Interactions Between the Lower, Middle and Upper Atmosphere

Space Sci Rev (2012) 168:1–21 DOI 10.1007/s11214-011-9791-y

Anne K. Smith

Received: 2 February 2011 / Accepted: 30 May 2011 / Published online: 19 June 2011 © Springer Science+Business Media B.V. 2011

Abstract This survey of atmospheric regions is based on a presentation given at the Work- shop on Coupling between Earth’s Atmosphere and its Plasma Environment. The workshop was sponsored by the International Space Science Institute (ISSI) during September 2010. This survey begins with descriptions of aspects of the troposphere, stratosphere, and mesosphere that affect vertical coupling. Then the tropopause, stratopause, and mesopause are considered, with a discussion of how they inhibit or allow the transmission of atmospheric mass or perturbations. Two transition interfaces are emphasized: the tropopause, which separates the troposphere from the middle atmosphere, and the mesopause, which separates the middle atmosphere from the thermosphere. The survey of coupling processes looks at radiative processes, wave propagation, and mass exchange. Each topic includes one or more speciﬁc examples drawn from recent literature.

Keywords Coupling · Solar variability

1 Introduction

This article introduces the vertical structure of the atmosphere and discusses the interfaces between regions or layers. The focus is on the interactions across regions and the impedi- ments to interaction. Brief descriptions of the lower and middle atmosphere are given and the regions are contrasted. Several examples of coupling across regions are presented. The additional articles in this volume include detailed discussions of aspects of these interactions and other processes not discussed here.

2 Deﬁnition of the Atmospheric Regions

Atmospheric scientists classify the neutral atmosphere in regions. The regions and the transitions between them are identiﬁed by their mean thermal structure. They have traditionally

A.K. Smith () NCAR, Boulder, CO, USA e-mail: [email protected] 2 A.K. Smith been considered to be layers but, since the atmosphere is not layered globally, the term regions will be used here. The regions are the troposphere (Earth’s surface to approximately 10–18 km), the stratosphere (10–18 km to about 50 km), the mesosphere (about 50 to 100 km) and the thermosphere (above 100 km). The transitions between these regions are the tropopause, the stratopause, and the mesopause. Because of common dynamical and chemical environments, the region that encompasses the stratosphere, mesosphere and sometimes the lowermost part of the thermosphere is also called the middle atmosphere.

2.1 The Troposphere

The troposphere is strongly affected by the Earth’s surface properties. – The orography of the surface affects the generation of atmospheric waves as winds flow- ing near the surface are displaced horizontally or vertically. – The landforms, vegetation, and snow cover of land surfaces affect absorption versus reflection of incoming solar radiation and the emission of outgoing infrared radiation. – The landforms, vegetation, and snow cover of land surfaces affect exchange of water and trace chemical species with the atmosphere. – The surface properties (for example, surface wave amplitudes) and ice cover of ocean and lake surfaces affect absorption versus reflection of solar radiation and the exchange of water and trace chemical species with the atmosphere. The troposphere is highly variable on short time scales; the details of this variability are often difficult to predict. Basic physical processes act to create unstable conditions. The responses to the unstable conditions are a variety of dynamical phenomena that act to bring the atmosphere back toward stability. An important mode of instability is convective instability. The bulk of the solar flux that enters the troposphere passes through the atmosphere and is absorbed at the Earth’s surface (IPCC 2007). A substantial portion of the energy is converted to heat. An additional portion of the energy evaporates water that then contributes to the humidity of the atmosphere. The transfer of the heat from the land or ocean surface to the atmosphere occurs by radiative (infrared), sensible, or latent heating. Reflection of solar radiation at the surface or by particles within the atmosphere also affects the energy exchange between the surface and the atmosphere. The heat input to the lower levels of the atmosphere can create convective instability. In dry convective instability (i.e., under conditions of low humidity), the air will be unstable wherever the temperature structure satisfies

∂T g < − ≡ d (1) ∂z cp where T is temperature, g is the acceleration of gravity and cp in the heat capacity of air at constant pressure. The so-called “dry adiabatic lapse rate” d described by (1)is−9.8K/km. If dry air is displaced upward, it will cool at the rate of 9.8 K/km. By definition, if dry air is convectively unstable, its temperature falls off more rapidly than that. As a result, air displaced upward will be warmer, and therefore lighter, than the surrounding air. The displaced air feels an upward force rather than a restoring force; it can continue to rise until it reaches an altitude where the temperature profile is no longer convectively unstable. This rising motion is convection. The opposite condition is an atmospheric profile that is convectively stable. In that case, air displaced upwards will cool and become colder than the Interactions Between the Lower, Middle and Upper Atmosphere 3 surrounding air. Gravitational force will then be a restoring force to accelerate the displaced air back towards its original altitude. When the humidity is high enough, the exchange of heat due to the condensation of moisture has an impact on the convection process. Air that ascends and cools can hold less moisture by virtue of its cooler temperature. As the relative humidity becomes high enough for the moisture to condense into cloud droplets, latent heat will be released. The latent heat release slows the rate at which the air cools. Under saturated conditions the moist adiabatic lapse rate s is

L dws s = d − (2) cp dz where ws is the mixing ratio at which water becomes saturated and L is the latent heat of condensation. The moist adiabatic lapse rate depends on both humidity and temperature. The lapse rate normally becomes steeper in the upper troposphere where both humidity and temperature are lower. The lapse rate is also expected to become less steep as the overall temperature increases due to global warming (Held and Soden 2006). Due to conservation of mass, convective upwelling must be balanced by downwelling of air in surrounding areas and the upwelling and downwelling air masses must be linked by horizontal motion. The combined effect of the upward and downward mass flow keep the mean atmosphere temperature structure over large parts of the Earth near the moist adiabatic lapse rate. The horizontal scale of convection is small (less than a few km) even though the heating that generates the instability can be widespread. The location and timing of the appearance of individual convective cells is difficult to predict. The atmospheric variables that best predict where convection will occur are discussed e.g. by Sherwood (1999). Convective instability is one of several important modes of instability. It is most important in the tropics where the heating is normally strong and the horizontal variations in angular momentum are the weakest. However, convection also occurs in midlatitudes, particularly over land surfaces during the warmer times of the year. The rotation of the Earth is associated with a change in the angular momentum from the equator (largest) to the poles (zero). The change in the angular momentum affects the horizontal path of air displaced either vertically or in latitude. Instabilities in which the conservation of angular momentum play a role are barotropic and baroclinic instability. Baroclinic instability is a key component of the generation of midlatitude storm systems. Baroclinic instability is an inevitable situation for the Earth’s atmosphere. The unstable situation develops because the atmosphere is heated more strongly in the tropics and therefore has an equator to pole temperature gradient. It also depends on the rotating Earth and the importance of the Coriolis effect on horizontal motion. Through the thermal wind relation, the meridional temperature gradient is associated with a vertical shear in the zonal wind. ∂u R ∂T ≈− (3) ∂z Hf ∂y where u is the zonal wind, R is the gas constant for air, f is the Coriolis term, and H is the atmospheric scale height. Small north-south perturbations in the air flow advect cooler air in the equatorward direction and warmer air poleward. Depending on the spatial scale and frequency of the perturbation and on the conditions of the background atmosphere, some disturbances will be unstable and their amplitudes can grow exponentially. During winter, baroclinic instability leads to the steady generation of midlatitude storm systems. With vigorous moist convection in the tropics, transport of air parcels from the bottom to the top of the troposphere can occur on timescales of minutes. The timescale for vertical 4 A.K. Smith transport is also fairly short (a few days) in midlatitudes. As a result, all but water and short- lived chemical species in the troposphere are well-mixed in the vertical. Due to instabilities, variations in external heating, and other processes, the winds, temperature, and humidity in the troposphere vary with space and time on multiple scales. Some of these variations damp out with time and have primarily a local impact. However, many perturbations can propagate in the horizontal and/or vertical direction. Propagating disturbances are atmospheric waves. Their behavior in space and time can be represented by a wave equation in one or more dimensions. The generation and propagation of waves is a dominant aspect of the atmosphere at all levels and is a key component linking different regions. Waves generated in the troposphere that have important impacts on the middle atmosphere include gravity waves, thermal tides, and planetary Rossby waves. Examples of the role of these processes in vertical coupling are presented in Sect. 4.2. The propagation conditions for all of these waves depend on the background winds, the thermal structure of the atmosphere, and characteristics of the wave such as wavelength and frequency. The linear wave equations, based on the assumption that the wave amplitudes are small, are often good predictors of wave propagation even when the conditions of linearity are violated. Although most of the waves that are important in the atmospheric system are generated in the troposphere, waves can also be forced in the stratosphere and mesosphere. Two examples are the generation of the quasi-two-day wave by baroclinic instability (Plumb 1983;Pfister 1985) and the forcing of the diurnal and semidiurnal tide by ozone heating (Hagan 1996).

2.2 The Stratosphere

2.2.1 Heating

The heating in the stratosphere occurs internally, unlike in the troposphere where much of the heating occurs at the interface with the surface. Most of the stratospheric heat comes from absorption of solar radiation by ozone. Although molecular oxygen is less important than ozone in the thermal budget, absorption of radiation in the Herzberg bands by O2 is the source of O which in turn is the source of ozone (ozone is formed from the reaction of O with O2). The solar ultraviolet radiation also photolyzes many other molecules and drives the active chemical system in the stratosphere. Throughout the middle atmosphere, CO2 radiative transfer is the dominant cooling process. The increase in CO2 since the industrial revolution is changing the radiative balance and is expected to lead to a continued cooling of the stratosphere. Water vapor and ozone also make contributions to the cooling. Ozone changes affect both the heating and cooling rates. Changes in these three gases over recent decades have led to temperature trends in the stratosphere (Ramaswamy et al. 2001). The stratospheric temperature structure is not in radiative balance. Heating occurs in the sunlit atmosphere whereas cooling occurs at all latitudes and all local times. There is no solar heating in the darkness of polar night. However, heating due to dynamically driven adiabatic descent occurs and keeps the polar temperature higher than that predicted from thermal equilibrium.

2.2.2 Dynamics

The adiabatic heating that is so important in the middle atmosphere is the result of a wave- driven circulation. Waves propagate from the troposphere. According to theory, waves that Interactions Between the Lower, Middle and Upper Atmosphere 5 are steady and not dissipating will not interact with the background atmosphere (Andrews and McIntyre 1976). The waves that are most important for the stratosphere are planetary Rossby waves, gravity waves, and equatorially trapped waves such as Kelvin waves. Planetary Rossby waves, which can normally propagate into the stratosphere only during winter (Charney and Drazin 1961), cause much of the variability in the wintertime middle atmosphere, including sudden stratospheric warmings (McIntyre 1982). Their propagation is affected by the background winds and temperature. They are forced in the troposphere and can be dissipated by long-wave radiative transfer or by Rossby wave breaking (McIntyre and Palmer 1983). Quasi-stationary planetary Rossby waves are always present during the winter months. They are readily observed by satellite and are well-simulated in global climate models. Their amplitudes tend to be larger and more variable in the NH. Gravity waves generated in the troposphere propagate into the stratosphere. Because they are more difficult to observe and cannot be represented in most global climate models, the global impact of gravity waves on the stratosphere is less well-established. Dissipating gravity waves are responsible for the structure of the stratopause and polar night jet in the high latitude winter (Hitchman et al. 1989; Siskind et al. 2007). Although equatorially trapped waves have a limited geographical domain, their impact in the tropics is important. In the classic simulation of the quasi-biennial oscillation (QBO) of the zonal wind over the equator, Lindzen and Holton (1968) and Holton and Lindzen (1972) were able to generate the oscillation from interactions of equatorial Kelvin waves, mixed-Rossby-gravity waves, and other gravity waves with the background wind. More recent simulations by Kawatani et al. (2010a, 2010b) showed that the spectrum of waves that contributed momentum to the driving of a simulated QBO was quite broad. Waves from planetary scale, such as Kelvin waves, to short horizontal wavelength gravity waves all contributed to the generation of the QBO in the model. The 2-dimensional mean circulation in the stratosphere is best viewed using the transformed Eulerian framework introduced by Andrews and McIntyre (1976). The transformed Eulerian mean is a good approximation of the zonal mean Lagrangian mass flux. The transformed Eulerian mean meridional and vertical mass fluxes are determined by adding the Stokes drift due to waves to the Eulerian mean meridional and vertical winds. For steady non-dissipative waves, the Eulerian mean winds and the Stokes drift cancel and there is no net mass flux. A residual remains for dissipating waves. Thus a wave-driven circulation is one that owes its existence to dissipating waves. The mass circulation is known as the transformed Eulerian mean or residual circulation. In the stratosphere, it is referred to as the Brewer-Dobson (B-D) circulation after pioneering inferences from observations of water by Brewer (1949) and of ozone column by Dobson (1956). In the lower stratosphere, the circulation goes from tropics to poles (Dunkerton 1978). Air enters the stratosphere through the tropopause in the tropics and slowly moves toward both poles. It can then reenter the troposphere in middle or high latitudes of both hemispheres. In the upper stratosphere, the B-D circulation has a global cell; it moves air from the summer pole toward the winter pole. One useful variable for assessing the B-D circulation is the age of air (Hall and Plumb 1994). The age of air measures the time that has elapsed since an air parcel entered the stratosphere. In the high latitude lower stratosphere the age can be several years. To determine the age of air from observations, studies have looked at the concentrations of long-lived anthropogenic gases that have a monotonic trend at the surface, such as CO2 and SF6. Although these tracers work well for determining the age of air, they are not ideal for investigating trends that might have occurred as a result of climate change. Variations in the production rate over the years make the analysis highly complicated and sometimes ambiguous (Garcia et al. 2011). 6 A.K. Smith

The tropopause is highest and coldest in the tropical region. Because cold air can hold little water, air entering the stratosphere is extremely dry. The seasonal cycle in the tropical tropopause temperature (coldest during NH winter) gives a seasonal cycle in the humidity of the air entering the stratosphere. The vertical mixing is sufﬁciently low that patterns of higher and lower water mixing ratio can be seen ascending in the lower stratosphere. This phenomenon, called the “atmospheric tape recorder”, was ﬁrst reported by Mote et al. (1996).

2.2.3 Composition

The distribution of the dominant molecules in the stratosphere is the same as in the troposphere: molecular nitrogen and oxygen are the dominant gases and are well-mixed (Brasseur and Solomon 2005). The CO2 concentrations are similar to those in the troposphere but with a lag in the growth by up to 4–5 years. As mentioned in Sect. 2.2.2, the air entering the stratosphere is very dry. Currently, roughly half of the water in the stratosphere enters through the tropopause and the other half is produced within the stratosphere from hydrogen released by the destruction of methane (CH4). The contribution of methane is believed to have been lower in preindustrial times. The most important trace gas in the stratosphere is ozone. Ozone protects the Earth’s surface from ultraviolet radiation, provides most of the heating of the stratosphere by solar insolation, and is highly active chemically. It has been extensively studied from satellite using infrared or microwave emissions, and ultraviolet absorption or scattering. The production of ozone depends on the solar ultraviolet ﬂux (which photolyzes molecular oxygen) and the temperature (which affects the kinetic rate of the reaction that makes ozone from molecular oxygen and atomic oxygen). The production rate is therefore not strongly inﬂuenced by global changes in the atmospheric composition. An exception to this generalization is the impact of lower temperature due to greenhouse gas increases. During daylit hours, ozone is lost by photolysis but is quickly regenerated when the O combines with O2 to form another ozone. However, there are a series of catalytic cycles that remove both O3 and O, and hence cause a net loss of ozone. These involve the molecules and atoms in the hydrogen, nitrogen, chlorine, and bromine families. Anthropogenic increases in chlorine and bromine are responsible for the severe springtime ozone loss in the lower stratosphere over Antarctica (the ozone hole), for ozone loss in the NH winter/spring lower stratosphere, and for ozone loss in the global upper stratosphere (WMO 2010). Many of the reactions involved in the catalytic ozone-loss cycles have reaction rates that depend on temperature. The overall effect of a decreasing temperature, with no changes to the background composition, would be to increase the ozone production rate and decrease the ozone loss rate. The combination leads to a marked anti-correlation of perturbations in temperature and ozone. This applies not only for perturbations of temperature and ozone due to waves on timescales of days to weeks but also for predictions of the response of the global ozone layer to greenhouse gas increases.

2.3 The Mesosphere

2.3.1 Heating

Like the stratosphere, the heating in the mesosphere occurs internally from solar ultraviolet radiation. The dominant absorber is molecular oxygen. Another molecule that contributes is water, which absorbs Lyman-α and Schumann-Runge radiation. The absorbed energy can Interactions Between the Lower, Middle and Upper Atmosphere 7 photolyze the absorbing molecule, excite it to a higher state of energy, and/or contribute to increasing its kinetic temperature. The energy is not necessarily available as heat at the time of absorption. Consider the absorption of ultraviolet radiation (<240 nm) by molecular oxygen. The photolysis gives two products: O(3P) and O(1D), i.e. atomic oxygen in the ground and excited states. The photolysis uses much of the energy but the remainder is available as heat. The excited O(1D) atom can participate in one of many chemical reactions (O(1D) is short- lived and is much more reactive than O(3P)); it can collisionally transfer its energy to excite another atom or molecule; it can be collisionally quenched and convert the energy to heat; or it can spontaneously emit a photon and drop to the ground state. The O(3P) atom has a lifetime that ranges from very short in the lower mesosphere to months in the upper mesosphere (Smith et al. 2010b). The reactions that either directly (reaction with O or O3) or indirectly (catalytic cycles with hydrogen-containing species) 3 convert O( P) to O2 are exothermic. This means, in simple terms, that the energy from the photon that photolyzed O2 is not available as heat until the products recombine into another O2 molecule. Substantial transport can occur in the interim. For example, O generated in the lower thermosphere can be transported downward into the upper mesosphere and provide a source of heating even though the energy from the solar radiation did not penetrate to this part of the atmosphere. The heating can be substantial during night due to the impact of exothermic reactions. Some of the energy from O2 or O3 photolysis contributes to sponta- 1 1 neous emission by O( D) and other species (notably O2( ) and vibrationally excited OH) as airglow. Much of the energy that is emitted as airglow is lost to space or to the ground and does not contribute to heating. The loss of energy through airglow is substantial (Mlynczak and Solomon 1993 estimate 20–40% for some processes) but is not precisely known. The delay between the absorption of energy and the heating effect of the absorbed energy has an additional impact on the dynamics. The net heating rate can shift away from having a pronounced daytime maximum. The two exothermic reactions that have the largest impact near the mesopause are O + O (i.e., the self-reaction of atomic oxygen) and O3 + H. The former is similar during day and night since O does not have a photochemically driven diurnal cycle. The latter is actually larger during night because of the very large diurnal cycle in the concentration of O3 (Smith and Marsh 2005). One net impact of the exothermic reactions is to damp the magnitude of the diurnal cycle in heating in the upper mesosphere.

2.3.2 Dynamics

A defining characteristic of the mesosphere is the cold summer mesopause. The mesopause at the summer pole is the coldest place in the Earth’s atmosphere. The mesopause is in- teresting for its cold temperature in summer and also because it is very far from radiative balance (the summer is much colder than the winter). The cold summer temperature provides an environment for polar mesospheric clouds (also called noctilucent clouds) even though the humidity is very low. The cold conditions are maintained by strong upwelling. The upwelling is part of a global circulation cell that moves air up at the summer pole, horizontally from the summer to the winter, and down at the winter pole (Smith et al. 2011). The mesospheric circulation is driven by dissipating and/or breaking gravity waves (Lindzen 1981; Holton 1982). The waves are generated within the troposphere and then filtered by winds in the stratosphere. A gravity wave cannot propagate when the wave phase speed is equal to the background wind. The point where the phase speed equals the background wind is know as a critical layer. Because the stratospheric zonal winds are stronger than the meridional winds, the critical layer filtering has the largest impact on waves propagating in the eastward and westward directions. The strong stratospheric jets in the solstice 8 A.K. Smith periods filter out westward waves during the summer and eastward waves during the winter. The waves that are not filtered continue to propagate vertically but eventually break or dis- sipate in the upper mesosphere. The momentum deposition from the resulting asymmetric spectrum of gravity waves causes the stratospheric jets to reverse direction in the mesosphere. The momentum deposition acts together with the Coriolis force to drive the summer to winter flow in both hemispheres. The strong dependence of the mesospheric temperature distribution on waves from below also means it is susceptible to differences or changes in the regions below. An example of the impact of changes in the lower stratosphere is given in Sect. 4.2.

2.3.3 Composition

The composition in the lower mesosphere is similar to that in the stratosphere. One difference is a shift in the catalytic cycles that destroy ozone. In the stratosphere, reactions involving chlorine species and nitrogen species are very important in determining the ozone balance. Beginning in the lower mesosphere, the catalytic cycles that destroy ozone are strongly dominated by those involving hydrogen species. The reactive hydrogen species H, OH and HO2 are products of the destruction of methane and water. Reactions involving species in the hydrogen and oxygen families continue to be important through the entire mesosphere. As noted above, many of the reactions are exothermic and they contribute to the distribution of heating. The proportion of heat from these reactions becomes larger in the upper mesosphere lower thermosphere (MLT) region. The composition in the MLT differs from that in the stratosphere and lower mesosphere in several important ways. – The gases are no longer well-mixed but are beginning to separate based on their molecular masses. – Ions begin to make a contribution to the chemical system. – Some free radicals, for example hydroxyl (OH), become longer-lived and play a more central role in the chemical system. – Atomic oxygen transitions from a very short lifetime below about 80 km to a long-lived and abundant gas in the upper mesosphere. – Excited states of atoms and molecules affect the chemistry and energetics. – Ozone has a large diurnal cycle (smallest during daytime) due to the very rapid photolysis rate. – Several important molecules, most notably CO2, are not in local thermodynamic equilibrium (LTE) in the mesosphere; see below. Non-LTE conditions occur when low atmospheric density leads to a low rate of collisions between molecules (Andrews et al. 1987). In general, strong cooling of the atmosphere occurs through emission by vibrationally excited CO2. However, under non-LTE conditions, the low collision rate limits the exchange of energy from kinetic to vibrational levels. As a result, the population of vibrational levels is not consistent with a Boltzmann distribution and the infrared CO2 cooling rates are substantially reduced over the LTE situation. Several of the characteristics of the mesosphere (extremely cold temperatures, importance of radical species and excited states) contribute to uncertainties in the chemical system. Laboratory measurement of the key reaction rates under realistic conditions is extremely difﬁcult. Satellite observations using atmospheric emissions are constrained by the very low densities and therefore weak signals of natural emissions. Retrieval of temperature or constituent densities using measured emissions from a non-LTE atmospheric state requires Interactions Between the Lower, Middle and Upper Atmosphere 9 knowledge of the energy structure of the molecule. For all of these reasons, the detailed knowledge of mesospheric composition lags that of the stratosphere. For example, observational estimates of one of the most important gases, atomic oxygen, differ substantially. As pointed out in Smith et al. (2010b), the O concentrations from different analyses are not compatible. O is derived during day and night using two different tech- niques and different radiance channels from the SABER instrument. When tidal variations are taken into account, the concentrations are consistent between day and night, support- ing the reliability of the concentrations retrieved from SABER. However, O derived from SABER is larger by about a factor of four than the concentrations derived from WINDII by Russell et al. (2004). The SABER values are also larger than the concentrations in the MSIS empirical model (Picone et al. 2002).

3 Interfaces Between the Atmospheric Regions

3.1 The Tropopause

The tropopause has traditionally been identified as the altitude of coldest temperature or, equivalently, the altitude where the decline of temperature with altitude ceases. However, for certain purposes, it has been more useful to fine-tune the definition. The tropopause can also be identified by isentropic surfaces or by potential vorticity (Holton et al. 1995). Chemically, the tropopause marks the location of sharp decreases in water and increases in ozone. It is the point where vigorous vertical mixing ceases. The altitude of the tropopause is higher in the tropics (equatorward of 30°) than at high latitudes. The altitude ranges from about 18 km near the equator to about 10 km at the poles and is discontinuous at ±30°. In the tropics, rather than a distinct tropopause boundary, the transition region is known as the tropical tropopause layer, or TTL (e.g., Fueglistaler et al. 2009). This takes into account the finite depth over which the transitions from troposphere to stratosphere occur. The TTL region separates two regions which have quite different dynamics. Since the air that enters the middle atmosphere passes through this region, its impact on transport of trace species is important. The processes responsible for the tropopause were discussed by Held (1982). The tropopause marks the transition between the vigorous dynamical mixing in the troposphere and the radiative balance in the stratosphere. A general idea of the tropospheric thermal structure and tropopause location can be deduced using the concept of radiative-convective balance. Solar visible radiation passes through the atmosphere and is absorbed at the Earth’s surface (land or ocean). Sensible and radiant heat from the surface warm the atmosphere from below. A purely radiative balance is not in general possible in the troposphere because the heating from below, particularly over land surfaces, leads to convective instability near the Earth’s surface. The temperature profile of an atmosphere in radiative convective balance is constrained to maintain stability, i.e., to have a lapse rate that is less than d for dry air or s for moist air. This concept has been used in simple radiative balance climate models for many decades. In these models, excess heat in the lower troposphere is redistributed in the vertical direction (as if by convection) until the profile achieves neutral stability. As discussed in Sect. 2.2, the heating in the stratosphere comes from absorption of solar ultraviolet radiation by ozone. The maximum temperature at the stratopause occurs near where the optical depth for the ozone heating (primarily the Hartley band) is unity. The tropopause then marks the transition between the negative temperature gradient of the tropospheric radiative-convective system and the positive gradient of the radiative balance system in the stratosphere. 10 A.K. Smith

Recent work has contributed to the understanding of the maintenance of the tropopause. The altitude and temperature of the tropopause are sensitive to the radiative balance in the upper troposphere and stratosphere and the stratospheric circulation (Thuburn and Craig 2000; Shepherd 2002; Schneider 2004;Birner2010). Water in the middle atmosphere depends on the transport of both water and methane across the tropical tropopause. As air passes through the very cold temperature of the tropical tropopause, much of the water condenses into ice particles and settles out. This “freeze-dried” air has water concentrations of only a few ppm. The water concentration increases above the tropopause due to formation from byproducts of methane destruction. The methane transport into the stratosphere depends on the tropospheric concentration and the net mass ﬂux across the tropopause; it is not directly affected by the tropopause temperature. Observations indicate that the concentration of water in the stratosphere has varied over decadal time scales. Part of the variation is a result of the increase of methane, which leads to more water production by photochemistry within the stratosphere. However, the observed variations in water are larger than can be explained by the methane increase and variations in the tropical tropopause temperature. Water increased for several decades until about 2000 and has decreased since then (Randel et al. 2006; Fujiwara et al. 2010) although the reasons are still not understood. Stratospheric water is important because of its impact on the global climate. The infrared effects of stratospheric water increases during 1980–2000 can account for a acceleration of climate warming at the Earth’s surface (Solomon et al. 2010). Likewise, the water decrease since 2000 has slowed down the surface warming compared to what it would otherwise be with unvarying stratospheric water. Changes in the ﬂux of water from the troposphere to the middle atmosphere also have an impact on the chemistry in the mesosphere (Marsh et al. 2003) and on the formation of polar mesospheric clouds (Thomas 1991).

3.1.1 The Stratopause

The transition between the stratosphere and mesosphere is defined by a change in the vertical gradient of the mean temperature from positive to negative. Although there are some differences between the stratosphere and mesosphere (see Sects. 2.2 and 2.3), the stratopause region is not normally a place of major transition in either dynamics or composition. This is the reason that the terminology now often uses middle atmosphere rather than stratosphere and mesosphere. The stratopause owes its altitude to two processes: the mean radiative balance and the thermal structure due to gravity waves. The radiative balance is due to ozone, which is produced in the stratosphere. Ozone absorbs solar ultraviolet radiation in the Hartley, Huggins and Chappuis bands. The largest heating occurs around 40–50 km; this is the mean altitude of the stratopause. Away from the high latitude winter, the stratopause altitude is well predicted by the radiative balance. This is approximately the altitude where the optical depth of ozone is unity. The maximum heating rate leads to a maximum in the temperature. In the high latitude winter, there is no sunshine so the heating is small. In this region, the stratopause is strongly influenced by the adiabatic warming due to sinking motion (Hitch- man et al. 1989). This is particularly the case in the NH stratosphere, where the winter dynamics is more active. Gravity wave driving of the circulation, which becomes dominant in the mesosphere, can also affect the stratospause region. Gravity wave driving of the polar winter circulation appears to be stronger in the SH. Another quality that distinguishes the winter stratopause from the tropopause and mesopause is the very large interannual variability in its location. Three well documented events Interactions Between the Lower, Middle and Upper Atmosphere 11 during recent NH winters (2003–2004, 2005–2006, and 2008–2009) show a displacement of the stratopause to much higher altitudes. All three cases were characterized by the sudden displacement upwards and slow return (over a period of weeks) for a region poleward of 70°N (Manney et al. 2005, 2008, 2009). These events highlight the importance of dynamics in controlling the stratospheric structure in winter high latitudes. The poleward flow within the stratosphere is driven primarily by quasi-stationary Rossby waves. The poleward flux is directly associated with downwelling motion in the polar winter. The stratospheric temperature is also affected by the larger-scale mean circulation that is characterized by downwelling through the entire winter polar middle atmosphere. Gravity waves are a prime driver of this circulation in the upper levels (above 40–50 km). The vertical propagation and eventual breakdown of gravity waves are strongly affected by the winds through which they propagate.

3.1.2 The Mesopause

At the mesopause, the vertical gradient of temperature switches sign again, from negative back to positive. The mesopause region is the coldest part of the Earth system. The overall temperature is low because of a lack of heating. Molecular oxygen is the leading absorber of solar radiation. In a climatological sense, there are two distinct mesopause positions. In the summer high latitudes, the altitude is about 83–89 km (von Zahn et al. 1996) while at all other latitudes and times, its position is close to 100 km. Model simulations indicate that the lower and colder position of the summer mesopause is due to very strong adiabatic cooling associated with strong upwelling. On top of this mean description are several details. First, there is a difference in the summer mesopause altitudes between the Northern (NH) and Southern Hemisphere (SH). The SH position is higher and not quite as cold (Xu et al. 2007). Polar mesospheric cloud abundances are lower in the SH (Bailey et al. 2007). Processes that have been proposed to account for this difference include the annual cycle in Earth-Sun distance (Chu 2003)and differences in gravity wave driving of the circulation due either to tropospheric sources or stratospheric filtering (for example, Dowdy et al. 2001). Another feature of the mesopause is the existence of temperature profiles that contain more than one minimum. The “mesospheric inversion” or “double mesopause” has been documented in observations and models (Meriwether and Gardner 2000). Analysis by Sassi et al. (2001) and Salby et al. (2002) show that, in winter midlatitudes, the double mesopause phenomenon is associated with planetary-scale variations in dynamics. The existence of planetary waves in the stratosphere and mesosphere affects the filtering, propagation, and breaking of gravity waves. The combined impact of planetary and gravity waves can give sharp changes in temperature and other dynamical fields in some phases of the planetary wave. From a ground-based observation, where only a single longitude can be observed, multiple “mesopauses” can appear in the vertical. Although single site observations of a double mesopause are common, this appears to be solely a local phenomenon. The mesopause is normally considered to be in the vicinity of the turbopause or homopause (Chabrillat et al. 2002). The turbopause is the altitude where molecular dissipation of waves is more important than turbulent dissipation. Molecular diffusion processes sup- presse wave breaking and turbulence, hence the term turbopause. The definition of the homopause is related: it is the altitude at which diffusive separation begins. Both the turbopause and homopause represent a transition to conditions where collisions are not frequent enough to keep the atmospheric constituents well mixed. Observations of wave properties (Offer- mann et al. 2007) and composition (such as CO2; see review by López Puertas et al. 2000) 12 A.K. Smith suggest that there is a gradual change toward diffusive separation rather than an abrupt onset for either the turbopause or the homopause.

4 Coupling Between Regions

As discussed in Sects. 2 and 3, the basic thermal structure and the interfaces between atmospheric regions are affected by dynamical and radiative processes through the depth of the atmosphere. In this sense, vertical coupling is the basis of the mean atmospheric structure. However, here the term “coupling” is used to refer not to the maintenance of the basic structure but to the response of the atmosphere to perturbations. This section gives brief summaries of some examples of vertical coupling between atmospheric regions.

4.1 Radiative Coupling

The penetration of incoming solar radiation through the Earth’s atmosphere varies strongly with wavelength. The absorption of the radiation depends on the concentrations of gases. Above the stratopause, the conversion of the absorbed energy to heat also depends on the concentrations of gases that are involved in exothermic reactions or in quenching excited states. The interfaces between atmospheric regions (the tropopause, the stratopause, etc.) do not in themselves pose any barrier to the transmission of radiation. The infrared radiation emitted by the Earth’s surface and by certain gases in the atmosphere can be reabsorbed and reemitted multiple times, particularly in the troposphere where the density of absorbing gases and aerosols is higher. Absorption can also be a heat source in the lower stratosphere. In the middle atmosphere, emission of infrared radiation by CO2 is the leading cooling process. Perturbations to the radiative balance can occur through – changes in the solar flux (such as the 11-year solar cycle); – changes to the absorption of solar radiation (such as reduced absorption in the stratosphere due to ozone loss); – changes in the reflection or absorption of radiation at the Earth’s surface (for example due to changes in snow-cover); – changes in the reflection or absorption of radiation in the Earth’s atmosphere (for example due to changes in aerosol or cloud amounts); – changes in the absorption of infrared radiation in the atmosphere (due to changes in water, CO2, or other radiatively active gases); – changes in the composition in the upper mesosphere that affect the energy loss through airglow. The radiative changes can lead directly to changes in the thermal structure. Changes in the thermal structure, particularly if the magnitude of the change varies with latitude, can lead directly to dynamical changes. Examples are given in Sects. 4.1.1 and 4.1.2.

4.1.1 Impact of Antarctic Ozone Loss in the Stratosphere on the Troposphere and Mesosphere

The primary heating in the stratosphere is absorption of solar ultraviolet radiation by ozone. Since 1980, the ozone over Antarctica during the spring has been substantially reduced from Interactions Between the Lower, Middle and Upper Atmosphere 13 levels prior to that year. The cause of this ozone hole is well-understood; it is due to ozone destruction by chlorine and bromine containing compounds that become particularly reactive during the special conditions of SH polar winter and spring. One effect of the ozone loss is a reduction of absorbed solar radiation, which leads to reduced heating. The negative trend in temperature in the lower SH stratosphere during winter is clearly evident in observations (Thompson and Solomon 2002;Randeletal.2009). Due to the thermal wind relation in the atmosphere (3), this polar temperature trend is associated with a trend in the mean jet strength in the SH spring and early summer. Much attention has been focused on the climate effects of the SH ozone loss (for example, Gillett and Thompson 2003; Baldwin et al. 2007; Son et al. 2010). A useful analysis tool for this is comparisons of past climate simulations (hind-casts) and future climate predictions between models with realistic ozone loss (and recovery in the future) and those without. These show that the effect of the ozone loss in the SH and of the future predicted ozone recovery are felt all the way to the Earth’s surface. Models with realistic ozone loss (either through interactive stratospheric chemistry or through speciﬁcation from observations) show substantially less warming over Antarctica during the 20th century than those that have non- varying ozone or that do not resolve the stratosphere. Changes in the SH stratosphere due to the ozone hole can also affect the mesosphere. The primary mechanism is changes in the propagation of gravity waves. The ozone hole led to trends in the lower stratosphere temperature and winds during late spring and early summer. The trend in the zonal mean wind, as simulated in the Whole Atmosphere Community Climate Model, have affected the propagation of gravity waves into the mesosphere (Smith et al. 2010a). As a result, the model simulations show that the strength of the summertime mean circulation in the SH decreased at the onset of the ozone hole around 1980. The model also simulated warmer summer mesopause temperatures due to the change in gravity wave interactions. The previous example illustrates the interlinked nature of some of the coupling processes. In this case, a chemical change in the stratosphere (ozone loss due to increased chlorine and bromine) leads to a radiative change (less heating due to lower ozone). The radiative change affects the temperature which, through the thermal wind relation, affects the zonal wind. The stratospheric wind change affects the propagation of gravity waves from the troposphere to the mesosphere. The breaking or dissipation of these waves then affects the strength of the circulation in the mesosphere and, through changes in the adiabatic heating, the polar mesopause temperature (Smith et al. 2010a).

4.1.2 Impact of Stratospheric Aerosols on Tropospheric Climate

Volcanic eruptions can inject trace chemical species directly into the atmosphere. One of these gases is SO2. The sulfur that is injected into the troposphere washes out in a matter of days to weeks, so its direct climate impact has limited duration. Over a period of weeks, the sulfur injected into the stratosphere by explosive volcanoes becomes incorporated into sulfate (H2SO4) aerosol particles. The eruption of Mt. Pinatubo, the largest in the satellite era, injected about 20 Mt of sulfur directly into the tropical lower stratosphere. The circulation in the lower stratosphere (the Brewer-Dobson circulation discussed in Sect. 2.2.2) tends to move air upwards in the tropics, poleward and then downward at the middle and high latitudes with a timescale of several years. Therefore, the sulfur injected by tropical volcanoes can persist for longer, become globally distributed, and can therefore have an impact over a wider area and longer period (Robock 2000; Stenchikov et al. 2006). 14 A.K. Smith

The best known climate impact of volcanic eruptions is cooling at the Earth’s surface, especially during summer. The cooling is caused by the radiative impact of aerosols. The particles absorb and reﬂect some of the solar radiation and therefore reduce the radiation at the Earth’s surface. The net effect of eruptions on the surface temperature is difﬁcult to determine because of few eruptions and large natural variability but is estimated to be 0.1–0.2 K (Robock 2000). The stratosphere is a crucial piece in the climate response. The timescale of several years for air to move through the stratosphere keeps the volcanic aerosols present in the system. The impact of volcanic aerosols on the heating also affects the stratospheric circulation. The enhanced heating in the lower tropical stratosphere can increase the B-D circulation (Tie et al. 1994). The altered circulation due to Mt. Pinatubo increased the strength of the tropical upwelling motion. This branch of the circulation brings up air that is low in ozone; therefore the net effect of a stronger upwelling was to decrease ozone in the tropical lower stratosphere. There is also a photochemical response of ozone to volcanoes; the aerosol particles enhance the rate of ozone loss due to chlorine and bromine. The distribution of aerosols in the stratosphere varies irregularly over multi-year timescales, due to intermittency in major volcanic eruptions and to variations in their latitude, time of year, and sulfur content. The impact that a given eruption can have depends on the transport that distributes the volcanic products within the stratosphere. If the stratospheric circulation accelerates due to climate change, as predicted by climate models (WMO 2010), this will affect the climate response to eruptions.

4.2 Coupling by Wave Processes

Vertical propagation of atmospheric waves can carry momentum and energy from one region of the atmosphere to another. To maintain constant kinetic energy, wave amplitude grows exponentially with decreasing density. Under favorable conditions for propagation, a small perturbation can grow to a large wave at a higher altitude. However, the converse implies that a perturbation at one altitude will have rapidly diminishing impact at lower altitudes. As discussed above, waves will not interact with the background atmosphere unless they are transient or dissipating (Andrews and McIntyre 1976). Sections 2.2.2 and 2.3.2 mentioned the importance of planetary waves and Kelvin waves in the stratosphere and of gravity waves throughout the middle atmosphere. While the waves themselves are generated in the troposphere, the propagation and dissipation depend on the atmosphere through which they propagate. The example of the impact of lower stratospheric changes on the wave processes in the mesosphere (Sect. 4.1.1) illustrates this. The message from this is that perturbations due to waves depend on the wave generation, the wave propagation, and the wave dissipation. Gravity waves cannot propagate beyond a critical line, where the background wind speed is equal to the phase speed of the waves. As discussed in Sect. 2.3, the critical layer effect leads to filtering of gravity waves by the mean stratospheric winds. This filtering then leads to an asymmetric spectrum of gravity waves. When those waves break in the mesosphere, the momentum forcing drives the mesospheric winds and circulation (Holton 1982). The winds in the winter stratosphere are never zonally symmetric because of the presence of large amplitude planetary waves. An individual small scale gravity wave will propagate vertically through the local “background” wind, not the zonally averaged wind. This means that the filtering of gravity waves changes with longitude. When those waves break, the momentum forcing will also change with longitude. The variations in momentum forcing will then lead to wavelike perturbations in the mesospheric winds. This mirror effect has Interactions Between the Lower, Middle and Upper Atmosphere 15 been seen in observations (Smith 1996, 1997) and in numerical simulations (Smith 2003). In observations, it is manifest as an anti-correlation between waves in stratospheric winds and those in mesospheric winds. Analysis has shown connections between variations in the middle atmosphere at northern and southern high latitudes. One of these phenomena is a relationship between the amplitude of the semidiurnal tide in the summer hemisphere and the planetary waves in the stratosphere of the other hemisphere (Portnyagin et al. 1998; Smith et al. 2007; Murphy et al. 2009). Simulations by Yamashita et al. (2002) suggest that this is due to a propagation of the tide from the winter to the summer hemisphere. Although the tides are global modes, the details of their generation and vertical or horizontal propagation in an atmosphere with spatial and temporal variations in temperature, winds, and heating are still not well understood. Another connection between the hemispheres is a correlation between planetary waves in the winter stratosphere and the polar mesospheric clouds at the summer mesopause (Karls- son et al. 2009). The processes that generate this link are not well understood. Analysis indicates that the signal probably goes from the stratosphere to the mesosphere of the winter hemisphere. Transfer of the signal to the opposite hemisphere may involve the mean circulation and/or impacts of wind changes on the propagation or breaking of the gravity waves that drive the circulation.

4.3 Coupling by Mass Exchange

Coupling among atmospheric regions can also occur by exchange of mass. One relevant process is the transport of chemically active gases from the mesosphere to the stratosphere. Another is the transport of water from the troposphere to the stratosphere. As discussed in Sect. 3.1, the tropopause marks the top of the region of rapid vertical mixing. Movement of mass across the tropopause occurs much more slowly than vertical movement of air within the troposphere. Flow across the tropopause is important for the composition of both the troposphere and the stratosphere. The impact on stratospheric water was discussed in Sect. 3.1. The advection of stratospheric air into the troposphere increases the concentrations of ozone and other gases (Holton 1990; Olsen et al. 2002). Although the production of chlorofluorocarbons (CFCs) has been phased out by the Mon- treal Protocol, these molecules still persist in the atmosphere. The most abundant CFCs have very long lifetimes in the troposphere but can be destroyed by photolysis and chemical reactions in the stratosphere (WMO 2010). The slow removal reflects the slow rate at which air from the troposphere is moved into the stratosphere and then returns to the troposphere. Transport of mesospheric air into the stratosphere has received much attention in the last decade. A motivating factor for this is to assess the impact of reactive nitrogen (NOx, which consists of NO and NO2) on stratospheric ozone. Downward motion by the mean circulation during winter (see Sects. 2.2 and 2.3) can transport trace species with positive vertical gradients. Perturbations in the composition can occur if the timescale for transport is short compared with the times for mixing with the surrounding air or significant photochemical production or loss. Perturbations to NO (Randall et al. 2006, 2009), NO2 (Hauchecorne et al. 2007), CO (Jin et al. 2005;Funkeetal.2009), ozone (Smith et al. 2009), and OH (Damiani et al. 2010) have been identified during periods of strong dynamical perturbation. It turns out that the assessment of NOx is complicated because NO can also be produced in situ in the polar mesosphere due to particle precipitation events (Randall et al. 2005). For this reason, NOx is not always a reliable indicator of enhanced transport between the mesosphere and stratosphere. 16 A.K. Smith

4.4 Observed Links with Unknown Mechanisms

For a number of decades, there have been studies reporting unexplained but statistically significant links in the atmosphere. Many of these have dropped out of view as additional data contradicted the previous relationship. The most persistent have been associated with solar variability, particularly the 11-year solar cycle. The first-order atmospheric response to solar variability is fairly straightforward: an increase in the solar flux will increase the heating (assuming absorbing gases or surface properties do not change) and will increase the photolysis rate of some compounds and can therefore affect the composition. The direct heating can be estimated based on satellite measurements of the solar irradiance and knowledge of the absorption as a function of altitude in the atmosphere. The predicted changes in the solar flux penetrating the troposphere due to the solar cycle observed or estimated over the previous century are quite small (IPCC 2007). Likewise, the variability in the solar flux reaching the lower stratosphere is small so little direct response is expected there. Nevertheless, observational analyses show that various climate indicators are well-cor- related with the solar flux. Gray et al. (2010) give a comprehensive overview of the current status of the observations and the search for understanding. One correlation that has received much attention is a relationship between the 11-year solar cycle, the QBO in tropical winds, and the middle and high latitude winter stratosphere. Analysis presented by Labitzke (1987) showed that the NH winter 30 hPa temperature was anticorrelated with the solar flux if one looked only at winters during which the QBO was in its westerly phase. Since then, many years of observations have been added and the correlation still persists (Labitzke 2005). To date, there has been no explanation of the physical mechanism for the observed relationship. The explanation is confounded by several factors. 1) The Holton-Tan relationship (Holton and Tan 1980) establishes a link between the tropical QBO and the high latitude stratospheric temperature. Separat- ing this mechanism from the solar cycle is not straightforward. 2) The two important volcanic eruptions during this period, El Chichon in 1982 and Mt. Pinatubo in 1991, both came near solar maximum. These had an impact on the stratosphere that, with the short duration of available observations, is not possible to separate statistically from the solar cycle. 3) The QBO is asynchronous with the annual cycle. The beating of these two cycles (the QBO at 24–30 months and the annual cycle at 12 months) gives a longer period variation that is difficult to separate statistically from the solar cycle (Salby and Shea 1991; Smith and Matthes 2008). 4) Numerical models have not succeeded in reproducing the observed relationship. The scientific community is divided about these correlation studies. Many researchers dismiss them as being spurious correlations. The thinking behind this is that, if one exam- ines enough time series (there are hundreds of possibilities) over what is, for an 11-year cycle, a fairly short data record, there will be some that show correlations. The possibil- ity of spurious but significant correlations is well-established statistically. The atmosphere has several modes of longer term variability (ENSO, decadal oscillations, etc.) that in some circumstances can be mistaken for a response to the solar cycle if extreme care is not taken. Despite the reservations by many in the community, much effort continues to be ex- pended in looking for additional correlations and trying to explain the correlations that turn up. Interactions Between the Lower, Middle and Upper Atmosphere 17

5 Conclusions

This document gives a brief description of atmospheric regions and some of the processes involved in their interactions. The coupling between regions has received much attention in recent years as both observations have improved and numerical models have provided better tools for analysis. The coupling described in this paper is categorized into three types of processes: radiative, wave propagation, and mass exchange. Vertical coupling occurs readily through radiative processes. Radiative processes can work both ways; the less dense parts of the atmosphere can have an impact on the underlying layers. Several examples were presented for the effect of stratospheric changes (aerosol distribution; ozone loss) on the troposphere. Note however that both of these involve local changes to the stratosphere itself (such as its temperature) and therefore include dynamical as well as radiative processes. Vertical coupling through wave processes also occurs readily. Most of the waves that are important to the middle atmosphere are generated in the troposphere. The conditions for generation of the waves and their propagation through the intervening atmosphere both lead to remote effects. The downward impact of waves was not discussed here. Because of the exponential decrease in wave amplitude with increasing atmospheric density, downward propagation of waves is not expected to be significant. One proposal for downward wave coupling is the impact of the middle atmosphere on the structure of planetary waves in the winter troposphere. For example, recent studies (Perlwitz and Harnik 2003, 2004; Kodera et al. 2006) have shown that, under some circumstances, upward propagating planetary waves can be reflected in the stratosphere. This reflection has an impact not only on the wave structure in the stratosphere but also on the wave within the troposphere. Coupling of atmospheric regions through mass exchange is also possible. In this case, we focus on the concentrations of trace chemical species that are perturbed due to changes in the transport. The examples are water transport into the stratosphere and mesospheric air transported into the NH polar winter stratosphere. This overview gives a description of the regions of the atmosphere and a taste of some of the important coupling processes between them. This survey is not comprehensive but instead reflects the experience and interests of the author. The knowledge of the atmosphere and its interactions is ever increasing due to increased measurements and improved numerical modeling capabilities. These will lead to better understanding of the interactions and better predictions of how they will affect the future of the atmosphere.

Acknowledgements The National Center for Atmospheric Research is sponsored by the National Science Foundation. Support for this work was also provided by the NASA Heliospheric Guest Investigator Program.

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