How does the temperature in the mountains change with altitude? The atmosphere of the earth and the physical properties of air. Change in air temperature

In the troposphere, air temperature decreases with altitude, as noted, by an average of 0.6 ºС for every 100 m of altitude. However, in the surface layer the temperature distribution can be different: it can decrease, increase, or remain constant. The vertical temperature gradient (VTG) gives an idea of ​​the distribution of temperature with height:

The value of VGT in the surface layer depends on weather conditions (in clear weather it is greater than in cloudy weather), time of year (more in summer than in winter) and time of day (more during the day than at night). Wind reduces the VGT, since when the air is mixed, its temperature at different altitudes is equalized. Above moist soil, the VGT in the ground layer sharply decreases, and above bare soil (fallow field) the VGT is greater than over dense crops or meadows. This is due to differences in temperature conditions these surfaces.

The change in air temperature with height determines the sign of the VGT: if VGT > 0, then the temperature decreases with distance from the active surface, which usually happens during the day and summer; if VGT = 0, then the temperature does not change with height; if VGT< 0, то температура увеличивается с высотой и такое распределение температуры называют инверсией.

Depending on the conditions for the formation of inversions in the surface layer of the atmosphere, they are divided into radiative and advective.

1. Radiation inversions occur during radiative cooling earth's 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.

2. Advective inversions are formed during advection (movement) warm air onto the cold underlying surface, which cools the adjacent layers of advancing air. These inversions also include snow inversions. They occur 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.

Air temperature measurement

At meteorological stations, thermometers are installed in a special booth, called a psychrometric booth, the walls of which are louvered. The rays of the Sun do not penetrate into such a booth, but at the same time air has free access to it.

Thermometers are installed on a tripod so that the reservoirs are located at a height of 2 m from the active surface.

Urgent air temperature is measured with a mercury psychrometric thermometer TM-4, which is installed vertically. At temperatures below -35°C, use a low-degree alcohol thermometer TM-9.

Extreme temperatures are measured using maximum TM-1 and minimum TM-2 thermometers, which are laid horizontally.

For continuous recording of air temperature, use thermograph M-16A, which is placed in a louvred recording booth. Depending on the rotation speed of the drum, thermographs are available for daily or weekly use.

In crops and plantings, the air temperature is measured without disturbing the vegetation cover. For this purpose, an aspiration psychrometer is used.

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 more high levels the energies of its constituent particles are more “populated” by particles than the lower ones

IN organic chemistry- process of saccharide breakdown

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 met in general outline with the vertical structure of the atmosphere and with temperature changes 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 fact, the vertical temperature gradient in temperate latitudes The earth is changeable. It depends on the seasons of the year, time of day, character atmospheric processes, and in the lower layers of the troposphere - mainly from 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 characteristic of central regions Antarctica. In winter they are found in Europe, especially in its eastern part, Canada and other areas. The magnitude of the temperature change with height (vertical temperature gradient) largely determines the weather conditions and types of air movements in the vertical direction.

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 exchanging heat with the environment, the 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 of the external force ceases, 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 gradient of the ambient air temperature was less than the dry adiabatic one, it turned out that the raised volume of air at a given altitude had more low 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 temperature gradient environment less than moist adiabatic. 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 will its temperature become equal temperature environment (-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 there is an 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 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 thing happens southern summer temperature increases with altitude.

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Sun rays falling on the surface of the earth heats it up. 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 downward. 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 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 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 the upward movement of air, falling into areas of more low pressure, expands and cools, and, conversely, with downward movement, the 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. Change 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 at clear sky. 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, I am also calm. Inversion layers extend to heights of several tens or hundreds of meters.

Inversions also occur due to the movement (advection) of 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 water flows in air masses to colder ones. Compression inversions occur when air is released from upper layers 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. During 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 aircraft. When measuring air temperature, it must be borne in mind that due to high speeds modern aircraft thermometers give errors. High speeds aircraft 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 characterizing its 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 a constant value and depends on the type of 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 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 lowering - formed in central parts slow-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 summer afternoon, 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. At wet soil VGT in the surface layer of the atmosphere sharply decreases. Over bare soil (fallow field) the VGT is greater than over developed crops or meadows. 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 water vapor pressure 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 characteristic of 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.

The daily and annual variation of relative humidity is almost everywhere opposite to the variation of temperature, 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.

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 downward. 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 solar heating, the temperature of the lower air layers varies throughout the year and day, reaching a maximum around 13-14 hours. The daily variation of air temperature on different days for the same place is 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 in winter 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 occur due to the movement (advection) of 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. During 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.

Air temperature is very important for weather forecasting.

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.

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 consider some interesting features of the 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 the chapter I, the air temperature decreases with altitude 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. The magnitude of the temperature change with height (vertical temperature gradient) largely determines the weather conditions and types of air movements in the vertical direction.

Stable and unstable atmosphere. The air in the troposphere is heated by the underlying surface. Air temperature changes with altitude and depending on atmospheric pressure. When this happens without exchanging heat with the environment, the 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 100m.

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 2.5-5.5km.

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 of the external force ceases, 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 there is an 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 vortices, squalls, etc. The so-called “bumpiness” of the aircraft, i.e. the aircraft bouncing 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, torrential rains, squalls and similar dangerous weather phenomena 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 (Central 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.

- Source-

Pogosyan, Kh.P. Atmosphere of the Earth / H.P. Poghosyan [and others]. – M.: Education, 1970.- 318 p.

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