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Unformatted text preview: Atmospheric Moisture and Precipitation • • • • • • • • • Hydrosphere and the Hydrologic Cycle At mospheric Humidity Dewpoint Temperature Adiabatic Processes At mospheric Stability and Instability Clouds Precipitation Processes Thunderstorms Fog Formation Processes The Hydrosphere Fig. 7.3 • Hydrosphere: includes water on the Earth in all its forms – 97.2% is saltwater, 2.8% is freshwater – only 0.03% of all planetary water resides in the atmosphere 1 The Hydrologic Cycle • Hydrologic Cycle: movement of water among the great global water reservoirs (oceans, lakes, groundwater, rivers, atmosphere) • Water moves from the land and ocean to the atmosphere as water vapor and eventually returns to the world ocean as precipitation and runoff. The Hydrologic Cycle (Fig 9.1) 3 3 Annual average data in km x 10 . Think of the latent heat amounts! 2 Three States of Water Humidity • Humidity: the amount of water vapor present in the air • The maximum quantity of water vapour that can be held at any time in the air is directly dependent on air temperature. Thus warm air can hold more water vapor than can cold air. • When air at a given temperature is holding the maximum possible amount of water vapour we say it is saturated. • When air is unsaturated it is termed ‘dry’, but this does not mean that the air contains no water vapour. 3 Concept of Vapour Pressure • Air is a mechanical mixture of gases, one of which is water vapour. • Water vapour thus exerts a partial pressure on a surface according to how much is present in the atmosphere. This is the meaning of the term ‘Vapour Pressure’. 2 • Pressures are measured in Newtons/m = Pascals. 2 1N/m = 1Pa. To convert to millibars (mb), divide Pa by 100. Thus 100,000 Pa = 100kPa = 1,000 mb Saturation Vapour Pressure vs. Temperature Fig 7.11 Relat ion between saturation vapour pressure, e and air temperature, T . s a Note the nonlinear relation: e.g., at 10°C, e = 12 mb. At 30°C, e = 43 mb. s s supersaturated All po ints on the red line are saturated – i.e., at given temperature, water vapour pressure is at its maximum value. Points below the line are unsaturated. A small degree o f supersaturation is possible with very pure air and water (up to about 3% above the saturated value). unsaturated 4 Saturation Vapour Pressure at Below 0°C It is quite common for pure water to become supercooled. This mea ns that pure water can exist in the liquid state at well below 0°C, provided it is left undisturbed (supercooling down to as low as 40°C has been achieved in the laboratory!). What the graph shows is that the saturation vapour pressure over ice is lower tha n that over water, at the given temperature. This mea ns that water vapour will spontaneously migrate from a water droplet to an ice crystal and will be deposited as solid ice (reverse sublimation). Incidentally, this explains why refr igerators have to be periodically defrosted. Specific Humidity Specific humidity: the mass of water vapor held by a given mass of air (g/kg), also known as Mixing Ratio. It is also a measure of the quantity of water in the atmosphere that is available for precipitation. Fig 6.4 Note: Nonlinear relation 5 Global Temperature and Specific Humidity Note how much more mo isture tropical air contains relative to Arctic air. This explains in large part why the humid tropical areas are so much more mo ist than higher lat itudes. Relative Humidity Relative Humidity (RH): is the actual amount of water vapor present, e , as a percentage of the maximum amount that the air a can hold at that temperature, e : s RH = e / e x 100% a s
where e = actual vapour pressure and e = saturation vapour a s pressure –e.g. if the air holds half the moisture possible at the given temperature, then e = 0.5 e , and RH = 50% a s When RH is 100%, e = e , the air holds the maximum a s possible moisture level (saturation), and is at its dewpoint temperature. 6 Relative Humidity Varies with Air Temperature Since the sat. vapour pressure varies with temperature, then the ratio R.H. = ea /e must change with temperature, s even though the specific humidity is unchanged. Important point. Fig 7.10 Relative humidity provides a “comfort index”, or a measure of how humid the air ‘feels’ to us. If the air is close to saturation (R.H. » 100%), then litt le evaporation is possible and we feel discomfort because evaporative cooling is slow or nonexistent. Humidex is a measure of air temperature which includes the effect of R.H. Note inverse relat ion between R.H. and air temperature Methods for Measuring Humidity: Summary Four main methods for measuring humidity are: • Vapour pressure determination: the change in total gas pressure within a sealed container due to the addition of water vapour (done in the lab under controlled conditions). • Specific Humidity: Humidity = mass of w.v./ mass of air • Relative Humidity = e / e x 100% a s • Measure with a psychrometer (Gr. psychro = moisture, whereas Gr. psycho = mind, so be careful!). (Next slide). 7 Humidity Measurement with a Psychrometer Dry Wet • Drybulb thermo meter measures air temperature T . Wetbulb a thermo meter measures wetbulb temperature T . w T w • If air is unsaturated, water evaporates from wetbulb wick, causing latent heat of vaporization to be extracted from wetbulb, causing cooling. So T < Ta w • If air is saturated, T = T a w • Psychrometric table gives R.H. in terms of (T – T ). a w T a water DewPoint Temperature • DewPoint Temperature, T , is the temperature of an air d mass at which it is just saturated with water vapour. • If air is slowly cooled, but the specific humidity remains constant, the air will eventually reach saturation because its ability to hold moisture is decreasing. Eventually Td is reached and condensation of water begins. • Moist air has a higher dewpoint temperature than drier air, thus condensation will occur sooner in moist air. 8 DewPoint Temperature Determination • Assumpt ion is that during cooling of air to dew point, no water vapour is added or removed from the air. • Air ‘parcel’ A has known Ta = 27°C and known actual vapour pressure, e = 8 mb. a • Since e is constant during a cooling, slide X to the left horizontally until it cuts the red line at Y. Now read off T from d temperature scale on graph. Y ● ● X T d T a Relative Humidity Determination • Air ‘parcel’ A has known T = a 27°C and known actual vapour pressure, e = 8 mb. a • Since T is known, run vertically a up through point X until red line is intersected at Y. • Read off saturation vapour pressure e fro m vertical scale s on graph. • Determine RH = e /e x 100% a s T a Y ● ● X 9 Precipitation • Precipitation: particles of liquid or solid water that fall fro m the atmosphere and may reach the ground/ocean surface as rain, snow, or hail. Particles fall because updrafts are unable to sustain their weight. • Before precipitation particles can form, moist air must first condense to form water droplets (or ice particles) in a cloud. • Principal mechanism for forming clouds is adiabatic cooling. The Adiabatic Process • Adiabatic process: a change of temperature within a gas volume (‘air parcel’) that occurs solely as a result of a pressure change. • When air parcels are lifted, they expand and cool – this is adiabatic cooling. • When air parcels subside, they are compressed and are heated – this is adiabatic warming. • There is assumed to be no gain or loss of sensible heat fro m the parcel as expansion/compression occurs. 10 Schematic of Adiabatic Cooling and Heating Dry Adiabatic Lapse Rate (DALR) • Dry adiabatic lapse rate: rate at which a rising ‘dry’ air parcel is cooled by expansion when no condensation is occurring (i.e. air is unsaturated). In meteorology, ‘dry’ means unsaturated, not desiccated. • Dry adiabatic lapse rate has a value of 1ºC decrease for every 100 m of parcel rise: DALR = 1ºC/100m • When a dry air parcel descends, it warms at the same rate of 1ºC/100m of descent. (The text uses DAR, but the rest of the world uses DALR). 11 Saturated Adiabatic Lapse Rate (SALR) • Saturated adiabatic lapse rate: rate at which a rising air parcel is cooled by expansion when condensation is occurring (i.e. air is saturated and latent heat is being released) o SALR ranges between 0.4 0.9 C/100 m and averages about 0.6°C/100 m of air parcel ascent. • An air parcel following the saturated adiabatic lapse rate will produce a cloud that begins to form at the lifting condensation level the level at which a rising air parcel is cooled to its dewpoint temperature and condensation begins. (Text uses MAR, but everybody uses SALR) DALR and SALR on a TemperatureHeight Diagram Adiabatic Cooling Note: ignore dewpoint lapse rate effect. 12 Environmental Lapse Rate (ELR) • Environmental lapse rate: rate of temperature change (increase or decrease) with height within the lower o atmosphere. Has an average value of about 0.6 C/100 m or 6° C/km, but it is very variable in both time and space. • Environmental lapse rate, ELR, determines whether or not an air parcel will spontaneously rise or sink close to the Earth’s surface, fro m basic buoyancy principles. The ELR is controlled by the ther mal history of an air mass, such as where it originated, and over what kinds of surfaces (e.g. cold land, warm water) it has travelled. Stable and Unstable Air • Stable: air parcel is cooler (thus denser) than the surrounding air no tendency to rise. For this to occur, DALR > ELR • Unstable: air parcel is warmer than the surrounding environment will tend to rise. For this to occur, DALR < ELR 13 Convection in Unstable Air: DALR>ELR Rising air parcel is warmer than environment at all levels, so it keeps on rising. Latent heat release in cloud Absolute Stability and Conditional Instability Absolute stability occurs when both the SALR and DALR are greater than the ELR. The air has no tendency to rise. All we need say is ELR<SALR. Conditional instability occurs when the ELR lies between DALR and SALR. The air is stable if ‘dry’ but unstable if saturated. That is: SALR<ELR<DALR. (Figs. 7.18c and 7.18b) 14 Clouds • A cloud is made up of water droplets or ice particles suspended in air (similar to fog) – clouds also may contain tiny “supercooled” water droplets which exist below freezing in a liquid state • Each cloud particle is formed on a tiny center of solid matter called a condensation nucleus – Two important sources of condensation nuclei are: salt fro m ocean spray and dust blown from the land surfaces Cloud Types • Clouds are classified into four fa milies according to the height above the surface where they occur: 1. 2. 3. 4. High middle low tall clouds with ‘vertical development’ • Clouds are also grouped into two major classes on the basis of form: 1. Stratiform (layer) clouds (e.g. cirrus, stratus, altostratus) • blanketlike layers and cover large areas 2. Cumuliform (‘lumpy’) clouds (e.g. cumulus, altocumulus, cumulo nimbus) • globular masses of varying height 15 Cloud Types Cont. Cloud Forms Cont. altocumulus cumulus with cirrus above high cirrus lenticular clouds 16 Methods of Vertical Air Movement Four main mechanisms for lift ing and adiabat ically cooling air: 1. Convergent lift ing: air streams mo ve towards each other forcing air to rise (e.g., at the ITCZ). 2. Convectional lift ing: the air is unstable and there is strong heat ing of the ground surface. Heated air parcels are less dense than the surrounding air and so rise, cool, and condense. 3. Orographic lift ing: results when a moist airmass is forced to rise up and over a mountain barr ier. Causes increased precipitation on windward slopes (known as ‘orographic enhancement’). 4. Frontal lift ing: results when a mass o f cooler, denser air slides under a mass of warmer, less dense air, and the warmer air is lifted as a layer and condenses. Fig 8.6 17 Precipitation: Raindrop Formation Processes Raindrops can form in two ways: A. In warm clouds, very small cloud water droplets, about 2μm in diameter, collide with one another and eventually coalesce to form raindrops 2,000 – 4,000μm (24 mm) in size that then fall as rain. This is the Collis io nCoalescence process. B. In colder clouds, which extend above the freezing level, a mixture of both ice crystals and supercooled water droplets typically exists. The latter instantaneously freeze against a colliding particle, so accretion can be rapid. Subsequent ly the particles fall below the freezing level, melt, and fall as rain. This is the Bergeron Process for raindrop formation, named after the Norwegian scient ist Thor Bergeron. (Sleet is simply ice that has not completely melted before hitting the ground). Precipitation: Snowflake Formation Snowflakes represent the growth of complexshaped ice crystals in clouds which extend above the freezing level. Freezing begins around a freezing nucleus, then progresses by reverse sublimat ion of water vapour onto the ice as well as by vapour transport fro m supercooled droplets (recall that ice has a lower saturation vapour pressure than water at the same temperature, so the ice phase will always gain mass at theexpense of the supercooled droplets). Fully developed snowflakes develop crystalline, branched structures (dendrites, stellar dendrites) displaying an overall hexagonal symmetry. Collisio n of supercooled droplets with snowflakes causes immediate freezing as amorphous (noncrystalline) ice known as rime. If the air below the cloud base is < 0°C then snow reaches the ground. If not, then it falls as eit her sleet or rather cold raindrops! 18 Photographs by Wilson Bentley, 1902 Free Convection and Precipitation • Free convection occurs when moist nearsurface air is either absolutely unstable (DALR<ELR) or conditionally unstable (SALR<ELR<DALR). Buoyant moist air rising in a convection cell is cooled adiabatically, leading to cumulus (or cumulonimbus) cloud formation and eventually precipitation, if the cloud is tall enough. – Convection cell: vertical motion of rising air found above warmed land surfaces 19 Convectional Precipitat ion From Cumuliform Clouds in Unstable Air Fig 7.21d cumulus no precipitation light showers heavy showers or thundershowers plus hail Thunderstorms: extreme convection phenomena • Thunderstorm: intense local storm associated with mo ist, unstable air and tall, dense cumulo nimbus clouds in which there are very strong updrafts of air. Heavy rain or hail are common precipitation forms. Two environmental condit ions encourage a thunderstorm: 1. Warm and mo ist air at the surface. 2. A very deep unstable layer, 810 km thick, which allows tall cumulo nimbus clouds to extend well above the freezing level. These are complex, mixed clouds which contain water droplets at low levels, and supercooled droplets and ice crystals above the freezing level. • 20 Hail • Hail is formed by accumulation of ice layers on ice pellets that are suspended wit hin the updrafts of a cumulo nimbus cloud. The stronger the updrafts, the larger the hailstones can grow. • An annular, or ringlike, concentric structure is o ften observed in hailstones, which can grow to the size of baseballs ! Killers! • Hailstones grow by colliding with supercooled droplets as well as by the freezing o f accreted liquid water. Opaque ice represents rapid freezing of supercooled droplets. Clear ice represents slower freezing o f liquid water. t i m e opaque ice clear ice Orographic Precipitation on Windward Slope of Mountains Air is forced to rise as it strikes the windward slope of a mountain barrier. Lift ing of air mass causes adiabatic cooling and the format ion of extensive cloud layers. Precipitat ion is heavy on the windward slopes (e.g. North Shore Mountains at Vancouver). On leeward slopes, subsidence of air promotes adiabatic warming, causing evaporation of clouds and thus an area of reduced precipitation (‘rain shadow area’), e.g. Okanagan Valley. 21 Note the strong relation bet ween precipitation and elevation. mm < 300 300 – 400 400 – 500 500 – 750 750 – 1000 1000 – 1500 1500 – 2500 2500 – 3500 > 3500 At las o f Brit ish Co lumbia Formation of Fog Fog is caused by the cooling of mo ist air to the dewpoint temperature at or near the Earth’s surface. Since mo ist air has a higher dew point than drier air, fog tends to form preferent ially close to major water bodies such as oceans and large lakes. Wet land areas such as deltas and marshes also experience higher frequencies of fogs than do drier upland areas. This is certainly the experience in the Greater Vancouver area. There are two main types of fog: • Radiation Fog: this forms when mo ist air is chilled to the dew point due to nocturnal radiative chilling o f air under calm, clear sky condit io ns. • Advection Fog: this forms when warm, mo ist air mo ves horizontally (advects) across a colder surface and is chilled to the dewpoint as it loses sensible heat to the colder surface. 22 Radiation Fog This is the type of fog which persisted for mor e than two weeks at Vancouver in January 2009. The r equir ements are as follows: • A prolonged r esidence of a highpressure cell (anticyclone), which is nor mally associated with clear skies and calm conditions. • Moist air which is fairly close to its dewpoint temperature. • Negative longwave budget which causes cooling of nearsurface air to the dewpoint. • Fog may persist for ma ny days because incoming solar energy is reflected off the top of the fog layer, so cannot heat the ground. • Fogs are usually dispersed by the arrival of a lowpressure system which brings winds to mix the saturated air upwards away from the gr ound. Radiation Fog under Te mperature Inversion Conditions Clear skies L¯ L- Fog layer Fog occurs within the inversio n layer due to radiat ive chilling of nearsurface air to the dewpoint temperature. 23 Radiat ion fog over English Bay: view fro m UBeastly, Nov 2007 Advection Fog (e.g., San Francisco, Halifax) Fog gradually evaporates as it mixes with warmer, drier air inland Air is chilled to the dewpoint as it crosses the cold water surface Advec tion warm, moist air warm water F o g Such fogs persist in especially in summer (why summer?) along the west coast from southern Vancouver Island to Central California, and are especially persistent in Oregon and from San Francisco as far south as Santa Barbara. cold water upwelling 24 Advection Fog at San Francisco Note how the fog evaporates as it mixes with warmer, drier air inland from the Pacific coast. Note how the warmer (lower albedo) downtown core is fog free. Oakland Downtown San Francisco Pac ific co ast
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This note was uploaded on 11/03/2009 for the course ECON 210 taught by Professor James during the Spring '09 term at The University of British Columbia.
- Spring '09