Dynamics of the Thermosphere
Jeffrey M. Forbes
Department of Aerospace Engineering Sciences
University of Colorado, Boulder, CO 80309, USA
Prepared for Submission to
Journal of the Meteorological Society of Japan
Special Anniversary Issue
November 2006; revised February, 2007
Abstract
Major advances in our understanding of the dynamics of Earth’s thermosphere (≈ 90500 km) during the past 25 years are reviewed. Since the thermosphere is primarily an
externally-forced system, a broad overview of the energy input, conversion and transport
mechanisms in the ionosphere-thermosphere system is first provided. This serves as
background and context for the non-specialist. Then, several broad areas of progress are
in turn discussed in some detail: (i) the role of solar thermal tides in imposing significant
longitudinal variability in the lower thermosphere (≈ 100-150 km), and affecting the zonal
mean circulation at these altitudes; (ii) the zonal mean circulation of the thermosphere, the
changes in O and N2 relative densities that accompany it, and the competing roles of solar
radiative heating and Joule (ohmic) heating in determining the overall structure of this
circulation; (iii) polar and auroral thermosphere dynamics, and connections to relevant
magnetosphere and ionosphere processes; and (iv) the global response to geomagnetic
disturbances, i.e., relatively sudden injections of energy and momentum from the
magnetosphere. The paper concludes with a personal assessment of future research
directions and scientific questions that remain to be addressed in forthcoming decades.
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1. Introduction
1.1 Scope of this Review
The purpose of this paper is to provide an overview of the neutral dynamics of the
thermosphere, defined here to be the region from about 90 to 500 km, and hereafter
referred to as the IT (Ionosphere-Thermosphere) system. The intended audience includes
dynamicists who study the lower regions of Earth’s atmosphere, and chemical aeronomers
and ionospheric physicists who desire broad exposure to the fundamentals. Consequently,
emphasis is placed on the features of the IT system that distinguish it from lower regions of
the atmosphere, and on the most fundamental concepts and processes without focusing
too strongly on details, or on phenomena that do not have global or broad impacts. In
keeping with the spirit of this special anniversary issue of the Journal of the Meteorological
Society of Japan, this paper particularly highlights those major advances achieved during
the past 25 years. Thus, the scope of the review is further limited, mostly, to those topics
whose maturity has developed over decades; hence some very interesting but more
focused contributions are not mentioned, as they would be in a different type of review. No
attempt is made here to provide a complete compilation of references; in fact, only the most
recent or relevant references that provide a pathway to earlier works are cited. In addition,
the selection of topics and scientific advances that are cited as most significant represent,
in part, biases of the author. For these limitations I apologize to readers and colleagues
who disagree with my choices.
This review is organized as follows. In Section 1, a broad overview of the IT system is
provided. Subsequent sections review those areas of thermosphere dynamics that reflect
the most significant advances over the past 2-3 decades. The first of these, the role of
solar thermal tides in imposing significant longitudinal variability in the lower thermosphere
(≈ 100-150 km), and in affecting the zonal mean circulation through tidal dissipation, is
described in Section 2. Section 3 deals with the zonal mean circulation of the
thermosphere, the accompanying changes in O and N2 densities, and the competing roles
of solar radiative heating and Joule (ohmic) heating in determining the overall structure of
this circulation. Section 4 summarizes our current knowledge of polar and auroral
thermosphere dynamics, and connections to relevant magnetosphere and ionosphere
processes. Section 5 deals with the global response to geomagnetic disturbances, i.e.,
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relatively sudden injections of energy and momentum from the magnetosphere. Finally, in
Section 6, a personal assessment of future research directions and remaining scientific
questions concerning thermosphere dynamics is provided.
1.2 Overview of the IT System
It is assumed that the reader has a basic knowledge of the thermal structure of the
thermosphere, its connections with variations in EUV radiation emanating from the Sun,
and the tendency for the main chemical constituents (N2, O2, O, He and H) to diffusively
separate according to their individual scale heights. In addition, in contrast to the lower
atmosphere, the thermosphere is an externally-forced deterministic system wherein
instabilities are strongly suppressed due to the dominance of molecular diffusion.
Therefore, to provide ample background for the following, we begin with an overview of the
various sources of external forcing to the thermosphere, as well as mechanisms for
transforming and removing the input energy from the system. Referring to Figure 1, energy
inputs from above and below the thermosphere are now considered in turn.
In addition to solar energy in the form of EUV radiation, the thermosphere is affected by
particle energy from the Sun in the form of the solar wind in the following way. The highlatitude ionosphere is electrically connected to the outer nightside magnetosphere (i.e.,
~10-100 Earth radii), by highly conductive magnetic field lines. When the interplanetary
magnetic field (IMF) in the solar wind has a negative Bz component (i.e., component of IMF
pointing downward in the ecliptic plane) energy is transferred from the solar wind to the
magnetosphere with predictable consequences. A plasma convection system is set up in
the magnetosphere that maps into the high-latitude ionosphere along conducting field lines,
and sets the neutral atmosphere into motion via ion-neutral collisions. Plasma particles are
energized on the convecting field lines, and a population of these precipitate into the highlatitude thermosphere, create the aurora, and also increase the conductivity of the
ionosphere through impact ionization of the neutral gas. This enhanced conductivity
facilitates the flow of current between the magnetosphere and ionosphere, with the net
effect of dissipating magnetospheric energy in the form of ohmic heating due to the
resistivity of the ionosphere. This process is more commonly called Joule dissipation of
ionospheric currents, and can significantly modify the energetics and dynamics of the
global IT system.
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Figure 1. Energy input, conversion and transport processes relevant to the IonosphereThermosphere (IT) system. Green indicates energy sources from above the thermosphere,
blue indicates influences of lower atmospheric regions on the thermosphere, and red indicates
energy conversion and transport processes within the thermosphere.
Energetic particle precipitation provides a second heat source to the high-latitude
thermosphere, but it plays the more important role of producing nitric oxide (NO) through a
series of chemical reactions between atomic oxygen and nitrogen species that are excited
to higher internal states through collisions by the precipitating particles (Gerard and Barth,
1977). NO is an efficient radiative cooler for the thermosphere, and serves to regulate the
magnitude of the thermosphere response to energy inputs of magnetospheric origin, as
well as recovery to pre-disturbed levels (Mlynczak et al., 2005). NO can also be
transported downward from the thermosphere, and ultimately reduce stratospheric ozone
levels through catalytic chemical reactions (Randall et al., 2005).
Waves often represent an important mechanism for transporting energy and momentum
from one point to another in an atmosphere. Gravity or buoyancy waves are excited in the
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lower atmosphere by flow over topography, convective activity, and shear instabilities.
Periodic absorption of solar radiation forces thermal tides at sub-harmonics of a solar day.
Longer-period waves can be excited by instabilities in the mean flow, by temporal variations
in convective activity (latent heating), and sometimes arise as resonant atmospheric
oscillations (normal modes). Many of the above waves are capable of propagating or
penetrating to higher altitudes where they undergo dissipation and deposit heat and
momentum into the mean flow.
1.3 Fundamentals of the Solar-Driven Circulation
The simple solar-driven undisturbed circulation of the IT system is now described.
Consider Figure 2, wherein temperature contours and wind vectors are plotted as a
function of latitude and longitude at 0000 UT near 300 km. These results correspond to
average solar cycle and quiet geomagnetic conditions. At high latitudes, the flow is
influenced by momentum transfer from the convecting ions to the neutrals, and winds attain
amplitudes up to ≈ 500 ms-1. At middle and low latitudes, the flow is less intense (≈ 50-150
ms-1) and notably tends to be across the isobars (or equivalently, isotherms). This stands
in contrast to the familiar geostrophic flow in the lower and middle atmosphere, where
balance between pressure gradient and Coriolis forces leads to flow along isobars. To
understand this, consider the following simplified horizontal momentum equation for the
thermosphere, where nonlinear terms, horizontal diffusion and ion drifts are neglected:


 
 µ ∂ 2U
∂U
1
+ 2Ω × U = − ∇P − ν niU +
(1)
∂t
ρ
ρ ∂z 2


where U = horizontal wind vector, Ω = Earth’s angular velocity, P = pressure, ν ni =
neutral-ion collision frequency for momentum transfer, z = altitude, and µ is the coefficient
of molecular viscosity. The last term represents vertical diffusion of horizontal momentum,
which is very fast in the thermosphere due to its inverse dependence on density. For typical
thermosphere conditions, the neutral-ion collision term in (1) is significantly larger than the
Coriolis term, so that deflection due to the Coriolis force does not dominate the flow.
Instead, the pressure gradient and ion-neutral collision terms more nearly balance to
produce cross-isobaric flow, while the main role of molecular diffusion is to strongly inhibit
vertical gradients in the wind field.
Note also that day-night temperature differences at low latitudes in Figure 2 are about
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200 K and peak near 15:30 h local time. Dickinson and Geisler (1968) showed that vertical
winds play an important role in determining the amplitude and phase of the temperature
response with respect to EUV heating, which maximizes at noon at the equator for equinox
conditions. By considering conservation of mass above a constant pressure surface, po,
they showed that the total vertical wind ( w ) may be expressed in terms of two components,
the barometric wind ( wB ) and the divergence wind ( wD ). The former is due to thermal
expansion/contraction of the atmosphere, and the latter is caused by diverging
Figure 2. Temperature contours (K) and wind vectors as a function of latitude
and longitude at 0000 UT near 300 km, corresponding to average solar cycle
and quiet geomagnetic conditions. (Rishbeth et al., 2000).
horizontal winds and the conservation of mass:
∞

1
⎛ ∂h ⎞
w = wB + wD = ⎜ ⎟ + ∇ ⋅ ∫ ρUdz
⎝ ∂t ⎠ p ρ h( p )
0
(2)
where h = height of a given pressure surface, P = pressure, ρ = total mass density, z =

1 Dp
altitude, and U is the horizontal wind vector. They also showed that wD = −
ρ g Dt
(vertical motion of air with respect to a pressure surface) where g is the acceleration due to
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gravity and D is the advective derivative, so that the thermodynamic equation
cp
DT
1 Dp
−
= J (where J is net heating from all sources, including conduction, T =
Dt
ρ Dt
temperature and cP is the specific heat at constant pressure) becomes c p
DT
+ gwD = J .
Dt
To a first approximation, to explain vertical structure, one can linearize about a mean basic
state dependent only on height (denoted by an overbar) so that
DT ∂T
∂T
∂T
J
≈
+ wD

= − wD Γ ,
Dt
∂t
∂z
∂t c p
where Γ =
(3)
∂T g
is the mean atmospheric stability. The adiabatic cooling term − wD Γ in
+
∂z c p
(3) serves as an additional “heat source” that is as important as EUV heating, and accounts
for much of the observed time shift of the temperature maximum to late afternoon with
respect to the noontime maximum in EUV heating.
2. Nonmigrating Tides
The global temperature, density and wind fields induced by the daily cyclic absorption of
solar energy in an atmosphere are referred to as solar thermal tides. Particularly within the
mesosphere-lower thermosphere (MLT) region (≈ 80-120 km), solar thermal tides dominate
the atmospheric dynamics; and, prior to the last ten years, the general perception has been
that solar tides were more or less longitude-independent. More recently, satellite-based
observations, increased knowledge of coupling processes with the troposphere, and
consideration of nonlinear effects has changed that view. The following contains a
summary of these accomplishments, beginning with a review of atmospheric tide
nomenclature.
Assuming continuity in space and time around a latitude circle, solar thermal tidal fields
are represented in the form
An,s cos( nΩt + sλ − φ n,s )
(4)
where t = time (days), Ω = rotation rate of the earth = 2π day-1, λ = longitude, n (= 1, 2, ...)
denotes a subharmonic of a solar day, s ( = .... -3, -2, ...0, 1, 2, ....) is the zonal
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wavenumber, and the amplitude An,s and phase φn,s are functions of height and latitude. In
this context, n = 1, 2, 3 represent oscillations with periods corresponding to 24 hours, 12
hours, 8 hours, and hence are referred to as diurnal, semidiurnal and terdiurnal tides,
respectively. Eastward (westward) propagation corresponds to s < 0 (s > 0). Phase is
defined as the time of maximum at zero longitude; in other words, the local time at
Greenwich. At any height and latitude the total tidal response is obtained as a sum over n
and s.
Rewriting (4) in terms of local time tLT = t + λ/Ω, we have
An, s cos ⎡⎣ nΩt LT + ( s − n ) λ − φn, s ⎤⎦
(5)
Solar radiation absorption by a zonally-symmetric atmosphere or surface yields daily (local
time) variations that are independent of longitude, i.e., s = n. From (4) such components
correspond to a zonal phase speed Cph = dλ/dt = - nΩ/s = -Ω, in other words westwardpropagating at the same speed as the apparent motion or ``migration'' of the Sun to a
ground-based observer. These sun-synchronous tidal components are referred to as
migrating tides.
Now consider the cyclic heating due to absorption of solar energy by a zonally
asymmetric (longitude-dependent) planetary atmosphere or surface. In response to this
heating, the local time structure of the atmosphere (at a given height and latitude) is
dependent on longitude. A common approach is to examine zonal wavenumber
components of the lowest-order local time harmonics (n = 1,2,3) that combine to give rise to
the salient features of this longitude dependence. In this case, Fourier representation of
each harmonic must involve a range of zonal wavenumbers of both sign, corresponding to
waves propagating to the east (s < 0) or west (s > 0) (Chapman and Lindzen 1970). This
approach offers the opportunity to relate results to tidal theory and numerical models, and
often to gain physical insight.
It is well known that latent heating associated with deep tropical convection possesses
strong variations with U.T., longitude, latitude and season. Studies (e.g., Hagan and
Forbes, 2002, 2003) demonstrate that this source of excitation leads to diurnal and
semidiurnal tides over a spectrum of zonal wavenumbers that propagate into the MLT and
achieve significant amplitudes in this height regime. These waves superimpose to yield
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different local time variabilies at different longitudes. Of particular interest is the wave-4
structure (i.e., four maxima) with respect to longitude that characterizes both the diurnal
and semidiurnal total tidal fields (Zhang et al., 2006). This feature is a result of the
predominant wave-4 topography/land-sea longitude dependence at the surface, which is
reflected in the diurnal and semidiurnal components of the latent heating rates due to deep
tropical convection (for details see e.g., Williams and Avery, 1996; Forbes et al., 2001;
Zhang et al., 2006). An example is shown in Figure 3, where longitude variations in the
diurnal and semidiurnal tides reflect wave-4 structures. For the diurnal tide, this mainly
arises due to constructive and destructive interference between the migrating diurnal tide (s
= 1) and the eastward-propagating diurnal tide with s = -3. In the case of the semidiurnal
tide, the predominant contributors are the migrating (s = 2) component, the eastward
propagating component with s = -2, and the westward-propagating component with s = 6.
The wave-4 structure has even been discovered in ionosphere plasma densities near 350400 km (Immel et al., 2006), which is presumably a manifestation of electric fields produced
by dynamo action of the above nonmigrating tides in the 100-170 km altitude region.
Figure 3. Latitude versus longitude distributions of tidal temperature amplitudes derived from
measurements by the SABER instrument on the TIMED spacecraft. Left: Diurnal tide at 88
km, 120-day mean centered on day 267 of 2004. Right: Semidiurnal tide at 110 km, 120-day
mean centered on day 115 of 2004. From Zhang et al. (2006).
Much evidence now exists that supports nonlinear wave-wave interactions as another
important source of nonmigrating tides. The mechanism works as follows (Teitelbaum and
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Vial, 1991). Given two oscillations with respective frequency-zonal wavenumber pairs (σ1,
s1) and (σ2, s2), and under the assumption of a so-called quadratic interaction between
these two ‘primary waves’, and neglecting self-interactions, “sum and difference”
secondary waves are generated with the frequency, zonal wavenumber pairs (σ1 + σ2, s1 +
s2) and (σ1 – σ2, s1 – s2). Teitelbaum and Vial (1991) invoked this mechanism as a
secondary means (beyond direct solar heating) of exciting the migrating terdiurnal tide (n =
3, s = 3), via interaction between the migrating diurnal (n = 1, s = 1) and semidiurnal (n = 2,
s = 2) tides. This mechanisms for exciting the terdiurnal tide has recently been studied
further by e.g., Akmaev (2001), Smith and Ortland (2001). Wave-wave nonlinear interaction
was first suggested by Bernard (1981) in order to explain longitudinal asymmetries in the
semidiurnal tide, and invoked by Forbes et al. (1995) to explain observation of the
semidiurnal tide with s = 1 over South Pole, both within the 90-100 km height regime.
Figure 4. Height versus latitude distributions of annual-mean meridional wind
diurnal tidal amplitudes derived from measurements by the TIDI instrument
on the TIMED spacecraft for s = 0 (top) and s = 2 (bottom). From Oberheide
et al. (2005).
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Recent modeling work in fact indicates that nonlinear interactions between the
stationary planetary wave with s = 1 and migrating tides lead to significant nonmigrating
diurnal and semidiurnal tidal signatures above about 80 km altitude (Hagan and Roble,
2001; Yamashita et al., 2002; Angelats i Coll and Forbes, 2002; Lieberman et al., 2004;
Grieger et al., 2004; Oberheide et al., 2002). The diurnal tidal components generated by
this mechanism are the westward-propagating diurnal tide with s = 2, and the zonallysymmetric (s = 0) diurnal oscillation. For the semidiurnal tide, the generated waves are the
westward-propagating components with s = 1 and s = 3. All of these waves have been
observed in recent satellite data analyses (Talaat and Lieberman, 1999; Forbes et al.,
2003; Angelats i Coll and Forbes, 2002; Forbes and Wu, 2006; Zhang et al., 2006;
Oberheide and Gusev, 2002; Oberheide et al., 2006). An example is provided in Figure 4,
which illustrates height versus latitude distributions of annual mean amplitudes of the s = 0
and s = 2 diurnal variations in meridional wind. The maximum values for each oscillation
are of order 8-10 ms-1, which can be compared with an annual mean amplitude for the
migrating tide of order 40 ms-1. Therefore, the interference effects of these waves, plus
other nonmigrating tides not illustrated here, give rise to quite substantial longitudinal
variability in the tidal amplitudes (cf. Yoshikawa and Miyahara, 2003).
Building upon the pioneering work of Miyahara (1978), more recent works have also
established that dissipation of both diurnal and semidiurnal migrating and nonmigrating
tides significantly affect the zonal mean circulation of the lower thermosphere, from about
100 to 150 km (e.g., Miyahara and Wu, 1989; Angelats i Coll and Forbes, 2002; Yoshikawa
and Miyahara, 2003; Forbes et al., 2006; Miyoshi, 2006). Figure 5 shows an example,
illustrating the relative importance of migrating and nonmigrating semidiurnal tidal
dissipation on the zonal mean flow. Zonal mean winds of similar magnitudes are produced
by dissipation of diurnal tidal oscillations and low latitudes, and are likely to exhibit temporal
variability reflective of convective activity in the tropics (Miyoshi, 2006).
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Figure 5. Zonal mean zonal (left) and meridional (right) winds induced by dissipation of
semidiurnal tides for the month of January as calculated in the modeling work of
Angelats i Coll and Forbes (2002). The top panels include effects due to both migrating
(s = 2) and nonmigrating (s = 1 and s = 3) tides. The middle panel illustrates
contributions due to the migrating component alone, and the bottom panel reveals the
contributions of the nonmigrating tides to results in the top panel. Contour intervals are
10 ms-1 for the zonal component, and 5 ms-1 for the meridional one.
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3. Zonal Mean Circulation and O-N2 Composition
The zonal-mean meridional circulation of the thermosphere driven by differential solar
heating is from summer to winter with upwelling in the summer hemisphere and
downwelling in the winter hemisphere. The circulation is closed by a weak return flow in
the lower thermosphere to maintain continuity of mass flow. This circulation system can
have an important effect on the distribution of major chemical species in the thermosphere.
The thermosphere composition primarily consists of O and N2 between about 120 and 500
km. The ratio of O to N2 densities (R[O/N2]) is important for the ionosphere, since photoionization of O is an important source of plasma, and molecular species such as N2 and O2
control the loss of ionosphere plasma. Processes such as lower thermosphere heating and
upwelling can carry N2-rich air to higher altitudes, drive the thermosphere from its diffusive
equilibrium state, and in addition enhance the loss of ionosphere plasma. In addition, the
N2-rich air can be transported by horizontal winds, thus affecting latitude regions outside
the heating zone. The upwelling referred to above is the divergence component of the
vertical wind field discussed in Section 1. The barometric part of the wind field is
connected with thermal expansion and contraction of the atmosphere, and does not
transport mass across pressure levels; upward divergent winds transport molecular-rich air
from lower to higher altitudes and decrease R[O/N2] aloft (e.g., Rishbeth et al., 1987;
Rishbeth and Müller-Wodarg, 1999).
The above principles can be used to understand changes in R[O/N2] with respect to
season, latitude, local time, and level of magnetic activity (e.g., Rishbeth et al., 1987;
Rishbeth and Müller-Wodarg, 1999; Rishbeth et al., 2004). To illustrate the effects of
vertical and horizontal winds on R[O/N2], consider Figure 6, applicable to the solar EUVdriven zonal mean circulation. Upwelling occurs in the summer hemisphere, transporting
N2 upward, reducing R[O/N2] and increasing the mean molecular weight. (This reduction in
R[O/N2] occurs despite the fact that production of O by photo-dissociation of O2 is greatest
in the summer hemisphere.) The atmosphere is not in diffusive equilibrium and tends
towards the mixed state. The molecular-rich thermosphere gas at upper levels is
transported into the winter hemisphere, and in the absence of significant upwelling,
diffusive balance is progressively restored. The diffusion time constant varies inversely
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with density, so diffusive separation is first restored at upper levels, and then after ~3-5
days is restored in the lower thermosphere at winter latitudes.
Figure 6. Schematic illustrating the zonal mean meridional circulation driven by differential
solar heating in the thermosphere (blue arrows), the transport of O and N2 (labeled arrows),
the latitudinal variation of R[O/N2], and a rough depiction of boundary between mixed and
diffusively-separated O-N2 composition with estimates of diffusive time constant (red).
(Figure inspired by Rishbeth et al., 2000).
Figure 7 illustrates how the above picture changes when high-latitude heating
processes are included. In the summer hemisphere, upwelling and equatorward transport
are assisted by the additional heat source, while a secondary circulation cell exists in the
winter hemisphere due to the upwelling and equatorward transport due to auroral/polar
heating. This can cause a local maximum in R[O/N2] near the boundary of these circulation
cells that is a function of geographic longitude since the high-latitude heating is ordered in
magnetic coordinates. Since the daytime peak electron density in the F-region tends to
vary with R[O/N2], the above arguments can be used to explain the first-order annual,
14
semiannual and longitudinal variations of the ionosphere at extra-tropical latitudes (e.g.,
Rishbeth et al., 2000; Zou et al., 2000).
Figure 7. Same as Figure 6, with effects of high-latitude heating due to magnetospheric
coupling processes.
For the solstitial conditions discussed above, the increased mean mass over regions of
upwelling and beyond reduces the total density scale height at a given altitude, essentially
“mixing” and “compressing” the atmosphere. Fuller-Rowell (1998) in fact likens this
circulation to one large turbulent eddy due to a huge “thermospheric spoon” that mixes the
thermosphere to make it more molecular than it otherwise would be. At equinox horizontal
pressure gradients are much weaker, since solar heating is more uniformly distributed and
balanced by high-latitude heating; the net result is weak prevailing meridional winds. The
equinoctial thermosphere is less mixed, and more nearly in diffusive equilibrium. The net
result is a semiannual variation in total thermosphere density, with maxima around the
solstices. This combined with composition variations in the lower thermosphere, possibly
connected with variations in gravity wave filtering and turbopause altitude (Maeda et al.,
15
1986; Fuller-Rowell and rees, 1992) appear to account for the observed semiannual
variation of thermosphere density. However, the increase in amplitude of the observed
semiannual variation with level of solar activity and its inter-annual variability (Bowman,
2004) remain unexplained.
Since R[O/N2] varies with height, it has been found convenient and illusory to define a
so-called compositional “P-parameter” (Rishbeth et al., 1987; Rishbeth and Muller-Wodarg,
1999). Under conditions of diffusive equilibrium, it can be shown that P ≈ 28 ln[O]-16ln[N2]
is nearly constant with height. Positive departures of P from this value correspond to
regions of enhanced heating and upwelling, whereas regions of depressed P are produced
by downwelling. In addition, the height gradient of P provides information on the differential
flow velocities of O and N2 that act to restore diffusive equilibrium. The P parameter has
been employed very successfully to diagnose both observations and thermosphere general
circulation models during both quiescent and magnetically active conditions (e.g., Rishbeth
and Muller-Wodarg, 1999; Rishbeth et al., 1987).
4. Polar-Auroral Circulation
Over the past 20-30 years our understanding of the high-latitude circulation system of
the thermosphere and its relationship with the magnetosphere and the interplanetary
magnetic field (IMF) has matured considerably. Observations from the DE-2 spacecraft
(Killeen and Roble, 1988) and complementary theory and modeling efforts are major
contributors to our current level of knowledge. Much of the research in this context was
devoted to an understanding of the relative importance of the EUV-driven circulation,
effects induced by Joule and particle heating, and that part of the circulation driven by
momentum transfer from the convecting ions (e.g., Killeen and Roble, 1988; Killeen et al.,
1992; Thayer and Killeen, 1991, 1993). Observations in combination with thermosphereionosphere general circulation models have been used to elucidate the morphology and
physics, often decomposing the flow field into nondivergent (irrotational) and divergent
(rotational) components and examining their relative behaviors and influences. In
particular, the neutral atmosphere circulations in the dawn and dusk sectors were found be
non-symmetric (in contrast to the salient features of the convecting ions), and to vary
considerably with altitude (e.g., Gundlach et al., 1988; McCormac et al., 1987; Fuller-
16
Rowell and Rees, 1984; Fuller-Rowell, 1995). In addition, the total mass density variations
that accompany these circulation systems have been a focus of interest (Crowley et al.,
1996; Schoendorf et al., 1996a,b). In the following, our current knowledge of these aspects
of the high-latitude circulation is described. The reader is cautioned that the described
circulation features represent a first-order view, and that there exists a whole body of
literature devoted to various nuances of the circulation system to be described.
Figure 8. Reconstituted wind fields from combined DE-2/TIGCM wind data in
geomagnetic coordinates for Northern High Latitudes for December solstice
conditions during periods of (a) Kp ≤ 3 and (b) 3+ ≤ Kp ≤ 6 (Thayer and
Killeen, 1993).
17
To expose the salient dynamical patterns, Thayer and Killeen (1993) merged averaged
DE-2 observations and TIGCM solutions (where observations were sparse) to arrive at
mean circulation characteristics for low (Kp ≤ 3) moderate (3+ ≤ Kp ≤ 6) magnetic activity
conditions. These are shown in Figure 8, and correspond to December solstice conditions
in the N. Hemisphere. The effects of convecting ions are clearly seen for both quiet and
moderate magnetic activity; the ions drag the neutrals in the anti-sunward direction over the
pole, with return flows in the dawn and dusk sectors above about 50-60° latitude. The
generally anti-sunward solar-driven circulation reinforces the convection-driven flow over
the pole, but impedes the convection-driven return flows in the dawn and dusk sectors. In
addition, the dusk circulation cell is more intense and well-defined than the dawn cell.
It is illustrative to examine the nondivergent and irrotational components of the above
wind field. Figure 9 illustrates these flow components with their respective stream function
and velocity potential contours, for moderate magnetic activity conditions. The nondivergent
wind component flows parallel to stream function contours, and increases in magnitude as
the spacing between contours decreases; it represents the sole source of vorticity to the
total wind field. The divergent wind component flows perpendicular to contours of constant
velocity potential, and increases in magnitude as the distance between contours
decreases. For the nondivergent component of the circulation in Figure 9a, the dawn and
dusk cells are now readily apparent, with the dusk cell characterized by tighter stream
function contours and thus higher wind speeds; sunward winds are weaker for the dawn
cell, which gives rise to the asymmetry in the wind pattern in Figure 8b. Momentum
transfer by the convecting ions is the dominating force driving nondivergent winds at high
latitudes, with Coriolis, advective and viscous forces playing secondary roles.
Figure 9b illustrates the irrotational wind component, which is characterized by flow in
the general anti-sunward direction over the polar cap. This flow primarily reflects the solaror EUV-driven circulation, with some additional intensification provided by Joule heating for
this level of magnetic activity. This divergent flow is mainly determined by a balance
between the pressure gradient force and the non-convective ion drag force, with the
Coriolis force secondary in importance.
It is also noted (Thayer and Killeen, 1991) that the vorticity of the high-latitude wind field
is several times greater than that of the divergence, but that this difference diminishes as
18
geomagnetic activity increases. This is possibly due to the fact that the momentum driving
of the neutral wind is proportional to the electric field E mapped from the magnetosphere,
whereas the joule heating, which drives a divergent wind field, is proportional to the square
of E.
Figure 9. Derived (a) nondivergent wind component and stream function
contours at an interval of 1 x 108 m2s-1 and (b) irrotational wind component
and velocity potential contours at and interval of 1 x 108 m2s-1 corresponding
to the reconstituted wind data in Figure 8b for 3+ ≤ Kp ≤ 6 (Thayer and
Killeen, 1993).
19
The tendency for a stronger and more persistent dusk cell (cf. Figure 9a) has been
described in terms of ‘inertial resonance’ (Fuller-Rowell and Rees, 1984; Fuller-Rowell et
al., 1994; Fuller-Rowell, 1995), analogous to inertial oscillations known to exist in the
oceans. In this case, the nonlinear curvature and Coriolis terms in the momentum
equations nearly balance, and in the absence of dissipation a vortex flow is self-sustaining,
and can persist for many hours. For this balance to occur, the Rossby number Ro = U fR
must be sufficiently large where U is a typical wind speed for the vortex, f is the Coriolis
parameter and R is the radius of the vortex. In addition, the flow must be clockwise in the
N. Hemisphere, and counter-clockwise in the S. Hemisphere. These conditions, i.e.,
sufficiently large vortex speeds at a sufficiently small radius, are met for the dusk cell driven
by ion convection. The extent to which this circulation extends down to the lower
thermosphere, i.e. ≈ 120 km, depends on whether ion densities there are sufficiently large
to transfer momentum from the convecting ions to the neutrals. Often this condition is met,
since precipitating energetic particles ionize this region of the atmosphere under
geomagnetically disturbed conditions. When magnetospheric forcing is suddenly
terminated, the vortex at higher regions of the thermosphere (≥ 300 km) dissipates within a
few hours due to molecular viscous effects (which increase inversely to the total mass
density). At altitudes near 120 km, where viscous effects are much smaller, the vortex
persists for up to ~12 hours in numerical simulations (Fuller-Rowell and Rees, 1984),
although only sparse experimental verification exists to support this prediction.
Neutral density cells associated with the above vortices have been discovered in
thermosphere general circulation simulations and in density measurements obtained by
satellite accelerometers (Crowley et al., 1996; Schoendorf and Crowley, 1996a,b), under
magnetically active conditions. At altitudes above about 170-200 km, the density structure
consists of low-density cells near dawn and dusk and high-density cells near noon and
midnight; at 200 km density variations from low to high cell are of order 40%. The low
density in the dawn cell is expected, as low pressure (density) systems are normally
associated within cyclonic circulations, i.e., counterclockwise flows in the N. Hemisphere.
However, the low densities associated with the anti-cyclonic dusk vortex were unexpected,
and have been explained in terms of “anomalous antibaric flow’ (Schoendorf et al., 1996b).
These authors determined the conditions for what is called an antibaric or anomalous low in
20
the meteorological context by following the conventional meteorological derivation except
with the addition of ion drag. Through manipulation of the momentum equations they
derive a pseudo Rossby number
RoP = Ro +
ν ni Vir
f U
(6)
where Vir is the ion speed in the radial direction and ν ni is the neutral-ion collision frequency.
In order for anticyclonic flow to exist around a pressure or density low, the following
condition must exist: RoP < −1 . Note that although the neutral wind flow U is assumed to be
in the tangential direction around the vortex, it is still acted upon by ions moving in the
radial direction, and has a divergent component. From (6), whether a density depression
exists within the dusk circulation cell depends on the relative magnitudes of ν ni , Vir and U,
and therefore on altitude and level of geomagnetic activity. Generally, for lower wind
speeds, i.e., quiet magnetic conditions at all altitudes and low altitudes (i.e., near 120 km)
during all geomagnetic conditions, an anomalous low does not exist, and high relative
densities exist within the anticyclonic vortex at dusk. However, for sufficiently high
magnetic activity and wind speeds RoP < −1 , and the dusk cell is characterized by a
depression in density. Under active magnetic conditions, joule heating tends to produce
density increases all over the polar cap. The vortex effects noted above generate density
depressions in the dawn and dusk sectors contrasting against density increases in the
noon and midnight sectors which then give the appearance of a 4-cell structure. At 120
km, below the peak level of Joule heating, only the dawn and dusk density cells appear,
with the characteristic low and high densities expected for “normal” cyclonic and
anticyclonic flow, respectively.
More recent studies (e.g., Richmond et al., 2003) have continued to examine the
behaviors of the above vortices with respect to variations in the horizontal component of the
IMF, By, among other aspects of the high-latitude circulation system, but these detailed
aspects of the dynamics are out of the scope of the present review.
5. Global Response to Geomagnetic Disturbances
The previous two sections address the response of the thermosphere to enhanced
levels of geomagnetic activity in a statistical sense. However, quite often energy and
21
momentum inputs from the magnetosphere occur rather precipitously, and induce a
transient response that spreads globally from the polar-auroral region. In this section, we
review our current knowledge of the global response, with emphasis on developments
during the past 20-25 years.
Our early knowledge of the global response to magnetic storms and substorms was
based on total mass densities derived from the orbital decay of satellites, followed by
composition measurements by neutral mass spectrometers on the Ogo, Esro-4, Aeros,
and Atmosphere Explorer satellites (see review by Prölss, 1980). More recently,
satellite-borne accelerometers have elucidated the thermosphere response to
geomagnetic storms [e.g., Berger and Barlier, 1981; Bruinsma et al., 2006; Forbes et
al., 1987, 1996, 2005; Liu and Lühr, 2005; Sutton et al., 2005]. In comparison, very few
measurements exist of the global response of the neutral wind system (e.g., Wu et al.,
1994; Marcos and Forbes, 1985; Killeen and Roble, 1988; Emmert et al., 2002; Sutton
et al., 2005). Observational data of the neutral thermosphere response to geomagnetic
storms, especially below 200 km, therefore remains rather meager. However, the data
that do exist have been sufficient to spawn significant advances in first-principles
modeling efforts (e.g., Fuller-Rowell et al., 1994, 1996; Burns et al., 1991, 2004;
Crowley et al., 2006; Fujiwara et al., 1996). These models show that the neutral
atmosphere response is dependent upon a complexity of interdependent processes
involving electrodynamics, mutual coupling between the neutral and plasma species,
chemistry and radiative cooling. Nevertheless, simpler models have been used to
elucidate the wavelike aspects of the thermosphere response to impulsive heating (e.g.,
Richmond and Matsushita, 1975; Richmond, 1979; Brinkman et al., 1992; Mayr et al.,
1990), and are able to explain some of the salient features of wave-like structures seen
in satellite data (e.g., Gross et al., 1984; Forbes et al., 1995, 2005; Bruinsma et al.,
2006). Our current general perception of the magnetic storm response of the
thermosphere based upon these and other modeling efforts noted below, as well as
recent observations, is now described.
In response to a sudden increase in high-latitude heating, a traveling atmospheric
disturbance (TAD) or “surge” is launched that travels towards the equator with a speed of
order 600-750 ms-1 near 400 km, but perhaps as low as half this speed at 200 km. The
22
disturbance consists of a spectrum of gravity waves that are filtered out such that the
apparent period of the disturbance varies with height and distance from the source. Typical
horizontal wavelengths are of order 1000-3000 km. Meridional winds of order 100-200 ms-1
and temperatures of order 50-100 K characterize the disturbance, which can propagate into
the opposite hemisphere. An illustration of a TAD as measured by the accelerometer on
the CHAMP satellite (e.g., Bruinsma et al., 2004) is provided in Figure 10.
Figure 10. Illustration of Traveling Atmospheric Disturbances (TADs) seen in densities
near 400 km near local noon measured by the accelerometer on the CHAMP satellite, in
connection with a geomagnetic disturbance on day 308 of 2003. The data are obtained at
nearly constant longitudes about every 1.5 hours, i.e., the time between consecutive
orbits. The disturbance was initiated between orbits 6 and 7, and takes roughly 4.5 hours
to reach the equator from both polar/auroral regions. The disturbances appear to
penetrate into opposite hemispheres from their origins. Figure courtesy of Dr. Sean
Bruinsma, CNES.
TADs due to energy injection at high latitudes have also been successfully simulated in
first principles thermosphere models (Fujiwara et al., 1996; Balthazor and Moffett, 1999;
Fujiwara and Miyoshi, 2006). TADs are a different manifestation of the same phenomenon
as traveling ionospheric disturbances (TIDs), and both have been interpreted in terms of
23
atmospheric gravity waves (see review by Hocke and Schlegel, 1996). Observations of
TIDs are more plentiful due to the global availability of ionospheric measurements, and thus
can provide important information on TADs. Recent observations indicate that TIDs can
also exist during magnetically undisturbed periods (Tsugawa et al., 2004; Hawlitschka,
2006) and may be generated by the terminator and other sharp gradients in the
thermosphere (Galuskho et al., 1998; Fujiwara and Miyoshi, 2006).
High-latitude heating also drives a global circulation system that redistributes mass,
momentum, and energy from high to low latitudes, and produces composition effects
similar to those depicted in Figure 7. The associated meridional winds at middle latitudes
can be of order 100 ms-1, but Coriolis effects are a limiting factor and veer the winds to the
west. The way in which the geomagnetically-forced circulation system interacts with the
prevailing solar-driven circulation (i.e., Figure 7) determines the equatorward penetration
and amplitude of the magnetic storm response in each hemisphere. Equatorward
penetration of disturbance effects originating in the summer polar-auroral region are
facilitated by the prevailing summer-to-winter circulation, but the same circulation impedes
equatorward penetratio of disturbances originating in the winter hemisphere. A similar
argument holds for the TADs discussed above. In addition, the meridional flow associated
with diurnal heating (i.e., equatorward at night and poleward during the day) can enhance
or diminish effects due to the seasonal circulation. On average, Joule heating is also
greater in the summer polar-auroral region due to enhanced ionization levels, which adds
to the prevailing (zonal mean) solar-driven circulation. All of these effects serve to produce
asymmetries of the magnetic storm response of the thermosphere with respect to the
equator.
An example is provided in Figure 11, which illustrates the thermosphere density
response near 400 km to the large and isolated magnetic storm occurring on day 324
(November 20) of 2003. The middle latitude daytime response (top panel) in the Southern
Hemisphere is almost a factor of 3 times the quiet-time levels representative of day 323,
and is of order 50% larger than the Northern Hemisphere (winter) response. It would
appear that the prevailing zonal-mean solar-driven circulation has acted to facilitate
expansion of the density disturbance out of the southern high-latitude region, but to inhibit
expansion out of the polar-auroral region in the winter hemisphere, as anticipated from the
24
previous discussion. Note that daytime tidal winds are expected to weaken the
disturbance-driven circulation in both hemispheres. Joule heating due to elevated
conductivities in the summer high latitudes might have further contributed to the enhanced
response in the summer hemisphere.
Figure 11. Densities inferred from the CHAMP accelerometer normalized to 410 km
during (a) day (top) and (b) night (bottom) versus geomagnetic latitude and UT during
days 323–325, 2003. The solid black lines denote the ap index. Magnetic local times are
indicated at the top of each panel. (Bruinsma et al., 2006).
25
With regard to the nighttime response (bottom panel) in Figure 11, we might have
anticipated emergence of the density disturbance to be facilitated in the summer (Southern
Hemisphere) as opposed to winter (Northern Hemisphere) on the basis of a summer to
winter prevailing circulation alone. However, the observed density response is more
symmetric-like than during daytime. In this case, diurnal tidal winds are equatorward in
both hemispheres, and thus diminish the importance of the prevailing circulation in the
winter hemisphere. This, combined perhaps with a disturbance-driven circulation of
exceptional magnitude for this extreme event, may have minimized the relative importance
of the zonal mean circulation, giving rise to a more symmetric response.
Figures 10 and 11 are illustrative of a new data source for the study of thermosphere
disturbances, that of the accelerometers on the CHAMP and GRACE satellites. CHAMP
was launched in July 2000 at 450 km altitude in a near-circular orbit with an inclination of
87.3°. The two identical satellites GRACE-A and GRACE-B were launched in March 2002
at approximately 500 km altitude, in near-circular 89.5° inclination orbits with GRACE-B
following approximately 220 km behind GRACE-A. The in-situ density data obtained from
the accelerometers on these satellites are proving invaluable for the study of solar
disturbance effects on the thermosphere, as well as diurnal, season, solar activity,
magnetospheric and ionospheric effects on the neutral thermosphere. For instance,
Forbes et al. (2005) investigated the disturbed period of 15–24 April 2002, and among
other aspects of the storm response, they followed the propagation of a TAD from high to
low latitudes and into the opposite hemisphere, following an impulsive injection of energy at
high latitudes. Liu et al. (2005) analyzed a year's worth of CHAMP data to discover density
enhancements associated with (1) the dayside cusp and (2) the equatorial ionization
anomaly. The global thermosphere density response to the intense storms of 29 October to
1 November 2003, as revealed by CHAMP measurements, were presented by Liu and Lühr
(2005) and Sutton et al. (2005). In the work of Sutton et al. (2005) the density response to
the solar EUV flare of 28 October 2003 was also detected (see also Sutton et al., 2006),
and in addition, these authors derived winds from the cross-track accelerometer that
revealed interesting similarities and differences with the empirical Horizontal Wind Model
(Hedin et al., 1996). Liu and Lühr (2005) and Bruinsma et al. (2006) also examined the
storm of 20–21 November 2003 and noted hemispheric asymmetries in the density
26
response that were interpreted in terms of solar EUV-driven wind patterns (cf. Fuller-Rowell
et al., 1996), as well as certain aspects of wave-like structures during these periods. In the
above works by Sutton et al., (2006) and Bruinsma et al. (2006), GRACE data were used
simultaneously with CHAMP data to provide an additional perspective on the storm
response. Considering that CHAMP and GRACE provide nearly continuous pole-to-pole
coverage with high resolution at four local times, it is anticipated that data from these
satellites will continue to provide unique opportunities to observe the global-scale response
of the neutral thermosphere to variations in solar photon and corpuscular radiation, and to
coupling processes within the magnetosphere-ionosphere-thermosphere system.
During recent decades much has also been learned about the recovery to quiescent
levels after magnetic storms. Nitric oxide (NO) production is greatly enhanced at high
latitudes, and the NO, which is an efficient radiative cooler, is transported equatorward by
the fortified wind system. Ultimately, much of the excess energy deposited into the
thermosphere during a magnetic storm is radiated to space in the infrared, mostly by NO.
Figure 12. Example of individual energy loss rate profiles derived on 14 April 2002
(quiescent conditions) and 19 April 2002 (disturbed conditions) at 52°N latitude. The peak
energy loss rate exceeds 2200 K/day at this latitude on 19 April. (Mlynczak et al., 2005).
27
but with small additional contributions from O and CO2. A detailed analysis of the flow of
energy within the thermosphere from the perspective of infrared radiation transport and
heat conduction was recently performed for the magnetically disturbed periods that
occurred during April 14-22, 2002 (Mlynczak et al., 2005), using data from the SABER
instrument on TIMED and a first-principles thermosphere model. The dominant radiative
response is associated with dramatically enhanced infrared emission from NO at 5.3 µm.
Energy loss rates due to NO emission exceed 2200 Kelvin per day (see Figure 12). In
comparison, energy loss from CO2 emission at 15 µm during the storms was only 2.3% that
of NO, and energy loss due molecular heat conduction was about 3.8% of that due to NO.
These authors refer to NO cooling as a “natural thermostat” by which storm energy is
primarily lost from the thermosphere.
6. Outstanding Questions and Prospects for the Future
Although considerable progress has been made in understanding the IT system, and in
particular the dynamics of the thermosphere, over the past 25 years, there is still much that
we do not know. Studies in this area still remain a fertile ground for students. Below, some
of the outstanding questions that remain, solely in the area of thermosphere dynamics, are
enumerated, with reference to works that address some aspect of the issue. (N.B. Many of
these topical areas were not addressed in previous sections since significant progress was
not made over the past 2-3 decades.) The list is by no means complete, but does serve to
illustrate the various pathways by which future research efforts are likely to follow.
•
What is the spectrum of gravity waves penetrating into the thermosphere from
below, and what are the consequences for thermosphere variability, and modification
of the mean state (e.g., Vadas and Fritts, 2006)?
•
What is the spatial and temporal distribution of turbopause altitude, what are the
causes, and what effects does this have on the global distribution and variability of
thermosphere composition (e.g., Maeda et al., 1986; Fuller-Rowell and Rees, 1992)?
•
What is the fundamental cause and what are the consequences of the extreme
wind shears observed in the MLT region (Larsen, 2002)?
28
•
What is the cause of oscillations of planetary wave periods in the upper
thermosphere revealed in densities derived from satellite drag (Bowman, private
communication, 2006)?
•
To what extent do gravity waves excited in the auroral region contribute to the
global redistribution of mass and momentum in connection with geomagnetic
disturbances (e.g., Brinkman et al., 1992)?
•
What is the cause of the semiannual variation in the thermosphere, including its
dependence on solar activity (Fuller-Rowell, 1998; Bowman, 2004)?
•
What is the cause of the apparent long-term decrease in thermosphere density,
spanning four decades (e.g., Marcos et al., 2005)?
•
What is the cause of long-term (decadal) changes in solar tides in the MLT region
(e,g., Jarvis, 2005)?
•
How is magnetospheric energy input into the polar-auroral region partitioned
between Joule heating, particle heating and momentum forcing?
•
What are the spatial and temporal scales of high-latitude heating rates that need to
be resolved in order to capture dynamical structures on more global spatial-temporal
scales (e.g., Codrescu et al., 1995)?
•
What is the impact of the ionosphere on neutral density and wind structures, e.g.,
equatorial anomaly (Liu et al., 2005) and midnight temperature anomaly (Fesen,
1996)?
The above questions and research areas are fundamental to more pragmatic issues,
such as response of the upper atmosphere to anthropogenic influences, the prediction of
satellite drag, and prediction of ionospheric conditions for a variety of communications
purposes. Answers to these questions will still require the usual suite of research
methodologies: basic theory, numerical simulation, and observations. Numerical
simulations are likely to benefit from advances in computing power, numerical techniques,
and realistic connections to the magnetosphere system and lower thermosphere. Major
recent advances have been the upward extension of a full meteorological GCM up to the
exobase (Miyoshi and Fujiwara, 2003, 2006), and the inclusion of self-consistent
29
dynamical, chemical, radiative, and electrodynamical couplings between the thermosphere
and mesosphere, 30-500 km (Roble and Ridley, 1994; Richmond et al., 1992). However, a
fully self-consistent chemical-dynamical-electrodynamics model extending from the ground
to the exobase, as well as self-consistent coupling with the plasmasphere and
magnetosphere, will be required before a comprehensive understanding and modeling
capability of the IT system is attained.
Future observational study of the IT system presents many challenges. Due to the
coupling between chemical, dynamical and electrodynamic processes, comprehensive,
simultaneous and global measurements covering all of these disciplines are of course
desirable. In terms of neutral dynamics alone, the main focus of this review, the most
pressing needs are (a) global measurement of gravity wave fluxes into the thermosphere
from below, the wave field within the thermosphere, and sufficient additional measurements
and modeling to understand the relations between gravity waves and the mean state of the
IT system; (b) measurement of the global wind field from 100 to 300 km, with coincident
measurements of chemical composition and ionosphere plasma densities; (c)
measurements that allow delineation of the dependence of magnetic storm response on
longitude and local time.
Acknowledgments: The author acknowledges support through grant ATM-0346218
from the National Science Foundation and the Glenn Murphy Professorship at the
University of Colorado in preparation of this review.
30
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