1. Introduction
Tropical cyclone (TC) landfall processes have been an important research problem because of the potentially huge destruction of lives and properties when the intense rainfall, severe winds, and storm surge batter coastal regions. Strongly asymmetric structures of wind and rainfall usually accompany landfall. Previous studies have documented the asymmetric convective activities during TC landfall. Generally speaking, rainfall patterns of landfalling TCs vary widely from case to case because various factors (e.g., land surface properties, topography, vertical wind shear) could drive the asymmetries. Some idealized modeling studies (e.g., Chan and Liang 2003) have been carried out to highlight the nonuniform surface fluxes as a factor to generate the asymmetries. A more detailed summary of those results can be found in Chan et al. (2004).
Some observational studies have documented the wind distribution of landfalling TCs (e.g., Powell 1982, 1987; Powell et al. 1991; Powell and Houston 1996, 1998). It was noted that maximum wind and gusts tend to occur near the most convective regions (e.g., Parrish et al. 1982). Possible mechanisms include the downward transport of high-velocity air by precipitation-induced downdrafts, the spreading of downbursts along the ground, and the formation of mesovortices (Willoughby and Black 1996). Powell (1987) argued that the wind asymmetry depends on the combined effects of land–sea roughness differences, background environmental flow, and storm translation.
The wind asymmetry has also been studied in a few idealized numerical simulations. Tuleya and Kurihara (1978) found that the roughness contrast between land and sea creates quasi-steady convergence (divergence) and negative (positive) relative vorticity zones along the coastline where the flow is onshore (offshore). Tuleya et al. (1984) noted the possibility of a temporary displacement of the surface circulation center from the pressure center at landfall. Jones (1987) found that the increase of rainfall to the left of the landfalling TC relative to a nonlandfalling one is related to the enhanced relative inflow in the left-front quadrant near landfall.
Kepert (2002) noted that the asymmetric structure of Hurricane Danny at landfall is similar to that induced by motion, where, in each case, it is forced by frictional asymmetry. Recently, Kepert (2006a,b) has studied Hurricanes Georges and Mitch and also found close analogy between the motion-induced and landfall-induced asymmetry. There would also be differences between the structures forced by motion and land friction (Kepert 2002) because the asymmetric friction would be discontinuous rather than smooth, as in the motion-induced case, and the asymmetry of friction is much stronger for landfall. Furthermore, the asymmetry in friction would not cover the whole storm if the TC center is some distance from the landfall position.
In an attempt to study the vortex drift associated with the asymmetric wind and convection near landfall, Wong and Chan (2006, hereafter WC06) found that the planetary boundary layer (PBL) convergence of the TC core could become strongly asymmetric even when the core TC circulation is completely away from the coast. They proposed the surface-induced asymmetric convergence as an important factor that affects the asymmetric convective activities of the TC core. Previous studies have not focused on such features, and a more thorough investigation has yet to be done.
This paper therefore focuses on understanding the asymmetric wind distribution of landfalling TCs. The hypothesis that the adjustment of air moving away from a roughness discontinuity could produce an asymmetry in the TC core is also addressed. Idealized numerical experiments are performed with the fifth-generation Pennsylvania State University (PSU)–National Center for Atmospheric Research (NCAR) Mesoscale Model (MM5: Dudhia 1993). The numerical experiments and simplifications of the model physics are described in section 2. Results on the wind asymmetries and the associated convergences are shown in section 3. Section 4 presents comparisons with a full-physics experiment, followed by the discussion and conclusions in section 5.
2. Model and design of experiments
a. MM5
The experimental design is the same as in WC06. The triply nested square domains have horizontal resolutions of 45, 15, and 5 km and span 6750, 2250, and 1200 km, respectively. There are 26 vertical layers. An f plane at 15°N is used. The surface fluxes and vertical diffusion are determined from a modified Mellor–Yamada level-2.5 PBL scheme (Janjić 1994). For the full-physics experiment, an explicit moisture prediction scheme (Dudhia 1989) and the Betts and Miller (1986) cumulus parameterization are also used.
The specified vortex is also the same, which includes the horizontal variation of the tangential wind profile first used by Wang and Li (1992) and the vertical variation of the wind used by Wong and Chan (2004). The vortex is then spun up for a 24-h period after which it attains an intensity of ∼976 hPa. The axisymmetric part of the mass and wind fields then form the initial conditions of the conceptual experiments (see below for details) with the TC center placed at the center of the model domain. In all the experiments, a portion of the surface is specified to be land of roughness length 0.5 m (sea roughness depends on wind strength) so that the TC is located at 50, 100, 150 km onshore and inland, and right on the straight coastline, giving a total of seven experiments (Table 1). The mass fields are held fixed during the 48-h simulation without feedback from moisture or heat sources. Two more experiments with sea-only and land-only surfaces are also performed as control. As it turns out, a quasi-steady state is reached for all the experiments. The hourly results of the second day (t = 24 to t = 48 h) are averaged and analyzed.
The use of the time-invariant axisymmetric mass fields during the numerical integration implies that the intensity of the TC is unchanged, the mass adjustment is negligible, and there is no TC movement. The results of the full-physics experiment RD of WC061 will illustrate that mass adjustment is not severe, and the results of the conceptual experiments would still be applicable to a nonsteady condition of a weakening TC (see section 4). Although TC motion is not considered, we expect that the roughness contrast would be the most dominant reason for PBL asymmetry at landfall except for unusually fast movement.
There are other reasons why we do not simply analyze the results of the RD experiment. The asymmetry in vertical motion in the RD experiment could result from vortex-scale vertical wind shear (see WC06 and Kepert 2006b) and not just the adjustment of air moving from one surface to another. By fixing the mass fields, we expect to remove the shear effect and determine more clearly the effect of roughness contrast on the asymmetries. Moreover, with the same mass fields for all of the conceptual experiments, comparisons of the magnitudes of winds among the experiments could be made. If the mass fields are allowed to change as in the RD experiment, the TC eventually evolves into different intensities at different times, making comparisons difficult.
Analytical solutions of the PBL flow over a homogeneous surface have been successfully obtained by Kepert (2001) for stationary and moving TCs. It might be possible that a similar analytical solution can be obtained for a stationary TC on a land–sea interface. However, the derivation would be much more difficult, so a numerical solution is opted instead.
b. Physical processes examined in the idealized experiments
3. Results for the conceptual experiments
a. Surface wind asymmetry
The “surface wind” in this study is referred to the 10-m level wind, which is a model output variable based on the surface layer similarity theory. The maximum surface gradient wind speed is about 31 m s−1 at a radius of 50 km. As expected, the surface wind field for the sea-only experiment (Fig. 1a) is axisymmetric and characterized by the presence of inflow and a reduction of the tangential wind speed to a value less than that of the gradient wind. The maximum surface tangential and total winds are 24 and 26 m s−1 respectively, both found at a radius of 50 km. The wind reduction factors, computed as the ratio between the wind speeds and the gradient wind speed, are 0.77 and 0.84, respectively. For the land-only experiment the inflow angle and the speed reduction are larger (Fig. 1b), and the radius of maximum wind (RMW) is smaller, at 40 km. The surface wind reduction factors for the tangential and total winds are, respectively, 0.43 and 0.57. These factors then decrease to nearly constant values at large radii. In addition, the maximum surface inflow is slightly larger in the land-only experiment (10.8 versus 9.6 m s−1) but is weaker at larger radii.
For the experiments with a land–sea interface, the most salient characteristic of the surface wind field is, as expected, the smaller surface wind speed (both tangential and radial components) over the land than the sea surface (Figs. 2a–g). Some other important features also require explanations. First, the winds near the TC core are asymmetric. When the TC is 150 km from the coast before landfall (Fig. 2a), the eye/eyewall region is located completely over the sea. However, the radial inflow is clearly stronger to the left/front left (relative to the direction facing land) of the TC and the tangential wind is also asymmetric and strongest to the rear left/rear. This implies that the local wind is not determined solely by the local surface roughness. On the contrary, the radial inflow for the postlandfall cases (L50, L100, L150) is weaker on the right/rear-right side where the flow is onshore and toward the TC (Figs. 2e–g, respectively), although the asymmetry of the core tangential wind is less clear. At the time of landfall, radial inflow is strong (weak) to the rear left (front right) of the TC (Fig. 2d). In this case, the offshore (onshore) flow is toward the rear-left (front right) quadrant. The maximum tangential wind occurs over the sea as the front half of the TC is over the land surface.
Second, for prelandfall (S150, S100, S50) and landfalling (LS) cases (Figs. 2a–d, respectively), the strongest inflow near the TC and the maximum tangential wind that occurs approximately 90° downstream at a smaller radius are associated with the offshore flow. This may seem counterintuitive as the offshore flow originates from the land surface at smaller wind speeds. The dynamics of this result is discussed further below.
Third, a comparison of the magnitude of the flows with the control experiments could reveal the importance of the “transitional” effect near the coast. In regions where the flow has moved over the same surface, the wind carries the characteristics of that surface. For example, in the LS case, the tangential and radial flows more than 100 km offshore are close to those in the sea-only experiment (cf. Figs. 1a and 2d), while the flows more than 100 km inland are close to those in the land-only experiment (cf. Figs. 1b and 2d). In the “transition region,” the flow never quite attains either the sea-only or land-only states because the boundary layer adjustment time 1/I (where I is the inertial stability; see Eliassen and Lystad 1977) is of similar order as the rotational period r/υ near the storm core. The maximum inflow for the LS case (Fig. 2d) has a magnitude >15 m s−1 and is actually larger than those in both the sea-only and land-only experiments. Moreover, the maximum tangential wind speed reaches over 25 m s−1 and is also larger than both control experiments. For the sea-only experiment, the surface tangential wind reduction factor (Fig. 3a) reaches 0.8 in only a small region, but that for the LS experiment has a region of near 1 (Fig. 3b), significantly larger than the control experiments.
These asymmetric features can be explained by (2.1) and (2.2). For the surface wind, which is often of primary interest, the sudden additional force (and acceleration) due to a change of the roughness would be in the direction of the wind for the offshore flow and in the opposite direction of the wind for the onshore flow. Therefore, if the land surface is rough, the inflow angle is large and the offshore flow would undergo strong inflow acceleration. This would explain why the maximum surface radial inflow associated with the offshore flow for prelandfall and landfalling cases (Figs. 2a–d) could be stronger than both control experiments (Figs. 1a,b). Furthermore, the maximum total wind could be associated with the offshore flow if the air could move very close to the TC center and attain more kinetic energy. In fact, a contraction of the left (the direction relative to that pointing from sea to land) eyewall of Hurricane Danny in 1997 with associated maximum wind and rainfall was observed (Blackwell 2000).
This inflow acceleration could also explain why the surface tangential wind is comparable to the gradient wind in the conceptual landfall experiments (Fig. 3). For small friction over sea (assumed negligible to facilitate the following discussion), the air that has just moved offshore is subjected to inward acceleration until the Coriolis and centrifugal forces balance the pressure gradient force in (2.1). However, the radial flow is still negative, which means that the tangential wind and total wind continue to increase according to (2.2). At this point, the total wind is already stronger than the (local) gradient wind owing to the presence of the inflow. Gradually the radial acceleration becomes positive (or the tangential wind has become supergradient) and the radial inflow ceases at some small radius. Note that there is also a region of outflow near the maximum tangential wind at a slightly smaller radius (e.g., Fig. 2b), consistent with the negative sign of dυ/dt downstream of the area of maximum surface tangential wind [see (2.2)].
To elaborate further, we consider the trajectory of the flow in the lowest model level (σ = 0.995; 43 m) in the LS case. Four streamlines are considered (Fig. 4). Each of the streamlines represents air movement for a 2-h period with the air located at 150 km either rear, right, front, or left of the TC center at the middle of the 2-h period. The results show that the onshore (offshore) flow associated with the right (left) path is accompanied by stronger tangential deceleration (inflow acceleration) than inflow deceleration (tangential acceleration) (Figs. 5a and 5b, respectively). The acceleration is primarily related to the change of surface roughness. Decomposition of the vertical diffusion of momentum into tangential and inflow components show that the evolution is consistent with the accelerations (Figs. 5c and 5d, respectively). For the left path, the tangential component of vertical diffusion is actually positive when the air is near the coast, which means that the tangential momentum flux from the upper layer is stronger than the tangential momentum flux at the surface (i.e., surface stress) and results in an acceleration.
As noted by Kepert (2002), an analogy between the motion-induced asymmetry could be made in order to explain the asymmetric features. For example, in the LS case the maximum inflow occurred at the rear left, and could be realized by motion from right to left. The asymmetry has, however, higher wavenumber (>1) components. This is partly due to the configuration of the land–sea interface and, partly, the large frictional asymmetry that causes significant deviation from the linear regime in deriving the solution in Kepert (2001). Moreover, for prelandfall and postlandfall TC positions the analogy is not obvious. For example, for prelandfall the inflow is strongest at the left or front left. Note also that the maximum tangential wind that occurred at inner radii has been explained as enhanced angular momentum advection by Kepert and is similar to the explanation of the maximum total wind in terms of the work done by pressure in the above.
b. Wind asymmetry above the surface
The study of the wind above the surface is also of great importance because the higher winds could be brought down to the surface as gusts, and the relationship between upper-level and surface winds has always been used for the estimation of surface winds based on upper-level measurement (e.g., radar, reconnaissance flight). The steady-state structure of the asymmetric flow is required to bring inertial waves to a halt (Kepert 2001).
To begin, we first investigate the maximum tangential wind above the surface for the sea-only and land-only experiments. The supergradient nature of the tangential flow near the top of the PBL of a TC is simulated (Kepert 2001; Kepert and Wang 2001). The maximum tangential2 wind of 39 (42) m s−1 for the sea- (land-) only experiment is located at σ = 0.94 (σ = 0.885), or 522 (1024) m MSL, about 50 km from the TC center (Figs. 6a and 6c). It is about 24% (37%) supergradient, higher than the 10%–25% range found in Kepert and Wang (2001). The vortex is inertially stable in the current study, and the higher winds would be related to the different PBL parameterization used here because additional sensitivity experiments show that the strength of the supergradient wind is sensitive to the choice of PBL parameterization (not shown). The figures should then be interpreted with caution. Vertical advection of the radial wind in (2.1) is found to be the dominant contributor to the supergradient flow (not shown). For the LS case, the height, depth, and strength of the azimuthal mean supergradient jet is in between the sea-only and land-only experiments (Fig. 6e). At 100 km from the TC center, the supergradient wind has become less significant for all cases (Figs. 6b,d,f).
The jet is clearly lower for the sea-only than the land-only experiment (cf. Figs. 6a and 6c). Moreover, the PBL is shallower at 50-km radius than at 100-km radius if we consider the depth of the radial inflow, and it is most evident in the sea-only case (cf. Figs. 6a and 6b). These results are consistent with Kepert (2001), who found that the jet and the PBL depth would be scaled as (2K/I)1/2, K being the vertical diffusivity and I the inertial stability. Note that one could also define the PBL top to be the height at which the turbulent kinetic energy has dropped below a threshold value, similar to Gayno et al. (1994). Contrary to the above depth scale, this definition gives a higher PBL height at 50 than at 100 km from the TC center due to stronger shear production of turbulent kinetic energy. This seemingly contrasting result is merely a matter of definition.
The winds above the surface layer for the LS case are also asymmetric. The asymmetries of tangential and radial winds rotate anticyclonically with height (Fig. 7). At σ = 0.915, radial inflow is strongest to the front left and has become weaker (Fig. 7c). As at the surface, the region of maximum tangential wind is about 90° downstream of the region of maximum inflow so that the region of maximum wind is approximately in between. Such a rotation is similar to that found in Kepert (2001) for a moving cyclone. The flow has become symmetric at σ = 0.825 (Fig. 7f).
Kepert (2001) and Kepert and Wang (2001) studied the dynamics of supergradient winds. The linear effect is similar to the classical Ekman spiral, where supergeostrophic wind occurs near the top of the PBL. Vertical advection is found to increase the level of supergradient winds. As in Kepert (2001) and Kepert and Wang (2001), we would like to see how the supergradient flow is maintained. At σ = 0.915 for the LS case, maximum “supergradienticity” (the excess of the Coriolis and centrifugal forces over the pressure gradient force) occurs to the left (Fig. 8a). The effects of radial advection and vertical diffusion are found to be small (not shown), so the tangential and vertical advections explain most of the asymmetries (Fig. 8b). Moreover, vertical advection is more dominant over the tangential advection (Figs. 8d and 8c, respectively).
c. Vorticity and divergence
An accompanying consequence of the asymmetric wind is the asymmetric vorticity and convergence. Except for the strongly positive relative vorticity in the TC core, the surface relative vorticity is characterized by a band of negative (positive) relative vorticity associated with the onshore (offshore) wind (not shown), in agreement with Tuleya and Kurihara (1978). The vorticity change is related to the friction term in the vorticity equation. Since the surface convergence (divergence) along the coast coincides with the negative (positive) vorticity, the magnitude of the vorticity band decreases downstream of the coast.
WC06 hypothesized that the enhanced inflow could drive the asymmetric convection. Moreover, the degree of asymmetry is stronger for prelandfall than for postlandfall TCs. As can be seen from the previous subsections, the inflow asymmetry would lead to asymmetric divergence at the TC core. However, the convergence is also contributed by the speed divergence of the tangential wind. Therefore, each of the effects has to be taken into account.
We investigate, for each of the seven cases in Table 1, the phase of the wavenumber-1 (WN1) convergence (both surface and PBL) of the horizontal wind (−∇ · Vh) in three radial bands: 1) 0–50, 2) 50–100, and 3) 100–500 km from the TC center. For simplicity the “PBL flow” is defined to be the averaged flow in the layer 1.0 > σ > 0.9. For the surface (Fig. 9, large symbols), stronger convergence occurs to the left in the 0–50-km band except for the LS case, the only case in which the coastline is within the radial band. In the 50–100-km band surface convergence is larger to the right for cases where the coastline is within the radial band (L50, LS, S50) and is shifted slightly to the front right for the other cases. Surface convergence is consistently stronger to the right in the 100–500-km radial band as the convergence (divergence) effect at the coastline is included in all experiments. Results using a stronger vortex are similar (not shown).
The phase of PBL convergence (Fig. 9, small symbols) is similar to that of the surface except for the LS case, where the PBL convergence is also stronger to the left as in the other experiments. In other words, the stronger surface convergence associated with the onshore wind does not result in stronger PBL convergence to the right. This could be explained by referring back to Figs. 7a–c, where the discontinuity in tangential wind along the coastline has become much weaker above the surface while the stronger inflow to the left (or front left) persists.
The asymmetric convergence is similar to that forced by motion (Kepert 2001), which for the boundary layer convergence (and hence the vertical velocity forced by the boundary layer) in the 0–50-km band is to the left. However, there are also distinct differences. For example, the surface convergence is largest to the right and the convergence for the 100–500-km band has a reverse relationship.
4. Comparison with the full-physics experiment
The results presented in section 3 assumed that the feedback from heat and moisture sources to the mass fields is small. This assumption may not be correct because most, if not all, TCs weaken upon landfall. Moreover, the asymmetric convergences and convective heating may distort the mass fields, and there may also be significant mass adjustments for the strong imbalance of the flow that has just moved over the surface of distinct roughness.
However, Kepert (2002) has noted that the simulated wind field matches quite well with observations, even when Hurricane Danny is not in a steady state. In this section the similarities and differences between the results presented in the previous section and those of the RD experiment of WC06 are discussed. The RD experiment is initiated with an intense TC 150 km offshore. Because of a large-scale asymmetry generated, the vortex is found to drift toward the coast. The focus in this section is on the effects of feedback on the mass fields and the wind in nonsteady intensity situations.
a. Sea level pressure and wind distributions
Past modeling and observational studies have shown that the surface pressure distribution remains fairly symmetric despite significant asymmetry of the wind (e.g., Tuleya and Kurihara 1978). The results of the RD experiment at t = 116 h and t = 122 h indicate that the sea level pressure had not undergone large asymmetric transition (Figs. 10a and 10b), and there is also no noticeable discontinuity of the pressure gradient across the coast. As revealed by the gradient winds (Figs. 10c and 10d), the pressure gradient is slightly higher to the right than to the left, probably related to the tendency of increasing surface pressure due to the convergence/divergence along the coastline at the storm scale, and is consistent with the landward drift of the vortex found in WC06.
The sea level pressure field of the RD experiment would therefore favor stronger winds to the right near landfall. However, the asymmetric surface inflow and tangential wind are found to be similar to those in the conceptual experiments (Figs. 10e and 10f; cf. Fig. 2). This shows that the effect of the asymmetric mass fields on the surface wind asymmetry is small compared to that forced by the land–sea interface. Note also that the TC is weakening, with central pressure rising from 910 to 923 hPa during this period, and the TC is much more intense than that in the conceptual experiments. This also supports the qualitative results obtained in the conceptual experiments, which are applicable to more general conditions.
To see more clearly the surface wind asymmetry and the similarities to the results in the conceptual experiments, the hourly results of the 144-h simulation are further examined. Just prior to landfall, the surface tangential wind is a maximum in the rear left (Fig. 11a), while the maximum surface inflow occurs to the left/front left (Fig. 11c). In other words, the maximum surface tangential wind occurs about 90° downstream of the maximum surface radial inflow, and both are associated with the offshore flow as in the conceptual experiments. Moreover, the radii of maximum tangential and maximum radial winds are smaller (∼20 km) just prior to landfall (Figs. 11b and 11d, respectively), which is also consistent with the arguments in section 3a concerning (2.3) and (2.4). The maximum surface gradient wind occurs to the right (Fig. 11e). After landfall, the vortex core breaks down with maximum surface gradient wind located in outer bands (Fig. 11f).
Furthermore, the surface wind reduction factor of the tangential wind increases when the TC is close to the coast (Fig. 12). It increases from about 0.65 when the TC is about 150 km from the coast to over 0.8 just prior to landfall. The results are also similar to those in the conceptual experiments (cf. Fig. 3).
b. Asymmetric convection
WC06 showed that the asymmetries of rainfall at inner and outer radii differ. Rainfall is larger over the left or front left within 100 km from the TC surface center over the 6-day simulation but, when the averaging distance is 300 km, rainfall is actually smallest to the left and strongest to the front, then gradually to the front right as the TC approaches the coast and finally moved over land. These results agree fairly well with the PBL convergence shown in Fig. 9 in the current study in two ways. First, the inner (0–50 km) and outer (100–500 km) PBL convergences are out of phase. Second, the phase of the PBL convergence in the middle (50–100 km) radial band rotates from front right to right as the vortex “approaches” the coast from the sea.
5. Summary and discussion
We have studied the asymmetric wind distribution of TCs forced by the different surface roughness between land and sea using MM5. Under a strong constraint—that the axisymmetric mass fields are time invariant during the model integration—and a straight coast separating land and sea, the center of the TCs are placed directly at the coast and at positions corresponding to prelandfall and postlandfall positions. The current study could be regarded as an extension of the study of Kepert (2001) and Kepert and Wang (2001), but using a surface of nonuniform roughness.
The results for the surface wind indicate that, apart from the onshore (offshore) convergence (divergence), significant asymmetry could also be triggered in the TC core. For prelandfall TCs, the acceleration of the offshore flow in the left-front quadrant of the TC, which is shown to be related to vertical diffusion, has a radial component that enhances the inflow. Work is done by the pressure gradient force and results in strong tangential wind and maximum wind associated with the offshore flow, which is located approximately in the left-rear and left quadrants. The maximum tangential wind is also shown to be stronger than the gradient wind. However, it should be noted that in reality the maximum wind for a landfalling TC does not necessarily have to be stronger than that without landfall since some TCs also become less intense as they approach the coast. For postlandfall TCs, radial acceleration of the onshore flow causes a reduced inflow in the right rear quadrant.
The winds above the surface layer are also asymmetric and the tangential component could also be supergradient. We have shown that vertical advection of the radial wind is the most important contribution to the supergradient tangential wind, in agreement with Kepert and Wang (2001).
Although many of the results are analogous to the motion-induced asymmetry, there are also differences. First, the motion-induced asymmetry is dominated by the wavenumber-1 component for slow translation speed or the linear approximation (e.g., Shapiro 1983; Kepert 2001), but the land–sea-induced asymmetry has forcing of higher (>1) wavenumbers and strong nonlinear interactions. Second, it is shown that the maximum surface and PBL convergences occur in distinct quadrants between inner and outer regions. While radial inflow gives stronger convergence to the left of the TC in the inner region (0–50 km), tangential acceleration (deceleration) gives stronger convergence to the right in the outer region (100–500 km). Such a radial dependence would not be true for motion-induced asymmetry.
Two processes could make a steady-state impossible. One is due to the mass-adjustment and intensity changes of the vortex, which involves a change of the mass fields and would occur even without land. We have shown in the RD experiment, even when it is not in steady state and we allow the free adjustment of mass, that the basic features in the conceptual experiments remain. These results are found to be similar to those from the full-physics version of the MM5, which suggests that in landfall situations the different roughness between land and sea leads to strong wind asymmetries that can be largely explained by the net acceleration associated with a deviation from gradient balance.
Another is due to the motion of the TC, which is not considered in this study. For a slow-moving vortex as in the RD experiment, the motion has negligible effect and the asymmetry is virtually related to the land–sea roughness contrasts alone. Kepert (2006b) has also shown with observations that the results remain true in the case of the slow-moving Hurricane Mitch. For larger translational speed, the modification to the land–sea-induced asymmetries would be related to the direction and speed of movement. For example, if the TC is moving from right to left (relative to the direction facing land in the Northern Hemisphere), the core asymmetries would be enhanced because the effect of motion could promote asymmetries of a similar azimuthal phase. On the other hand, the core asymmetries would be reduced (or even reversed in azimuthal phase for fast movement) if the TC is moving from left to right.
In reality a landfalling TC would also interact with topographic features of various complexity. An obvious follow-up work of the current study would be to consider topographic effects, perhaps in idealized simulations as in the current study or using real terrain data (e.g., South China).
Acknowledgments
This research is sponsored by the Research Grant Council of the Hong Kong Special Administrative Region, China Grant CityU 100203. Thanks are due to Jeff Kepert and the anonymous reviewers for their helpful reviews.
REFERENCES
Betts, A. K., and M. Miller, 1986: A new convective adjustment scheme. Part II: Single column tests using GATE wave, BOMEX, ATEX and arctic air-mass data sets. Quart. J. Roy. Meteor. Soc., 112 , 693–709.
Blackwell, K. G., 2000: The evolution of Hurricane Danny (1997) at landfall: Doppler-observed eyewall replacement, vortex contraction/intensification, and low-level wind maxima. Mon. Wea. Rev., 128 , 4002–4016.
Chan, J. C. L., and X. Liang, 2003: Convective asymmetries associated with tropical cyclone landfall. Part I: f plane simulations. J. Atmos. Sci., 60 , 1560–1576.
Chan, J. C. L., K. S. Liu, S. E. Ching, and E. S. T. Lai, 2004: Asymmetric distribution of convection associated with tropical cyclones making landfall along the South China coast. Mon. Wea. Rev., 132 , 2410–2420.
Dudhia, J., 1989: Numerical study of convection observed during the Winter Monsoon Experiment using a mesoscale two-dimensional model. J. Atmos. Sci., 46 , 3077–3107.
Dudhia, J., 1993: A nonhydrostatic version of the Penn State–NCAR Mesoscale Model: Validation tests and simulation of an Atlantic cyclone and cold front. Mon. Wea. Rev., 121 , 1493–1513.
Eliassen, A., and M. Lystad, 1977: The Ekman layer of a circular vortex: A numerical and theoretical study. Geophys. Norv., 31 , 1–16.
Gayno, G. A., N. L. Seaman, A. M. Lario, and D. R. Stauffer, 1994: Forecasting visibility using a 1.5-order closure boundary layer scheme in a 12-km non-hydrostatic model. Preprints, 10th Conf. on Numerical Weather Prediction, Portland, OR, Amer. Meteor. Soc., 18–20.
Janjić, Z. I., 1994: The step-mountain eta coordinate model: Further developments of the convection, viscous sublayer, and turbulence closure schemes. Mon. Wea. Rev., 122 , 927–945.
Jones, R. W., 1987: A simulation of hurricane landfall with a numerical model featuring latent heating by the resolvable scales. Mon. Wea. Rev., 115 , 2279–2297.
Kepert, J. D., 2001: The dynamics of boundary layer jets within the tropical cyclone core. Part I: Linear theory. J. Atmos. Sci., 58 , 2469–2484.
Kepert, J. D., 2002: The impact of landfall on tropical cyclone boundary layer winds. Extended Abstracts, 25th Conf. on Hurricanes and Tropical Meteorology, San Diego, CA, Amer. Meteor. Soc., 335–336.
Kepert, J. D., 2006a: Observed boundary layer wind structure and balance in the hurricane core. Part I: Hurricane Georges. J. Atmos. Sci., 63 , 2169–2193.
Kepert, J. D., 2006b: Observed boundary layer wind structure and balance in the hurricane core. Part II: Hurricane Mitch. J. Atmos. Sci., 63 , 2194–2211.
Kepert, J. D., and Y. Wang, 2001: The dynamics of boundary layer jets within the tropical cyclone core. Part II: Nonlinear enhancement. J. Atmos. Sci., 58 , 2485–2501.
Parrish, J. R., R. W. Burpee, F. D. Marks Jr., and R. Grebe, 1982: Rainfall patterns observed by digitized radar during the landfall of Hurricane Frederic (1979). Mon. Wea. Rev., 110 , 1933–1944.
Powell, M. D., 1982: The transition of the Hurricane Frederic boundary-layer wind field from the open Gulf of Mexico to landfall. Mon. Wea. Rev., 110 , 1912–1932.
Powell, M. D., 1987: Changes in the low-level kinematic and thermodynamic structure of Hurricane Alicia (1983) at landfall. Mon. Wea. Rev., 115 , 75–99.
Powell, M. D., and S. H. Houston, 1996: Hurricane Andrew’s landfall in South Florida. Part II: Surface wind fields and potential real-time applications. Wea. Forecasting, 11 , 329–349.
Powell, M. D., and S. H. Houston, 1998: Surface wind fields of 1995 Hurricanes Erin, Opal, Luis, Marilyn, and Roxanne at landfall. Mon. Wea. Rev., 126 , 1259–1273.
Powell, M. D., P. P. Dodge, and M. L. Black, 1991: The landfall of Hurricane Hugo in the Carolinas: Surface wind distribution. Wea. Forecasting, 6 , 379–399.
Shapiro, L. J., 1983: The asymmetric boundary layer flow under a translating hurricane. J. Atmos. Sci., 40 , 1984–1998.
Tuleya, R. E., and Y. Kurihara, 1978: A numerical simulation of the landfall of tropical cyclones. J. Atmos. Sci., 35 , 242–257.
Tuleya, R. E., M. A. Bender, and Y. Kurihara, 1984: A simulation study of the landfall of tropical cyclones. Mon. Wea. Rev., 112 , 124–136.
Wang, B., and X. Li, 1992: The beta drift of three-dimensional vortices: A numerical study. Mon. Wea. Rev., 120 , 579–593.
Willoughby, H. E., 1990: Gradient balance in tropical cyclones. J. Atmos. Sci., 47 , 265–274.
Willoughby, H. E., and P. G. Black, 1996: Hurricane Andrew in Florida: Dynamics of a disaster. Bull. Amer. Meteor. Soc., 77 , 543–549.
Wong, M. L. M., and J. C. L. Chan, 2004: Tropical cyclone intensity in vertical wind shear. J. Atmos. Sci., 61 , 1859–1876.
Wong, M. L. M., and J. C. L. Chan, 2006: Tropical cyclone motion in response to land surface friction. J. Atmos. Sci., 63 , 1324–1337.
Tangential (contour) and radial (shaded) surface winds for the (a) sea-only and (b) land-only experiments. The TC center is located at the origin. (Unit: m s−1)
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
As in Fig. 1 except for the experiments with a land–sea interface (see Table 1). The vertical dashed line indicates the coastline with land (sea) to its left (right).
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Surface tangential wind reduction factor for the (a) sea-only and the (b) LS experiments. The TC center is at the origin. Values larger than 1 are shaded and the vertical dashed line in (b) indicates the coastline with land (sea) to its left (right).
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Four streamlines in the LS case, at the lowest model level (σ = 0.995; 43 m). The streamlines represent airflow for a period of 2 h and the air is located at 150 km to the rear, right, front, and left of the TC center (TC symbol) at the middle of the period (marked by 
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
(a) Tangential and (b) inflow accelerations (i.e., dυ/dt and −du/dt) along the four streamlines in Fig. 4; (c) and (d) the respective accelerations due to vertical diffusion, which includes surface friction (unit is 10−3 m s−2). The value of 0 on the time axis corresponds to the middle of the 2-h period.
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Azimuthal-mean tangential, radial, and gradient winds for the (a) sea-only, (c) land-only, and (e) LS experiments at a radius of 50 km from the respective TC center. The lowest level shown is σ = 0.995 (43 m). (b), (d), (f) The results at a radius of 100 km from the respective TC center.
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Tangential (contour) and radial (shaded) winds at various levels above the surface for the LS case: (a) σ = 0.975, (b) σ = 0.94, (c) σ = 0.915, (d) σ = 0.885, (e) σ = 0.855, and (f) σ = 0.825. The vertical dashed line indicates the coastline, with land (sea) to its left (right).
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Combination of various terms in (2.1) at σ = 0.915 for the LS case: (a) Coriolis + centrifugal − pressure gradient, (b) tangential advection + vertical advection, (c) tangential advection, and (d) vertical advection. Unit is 10−3 m s−2. Positive values are shaded. The TC center is at the origin and the vertical dashed line represents the coastline with land (sea) to its left (right).
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Azimuthal location (relative to the direction pointing from sea to land) of the maximum wavenumber-1 component of the surface (large symbols) and PBL (small symbols) convergence in the 0–50-km (squares), 50–100-km (triangles), and 100–500-km (circles) radial bands. The results of a particular experiment is shown at the corresponding x position (e.g., L150 results along x = −150 km).
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
(a) Sea level pressure (hPa), (c) surface gradient wind speed (m s−1), and (e) surface radial (shaded) and tangential (contour) winds (m s−1) at t = 116 h in the RD experiment of WC06; (b), (d), (f) respective results at t = 122 h. The TC center is at the origin and the vertical dashed line represents the coastline with land (sea) to its left (right).
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
(a) Azimuthal location (relative to the direction pointing from sea to land) and (b) radial distance of the maximum surface tangential wind in the RD experiment of WC06. Each symbol represents the result at a particular hour of the 6-day (144 h) simulation. (c), (d) Respective results for the surface inflow; (e), (f) the surface gradient wind.
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
Ratio of the maximum surface tangential wind to the gradient wind in the RD experiment of WC06. The vertical dashed line marks the time of landfall at t = 122 h.
Citation: Journal of the Atmospheric Sciences 64, 9; 10.1175/JAS4027.1
List of numerical experiments. The labels S and L stand for sea and land, respectively.
In the RD experiment of WC06, the mass fields are free to change. The TC center, which is initially 150 km offshore, drifts forward left (relative to the onshore direction) toward the rough and dry land surface.
The difference between the maximum total wind and maximum tangential wind is found to be small.
