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1. Introduction Understanding the winds within a valley and their interactions with the larger-scale forcings is of interest for a number of reasons. For example, the dispersion of pollutants in a valley depends strongly on local valley circulations (e.g., Whiteman 1989 ; Fast et al. 2006 ); nocturnal minimum surface temperatures depend strongly on the near-surface wind speed (e.g., Estournel and Guedalia 1985 ; Steeneveld et al. 2006 ) and hence on the strength of the valley wind; land
1. Introduction Understanding the winds within a valley and their interactions with the larger-scale forcings is of interest for a number of reasons. For example, the dispersion of pollutants in a valley depends strongly on local valley circulations (e.g., Whiteman 1989 ; Fast et al. 2006 ); nocturnal minimum surface temperatures depend strongly on the near-surface wind speed (e.g., Estournel and Guedalia 1985 ; Steeneveld et al. 2006 ) and hence on the strength of the valley wind; land
the MAPR site. 3. Case studies A pressure gradient forces air to flow over or around a mountain range. We discern two major factors, namely the synoptic-scale pressure field, hence called the dynamic forcing because it causes the synoptic wind field, and the contribution from differences in temperature between the upstream and downstream side of the barrier, hence called the hydrostatic forcing. To distinguish the magnitude of the two forcings for the four cases, we compare the vertical structure
the MAPR site. 3. Case studies A pressure gradient forces air to flow over or around a mountain range. We discern two major factors, namely the synoptic-scale pressure field, hence called the dynamic forcing because it causes the synoptic wind field, and the contribution from differences in temperature between the upstream and downstream side of the barrier, hence called the hydrostatic forcing. To distinguish the magnitude of the two forcings for the four cases, we compare the vertical structure
. 2008b ). Nevertheless, it is a reasonable expectation that mountain-wave activity generated by westerly flow over the Sierra Nevada can lead to various degrees of penetration of momentum to the valley floor on the lee side ( Jiang and Doyle 2008 ; Strauss et al. 2016 ). Dynamical and thermal forcing mechanisms are known to cause downslope winds on the lower mountain slopes (or the valley bottom or leeside plains) at many locations around the world. The connection between large-amplitude mountain
. 2008b ). Nevertheless, it is a reasonable expectation that mountain-wave activity generated by westerly flow over the Sierra Nevada can lead to various degrees of penetration of momentum to the valley floor on the lee side ( Jiang and Doyle 2008 ; Strauss et al. 2016 ). Dynamical and thermal forcing mechanisms are known to cause downslope winds on the lower mountain slopes (or the valley bottom or leeside plains) at many locations around the world. The connection between large-amplitude mountain
valley atmosphere at the same elevation. Furthermore, it has a nonzero initial momentum ( u s ~ 3 m s −1 , see Fig. 9a ) as a result of the synoptic forcing. As the parcel plunges down-slope, it is accelerated by greater buoyancy forcing as a result of its initial temperature deficit. When it reaches the elevation of its neutral buoyancy, the excess momentum gained from the downslope acceleration, together with its initial momentum, allows the parcel to continue its down-slope motion into a colder
valley atmosphere at the same elevation. Furthermore, it has a nonzero initial momentum ( u s ~ 3 m s −1 , see Fig. 9a ) as a result of the synoptic forcing. As the parcel plunges down-slope, it is accelerated by greater buoyancy forcing as a result of its initial temperature deficit. When it reaches the elevation of its neutral buoyancy, the excess momentum gained from the downslope acceleration, together with its initial momentum, allows the parcel to continue its down-slope motion into a colder
slopes and neighboring plateaus, most of which assume “undisturbed” conditions; that is, the large-scale flow is relatively weak (e.g., Gleeson 1951 ; Whiteman 1990 ; Rampanelli et al. 2004 ), with a few exceptions. For example, the relationship between valley winds and the synoptic-scale flow has been examined by Whiteman and Doran (1993) using a numerical model, and mechanisms including thermal forcing, downward mixing of momentum, valley channeling, and pressure-driven channeling have been
slopes and neighboring plateaus, most of which assume “undisturbed” conditions; that is, the large-scale flow is relatively weak (e.g., Gleeson 1951 ; Whiteman 1990 ; Rampanelli et al. 2004 ), with a few exceptions. For example, the relationship between valley winds and the synoptic-scale flow has been examined by Whiteman and Doran (1993) using a numerical model, and mechanisms including thermal forcing, downward mixing of momentum, valley channeling, and pressure-driven channeling have been
thermal forcing. As expected from the theory of thermally driven circulations, daytime winds generally blow up valley and upslope while nighttime winds generally blow down valley and downslope. Figure 3 shows daytime and nighttime wind roses at Keeler and Lone Pine using all available data from their individual periods of record. The daytime winds at Keeler are mostly from the southerly quadrants indicating up-valley and upslope flows. At night, wind directions are generally down valley (from the
thermal forcing. As expected from the theory of thermally driven circulations, daytime winds generally blow up valley and upslope while nighttime winds generally blow down valley and downslope. Figure 3 shows daytime and nighttime wind roses at Keeler and Lone Pine using all available data from their individual periods of record. The daytime winds at Keeler are mostly from the southerly quadrants indicating up-valley and upslope flows. At night, wind directions are generally down valley (from the
“still air” rise–fall rate. The latter can be computed based on fluid dynamics and characteristics of radiosonde and dropsonde systems (see below). The VV is estimated by subtracting the still-air rise–fall rate from that in the actual atmosphere. The radiosonde rise rate in the still air is calculated based on the balance of the buoyancy and drag force exerted on the sonde (e.g., Johansson and Bergström 2005 ). The buoyancy force (BF) is the difference between the free lifting and the total weight
“still air” rise–fall rate. The latter can be computed based on fluid dynamics and characteristics of radiosonde and dropsonde systems (see below). The VV is estimated by subtracting the still-air rise–fall rate from that in the actual atmosphere. The radiosonde rise rate in the still air is calculated based on the balance of the buoyancy and drag force exerted on the sonde (e.g., Johansson and Bergström 2005 ). The buoyancy force (BF) is the difference between the free lifting and the total weight
between the minimum and maximum vertical grid spacing was given by where Δ z min = 20 m, Δ z m = 110 m, a = (1 + n )/2, α = 0.5, and n = 20. The lateral boundary conditions are periodic. A Rayleigh sponge layer, starting at 5 km, was specified as the top boundary condition. All simulations were run with the Coriolis force turned off. The models were integrated for 12 h from sunrise at 0600 local time (LT) to sunset at 1800 LT. The temporal evolution of surface sensible heat flux is determined
between the minimum and maximum vertical grid spacing was given by where Δ z min = 20 m, Δ z m = 110 m, a = (1 + n )/2, α = 0.5, and n = 20. The lateral boundary conditions are periodic. A Rayleigh sponge layer, starting at 5 km, was specified as the top boundary condition. All simulations were run with the Coriolis force turned off. The models were integrated for 12 h from sunrise at 0600 local time (LT) to sunset at 1800 LT. The temporal evolution of surface sensible heat flux is determined
underlying terrain. For example, mesoscale predictions of landfalling fronts were found to be very sensitive to small changes in incident flow, as deduced through simulations made with small modifications to the topography orientation by Nuss and Miller (2001) . Two-dimensional idealized adjoint ( Doyle et al. 2007 ) and ensemble ( Doyle and Reynolds 2008 ) model results indicate large sensitivity to the initial state as the mountain height increases, forcing wave breaking, where perturbation growth
underlying terrain. For example, mesoscale predictions of landfalling fronts were found to be very sensitive to small changes in incident flow, as deduced through simulations made with small modifications to the topography orientation by Nuss and Miller (2001) . Two-dimensional idealized adjoint ( Doyle et al. 2007 ) and ensemble ( Doyle and Reynolds 2008 ) model results indicate large sensitivity to the initial state as the mountain height increases, forcing wave breaking, where perturbation growth
, with the fifth- and sixth-order schemes yielding values, 1% and 3% stronger than the corresponding continuous solution. (Despite the implicit numerical diffusion associated with the fifth-order scheme, the wave amplitude is slightly stronger than in the continuous solution.) We now consider the impact of discretization on hydrostatic mountain waves. Figure 4a shows the normalized vertical velocity in the continuous Boussinesq solution for δ = 10. The influence of the Coriolis force is included
, with the fifth- and sixth-order schemes yielding values, 1% and 3% stronger than the corresponding continuous solution. (Despite the implicit numerical diffusion associated with the fifth-order scheme, the wave amplitude is slightly stronger than in the continuous solution.) We now consider the impact of discretization on hydrostatic mountain waves. Figure 4a shows the normalized vertical velocity in the continuous Boussinesq solution for δ = 10. The influence of the Coriolis force is included
