Lower-thermospheric neutral winds at high latitude determined from incoherent scatter measurements: a review of techniques and observations

Adv. Space Res. Vot. 10, No. 6, pp. (6)261—(6)275, 1990 Printed in Great Britain. All rights reserved

0273—1177/90 $0.00 + .50 Copyright © 1989 COSPAR

LOWER-THERMOSPHERIC NEUTRAL WINDS AT HIGH LATITUDE DETERMINED FROM INCOHERENT SCATTER MEASUREMENTS: A REVIEW OF TECHNIQUES AND OBSERVATIONS R. M. Johnson Geoscience and Engineering Center, SRI International, 333 Ravenswood A venue, Menlo Park, CA 94025, U.S.A.

ABSTRACT Neutral winds can be deduced from E-region incoherent scatter measurements of the ion velocity, in conjunction with measurements of the electric field and knowledge of the ion-neutral collison frequency. Measurement techniques and methods used to derive lower-thermospheric neutral winds are presented and compared. The sensitivity of the results to the ion-neutral collision frequency is discussed. Observations of the lower thermosphere at mid- and high latitudes show that neutral dynamics are influenced by both tidal oscillations and magnetospheric forcing. The tidal results obtained using Mesosphere-StratosphereTroposphere (MST), Medium-Frequency (MF), meteor, and Incoherent Scatter (IS) radars are generally consistent, and show the dominance of semidiurnal oscillations. At Chatanika, Alaska (65°N,65A), the average summer semidiurnal oscillation reaches a peak amplitude of approximately 60 rn/s at 110-km altitude. Recent measurements obtained using the IS radar at SØndre StrØmf]ord, Greenland (67°N,74.5A), show a similar strong semidiurnal component in the neutral winds. The response of the neutral winds to magnetospheric forcing is in qualitative agreement with the results of three-dimensional global model simulations of the thermosphere, which predict deviations in the normal flows as a result of frictional coupling to the convecting ions. Auroral zone measurements at Chatanika show that the response. is strongest in the morning sector, where enhanced zonal flow is detected during geomagnetically disturbed conditions. Recent polar cap measurements at Sondrestrom likewise show a response to enhanced ion convection during a highly disturbed interval. INTRODUCTION The neutral atmosphere at lower thermospheric heights (90 tol3O km) is influenced by multiple physical processes which play a role in determining its dynamic behavior. The lower atmosphere exerts its influence mainly through the upward propagation of gravity waves and thermal tides. Thermal tides are also driven in situ in the thermosphere, which is additionally influenced by frictional coupling of the neutral gas to the motions of the ionized constituents. This frictional coupling produces a source of energy and momentum for the thermosphere, which can become particularly important at high latitudes during intervals of geomagnetic activity. To some extent, the major processes may interact; for instance, changes of the mean flow structure can alter the upward propagation of gravity waves and tides from the lower atmosphere. Only recently have observational tools been developed which allow reasonably regular measurements of neutral dynamical parameters in the lower thermosphere. Rocket experiments /1—4/, optical techniques /5/, MST /6,7/, meteor /8/, medium-frequency (MF) /9,10/, and incoherent scatter (IS) /11—32/ radars can all provide some information about the dynamical state of the lower thermosphere with varying degrees of temporal and spatial resolution. Other than by remote optical satellite studies /33/, this region is not accessible for in situ satellite measurements because of strong atmospheric drag. The incoherent scatter technique allows a wide variety of ionospheric and thermospheric quantities to be measured (electron density, electron and ion temperatures, line-of-sight ion velocity, and ion-neutral collision frequency) or derived (vector ion velocity, electric field, vector neutral wind, temperature, and density, ionospheric conductivities, Joule heating, and electrojet currents, to name a few) from the return signal power and spectrum. Altitude resolution of the order of a kilometers is now available at most IS radars, and time resolution up to the order of minutes allows measurements of a wide spectrum of atmospheric oscillations. At high latitudes, the ability to observe the dynamics of both the ionized and neutral components simultaneously makes the incoherent scatter technique unique for the study of magnetosphereionosphere-thermosphere coupling. In the next section, measurement techniques and methods used to determine the neutral wind are described, with emphasis on the importance of the ion-neutral collision frequency in the derivation of the neutral wind. This section is followed by a review of lower-thermospheric vector neutral wind results derived from

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R. M. Johnson DERWATION OF THE NEUTRAL WIND

The incoherent scatter return signal spectrum is dependent on multiple ionospheric parameters: electron density, ion and electron velocities and temperatures, and ion composition /34/. To obtain estimates of these parameters, a nonlinear least squares fit of theoretical autocorrelation functions is made to the measured autocorrelation function of the return signal. In practice, it is not possible to obtain wellconstrained values for all of these quantities because of the uncertainties of the measurements. In addition to measurement errors, some parameters have correllated effects on the spectrum, e.g. 1’~and m~.Instead of attempting to fit for all the parameters simultaneously, reasonable assumptions are used to limit the number of parameters fitted for within any given altitude range. For E region measurements, an iterative approach can be used to extend the number of parameters that can be obtained. In the analysis of Chatanika and Sondrestrom E-region autocorrelation functions, we assume thermal equilibrium between electrons and ions and use an ion-neutral collision frequency based on empirical ion and neutral composition models for the first fit. With these assumptions, data quality is sufficient to allow estimation of the line-of-sight ion velocity v 1, ion (and electron) temperature T1 = electron density ne, and their uncertainties. The total ion-neutral collision frequency is approximated by the weighted average of the contributions from the dominant ionic constituents NO~,O~,and O~due to collisional interactions with N2, 02, and 0, ~‘ij,



(1)

i=NO~,O~ ,o~ j=N2,02 ,O

where n is the number density of the ith ion species and N = = n~is the total ionization density. The number densities of the different ionic species are obtained using an empirical model of the E-region ionospheric composition based on measurements of n0÷/ne and an assumed molecular composition of 75% NO~and 25% O~/35/. The individual collision frequencies ii,j are those given by Schunk and Walker /36/, and are dependent on the number density of the different neutral constituents (as well as the ion and neutral temperature for O~—Ocollisions and, when 2’,. (T~+ T,5)/2 > 800 K, for 0~—O2collisions). Throughout the lower thermosphere, the dominant contributions to t’,,~come from collisions involving N2. The neutral number densities and temperatures are obtained from an empirical model atmosphere, and temperature equilibrium is assumed between ions and neutrals (Tr = Tft). Using the estimate of v~obtained from the first fit, a second fit can then be performed to obtain the ionneutral collision frequency v,,~.Alternatively, the electron and ion temperatures T~and T1 can be allowed to vary independently using the model i’,,~and estimate of v determined from the first fit. In order to derive vector neutral winds from monostatic (e.g., from Sondrestrom) incoherent scatter radar measurements at E-region heights, the line-of-sight ion velocities along three successive beam positions (a,b, and c) have typically been used to form the ion velocity vector (see Figure la). In each position, the beam remains fixed for a dwell time, typically 3 to 5 mm. This implies that the vector neutral wind derived from monostatic radar measurements are limited in temporal resolution by the time required for the beam to complete at least one full sequence of three beam positions (typically 9 to 15 mm). Inherent in the results is the assumption that each line-of-sight velocity, v4 (t1), Vb(t2), v~(t3),remains fixed for the full interval required to form a vector (Figure ib), and that the velocity vector does not vary significantly over the spatial region described by the three-position measurement. In the analysis of E-region measurements from Sondrestrom and Chatanika, we do not assume that each line-of-sight quantity remains fixed for the interval required to make a three-position measurement; rather, the line-of-sight quantities can be allowed to vary linearly in time (Figure lc). Five successive line-of-sight measurements (Va(tl),Vb(t2),Vc(t3),va(i4), and vb(t5)) obtained in a three-position mode are used, and the line-of-sight quantities from two positions (a and b) are interpolated to the time of the central observation (c). These two approaches both produce some amount of temporal and spatial smearing of the results. For example, using three line-of-sight measurements obtained at 70°elevation angle with a 5-mm dwell time per position, the measurements at 110 km are obtained over 1°latitude and incorporate local time smearing equivalent to — 4°longitude. Using five line-of-sights, the longitudinal span of the measurements increases to -~ 6°.However, the latter is an improvement on the former in that it takes into account the temporal variability of the line-of-sight quantities. All the quantities derived from the measurement can be treated in a similar fashion, including ion velocity, ion temperature, electron density, electron temperature, and electric field. For a tristatic radar such as EISCAT, the beam position of the transmitter (and its associated receiver) is fixed, and the two remote-receiving-station beams are synchronously moved to intersect at a sequence of altitudes (Figure id). The combination of the three receiving beams intersecting in the same volume of space allows the vector ion velocity to be determined directly from the triad of line-of-sight measurements (Figure le). However, the electric field, obtained from ion velocity measurements in the F region, is measured at a time different from the measurements of ion velocity in the E region, and thus temporal smearing of the derived neutral winds over the cycle time of the experiment (typically about 10 miii) also occurs.

c

b

/

/

Lower Thermospheric Neutral Winds

(6)263

a ~

(receiver) 23’E

(a)

-

~

69 ‘N

67’ (recetver)

(d) Constant

Constant

_—>~T

-~-—

V~(t,). V 5(i,), V~(t,). V,,(t,). Vb(tg), V,(t6). _________________I

V~(t,). V~(ti). V~(ti), ~‘T(t1). V~(t2), V5(iz),

...

I

_________________

I

I. —

V(z5).

Y(t,).

P(t,),

...

(b)

(e)

interpolated to

—~4~

V~(s1). V5(t2). V~(e,), V,(e,), V5(t5), V~(e6).

L (c)

Fig. 1. Illustration of techniques used to measure line-of-sight ion velocities and to derive vector velocities from line-of-sight measurements: (a) three-position mode used for a monostatic radar; (b) vector velocity determination using three monostatic line-of-sight measurements; (c) vector velocity determination using five monostatic line-of-sight measurements; (d) example of line-of-sight positions used for a tristatic system; (e) vector velocity determination using tristatic line-of-sight measurements. iii(z)

=

ifj(z)

—

(~)(2~.

+ Uijz) x

a),

(2)

where E~ is the ionospheric electric field, iijjz) is the height-dependent E-region ion velocity perpendicular to B, and 1l~ = eB/m1 is the ion gyrofrequency. Assuming that ion-neutral collisions have negligible influence on ion motion above —200 km, E~can be determined from the ion drift at F-region altitudes using E~= ~ x B. In the analysis of Chatanika and Sondrestrom measurements, 12~is calculated using an ion mass equal to the weighted average determined from the ionospheric composition model described above. Parallel to

.a, the ion equation of motion is

u11(z)

=

v11(z)

—

vamit(z),

(3)

where v~,,,gj(z)is the height-dependent ambipolar diffusion velocity /37/. At E-region altitudes, vamgj(z) ~5 usually negligible relative to v11, and the ionized and neutral constituents move together along the magnetic field. The components of the neutral wind (in a right-handed coordinate system with the z-axis aligned anti-parallel to B in the northern hemisphere and the y-axis in the direction ofmagnetic north but inclined out of the horizontal nlane’~can be reexnressed in terms of the neutral wind comoonents in local ~eo~raohic

R. M. Johnson

(6)264

and =

—u

5sinScosl — u,,cosöcosl + u,~,,sinI, (4c) where the subscripts e, n, and up refer to local geographic east, north, and vertical, 5 is the magnetic declination, and I is the inclination of the magnetic field from the horizontal plane. Assuming the vertical neutral wind, u,~,,is negligible relative to horizontal neutral winds, u~and u,,, the neutral wind parallel to B is produced solely through projection of horizontal winds along fr With these assumptions and equation (3), the component of the horizontal neutral wind in the magnetic meridian, Urn, can be determined from measurements parallel to B, (5)

tLm~.

Equation (5) has been exploited to obtain meridional neutral wind measurements at E-region heights with the St. Santin (47°N,2°E)-Nancay(45°N,2°E)radar facility /11,12,15,17,18,24,29/. As this review focuses on derivation and observations of vector neutral winds, meridional wind results obtained using this bistatic single-position technique will not be described in detail here. 11mrn The smaller L’j,, becomes, Equation the more the (2) ion shows velocity that the diverges neutral from wind thevelocity neutral wind is inversely becausedependent of the t1/v~,, on weighting factor. The inverse dependence on p 1,, causes uncertainties in ~ and ~j to result in proportionately larger uncertainties in iij(z) at greater heights. It is therefore important to determine these quantities as accurately as possible. In an analysis of Chatanika summer experiments, Johnson et al. /32/ interpolated in time in the manner described above (Figure lc) to find a better estimate of the inputs to equation (2) than can be obtained assuming fixed line-of-sight velocity measurements. Additionally, the electric field was determined from F-region measurements iteratively so as to minimize the uncertainty of the resulting average. Figures 2 and 3 show how uncertainty in L’~,,can affect the resultant neutral wind obtained through equation (2). The measurements analyzed here were obtained with the Chatanika radar on 11 June 1980 when conditions were highly disturbed. Three different neutral density profiles were calculated, corresponding to 75%, 100%, and 125% of the MSIS-.83 value /38/. The left panel of Figure 2 compares the profiles of the zonal neutral wind determined using these density profiles at a time when contributions from the electric field and the meridional component of the ion velocity dominate the calculation. On the righthand side of Figure 2, the zonal ion velocity dominates, while the contributions from Ej and ~ x B are comparatively minor. Thus, depending on the relative magnitudes of the ion velocity components and the electric field, the uncertainty produced as a result of uncertainty in i’~,, can range from being negligible to the order of the neutral velocity determined using equation (2). In Figure 3, this same calculation is shown for the entire day, again for the zonal neutral wind, but now only at 115 km. Superimposed is a curve indicating the average for Kp< 2 conditions during summer months /32/. Even though the derived neutral velocities show substantial differences between the curves for different MSIS—83 density profiles, they all differ drastically from that determined for geomagnetically quiet conditions. ZONAL NEUTRAL WIND 130 800611

155632

800611

200051

100

90

.400

—200

0

200

400 —400

VELOCITY (m/s) •—.

0.75 p

.200 .

.—

1.00 p

0

200

400

VELOCITY (mit) •

.

1.25 p

Fig. 2. Altitude profiles of the zonal neutral wind on 11 June 1980 at two different universal times. The three curves on each panel show the neutral wind obtained using the MSIS-83 density profiles multiplied by 0.75 (dashed), 1.00 (solid), and 1.25 (dotted). Figure 4 compares the average error in the derived zonal neutral wind resulting from a 25% uncertainty in the neutral density with the average error produced from uncertainty in the line-of-sight measurements. i’s... c......... .. ~ ~ s........ ~ ,~ s.~.... s.;.,.,~.i .~ ~ ~..c~ ,A

Lower Thermospheric Neutral Winds I

F

500

400

—

— ——

300

—

200

-

—100

-

—200

-

I

I

75% MSIS 100% MSIS 125% MSIS <

—

/.~

2 average

—

-

k4:.~. ..........

—

—

I

—300 0

(6)265

2

4

6

8

10

I

I

12 UT (hours)

14

16

I

I

I

15

20

22

24

Fig. 3. Zonal neutral wind (geomagnetic coordinates) at 115 km on 11 June 1980 obtained using the MSIS-83 density multiplied by 0.75 (solid), 1.00 (long dash), and 1.25 (short dash). The dotted curve shows the average zonal neutral wind (geomagnetic coordinates) at 115 km obtained for quiet geomagnetic conditions when daily averaged Kp less than 2.

./

~

0

25

“

..

.._~

120km

~

-

111km

0

25

-

-

102 km 0

~

25

-‘‘—

--.

..- .:.~

.

-

93 km

0

-

~

~-

-.

(6)266

R. M. Johnson

Figures 2 through 4 demonstrate the critical importance of L’s,, to the calculation of the neutral wind using this technique. The question is: How accurate are our estimates of v 1,,? Recent results /39/ have indicated that the r’~,,determined from incoherent scatter measurements at EISCAT are a factor of 2 to 3 times greater than those determined using the MSIS—83 model. Such an underestimate would mean that neutral and ion motions are more tightly coupled than has previously been assumed, and that the same value of Lii,, is obtained roughly one scale height higher (— 6 km /25,28,30,31/). Measurements of neutral composition indicate variations in number densities with time. Although the results of Offerman et al. /40/ show relatively small variations in N2] and [021 (— 5%), [0] was observed to vary by factors of 4 to 5. Finally, at high latitudes, large e ectric fields and particle precipitation can produce elevated ion (and neutral) temperatures, and thereby change the temperature dependent components of the collision frequency /28/. Model results show that Joule heating produces upwelling of the neutral atmosphere /41/, enriching the number density of the heavier species N2 and 02 relative to atomic oxygen. Thus, at least for atomic oxygen, the temperature dependence of i’~+_0 is somewhat mitigated by a reduction in [0] produced from upwelling. This mitigation does not occur for Ot—Os, since [021 and 1’,. increase in concert. The neutral winds determined using this technique must therefore be interpreted with an awareness of these considerations. The derived neutral wind is probably most accurate during quiet geomagnetic conditions, when temperatures and densities and therefore r’~,,are more likely to remain stable with time. Further work on precise measurements of time dependent Lii,, are highly desirable in order to allow neutral wind calculations of improved accuracy. Previous measurements of ii~,, using the incoherent scatter technique have unfortunately had considerable error bars, making it difficult to obtain a well-constrained experimental value. Instead ui,, measurements have been obtained by time averaging over large data sets /28/ or through empirical models. OBSERVATIONS The incoherent scatter technique has been used to derive vector neutral winds at lower thermospheric heights for over 15 years. The Chatanika radar, which operated from 1971 to 1982 before being relocated at Søndre Strømfjord, acquired measurements at E-region altitudes using several different observation modes with varying degrees of resolution. The first results /13,16,22/ were obtained using a 320 ss pulse length, resulting in a nonuniformly weighted average ion velocity determined over an effective range gate of -~45-km width. The resulting neutral winds represented an average over the E region, smearing over any fine-scale structures such as large wind shears /1,2,3/ which are commonly observed over tens of kilometers in this region. Figure 5 shows the average neutral wind pattern obtained from six experiments using this observation mode / 13/ during the months of February (24 hr), May (25 hr), and July (54 hr) of 1972. The average E-region neutral wind was found to blow in an antisolar direction across the polar cap from the hot dayside. The eastward and westward turning of the wind vector during morning and evening hours (when the Chatanika site emerges from and enters the auroral zone, respectively) was interpreted as evidence of coupling to the ions convecting under the influence of the electric field. MILLSTO~’JE(131 km)

12

tCHATANIKA (110-115 km)

a

(

\~

/1

IS

‘,,

.../.

06 LOCAL TIME

43’

00

Fig. 5. Average neutral wind pattern obtained using long-pulse measurements from February, May, and July of 1972 at Chatanika in the E-region. From /13/. Measurements obtained at Chatanika using this observation mode during the 3—9 August 1972 magnetospheric storm were analyzed to obtain lower thermospheric neutral winds /16/. The data included a quiet interval prior to the storm onset, as well as five days of subsequent measurements during the main phase of the storm and its decay. A similar neutral wind pattern was found during the storm interval as that

Lower Thermospheric Neutral Winds

(6)267

multipulse correlator in six experiments utilizing pulse modes with either 24- or 9-km range resolution. The critical dependence of the derived neutral winds on iii,, was noted. The general pattern of pressure-driven flow from the dayside to the nightside was again detected, as well as evidence of ion drag effects on the neutral winds above 110 km. During a geomagnetically disturbed interval, large equatorward meridional flows were detected during the evening and large meridional flows in the morning sector. Lower therinospheric neutral winds were determined from 46 days of midlatitude measurements obtained with the Millstone Hill (42.6°N, 56°A)incoherent scatter radar /26,27/ over a 17-month period during sunspot maximum. Three-position E- (20-km altitude resolution) and F-region measurements were used to derive the lower thermospheric neutral wind /43/ during daylight hours from 105 to 135 km, in a manner essentially equivalent to that utilized in /23/ and described above. The winds were studied for evidence of both seasonal and geomagnetic activity effects. The semidiurnal oscillation was found to predominate in the observations, although a non-negligible diurnal component was also evident. The semjdiurnal oscillation was found to maximize near 125 km with an amplitude between ranging from 50 to 80 rn/s during the equinoxes and from 40 to 50 m/s during the soistices. In keeping with a model of a propagating semidiurnal tidal mode, the phase difference between the wind components was found to be —~ 3 hr. A vertical wavelength of 70 to 100 km was determined from the phase change of the semidiurnal oscillation between 105 and 135 km. A depression in the semidiurnal tidal amplitude was found during disturbed geomagnetic conditions at 105and 115-km altitude. Additionally, the vertical wavelength of the semidiurnal tide apparently decreased. Possible mechanisms invoked to affect the semidiurnal tide included the existence of an in situ generation mechanism and/or a dissipation mechanism related to the level of geomagnetic activity, with the ability to modify the upwardly propagating tidal component. Prevailing winds became more southwestward by — 25 rn/s during disturbed intervals at 125 and 135 km, while at 105 km a marginally significant southeastward wind shift of — 5 rn/s was detected. The upper-level results are in agreement with the notion that highlatitude Joule heating induces a meridional circulation cell in the thermosphere in which equatorward flow is deflected toward the west through the Coriolis force (in the northern hemisphere) with a return flow at lower altitudes. Evidence for this circulation cell was also found /18/ in an analysis of mid-latitude meridional wind results obtained at St. Santin. Recently, Johnson et al. /32/ presented results from a study of fourteen 24-hr experiments performed during summer months from 1976 to 1982 at Chatanika with altitude resolution ranging from 9 to 24 km. Lower thermospheric neutral winds were derived in the manner described above (Figure ic) at altitudes from 90 to 125 km. The average summer wind structure was characterized by the mean flow, as well as diurnal and semidiurnal tidal components. The mean flow was found to rotate in a counterclockwise sense through 180°from southeastward at 90 km to northwestward at 125 km. The eastward flows that are present from 90 to 125 km may represent the effects of momentum deposition by upwardly propagating gravity waves /44/. Figure 6 from /32/ compares summer Chatanika results for the sernidiurnal tide with those obtained during summer months with the Poker Flat, Alaska, MST radar /7/, the Millstone Hill radar /26/, and medium frequency radar results obtained at Saskatoon /9/. As in the study of Wand /26/, the semidiurnal tide was found to be dominant in this altitude range, although the altitude of maximum amplitude was slightly lower (—~60 m/s at 110 kin) than the Millstone Hill results indicate (Figures 6a and c). Similarly, a 3-hr phase shift between the zonal and meridional components of the same sense was determined from 90 to 115 km, indicating the presence of a strong propagating semidiurnal component (Figures 6b and d). However, the vertical wavelength deduced from the phase change from 90 to 125 km of — 50 km was significantly shorter than that found at Millstone Hill (— 70 to 100 km), and greater than that deduced from Poker Flat MST observations (— 30 km). Figure 7 from /32/ summarizes the results of the mean flow and tidal analysis of the summer Chatanika data. Vectors representing the neutral wind component contributed by the diurnal (7a), semidiurnal (7b), net tides (7c) and net tides plus mean flow (7d) are indicated as a function of local time. Figure 7d is comparable with the results obtained by Brekke et al. /13/ (Figure 5), except that the altitude resolution of Figure 7 is greater (—~10km versus -~ 50 km). Also, Figure 7 includes only mean flow and tidal components determined from a least_squares fit to the average of the summer experiments, while Figure 5 represents an average of the observed neutral winds in February, May, and July 1972. The similarity of the two figures is striking, considering the differing altitude resolution and seasonal distribution of the contributing data. This similarity suggests that the weighting function of the long-pulse measurements was strongly peaked near this altitude /13/, and that seasonal variations in the neutral winds are not dramatic. However, the evidence for an upwardly propagating semidiurnal tide seen in Figure 6 (the steady phase progression with height from 90 to 115 km with a 3-hr phase lag between the two components), coupled with the neutral wind vectors shown in Figure 7, indicates that southeastward flows at dawn and southward to southwestward flows in the evening sector are produced at least in part through the combination of mean flow and tidal components. In order to determine whether the lower thermospheric neutral winds during summer months are influenced by auroral zone forcing, Johnson et al. /32/ divided the data set as a function of Kp. Consistent differences were found between average zonal and meridional winds calculated for Kp<2, Kp=2, and Kp>2 conditions above inn km. ndia,at.inc- a svstematk nrnor.ie.enn of inere2.qin~ effette aQ art~vitv inrreaaes The stonal

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R. M. Johnson Average Zonal 12-hr Amplitude

Average Zonal 12-hr Phase

F

I

I

Average Meridional 12-hr Amplitude

~

~:::

:

::~,E~

~‘:: 80

.‘

0

I

I

25

50

VELOCITY (mis)

I

I

ill

Average Meridional 12-hr Phase

75

8

~

(d)

I

I

10

12

I ‘6 14

16

18

20

LOCAL TIME OF MAXIMUM (hours)

—. —. —

1979—1980 Poker Flat. MST. 65°N /7/ 1980 Saskatoon. 52’N /9/ 1976—1977 Millstone Hill. 43’~4/26/

o—O

1976—1981 Chatanika. 65°N/32/

Fig. 6. Comparison of semidiumal tidal results during summer months from radars at Chatankia (open circles), Millstone Hill (dot-dashed curve, taken from /26/), Saskatoon (dashed curve, taken from /9/), and Poker Flat (dotted curve, taken from /7/): (a) zonal amplitudes from 80 to 130 km; (b) local time of maximum eastward amplitude; (c) meridional amplitude; (d) local time of maximum northward amplitude. From /32/. In the same study /32/, the Thermospheric General Circulation Model (TGCM) at the National Center for Atmospheric Research (NCAR) was used to simulate the response of the lower thermosphere to three levels of steady-state magnetospheric forcing. The magnetospheric forcing was characterized by the average ionospheric convection patterns determined by Foster et a!. /45/ for particle precipitation levels 4, 7, and 9—10. In Figure 8, the difference between the average radar derived neutral winds at 115 km determined for Kp>2 and Kp<2 conditions is compared with the difference between the model results for high-latitude forcing at the 9—10 level and level 4 near 115 km. The model does a reasonable job of reproducing most of the differences apparent in the data, with the exception of the absolute magnitude of the differences and slight shifts in time between the model and the data. Both model and data show enhanced poleward flows before dawn, enhanced eastward flows after dawn through the dayside, and enhanced westward flows after dusk on the nightside. The most significant differences occur for the meridional component, where the model predicts more poleward flow, while the data displayed here and obtained in previous studies /13,16,23,46/ show a more equatorward flow. This difference, as well as the time shift between the model and data, can be attributed, at least in part, to imperfect representation of the high-latitude convection and precipitation pattern used in the model. Additionally, the model results were obtained using tidal forcing at the lower boundary appropriate to equinox conditions, since solstice values were not yet available. During a highly disturbed interval when Ap=54 (11 June 1980), morning sector zonal winds at 115 km exceeded those obtained on the average for Kp<2 conditions by — 250 m/s for an extended interval (from 1400 to 2100 UT) of intense particle precipitation and — 50 mV/rn electric fields, while meridional flows remained more equatorward throughout thifl same interval by — 100 rn/s (Figure 9). Although this interval was the most disturbed of the day, 10 and 11 June were the two most disturbed days of the month. The interval can therefore be considered to be one of continuous geomagnetic activity, although a constant level of disturbance was not maintained throughout. The winds during this disturbed interval were more diurnal in character than those observed during quiet intervals, when a strong semidiurnal variation is evident. Rather than adopting an explanation based on changes in tidal dissipation rates /27/ or in situ generation of tidal modes /27,47/, Johnson et al /32/ suggested that the apparent tidal changes can be understood in terms of the balance obtained between upwardly propagating semidiurnal tides and ion-drag forcing. Ion drag acts in concert with the semidiurnal comnonent during the evening sector after dusk, and against

Lower Thermospheric Neutral Winds -

(a)

(6)269 (c)

100 rn/s

100 mis

D~ALHTIDE~,

-

-

-

(b) 12

(dl 100 mis

•

NET TIDES

+

12

100 m/s

_________

Fig. 7. Polar plots showing (a) the diurnal, (b) semidiurnal, (c) net tides, and (d) tides plus mean flow calculated from a least-squares analysis of Chatanika average summer neutral wind velocities at 115 km. From /32/. Recently, Virdi and Williams /48/ have analyzed EISCAT measurements to deduce lower thermospheric neutral winds during 1985—1987 for conditions of magnetic quiet. These high-latitude data also showed the presence of a strong semidiurnal tide, with a wavelength most consistent with the (2,4) mode below 120 km. LTCS-1 OBSERVATIONS AT SONDRESTROM-AN OVERVIEW The incoherent scatter radar at Sondrestrorn obtained both E- and F-region measurements during the first core observation campaign of the Lower Thermospheric Coupling Study from 21 September at 1000 UT through 26 September at 1700 UT. The E-region measurements had a range resolution of 24 km, and a dwell-time of 5 min per position in the three-position sequence that was used. A range gate spacing of 6 km was used, with the first gate set at 90-km range. Thus between 90 and 126 km there are four different altitudes with independent measurements. The radar beam azimuths were 141°,261°,and 21°,and a common elevation of 700 was used. Overall, the experiment ran smoothly, with only two short intervals when no data were obtained. The geomagnetic conditions during LTCS-1 were variable, as is seen in Figure 10, where the ionospheric electric field measured at Sondrestrom is shown along with Kp and Ap throughout the experiment. The first interval of activity on the 22—23 September (indicated by Kp and Ap) was not accompanied by the large polar cap electric fields that were measured during the second interval of activity on 25—26 September. Neutral winds were derived from the measured ion velocities and electric field using MSIS—83 values to determine z’j,, as shown in equation (5) and Figure ic. Figure 11 shows the zonal and meridional winds determined on 25 September (solid curves) binned in 1-hr intervals, in conjunction with the average winds determined on the two preceding days, 23—24 September (dashed curves). The two latter days were significantly less active geomagnetically than the former day, when up to 100-mV/rn electric fields were measured at Sondrestrom (Figure 10). The neutral winds on the 23rd and 24th vary with a semidiurnal oscillation which is most clearly apparent at the lowest altitudes. Over the 90- to 120-km altitude range, a phase change of -~ 6 hr is evident, corresponding roughly to a vertical wavelength of 60 kin, and the meridional component leads the zonal component by roughly 2.5 hr. The data from the 25th also show the presence of a semidiurnal wave at 90 and 100 km, but above this altitude, the neutral winds diverge from those determined during quiet conditions. An enhanced westward flow developed after dawn (0900 UT),

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R. M. Johnson DIFFERENCE VELOCITIES (115 km)

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Fig. 8. Comparison of the difference velocities obtained from the derived neutral winds and from the steady-state TGCM results. The dashed curves represent the difference between average neutral winds calculated for daily averaged Kp greater than, and less than 2. The solid curves show the difference between model results obtained using magnetospheric forcing appropriate to /45/ levels 9—10 and 4. CONCLUSIONS Neutral winds at lower thermospheric heights can be derived from combined E- and F-region incoherent scatter measurements. In order to obtain vector neutral winds, observations along at least three different lines-of-sight are required. Spatial and temporal smearing of the derived neutral winds results from the techniques used to obtain the three line-of-sight measurements as well as the electric field. The derived neutral winds are sensitive to the value of the ion-neutral collision frequency, particularly as altitude increases above -~ 110 km. Time-dependent composition changes, resulting, for example, from localized high-latitude heating, could produce variations in ii,,, that are not taken into account in the value obtained using an empirical model such as MSIS-83. Mid- to high-latitude observations show that the semidiurnal tide is a dominant feature at lower thermospheric heights. During summer months (when incoherent scatter data is most plentiful for comparison), the semidiurnal oscillation maximizes near 125 km at mid-latitudes /26/ with an amplitude between 40 and 50 m/s, while at high latitude the oscillation maximizes near 110 km at -~ 75 rn/s /32/. Consistent and systematic differences were found between neutral winds derived from Chatanika radar measurements under geomagnetically quiet and active conditions /32/. At this auroral zone latitude, the zonal wind showed the greatest sensitivity in the development of enhanced eastward flows through the morning sector during active conditions. Comparison with steady-state TGCM results for different levels of magnetospheric forcing showed that, in general, the model reproduces the average response observed in ~

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R. M. Johnson

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REFERENCES 1. E. Pereira, M. C. Kelley, D. Rees, I. S. Mikkelsen, T. S. Jørgensen, and T. J. Fuller-Rowell, Observations of neutral wind profiles between 115-and 175-km altitude in the dayside aurora! oval, J. Geophys. Res. 85, #A6, 2935—2940 (1980). 2. I. S. Mikkelsen, T. S. Jørgensen, M. C. Kelley, M. F. Larsen, E. Pereira, and J. Vickrey, Neutral winds and electric fields in the dusk auroral oval 1. Measurements, J. Geophys. Res. 86, #A3, 1513—1524 (1981). 3. I. S. Mikkelsen, M. F. Larsen, M. C. Kelley, J. Vickrey, E. Friis-Christensen, J. Meriwether, and P. Shih, Simultaneous measurements of the thermospheric wind profile at three separate positions in the dusk auroral oval, J. Geophys. Res. 92, #A5, 4639—4648 (1987). 4. J. P. Heppner and M. L. Miller, Thermospheric winds at high latitudes from chemical release observations, J. Geophys. Res. 87, #A3, 1633—1647 (1982). 5. L. L. Cogger, J. S. Murphree, C. A. Tepley, and J. W. Meriwether, Jr., Measurements of the E region neutral wind field, Planet. Space Sci. 33, #4, 373—379 (1985). 6. B. B. Balsley, W. L. Ecklund, D. A. Carter, and P. E. Johnston, The MST radar at Poker Flat, Alaska, Radio Sci. 15, #2, 213—223 (1980). 7. D. A. Carter and B. B. Balsley, The summer wind field between 80 and 93 km observed by the MST radar at Poker Flat, Alaska (65°N),J. Atmos. Sci. 39, #12 (1982). 8. D. Tetenbaum, S. K. Avery, and A. C. Riddle, Observations of mean winds and tides in the upper mesosphere during 1980—1984 using the Poker Flat, MST radar as a meteor radar, J. Geophys. Res. 91, #D13, 14539—14555 (1986). 9. A. H. Manson, C. E. Meek, and J. B. Gregory, Winds and waves (10 mm—3D days) in the mesosphere and lower thermosphere at Saskatoon (52°N,107°W, L=4.3) during the year, October 1979 to July 1980, J. Geophys. Res. 86, #C10, 9615—9625 (1981). 10. A. H. Manson and C. E. Meek, The dynamics of the mesosphere and lower thermosphere at Saskatoon (52°N),J. Atmos. Sci. 43, #3, 276—284 (1986). 11. R. Bernard and A. Spizzichino, Semidiurnal wind and temperature oscillations in the E-region observed by the Nancay Incoherent Scatter Experiment, J. Atmos. Terr. Phys. 33, 1345—1352 (1971). 12. P. Amayenc and C. A. Reddy, Height structure of tidal winds as inferred from incoherent scatter observations, Planet. Space Sci. 20, 1269—1279 (1972). 13. A. Brekke, J. R. Doupnik, and P. M. Banks, A preliminary study of the neutral wind in the auroral E region, J. Geophys. Res. 78, #34, 8235—8250 (1973). 14. J. E. Salah, J. V. Evans, and R. H. Wand, Seasonal variations in the thermosphere above Millstone Hill, Radio Sci. 9, #2, 231—238 (1974). 15. P. Amayenc, Tidal oscillations of the meridional neutral wind at midlatitudes, Radio Sci. 9, #2, 281—293 (1974). 16. A. Brekke, J. R. Doupnik, and P. M. Banks, Observations of neutral winds in the aurora! E region during the magnetospheric storm of August 3—9, 1972, J. Geophys. Res. 79, #16, 2448—2456 (1974). 17. R. Bernard, Tides in the E-region observed by incoherent scatter over Saint Santin, J. Atmos. Terr. Phys. 36, 1105—1120 (1974). 18. C. A. Reddy, Evidence of a meridional circulation cell in the lower thermosphere during a magnetic storm, J. Atmos. Terr. Phys. 36, 1561—1564 (1974). 19. J. E. Salah, Daily oscillations of the mid-latitude thermosphere studied by incoherent scatter at Millstone Hill, J. Atmos. Terr. Phys. 36, 1891—1909 (1974). 20. J. E. Salah, R. H. Wand, and J. V. Evans, Tidal effects in the E region from incoherent scatter radar observations, Radio Sci. 10, #3, 347—355 (1975). 21. J. E. Salah, J. V. Evans, and R. H. Wand, E-region temperature measurements at Millstone Hill, J. Atmos. Terr. Phys. 37, 461—489 (1975). 22. R. H. Comfort, S. T. Wu, and G. R. Swenson, An analysis of aurora! E region neutral winds based on incoherent scatter radar observations at Chatanika, Planet. Space Sci. 24, 541—560 (1976). 23. C. L. Rino, A. Brekke, and M. J. Baron, High-resolution auroral zone E region neutral wind and current

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25. K. Schlegel, H. Kohl, and K. Rinnert, Temperatures and collision frequency in the polar E region measured with the incoherent scatter technique, J. Geophys. Res. 85, #A2, 710—714 (1980). 26. R. H. Wand, Seasonal variations of lower thermospheric winds from the Millstone Hill incoherent scatter radar, J. Geophys. Res. 88, #A11, 9227—9241 (1983). 27. R. H. Wand, Geomagnetic activity effects on semidiurnal winds in the lower thermosphere, J. Geophys. R 88, #A11, 9243—9248 (1983). 28. C. Lathuillere, V. B. Wickwar, and W. Kofman, Incoherent scatter measurements of ion-neutral collision frequencies and temperatures in the lower thermosphere of the aurora! region, J. Geophys. Res. 88, #Ab2, 10137—10144 (1983). 29. C. Ma.zaudier and R. Bernard, Saint-Santin radar observations of lower thermosphericstorms, J. Geophys 00, #A3, 2885—2895 (1985). 30. W. Kofman and C. Lathuillere, Eiscat multipulse technique and its contribution to aurora! ionosphere and thermosphere description, J. Geophys. Res. 90, #A4, 3520—3524 (1985). 31. W. Kofman, C. Lathuillere, and B. Pibaret, Neutral atmosphere studies in the altitude range 90—110 km by EISCAT, J. Atmos. Terr. Phys. 48, 837—848 (1986). 32. R. M. Johnson, V. B. Wickwar, R. G. Roble, and J. G. Luhmann, Lower-thermospheric winds at high latitude: Chatanika radar observations, Ann. Geophys. 5A, #6, 383—404 (1987). 33. T. L. Killeen, B. Nardi, F. G. McCormac, J. W. Meriwether, Jr., J. P. Thayer, R. G. Roble, T. J. Fuller-Rowe!!, and D. Rees, Lower thermospheric structure and dynamics inferred from satellite and ground-based Fabry-Perot observations of the O(’S) green line emission, this issue. 34. J. V. Evans, Theory and practice of ionosphere study by Thomson scatter radar, Proc. IEEE 57, # 4, 496—530 (1969). 35. J. D. Kelly and V. B. Wickwar, Radar measurements of high-latitude ion composition between 140 and 300 km altitude, J. Geophys. Res. 86, #A9, 7617—7626 (1981). 36. R. W. Schunk and J. C. G. Walker, Theoretical ion densities in the lower ionosphere, Planet. Space Sci. 21, 1875—1896 (1973). 37. V. B. Wickwar, 3. W. Meriwether, Jr., P. B. Hays, and A. F. Nagy, The meridional thermospheric neutral wind measured by radar and optical techniques in the aurora! region, 3. Geophys. Res. 89, #A12, 10987—10998 (1984). 38. A. Hedin, A revised thermospheric model based on mass spectrometer and incoherent scatter data: MSIS—83, J. Geophys. Res. 88, #A12, 10170—10188 (1983). 39. W. Kofman, F and E regions studies by incoherent scatter radars, this issue. 40. D. Offerman, V. Friedrich, P. Ross, and U. Von Zahn, Neutral gas composition measurements between 80 and 120 km, Planet. Space Sci. 29, 747—764 (1981). 41. R. C. Roble, R. E. Dickinson, and E. C. Ridley, Global circulation and temperature structure of thermosphere with high-latitude plasma convection, 3. Geophys. Res. 87, #A3, 1599—1614 (1982). 42. C. L. Rino, M. J. Baron, C. H. Burch, and 0. de Ia Beaujardière, A multipulse correlator design for incoherent-scatter radar, Radio Sci. 9, 1117—1127 (1974). 43. R. H. Wand and J. V. Evans, Seasonal and magnetic activity variations of ionospheric electric fields over Millstone Hill, J. Geophys. Res. 86, #A1, 103—118 (1981). 44. R. R. Garcia and S. Solomon, The effect of breaking gravity waves on the dynamics and chemical composition of the mesosphere and lower thermosphere, J. Geophys. Res. 90, #D2, 3850—3868 (1985). 45. 3. C. Foster, J. M. Holt, R. G. Musgrove, and D. S. Evans, Ionospheric convection associated with discrete levels of particle precipitation, Geophys. Res. Lett. 13, #7, 656—659 (1986). 46. D. Rees, Winds and temperatures in the aurora! zone and their relations to geomagnetic activity, Phil. Trans. Roy. Soc. Lond. A. 271, 563—575 (1972). 47. D. Rees, Rocket measurements of annual mean prevailing diurnal and semi-diurnal winds in the lower thermosphere at mid-latitudes, J. Geomag. Geoelectr. 31, 253—265 (1979). 48. T. S. Virdi and P. 3. S. Williams, EISCAT Observations of Tides in the E-region, this issue.