The sun's rays, when passing through transparent substances, heat them very weakly. This is explained by the fact that direct sunlight produces virtually no heat. atmospheric air, but they strongly heat the earth’s surface, capable of transmitting thermal energy adjacent layers of air. As the air heats up, it becomes lighter and rises higher. In the upper layers, warm air mixes with cold air, giving it part of the thermal energy.
The higher the heated air rises, the more it cools. The air temperature at an altitude of 10 km is constant and amounts to -40-45 °C.
A characteristic feature of the Earth's atmosphere is a decrease in air temperature with height. Sometimes there is an increase in temperature as altitude increases. The name of this phenomenon is temperature inversion (temperature rearrangement).
Temperature change
The appearance of inversions may be due to cooling earth's surface and the adjacent layer of air in a short period of time. This is also possible when dense cold air moves from mountain slopes to valleys. During the day, the air temperature continuously changes. During the daytime, the earth's surface heats up and heats the lower layer of air. At night, along with the cooling of the earth, the air cools. It is coolest at dawn and warmest in the afternoon.
IN equatorial belt There is no daily temperature fluctuation. Night and day temperatures have the same values. Daily amplitudes on the coasts of seas, oceans and above their surface are insignificant. But in the desert zone, the difference between night and day temperatures can reach 50-60 °C.
In the temperate zone, the maximum amount of solar radiation on Earth occurs on the days of the summer solstices. But the hottest month is July in the Northern Hemisphere and January in the Southern. This is explained by the fact that despite the fact that solar radiation is less intense during these months, a huge amount of thermal energy is given off by the highly heated earth's surface.
The annual temperature range is determined by the latitude of a particular area. For example, at the equator it is constant and amounts to 22-23 °C. The highest annual amplitudes are observed in areas of mid-latitudes and in the interior of continents.
Any area is also characterized by absolute and average temperatures. Absolute temperatures are determined through long-term observations at weather stations. The hottest area on Earth is the Libyan Desert (+58 °C), and the coldest is the Vostok station in Antarctica (-89.2 °C).
Average temperatures are established by calculating the arithmetic mean values of several thermometer indicators. This is how average daily, average monthly and average annual temperatures are determined.
In order to find out how heat is distributed on Earth, temperature values are plotted on a map and points with the same values are connected. The resulting lines are called isotherms. This method allows us to identify certain patterns in the temperature distribution. So, most high temperatures are recorded not at the equator, but in tropical and subtropical deserts. Temperatures decrease from the tropics to the poles in the two hemispheres. Taking into account the fact that in Southern Hemisphere reservoirs occupy a larger area than land, temperature amplitudes between the hottest and coldest months there are less pronounced than in the North.
Based on the location of the isotherms, seven thermal zones are distinguished: 1 hot, 2 moderate, 2 cold, 2 permafrost areas.
Related materials:
inversion
air temperature increases with altitude instead of the usual decrease
Alternative descriptions
An excited state of a substance in which the number of particles is at a higher energy. level exceeds the number of particles at a lower level (physics)
Changing direction magnetic field Earth reversed, observed at time intervals from 500 thousand years to 50 million years
Changing the normal position of elements, placing them in reverse order
Linguistic term meaning a change in the usual word order of a sentence
Reverse order, reverse order
Logical operation "not"
Chromosomal rearrangement associated with a rotation of individual chromosome sections by 180
Conformal transformation of the Euclidean plane or space
Rearrangement in mathematics
Dramatic device demonstrating the outcome of the conflict at the beginning of the play
In metrology, an anomalous change in a parameter
A state of matter in which higher energy levels of its constituent particles are more "populated" by particles than lower ones
In organic chemistry, the process of breaking down a saccharide
Changing the order of words in a sentence
Changing word order for emphasis
White trail behind the plane
Changing word order
Reverse element order
Changing the usual word order in a sentence to enhance expressiveness of speech
In the first sections we became acquainted in general terms with the vertical structure of the atmosphere and with changes in temperature with altitude.
Here we will look at some interesting features temperature regime in the troposphere and in the overlying spheres.
Temperature and humidity in the troposphere. The troposphere is the most interesting area, since rock-forming processes are formed here. In the troposphere, as already indicated in Chapter I, the air temperature decreases with height by an average of 6° for each kilometer rise, or by 0.6° per 100 m. This value of the vertical temperature gradient is most often observed and is defined as the average of many measurements. In reality, the vertical temperature gradient in the Earth's temperate latitudes is variable. It depends on the seasons of the year, time of day, the nature of atmospheric processes, and in the lower layers of the troposphere - mainly on the temperature of the underlying surface.
In the warm season, when the layer of air adjacent to the surface of the earth is sufficiently heated, the temperature decreases with height. When the surface layer of air is strongly heated, the magnitude of the vertical temperature gradient exceeds even 1° for every 100 m raising.
In winter, with strong cooling of the earth's surface and the ground layer of air, instead of a decrease, an increase in temperature is observed with height, i.e., a temperature inversion occurs. The strongest and most powerful inversions are observed in Siberia, especially in Yakutia in winter, where clear and calm weather prevails, promoting radiation and subsequent cooling of the surface layer of air. Very often the temperature inversion here extends to a height of 2-3 km, and the difference between the air temperature at the surface of the earth and the upper boundary of the inversion is often 20-25°. Inversions are also typical for the central regions of Antarctica. In winter they are found in Europe, especially in its eastern part, Canada and other areas. From the magnitude of temperature change with height (vertical temperature gradient) in to a large extent weather conditions and types of air movements in the vertical direction depend.
Stable and unstable atmosphere. The air in the troposphere is heated by the underlying surface. Air temperature varies with altitude and depending on atmospheric pressure. When this happens without heat exchange with environment, then such a process is called adiabatic. Rising air produces work due to internal energy, which is spent on overcoming external resistance. Therefore, as the air rises, it cools, and as it descends, it heats up.
Adiabatic temperature changes occur according to dry adiabatic And moist adiabatic laws.
Accordingly, vertical gradients of temperature changes with height are also distinguished. Dry adiabatic gradient- is the change in temperature of dry or humid unsaturated air for every 100 m raising and lowering it by 1 °, A moist adiabatic gradient- is a decrease in the temperature of moist saturated air for every 100 m elevation less than 1°.
When dry or unsaturated air rises or falls, its temperature changes according to the dry-adiabatic law, i.e., it falls or rises, respectively, by 1° every 100 m. This value does not change until the air, when rising, reaches a state of saturation, i.e. condensation level water vapor. Above this level, due to condensation, latent heat of vaporization begins to be released, which is used to heat the air. This additional heat reduces the amount of cooling the air receives as it rises. Further rise of saturated air occurs according to the moist-adiabatic law, and its temperature decreases by no more than 1° per 100 m, but less. Since the moisture content of air depends on its temperature, the higher the air temperature, the more heat is released during condensation, and the lower the temperature, the less heat. Therefore, the moisture-adiabatic gradient in warm air is less than in cold air. For example, at a temperature at the surface of the earth of rising saturated air +20°, the moist adiabatic gradient in the lower troposphere is 0.33-0.43° per 100 m, and at a temperature of minus 20° its values range from 0.78° to 0.87° by 100 m.
The moist adiabatic gradient also depends on air pressure: the lower the air pressure, the lower the moist adiabatic gradient at the same initial temperature. This happens because at low pressure the air density is also less, therefore, the released heat of condensation goes to heat a smaller mass of air.
Table 15 shows the averaged values of the moisture-adiabatic gradient at various temperatures and values
pressure 1000, 750 and 500 mb, which approximately corresponds to the surface of the earth and heights of 2.5-5.5 km.
In the warm season, the vertical temperature gradient is on average 0.6-0.7° per 100 m raising.
Knowing the temperature at the earth's surface, it is possible to calculate approximate temperature values at various altitudes. If, for example, the air temperature at the surface of the earth is 28°, then, assuming that the vertical temperature gradient is on average 0.7° per 100 m or 7° per kilometer, we get that at an altitude of 4 km temperature is 0°. The temperature gradient in winter in mid-latitudes over land rarely exceeds 0.4-0.5° per 100 m: There are often cases when in certain layers of air the temperature almost does not change with height, i.e., isothermia occurs.
By the magnitude of the vertical gradient of air temperature, one can judge the nature of the equilibrium of the atmosphere - stable or unstable.
At stable equilibrium atmosphere, air masses do not tend to move vertically. In this case, if a certain volume of air is displaced upward, it will return to its original position.
Stable equilibrium occurs when the vertical temperature gradient of unsaturated air is less than the dry adiabatic gradient, and the vertical temperature gradient of saturated air is less than the moist adiabatic one. If, under this condition, a small volume of unsaturated air is raised to a certain height by external influence, then as soon as the action ceases external force, this volume of air will return to its previous position. This happens because the raised volume of air, having spent internal energy on its expansion, cooled by 1° for every 100 m(according to the dry adiabatic law). But since the vertical temperature gradient of the surrounding air was less than the dry adiabatic one, it turned out that the raised volume of air at a given altitude had a lower temperature than the surrounding air. Having a higher density compared to the density of the surrounding air, it must sink until it reaches its original state. Let's show this with an example.
Let us assume that the air temperature at the earth's surface is 20°, and the vertical temperature gradient in the layer under consideration is 0.7° per 100 m. With this gradient value, the air temperature at an altitude of 2 km will be equal to 6° (Fig. 19, A). Under the influence of an external force, a volume of unsaturated or dry air raised from the surface of the earth to this height, cooling according to the dry adiabatic law, i.e. by 1° per 100 m, will cool by 20° and take on a temperature equal to 0°. This volume of air will be 6° colder than the surrounding air, and therefore heavier due to its higher density. So he will start
descend, trying to reach the original level, i.e., the surface of the earth.
A similar result will be obtained in the case of rising saturated air, if the vertical gradient of the ambient temperature is less than the moist adiabatic one. Therefore, in a stable state of the atmosphere in a homogeneous mass of air, the rapid formation of cumulus and cumulonimbus clouds does not occur.
The most stable state of the atmosphere is observed at small values of the vertical temperature gradient, and especially during inversions, since in this case warmer and lighter air is located above the lower cold, and therefore heavy, air.
At unstable atmospheric equilibrium The volume of air raised from the surface of the earth does not return to its original position, but maintains its upward movement to a level at which the temperatures of the rising and surrounding air are equalized. The unstable state of the atmosphere is characterized by large vertical temperature gradients, which are caused by heating of the lower layers of air. At the same time, the heated air masses below, being lighter, rush upward.
Suppose, for example, that unsaturated air in the lower layers up to a height of 2 km stratified unstably, i.e. its temperature
decreases with altitude by 1.2° for every 100 m, and above the air, having become saturated, has a stable stratification, i.e. its temperature drops by 0.6° for every 100 m uplifts (Fig. 19, b). Once in such an environment, the volume of dry unsaturated air will rise according to the dry adiabatic law, i.e., cool by 1° per 100 m. Then, if its temperature at the surface of the earth is 20°, then at an altitude of 1 km it will become equal to 10°, while the ambient temperature is 8°. Being 2° warmer, and therefore lighter, this volume will rush higher. At altitude 2 km it will be warmer than the environment by 4°, since its temperature will reach 0°, and the ambient air temperature is -4°. Being lighter again, the volume of air in question will continue to rise to a height of 3 km, where its temperature becomes equal to the ambient temperature (-10°). After this, the free rise of the allocated volume of air will stop.
To determine the state of the atmosphere are used aerological diagrams. These are diagrams with rectangular coordinate axes along which the characteristics of the state of the air are plotted.
Families are shown on aerological diagrams dry And wet adiabats, i.e., curves graphically representing the change in the state of air during dry adiabatic and wet adiabatic processes.
Figure 20 shows such a diagram. Here, isobars are depicted vertically, isotherms (lines of equal air pressure) are shown horizontally, inclined solid lines are dry adiabats, inclined broken lines are wet adiabats, dotted lines specific humidity The diagram below shows curves of changes in air temperature with height at two points at the same observation period - 15 hours on May 3, 1965. On the left is the temperature curve according to the radiosonde data released in Leningrad, on the right - in Tashkent. From the shape of the left curve of temperature change with height it follows that in Leningrad the air is stable. Moreover, up to the isobaric surface 500 mb the vertical temperature gradient is on average 0.55° per 100 m. In two small layers (on surfaces 900 and 700 mb) isothermia registered. This indicates that over Leningrad at altitudes of 1.5-4.5 km located atmospheric front, separating the cold air masses in the lower one and a half kilometers from the warm air located above. The height of the condensation level, determined by the position of the temperature curve in relation to the wet adiabat, is about 1 km(900 mb).
In Tashkent, the air had an unstable stratification. Up to height 4 km the vertical temperature gradient was close to adiabatic, i.e. for every 100 m As the temperature rose, the temperature decreased by 1°, and above that, to 12 km- more adiabatic. Due to the dry air, cloud formation did not occur.
Over Leningrad, the transition to the stratosphere occurred at an altitude of 9 km(300 mb), and above Tashkent it is much higher - about 12 km(200 MB).
With a stable state of the atmosphere and sufficient humidity, stratus clouds and fogs can form, and with an unstable state and high moisture content of the atmosphere, thermal convection, leading to the formation of cumulus and cumulonimbus clouds. The state of instability is associated with the formation of showers, thunderstorms, hail, small whirlwinds, squalls, etc.
n. The so-called “bumpiness” of the aircraft, i.e. the aircraft throwing during flight, is also caused by the unstable state of the atmosphere.
In summer, atmospheric instability is common in the afternoon, when layers of air close to the earth's surface heat up. Therefore, heavy rains, squalls and the like dangerous phenomena weather conditions are more often observed in the afternoon, when strong vertical currents arise due to breaking instability - ascending And descending air movement. For this reason, aircraft flying during the day at an altitude of 2-5 km above the surface of the earth, they are more subject to “bumpiness” than during a night flight, when, due to the cooling of the surface layer of air, its stability increases.
Air humidity also decreases with altitude. Almost half of all humidity is concentrated in the first one and a half kilometers of the atmosphere, and the first five kilometers contain almost 9/10 of all water vapor.
To illustrate the daily observed nature of temperature changes with height in the troposphere and lower stratosphere in different regions of the Earth, Figure 21 shows three stratification curves up to a height of 22-25 km. These curves were constructed based on radiosonde observations at 3 pm: two in January - Olekminsk (Yakutia) and Leningrad, and the third in July - Takhta-Bazar ( middle Asia). The first curve (Olekminsk) is characterized by the presence of a surface inversion, characterized by an increase in temperature from -48° at the earth's surface to -25° at an altitude of about 1 km. At this time, the tropopause above Olekminsk was at an altitude of 9 km(temperature -62°). In the stratosphere, an increase in temperature was observed with altitude, the value of which was at 22 km was approaching -50°. The second curve, representing the change in temperature with height in Leningrad, indicates the presence of a small surface inversion, then isotherm in a large layer and a decrease in temperature in the stratosphere. At level 25 km the temperature is -75°. The third curve (Takhta-Bazar) is very different from the northern point - Olekminsk. The temperature at the earth's surface is above 30°. The tropopause is located at an altitude of 16 km, and above 18 km The usual temperature increase with height for southern summer occurs.
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The sun's rays falling on the surface of the earth heat it. Heating of the air occurs from the bottom up, i.e. from the earth's surface.
The transfer of heat from the lower layers of air to the upper layers occurs mainly due to the rise of warm, heated air upward and the lowering of cold air downwards. This process of heating air is called convection.
In other cases, upward heat transfer occurs due to dynamic turbulence. This is the name given to random vortices that arise in the air as a result of its friction against the earth's surface during horizontal movement or when different layers of air rub against each other.
Convection is sometimes called thermal turbulence. Convection and turbulence are sometimes combined under the common name - exchange.
Cooling of the lower atmosphere occurs differently than heating. The earth's surface continuously loses heat into the atmosphere surrounding it by emitting heat rays invisible to the eye. The cooling becomes especially severe after sunset (at night). Thanks to thermal conductivity, the air masses adjacent to the ground are also gradually cooled, then transferring this cooling to the overlying layers of air; in this case, the lowest layers are cooled most intensively.
Depending on the solar heating the temperature of the lower air layers varies throughout the year and day, reaching a maximum around 13-14 hours. Daily cycle air temperatures on different days for the same place are not constant; its magnitude depends mainly on weather conditions. Thus, changes in the temperature of the lower layers of air are associated with changes in the temperature of the earth's (underlying) surface.
Changes in air temperature also occur from its vertical movements.
It is known that air cools when it expands, and heats up when compressed. In the atmosphere, during upward movement, air, falling into areas of lower pressure, expands and cools, and, conversely, during downward movement, air, compressing, heats up. Changes in air temperature during its vertical movements largely determine the formation and destruction of clouds.
Air temperature usually decreases with height. The change in average temperature with altitude over Europe in summer and winter is given in the table “Average air temperatures over Europe”. 
The decrease in temperature with height is characterized by a vertical temperature gradient. This is the name for the change in temperature for every 100 m of altitude. For technical and aeronautical calculations, the vertical temperature gradient is taken equal to 0.6. It must be kept in mind that this value is not constant. It may happen that in some layer of air the temperature does not change with height.
Such layers are called isothermal layers.
Quite often in the atmosphere there is a phenomenon when in a certain layer the temperature even increases with height. These layers of the atmosphere are called layers of inversion. Inversions occur for various reasons. One of them is cooling the underlying surface by radiation at night or winter time under clear skies. Sometimes, in the case of calm or weak wind, the surface air also cools and becomes colder than the overlying layers. As a result, the air at altitude is warmer than at the bottom. Such inversions are called radiation. Strong radiation inversions are usually observed over snow cover and especially in mountain basins, and also during calm conditions. Inversion layers extend to heights of several tens or hundreds of meters.
Inversions also arise due to movement (advection) warm air onto a cold underlying surface. These are the so-called advective inversions. The height of these inversions is several hundred meters.
In addition to these inversions, frontal inversions and compression inversions are observed. Frontal inversions occur when warm air masses flow over colder ones. Compression inversions occur when air descends from the upper layers of the atmosphere. In this case, the descending air sometimes heats up so much that its underlying layers turn out to be colder.
Temperature inversions are observed at various altitudes in the troposphere, most often at altitudes of about 1 km. The thickness of the inversion layer can vary from several tens to several hundred meters. The temperature difference during inversion can reach 15-20°.
Inversion layers play a big role in weather. Because the air in the inversion layer is warmer than the underlying layer, the air in the lower layers cannot rise. Consequently, inversion layers retard vertical movements in the underlying air layer. When flying under an inversion layer, a bump (“bumpiness”) is usually observed. Above the inversion layer, the flight of an aircraft usually occurs normally. So-called wavy clouds develop under the inversion layers.
Air temperature influences piloting technique and equipment operation. At ground temperatures below -20°, the oil freezes, so it must be poured in a heated state. In flight at low temperatures The water in the engine cooling system is intensively cooled. At elevated temperatures (above +30°), the motor may overheat. Air temperature also affects the performance of the aircraft crew. At low temperatures, reaching -56° in the stratosphere, special uniforms are required for the crew.
The air temperature is very great importance for weather forecast.
Air temperature is measured during an airplane flight using electric thermometers attached to the airplane. When measuring air temperature, it is necessary to keep in mind that due to the high speeds of modern aircraft, thermometers give errors. High aircraft speeds cause an increase in the temperature of the thermometer itself, due to the friction of its reservoir with the air and the influence of heating due to air compression. Heating from friction increases with increasing aircraft flight speed and is expressed by the following quantities:
Speed in km/h…………. 100 200 З00 400 500 600
Heating from friction……. 0°.34 1°.37 3°.1 5°.5 8°.6 12°.b
Heating from compression is expressed by the following quantities:
Speed in km/h…………. 100 200 300 400 500 600
Heating from compression……. 0°.39 1°.55 3°.5 5°.2 9°.7 14°.0
The distortion of the readings of a thermometer installed on an airplane when flying in the clouds is 30% less than the above values, due to the fact that part of the heat generated by friction and compression is spent on evaporating water condensed in the air in the form of droplets.
Air temperature. Units of measurement, temperature change with altitude. Inversion, isothermy, Types of inversions, Adiabatic process.
Air temperature is a quantity that characterizes it thermal state. It is expressed either in degrees Celsius (ºС on the centigrade scale or in Kelvin (K) on the absolute scale. The transition from temperature in Kelvin to temperature in degrees Celsius is carried out according to the formula
t = T-273º
The lower layer of the atmosphere (troposphere) is characterized by a decrease in temperature with height, amounting to 0.65ºС per 100 m.
This change in temperature with height per 100m is called the vertical temperature gradient. Knowing the temperature at the surface of the earth and using the value of the vertical gradient, you can calculate the approximate temperature at any altitude (for example, at a temperature at the surface of the earth +20ºС at an altitude of 5000 m, the temperature will be equal to:
20º- (0.65*50) = - 12.5.
The vertical gradient γ is not constant and depends on the type air mass, time of day and season of the year, the nature of the underlying surface and other reasons. When the temperature decreases with height, γ is considered positive; if the temperature does not change with height, then γ = 0 layers are called isothermal. Layers of the atmosphere where temperature increases with height (γ< 0), называются inversion. Depending on the magnitude of the vertical temperature gradient, the state of the atmosphere can be stable, unstable, or indifferent in relation to dry (unsaturated) or saturated air.
The air temperature decreases as it rises adiabatically, that is, without heat exchange of air particles with the environment. If an air particle rises upward, then its volume expands, and the internal energy of the particle decreases.
If a particle descends, it contracts and its internal energy increases. It follows from this that when the volume of air moves upward, its temperature decreases, and when it moves downward, it increases. These processes play an important role in the formation and development of clouds.
The horizontal gradient is the temperature expressed in degrees over a distance of 100 km. When moving from a cold VM to a warm one and from a warm one to a cold one, it can exceed 10º per 100 km.
Types of inversions.
Inversions are retarding layers, they dampen vertical air movements, under them there is an accumulation of water vapor or other solid particles that impair visibility, the formation of fog and various forms clouds Inversion layers are also braking layers for horizontal air movements. In many cases, these layers are wind break surfaces. Inversions in the troposphere can be observed near the earth's surface and on high altitudes. A powerful layer of inversion is the tropopause.
Depending on the causes of occurrence, the following types of inversions are distinguished:
1. Radiation - the result of cooling of the surface layer of air, usually at night.
2. Advective - when warm air moves to a cold underlying surface.
3. Compression or subsidence - formed in the central parts of low-moving anticyclones.
13. Change in air temperature with altitude.
The vertical distribution of temperature in the atmosphere forms the basis for dividing the atmosphere into five main layers. For agricultural meteorology, the patterns of temperature changes in the troposphere, especially in its surface layer, are of greatest interest.
Vertical temperature gradient
The change in air temperature per 100 m of altitude is called the vertical temperature gradient (VHT depends on a number of factors: time of year (less in winter, more in summer), time of day (less at night, more during the day), location of air masses (if at any altitudes above a cold layer of air is located in a layer of warmer air, then the VGT reverses sign).The average value of VGT in the troposphere is about 0.6 °C/100 m.
In the surface layer of the atmosphere, the VGT depends on the time of day, weather and the nature of the underlying surface. During the day, the VGT is almost always positive, especially in summer over land, but in clear weather it is tens of times greater than in cloudy weather. On a clear afternoon in summer, the air temperature at the soil surface can be 10 °C or more higher than the temperature at a height of 2 m. As a result, the VGT in a given two-meter layer in terms of 100 m is more than 500 °C/100 m. Wind reduces the VGT, since at When the air is mixed, its temperature at different altitudes is equalized. Cloudiness and precipitation reduce VGT. When the soil is wet, the VGT in the surface layer of the atmosphere sharply decreases. Above bare soil ( steam field) VHT is greater than over a developed crop or meadow. In winter, above the snow cover, the VGT in the surface layer of the atmosphere is small and often negative.
With height, the influence of the underlying surface and weather on the VGT weakens and the VGT decreases compared to its values in the surface layer of air. Above 500 m, the influence of the daily variation of air temperature fades. At altitudes from 1.5 to 5-6 km, the VGT is within 0.5-0.6 ° C/100 m. At an altitude of 6-9 km, the VGT increases and is 0.65-0.75 ° C/100 m. in the upper layer of the troposphere, the VGT again decreases to 0.5-0.2° C/100 m.
Data on VGT in various layers of the atmosphere are used in weather forecasting, in meteorological services for jet aircraft and in launching satellites into orbit, as well as in determining the conditions for the release and distribution of industrial waste in the atmosphere. Negative VGT in the surface layer of air at night in spring and autumn indicates the possibility of frost.
17. Characteristics of air humidity. Daily and annual variations in partial pressure of water vapor and relative humidity.
Atmospheric water vapor pressure - partial pressure of water vapor in the air
The Earth's atmosphere contains about 14 thousand km 3 of water vapor. Water enters the atmosphere as a result of evaporation from the underlying surface. In the atmosphere, moisture condenses, moves with air currents and falls again in the form of various precipitation on the surface of the Earth, thus completing a constant water cycle. The water cycle is possible thanks to the ability of water to be in three states (liquid, solid, gaseous (vapor)) and easily move from one state to another. Moisture circulation is one of the most important climate formation cycles.
To quantify the content of water vapor in the atmosphere, various characteristics of air humidity are used. The main characteristics of air humidity are the elasticity of water vapor and relative humidity.
Elasticity (actual) of water vapor (e) - the pressure of water vapor in the atmosphere is expressed in mmHg. or in millibars (mb). Numerically, it almost coincides with absolute humidity (the content of water vapor in the air in g/m3), which is why elasticity is often called absolute humidity. Saturation elasticity (maximum elasticity) (E) is the limit of water vapor content in the air at a given temperature. The value of saturation elasticity depends on the air temperature; the higher the temperature, the more water vapor it can contain.
The daily variation of humidity (absolute) can be simple or double. The first coincides with the daily variation of temperature, has one maximum and one minimum and is typical for places with sufficient moisture. It is observed over the oceans, and over land in winter and autumn.
The double move has two maximums and two minimums and is typical for the summer season on land: maximums at 9 and 20-21 hours, and minimums at 6 and 16 hours.
The morning minimum before sunrise is explained by weak evaporation during the night hours. With increasing radiant energy, evaporation increases, and the water vapor pressure reaches a maximum at about 9 hours.
As a result of heating the surface, air convection develops; moisture transfer occurs faster than its entry from the evaporating surface, so at about 16 o'clock a second minimum occurs. By evening, convection stops, but evaporation from the heated surface is still quite intense and moisture accumulates in the lower layers, providing a second maximum at about 20-21 hours.
The annual variation of water vapor pressure corresponds to the annual variation of temperature. In summer the pressure of water vapor is greater, in winter it is less.
Daily allowance and annual course relative humidity is almost everywhere opposite to the temperature trend, since the maximum moisture content with increasing temperature increases faster than the elasticity of water vapor. The daily maximum of relative humidity occurs before sunrise, the minimum - at 15-16 hours.
During the year, the maximum relative humidity usually occurs in the coldest month, and the minimum in the warmest month. The exception is in regions where humid winds blow from the sea in summer and dry winds from the mainland in winter.
Absolute humidity = the amount of water in a given volume of air, measured in (g/m³)
Relative humidity = percentage of the actual quantity of water (water vapor pressure) to the vapor pressure of water at that temperature under saturated conditions. Expressed as a percentage. Those. 40% humidity means that at this temperature, another 60% of the total water can evaporate.
Change in air temperature with altitude
Exercise 1. Determine what temperature the air mass will have, not saturated with water vapor and rising adiabatically at an altitude of 500, 1000, 1500 m, if its temperature at the surface of the earth was 15 degrees.
The temperature changes by 1° when the air mass rises for every 100 m. This value is called dry adiabatic temperature gradient. When air saturated with water vapor rises, its cooling rate decreases somewhat, since condensation of water vapor occurs, during which latent heat of vaporization is released (600 cal per 1 g of condensed water), which is used to heat this rising air. The adiabatic process occurring inside rising saturated air is called moist adiabatic. The amount of temperature decrease (increase) for every 100 m in a rising moist saturated air mass is called moist adiabatic temperature gradient g V , and the graph of temperature changes with height in such a process is called wet adiabatic. In contrast to the dry adiabatic gradient g a, the wet adiabatic gradient g b is a variable value, depending on temperature and pressure, and ranges from 0.3° to 0.9° per 100 m of height (on average 0.6° per 100 m. ). The more moisture condenses as the air rises, the smaller the value of the moisture-adiabatic gradient; with a decrease in the amount of moisture, its value approaches the dry adiabatic gradient.
The vertical temperature gradient at an altitude of 500 meters should be = 12 °. The vertical temperature gradient at an altitude of 1000 meters should be = 9 °. The vertical temperature gradient at an altitude of 1500 meters should be = 6 °. But as soon as the air begins to rise, it will become colder than the surrounding air, and the temperature difference increases with altitude.
But cold air, being heavier, tends to descend, i.e. take the original position. Since the air is unsaturated, as it rises the temperature should decrease by 1°C per 100 m.
Therefore, the temperature of the air mass at an altitude of 500 meters will be = 10°C. Therefore, the temperature of the air mass at an altitude of 1000 meters will be = 5°C. Therefore, the temperature of the air mass at an altitude of 1500 meters will be = 0°C.
Determination of the height of condensation and sublimation levels
Exercise 1. Determine the height of the level of condensation and sublimation of adiabatically rising air not saturated with water vapor, if its temperature (T) and water vapor pressure (e) are known; T = 18є, e = 13.6 hPa.
The temperature of rising air, not saturated with water vapor, changes by 1° every 100 meters. First, using the curve of maximum vapor pressure versus air temperature, you need to find the dew point (φ). Then determine the difference between the air temperature and the dew point (T - f). Multiply this value by 100 m to find the condensation level. To determine the level of sublimation, you need to find the temperature difference from the dew point to the sublimation temperature and multiply this difference by 200 m.
The condensation level is the level to which it must rise before the water vapor contained in the air during adiabatic rise reaches a state of saturation (or 100% relative humidity). The height at which water vapor in the rising air becomes saturated can be found using the formula: , where T is the air temperature; f - dew point.
f = 2.064 (according to the table)
18 є - 2.064 = 15.936 є x 122 = 1994 m saturation height of water vapor.
Sublimation occurs at a temperature of - 10°.
2.064 - (-10) = 12.064 x 200 = 2413m sublimation level.
Task 2 (B). Air having a temperature of 12°C and a relative humidity of 80% passes over mountains 1500 m high. At what altitude will clouds begin to form? What is the temperature and relative humidity at the top of the ridge and behind the ridge?
If the relative air humidity r is known, then the height of the condensation level can be determined using Ippolitov’s formula: h = 22 (100-r) h = 22 (100-80) = 440 m the beginning of the formation of stratus clouds.
The process of cloud formation begins with the fact that a certain mass of sufficiently moist air rises upward. As you rise, the air will expand. This expansion can be considered adiabatic, since the air rises quickly, and if its volume is large enough, the heat exchange between the air in question and the environment during the rise simply does not have time to occur.
When a gas expands adiabatically, its temperature decreases. So, rising up wet air will cool down. When the temperature of the cooling air drops to the dew point, the process of condensation of the steam contained in the air becomes possible. If there are a sufficient number of condensation nuclei in the atmosphere, this process begins. If there are few condensation nuclei in the atmosphere, condensation begins not at a temperature equal to the dew point, but at lower temperatures.
Having reached a height of 440m, the rising moist air will cool and condensation of water vapor will begin. Height 440m is the lower boundary of the forming cloud. The air that continues to flow from below passes through this boundary, and the process of vapor condensation will occur above the specified boundary - the cloud will begin to develop in height. The vertical development of the cloud will stop when the air stops rising; in this case, the upper boundary of the cloud will form.
The temperature at the top of the ridge is +3 °C and the relative air humidity is 100%.
local time dry adiabatic gradient
The air temperature in the troposphere as a whole decreases by an average of 0.6 °C for every 100 m of altitude. However, in the lower layer (up to 100-150 m), the temperature distribution can be different: it can increase, remain constant or decrease.
When the temperature decreases with distance from the active surface, such a distribution, as noted in Section. 3.4, called insolation. In the air over land this happens in the warm season during the daytime in clear weather. During insolation, favorable conditions are created for the development of thermal convection (see Section 4.1) and the formation of clouds.
When the air temperature does not change with altitude, this condition is called "isothermia". Temperature isothermia is observed in cloudy, calm weather.
If the air temperature increases with distance from the surface, this temperature distribution is called inversion.
Depending on the conditions for the formation of inversions in the surface layer of the atmosphere, they are divided into radiative and advective.
Radiative inversions arise during radiation cooling of the active surface. Such inversions form at night during the warm season, and are also observed during the day in winter. Therefore, radiation inversions are divided into nighttime (summer) and winter.
Night inversions are established in clear, quiet weather after the radiation balance passes through zero 1.0... 1.5 hours before sunset. During the night they intensify and reach their greatest power before sunrise. After sunrise, the active surface and air warm up, which destroys the inversion. The height of the inversion layer is most often several tens of meters, but under certain conditions (for example, in closed valleys surrounded by significant elevations) it can reach 200 m or more. This is facilitated by the flow of cooled air from the slopes into the valley. Cloudiness weakens the inversion, and wind speed is more
2.5...3.0 m/s destroys it. Under the canopy of dense grass, crops, and gardens in summer, inversions are also observed during the day (Fig. 4.4, b).
Night radiation inversions in spring and autumn, and in some places in summer, can cause a decrease in soil and air surface temperatures to negative values (freezing), which causes damage to cultivated plants.
Winter inversions occur in clear, calm weather in conditions short day, when the cooling of the active surface is continuous
Rice. 4.4.
1 - at night; 2 - during the day it increases every day. They can persist for several weeks, weakening slightly during the day and getting stronger again at night.
Radiation inversions are especially intensified under highly heterogeneous terrain. The cooling air flows into lowlands and basins, where weakened turbulent mixing contributes to its further cooling. Radiation inversions associated with terrain features are usually called orographic. They are dangerous for fruit trees and berry bushes, since the air temperature during such inversions can drop to critical levels.
Advective inversions are formed by the advection of warm air onto a cold underlying surface, which cools the adjacent layers of advancing air. These inversions also include snow inversions. They arise when air with a temperature above 0 °C advects onto a surface covered with snow. The decrease in temperature in the lowest layer in this case is associated with the heat consumed by snow melting.
