What is true about the region of Earths atmosphere known as the stratosphere?

1.1.2 Stratosphere

The stratosphere contains ~9.9% of air mass over the Earth, and ranges from ~10 to ~50 km ASL with ascending temperature up to ~270 K. Due to precipitation in the troposphere, H2O can scarcely survive through vertical transport to reach the stratosphere. In the stratosphere, O2 may be photolyzed by solar ultraviolet radiation to form ozone (O3), which results in the so-called “O3 layer”. The O3 layer itself absorbs solar ultraviolet radiation with a little longer wavelength, to close the Chapman cycle of O3 formation in the natural stratosphere. As a result, solar UV radiation at the top of the stratosphere is much stronger than at the bottom of the stratosphere for wavelengths less than ~300 nm. Thus, humans and animals are effectively protected by the O3 layer from harmful, solar UV radiation with wavelengths shorter than ~300 nm.

In the modern atmosphere, chemicals such as N2O and chlorofluorocarbons (CFCs), which decompose slowly in the troposphere, may accumulate to a significant amount and enter the stratosphere via stratosphere–troposphere exchange events. Due to strong solar radiation in the stratosphere, these chemicals photolyze to form NO and halogen radicals, which then perturb the Chapman cycle to affect the thickness of the O3 layer. The most important observation related to the O3 layer in the stratosphere is the so-called “O3 hole”, initially observed over Antarctica during early spring in the early 1980s. Table 1.3 lists typical seasonal column O3 over the Equator, and global distributions of column O3 are presented in Chapter 5; ~90% of the column O3 stays in the stratosphere.

Table 1.3. Column O3 in Dobson units over the Equator

MonthO3MonthO3Jan251July265Feb252Aug269Mar257Sept271Apr260Oct265May260Nov258June262Dec253

Due to strong photochemical reactions in the stratosphere, mixing ratios of surface-originated compounds, such as CH4, N2O, and CFCs, are lower than in the troposphere. Meanwhile, their reaction products, such as NO, OH, and halogen radicals, may be more abundant in the stratosphere than part of the troposphere.

1 Introduction

The stratosphere is the layer of highly stratified air that extends for roughly 40 km above the tropopause and contains approximately 20% of the mass of the atmosphere. The climatology, seasonal evolution, and variability of the stratospheric circulation are strongly governed by the combined influences of solar and infrared radiation, ozone chemistry, and transport and momentum transport by Rossby and gravity waves that propagate upward from the troposphere below. While it contains a smaller fraction of atmospheric mass than the troposphere, the stratosphere is far from being a passive bystander to tropospheric influences. It exhibits a diverse range of variability on a spectrum of timescales with, in many cases, a well-established influence on the tropospheric circulation below. As a result, knowledge of the state of the stratosphere has the potential to enhance the predictability of the troposphere on sub-seasonal to seasonal (S2S) timescales and beyond.

This chapter reviews our knowledge of the coupling between the stratosphere and troposphere in the tropics (Section 2) and the extratropics (Section 3) to provide a clear understanding of where and when coupling is important. In Section 4, we review the progress to date in trying to harness stratosphere-troposphere coupling to enhance predictability on the S2S timescale, a key focus of the World Climate Research Programme/Stratosphere-Troposphere Processes And Their Role in Climate (WCRP/SPARC) Stratospheric Network for the Assessment of Predictability (SNAP) project. Finally, in Section 5, we examine a number of open questions and provide some perspective on where and how improved understanding and simulation of stratosphere-troposphere coupling are most likely to lead to improved skill. Throughout the chapter, it is important to emphasize that one of the significant difficulties in assessing and understanding stratosphere-troposphere coupling (in common with other low-frequency phenomena, such as Deser et al., 2017) is the relatively short observational record that exists for the stratosphere.

Fig. 1 highlights the major phenomena relevant to coupling between the stratosphere and troposphere, including the Quasi-Biennial Oscillation (QBO), solar variability, ozone, and the role of tropospheric planetary-scale waves.

The Subtropics

Stratosphere–troposphere exchange in this region occurs between the upper and mid-equatorial troposphere and the lowermost stratosphere. The subtropical tropopause drops rapidly near 30° from tropical heights to the level of the extratropical tropopause (approximately from 100 hPa to 300 hPa) (Figure 1). Trajectories from analyzed winds suggest very little STE occurs across this portion of the tropopause during the winter months but that considerable STE occurs during the summer months.

The subtropical tropopause cuts through the subtropical jet stream. This jet undergoes a substantial annual cycle in amplitude with the strongest winds occurring during the winter season. When the jet is strong, mixing across it between the troposphere and the stratosphere is inhibited; inhibited both through the large potential vorticity gradients associated with the jet, and the fact that breaking of tropospheric waves and the resulting mixing is unlikely to penetrate the jet core. Indeed, in the case of a strong jet the wind speeds are substantially larger than those associated with most tropospheric waves, implying that the critical layers (where the phase speed of the wave is equal to that of the large-scale flow field and therefore where the wave is unstable and breaks) will occur away from the jet core. During the summer months the subtropical jet weakens considerably, allowing mixing across the jet to be enhanced. Not only do the critical layers occur closer to the jet core during the summer months, but the smaller gradients of potential vorticity associated with the summer jet make for wider critical layers and weaker barriers to mixing.

The transport across the summertime subtropical jet is primarily associated with the Asian monsoon (Figure 2), and to a lesser extent the Mexican monsoon. While the monsoons of South America, Africa, and Australia probably play a similar role during the austral summer, their comparatively weak circulations are much less effective in transporting air across the tropopause. As indicated by the arrows in Figure 2, monsoon circulations are able to tap a particularly rich source of water vapor in the midlatitudes. The resulting STE is believed to be of primary importance to the seasonal cycle of water vapor in the extratropical lowermost stratosphere and does not involve the pronounced dehydration that occurs in the tropics.

The tropopause is elevated over monsoon regions with the associated anticyclonic circulation penetrating into the lowermost stratosphere. A steady state monsoon circulation will not in itself result in STE. However, due to the proximity of the monsoon circulation to the jet core, perturbations in the circulation are likely to be important, resulting in isentropic mixing between the troposphere and the stratosphere (Figure 2) and associated STE. Moreover, the interaction between monsoon and midlatitude synoptic disturbances or large-scale low-frequency transients will act to transport species across the tropopause. It has been demonstrated in the case of the Asian monsoon that the interaction can act to pull filaments (see the following section on extratropical STE) of moist tropospheric air into the stratosphere, and filaments of dry stratospheric air into the troposphere.

Stratospheric Waves

The stratosphere is normally about twice as stable as the troposphere (in terms of its Brunt–Väisälä frequency, N) and permits the vertical propagation of the high-frequency waves generated in the troposphere by convection. These highest-frequency waves have phase lines that are oriented closer to the vertical than lower-frequency waves and the vertical component of their group velocity can be similar in magnitude to the horizontal component. Consequently, these waves often occur in the stratospheric air directly above convective systems. Nonetheless, as the waves propagate horizontally as well, the total horizontal area affected by high-frequency gravity waves eventually exceeds the horizontal area of the convective system that generated them.

Figure 5 illustrates the pattern of waves in the stratosphere above a numerically simulated isolated midlatitude convective system. (The tropopause is at approximately 12 km altitude in this case). The wave fronts are approximately circular, with the center of the fronts corresponding to the convective system itself. As is typical, the horizontal wavelength of the highest-frequency waves is smaller than the width of the convective system and corresponds to the size of individual convective updrafts. Gravity waves with horizontal wavelengths as small as 5 km have been observed directly above active convection by aircraft, consistent with those shown here. Organized MCSs comprises multiple regions of active convection and often show separate sets of circular wave fronts; each set of wave fronts is centered on a different region of active convection within the system. The result can be complex interference patterns in the wave field in the stratosphere, with additional complexity arising from source intermittency as convective regions become active and decay with time.

Convective clouds generate a spectrum of waves and lower-frequency waves that are not directly attributable to overshooting updrafts that also reach the stratosphere. For example, n = 1 and n = 2 waves are only partially reflected at the tropopause and propagate into the stratosphere as well. Typical wave periods above isolated convection can range from 10 min to a few hours, horizontal wavelengths range from ∼5 to >50 km, and there is a corresponding range of horizontal phase speeds. For example, Figure 6 shows an example spectrum of the vertical flux of horizontal momentum (ρuw, where ρ is the density and u and w are the perturbation horizontal and vertical velocities, respectively) at 20 km altitude above a numerically simulated midlatitude squall line. For upward propagating waves the sign of the momentum flux is equal to the sign of the phase speed relative to the mean flow, and the fluxes in this figure are consistent with upward propagating waves in an environment moving at approximately 10 m s−1. In this example, peaks in the spectrum exist at phase speeds between approximately 10 and 30 m s−1 in magnitude relative to the 10 m s−1 background flow.

Although the examples presented earlier in Figures 1 and 2 showed symmetric wave patterns, the waves shown in Figures 5 and 6 show that the waves and their spectrum are different for the different propagation directions. This asymmetry occurs for two main reasons: wave generation and wave propagation.

Long-lived convective systems, like squall lines, usually feature convective updrafts that are tilted in the vertical. Studies have shown that although tilted updrafts generate gravity waves propagating in all horizontal directions, the largest amplitude waves generated are normally those that propagate in the direction of the updraft tilt. The tilted structures project onto that part of the wave spectrum more efficiently than other parts, generating an asymmetric spectrum of gravity waves.

The spectrum of waves at a given altitude in the stratosphere is also shaped by changes in wind shear and stability between that altitude and the convective wave source, i.e., wave filtering. Filtering by wind shear can create significant asymmetries in the stratospheric wave spectrum, which is closer to symmetric in the absence of shear. Moreover, as illustrated by Figure 5, the shortest gravity waves are often filtered in the upper troposphere and lower stratosphere, leaving only the longer wavelength waves further aloft.

The motion of convective systems, viz. advection and propagation, complicates interpretation of the dynamics underlying the wave generation. Analyzing the wave spectrum in a frame of reference moving with the convective source (as opposed to the ground-based reference frame, e.g., Figure 6) is sometimes illustrative. However, choosing this reference frame is complicated by the fact that individual convective cells often propagate at a different velocity to the MCSs they are embedded within. Some studies have shown, however, that when the convective updrafts move with the mean wind at their equilibrium level, the peaks in the spectrum occur at similar phase speeds in all directions relative to that reference frame. Sheared flows with an ensemble of cloud top heights can broaden the spectrum considerably because the ensemble of wave sources moves at different speeds.

Organized convective systems, viz. systems larger than a single convective cell that persist for multiple convective overturn times, generate gravity waves with horizontal wavelengths beyond the scale of individual clouds. As convective clouds become organized, the spectrum of gravity waves broadens to encompass longer wavelengths and longer periods that are related to the spatial scale of the organized system and its lifetime. For example, regions of organized convective activity in the tropics can contain a rich spectrum of multiscale convective systems and may include individual clouds, MCSs, mesoscale convective complexes, clusters, and superclusters. Each of these systems generates gravity waves leading to a spectrum that contains notable contributions from horizontal wavelengths of 10 km through to scales beyond 1000 km. Waves with these longer wavelengths can be influenced by the Earth's rotation and are then classified as ‘inertia-gravity waves.’ The intrinsic period of inertia-gravity waves can be tens of hours. (The minimum intrinsic frequency of inertia-gravity waves is the local inertial frequency). Accordingly, the phase lines of inertia-gravity waves are oriented close to the horizontal and can propagate significant horizontal distances from their source. Indeed, inertia-gravity waves observed in the lower stratosphere have been traced back to convection more than 1000 km away.

Tropical cyclones and hurricanes are also notable sources of gravity waves. Patterns of vertical velocity above these systems show a combination of short-scale gravity waves from individual thunderstorms and long-scale (∼100 km horizontal wavelength) waves that are associated with the evolving eye wall and spiral rainbands. Numerical simulations have found coherent spiral wave fronts above the convectively active cyclone rainbands. Notable variations in wave amplitude from these separate sources have also been linked to different stages of the cyclone intensification and decay cycle. The rotation of the wind around cyclones affects the wave propagation as well; the subsequent wave filtering, combined with different sectors of the storm featuring more active convection than others, creates considerable spatial variability in the gravity wave activity above cyclones.

Control of the Mean Distribution and Variability of Stratospheric Water Vapor

It is one of the fascinations of the study of stratospheric humidity that, while this extreme aridity and the overall mechanisms causing it have been known for more than half a century, the detailed understanding of precisely how this state is maintained remains elusive. This basic theory of the humidity of the stratosphere has survived and today forms the basis of our understanding. However, we now realize that there are many important details that modify this model.

The key mechanisms that contribute to the control of stratospheric water vapor concentrations are stratospheric photochemistry and transport, tropical tropopause dehydration, and polar dehydration. Additionally, transport from troposphere to stratosphere at midlatitudes also appears to impact stratospheric water vapor.

Introduction

The troposphere and the stratosphere are separated by a boundary called the tropopause, whose altitude varies from about 16 km in the tropics to about 8 km near the poles. The troposphere is characterized by rapid vertical transport and mixing caused by weather disturbances; the stratosphere is characterized by very weak vertical transport and mixing. The tropopause thus represents a boundary between the troposphere, where chemical constituents tend to be well mixed; and the stratosphere, where chemical constituents tend to have strong vertical gradients. The two-way exchange of material that occurs across the tropopause is important for determining the climate and chemical composition of the upper troposphere and the lower stratosphere. This cross-tropopause transport is referred to as stratosphere–troposphere exchange. The upward transport of tropospheric constituents into the stratosphere occurs primarily in the tropics, and initiates much of the chemistry that is responsible for global ozone depletion. The downward transport of stratospheric constituents into the troposphere occurs mostly in the extratropics and not only serves as the major sink for some of the constituents involved in stratospheric ozone depletion, but also provides a source of upper tropospheric ozone.

This pattern of upward cross-tropopause transport in the tropics and downward cross-tropopause transport in the extratropics is part of a global mass circulation in the stratosphere that occurs as an indirect response to zonal (westward) forcing in the stratosphere, which is caused by the breaking of large-scale waves propagating from the troposphere. The magnitude and variability of this stratospheric mass circulation, and its consequences for atmospheric chemistry, are primary considerations in the study of stratosphere–troposphere exchange.

Stratospheric Sulfur Compounds

Sulfate aerosols in the stratosphere are important to the albedo of the Earth (Warneck 2000). A layer of sulfate aerosols, known as the Junge layer, is found in the stratosphere at about 20 to 25 km altitude. Its origin is twofold. Large volcanic eruptions can inject SO2 into the stratosphere, where it is oxidized to sulfate (Eqs. 3.27 and 3.28). Large eruptions have the potential to increase the abundance of stratospheric sulfate 100-fold (Arnold and Bührke 1983, Hofmann and Rosen 1983), and the sulfate aerosols persist in the stratosphere for several years, cooling the planet (McCormick et al. 1995, Briffa et al. 1998). During periods without volcanic activity, the dominant source of stratospheric sulfate derives from carbonyl sulfide (COS)8 that mixes up from the troposphere, where it originates from a variety of sources (Chapter 13). Most sulfur gases are so reactive that they do not reach the stratosphere, but COS has a mean residence time of about 5 years in the atmosphere (refer to Table 3.5), so about one-third of the annual production mixes to the stratosphere. Additional COS may be lofted to the stratosphere in the smoke plumes of large wildfires (Notholt et al. 2003).

Carbonyl sulfide that reaches the stratosphere is oxidized by photolysis, forming sulfate aerosols which contribute to the Junge layer (Chin and Davis 1993). Eventually, these aerosols are removed from the stratosphere by downward mixing of stratospheric air. The concentration of COS in the atmosphere today is higher than in preindustrial times (Montzka et al. 2004), although it seems to have declined slightly in recent years (Rinsland et al. 2002, Sturges et al. 2001).

4.6.4 Stratospheric Ozone Chemistry

In the stratosphere, ozone is produced when ultraviolet light with wavelengths between 200 and 240 nm breaks the bond between two oxygen atoms in a molecule of oxygen:

(4.43)O2→hν2O.

In turn, the highly reactive oxygen atom reacts with other molecules of oxygen (as previously shown in Eq. 4.41) to form O3:

(4.44)O+O2→O3.

Ozone, in turn, can be destroyed by interaction with another photon that breaks it into an oxygen molecule (O2) and an oxygen atom (O). Stratospheric ozone also can be destroyed by reaction with other species, such as nitric oxide (NO), as shown in Eq. (4.42), and halogen atoms, such as chlorine and bromine. Chlorine and bromine atoms are released into the stratosphere from the photodegradation of haloalkanes, often called halons. Classes of haloalkanes that impact ozone chemistry include CFCs and hydrochlorofluorocarbons (HCFCs). The net concentration of ozone in the stratosphere is established by the rates of both the production and the destruction reactions.

Stratospheric ozone is produced at the highest rate in equatorial regions, where solar radiation is most intense. Stratospheric ozone does not occur as a sharply defined layer, but instead as a broadly distributed gas whose peak concentration occurs in mid-stratosphere. Ozone shields the Earth from UV-B radiation; it is the total mass of ozone per unit of Earth’s surface area, however, and not the ozone’s vertical distribution, that determines the effectiveness of the shielding.

The amount of ozone present in the atmosphere is expressed in Dobson units; one Dobson unit of ozone, if gathered together in a thin layer covering Earth’s surface at a pressure of 1 atm, would occupy a thickness of 1/100 of a millimeter (10 μm). Ozone coverage varies spatially and temporally but averages 300 Dobson units. Thus, the ozone shield, which protects life on Earth from damage by the UV-B radiation of the Sun (ultraviolet radiation in the 280-320 nm range), is equivalent to a layer of ozone only 3 mm thick at sea level pressure. When the levels of UV-B radiation at Earth’s surface increase, so does biological damage; humans are at a higher risk of skin cancer and certain eye disorders, and damage to crops, marine phytoplankton, and other organisms can occur.

Widespread concern about the diminishment of stratospheric ozone began in the 1970s. Molina and Rowland (1974) proposed that the release of CFCs through human activities played a major role in O3 depletion in the stratosphere. At that time, CFCs were extensively used as refrigerants, solvents, aerosol can propellants, and blowing agents for plastic foam manufacturing, as well as in fire extinguishers. In 1978, in response to scientific discoveries, the United States banned the use of CFCs as propellants in aerosol sprays. In 1987, the Montreal Protocol on Substances that Deplete the Ozone Layer was established to halt, by the year 2000, the production of most O3-destroying CFCs by industrialized countries. The Montreal Protocol was subsequently revised several times, in part to completely halt the production and consumption of the five major CFCs by 1996 and to completely phase out ten other CFCs and three bromine-containing halons by 2010. HCFCs, used as CFC replacements for refrigerants, solvents, blowing agents for plastic foam manufacturing, and in fire extinguishers, will also ultimately be phased out under the Montreal Protocol; reduction in consumption and production of HCFCs is planned for 2015. Although the Montreal Protocol is widely considered to be a successful international environmental treaty, the haloalkanes limited or banned by the Protocol will persist in the atmosphere for years to decades due to their chemical stability. Note the Protocol does not address two related classes of compounds, hydrofluorocarbons (HFCs) and perfluorocarbons (PFCs), because they do not deplete the ozone layer; they are greenhouse gases, however, just like CFCs and HCFCs, and therefore are also of concern from the standpoint of the Earth’s heat balance (see Section 4.7).

Concrete evidence for a dramatic decrease in stratospheric ozone over the South Pole was first presented in 1985, when a team from the British Antarctic Survey, using ground-based instrumentation, reported that average ozone concentrations over Antarctica for the month of October (polar spring) were decreasing (Farman et al., 1985). In the early 1970s at one site, Halley Bay, the October average was approximately 300 Dobson units, whereas in 1984, it was 180 Dobson units. Based on this finding, American investigators reanalyzed data obtained from the Nimbus 7 satellite from 1978 to 1986 (Stolarski et al., 1986) and confirmed the existence of the Antarctic ozone hole. The existence of the ozone hole had not been recognized earlier from the satellite data because the entire analysis was done by computer, and programmers had written the code to prefilter the data; all values less than 190 Dobson units were rejected as being impossibly low and therefore in error! Direct observations of Antarctic atmospheric ozone levels were made during the Antarctic Airborne Ozone Expedition in 1987, using ER-2 research aircraft (a derivative of the famous U-2 spy plane, also called the “Dragon Lady,” capable of flying to an altitude of 70,000 ft) and a lower-flying DC-8 aircraft. Ozone depletion over the North Pole has also now been documented. Figure 4.45 shows one of the earliest published maps of the ozone hole over Antarctica. Figure 4.46 shows the historical trends in ozone levels for both polar regions.

The current Antarctic ozone hole, an ozone-depleted region between approximately 12 and 30 km in altitude, is the result of the catalytic degradation of ozone by chlorine and bromine, which are released into the atmosphere through human production and use of CFCs and other halons. CFCs and other halons affect stratospheric chemistry because their chemical stability and resultant long lifetime in the troposphere allow them to mix into the stratosphere, where they are degraded by shortwave ultraviolet light and release chlorine and bromine. Much of the chlorine becomes incorporated in reservoir compounds, such as hydrochloric acid (HCl) and chlorine nitrate (ClONO2), which are not directly reactive with ozone. Free chlorine atoms, however, react directly with ozone molecules, destroying the ozone molecules and forming chlorine monoxide (ClO) and molecular oxygen:

(4.45)Cl+O3→ClO+O2.

A chlorine monoxide molecule in turn can react with an oxygen atom to form more molecular oxygen and a free chlorine atom:

(4.46)ClO+O→Cl+O2.

Alternatively, a chlorine monoxide molecule may react with another chlorine oxide molecule to form chlorine peroxide (Cl2O2):

(4.47)ClO+ClO→Cl2O2.

Chlorine peroxide is readily photodissociated, releasing chlorine atoms that can then attack more ozone:

(4.48)Cl2O2→hν2Cl+O2.

A single chlorine atom can destroy thousands of ozone molecules during its residence in the stratosphere (Dameris, 2010).

Bromine-containing halons also significantly contribute to the depletion of stratospheric ozone when these halons photodissociate and release bromine atoms. Just as a chlorine atom destroys ozone, as shown in Eq. (4.45), a bromine atom can also destroy ozone:

(4.49)Br+O3→BrO+O2.

The preceding reactions occur throughout the stratosphere and are responsible for depletion of ozone on a global scale. The extreme ozone depletion at the South Pole, however, did not at first seem to be consistent with the kinetics of the known ozone-destroying reactions and the known free chlorine concentrations. The discrepancy turned out to be due to heterogeneous chemistry (Fig. 4.47). Under wintertime conditions when polar stratospheric temperatures drop below approximately − 78 °C, polar stratospheric clouds (PSCs) are formed. These clouds can form from H2SO4/H2O droplets, which take up HNO3 under cold temperatures; from ice crystals formed from the condensation of water; and from solid hydrates of nitric acid, such as nitric acid trihydrate (NAT). In PSCs, three heterogeneous reactions occur, whose net effect is the release of chlorine from reservoir compounds:

(4.50)HCl+ClONO2→HNO3+Cl2,

(4.51)H2O+ClONO2→HNO3+HOCl,

and

(4.52)HOCl+HCl→Cl2+H2O.

PSCs thus catalyze the release of chlorine molecules from reservoir compounds, leading to the eventual formation, when springtime sunlight arrives, of chlorine atoms. PSCs also have a secondary effect, by redistributing the nitric acid (HNO3) concentration in the stratosphere through the formation and sedimentation of large NAT particles. As the NAT particles settle out of the lower stratosphere, nitrogen oxides are removed, preventing them from capturing chlorine atoms in the reservoir compound ClONO2; thus, ozone destruction can occur for a longer period of time.

Methane can also impact the stratospheric ozone cycle through two processes having opposite effects on ozone destruction. First, methane can convert Cl and ClO to HCl, which does not directly deplete O3. Second, the oxidation of methane in the stratosphere produces water vapor, enhancing the formation of PSCs which facilitate O3 destruction.

What is true about the region of Earth's atmosphere known as the stratosphere?

What are 3 facts about the stratosphere?

What is the correct describes the stratosphere?

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