundary layer flow and land surface characteristics.

1. Introduction 2004). A “severe” TC is one that produces sustained winds of at least 118 km hϪ1 (33 m sϪ1), with gusts of Tropical cyclones are an annual threat to Australia 170–225 km hϪ1 (47–63 m sϪ1). An average of 12.5 TCs each year from November to April, often with devas- occur in the Australian region each year, with about 5 tating consequences. A tropical cyclone (TC) in the making landfall. Australian region (90°–160°E in the Southern Hemi- The far northeastern region of Australia (see loca- sphere) is defined by the Australian Bureau of Meteo- tion map Fig. 1a) is significant because the complex rology as “a nonfrontal, synoptic-scale cyclone that has terrain in close proximity to the coastline acts to en- developed over tropical waters, with a 10-min average hance the destructive potential of landfalling TCs. Be- V Ն Ϫ1 Ϫ1 wind speed, 63 km h [17 m s ] near the center tween one and two TCs make landfall on the northeast of the organized wind circulation” (Dare and Davidson Australian coast each year (Dare and Davidson 2004). While less prone to TCs than more sparsely populated northwestern Australia, TCs that impact the far north- Corresponding author address: Hamish A. Ramsay, CIMMS, University of Oklahoma, National Weather Center, Suite 2100, east coast pose an extreme threat to large cities such as 120 David L. Boren Blvd., Norman, OK 73072. and Townsville, Australia, smaller coastal com- E-mail: [email protected] munities, and the agricultural industry of the region.

DOI: 10.1175/2008MWR2429.1

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FIG. 1. (a) Location map showing the far northeastern region of Australia, delineated by the black rectangle. (b) Map showing the region of complex terrain affected by TC Larry, including the town of Innisfail where the eye crossed during landfall. Terrain contours are every 100 m. The tallest peaks of and Bellenden Ker are marked with the initials BF and BK, respectively.

Severe TC Larry crossed the northeastern Australian caused by TC Agnes at Cairns was comparable to the coast near the town of Innisfail, Australia (Fig. 1b) on damage near where the eye crossed the coast at Towns- the morning of 20 March 2006, causing significant dam- ville. In 1986 TC Winifred, though smaller than Agnes age to infrastructure and crops, with an estimated cost in size, crossed the coast just south of Innisfail and of about $500 million Australian dollars. Analyses of caused scattered westerly wind damage to the Cairns damage to the region suggested that the local topogra- region. phy played a significant role in modifying Larry’s wind Previous studies of the influence of orography on TC field, as evidenced by varying degrees of damage across structure and intensity have focused mainly on ideal- relatively small distances. A preliminary damage inves- ized numerical simulations (e.g., Chang 1982; Bender et tigation by engineers at James Cook University, al. 1985), or on TCs that interact with the Central Townsville, Australia, found that wind speed up over Mountain Range (CMR) of Taiwan (e.g., Bender et al. topographic ridges resulted in a significant increase in 1987; Chang et al. 1993; Yeh and Elsberry 1993a,b; Lin damage to buildings, whereas those structures sheltered et al. 1999, 2002, 2005, 2006; Wu and Kuo 1999; Wu by topography suffered much less damage. The town of 2001; Wu et al. 2002). The Sierra Madre Mountains of , for example, to the northwest of Innisfail (Fig. Mexico have also been shown to influence TCs that 1b), sustained disproportionate damage relative to its approach from the Gulf of Mexico (e.g., Zehnder 1993; distance from Larry’s eye by the funneling of westerly Farfan and Zehnder 2001). Other mountainous regions downslope winds. Westerly downslope winds also affected by TCs include Luzon in the northern Philip- caused moderate damage in and around the Cairns re- pines (Brand and Blelloch 1973) and the Caribbean gion (Fig. 1b). Islands of Cuba, Hispaniola, and Puerto Rico (Bender Cairns is often affected by these damaging westerly et al. 1987). winds when TCs make landfall south of the region. Cal- The above studies have provided considerable under- laghan (2003) noted at least three historical TC events standing of how orography affects the track and struc- that resulted in substantial westerly wind damage in ture of TCs. While the importance of orographic influ- Cairns, the most extreme of which was TC Agnes in ences on landfalling TCs over northeastern 1956. TC Agnes crossed the coast near Townsville, has long been recognized (e.g., Whittingham 1964; about 280 km south of Cairns, with a central pressure of Harper 1999; Callaghan 2003), to our knowledge there 960 hPa. Despite its modest intensity at landfall, have been no comprehensive high-resolution numerical Agnes’s very large circulation resulted in extensive simulations that have explored the influence of orogra- westerly wind damage in the Cairns region, with wind phy on Australian TCs. speeds varying from nearly calm to 43 m sϪ1 over very Observations of TC Larry (more information avail- short time periods (Whittingham 1964). The damage able online at http://www.bom.gov.au/weather/qld/

Unauthenticated | Downloaded 10/05/21 10:04 PM UTC 4336 MONTHLY WEATHER REVIEW VOLUME 136 cyclone/tc_larry; Henderson et al. 2006) motivated this high-resolution numerical simulation. Our main objec- tive is to understand how the complex terrain of the northeastern Australian region affects TC track, winds, and precipitation, and to assess how different the im- pact of TC Larry would have been if the region was flat. Section 2 is a summary of the major observational aspects of TC Larry, including its track and its charac- teristics before, during, and after landfall. Section 3 de- scribes the fifth-generation Pennsylvania State Univer- sity–National Center for Atmospheric Research (PSU– NCAR) Mesoscale Model (MM5) model configuration and the proposed numerical experiments. The results are given in section 4, and section 5 summarizes the findings and suggests possible future work.

2. Observations of TC Larry FIG. 2. The quadruply nested grid configuration used for the TC Larry developed from a tropical low in the Coral simulations. The grid spacing for each domain is indicated next to Sea (see location in Fig. 1a), initially located about 1500 the domain number. km east of Cairns at 0600 UTC 17 March 2006. The low was upgraded to an Australian category 1 TC 12 h later duced extensive flooding in coastal rivers. As TC Larry (1800 UTC 17 March 2006) and continued to steadily moved inland toward the low-lying area south of the intensify as it moved westward, reaching the Australian Gulf of Carpentaria very heavy continued, with the severe TC classification during the morning of 18 highest recorded total of 436 mm at Gereta Station March 2006. Ten hours prior to landfall (1110 UTC 19 Ϫ (about 700 km inland) in the 24 h preceding 2300 UTC March 2006) a wind gust of 59 m s 1 was recorded at 21 March 2006. Flinders Reef, located about 260 km east of Innisfail. Larry was a small TC with destructive wind gusts of An analysis of high-resolution microwave satellite im- greater than 47 m sϪ1 extending no more than 50 km agery suggests that Larry may have briefly reached from the center. It had an unusually fast forward speed Australian category 5 intensity (i.e., gusts of greater Ϫ1 Ϫ of about 8 m s as it tracked to the west-northwest than 78 m s 1) before making landfall as a category 4 Ϫ after landfall. Tropical Cyclone Larry decayed fairly storm (with gusts 63–78 m s 1) near the town of Innis- rapidly as it moved inland, though it maintained cat- fail (Fig. 1b). Radar scans from Willis Island showed a egory 1 strength or higher for several hundred kilome- persistent convective asymmetry in TC Larry well be- ters. It was downgraded to a tropical low at 1500 UTC fore landfall, with the deepest convection located on 20 March 2006, 18 h after landfall and about 500 km the northern side the eyewall (P. Otto, Bureau of Me- inland from the coast. teorology 2008, personal communication). The highest official observed wind gust near the time of landfall (2020–2120 UTC 19 March 2006) was 50 3. Model configuration and experimental design msϪ1 recorded by the South Johnstone Automatic Weather Station (AWS), located about 11 km south- Version 3.7 of MM5 was used for all simulations. A west of Innisfail (Fig. 1b) and 14 km inland. However, detailed description of the properties of MM5 can be an analysis of damage to the regions most affected sug- found in Grell et al. (1994). gested that peak wind gusts were considerably higher, The MM5 simulations were computed on a quadru- being estimated at 66 m sϪ1, which is consistent with an ply nested, two-way interactive grid, as shown in Fig. 2. Australian category 4 TC. The highest unofficial mea- The four domains have dimensions and grid spacings of sured wind gust was 82 m sϪ1 at Bellenden Ker Tower 93 ϫ 100, 27 km (D1); 220 ϫ 210, 9 km (D2); 445 ϫ 286, near the peak of Mt. Bellenden Ker (see Fig. 1b). The 3 km (D3); and 385 ϫ 268, 1 km (D4). In total, 46 tower elevation is approximately 1450 m and is 30 km vertical half-sigma levels were used with higher resolu- northwest of Innisfail. Cairns Airport reported a maxi- tion in the boundary layer and upper troposphere to mum wind gust of 30 m sϪ1 at 2201 UTC. more accurately resolve the inflow and outflow layers Heavy rainfall, with 3-h totals up to 139 mm, pro- of the TC. The model top is set at 50 hPa. High-resolu-

Unauthenticated | Downloaded 10/05/21 10:04 PM UTC NOVEMBER 2008 RAMSAY AND LESLIE 4337 tion terrain data with horizontal resolution of 900 m 18 March to 1200 UTC 20 March 2006 when Larry was was used in D4 to resolve the interaction of the TC officially downgraded, 12 h after landfall. Varying the wind field with topography. The complex nature of this size and strength of the initial vortex by specifying dif- topography in the region where Larry made landfall is ferent values of the radius of maximum wind (RMW) shown in Fig. 1b. The steepest elevated regions of and the maximum tangential wind speed led to rela- Mount Bartle Frere and Mount Bellenden Ker, at 1622 tively large changes in simulated intensity, though only and 1593 m, respectively, were reproduced well in D4 minor variations in track. Stronger initial bogus vortices with the highest-resolution model terrain rising to 1600 generally resulted in more southerly storm tracks rela- and 1484 m. tive to Larry’s best track. Of more importance for the All simulations were initialized at 1200 UTC 17 current study for the purpose of studying the effect of March and run for 72 h until the storm decayed roughly the local terrain on the TC’s wind field, are the time and to the Australian category 1 strength (i.e., gusts to 35 location of landfall. Both of these agree very well with msϪ1) after landfall. The initial and lateral boundary observation (Figs. 3 and 4). The simulated TC with to- conditions were obtained from the National Centers for pography crosses the coast about 2 h after the observed Environmental Prediction (NCEP) Final Analysis time of Larry’s landfall, with the southern half of the (FNL) dataset, with 1 ϫ 1 degree horizontal grid spac- eye passing directly over Innisfail. The TC with no to- ing and 26 mandatory pressure levels from 1000 to 10 pography shares the same track as the CTRL TC until hPa. In addition, a bogus vortex was inserted into the t ϭ 36 h, after which it develops a more southerly com- initial state, following the TC bogussing scheme of ponent of motion that results in a landfall location Low-Nam and Davis (2001). The model failed to gen- about 80 km to the south of the CTRL TC and the erate a TC when no bogus vortex was used, owing to observed landfall location of Larry (Fig. 3). Until about the relatively coarse resolution of the FNL dataset. t ϭ 30 h, the low-level and upper-level wind field PBL processes were parameterized using the Eta patterns for the CTRL TC and NOTOPOG TCs are Model Mellor–Yamada 1.5-order local closure scheme in close agreement, however by t ϭ 36 h some nota- (Janjic 1990, 1994) that includes a prognostic equation ble differences begin to emerge. Analyses of the for turbulence kinetic energy (TKE). Betts–Miller cu- NOTOPOG TC’s wind field indicate increased mid-to- mulus parameterization (Betts 1986; Betts and Miller upper-level flow in the rear-left quadrant relative to the 1986) was used for D1 and D2 only, while cloud- CTRL TC, resulting in greater advection of potential microphysical processes were explicitly resolved in all vorticity (toward the south) and a concomitant south- domains using the Reisner mixed-phase scheme (Reis- erly component of the TC’s motion. While difficult to ner et al. 1998). Both shortwave and longwave radiation establish a direct link between the removal of orogra- were computed using the rapid radiation transfer model phy and TC track differences prior to landfall, the re- (RRTM; Mlawer et al. 1997). lationship between orography and storm motion after To study the effect of orography on TC Larry, a sen- landfall is more obvious. As the CTRL TC passes over sitivity experiment (NOTOPOG) was carried out in the coastal range, the easterly flow on the left side of which the terrain elevation was set to zero. Special care the storm descends and warms while westerly flow on was taken to ensure that the sensitivity experiment had the right side of the storm rises and cools (not shown). no memory of the original terrain data by allowing the Therefore, storm-relative cyclonic vorticity is increased model to run for a sufficiently long time (ϳ60 h) prior (decreased) on the left (right) side of the storm by to TC landfall. By 48 h the surface and upper-level stretching (compression) of vortex tubes. Hence, the fields over land showed no sign of the orography that vortex moves toward the southwest like the observed was removed at initialization. storm (Fig. 3), while the NOTOPOG TC maintains its westerly motion. In the Northern Hemisphere this same combination of adiabatic warming and vortex 4. Results tube stretching in downslope flow results in TC move- ment toward the northwest (e.g., Lin et al. 2002, p. a. TC track and intensity 2934). The simulated TC track for the realistic terrain ex- The removal of orography results in a more intense periment (hereafter CTRL) is in very good agreement TC relative to the CTRL TC. Figure 4 shows the esti- with the observed track of TC Larry, as shown in Fig. 3. mated best-track intensity versus the simulated inten- After an initial 24-h spinup period, the simulated TC sities for the CTRL and NOTOPOG simulations, re- developed a more westerly track, following the esti- spectively. After the 24-h spinup period, intensification mated best-track positions very closely from 1200 UTC rates for the CTRL and NOTOPOG TCs follow TC

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FIG. 3. (a) Simulated storm tracks for the CTRL (white triangles) and NOTOPOG (white circles) experiments as well the best track of TC Larry (black squares). The white symbols indicate the position of central pressures every 3 h from 1712 UTC 17 Mar to 2012 UTC 20 Mar 2006.

Larry’s rate of intensification closely. The central pres- prior to landfall; 18 hPa lower than the minimum cen- sures in D2 are about 5–7 hPa higher than D3, however tral pressure (929 hPa) for the CTRL TC. Both TCs there is almost no difference between the pressures in exhibit weakening in the 4 h leading up to landfall. D3 and D4, indicating that the TC eye is well resolved The time at which the intensities of the CTRL and for grid spacing less than about 3 km. Rapid intensifi- NOTOPOG TCs begin to diverge (t ϭ 30 h) is the same cation from t ϭ 36htot ϭ 45 h is captured well by the as when the TC tracks separate, suggesting differences model. By t ϭ 51 h the best track and CTRL central in storm structure and/or the ambient environment. pressures converge and remain close until landfall. The sea surface temperature is the same for both simu- Note that a substantial data-void period exists in the lations so should not cause differences in intensity. estimated best-track intensity between 1800 UTC 19 Rather, the difference in intensity appears to be the March and 0000 UTC 20 March 2006 (t ϭ 54 and t ϭ 60 result of a distinct increase in the deep-layer (850–200 in Fig. 4) including the time at which Larry crossed the hPa) near-core vertical wind shear experienced by the coast near Innisfail around 2100 UTC 19 March 2006. CTRL TC from about t ϭ 27 h to t ϭ 36 h. This increase The lowest official central pressure near landfall was causes the CTRL vortex to tilt in the downshear direc- 957 hPa at South Johnstone automatic weather station, tion (i.e., toward the southeast) such that the upper- which was near the southern edge of the eyewall, so the level circulation is displaced about 30 km from the low- actual minimum pressure probably was lower. Starting level circulation. The NOTOPOG vortex shows much from model initialization at 1712 UTC 17 March 2006, less vertical tilt in accordance with the somewhat the intensification rate of the simulated TCs is the same weaker deep-layer vertical shear. These results agree until t ϭ 30 h, after which the NOTOPOG TC deepens with previous studies that have investigated the effects faster than the CTRL TC. The NOTOPOG TC attains of vertical wind shear on TC intensity (e.g., DeMaria a minimum central pressure in D3 of 911 hPa about 4 h 1996; Frank and Ritchie 2001; Braun et al. 2006).

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FIG. 4. Time series of best-track central pressure for TC Larry (black squares) and simulated central pressure for the CTRL (solid line with triangles) and NOTOPOG simulations (dashed line with circles) using data from D2. The vertical lines indicate the time of landfall for each simulation.

b. TC structure during landfall shape of the eye that is consistent with radar observa- tions during Larry’s landfall (P. Otto, Bureau of Me- In addition to the structure and intensity differences teorology 2008, personal communication). While not an prior to landfall discussed above, distinct differences integral part of this paper, we note also that the CTRL also exist between the CTRL and NOTOPOG TC’s simulation produces polygonal-shaped eyewalls in the finescale characteristics during landfall. The CTRL TC hours leading up to landfall that match very closely with makes landfall around 2300 UTC 19 March 2006 with a Larry’s polygonal eyewall structure observed by radar central pressure of 935 hPa (D4). The center of the eye (Fig. 7). In Fig. 7 both the radar and the model eye walls crosses the coast 3 km north of Innisfail (Fig. 5). The appear to be pentagonal. However, at other times the diameter of the eye, as shown by the 1.8-km simulated eyewall takes other polygonal shapes (not shown). Pre- radar reflectivity in Fig. 5c, is about 22 km; only slightly vious studies have shown that polygonal eyewalls are smaller than Larry’s actual eye, estimated to be about the result of a breakdown of strong vorticity within the 25 km in diameter during landfall (Henderson et al. eyewall into discrete pools and that these play an im- 2006). port role in the inner-core dynamics of TCs (e.g., Schu- In terms of finescale inner-core structure, we find bert et al. 1999; Kossin and Schubert 2001; Wang ␪ that the equivalent potential temperature ( e)at1km 2002a,b). within the eye is fragmented into three separate Throughout this study, the term “surface wind” de- maxima, with values approaching 375 K concentrated notes the 10-m wind, which is calculated by the model around sections the eyewall (Fig. 6). Maxima in cy- based on boundary layer and land surface characteris- clonic vertical vorticity of up to Ϫ1.8 ϫ 10Ϫ2 sϪ1 are tics. For the CTRL TC, a region of maximum surface ␪ found within the tight e gradient along the southern wind speed is evident in the rear-left quadrant of the and western sides of the eyewall (Fig. 6). This distribu- circulation just offshore over the water, with velocities Ϫ tion of cyclonic vorticity reveals the distinct elliptical ranging between 50 and 58 m s 1. A second area of very

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Ϫ FIG. 5. (a) Surface wind magnitude (gray shading with contour interval of 8 m s 1) and velocity vectors during landfall for the CTRL simulation. The terrain height is given by solid black contours with an interval of 200 m. (b) Corresponding east–west vertical cross section through the center of storm [marked by the black line in (a)] showing the ␷ component of the wind with contour interval of 5 msϪ1. Solid (dashed) white contours indicate southerly (northerly) winds. (c) Simulated radar reflectivity at 1.8 km with contour interval of 5 dBZ. (d) East–west vertical cross section of the radar reflectivity corresponding to the position of the solid black line in (c). The freezing line is plotted in white. strong surface winds, approaching 52 m sϪ1, is evident distinct maxima. The first is collocated with the eyewall in the core region to the north of the eye over the with wind speeds up to 34 m sϪ1. The second region ocean. The surface wind speed over land is significantly farther to the west is stronger, with speeds approaching reduced owing to the increased surface roughness, and 44 m sϪ1, despite a greater radial distance (ϳ25 km) shows much more spatial variability as a result of the from the eye. This second band of winds follows the steep and complex terrain (Fig. 5a). The strongest contours of elevated terrain to the west, suggesting the winds over land (ϳ50 m sϪ1) are restricted to a small TC’s tangential flow is accelerated by the orography, at coastal area south of the eye exposed to the onshore least locally. No such strong, secondary inland wind flow on the left side of the storm. This onshore flow maximum is evident in the NOTOPOG TC’s wind field diminishes quickly with distance from the coast, slow- during landfall (Fig. 8a). In addition, there are several ing from 50 to 34 m sϪ1 over a distance of 3 km. Analy- smaller, localized, pockets of strong winds located on ses of the tangential wind speed over land reveals two the upwind sides of Mt. Bartle Frere and Mt. Bellenden

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FIG. 6. Horizontal distribution of equivalent potential temperature (black contours with contour interval of 2 K) and cyclonic vertical vorticity starting from Ϫ1000 ϫ 10Ϫ5 sϪ1 [gray shading with contour interval of (a) 100 and (b) 150] at z ϭ 1 km for (a) CTRL simulation at landfall and (b) NOTOPOG simulation at landfall. Terrain elevation contours are designated by the black dashed lines.

Ker (Fig. 5a; see location map Fig. 1b). Surface winds of m above ground level (AGL) and 21 km to the west of up to 38 m sϪ1 are evident on their southern slopes the TC center (Fig. 5b). On the eastern side of the TC, where the tangential flow is partially blocked so it is a similar northerly wind maximum of 77 m sϪ1 is forced to go around rather than over the mountains. present about 500 m MSL and 20 km east of the eye. The surface winds on the sheltered lee sides of these Relative to the CTRL TC, the NOTOPOG TC is mountains are significantly lower (ϳ8msϪ1). more intense at landfall with a central pressure of 928 An east–west cross section of the meridional winds hPa (D4); however, it contains only slightly stronger through the center of the TC during landfall reveals surface winds with maximum values approaching 60 only a very slight asymmetry in the tangential winds, msϪ1 (Fig. 8a). The finescale inner-core structure is with a maximum southerly wind of 73 m sϪ1 about 600 characterized by a strong warm core with a maximum

FIG. 7. (a) Observed reflectivity of TC Larry at 1850 UTC 19 Mar 2006 showing a distinct pentagonal eyewall and (b) simulated reflectivity for the CTRL simulation (D4) at z ϭ 2 km showing pentagonal eyewall structure at 2109 UTC 19 Mar 2006.

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FIG. 8. As in Fig. 5, but for the NOTOPOG simulation. equivalent potential temperature of 380 K near the cen- The east–west cross section of the meridional wind field ter of the eye (5 K higher than the 375 K for the CTRL through the center of the TC reveals two distinct low- TC) decreasing sharply with radial distance out from level jets at a radius of about 10 km from the axis of the center (Fig. 6b). The stronger, more symmetric minimum winds in the eye. These jets occur at a height warm core of the NOTOPOG TC relative to the CTRL of about 250 m MSL. Marked asymmetry is present TC is consistent with its lower central pressure at land- with a maximum of 82 m sϪ1 in the northerly flow east fall. of the eye and 67 m sϪ1 in the southerly flow west of the The maximum surface winds are located in the east- eye over land (Fig. 8b). Several other, weaker, low-level ern half of the circulation over water, collocated with a wind maxima are evident to the west of the eye in Fig. maximum in low-level cyclonic vorticity and a very 8b, extending as far as 50 km inland. strong gradient of equivalent potential temperature in In addition to the differences in wind structure and the eyewall (Figs. 8a and 6b). An analysis of the sea intensity, the CTRL and NOTOPOG TCs show signif- level pressure field and the surface wind vectors reveals icant differences in their simulated reflectivities at land- substantial cross-isobaric flow over land where the sur- fall. The NOTOPOG TC is characterized by a region of face wind speed is generally 15–20 m sϪ1 less than over relatively high reflectivity in the eastern half of the cir- the ocean owing to an increase in surface roughness. culation offshore, with 55 dBZ up to 5 km MSL (Fig.

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Ϫ FIG. 9. (a) Distribution of TKE (J kg 1) calculated at the lowest model sigma level (ϳ19 m) at the time of landfall for CTRL simu- lation. Terrain elevation contours are in black with interval of 200 m. (b) Distribution of TKE (gray shading) and the magnitude of the vertical shear in the lowest 100 m (black contours) at the time of landfall for the CTRL TC. (c) As in (b), but for the NOTOPOG TC.

8c). In contrast, the CTRL TC (Fig. 5c) is more asym- surface that may cause significant damage. Figure 9 metrical, with similar high values of reflectivity, ap- shows the turbulence kinetic energy for the lowest proaching 55 dBZ, in the western half of the circula- model sigma level (ϳ19 m) during landfall for both tion over land. This region of high reflectivity results CTRL (Figs. 9a,b) and NOTOPOG simulations (Fig. from a combination of strong frictional convergence 9c). For the CTRL simulation, a marked difference in upstream (nearer to the coast) and enhanced vertical TKE is evident between land and water surfaces, with motion as onshore flow is forced upward by the orog- values up 180 J kgϪ1 over land compared with less than raphy (not shown). A band of enhanced reflectivity 80JkgϪ1 over water. This difference is largely the re- caused by the melting of precipitation near the freezing sult of enhanced shear production of TKE over land level (ϳ5 km) is present in both simulations (Figs. 5d where the increased surface friction acts to reduce the and 8d). The upward bulge of the freezing level within surface wind speed while the flow aloft maintains its the eye signifies warming caused by strong subsidence. velocity. Indeed, the spatial distribution of the vertical This feature is more pronounced in the NOTOPOG TC shear in the lowest 100 m is in very close agreement during landfall and is consistent with both the stronger with the spatial distribution of TKE. This increase in warm core and lower central pressure relative to the the boundary layer vertical wind shear, caused by flow CTRL TC. transition from smooth to rough surfaces, is well docu- mented in the literature (e.g., Powell 1982, 1987). There c. Boundary layer turbulence are several TKE maxima evident in Fig. 9a. The first of The production of turbulence within the TC bound- these is located on the stretch of coast exposed to the ary layer is linked closely to maximum wind gusts at the very strong onshore flow south of the TC eye. The TKE

Unauthenticated | Downloaded 10/05/21 10:04 PM UTC 4344 MONTHLY WEATHER REVIEW VOLUME 136 there is maximized only in the immediate vicinity of the the wind by orography are later exposed to locally high coast and is consistent with very large shear in the low- wind speeds as the TC moves westward and the flow est100mupto30msϪ1. A second region of large TKE changes direction. is evident in a narrow band about 20 km west of the eye As the TC approaches Mt. Bellenden Ker (Fig. 1b) associated with strong vertical wind shear. Finally, sev- from the southwest, westerly winds impinging on the eral local maxima of TKE are evident over the wind- western slope of the mountain range acquire enough ward slopes of Mt. Bartle Frere and Mt. Bellenden Ker, momentum and depth to pass freely over the top of the both of which are located outside the eyewall at 2300 ridge and down the other side. Figure 11 is an east–west UTC. The elevated terrain acts to speed up the winds in cross section of the near-zonal winds over Mt. Bel- the boundary layer and increase the production of lenden Ker 20 min after landfall (2320 UTC). A maxi- shear-induced turbulence locally. For instance, the mum wind speed of 68 m sϪ1 is evident over the top of 50-m wind speed over the eastern slopes of Mt. Bartle the ridge where the flow is locally accelerated. An ob- Frere (elevation ϳ1200 m) is 56 m sϪ1 whereas the sur- served westerly wind gust of 82 m sϪ1 was recorded face wind is only 32 m sϪ1, resulting in extremely high during the passage of Larry at roughly the same loca- shear in the lowest 50 m. In contrast, over the sheltered tion, suggesting good agreement between the simulated northern slopes of the mountain (elevation ϳ960 m) winds over elevated terrain and observations, despite the 50-m wind is only 16 m sϪ1. Similar speedup/ the smoother model terrain compared with actual orog- sheltering effects that coincide with distinct maxima raphy. Significant downslope winds are also evident on and minima of TKE are also noted over and around Mt. the lee side of the range where the near-surface west- Bellenden Ker farther to the north. erly wind speeds range between 50 and 60 m sϪ1. Con- The NOTOPOG simulation also reveals distinct re- sequently, the mountains play an important role by fa- gions of locally high TKE values during TC landfall, cilitating the transport of pockets of higher-momentum measuring up to 180 J kgϪ1 where the maximum tan- flow down toward the surface. The Cairns region, lo- gential winds to the south of the eye experience a sud- cated about 35 km to the north of Mount Bellenden den increase in surface roughness (Fig. 9c). The overall Ker, also experiences these downslope winds but to a structure is characterized by a ring of high TKE around lesser degree because of the smaller mountains and the inner core of the circulation, particularly over land, larger distance from the TC eye. The simulated west- and several smaller yet distinct bands of lesser TKE at erly winds in the Cairns region range from about 25 to larger radii. A noteworthy aspect of the TKE distribu- 30 m sϪ1 as the TC passes to the south (not shown). tion in the NOTOPOG simulation is the apparent e. Downslope winds in the Port Douglas region banding structure over land. These bands contain lo- cally high values of TKE and are even more distinctive The morphology of the north–south-oriented moun- as the TC moves farther inland, particularly when the tain range situated about 13 km west of the popular core of the circulation is entirely over land (not shown). tourist town of Port Douglas (Fig. 1b), with its gentle The physical mechanisms behind these finescale bands, windward rise and steep leeward slope, is well-suited including their role in the exchange of heat and mo- for generating severe downslope winds (e.g., Smith mentum within the TC boundary layer, are beyond the 1977; Lilly and Klemp 1979; Hoinka 1985). Indeed, scope of the current study. CTRL successfully captures such an event. At 2300 UTC, the u-w component of the wind ranges from 4–8 d. Influence of orography on TC winds Ϫ Ϫ ms 1 on the windward side of the range to 16–20 m s 1 As introduced in the previous section, the interaction on the leeward side (Fig. 12a). As the flow in the lowest between the complex terrain and the TC wind field in 3 km approaches the mountain ridge from the west, it the CTRL simulation results in substantial variability of accelerates and is forced downward toward the surface low-level wind speed and direction. The terrain acts to on the leeward side. By 0030 UTC the near-surface alter the surface wind field before, during, and after wind speed reaches its maximum sustained value of ap- landfall, as shown by the streamlines in Fig. 10. Block- proximately 24 m sϪ1 (Fig. 12b). The associated moun- ing on the windward side of mountains, funneling tain wave, as depicted by the large vertical displace- through gaps, and leeside sheltering are all evident de- ment of lines of constant potential temperature, is also pending on angle of approach of the wind as well as the amplified during this time. Vertical velocities up to 4 amount of forcing. The strongest forcing is found near msϪ1 at 2000 m AGL are noted on the leeward side of the center of the vortex where the tangential flow is the range. The vertically propagating gravity wave is maximized, as should be expected. It is also apparent trapped by a critical layer at 10 km AGL where the from Fig. 10 that some regions initially sheltered from zonal wind goes to zero (not shown). Critical layers

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FIG. 10. Surface streamlines and terrain height (shaded with a 250-m contour interval) for the CTRL simulation for times (a) 2100, (b) 2300, (c) 0100, and (d) 0300 UTC. have been shown to play an important role in the am- downslope windstorm is possible if a TC makes landfall plification of mountain waves and subsequent intensi- to the south of the mountain range such that moderate- fication of severe downslope windstorms, both in nu- to-strong westerly flow impinges on the mountain crest. merical and analytical studies (see e.g., Clark and f. Wind gusts Peltier 1984; Smith 1985). While the downslope winds generated here are not severe, a slightly stronger TC at Wind gusts, as opposed to 1- or 10-min-averaged sus- landfall with stronger westerly winds at the crest of the tained winds, last only a few seconds but produce the mountain produces greater downslope winds of up to majority of structural damage. The wind gust estimate 36 m sϪ1 (not shown). The results suggest that a severe (WGE) method of Brasseur (2001) is used here to es-

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FIG. 11. (a) Surface streamlines and terrain height (gray shading with an interval of 200 m) at 2320 UTC 19 Mar 2006. (b) Vertical cross section of u–w wind component (gray shading with contour interval of 10 m sϪ1), circulation vectors (white vectors) and potential temperature (white contours every 3 K) at 2320 UTC 19 Mar 2006. The orientation of the cross section relative to the center of the TC is shown by the solid white line in (a). timate maximum surface gusts given certain character- homogeneous, even in the absence of topography, and istics of the model boundary layer. The assumptions of that the estimated maximum gust does not decrease WGE are (i) wind gusts at the surface result from de- simply with distance from the center of the TC, but flection of parcels flowing in the (whole) boundary instead depends on both the shape and height of the layer, and (ii) the deflection process is accomplished by underlying terrain, and the internal characteristics of large turbulent eddies that transport air parcels to the the TC boundary layer. surface. Analyses of these parameterized wind gusts for For CTRL during landfall a region of maximum sur- both the CTRL and NOTOPOG simulations reveal face gusts with speeds ranging from 80 to 90 m sϪ1 is that the spatial distribution of wind gusts is far from evident to the south of the eye over land and coincides

Ϫ FIG. 12. East–west vertical cross section of u–w wind component (gray shading with contour interval of 4 m s 1), potential tempera- ture (white contours every 3 K), and circulation vectors (white vectors) over the mountain range west of Port Douglas (see location map Fig. 1b) at time (a) 2300 and (b) 0030 UTC.

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Ϫ FIG. 13. Parameterized maximum surface wind gusts (gray shading with contour interval of 10 m s 1) for the CTRL simulation at time (a) 2300 (landfall) and (b) 0100 UTC. The maximum surface wind gusts for the NOTOPOG simulation at time (c) 0009 (landfall) and (d) 0209 UTC. The letters A and B in (a) and (c) show the location of the wind profiles in Fig. 14. with the area of maximum TKE shown in Fig. 9a. The west of the eye (as shown in Fig. 14) reveals accordance maximum gust is roughly twice the speed of the surface between the maximum wind within the PBL and the wind for the same time and location. An analysis of maximum estimated gust at the surface. The profile wind speed profiles in the lowest 2 km to the south and labeled A in Fig. 13, west of the eye, is characterized by

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over an elevated plateau where tangential flow and translation speed are maximized. These wind gusts are about 10 m sϪ1 stronger than the gusts produced in the NOTOPOG TC for the same storm-relative location (i.e., south of the eye) and time after landfall, despite the lower central pressure of the NOTOPOG TC. Al- though the strongest gusts are in the southern half of the circulation after landfall, gusts up to 40 m sϪ1, ca- pable of causing damage, extend 70 km northeast of the eye to Cairns and south along the coast. In contrast to the CTRL TC, the parameterized wind gusts for the NOTOPOG TC during landfall are stron- gest in the eastern half of the circulation over the ocean and range from 70 to 80 m sϪ1 (Fig. 13c). These maxi- mum gusts are about 1.6 times greater than the surface wind speed for the same time and location. As the TC moves inland, the region of maximum gusts moves cy- clonically (clockwise) around the vortex to the southern half of the circulation 2 h after landfall (Fig. 13d). This transition in the location of maximum surface gusts is not surprising given the westward motion of the TC, resulting in an increase of the tangential winds south of the eye due to the translation speed of the vortex. Gusts of up to 50 m sϪ1 extend at least twice as far south of the eye as they do north (Fig. 13d). g. Rainfall

FIG. 14. Vertical profiles of wind speed in the lowest 2000 m Significant differences in the accumulated rainfall are corresponding to locations A and B as shown in Figs. 13a,c, for the noted between the CTRL and NOTOPOG run, both CTRL (solid line with triangles) and NOTOPOG (dashed line before and after landfall, as shown in Figs. 15 and 16. In with circles) simulations. the 12 h prior to landfall the heaviest precipitation oc- curs to the north of the storm track in the deep con- a particularly strong jet of 83 m sϪ1 at 800 m AGL (Fig. vection surrounding the TC eye. This bias is evident in 14a), suggesting a potential for these locally very strong both simulations, though it is particularly pronounced winds to be transported to the surface by turbulent ed- in the CTRL simulation. It is also in agreement with the dies and/or convectively driven downdrafts. Moreover, radar-observed convective asymmetries. Both the the wind increases from 26 m sϪ1 at the surface to 76 CTRL and NOTOPOG simulations are characterized msϪ1 at 500 m AGL resulting in favorable environmen- by an almost continuous swath of accumulated rainfall tal shear for tornadogenesis (e.g., Novlan and Gray of greater than 200 mm to the north of the TC track, 1974; McCaul 1991). Although no tornadoes were re- with embedded isolated amounts of more than 300 mm. ported, damage patterns were consistent with tornadic The apparent asymmetry in the accumulated rainfall is winds (more information available online at http:// likely due to the deep-layer ambient environmental www.ga.gov.au/image_cache/GA8468.pdf). Wong and shear. An analysis of the 850–200-hPa deep-layer shear Chan (2007) showed that supergradient tangential reveals that the TCs in each simulation are embedded winds near the top of the PBL are primarily the result in generally weak (Ͻ5msϪ1) easterly shear in the 12 h of the vertical advection of the radial wind, in agree- leading up to landfall. Chen et al. (2006) show that for ment with Kepert and Wang (2001). Inspection of the TCs in the Southern Hemisphere, enhanced precipi- parameterized wind gusts at landfall reveals two dis- tation is favored to the right of the deep-layer en- tinct maxima: one south of the TC eye in the onshore vironmental shear vector, in agreement with the north- flow and the other west of the eye and over land. ward bias of heavy precipitation in Fig. 15. For CTRL, The asymmetrical distribution of wind gusts persists analyses of the column-integrated cloud liquid water as the TC moves inland (Fig. 13b), with the strongest content indicates maximum values occur generally in gusts confined to the southern side of the circulation the front-left quadrant of the vortex, upstream of the

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FIG. 15. The 12-h accumulated explicit precipitation in D3 prior to landfall (gray shading with a 50-mm contour interval) for (a) the CTRL simulation and (b) NOTOPOG simulation. The 100-mm contour is outlined in black and the 200- and 300-mm contours are outlined in white. Terrain elevation is shown by dark gray contours with a contour interval of 100 m. heavy precipitation in the front-right/rear-right quad- amount of postlandfall precipitation. The highest 3-h rants. It is likely that strong cyclonic advection around observed rainfall total to 2300 UTC 19 March 2006 was the eye plays an important role in the final distribution 139 mm at The Boulders on Babinda Creek, at the base of precipitation. The orography in CTRL has a signif- of Mt. Bellenden Ker. The modeled 3-h accumulated icant influence on the location and amount of rainfall precipitation is 150 mm, in close agreement with the before, during and after landfall. In the 3 h prior to observed total. By 0500 UTC 20 March 2006, with the landfall (2000–2300 UTC) the heaviest precipitation oc- TC well inland, the accumulated rainfall amounts de- curs to the north of the TC track with a distinct swath crease substantially, ranging generally from 75 to 100 of more than 200 mm (Fig. 16a). Isolated amounts of mm (Fig. 16c). Local rainfall amounts exceeding 100 between 75 and 100 mm are evident over the windward mm occur over the mountains south of Innisfail, far slopes of the coastal range where strong tangential flow from the TC core, and are due to strong onshore winds. is forced upward resulting in increased vertical motion The amount and pattern of accumulated rainfall in (not shown). From 2300 UTC 19 March to 0200 UTC 20 NOTOPOG differs substantially from the CTRL simu- March 2006, maximum accumulated rainfall values of lation. In the 3 h prior to landfall, the heaviest precipi- up to 225 mm are evident over the highest orography to tation occurs south of the eye with amounts exceeding the north and south of the eye (Fig. 16b). The maxi- 200 mm (Fig. 16d). This southward bias, in contrast to mum on the southern side of the circulation is absent in CTRL, is due partly to the enhanced low-level conver- the preceding 3-h period (Fig. 16a) highlighting the im- gence caused by the shape of the coastline, including portance of orography in determining the location and Hinchinbrook Island. The rainfall amounts in the 3 h

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FIG. 16. The 3-hourly accumulated precipitation in D4 (gray shading with a 25-mm contour interval) for the CTRL TC at (a) 2300, (b) 0200, and (c) 0500 UTC, and for the NOTOPOG TC at (d) 0009, (e) 0309, and (f) 0609 UTC. The 50- and 100-mm contours are outlined in black. The 150- and 200-mm contours are outlined in white. Terrain elevation is shown by dark gray contours with a contour interval of 100 m.

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Large differences are evident in the TCs respective decay rates, from hour to hour (Fig. 17). The influence of orography on the decay of the CTRL TC is indicated by the nonsteady nature of the decrease in maximum surface wind speed as a function of time, compared to the relatively smooth downward decay rate for the NOTOPOG TC. As the CTRL TC moves westward over the Atherton Tablelands the maximum surface wind speed actually increases from 32 to 35 m sϪ1 over a 2-h period. These results suggest that simple TC decay models (e.g., Kaplan and DeMaria 1995) should be used with caution when being applied to TCs that make landfall in regions of complex terrain.

FIG. 17. Time series of maximum surface wind speed for the CTRL simulation (solid line with black triangles) and 5. Discussion and conclusions NOTOPOG simulation (dashed line with white circles) plotted every hour from the time landfall (0 h) to 12 h after landfall. The MM5 mesoscale model has been used to inves- tigate the effects of complex terrain on TC Larry that made landfall on the northeastern Australian coast as a after landfall (up to 175 mm; Fig. 16e) are less than severe category 4 TC. The simulated track, speed of in the CTRL simulation (225 mm), despite the movement, strength, size, and decay rate of the simu- NOTOPOG TC’s greater intensity. As the TC moves lated storm (CTRL) matched the observations closely, inland the precipitation increases relative to the CTRL and reproduced many of the features of TC Larry over TC with a large region of 75–100 mm and an embedded land. To compare with CTRL a contrasting simulation band of more than 100 mm (Fig. 16f). An analysis of the (NOTOPOG) was carried out in which the orography total precipitable water field suggests that the absence was completely removed, but other surface properties of steep coastal terrain allows the NOTOPOG TC to were retained. Both experiments were initialized using retain its moisture content farther inland compared to a bogus vortex to produce a TC with intensity close to the CTRL TC. the observed intensity of TC Larry as it approached landfall. Tropical cyclone genesis occurred only with a h. Decay after landfall bogus initial vortex, owing to relatively coarse model The overall TC decay rate from landfall to 12 h after input data. landfall is relatively insensitive to the underlying ter- Previous studies on the orographic influence of TCs rain, though there are marked differences in the trends have focused mainly on the effects of mountain ranges of maximum surface wind speed from hour to hour on TC motion. Examples of mountain ranges that affect (Fig. 17). The CTRL TC makes landfall with a central TC motion include the Central Mountain Range of Tai- pressure of 935 hPa; 12 h after landfall the central pres- wan and the Sierra Madre of Mexico (e.g., Chang 1982; sure is 991 hPa—an increase of 56 hPa. The CTRL Zehnder 1993; Yeh and Elsberry 1993a, b; Zehnder and TC’s central pressure after landfall is in good agree- Reeder 1997; Lin et al. 2002, 2005, 2006). The problem ment with Larry’s best-track central pressure (Fig. 4). of orographically enhanced mesoscale precipitation as- The NOTOPOG TC makes landfall with a central pres- sociated with landfalling TCs has been the focus of a sure of 928 hPa that increases by 55 hPa to 983 hPa 12 number of studies (e.g., Geerts et al. 2000; Wu 2001; Lin h later. Thus, the simulated TCs decay at a rate relative et al. 2002; Wu et al. 2002). Here, the focus is not only to their respective central pressures at landfall, consis- on the primary variables such as wind, pressure, and tent with several previous studies (e.g., Schwerdt et al. precipitation, but also on how complex terrain acts to 1979; Ho et al. 1987; Kaplan and DeMaria 1995). modify the TC boundary layer. This is an important For both CTRL and NOTOPOG simulations, the issue because the damaging winds associated with land- largest decrease in surface wind speed occurs in the first falling TCs are intrinsically related to characteristics of hour or two following landfall, as expected owing to the the underlying surface. The present study explores sudden increase in surface roughness over land. The these complex interactions using high-resolution terrain sudden decrease in maximum surface wind is particu- data (ϳ900 m) and two fine nest domains (D3 and D4) larly pronounced in CTRL, where a reduction of 30 with storm-scale grid spacing (3 and 1 km, respectively). msϪ1 occurs in the first hour after landfall. It was found that complex terrain has a significant

Unauthenticated | Downloaded 10/05/21 10:04 PM UTC 4352 MONTHLY WEATHER REVIEW VOLUME 136 impact on the wind and rainfall distribution of a land- the NOTOPOG TC were lower immediately following falling TC. In the CTRL simulation, the orography is landfall, but increased relative to CTRL as the system found to affect the TC track, intensity, and structure moved farther inland. well prior to landfall by modifying the three-dimen- Parameterized maximum surface wind gusts were sional environmental wind upstream. Differences in spatially consistent with regions of locally high wind the deep-layer vertical wind shear close to the TC core speeds in the TC boundary layer. The distribution and had a large impact on intensity, track and rainfall intensity of these gusts appear to be controlled primar- distribution prior to landfall. As the NOTOPOG TC ily by vertical mixing associated with strong boundary approached landfall, a change in the near-core deep- layer vertical wind shear that is modified by the shape layer shear from circularly asymmetric to circularly and height of the underlying terrain. Locally high val- symmetric occurred, resulting in a transition of the spa- ues of TKE occurred over ridges where acceleration of tial distribution of heaviest precipitation. the TC’s tangential winds increased vertical wind shear. The direct influences of orography on the three-di- However, given the simplistic assumptions of the WGE mensional wind field of the CTRL TC are pronounced. scheme the maximum surface gusts are probably over- Locally high surface wind speeds, ranging from 36 to 40 estimated. msϪ1, occur over Mount Bellenden Ker and Mount Finally, an analysis of the postlandfall boundary layer Bartle Frere. The 50-m winds above the peaks were wind field for the NOTOPOG simulation revealed fine- higher, reaching 68 m sϪ1 over Mount Bellenden Ker as scale banding features that appear related to the in- the TC passed to its south. On the sheltered lee sides, creased surface roughness over land and the resultant wind speeds were much lower (ϳ8msϪ1). Analyses of strong low-level vertical wind shear. Similar boundary the winds above the surface for CTRL during landfall layer features have recently been documented in theo- revealed several wind speed maxima within the bound- retical and observational studies (e.g., Foster 2005; ary layer, the strongest being 86 m sϪ1 at 631 m. These Morrison et al. 2005; Romine and Wilhelmson 2006) boundary layer jets produced strong low-level vertical and underline the need for further investigation regard- wind shear (Ͼ30 m sϪ1); a favorable environment for ing their possible connection with surface wind gusts. tornadogenesis associated with landfalling TCs. How- ever, the issue of how these boundary layer wind Acknowledgments. This research was supported by maxima are transported to the surface, resulting in funding from the Insurance Australia Group, Sydney, wind gusts, remains an elusive area in TC research. The Australia. We thank Dr. David Stensrud (NOAA/ boundary layer parameterization schemes used in me- National Severe Storms Laboratory) for valuable dis- soscale models are too crude to accurately reproduce cussions as well as Dr. Bruce Buckley, Jeff Callaghan, the small-scale exchanges of heat and momentum that and Peter Otto of the Australian Bureau of Meteorol- result in wind gusts within the TC boundary layer, par- ogy for providing details of the observed characteristics ticularly over complex terrain. of TC Larry. We also wish to thank the three anony- Strong downslope winds were produced in the CTRL mous reviewers for many beneficial suggestions, which simulation when westerly flow from the outer TC cir- helped to improve the manuscript. Finally, we acknowl- culation impinged on the north–south-oriented moun- edge Greg Pearson and Norm Henry of the New Zea- tain range west of Port Douglas. The shape of this range land Meteorological Service for providing the code to is well-suited for generating severe downslope winds, compute wind gusts in section 4f. with its steep leeside slope and gentle windward rise. During the most amplified phase of the mountain wave, REFERENCES the near-surface wind speed at the base of the range approached 24 m sϪ1. Even stronger surface winds of Bender, M. A., R. E. Tuleya, and Y. Kurihara, 1987: A numerical up to 36 m sϪ1 were observed in a separate simulation study of the effect of island terrain on tropical cyclones. Mon. Wea. 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