Formation of the atmosphere. Today, the Earth's atmosphere is a mixture of gases - 78% nitrogen, 21% oxygen and small amounts of other gases, such as carbon dioxide. But when the planet first appeared, there was no oxygen in the atmosphere - it consisted of gases that originally existed in the solar system.
Earth arose when small rocky bodies made of dust and gas from the solar nebula, known as planetoids, collided with each other and gradually took the shape of a planet. As it grew, the gases contained in the planetoids burst out and enveloped the globe. After some time, the first plants began to release oxygen, and the primordial atmosphere developed into the current dense air envelope.
Origin of the atmosphere
- A rain of small planetoids fell on the nascent Earth 4.6 billion years ago. Gases from the solar nebula trapped inside the planet burst out during the collision and formed the Earth's primitive atmosphere, consisting of nitrogen, carbon dioxide and water vapor.
- The heat released during the formation of the planet is retained by a layer of dense clouds in the primordial atmosphere. "Greenhouse gases" such as carbon dioxide and water vapor stop the radiation of heat into space. The surface of the Earth is flooded with a seething sea of molten magma.
- When planetoid collisions became less frequent, the Earth began to cool and oceans appeared. Water vapor condenses from thick clouds, and rain, lasting for several eons, gradually floods the lowlands. Thus the first seas appear.
- The air is purified as water vapor condenses to form oceans. Over time, carbon dioxide dissolves in them, and the atmosphere is now dominated by nitrogen. Due to the lack of oxygen, the protective ozone layer does not form, and ultraviolet rays from the sun reach the earth's surface without hindrance.
- Life appears in ancient oceans within the first billion years. The simplest blue-green algae are protected from ultraviolet radiation by seawater. They use sunlight and carbon dioxide to produce energy, releasing oxygen as a byproduct, which gradually begins to accumulate in the atmosphere.
- Billions of years later, an oxygen-rich atmosphere forms. Photochemical reactions in the upper atmosphere create a thin layer of ozone that scatters harmful ultraviolet light. Life can now emerge from the oceans onto land, where evolution produces many complex organisms.
Billions of years ago, a thick layer of primitive algae began releasing oxygen into the atmosphere. They survive to this day in the form of fossils called stromatolites.
Volcanic origin
1. Ancient, airless Earth. 2. Eruption of gases.
According to this theory, volcanoes were actively erupting on the surface of the young planet Earth. The early atmosphere likely formed when gases trapped in the planet's silicon shell escaped through volcanoes.
Accumulation of O 2 in the Earth's atmosphere:
1
. (3.85-2.45 billion years ago) - O 2 was not produced
2
. (2.45-1.85 billion years ago) O 2 was produced but absorbed by the ocean and seafloor rocks
3
. (1.85-0.85 billion years ago) O 2 leaves the ocean, but is consumed during the oxidation of rocks on land and during the formation of the ozone layer
4
. (0.85-0.54 billion years ago) all rocks on land are oxidized, accumulation of O 2 in the atmosphere begins
5
. (0.54 billion years ago - present) modern period, O 2 content in the atmosphere has stabilized
Oxygen disaster(oxygen revolution) - a global change in the composition of the Earth's atmosphere that occurred at the very beginning of the Proterozoic, about 2.4 billion years ago (the Siderian period). The result of the Oxygen Catastrophe was the appearance of free oxygen in the atmosphere and a change in the general character of the atmosphere from reducing to oxidizing. The assumption of an oxygen catastrophe was made based on a study of a sharp change in the nature of sedimentation.
Primary composition of the atmosphere
The exact composition of the Earth's primary atmosphere is currently unknown, but it is generally accepted that it was formed as a result of degassing of the mantle and was of a reducing nature. It was based on carbon dioxide, hydrogen sulfide, ammonia, and methane. This is supported by:
- unoxidized sediments formed clearly on the surface (for example, river pebbles from oxygen-labile pyrite);
- absence of known significant sources of oxygen and other oxidizing agents;
- study of potential sources of the primary atmosphere (volcanic gases, composition of other celestial bodies).
Causes of the oxygen catastrophe
The only significant source of molecular oxygen is the biosphere, or more precisely, photosynthetic organisms. Having appeared at the very beginning of the existence of the biosphere, photosynthetic archaebacteria produced oxygen, which was almost immediately spent on the oxidation of rocks, dissolved compounds and atmospheric gases. A high concentration was created only locally, within bacterial mats (so-called “oxygen pockets”). After the surface rocks and gases of the atmosphere became oxidized, oxygen began to accumulate in the atmosphere in free form.
One of the likely factors influencing the change in microbial communities was a change in the chemical composition of the ocean caused by the extinction of volcanic activity.
Consequences of the oxygen catastrophe
Biosphere
Since the overwhelming majority of organisms of that time were anaerobic, unable to exist at significant oxygen concentrations, a global change in communities occurred: anaerobic communities were replaced by aerobic ones, previously limited only to “oxygen pockets”; anaerobic communities, on the contrary, were pushed into “anaerobic pockets” (figuratively speaking, “the biosphere turned inside out”). Subsequently, the presence of molecular oxygen in the atmosphere led to the formation of an ozone screen, which significantly expanded the boundaries of the biosphere and led to the spread of more energetically favorable (compared to anaerobic) oxygen respiration.
Lithosphere
As a result of the oxygen catastrophe, virtually all metamorphic and sedimentary rocks that make up most of the Earth's crust are oxidized.
The marked increase in free oxygen in the Earth's atmosphere 2.4 billion years ago appears to have resulted from a very rapid transition from one equilibrium state to another. The first level corresponded to an extremely low concentration of O 2 - about 100,000 times lower than what is observed now. The second equilibrium level could have been achieved at a higher concentration, no less than 0.005 of the modern one. The oxygen content between these two levels is characterized by extreme instability. The presence of such “bistability” makes it possible to understand why there was so little free oxygen in the Earth’s atmosphere for at least 300 million years after cyanobacteria (blue-green “algae”) began to produce it.
Currently, the Earth's atmosphere consists of 20% free oxygen, which is nothing more than a by-product of photosynthesis by cyanobacteria, algae and higher plants. A lot of oxygen is released by tropical forests, which in popular publications are often called the lungs of the planet. At the same time, however, it is silent that during the year tropical forests consume almost as much oxygen as they produce. It is spent on the respiration of organisms that decompose finished organic matter - primarily bacteria and fungi. For that, In order for oxygen to begin to accumulate in the atmosphere, at least part of the substance formed during photosynthesis must be removed from the cycle- for example, get into bottom sediments and become inaccessible to bacteria that decompose it aerobically, that is, with the consumption of oxygen.
The total reaction of oxygenic (that is, “giving oxygen”) photosynthesis can be written as:
CO 2 + H 2 O + hν→ (CH 2 O) + O 2,
Where hν is the energy of sunlight, and (CH 2 O) is the generalized formula of organic matter. Breathing is the reverse process, which can be written as:
(CH 2 O) + O 2 → CO 2 + H 2 O.
At the same time, the energy necessary for organisms will be released. However, aerobic respiration is possible only at an O 2 concentration of no less than 0.01 of the modern level (the so-called Pasteur point). Under anaerobic conditions, organic matter decomposes through fermentation, and the final stages of this process often produce methane. For example, the generalized equation for methanogenesis through acetate formation looks like:
2(CH 2 O) → CH 3 COOH → CH 4 + CO 2.
If we combine the process of photosynthesis with the subsequent decomposition of organic matter under anaerobic conditions, then the overall equation will look like:
CO 2 + H 2 O + hν→ 1/2 CH 4 + 1/2 CO 2 + O 2.
It was precisely this path of decomposition of organic matter that apparently was the main one in the ancient biosphere.
Many important details of how the modern balance between oxygen supply and removal from the atmosphere was established remain unclear. After all, a noticeable increase in oxygen content, the so-called “Great Oxidation of the Atmosphere,” occurred only 2.4 billion years ago, although it is known for sure that cyanobacteria carrying out oxygenic photosynthesis were already quite numerous and active 2.7 billion years ago, and they arose even earlier - perhaps 3 billion years ago. Thus, within for at least 300 million years, the activity of cyanobacteria did not lead to an increase in oxygen content in the atmosphere.
The assumption that, for some reason, there suddenly was a radical increase in net primary production (that is, the increase in organic matter formed during the photosynthesis of cyanobacteria) did not stand up to criticism. The fact is that during photosynthesis, the light isotope of carbon 12 C is predominantly consumed, and in the environment the relative content of the heavier isotope 13 C increases. Accordingly, bottom sediments containing organic matter must be depleted in the isotope 13 C, which accumulates in water and goes for the formation of carbonates. However, the ratio of 12 C to 13 C in carbonates and in organic matter of sediments remains unchanged despite radical changes in the concentration of oxygen in the atmosphere. This means that the whole point is not in the source of O 2, but in its, as geochemists put it, “sink” (removal from the atmosphere), which suddenly decreased significantly, which led to a significant increase in the amount of oxygen in the atmosphere.
It is usually believed that immediately before the “Great Oxidation of the Atmosphere,” all the oxygen then formed was spent on the oxidation of reduced iron compounds (and then sulfur), which were quite abundant on the Earth’s surface. In particular, the so-called “banded iron ores” were formed then. But recently Colin Goldblatt, a graduate student in the School of Environmental Sciences at the University of East Anglia (Norwich, UK), together with two colleagues from the same university, came to the conclusion that the oxygen content in the earth's atmosphere can be in one of two equilibrium states: it can be either very small - about 100 thousand times less than now, or already quite a lot (although from the position of a modern observer it is small) - no less than 0.005 of the modern level.
In the proposed model, they took into account the entry into the atmosphere of both oxygen and reduced compounds, in particular paying attention to the ratio of free oxygen and methane. They noted that if the oxygen concentration exceeds 0.0002 of the current level, then some of the methane can already be oxidized by methanotroph bacteria according to the reaction:
CH 4 + 2O 2 → CO 2 + 2H 2 O.
But the rest of the methane (and there is quite a lot of it, especially at low oxygen concentrations) enters the atmosphere.
The entire system is in a nonequilibrium state from the point of view of thermodynamics. The main mechanism for restoring the disturbed equilibrium is the oxidation of methane in the upper layers of the atmosphere by hydroxyl radical (see Fluctuations of methane in the atmosphere: man or nature - who wins, "Elements", 10/06/2006). The hydroxyl radical is known to be formed in the atmosphere under the influence of ultraviolet radiation. But if there is a lot of oxygen in the atmosphere (at least 0.005 of the current level), then an ozone screen is formed in its upper layers, which well protects the Earth from hard ultraviolet rays and at the same time interferes with the physicochemical oxidation of methane.
The authors come to the somewhat paradoxical conclusion that the existence of oxygenic photosynthesis itself is not a sufficient condition either for the formation of an oxygen-rich atmosphere or for the emergence of an ozone screen. This circumstance should be taken into account in cases where we are trying to find signs of the existence of life on other planets based on the results of a survey of their atmosphere.
The atmosphere began to form along with the formation of the Earth. During the evolution of the planet and as its parameters approached modern values, fundamentally qualitative changes occurred in its chemical composition and physical properties. According to the evolutionary model, at an early stage the Earth was in a molten state and about 4.5 billion years ago formed as a solid body. This milestone is taken as the beginning of the geological chronology. From that time on, the slow evolution of the atmosphere began. Some geological processes (for example, lava outpourings during volcanic eruptions) were accompanied by the release of gases from the bowels of the Earth. They included nitrogen, ammonia, methane, water vapor, CO oxide and carbon dioxide CO 2. Under the influence of solar ultraviolet radiation, water vapor decomposed into hydrogen and oxygen, but the released oxygen reacted with carbon monoxide to form carbon dioxide. Ammonia decomposed into nitrogen and hydrogen. During the process of diffusion, hydrogen rose upward and left the atmosphere, and heavier nitrogen could not evaporate and gradually accumulated, becoming the main component, although some of it was bound into molecules as a result of chemical reactions ( cm. CHEMISTRY OF THE ATMOSPHERE). Under the influence of ultraviolet rays and electrical discharges, a mixture of gases present in the original atmosphere of the Earth entered into chemical reactions, which resulted in the formation of organic substances, in particular amino acids. With the advent of primitive plants, the process of photosynthesis began, accompanied by the release of oxygen. This gas, especially after diffusion into the upper layers of the atmosphere, began to protect its lower layers and the surface of the Earth from life-threatening ultraviolet and X-ray radiation. According to theoretical estimates, the oxygen content, 25,000 times less than now, could already lead to the formation of an ozone layer with only half the concentration than now. However, this is already enough to provide very significant protection of organisms from the destructive effects of ultraviolet rays.
It is likely that the primary atmosphere contained a lot of carbon dioxide. It was used up during photosynthesis, and its concentration must have decreased as the plant world evolved and also due to absorption during certain geological processes. Because the Greenhouse effect associated with the presence of carbon dioxide in the atmosphere, fluctuations in its concentration are one of the important reasons for such large-scale climate changes in the history of the Earth as ice ages.
The helium present in the modern atmosphere is mostly a product of the radioactive decay of uranium, thorium and radium. These radioactive elements emit a particles, which are the nuclei of helium atoms. Since during radioactive decay an electric charge is neither formed nor destroyed, with the formation of each a-particle two electrons appear, which, recombining with the a-particles, form neutral helium atoms. Radioactive elements are contained in minerals dispersed in rocks, so a significant part of the helium formed as a result of radioactive decay is retained in them, escaping very slowly into the atmosphere. A certain amount of helium rises upward into the exosphere due to diffusion, but due to the constant influx from the earth's surface, the volume of this gas in the atmosphere remains almost unchanged. Based on spectral analysis of starlight and the study of meteorites, it is possible to estimate the relative abundance of various chemical elements in the Universe. The concentration of neon in space is approximately ten billion times higher than on Earth, krypton - ten million times, and xenon - a million times. It follows that the concentration of these inert gases, apparently initially present in the Earth’s atmosphere and not replenished during chemical reactions, decreased greatly, probably even at the stage of the Earth’s loss of its primary atmosphere. An exception is the inert gas argon, since in the form of the 40 Ar isotope it is still formed during the radioactive decay of the potassium isotope.
Barometric pressure distribution.
The total weight of atmospheric gases is approximately 4.5 10 15 tons. Thus, the “weight” of the atmosphere per unit area, or atmospheric pressure, at sea level is approximately 11 t/m 2 = 1.1 kg/cm 2. Pressure equal to P 0 = 1033.23 g/cm 2 = 1013.250 mbar = 760 mm Hg. Art. = 1 atm, taken as the standard average atmospheric pressure. For the atmosphere in a state of hydrostatic equilibrium we have: d P= –rgd h, this means that in the height interval from h before h+d h occurs equality between the change in atmospheric pressure d P and the weight of the corresponding element of the atmosphere with unit area, density r and thickness d h. As a relationship between pressure R and temperature T The equation of state of an ideal gas with density r, which is quite applicable to the earth’s atmosphere, is used: P= r R T/m, where m is the molecular weight, and R = 8.3 J/(K mol) is the universal gas constant. Then dlog P= – (m g/RT)d h= – bd h= – d h/H, where the pressure gradient is on a logarithmic scale. Its inverse value H is called the atmospheric altitude scale.
When integrating this equation for an isothermal atmosphere ( T= const) or for its part where such an approximation is permissible, the barometric law of pressure distribution with height is obtained: P = P 0 exp(– h/H 0), where the height reference h produced from ocean level, where the standard mean pressure is P 0 . Expression H 0 = R T/ mg, is called the altitude scale, which characterizes the extent of the atmosphere, provided that the temperature in it is the same everywhere (isothermal atmosphere). If the atmosphere is not isothermal, then integration must take into account the change in temperature with height, and the parameter N– some local characteristic of atmospheric layers, depending on their temperature and the properties of the environment.
Standard atmosphere.
Model (table of values of the main parameters) corresponding to standard pressure at the base of the atmosphere R 0 and chemical composition is called a standard atmosphere. More precisely, this is a conditional model of the atmosphere, for which the average values of temperature, pressure, density, viscosity and other characteristics of air at altitudes from 2 km below sea level to the outer boundary of the earth’s atmosphere are specified for latitude 45° 32ў 33І. The parameters of the middle atmosphere at all altitudes were calculated using the equation of state of an ideal gas and the barometric law assuming that at sea level the pressure is 1013.25 hPa (760 mm Hg) and the temperature is 288.15 K (15.0 ° C). According to the nature of the vertical temperature distribution, the average atmosphere consists of several layers, in each of which the temperature is approximated by a linear function of height. In the lowest layer - the troposphere (h Ј 11 km) the temperature drops by 6.5 ° C with each kilometer of rise. At high altitudes, the value and sign of the vertical temperature gradient changes from layer to layer. Above 790 km the temperature is about 1000 K and practically does not change with altitude.
The standard atmosphere is a periodically updated, legalized standard, issued in the form of tables.
| Table 1. STANDARD MODEL OF THE EARTH'S ATMOSPHERE. The table shows: h– height from sea level, R- pressure, T– temperature, r – density, N– number of molecules or atoms per unit volume, H– height scale, l– free path length. Pressure and temperature at an altitude of 80–250 km, obtained from rocket data, have lower values. Values for altitudes greater than 250 km obtained by extrapolation are not very accurate. | ||||||
| h(km) | P(mbar) | T(°C) | r (g/cm 3) | N(cm –3) | H(km) | l(cm) |
| 0 | 1013 | 288 | 1.22 10 –3 | 2.55 10 19 | 8,4 | 7.4·10 –6 |
| 1 | 899 | 281 | 1.11·10 –3 | 2.31 10 19 | 8.1·10 –6 | |
| 2 | 795 | 275 | 1.01·10 –3 | 2.10 10 19 | 8.9·10 –6 | |
| 3 | 701 | 268 | 9.1·10 –4 | 1.89 10 19 | 9.9·10 –6 | |
| 4 | 616 | 262 | 8.2·10 –4 | 1.70 10 19 | 1.1·10 –5 | |
| 5 | 540 | 255 | 7.4·10 –4 | 1.53 10 19 | 7,7 | 1.2·10 –5 |
| 6 | 472 | 249 | 6.6·10 –4 | 1.37 10 19 | 1.4·10 –5 | |
| 8 | 356 | 236 | 5.2·10 -4 | 1.09 10 19 | 1.7·10 –5 | |
| 10 | 264 | 223 | 4.1·10 –4 | 8.6 10 18 | 6,6 | 2.2·10 –5 |
| 15 | 121 | 214 | 1.93·10 –4 | 4.0 10 18 | 4.6·10 –5 | |
| 20 | 56 | 214 | 8.9·10 –5 | 1.85 10 18 | 6,3 | 1.0·10 –4 |
| 30 | 12 | 225 | 1.9·10 –5 | 3.9 10 17 | 6,7 | 4.8·10 –4 |
| 40 | 2,9 | 268 | 3.9·10 –6 | 7.6 10 16 | 7,9 | 2.4·10 –3 |
| 50 | 0,97 | 276 | 1.15·10 –6 | 2.4 10 16 | 8,1 | 8.5·10 –3 |
| 60 | 0,28 | 260 | 3.9·10 –7 | 7.7 10 15 | 7,6 | 0,025 |
| 70 | 0,08 | 219 | 1.1·10 –7 | 2.5 10 15 | 6,5 | 0,09 |
| 80 | 0,014 | 205 | 2.7·10 –8 | 5.0 10 14 | 6,1 | 0,41 |
| 90 | 2.8·10 –3 | 210 | 5.0·10 –9 | 9·10 13 | 6,5 | 2,1 |
| 100 | 5.8·10 –4 | 230 | 8.8·10 –10 | 1.8 10 13 | 7,4 | 9 |
| 110 | 1.7·10 –4 | 260 | 2.1·10 –10 | 5.4 10 12 | 8,5 | 40 |
| 120 | 6·10 –5 | 300 | 5.6·10 –11 | 1.8 10 12 | 10,0 | 130 |
| 150 | 5·10 –6 | 450 | 3.2·10 –12 | 9 10 10 | 15 | 1.8 10 3 |
| 200 | 5·10 –7 | 700 | 1.6·10 –13 | 5 10 9 | 25 | 3 10 4 |
| 250 | 9·10 –8 | 800 | 3·10 –14 | 8 10 8 | 40 | 3·10 5 |
| 300 | 4·10 –8 | 900 | 8·10 –15 | 3 10 8 | 50 | |
| 400 | 8·10 –9 | 1000 | 1·10 –15 | 5 10 7 | 60 | |
| 500 | 2·10 –9 | 1000 | 2·10 –16 | 1·10 7 | 70 | |
| 700 | 2·10 –10 | 1000 | 2·10 –17 | 1 10 6 | 80 | |
| 1000 | 1·10 –11 | 1000 | 1·10 –18 | 1·10 5 | 80 | |
Troposphere.
The lowest and most dense layer of the atmosphere, in which the temperature decreases rapidly with height, is called the troposphere. It contains up to 80% of the total mass of the atmosphere and extends in the polar and middle latitudes to altitudes of 8–10 km, and in the tropics up to 16–18 km. Almost all weather-forming processes develop here, heat and moisture exchange occurs between the Earth and its atmosphere, clouds form, various meteorological phenomena occur, fog and precipitation occur. These layers of the earth's atmosphere are in convective equilibrium and, thanks to active mixing, have a homogeneous chemical composition, mainly consisting of molecular nitrogen (78%) and oxygen (21%). The vast majority of natural and man-made aerosol and gas air pollutants are concentrated in the troposphere. The dynamics of the lower part of the troposphere, up to 2 km thick, strongly depends on the properties of the underlying surface of the Earth, which determines the horizontal and vertical movements of air (winds) caused by the transfer of heat from warmer land through the infrared radiation of the earth's surface, which is absorbed in the troposphere, mainly by vapors water and carbon dioxide (greenhouse effect). The temperature distribution with height is established as a result of turbulent and convective mixing. On average, it corresponds to a temperature drop with height of approximately 6.5 K/km.
The wind speed in the surface boundary layer initially increases rapidly with height, and above it continues to increase by 2–3 km/s per kilometer. Sometimes narrow planetary flows (with a speed of more than 30 km/s) appear in the troposphere, western in the middle latitudes, and eastern near the equator. They are called jet streams.
Tropopause.
At the upper boundary of the troposphere (tropopause), the temperature reaches its minimum value for the lower atmosphere. This is the transition layer between the troposphere and the stratosphere located above it. The thickness of the tropopause ranges from hundreds of meters to 1.5–2 km, and the temperature and altitude, respectively, range from 190 to 220 K and from 8 to 18 km, depending on the latitude and season. In temperate and high latitudes in winter it is 1–2 km lower than in summer and 8–15 K warmer. In the tropics, seasonal changes are much less (altitude 16–18 km, temperature 180–200 K). Above jet streams tropopause breaks are possible.
Water in the Earth's atmosphere.
The most important feature of the Earth's atmosphere is the presence of significant amounts of water vapor and water in droplet form, which is most easily observed in the form of clouds and cloud structures. The degree of cloud coverage of the sky (at a certain moment or on average over a certain period of time), expressed on a scale of 10 or as a percentage, is called cloudiness. The shape of clouds is determined according to the international classification. On average, clouds cover about half of the globe. Cloudiness is an important factor characterizing weather and climate. In winter and at night, cloudiness prevents a decrease in the temperature of the earth's surface and the ground layer of air; in summer and during the day, it weakens the heating of the earth's surface by the sun's rays, softening the climate inside the continents.
Clouds.
Clouds are accumulations of water droplets suspended in the atmosphere (water clouds), ice crystals (ice clouds), or both together (mixed clouds). As droplets and crystals become larger, they fall out of the clouds in the form of precipitation. Clouds form mainly in the troposphere. They arise as a result of condensation of water vapor contained in the air. The diameter of cloud drops is on the order of several microns. The content of liquid water in clouds ranges from fractions to several grams per m3. Clouds are classified by height: According to the international classification, there are 10 types of clouds: cirrus, cirrocumulus, cirrostratus, altocumulus, altostratus, nimbostratus, stratus, stratocumulus, cumulonimbus, cumulus.
Pearlescent clouds are also observed in the stratosphere, and noctilucent clouds are observed in the mesosphere.
Cirrus clouds are transparent clouds in the form of thin white threads or veils with a silky sheen that do not provide shadows. Cirrus clouds are composed of ice crystals and form in the upper troposphere at very low temperatures. Some types of cirrus clouds serve as harbingers of weather changes.
Cirrocumulus clouds are ridges or layers of thin white clouds in the upper troposphere. Cirrocumulus clouds are built from small elements that look like flakes, ripples, small balls without shadows and consist mainly of ice crystals.
Cirrostratus clouds are a whitish translucent veil in the upper troposphere, usually fibrous, sometimes blurred, consisting of small needle-shaped or columnar ice crystals.
Altocumulus clouds are white, gray or white-gray clouds in the lower and middle layers of the troposphere. Altocumulus clouds have the appearance of layers and ridges, as if built from plates, rounded masses, shafts, flakes lying on top of each other. Altocumulus clouds form during intense convective activity and usually consist of supercooled water droplets.
Altostratus clouds are grayish or bluish clouds with a fibrous or uniform structure. Altostratus clouds are observed in the middle troposphere, extending several kilometers in height and sometimes thousands of kilometers in the horizontal direction. Typically, altostratus clouds are part of frontal cloud systems associated with upward movements of air masses.
Nimbostratus clouds are a low (from 2 km and above) amorphous layer of clouds of a uniform gray color, giving rise to continuous rain or snow. Nimbostratus clouds are highly developed vertically (up to several km) and horizontally (several thousand km), consist of supercooled water droplets mixed with snowflakes, usually associated with atmospheric fronts.
Stratus clouds are clouds of the lower tier in the form of a homogeneous layer without definite outlines, gray in color. The height of stratus clouds above the earth's surface is 0.5–2 km. Occasionally, drizzle falls from stratus clouds.
Cumulus clouds are dense, bright white clouds during the day with significant vertical development (up to 5 km or more). The upper parts of cumulus clouds look like domes or towers with rounded outlines. Typically, cumulus clouds arise as convection clouds in cold air masses.
Stratocumulus clouds are low (below 2 km) clouds in the form of gray or white non-fibrous layers or ridges of round large blocks. The vertical thickness of stratocumulus clouds is small. Occasionally, stratocumulus clouds produce light precipitation.
Cumulonimbus clouds are powerful and dense clouds with strong vertical development (up to a height of 14 km), producing heavy rainfall with thunderstorms, hail, and squalls. Cumulonimbus clouds develop from powerful cumulus clouds, differing from them in the upper part consisting of ice crystals.
Stratosphere.
Through the tropopause, on average at altitudes from 12 to 50 km, the troposphere passes into the stratosphere. In the lower part, for about 10 km, i.e. up to altitudes of about 20 km, it is isothermal (temperature about 220 K). It then increases with altitude, reaching a maximum of about 270 K at an altitude of 50–55 km. Here is the boundary between the stratosphere and the overlying mesosphere, called the stratopause. .
There is significantly less water vapor in the stratosphere. Still, thin translucent pearlescent clouds are sometimes observed, occasionally appearing in the stratosphere at an altitude of 20–30 km. Pearlescent clouds are visible in the dark sky after sunset and before sunrise. In shape, nacreous clouds resemble cirrus and cirrocumulus clouds.
Middle atmosphere (mesosphere).
At an altitude of about 50 km, the mesosphere begins from the peak of the broad temperature maximum . The reason for the increase in temperature in the region of this maximum is an exothermic (i.e. accompanied by the release of heat) photochemical reaction of ozone decomposition: O 3 + hv® O 2 + O. Ozone arises as a result of the photochemical decomposition of molecular oxygen O 2
O 2 + hv® O + O and the subsequent reaction of a triple collision of an oxygen atom and molecule with some third molecule M.
O + O 2 + M ® O 3 + M
Ozone voraciously absorbs ultraviolet radiation in the region from 2000 to 3000 Å, and this radiation heats the atmosphere. Ozone, located in the upper atmosphere, serves as a kind of shield that protects us from the effects of ultraviolet radiation from the Sun. Without this shield, the development of life on Earth in its modern forms would hardly have been possible.
In general, throughout the mesosphere, the atmospheric temperature decreases to its minimum value of about 180 K at the upper boundary of the mesosphere (called mesopause, altitude about 80 km). In the vicinity of the mesopause, at altitudes of 70–90 km, a very thin layer of ice crystals and particles of volcanic and meteorite dust may appear, observed in the form of a beautiful spectacle of noctilucent clouds shortly after sunset.
In the mesosphere, small solid meteorite particles that fall on the Earth, causing the phenomenon of meteors, mostly burn up.
Meteors, meteorites and fireballs.
Flares and other phenomena in the upper atmosphere of the Earth caused by the intrusion of solid cosmic particles or bodies into it at a speed of 11 km/s or higher are called meteoroids. An observable bright meteor trail appears; the most powerful phenomena, often accompanied by the fall of meteorites, are called fireballs; the appearance of meteors is associated with meteor showers.
Meteor shower:
1) the phenomenon of multiple falls of meteors over several hours or days from one radiant.
2) a swarm of meteoroids moving in the same orbit around the Sun.
The systematic appearance of meteors in a certain area of the sky and on certain days of the year, caused by the intersection of the Earth's orbit with the common orbit of many meteorite bodies moving at approximately the same and identically directed speeds, due to which their paths in the sky appear to emerge from a common point (radiant) . They are named after the constellation where the radiant is located.
Meteor showers make a deep impression with their light effects, but individual meteors are rarely visible. Much more numerous are invisible meteors, too small to be visible when they are absorbed into the atmosphere. Some of the smallest meteors probably do not heat up at all, but are only captured by the atmosphere. These small particles with sizes ranging from a few millimeters to ten thousandths of a millimeter are called micrometeorites. The amount of meteoric matter entering the atmosphere every day ranges from 100 to 10,000 tons, with the majority of this material coming from micrometeorites.
Since meteoric matter partially burns in the atmosphere, its gas composition is replenished with traces of various chemical elements. For example, rocky meteors introduce lithium into the atmosphere. The combustion of metal meteors leads to the formation of tiny spherical iron, iron-nickel and other droplets that pass through the atmosphere and settle on the earth's surface. They can be found in Greenland and Antarctica, where ice sheets remain almost unchanged for years. Oceanologists find them in bottom ocean sediments.
Most meteor particles entering the atmosphere settle within approximately 30 days. Some scientists believe that this cosmic dust plays an important role in the formation of atmospheric phenomena such as rain because it serves as condensation nuclei for water vapor. Therefore, it is assumed that precipitation is statistically related to large meteor showers. However, some experts believe that since the total supply of meteoric material is many tens of times greater than that of even the largest meteor shower, the change in the total amount of this material resulting from one such rain can be neglected.
However, there is no doubt that the largest micrometeorites and visible meteorites leave long traces of ionization in the high layers of the atmosphere, mainly in the ionosphere. Such traces can be used for long-distance radio communications, as they reflect high-frequency radio waves.
The energy of meteors entering the atmosphere is spent mainly, and perhaps completely, on heating it. This is one of the minor components of the thermal balance of the atmosphere.
A meteorite is a naturally occurring solid body that fell to the surface of the Earth from space. Usually a distinction is made between stony, stony-iron and iron meteorites. The latter mainly consist of iron and nickel. Among the meteorites found, most weigh from a few grams to several kilograms. The largest of those found, the Goba iron meteorite weighs about 60 tons and still lies in the same place where it was discovered, in South Africa. Most meteorites are fragments of asteroids, but some meteorites may have come to Earth from the Moon and even Mars.
A bolide is a very bright meteor, sometimes visible even during the day, often leaving behind a smoky trail and accompanied by sound phenomena; often ends with the fall of meteorites.
Thermosphere.
Above the temperature minimum of the mesopause, the thermosphere begins, in which the temperature, first slowly and then quickly begins to rise again. The reason is the absorption of ultraviolet radiation from the Sun at altitudes of 150–300 km, due to the ionization of atomic oxygen: O + hv® O + + e.
In the thermosphere, the temperature continuously increases to an altitude of about 400 km, where it reaches 1800 K during the day during the epoch of maximum solar activity. During the epoch of minimum solar activity, this limiting temperature can be less than 1000 K. Above 400 km, the atmosphere turns into an isothermal exosphere. The critical level (the base of the exosphere) is at an altitude of about 500 km.
Polar lights and many orbits of artificial satellites, as well as noctilucent clouds - all these phenomena occur in the mesosphere and thermosphere.
Polar lights.
At high latitudes, auroras are observed during magnetic field disturbances. They may last a few minutes, but are often visible for several hours. Auroras vary greatly in shape, color and intensity, all of which sometimes change very quickly over time. The spectrum of auroras consists of emission lines and bands. Some of the night sky emissions are enhanced in the aurora spectrum, primarily the green and red lines l 5577 Å and l 6300 Å oxygen. It happens that one of these lines is many times more intense than the other, and this determines the visible color of the aurora: green or red. Magnetic field disturbances are also accompanied by disruptions in radio communications in the polar regions. The cause of the disruption is changes in the ionosphere, which mean that during magnetic storms there is a powerful source of ionization. It has been established that strong magnetic storms occur when there are large groups of sunspots near the center of the solar disk. Observations have shown that storms are not associated with the sunspots themselves, but with solar flares that appear during the development of a group of sunspots.
Auroras are a range of light of varying intensity with rapid movements observed in high latitude regions of the Earth. The visual aurora contains green (5577Å) and red (6300/6364Å) atomic oxygen emission lines and molecular N2 bands, which are excited by energetic particles of solar and magnetospheric origin. These emissions usually appear at altitudes of about 100 km and above. The term optical aurora is used to refer to visual auroras and their emission spectrum from the infrared to the ultraviolet region. The radiation energy in the infrared part of the spectrum significantly exceeds the energy in the visible region. When auroras appeared, emissions were observed in the ULF range (
The actual forms of auroras are difficult to classify; The most commonly used terms are:
1. Calm, uniform arcs or stripes. The arc typically extends ~1000 km in the direction of the geomagnetic parallel (toward the Sun in polar regions) and has a width of one to several tens of kilometers. A stripe is a generalization of the concept of an arc; it usually does not have a regular arc-shaped shape, but bends in the form of the letter S or in the form of spirals. Arcs and stripes are located at altitudes of 100–150 km.
2. Rays of the aurora . This term refers to an auroral structure elongated along magnetic field lines, with a vertical extent of several tens to several hundred kilometers. The horizontal extent of the rays is small, from several tens of meters to several kilometers. The rays are usually observed in arcs or as separate structures.
3. Stains or surfaces . These are isolated areas of glow that do not have a specific shape. Individual spots may be connected to each other.
4. Veil. An unusual form of aurora, which is a uniform glow that covers large areas of the sky.
According to their structure, auroras are divided into homogeneous, hollow and radiant. Various terms are used; pulsating arc, pulsating surface, diffuse surface, radiant stripe, drapery, etc. There is a classification of auroras according to their color. According to this classification, auroras of the type A. The upper part or the entire part is red (6300–6364 Å). They usually appear at altitudes of 300–400 km with high geomagnetic activity.
Aurora type IN colored red in the lower part and associated with the glow of the bands of the first positive system N 2 and the first negative system O 2. Such forms of auroras appear during the most active phases of auroras.
Zones polar lights – These are the zones of maximum frequency of auroras at night, according to observers at a fixed point on the Earth's surface. The zones are located at 67° north and south latitude, and their width is about 6°. The maximum occurrence of auroras, corresponding to a given moment of geomagnetic local time, occurs in oval-like belts (oval auroras), which are located asymmetrically around the north and south geomagnetic poles. The aurora oval is fixed in latitude – time coordinates, and the aurora zone is the geometric locus of the points of the oval’s midnight region in latitude – longitude coordinates. The oval belt is located approximately 23° from the geomagnetic pole in the night sector and 15° in the daytime sector.
Aurora oval and aurora zones. The location of the aurora oval depends on geomagnetic activity. The oval becomes wider at high geomagnetic activity. Auroral zones or auroral oval boundaries are better represented by L 6.4 than by dipole coordinates. Geomagnetic field lines at the boundary of the daytime sector of the aurora oval coincide with magnetopause. A change in the position of the aurora oval is observed depending on the angle between the geomagnetic axis and the Earth-Sun direction. The auroral oval is also determined on the basis of data on precipitation of particles (electrons and protons) of certain energies. Its position can be independently determined from data on Kaspakh on the dayside and in the tail of the magnetosphere.
The daily variation in the frequency of occurrence of auroras in the aurora zone has a maximum at geomagnetic midnight and a minimum at geomagnetic noon. On the near-equatorial side of the oval, the frequency of occurrence of auroras sharply decreases, but the shape of the daily variations is preserved. On the polar side of the oval, the frequency of auroras decreases gradually and is characterized by complex diurnal changes.
Intensity of auroras.
Aurora intensity determined by measuring the apparent surface brightness. Luminosity surface I aurora in a certain direction is determined by the total emission of 4p I photon/(cm 2 s). Since this value is not the true surface brightness, but represents the emission from the column, the unit photon/(cm 2 column s) is usually used when studying auroras. The usual unit for measuring total emission is Rayleigh (Rl) equal to 10 6 photons/(cm 2 column s). More practical units of auroral intensity are determined by the emissions of an individual line or band. For example, the intensity of auroras is determined by the international brightness coefficients (IBRs) according to the intensity of the green line (5577 Å); 1 kRl = I MKY, 10 kRl = II MKY, 100 kRl = III MKY, 1000 kRl = IV MKY (maximum intensity of the aurora). This classification cannot be used for red auroras. One of the discoveries of the era (1957–1958) was the establishment of the spatiotemporal distribution of auroras in the form of an oval, shifted relative to the magnetic pole. From simple ideas about the circular shape of the distribution of auroras relative to the magnetic pole there was The transition to modern physics of the magnetosphere has been completed. The honor of the discovery belongs to O. Khorosheva, and the intensive development of ideas for the auroral oval was carried out by G. Starkov, Y. Feldstein, S. I. Akasofu and a number of other researchers. The auroral oval is the region of the most intense influence of the solar wind on the Earth's upper atmosphere. The intensity of the aurora is greatest in the oval, and its dynamics are continuously monitored using satellites.
Stable auroral red arcs.
Steady auroral red arc, otherwise called mid-latitude red arc or M-arc, is a subvisual (below the limit of sensitivity of the eye) wide arc, stretching from east to west for thousands of kilometers and possibly encircling the entire Earth. The latitudinal length of the arc is 600 km. The emission of the stable auroral red arc is almost monochromatic in the red lines l 6300 Å and l 6364 Å. Recently, weak emission lines l 5577 Å (OI) and l 4278 Å (N+2) were also reported. Sustained red arcs are classified as auroras, but they appear at much higher altitudes. The lower limit is located at an altitude of 300 km, the upper limit is about 700 km. The intensity of the quiet auroral red arc in the l 6300 Å emission ranges from 1 to 10 kRl (typical value 6 kRl). The sensitivity threshold of the eye at this wavelength is about 10 kRl, so arcs are rarely observed visually. However, observations have shown that their brightness is >50 kRL on 10% of nights. The usual lifespan of arcs is about one day, and they rarely appear in subsequent days. Radio waves from satellites or radio sources crossing persistent auroral red arcs are subject to scintillation, indicating the existence of electron density inhomogeneities. The theoretical explanation for red arcs is that the heated electrons of the region F The ionosphere causes an increase in oxygen atoms. Satellite observations show an increase in electron temperature along geomagnetic field lines that intersect persistent auroral red arcs. The intensity of these arcs is positively correlated with geomagnetic activity (storms), and the frequency of occurrence of arcs is positively correlated with sunspot activity.
Changing aurora.
Some forms of auroras experience quasi-periodic and coherent temporal variations in intensity. These auroras with approximately stationary geometry and rapid periodic variations occurring in phase are called changing auroras. They are classified as auroras forms R according to the International Atlas of Auroras A more detailed subdivision of the changing auroras:
R 1 (pulsating aurora) is a glow with uniform phase variations in brightness throughout the aurora shape. By definition, in an ideal pulsating aurora, the spatial and temporal parts of the pulsation can be separated, i.e. brightness I(r,t)= I s(r)· I T(t). In a typical aurora R 1 pulsations occur with a frequency from 0.01 to 10 Hz of low intensity (1–2 kRl). Most auroras R 1 – these are spots or arcs that pulsate with a period of several seconds.
R 2 (fiery aurora). The term is usually used to refer to movements like flames filling the sky, rather than to describe a distinct form. The auroras have the shape of arcs and usually move upward from a height of 100 km. These auroras are relatively rare and occur more often outside the aurora.
R 3 (shimmering aurora). These are auroras with rapid, irregular or regular variations in brightness, giving the impression of flickering flames in the sky. They appear shortly before the aurora disintegrates. Typically observed frequency of variation R 3 is equal to 10 ± 3 Hz.
The term streaming aurora, used for another class of pulsating auroras, refers to irregular variations in brightness moving quickly horizontally in auroral arcs and streaks.
The changing aurora is one of the solar-terrestrial phenomena that accompany pulsations of the geomagnetic field and auroral X-ray radiation caused by the precipitation of particles of solar and magnetospheric origin.
The glow of the polar cap is characterized by high intensity of the band of the first negative system N + 2 (l 3914 Å). Typically, these N + 2 bands are five times more intense than the green line OI l 5577 Å; the absolute intensity of the polar cap glow ranges from 0.1 to 10 kRl (usually 1–3 kRl). During these auroras, which appear during periods of PCA, a uniform glow covers the entire polar cap up to a geomagnetic latitude of 60° at altitudes of 30 to 80 km. It is generated predominantly by solar protons and d-particles with energies of 10–100 MeV, creating a maximum ionization at these altitudes. There is another type of glow in aurora zones, called mantle aurora. For this type of auroral glow, the daily maximum intensity, occurring in the morning hours, is 1–10 kRL, and the minimum intensity is five times weaker. Observations of mantle auroras are few and far between; their intensity depends on geomagnetic and solar activity.
Atmospheric glow is defined as radiation produced and emitted by a planet's atmosphere. This is non-thermal radiation of the atmosphere, with the exception of the emission of auroras, lightning discharges and the emission of meteor trails. This term is used in relation to the earth's atmosphere (nightglow, twilight glow and dayglow). Atmospheric glow constitutes only a portion of the light available in the atmosphere. Other sources include starlight, zodiacal light, and daytime diffuse light from the Sun. At times, atmospheric glow can account for up to 40% of the total amount of light. Atmospheric glow occurs in atmospheric layers of varying height and thickness. The atmospheric glow spectrum covers wavelengths from 1000 Å to 22.5 microns. The main emission line in the atmospheric glow is l 5577 Å, appearing at an altitude of 90–100 km in a layer 30–40 km thick. The appearance of luminescence is due to the Chapman mechanism, based on the recombination of oxygen atoms. Other emission lines are l 6300 Å, appearing in the case of dissociative recombination of O + 2 and emission NI l 5198/5201 Å and NI l 5890/5896 Å.
The intensity of airglow is measured in Rayleigh. Brightness (in Rayleigh) is equal to 4 rv, where b is the angular surface brightness of the emitting layer in units of 10 6 photons/(cm 2 ster·s). The intensity of the glow depends on latitude (different for different emissions), and also varies throughout the day with a maximum near midnight. A positive correlation was noted for airglow in the l 5577 Å emission with the number of sunspots and solar radiation flux at a wavelength of 10.7 cm. Airglow is observed during satellite experiments. From outer space, it appears as a ring of light around the Earth and has a greenish color.
Ozonosphere.
At altitudes of 20–25 km, the maximum concentration of an insignificant amount of ozone O 3 is reached (up to 2×10 –7 of the oxygen content!), which arises under the influence of solar ultraviolet radiation at altitudes of approximately 10 to 50 km, protecting the planet from ionizing solar radiation. Despite the extremely small number of ozone molecules, they protect all life on Earth from the harmful effects of short-wave (ultraviolet and x-ray) radiation from the Sun. If you deposit all the molecules to the base of the atmosphere, you will get a layer no more than 3–4 mm thick! At altitudes above 100 km, the proportion of light gases increases, and at very high altitudes helium and hydrogen predominate; many molecules dissociate into individual atoms, which, ionized under the influence of hard radiation from the Sun, form the ionosphere. The pressure and density of air in the Earth's atmosphere decrease with altitude. Depending on the temperature distribution, the Earth's atmosphere is divided into the troposphere, stratosphere, mesosphere, thermosphere and exosphere. .
At an altitude of 20–25 km there is ozone layer. Ozone is formed due to the breakdown of oxygen molecules when absorbing ultraviolet radiation from the Sun with wavelengths shorter than 0.1–0.2 microns. Free oxygen combines with O 2 molecules and forms ozone O 3, which greedily absorbs all ultraviolet radiation shorter than 0.29 microns. O3 ozone molecules are easily destroyed by short-wave radiation. Therefore, despite its rarefaction, the ozone layer effectively absorbs ultraviolet radiation from the Sun that has passed through higher and more transparent atmospheric layers. Thanks to this, living organisms on Earth are protected from the harmful effects of ultraviolet light from the Sun.
Ionosphere.
Radiation from the sun ionizes the atoms and molecules of the atmosphere. The degree of ionization becomes significant already at an altitude of 60 kilometers and steadily increases with distance from the Earth. At different altitudes in the atmosphere, sequential processes of dissociation of various molecules and subsequent ionization of various atoms and ions occur. These are mainly molecules of oxygen O 2, nitrogen N 2 and their atoms. Depending on the intensity of these processes, the various layers of the atmosphere lying above 60 kilometers are called ionospheric layers , and their totality is the ionosphere . The lower layer, the ionization of which is insignificant, is called the neutrosphere.
The maximum concentration of charged particles in the ionosphere is achieved at altitudes of 300–400 km.
History of the study of the ionosphere.
The hypothesis about the existence of a conducting layer in the upper atmosphere was put forward in 1878 by the English scientist Stuart to explain the features of the geomagnetic field. Then in 1902, independently of each other, Kennedy in the USA and Heaviside in England pointed out that to explain the propagation of radio waves over long distances it was necessary to assume the existence of regions of high conductivity in the high layers of the atmosphere. In 1923, academician M.V. Shuleikin, considering the features of the propagation of radio waves of various frequencies, came to the conclusion that there are at least two reflective layers in the ionosphere. Then in 1925, English researchers Appleton and Barnett, as well as Breit and Tuve, first experimentally proved the existence of regions that reflect radio waves, and laid the foundation for their systematic study. Since that time, a systematic study has been carried out of the properties of these layers, generally called the ionosphere, which play a significant role in a number of geophysical phenomena that determine the reflection and absorption of radio waves, which is very important for practical purposes, in particular for ensuring reliable radio communications.
In the 1930s, systematic observations of the state of the ionosphere began. In our country, on the initiative of M.A. Bonch-Bruevich, installations for its pulse probing were created. Many general properties of the ionosphere, heights and electron concentration of its main layers were studied.
At altitudes of 60–70 km layer D is observed, at altitudes of 100–120 km layer E, at altitudes, at altitudes of 180–300 km double layer F 1 and F 2. The main parameters of these layers are given in Table 4.
| Table 4. | ||||||
| Ionospheric region | Maximum height, km | T i , K | Day | Night n e , cm –3 | a΄, ρm 3 s – 1 | |
| min n e , cm –3 | Max n e , cm –3 | |||||
| D | 70 | 20 | 100 | 200 | 10 | 10 –6 |
| E | 110 | 270 | 1.5 10 5 | 3·10 5 | 3000 | 10 –7 |
| F 1 | 180 | 800–1500 | 3·10 5 | 5 10 5 | – | 3·10 –8 |
| F 2 (winter) | 220–280 | 1000–2000 | 6 10 5 | 25 10 5 | ~10 5 | 2·10 –10 |
| F 2 (summer) | 250–320 | 1000–2000 | 2·10 5 | 8 10 5 | ~3·10 5 | 10 –10 |
| n e– electron concentration, e – electron charge, T i– ion temperature, a΄ – recombination coefficient (which determines the value n e and its change over time) | ||||||
Average values are given because they vary at different latitudes, depending on the time of day and seasons. Such data is necessary to ensure long-distance radio communications. They are used in selecting operating frequencies for various shortwave radio links. Knowledge of their changes depending on the state of the ionosphere at different times of the day and in different seasons is extremely important to ensure the reliability of radio communications. The ionosphere is a collection of ionized layers of the earth's atmosphere, starting from altitudes of about 60 km and extending to altitudes of tens of thousands of km. The main source of ionization of the Earth's atmosphere is ultraviolet and X-ray radiation from the Sun, which occurs mainly in the solar chromosphere and corona. In addition, the degree of ionization of the upper atmosphere is influenced by solar corpuscular streams that occur during solar flares, as well as cosmic rays and meteor particles.
Ionospheric layers
- these are areas in the atmosphere in which maximum concentrations of free electrons are reached (i.e., their number per unit volume). Electrically charged free electrons and (to a lesser extent, less mobile ions) resulting from the ionization of atoms of atmospheric gases, interacting with radio waves (i.e., electromagnetic oscillations), can change their direction, reflecting or refracting them, and absorb their energy. As a result of this, when receiving distant radio stations, various effects may occur, for example, fading of radio communications, increased audibility of remote stations, blackouts and so on. phenomena.
Research methods.
Classical methods of studying the ionosphere from Earth come down to pulse sounding - sending radio pulses and observing their reflections from various layers of the ionosphere, measuring the delay time and studying the intensity and shape of the reflected signals. By measuring the heights of reflection of radio pulses at various frequencies, determining the critical frequencies of various areas (the critical frequency is the carrier frequency of a radio pulse, for which a given region of the ionosphere becomes transparent), it is possible to determine the value of the electron concentration in the layers and the effective heights for given frequencies, and select the optimal frequencies for given radio paths. With the development of rocket technology and the advent of the space age of artificial Earth satellites (AES) and other spacecraft, it became possible to directly measure the parameters of near-Earth space plasma, the lower part of which is the ionosphere.
Measurements of electron concentration, carried out on board specially launched rockets and along satellite flight paths, confirmed and clarified data previously obtained by ground-based methods on the structure of the ionosphere, the distribution of electron concentration with height above various regions of the Earth and made it possible to obtain electron concentration values above the main maximum - the layer F. Previously, this was impossible to do using sounding methods based on observations of reflected short-wave radio pulses. It has been discovered that in some areas of the globe there are quite stable areas with a reduced electron concentration, regular “ionospheric winds”, peculiar wave processes arise in the ionosphere that carry local ionospheric disturbances thousands of kilometers from the place of their excitation, and much more. The creation of particularly highly sensitive receiving devices made it possible to receive pulse signals partially reflected from the lowest regions of the ionosphere (partial reflection stations) at ionospheric pulse sounding stations. The use of powerful pulsed installations in the meter and decimeter wavelength ranges with the use of antennas that allow for a high concentration of emitted energy made it possible to observe signals scattered by the ionosphere at various altitudes. The study of the features of the spectra of these signals, incoherently scattered by electrons and ions of the ionospheric plasma (for this, stations of incoherent scattering of radio waves were used) made it possible to determine the concentration of electrons and ions, their equivalent temperature at various altitudes up to altitudes of several thousand kilometers. It turned out that the ionosphere is quite transparent for the frequencies used.
The concentration of electric charges (the electron concentration is equal to the ion concentration) in the earth's ionosphere at an altitude of 300 km is about 10 6 cm –3 during the day. Plasma of such density reflects radio waves with a length of more than 20 m, and transmits shorter ones.
Typical vertical distribution of electron concentration in the ionosphere for day and night conditions.
Propagation of radio waves in the ionosphere.
Stable reception of long-distance broadcasting stations depends on the frequencies used, as well as on the time of day, season and, in addition, on solar activity. Solar activity significantly affects the state of the ionosphere. Radio waves emitted by a ground station travel in a straight line, like all types of electromagnetic waves. However, it should be taken into account that both the surface of the Earth and the ionized layers of its atmosphere serve as the plates of a huge capacitor, acting on them like the effect of mirrors on light. Reflecting from them, radio waves can travel many thousands of kilometers, circling the globe in huge leaps of hundreds and thousands of kilometers, reflecting alternately from a layer of ionized gas and from the surface of the Earth or water.
In the 20s of the last century, it was believed that radio waves shorter than 200 m were generally not suitable for long-distance communications due to strong absorption. The first experiments on long-distance reception of short waves across the Atlantic between Europe and America were carried out by English physicist Oliver Heaviside and American electrical engineer Arthur Kennelly. Independently of each other, they suggested that somewhere around the Earth there is an ionized layer of the atmosphere capable of reflecting radio waves. It was called the Heaviside-Kennelly layer, and then the ionosphere.
According to modern concepts, the ionosphere consists of negatively charged free electrons and positively charged ions, mainly molecular oxygen O + and nitric oxide NO +. Ions and electrons are formed as a result of the dissociation of molecules and ionization of neutral gas atoms by solar X-rays and ultraviolet radiation. In order to ionize an atom, it is necessary to impart ionization energy to it, the main source of which for the ionosphere is ultraviolet, x-ray and corpuscular radiation from the Sun.
While the gaseous shell of the Earth is illuminated by the Sun, more and more electrons are continuously formed in it, but at the same time some of the electrons, colliding with ions, recombine, again forming neutral particles. After sunset, the formation of new electrons almost stops, and the number of free electrons begins to decrease. The more free electrons there are in the ionosphere, the better high-frequency waves are reflected from it. With a decrease in electron concentration, the passage of radio waves is possible only in low frequency ranges. That is why at night, as a rule, it is possible to receive distant stations only in the ranges of 75, 49, 41 and 31 m. Electrons are distributed unevenly in the ionosphere. At altitudes from 50 to 400 km there are several layers or regions of increased electron concentration. These areas smoothly transition into one another and have different effects on the propagation of HF radio waves. The upper layer of the ionosphere is designated by the letter F. Here the highest degree of ionization (the fraction of charged particles is about 10 –4). It is located at an altitude of more than 150 km above the Earth's surface and plays the main reflective role in the long-distance propagation of high-frequency HF radio waves. In the summer months, region F splits into two layers - F 1 and F 2. Layer F1 can occupy heights from 200 to 250 km, and layer F 2 seems to “float” in the altitude range of 300–400 km. Usually layer F 2 is ionized much stronger than the layer F 1 . Night layer F 1 disappears and the layer F 2 remains, slowly losing up to 60% of its degree of ionization. Below layer F at altitudes from 90 to 150 km there is a layer E ionization of which occurs under the influence of soft X-ray radiation from the Sun. The degree of ionization of the E layer is lower than that of the F, during the day, reception of stations in the low-frequency HF ranges of 31 and 25 m occurs when signals are reflected from the layer E. Typically these are stations located at a distance of 1000–1500 km. At night in the layer E Ionization decreases sharply, but even at this time it continues to play a significant role in the reception of signals from stations on the 41, 49 and 75 m ranges.
Of great interest for receiving signals of high-frequency HF ranges of 16, 13 and 11 m are those arising in the area E layers (clouds) of highly increased ionization. The area of these clouds can vary from a few to hundreds of square kilometers. This layer of increased ionization is called the sporadic layer E and is designated Es. Es clouds can move in the ionosphere under the influence of wind and reach speeds of up to 250 km/h. In summer in mid-latitudes during the daytime, the origin of radio waves due to Es clouds occurs for 15–20 days per month. Near the equator it is almost always present, and in high latitudes it usually appears at night. Sometimes, during years of low solar activity, when there is no transmission on the high-frequency HF bands, distant stations suddenly appear on the 16, 13 and 11 m bands with good volume, the signals of which are reflected many times from Es.
The lowest region of the ionosphere is the region D located at altitudes between 50 and 90 km. There are relatively few free electrons here. From the area D Long and medium waves are well reflected, and signals from low-frequency HF stations are strongly absorbed. After sunset, ionization disappears very quickly and it becomes possible to receive distant stations in the ranges of 41, 49 and 75 m, the signals of which are reflected from the layers F 2 and E. Individual layers of the ionosphere play an important role in the propagation of HF radio signals. The effect on radio waves occurs mainly due to the presence of free electrons in the ionosphere, although the mechanism of radio wave propagation is associated with the presence of large ions. The latter are also of interest when studying the chemical properties of the atmosphere, since they are more active than neutral atoms and molecules. Chemical reactions occurring in the ionosphere play an important role in its energy and electrical balance.
Normal ionosphere. Observations made using geophysical rockets and satellites have provided a wealth of new information indicating that ionization of the atmosphere occurs under the influence of a wide range of solar radiation. Its main part (more than 90%) is concentrated in the visible part of the spectrum. Ultraviolet radiation, which has a shorter wavelength and higher energy than violet light rays, is emitted by hydrogen in the Sun's inner atmosphere (the chromosphere), and X-rays, which have even higher energy, are emitted by gases in the Sun's outer shell (the corona).
The normal (average) state of the ionosphere is due to constant powerful radiation. Regular changes occur in the normal ionosphere due to the daily rotation of the Earth and seasonal differences in the angle of incidence of the sun's rays at noon, but unpredictable and abrupt changes in the state of the ionosphere also occur.
Disturbances in the ionosphere.
As is known, powerful cyclically repeating manifestations of activity occur on the Sun, which reach a maximum every 11 years. Observations under the International Geophysical Year (IGY) program coincided with the period of the highest solar activity for the entire period of systematic meteorological observations, i.e. from the beginning of the 18th century. During periods of high activity, the brightness of some areas on the Sun increases several times, and the power of ultraviolet and X-ray radiation increases sharply. Such phenomena are called solar flares. They last from several minutes to one to two hours. During the flare, solar plasma (mostly protons and electrons) is erupted, and elementary particles rush into outer space. Electromagnetic and corpuscular radiation from the Sun during such flares has a strong impact on the Earth's atmosphere.
The initial reaction is observed 8 minutes after the flare, when intense ultraviolet and X-ray radiation reaches the Earth. As a result, ionization increases sharply; X-rays penetrate the atmosphere to the lower boundary of the ionosphere; the number of electrons in these layers increases so much that the radio signals are almost completely absorbed (“extinguished”). The additional absorption of radiation causes the gas to heat up, which contributes to the development of winds. Ionized gas is an electrical conductor, and when it moves in the Earth's magnetic field, a dynamo effect occurs and an electric current is created. Such currents can, in turn, cause noticeable disturbances in the magnetic field and manifest themselves in the form of magnetic storms.
The structure and dynamics of the upper atmosphere are significantly determined by non-equilibrium processes in the thermodynamic sense associated with ionization and dissociation by solar radiation, chemical processes, excitation of molecules and atoms, their deactivation, collisions and other elementary processes. In this case, the degree of nonequilibrium increases with height as the density decreases. Up to altitudes of 500–1000 km, and often higher, the degree of nonequilibrium for many characteristics of the upper atmosphere is quite small, which makes it possible to use classical and hydromagnetic hydrodynamics, taking into account chemical reactions, to describe it.
The exosphere is the outer layer of the Earth's atmosphere, starting at altitudes of several hundred kilometers, from which light, fast-moving hydrogen atoms can escape into outer space.
Edward Kononovich
Literature:
Pudovkin M.I. Fundamentals of Solar Physics. St. Petersburg, 2001
Eris Chaisson, Steve McMillan Astronomy today. Prentice-Hall, Inc. Upper Saddle River, 2002
Materials on the Internet: http://ciencia.nasa.gov/
