The Influence of Mountains on Airflow, Clouds, and Precipitation

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<ul><li><p>The Influence of Mountains on Airflow, Clouds, and Precipitation </p></li><li><p>Severe Downslope Windstorms </p></li><li><p>Figure 11.6. West to east vertical cross section of potential temperature across the Sierra Nevada. Dashed line represents sailplane soundings. Observed Chinook arch or Foehn wall cloud is illustrated over barrier crest, as well as rotor cloud at low levels to the east and lenticular cloud at higher levels. [From Holmboe and Klieforth (1957).] </p></li><li><p>Figure 11.9. Contours of horizontal velocity (m s-1) along an east-west line through Boulder, Colorado, as derived from the NCAR Sabreliner data on 11 January 1972. The analysis below 500 mbar was partially obtained from vertical integration of the continuity equation, assuming two-dimensional steady-state flow. [From Klemp and Lilly (1975).] </p></li><li><p>. Over the mountain crest is a rather typical flow pattern over a broad mountain. A stationary orographic cloud exists over the highest peaks. Directly to the lee of the higher peaks, the flow descends abruptly to the plains elevation. Evidence of trapped lee waves, including a lenticular cloud, can be seen over the plains. At higher levels is a deep trough in which air originating near stratospheric heights descends to below 500 mbar. This very high-amplitude wave is believed to be instrumental in causing surface winds in excess of 50m/s. </p><p>Maximum wind speeds occur along the lee slope of the mountain barrier at low levels.</p></li><li><p>Figure 11.11. Total horizontal velocity field for the Boulder, Colorado, windstorm simulation. Times are (a) 3200, (b) 4160, (c) 5120, (d) 6020, (e) 7040, and (f) 8000 sec. Contour interval is 8 m s-1. In (f) the horizontal wind maximum in the lee of the peak is in excess of 60 m s-1. [From Peltier and Clark (1979).] </p></li><li><p>Figure 11.12. Total horizontal velocity field for the Boulder, Colorado, windstorm simulation. Times are (a) 3200, (b) 4160, (c) 5120, (d) 6020, (e) 7040, and (f) 8000 sec. Contour interval is 8 m s-1. In (f) the horizontal wind maximum in the lee of the peak is in excess of 60 m s-1. [From Peltier and Clark (1979).] </p></li><li><p>Klemp and Lilly (1975) first explained the severe downslope wind phenomena with a two-dimensional, linearized, hydrostatic model in isentropic coordinates. They concluded that the mechanism leading to strong amplification of the wave is associated with the partial reflection of upward-propagating wave energy by variations in thermal stability. They argued that a strong wave response occurs whenever the mean vertical wavelength is such that an integral number of half-wavelengths can be confined between the ground and the tropopause.</p></li><li><p>Peltier and Clark found that the wave actually broke, leading to a local wind reversal and a layer of constant potential temperature. As a consequence, wave energy reflected from the earth's surface became trapped between the resultant critical layer and the ground. This reflection cavity produced the large-amplitude streamline deflections that resulted in the strong surface winds. When the stratospheric wind profile was modified to prevent wave breaking, the final phase of wave amplification did not occur, and the results were then similar to linear theory. </p></li><li><p>Effects of moisture on less waves</p></li><li><p>Cloud processes can influence mountain-wave flow and thereby feed back on the formation of precipitation in orographic clouds. </p></li><li><p>Trapping of g-wavesRecall that when l**2 is less than k**2, where l**2 is the Scorer parameter given by </p><p> Moisture alters N**2</p></li><li><p>Figure 11.14. Absolutely stable atmosphere favorable for the development of dry lee waves. (a) Temperature and wind speed profiles; dry adiabats are marked with a short-dash line; moist pseudoadiabats are marked with a Iong-dash line. (b) Scorer parameter l2 profiles; the dry l2 is marked with a solid line, the equivalent saturated l2 is a dashed line. [From Durran and Kemp (1982b).] </p></li><li><p>Figure 11.15. Streamlines produced by a 300-m-high mountain in the flow for relative humidity (RH): (a) RH =0%, (b) RH =90% (c) RH = 100%, and (d) RH = 100% with 0.2 g kg-1 of cloud, in the lowest layer upstream. Cloudy regions are shaded. [From Durran and Klemp, 1982b.] </p></li><li><p>The inclusion of the effects of clouds results in less stable flow because the buoyancy-restoring force is decreased, the amplitude of the mountain wave under certain conditions can be significantly weakened. </p></li><li><p>Figure 11.16. Streamlines produced by a 300-m-high mountain in the flow. (a) Steady solution for RH = 0%. Time-dependent flow for RH = 90% in the lowest upstream layer at (b) t = 8000 s, (c) t = 12,000 s, and (d) t = 16,000 s. Cloud regions are shaded; dark shading indicates cloud densities exceeding 0.3 g kg-1. [From Durran and Klemp, 1982b.] </p></li><li><p>In the dry atmosphere, a distinct trapped lee wave is evident in their solutions. The addition of a layer with 90% relative humidity results in the formationof clouds over the mountain crest and in the regions of upward motion of the trapped lee waves. The wave structure is modified somewhat. As a result of adding a 100% saturated layer, the flow is modified so that the wavelength of the partially trapped waves is increased significantly.Finally, by adding 0.2g/kg of liquid water to the saturated layer, a cloud could be maintained in the wave troughs as well as the wave crests. This so altered the resultant vertical profile of the Scorer parameter that the lee waves became untrapped.</p></li><li><p>For 11 January 1972 severe downslope windstorm over Boulder, Colorado. Durran and Klemp noted the addition of moisture to their model decreased the downslope wind speed from 45 to 25m/s, a result of a weakened mountain-wave amplitude. Durran and Klemp also noted that lee-side warming, referred to as the Chinook or Alpine foehn, is often attributed to the release of latent heat on the windward side of the barrier in precipitating clouds and to dry adiabatic descent on the lee side. In a precipitating cloud simulation, they noted that the lee-side temperatures were several degrees cooler than those in nonprecipitating flow. This suggests that the most important factor influencing Chinook or foehn wind lee-side temperatures is the amplitude of the mountain wave, which is larger in the dry case.</p></li><li><p>Lilly and Durran (1983) extended the Durran-Klemp calculations to precipitating clouds as well. Using a simple Kessler-type warm rain parameterization, they investigated the effects of cloud processes, including precipitation, on vertical momentum fluxes over orographic barriers. The calculated vertical momentum fluxes for (a) a case having low clouds and (b) a case saturated everywhere. The fluxes are normalized to fluxes expected for linear mountain wave theory </p></li><li><p>Figure 11.17. The effects of rain on the vertical profiles of momentum flux produced by upstream moisture profiles in which (a) there are low clouds between the heights of 667 and 3000 m, and (b) RH = 100 % everywhere. The fluxes are normalized by MLC, the flux associated with linear mountain waves. [From Lilly and Durran (1983).] </p></li><li><p>The seeder-feeder process</p></li><li><p>Figure 11.21. Conceptual model illustrating the orographic enhancement of rain. [From Browning's (1979) adaptation of Bergeron's (1965) figure.] </p></li><li><p>Blocking of low-level flowFroude Number: Fr=U/(N*h), where U is the speed of the incoming flow, N is the Brunt-Vaisala frequency, and h is the height of the mountain. The Froude number(Fr) represents the ratio of the square root of the kinetic energy of the horizontal flow impinging on a mountain barrier to the energy required to lift an air parcel from the base of a mountain to its top in a stably stratified environment. Therefore blocking is more likely to occur when winds are weak or stabililty is large. </p></li><li><p>Figure 11.23. Conceptual model for the working hypothesis that low-level decoupled flow (stippled area) acts as an extension of the mountain barrier for orographic lift purposes which would then alter the location of condensate production and hence precipitation. A small amount of decoupled low-level flow (a) allows parcel lift to occur near the barrier while a large amount of low-level decoupled flow (b) forces parcel lift to occur upstream of the barrier. [From Peterson et al., 1991.] </p></li><li><p>Figure 11.24. A schematic depiction of the position of a cold front, at 2-h intervals, as it approaches and is influenced by a mountain range. The distortion of the frontal surface is from slowing of the low-level flow by the mountain and the acceleration aloft. This differential advection causes the cold air behind the front to override the warm air, producing an unstable air column. The resulting small-scale convection enhances precipitation upstream of the mountain and on its windward slopes. This diagram is constructed for u0 =10 m s-1, N = 0.01 s-1, b = 20 km, h = 800 m, x0 = -100 km, and a = 1/50. The vertical exaggeration is 12:1. [From Smith (1982).] </p></li><li><p>Note that blocking of low-level flow on the windward side of the barrier causes differential thermal advection and makes the flow on the lee-side unstable as cooler air rides over relatively warmer air resulting small-scale convection</p></li><li><p>Figure 11.25. Schematic portrayal of a split front with the warm conveyor belt undergoing forward-sloping ascent, but drawing attention to the split-front characteristic and the overall precipitation distribution: (a) plan view, (b) vertical section along AB in (a). In (a) UU represents the upper cold front. The hatched shading along UU and ahead of the warm front represents precipitation associated with the upper cold front and warm front, respectively. Numbers in (b) represent precipitation type as follows: (1) warm-frontal precipitation; (2) convective precipitation-generating cells associated with the upper cold front;(3) precipitation for the upper cold-frontal convection descending through an area of warm advection; (4) shallow moist zone between the upper and surface cold fronts characterized by warm advection and scattered outbreaks of mainly light rain and drizzle;(5) shallow precipitation at the surface cold front itself. [Adapted from Browning and Monk (1982), Browning (1985), and Reynolds and Dennis (1986). Reproduced with the permission of the Controller of Her Britannic Majesty's Stationery Office.] </p></li><li><p>Precipitation Efficiency in orographic cloudsDirks (1973) employed aircraft data to measure directly precipitation efficiencies over a relatively isolated mountain(Elk Mountain) in Wyoming. Aircraft soundings were used to calculate condensate production and precipitation rates. Precipitation efficiencies ranged from 25 to 80%. The lowest efficiencies were at very cold temperatures and strong winds, and the highest efficiencies were under moderately cold cloud top temperatures and moderate wind speeds.</p></li><li><p>Evidence suggests that a significant contributor to the water content of the snowpack arises from rime deposits on trees and other objects located on mountain peaks; this deposit is then shed to the snow surface. Hindman(1982), for example, estimated that between 4 and 11% of the water content of snowpack near the peak of the Park Range, Colorado, was the result of rime deposits.</p></li><li><p>Smith et al.(2003) proposed an alternate strategy for estimating precipitation efficiencies. They propose using the drying ratio (DR), DR=P/I= area integrated total precipitation/impinging horizontal vapor flux</p></li><li><p>Kirshbaum and Smith(2009) proposed an alternate expression for DR in the formDR=CR X PE , where CR is the condensation ratio (C/I) where C is the area integrated condensation rate.</p></li><li><p>Because PE requires an estimate of vertical air velocity at cloud base, it is subject to rather large measurement errors. Since DR uses the ratio of precipitation to water vapour flux, it is easier to quantify than PE. An important advantage of DR is that it can be estimated using hydrogen and oxygen isotope analysis, which permits evaluation using streamflow or sapwood collected near a stream. </p></li><li><p>The idea is that condensation preferentially removes the heavier isotopes so that the remaining vapor becomes progressively lighter such that the ratios of heavy-to-light isotope concentrations in the final and initial state of vapor, is the fraction of water vapor remaining. The drying ratio can thus be computed using the ratios of heavy to light isotope concentrations in the upwind and downwind precipitation. </p></li><li><p>DR estimates using isotope analysis range from 43% for the combination of the coastal and Cascade mountain ranges in western Oregon and 43% for a case in the of flow against the Alps(Smith et al., 2003) and nearly 50% for the southern Andes (Smith and Evans 2006).</p></li><li><p>Aerosol Influences on orographic precipitationThere is observational evidence(Borys et al.,2000;Borys et al.,2003) that pollution can delay precipitation in winter orographic clouds in the Rocky Mountains. They show that pollution increases the concentration of CCN and therefore cloud drops, leading to the formation of smaller cloud drops and less efficient riming. Reduced riming results in smaller, more pristine ice crystals, with smaller fall velocities, and less snowfall.</p></li><li><p>Figure 11.36. Light riming of ice crystals in clouds affected bypollution (left) compared to heavier riming in non-polluted clouds(right) (Borys et al. 2003). </p></li><li><p>Effect of pollution on snow fall from orographic cloudsBorys et al, GRL, 2003Polluted case, smaller drops, less rimingClean case,larger drops,more riming</p></li><li><p>Table 1. Chemical and Physical Properties of Cloud Droplets and Snow During Two Precipitation Events </p><p>February 15 (Polluted) 19 (clean)</p><p>Major Habit Planar Dendrite Planar Dendrite Rime Category Unrimed (0.5) Moderate (2.0) Rime Mass Frac. 5% 51% SPL Precip. Rate 0.02 mm hr1 0.38 mm hr1 ISS Precip. Rate 0 to 0.1 mm hr1 1.1 mm hr1 SPL Temperature 13C 4C Snow d18O 22.1 16.5 Cloud d18O 21.1 16.2 d18O Snow Mass 14C 4.8C Temp. Of Origin </p><p>Cloud Top Temp 19C 22C Snow CAE SO4 = 0.011 mgm 3 0.072 mgm 3 Cloud CAE SO4 = 1.1 mgm 3 0.12 mgm 3 Droplet Mean Dia. 8.3 mm 13.6 mm Droplet Conc. 310 cm 3 74 cm 3 Cloud SCLW 0.13 g m 3 0.14 g m 3 </p><p>Borys et al, GRL, 2003</p></li><li><p>Orographic enhancement factor (Ro)Orographic enhancement factor (Ro): ratio between precipitation amounts over hills to precipitation amounts in upwind lowland (Givati and Rosenfeld 2004, 2005)Suppression rate downwind of coastal urban areas in California and Israel 15-25% of annual precipitation in hillsOccurred mainly in relatively shallow orographic clouds within cold air mass of cyclonesRole of pollution aerosols in decreasing Ro over mountains to E of Salt Lake City (Griffith et al. 2005) but upwind increased!Similar decreasing trend...</p></li></ul>