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Interactions Between the Lower, Middle and Upper Atmosphere Space Sci Rev (2012) 168:1–21 DOI 10.1007/s11214-011-9791-y Interactions Between the Lower, Middle and Upper Atmosphere 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 meso- sphere 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 sep- arates the troposphere from the middle atmosphere, and the mesopause, which separates the middle atmosphere from the thermosphere. The survey of coupling processes looks at ra- diative 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 tran- sitions 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 approxi- mately 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 re- flection 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 par- ticles 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 im- portant 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, par- ticularly 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 there- fore 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 re- lation, 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, tem- perature, 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 distur- bances 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.
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