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Altitude variations in the amplitude and phase of tidal oscillations at high latitude
Juurnulo/Afmospher~and Terrrsrrrul
Phpm,
Vol.
55, No.
4/5, pp.
697-117,
1993.
0021-9169/93 $6.00+ .OO CC 1993 Per@mon Press Ltd
Printed in Great Bnkin.
Altitude variations in the amplitude and phase of tidal oscillations at high latitude T. S. VIRDI and P. J. S. WILLIAMS Physics Department (Adran Ffiseg), University College of Wales (Coleg Prifysgol Cymru), Aberystwyth SY23 3BZ, Wales, U.K. (Received infinalform
23 January 1992; accepted 2 June 1992)
Abstract-Measurements of ion velocity from EISCAT Common Programme Two (CP-2) have been used to determine the profile of neutral winds in the high latitude E-region, following the method proposed by BREKKEet al. [(1974) J. geophys. Res. 78, 82351. CP-2 has a number of advantages over the previouslyused Common Programme One (CP- 1) : in particular it provided a complete profile of neutral wind velocity whereas CP-1 only determined values at four separate heights. The time variation of neutral wind velocity at each height in turn was analysed into diurnal, semi-diurnal and ter-diurnal components. The height profile of amplitude and phase for each component was interpreted in terms of tidal modes and the results were compared with previous results from EISCAT and the Chatanika incoherent-scatter radar and the theoretical model derived by FORBES[(1982) J. geophys. Res. 87, 5222 and 52411.
1. INTRODIJCIION EISCAT studies of tidal oscillations in the lower thermosphere have been described in a number of recent papers (VIRDI et al., 1986 ; VIRDI and WILLIAMS, 1989 ; WILLIAMSand VIRDI, 1989 ; JOHNSONand VIRDI, 1991; KUNITAKE and SCHLEGEL, 1991; HUUSKONEN et al., 199 1). These studies were mainly based on data taken by EISCAT using Common Programme One (CP-1). In this experimental routine the Tromss antenna was fixed, pointing along the magnetic field line and the estimates of ion velocity in the lower thermosphere were based on tristatic measurements at four heights in the E-region. Over the same period, EISCAT also operated regularly in the Common Programme Two (CP-2) mode, where the Tromss antenna scanned through four different pointing directions in a 6-min cycle. In the present study data from CP-2 have been used to measure tidal oscillations. CP-2 has a number of important advantages over CP-1 and in Section 2 the two experiments are compared in detail. Section 3 describes how the velocity of the neutral wind was calculated from the measured ion velocity and the limitations of this method are assessed. Section 4 explains how the data were selected, while Section 5 describes the analysis of the data into separate tidal components, including the ter-diurnal as well as the diurnal and semi-diurnal modes. Finally in Section 6 the results are compared with previous studies
at EISCAT, at Chatanika incoherent-scatter radar (JOHNSON et al., 1987) and the values predicted by FORBES (1982a, b).
2. EISCAT CP-1-F AND
CP-2-D EXPERIMENTS
Since 1981 EISCAT has assembled an impressive collection of synoptic data using various Common Programmes, which include CP-1 and CP-2-two experimental routines which are both suitable for monitoring tidal oscillations in the lower thermosphere. In the version CP-1-F the Tromso antenna pointed in a fixed direction approximately parallel to the geomagnetic field lines. The remote antennae at Kiruna and SodankylH scanned up and down the Tromss beam in a cycle lasting 10 min during which they received signals from five different scattering volumes, at altitudes of approximately 100, 110, 120, 130 and 279 km. The dwell time of the remote antennae at each E-region altitude was 90 s while at 279 km the remote sites continued measurements for about 5 min. The components of velocity measured simultaneously at the three sites were then combined to give full vectors of ion velocity at each height in turn. The ion velocity at 279 km provided a measure of the electric field perpendicular to the magnetic field line and knowing this it was possible to make the appropriate correction to the measured ion velocities in the E-
691
698
T. S. VIRDI and P. J. S. WILLIAMS
region and hence determine the neutral winds in this height range. These experiments had a number of limitations. In the first place the velocity measurements at different altitudes were made at different times. As the neutral wind was determined after correcting for the influence of electric field on the observed E-region ion velocity it was preferable that both should be measured simultaneously but this was not possible with CP-1. A second limitation of CP-1 was that for each of the four E-region altitudes, the remote site antennae made observations for only 90 s in each 10 min cycle. During the night time when the ion density at these altitudes was low this dwell time was not sufficient to allow a reliable measurement of velocity, especially at Sodankyli where the signal-to-noise ratio was lowest. As a result there were often large gaps in the data set. A third limitation of CP-1 was that it only provided velocities at four altitudes in the E-region, separated by about 10 km. However, to obtain an unambiguous description of the altitudinal variations of tidal modes in the E-region, the whole profile should be observed. In January 1987 a new version of CP-2 was introduced. The 6-min scanning cycle was unchanged from previous CP-2 versions, with the Tromss antenna providing line-of-sight measurements in four directions in sequence (vertical, south, southeast and along the field line) while the remote site antennae at Kiruna and Sodankyla tracked the Tromso beam at 279 km altitude to provide continuous tristatic measurements of the electric field. However, CP-2-D incorporated the GEN routines which used nested multi-pulse coding to improve the range- and time-resolution of E-region measurements (TURUNEN, 1986). Ion velocities can be determined from CP-2 data in two different ways (WILLIAMS et al., 1984). Using the tristatic method, signals from a height of 279 km are received simultaneously at Tromss, Kiruna and SodankylH and the analysed results combined to give an unbiased estimate of the F-region ion velocity and hence the electric field. Alternatively, Tromss line-ofsight measurements at a constant height for three or more pointing directions can be combined to give an estimate of three orthogonal components of velocity. This is called the monostatic method and it is valid provided the velocity is constant with space over the area covered by the scan and constant with time over the 6-min cycle. If, however, there are significant changes in either of these factors then the estimated plasma velocity will contain a spurious component. It is now recognised that the electric field at high latitudes often varies substantially over a period of 6 min or less (WILLIAMS et al., 1990). This means that the monostatic method cannot be used to measure ion
velocity during any period of aurora1 disturbance and only simultaneous tristatic measurements can provide a valid estimate of the electric field. For the same reason it is essential that the electric field measurement in the F-region should be made at exactly the same time as the measurement of ion velocity in the Eregion : if there is a time-lag between the two measurements then any change in the electric field during this time will lead to serious errors when the E-region ion velocity is corrected to give an estimate of the neutral wind. CP-2 has the advantage that electric field measurements are made continually and although these measurements are made at four different locations, separated by up to 120 km, there is good evidence that short-lived bursts in plasma velocity are usually well correlated over such distances, so even if the magnitude of such a burst varies from position to position the presence of the burst will almost always be recognised and if it is thought that an accurate correction for electric field is not practical then the data point can be rejected. In CP-2 no E-region measurements are made at Kiruna and Sodankyla so velocities at these heights can only be obtained using the monostatic method. As it happens, the tidal components we wish to measure have periods of 8 h or longer so provided the electric$eld is monitored continually and any intervals of large electricfield eliminated the monostatic method is entirely valid for measuring E-region neutral winds. Moreover, for the study of neutral winds CP-2-D had several other advantages over CP-1. For example, the Tromss E-region data in CP-2 has, on average, a better signal-to-noise ratio than the Kiruna and Sodankyla data in CP- 1. This is the net result of several factors. Tromsa suffers a higher system noise than the remote antennae and nested multiphase codes at different frequencies are at a basic disadvantage compared with long pulses as only some cross-products of the received data samples provide valid information. However, these disadvantages are more than offset by the greater scattering volume observed by the monostatic method and the shorter distance between the scattering volume and the receiving antenna. As a result there were fewer large gaps in the time sequence of E-region data provided by CP-2-D when compared with CP-1-F. Finally CP-2-D provided plasma velocities over the whole altitude range between 100 and 200 km with an altitude resolution of 4.5 km which must be compared with only four ‘spot’ values of tristatic velocities from CP-1 at a spacing of 10 km. For these reasons data from CP-2-D were used in this new study.
Tidal oscillations at high latitude 3. DETERMINATION
OF NEUTRAL
CP-2 MEASUREMENTS
WINDS
vin =4.34x
FROM
OF ION VELOCITY
699
IO-‘%[N>]+4.28x
lo-‘%[OJ +2.44x
lo-‘%[O]
(2)
EISCAT does not measure
neutral winds directly but instead determines ion velocities. At high latitudes the E-region ion population is driven by two forces : magnetospheric electric fields and neutral winds. The relative importance of the two depends on the ionneutral collision frequency which varies considerably with height within the E-region. Thus at altitudes above about 130 km, the ion-neutral collision frequency is smaller than the ion-gyro frequency and so the ions move readily under the influence of electric field while at lower altitudes the neutral wind is the main driving mechanism. To use ion velocity as a tracer for neutral wind the effect of electric field on the ions must be calculated. This was first done by Brekke et al. in 1974. They showed that the measured ion velocity at an appropriate altitude can be used to determine the velocity of the neutral wind providing the electric field and ion-neutral collision frequency are known. If the two components of ion velocity perpendicular to the magnetic field line are represented by v~*,~ and vln.E in the E-region and v~,~,~and vl+ in the Fregion, the corresponding components of neutral velocity, u, and u, are given by :
us == VI,E
+ K,
’ VL~,F -
K,
* vL.~,F
(14
and
n, = v,,,,fK,.v,,~,--K,.v,,,~
(lb)
where K, = wi~vi,/(w’+v~) and
where the number densities n[N,], n[O,] and n[O] are given in units of molecules m- ‘. This method of determining neutral winds should be limited to relatively quiet geomagnetic conditions. At high latitudes geomagnetic activity can play an important part in the dynamics of the lower thermosphere. In addition to the problem of measuring neutral winds in the presence of strong electric fields, during periods of high activity the neutral winds themselves are affected in two different ways. Firstly under such conditions the ion motion perpendicular to the magnetic field lines in the aurora1 zone is greatly enhanced due to the electric field and through collisions with neutral molecules the ion population transfers substantial momentum to the neutral atmosphere. Secondly the energy input from the magnetosphere heats the lower thermosphere via joule heating and this also perturbs the neutral wind. The overall effect on the neutral wind in the aurora1 zone will have strong diurnal and semi-diurnal components (MIKKELSEN et al., 1981). It follows that to study genuine tidal oscillations, which are generated by solar UV energy input at lower altitudes, incoherent-scatter data at high latitudes should only be used for periods of low to moderate geomagnetic activity. As a rule data were rejected if the electric field had values greater than 25 mV mm ‘. However, the nature of the measurements imposes another limitation. Under very quiet conditions, the ion density in the E-region at night is usually insufficient to produce scattering and large gaps occur in the data. In this study, therefore, suitable data sets were only available for days of moderate geomagnetic activity. 4. DATA
Kz = c$/(c$
+ v,‘,).
wi and v,, are, respectively, the ion-gyro and ionneutral collision frequencies. The value of oi used in this study was calculated for a mean ion mass of 30.5 a.m.u. The height profile of vin is difficult to measure directly using routine EISCAT observations but several authors have estimated this parameter using special experimental modes (KIRKWOOD, 1986; HUUSKONEN et al., 1986). Alternatively this parameter can be estimated by using empirical models of the neutral atmosphere, such as MSIS-86 (HEDIN, 1987), to indicate the number densities of N,, O2 and 0 at each height and then calculating vi,, from the equation defined by SCHUNK and WALKER (1973) :
Between 1987 and 1990 the CP-2-D experiment was run on 20 occasions for a total of over 33 days. On most occasions the experiment ran continuously for at least 36 h. Twice, in April 1988 and in August 1989, CP-2-D was run for four days continuously as part of the Worldwide Atmospheric Gravity-wave Study (WAGS). The days when EISCAT ran in CP-2-D mode between 1987 and 1990 are listed in Table 1 with the corresponding Z Kp values. During seven of the twenty runs the magnetic activity and the electric fields were too large to allow tidal and magnetospheric effects to be separated. On five occasions, in contrast, very low ion densities in the E-region, especially at night times, or technical problems resulted in large gaps in the data.
T. S. VIRDI and P. J. S. WILLIAMS
700
Table 1. Days between 1987 and 1990 when EISCAT ran CP-2-D. Days in italic typescript are those used in the present study. The right column gives C Kp for the corresponding day 1987 January (20,21) February (17, 18) March (If, 18) May (OS,06) June (23) October (20, 21, 22)
25,18+ 17-, 1517-, 19+ 9-, 19f 16 13,13+, 22-
1988 April (II, 12, 13, 14, 15)
June (13.14.15) Au& (09, 10)’ August (16, I?, 18)
November (15, 16, 17)
17+, 23,18+, 18, II+ 7, lo+, 27 22-, 1616, io-, 1722,25+, 19
1989 Febnravy (24, 15)
April (25, 26) August (28,29, 30,SI) September (01) October (23,24,25)
21,20+ 32-, 46 21+, 38-‘, 24’) 12+ 15+ 26-, 2.5+, 25
1990
March (20,21, 22) May (21, 22) October (24,25)
November (13, 14, 15) November (20, 21)
30-, 46,32-l 26+, 3429, 16 6,2+, 6 13+, 18
The remaining eight runs, corresponding to 16 days of data, were used in the present study. These days are listed in Table 1 in italic typeface. Unfortunately they were not spread evenly over the year and most are clustered around April and August. Therefore only a tentative study of the seasonal variations could be conducted with this data set. Ion velocities were determined between 100 and 135 km at 5 km intervals and at 279 km using the monostatic method ~WILLIAMS et al., 1984). As Tromsa antenna positions 1, 2 and 3 (pointing vertically, to the south and to the southeast) presented the best available geometry, line-of-sight velocities from these positions were used to calculate the ion velocities perpendicular to the field line. The velocities were then averaped over intervals of 10 min. The values derived at a height of 279 km were used to make the appropriate correction to the E-region ion velocities [equations (la) and (lb)] to calculate neutral wind (Section 3). Tristatic measurements of ion velocity, made simultaneously at 279 km, were used to monitor the electric field. When the electric field was large, subject to rapid changes, the corresponding data were rejected.
5. NEUTRAL WINDS
AND TIDAL
COMPONENTS
There are two ways in which the average tidal components can be determined from EISCAT measurements of neutral winds. The data from each day can be analysed separately and the amplitudes and phases determined in this way can be averaged over the whole data set (or an appropriate subset). Alternatively, the original measurements of neutral wind can be averaged first and the analysis into tidal components performed subsequently. In our previous studies only the first method was used. This did pose a problem on one or two occasions when large genuine variations in the phase of a component combined with large errors in measurement to give a value of the phase which was ambiguous (i.e. it was uncertain whether on a few exceptional days a particular component was early or late when compared with the tidal components indicated by the rest of the data, see VIRDI and WILLIAMS, 1989 ; WILLIAMS and VIRDI, 1989). This problem does not arise using the second method as the actual velocity measurements at a given local time are unambiguous. Both
methods were used in the present study. For all the days listed in Table 1 as suitable for studying tidal components, the measurements of neutral wind at a given local time and a given height were averaged to indicate the prevailing diurnal variation of the neutral wind in the lower thermosphere above Tromss. Figure I shows the zonal and the meridional winds in contour form where positive values correspond to winds towards the west and south, respectively. For clarity of presentation, the data used in this figure were averaged again, this time over intervals of 1 h in local time. Diurnal and semi-diurnal variations in the two wind components are clearly visible and the descending phase-fronts of a tidal mode are also prominent, especially for the semi-diurnal component. The process was then repeated after dividing the original data set into subsets according to season. Unfortunately, the data available were not sufficient to allow division into summer, winter and equinox subsets. Consequently, data from June, August and early September were grouped together to represent summer months and data from October, February and April were combined to represent winter and equinox. In each case, the average time variations of the wind at a given altitude were fitted to a mathematical function consisting of a mean wind A,, a diurnal variation with amplitude AZ4 and phase +24, a semi-diurnal variation with amplitude A, z and phase 4, 2 and a terdiurnal variation with amplitude A, and phase rt)*.
Tidal oscillations
Uns (m/s)
701
135 %25 5 2 115
67 2; -33 -67
at high latitude
,’
$105 d”
Uew (m/s)
1
I
I
k
bl
I
I
I
135 g125
?3 -3:
g 115
-67
@ 105 I”
I
0
;6
10
Fig. I. Contours of the meridional (top) and zonal (bottom) winds with respect to time and height. Positive values are towards south and west, respectively. For clarity. the data are averaged over one hour.
b
Tidal oscillations
Thus, in the meridional
J4,fit= &,+A,,,*
plane :
~0s wf
-cos (2n[r-$,,,1/12)+A,,~cos
- &24.71/24)+ A, 2r (27c[f--&J/8).
(3)
A similar expression was used for the zonal component u,,~,. A least-squares method was used in fitting. The results are shown in Figs 2(a-d), 3(ad) and 4(a-d) for the whole data set and for summer and winter/ equinox, respectively. In the meridional plane the mean wind below 115 km is very small in all cases, but above 120 km there is a pronounced southward wind which is especially strong in summer. In the zonal plane there is a clear reversal of direction with height, with an eastward wind below 115 km and a westward wind above. Again these winds appear to be stronger in summer. The diurnal variation also shows a similar pattern throughout the year. The meridional component has a maximum amplitude at 115 km and there is a slow variation of phase with height. For the zonal component the maximum amplitude is higher and there is a very sharp phase change between 110 and 120 km. This phase profile demonstrates one of the advantages of using CP-2. With the height-profile sampled every 5 km it was possible to follow the phase profile unambiguously. Using CP-1, in contrast, it was only possible to determine the phase at 110 and 120 km so that the phase profile between these heights was ambiguous. It now seems that in a previous study (VIRDI and WILLIAMS, 1989 ; WILLIAMS and VIRDI, 1989), of two alternative profiles linking the data at 110 and 120 km, the wrong one was chosen. In the case of the semi-diurnal tide there is evidence of a difference between summer and winter equinox. Below 120 km both show a mode which reaches a maximum amplitude at 105 km in the meridional plane and at 110 km in the zonal plane, but drops sharply in amplitude at greater heights. This mode shows a fairly rapid variation of phase with height, with a relatively short vertical wavelength, which is compatible with the (2,4) mode. Above 120 km, however, there is some evidence that in winter there is a second mode. HUUSKONEN et al. (1991) observed a similar pattern in an extended series of observations during March 1988 and pointed out that the phase variation of the semi-diurnal component above 120 km resembled that of a (2,2) tidal model. This mode, generated by solar heating at lower heights in the atmosphere, is attenuated sharply in the mesosphere, where it is evanescent and only a vestigial component of the (2,2) mode ‘tunnels’ through to the thermo-
at high latitude
703
sphere. Once in the lower thermosphere, however, this mode is again free to propagate with an amplitude which initially increases exponentially with height until at 130 km it becomes the dominant mode. There is no evidence of a similar mode in the summer observations which may reflect even greater attenuation of the (2, 2) mode in the summer mesosphere. The semi-diurnal component above 120 km in summer seems much closer to a tidal extension mode of the (2,4) component, as predicted theoretically by FORBESand HAGAN (1982). Finally, the CP-2 data, averaged over many days, were of a high enough quality to allow us to fit a terdiurnal component. The most important result is that the ter-diurnal mode is much weaker than the others (note that the scale in Figs 2d, 3d and 4d has been increased by a factor of 3). There is evidence that this
MO-
130 -
T S
120 -
5 f
110 -
100 -
go-100
140
-50 &,s
0 (m/s>
50
loo
-50 &w
0 (m/s)
50
loo
1
130 -
T c 2
.IF
120 -
110 -
P 100 -
go-100
Fig. 2a. Height profiles of the mean meridional (top) and zonal (bottom) winds averaged over all days.
704
T. S.
VIRDI
and P. J. S. WILLIAMS
MOX30-
g
l30-
‘2O-
.r;l 1102
g
‘20-
$
llo-
P
100 -
90 , 0
lOO, 30
, f 60 90 A24s (m/s)
1
120
ml
I
150
12 “$24~
140 -
WO-
130 -
l30-
‘-i 55
t20-
$5 P
110 100 -
g
‘xl-
z .P 3
no -
&)
6
t2
loo -
901 12 18fi24w
&) ’
t2
Fig. 2b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds averaged over all days.
Tidal oscillations at high latitude
l40l30-
Emz ?.S
120 -
g
*o-
5 z!
110 -
zi I”
llo-
lOO-
l0.i
90’ 0
g z .rr P
90
30
60 90 A72s (m/s)
l20
I
0
150
6p2s
140
l40-
130
UO-
120
g
fLo-
110
ZE .P 2
110 -
i
(E) l8
t
O
lOO-
100 I)
901 0
I
90’ 30
60 90 A12w (m/s)
120
150
0 6 $lZw (ET,
l8
O
Fig. 2c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds averaged over all days.
T. S. VIRDI and
P. J. S. WILLIAMS
l40130
l30-
T‘ &
120
s m-
g
110
5
llo-
2
L
i
x)0-
100
go6
9 Ps
z .P f
140
wo-
130
l30-
110
PO-
5
llo-
15
18
2
100 s
g
12 (LT)
120 1
XIO-
goI 6 g
Fig. 2d. Height profiles of the ter-diurnal amplitude and phase of the meridional winds averaged over all days.
(IF) l5
@3w
l8
(top) and zonal (bottom)
707
Tidal oscillations at high latitude
90
I I -50 A=s &,
-100
I 50
90
I loo
-100
I I -50 Aaow &‘s,
I 50
I 100
Fig. 3a. Height profiles of the mean meridional (left) and zonal (right) winds on summer days.
3 5 P
140-l
'40-j
130 -
l30-
120 nOloo-
‘::
901 0
90
30
60 90 A249 (m/s)
120
150 12
140,
1407
130
130 -
g
mJ-
2
no -
& 5 P T
110
.P
la@249 &)
6
l2
%$24w &I
6
l2
f .._i_-:-
100 120 1!
Qoh
90 0
30
60 A24w @I-$:]
120
150 12
Fig. 3b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonai (bottom) winds on summer days.
7.
708
S. VIRDI and P. J. S. WILLIAMS
MO130 -
g
?20-
g
m-
5 ii
110 -
,F
no-
100-
E
i\-
UYO-
90’ 306090
0
A12s (m/s)
120
140 1
MO-
130 -
l30-
120 -
2 f$
901 0
150
110 loo-90 -0
;'
g
d
p
no-
B$12s(~)
la
O
6 @2w
la
O
MO-
f 30
I I 60 (rn,:;) A12w
120
150
90’
0
(is)
Fig. 3c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on summer days.
Tidal oscillations at high latitude
709
130
loo
I
I
7
6
140 130
5 P
110
100
QOi 6 Fig. 3d. Height profiles of the ter-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on summer days.
710
T. S. VIRDI and P. J. S. WILLIAMS
:;g
{
110
90 -100
I '0
I -50
A&
)
1.
I
I
50
loo
90
I
-100
(m/s>
-50 A,w
1
I
(:/a,
50
1
100
Fig. 4a. Height profiles of the mean meridional (left) and zonal (right) winds on winter/equinox days.
g 5 f
140-l
WI-
130 -
l30-
120 -
g
PO-
110 -
&
tlo-
P
joo-
3
x)0-
goI
0
30
60
90
A24s (m/s)
120
go12
150
140
140
730 -
DO-
g
120 -
S 120
5 P
110 -
.P
2
P 100 -
0
(k)
6
l2
18qJ24w LT)
6
l2
no -.i
x)0-
901
"$249
901 30
60 A24w (m,$
I20
1.50
72
Fig. 4b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on winter/equinox days.
Tidal oscillations at high latitude
711
140 l30
130
3$ I”
no
1
120 ~
100
901 0
30
60 90 AlZs (m/s)
120
750
0
I 30
1 I 60 90 A12w (m/s)
I 120
1 150
90
9o”l----0
90 0
6$12s(~)l6
O
I 6
1
I 12 $12~ (LT)
I 18
0
Fig. 4c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on winter/equinox days.
T. S. VIRDI and P. J. S. WILLIAMS
712
z 5 P
'9
l40-
13
?30-
120
g
m-
110
5
llo-
100 ~
2
i
x)0-
go’
6
901 6
Fig. 4d. Height profiles of the ter-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on winter/equinox days.
713
Tidal oscillations at high latitude mode reaches a maximum at about 120 km but the phase variation with height does not show any clear pattern and the errors in phase are always large.
6. DISCUSSION OF RESULTS To conclude our study we compared the results of the present work with our previous results using CP1 and with the results obtained at Chatanika, an incoherent-scatter radar located at a similar geographic latitude as EISCAT (JOHNSONet al., 1987). The amplitude and phase of the diurnal and semi-diurnal components were also compared with the theoretical values predicted by FORBES(1982a, b). The main comparisons, based on the analysis of the mean wind profile averaged over the whole data set, are summarised in Fig. 5(a-c). For the mean wind there is extremely good agreement between all three sets of data, except for the meridional component above 115 km where the Chatanika data indicate a northward wind while both sets of EISCAT data agree in showing a southward wind which increases with height. There is also a reassuring agreement in the case of the diurnal component-including the sharp phase change in the zonal component between 110 and 120 km. The new data from EISCAT agree well with the Chatanika result and suggest a re-interpretation of the CP-1 data which were published previously (VIRDI and WILLIAMS, 1989 ; WILLIAMS and VIRDI, 1989). The diurnal amplitude and phase values predicted by FORBES’(1982a) model compare well with the three data sets below about 110 km. Above this altitude the predicted amplitudes in both directions are smaller than the measured values. The predicted meridional phase agrees with observation but the zonal phase above 110 km shows large disagreement. Finally, there is very good agreement for the semidiurnal tidal mode, especially in the zonal plane, but two features deserve comment. The new results indicate a very similar profile of amplitude with height for both meridional and zonal components, with a maximum at 110 km. The previous results at EISCAT and the results from Chatanika, suggest that the amplitude profile in the meridional plane reaches a maximum at about 120 km. The difference may be partly due to seasonal effects combined with the uneven distribution of data throughout the year. It must be added, however, that the results quoted for heights above 120 km in the previous EISCAT study were based on very few data points and so the present results must be regarded as more reliable. This may
also explain the second feature: the phase profile of the meridional semi-diurnal mode in the present study agrees perfectly with the Chatanika results but not with the previous EISCAT results at 120 and 130 km. Once again, it is likely that the present results are more reliable. The semi-diurnal amplitude and phase are predicted reasonably well by FORBES’(1982b) model. Meridional and zonal phase show a similar variation with height although the times of the maxima are approximately 2 h later. The model predicts peaks in both meridional and zonal amplitudes at about 110 km. These peaks are observed in the measurements but with much smaller values. The final result is based on the analysis of the tidal components for each separate day. WILLIAMS and
-75
-50
-25 0 A,” (mls)
25
50
75
-75
-50
-25 0 A2 (m/s)
25
50
75
Fig. 5a. Height profiles of the mean meridional (top) and zonal (bottom) winds from the current study (open circles), from JOHNSONet al. (1957) (squares) and from WILLIAMS and VIRDI(1989) (crosses).
T. S. VIRDI and
714
P. J. S. WILLIAMS
140 -
120
120 -
-z
g
25 E
E
.a 2
.o,
2
100
0
25
0
25
50 A24s (m/s)
75
loo-
loo
140
120
E E
ca f
100
80
I
I 50 A24w (m/s)
1 75
I 100
80 12 324w(-$
6
l2
Fig. 5b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds from the current study (open circles), from JOHNSON et al. (1987) (squares), from WILLIAMS and VIRDI (1989) (crosses) and from FORBES’ (1982a) theoretical model (stars).
Tidal oscillations
715
at high latitude
140
120
g
E .En f
lcm
80
I
0
25
0
25
80
I
I
50 Al 2s (m/s)
/
50 Al 2w (m/s)
I
75
I
75
I
I
I
I
I
100
I
100
00
0
Fig. 5c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds from the current study (open circles), from JOHNSON et al. (1987) (squares), from WILLIAMS and VIRDI (1989) (crosses) and from FORBES’ (1982b) theoretical model (stars).
T. S. VIRDI and
716
P. J. S. WILLIAMS 120 -
Table 2. Correlation coefficient between the day-to-day variations in the measured amplitude of the meridional and zonal components of the different tidal modes at different altitudes
lGQ-
Height (56)
Mean
Diurnal
(%)
(%)
105 110 115 120 125 130
81 57 31 39 39 44
16 44 21 27 20 28
Semi-diurnal (X)
0
Ter-diurnal
12 60 50 39 36 36
%
:
(“/) 58 36 17 14 17 15
.
.
0
20-
VIRDI (1989) had examined
the correlation between the day-to-day variation in the amplitude of the meridional and zonal components of each tidal mode. As a result they tentatively concluded that there was very little correlation in the case of the diurnal variation but a much higher correlation in the case of the semidiurnal component, especially at heights below 120 km. This result is confirmed by the present set of results. Table 2 indicates the correlation coefficient between the amplitudes of the meridional and zonal components of each mode at different heights. There appears to be a relatively high correlation at lower heights for the mean wind and for the semi-diurnal component, but much lower correlation for the diurnal and ter-diurnal modes. In all cases the correlation coefficient tends to drop with height, but it is not yet certain whether this is an intrinsic property of the tidal modes or an effect due to the larger errors in the analysis of tidal modes at greater heights. The correlation in the case of the semi-diurnal mode at 105, 110 and 115 km is illustrated in Fig. 6. At these heights the phase of the two components of the semi-diurnal mode differs by about 3 h. To express this in very simple terms, we can say that the semi-diurnal mode below 120 km usually has a similar amplitude in the meridional and zonal planes and the two components are in quadrature ; in other words, the mode shows a There is no similar conclusion circular polarisation.
0
I
0
/ 20
I
I
loo
1
120
A12 W:'t (m/s?
Fig. 6. The correlation between daily values of the amplitude of the semi-diurnal tide in the meridional and zonal planes as measured by EISCAT CP-2 at 105 (squares), 110 (crosses) and 115 km (open circles).
to be drawn for -the diurnal mode where the day-today variations in the amplitude of the meridional and zonal components seem almost independent and the phase variations are much larger. It is hoped that in the near future EISCAT will introduce alternating codes into the operation of CP2 and this will provide E-region data of an even higher quality. There are also plans to organise even longer continuous runs which will allow a better study of the day-to-day variability of the tidal modes. In the meantime the present study removes some of the discrepancies noted in previous work and provides a reliable outline of the tidal modes in the lower thermosphere at high latitudes. Acknowledgenzentsm-The authors would like to thank the EISCAT Director and his staff for their help in measurement and analysis of the data. EISCAT is a scientific association
supporteh by the research councils of Finland, France, Germany. . , was ., Norwav. __Sweden and the U.K. One of us (TSV) supported by the SERC carried out.
during
REFERENCES BREKKE A., DOUPNIK J. R. and BANKS P. M. FORBES J. M. FORBES J. M. FORBES J. M. and HAGAN A. E. HEDIN A. E. HUUSKONEN A., NYGREN T., JALONEN T., TLIRUNENT. and GLEN J. HUUSKONENA., VIRDI T. S., JONES G. 0. L. and WILLIAMS P. J. S.
I
8
40
1974 1982a 1982b 1982 1987 1986
J. J. J. J. J. J.
geophys. Res. 78, 8235. geophys. Res. 87, 5222. geophys. Res. 87,5241. geophys. Res. 81, 5253. geophys. Ref. 92,4649. atmos. ierr. Phys. 48, 827.
1991
Ann. Geophys. 9,407.
the period
this study
was
Tidal oscillations at high latitude JOHNSON R. M. and VIRLX T. S. JOHNSON R. M., WICKWAR V. B., ROBLER. G. and LUHMANNJ. G. KIRKW~~D S. KUNITAKEM. and SCHLEGEL K. MIKKEUENI. S., JORGENSEN T. S., KELLEYM. C., LARSENM. E., PEIREIRA E. and VICICREY J. SCHUNKR. W. and WALKERJ. C. G. TURUNENT. VIRDIT. S., JONESG. 0. L. and WILLIAMSP. J. S. VIRDI T. S. and WILLIAMSP. J. S. WILLIAMSP. J. S., JONESG. 0. L. and JAINA. R. WILLIAMSP. J. S., JONEXG. 0. L., JONESB., OPGENNCKIRTH H. and HAGGSTR~MI. WILLIAMSP. J. S. and VIRDI T. S.
1991 1987
J. geophys. Res. %, 1099. Ann. Geophys. SA, 383.
1986 1991 1981
J. atmos. terr. Phys. 48,817. Ann. Geophys. 9, 143. J. geophys. Res. 86, 15 13.
1973 1986 1986 1989 1984 1990
Planet. Space Sci. 21, 1875. J. atmos. terr. Phys. 48, 777. Nature 324,354. Adv. Space Res. 9(5), 83. J. atmos. terr. Phys. 86, 521. J. atmos. terr. Phys. 52,439.
1989
J. atmos. terr. Phys. 51, 569.
717
Phpm,
Vol.
55, No.
4/5, pp.
697-117,
1993.
0021-9169/93 $6.00+ .OO CC 1993 Per@mon Press Ltd
Printed in Great Bnkin.
Altitude variations in the amplitude and phase of tidal oscillations at high latitude T. S. VIRDI and P. J. S. WILLIAMS Physics Department (Adran Ffiseg), University College of Wales (Coleg Prifysgol Cymru), Aberystwyth SY23 3BZ, Wales, U.K. (Received infinalform
23 January 1992; accepted 2 June 1992)
Abstract-Measurements of ion velocity from EISCAT Common Programme Two (CP-2) have been used to determine the profile of neutral winds in the high latitude E-region, following the method proposed by BREKKEet al. [(1974) J. geophys. Res. 78, 82351. CP-2 has a number of advantages over the previouslyused Common Programme One (CP- 1) : in particular it provided a complete profile of neutral wind velocity whereas CP-1 only determined values at four separate heights. The time variation of neutral wind velocity at each height in turn was analysed into diurnal, semi-diurnal and ter-diurnal components. The height profile of amplitude and phase for each component was interpreted in terms of tidal modes and the results were compared with previous results from EISCAT and the Chatanika incoherent-scatter radar and the theoretical model derived by FORBES[(1982) J. geophys. Res. 87, 5222 and 52411.
1. INTRODIJCIION EISCAT studies of tidal oscillations in the lower thermosphere have been described in a number of recent papers (VIRDI et al., 1986 ; VIRDI and WILLIAMS, 1989 ; WILLIAMSand VIRDI, 1989 ; JOHNSONand VIRDI, 1991; KUNITAKE and SCHLEGEL, 1991; HUUSKONEN et al., 199 1). These studies were mainly based on data taken by EISCAT using Common Programme One (CP-1). In this experimental routine the Tromss antenna was fixed, pointing along the magnetic field line and the estimates of ion velocity in the lower thermosphere were based on tristatic measurements at four heights in the E-region. Over the same period, EISCAT also operated regularly in the Common Programme Two (CP-2) mode, where the Tromss antenna scanned through four different pointing directions in a 6-min cycle. In the present study data from CP-2 have been used to measure tidal oscillations. CP-2 has a number of important advantages over CP-1 and in Section 2 the two experiments are compared in detail. Section 3 describes how the velocity of the neutral wind was calculated from the measured ion velocity and the limitations of this method are assessed. Section 4 explains how the data were selected, while Section 5 describes the analysis of the data into separate tidal components, including the ter-diurnal as well as the diurnal and semi-diurnal modes. Finally in Section 6 the results are compared with previous studies
at EISCAT, at Chatanika incoherent-scatter radar (JOHNSON et al., 1987) and the values predicted by FORBES (1982a, b).
2. EISCAT CP-1-F AND
CP-2-D EXPERIMENTS
Since 1981 EISCAT has assembled an impressive collection of synoptic data using various Common Programmes, which include CP-1 and CP-2-two experimental routines which are both suitable for monitoring tidal oscillations in the lower thermosphere. In the version CP-1-F the Tromso antenna pointed in a fixed direction approximately parallel to the geomagnetic field lines. The remote antennae at Kiruna and SodankylH scanned up and down the Tromss beam in a cycle lasting 10 min during which they received signals from five different scattering volumes, at altitudes of approximately 100, 110, 120, 130 and 279 km. The dwell time of the remote antennae at each E-region altitude was 90 s while at 279 km the remote sites continued measurements for about 5 min. The components of velocity measured simultaneously at the three sites were then combined to give full vectors of ion velocity at each height in turn. The ion velocity at 279 km provided a measure of the electric field perpendicular to the magnetic field line and knowing this it was possible to make the appropriate correction to the measured ion velocities in the E-
691
698
T. S. VIRDI and P. J. S. WILLIAMS
region and hence determine the neutral winds in this height range. These experiments had a number of limitations. In the first place the velocity measurements at different altitudes were made at different times. As the neutral wind was determined after correcting for the influence of electric field on the observed E-region ion velocity it was preferable that both should be measured simultaneously but this was not possible with CP-1. A second limitation of CP-1 was that for each of the four E-region altitudes, the remote site antennae made observations for only 90 s in each 10 min cycle. During the night time when the ion density at these altitudes was low this dwell time was not sufficient to allow a reliable measurement of velocity, especially at Sodankyli where the signal-to-noise ratio was lowest. As a result there were often large gaps in the data set. A third limitation of CP-1 was that it only provided velocities at four altitudes in the E-region, separated by about 10 km. However, to obtain an unambiguous description of the altitudinal variations of tidal modes in the E-region, the whole profile should be observed. In January 1987 a new version of CP-2 was introduced. The 6-min scanning cycle was unchanged from previous CP-2 versions, with the Tromss antenna providing line-of-sight measurements in four directions in sequence (vertical, south, southeast and along the field line) while the remote site antennae at Kiruna and Sodankyla tracked the Tromso beam at 279 km altitude to provide continuous tristatic measurements of the electric field. However, CP-2-D incorporated the GEN routines which used nested multi-pulse coding to improve the range- and time-resolution of E-region measurements (TURUNEN, 1986). Ion velocities can be determined from CP-2 data in two different ways (WILLIAMS et al., 1984). Using the tristatic method, signals from a height of 279 km are received simultaneously at Tromss, Kiruna and SodankylH and the analysed results combined to give an unbiased estimate of the F-region ion velocity and hence the electric field. Alternatively, Tromss line-ofsight measurements at a constant height for three or more pointing directions can be combined to give an estimate of three orthogonal components of velocity. This is called the monostatic method and it is valid provided the velocity is constant with space over the area covered by the scan and constant with time over the 6-min cycle. If, however, there are significant changes in either of these factors then the estimated plasma velocity will contain a spurious component. It is now recognised that the electric field at high latitudes often varies substantially over a period of 6 min or less (WILLIAMS et al., 1990). This means that the monostatic method cannot be used to measure ion
velocity during any period of aurora1 disturbance and only simultaneous tristatic measurements can provide a valid estimate of the electric field. For the same reason it is essential that the electric field measurement in the F-region should be made at exactly the same time as the measurement of ion velocity in the Eregion : if there is a time-lag between the two measurements then any change in the electric field during this time will lead to serious errors when the E-region ion velocity is corrected to give an estimate of the neutral wind. CP-2 has the advantage that electric field measurements are made continually and although these measurements are made at four different locations, separated by up to 120 km, there is good evidence that short-lived bursts in plasma velocity are usually well correlated over such distances, so even if the magnitude of such a burst varies from position to position the presence of the burst will almost always be recognised and if it is thought that an accurate correction for electric field is not practical then the data point can be rejected. In CP-2 no E-region measurements are made at Kiruna and Sodankyla so velocities at these heights can only be obtained using the monostatic method. As it happens, the tidal components we wish to measure have periods of 8 h or longer so provided the electric$eld is monitored continually and any intervals of large electricfield eliminated the monostatic method is entirely valid for measuring E-region neutral winds. Moreover, for the study of neutral winds CP-2-D had several other advantages over CP-1. For example, the Tromss E-region data in CP-2 has, on average, a better signal-to-noise ratio than the Kiruna and Sodankyla data in CP- 1. This is the net result of several factors. Tromsa suffers a higher system noise than the remote antennae and nested multiphase codes at different frequencies are at a basic disadvantage compared with long pulses as only some cross-products of the received data samples provide valid information. However, these disadvantages are more than offset by the greater scattering volume observed by the monostatic method and the shorter distance between the scattering volume and the receiving antenna. As a result there were fewer large gaps in the time sequence of E-region data provided by CP-2-D when compared with CP-1-F. Finally CP-2-D provided plasma velocities over the whole altitude range between 100 and 200 km with an altitude resolution of 4.5 km which must be compared with only four ‘spot’ values of tristatic velocities from CP-1 at a spacing of 10 km. For these reasons data from CP-2-D were used in this new study.
Tidal oscillations at high latitude 3. DETERMINATION
OF NEUTRAL
CP-2 MEASUREMENTS
WINDS
vin =4.34x
FROM
OF ION VELOCITY
699
IO-‘%[N>]+4.28x
lo-‘%[OJ +2.44x
lo-‘%[O]
(2)
EISCAT does not measure
neutral winds directly but instead determines ion velocities. At high latitudes the E-region ion population is driven by two forces : magnetospheric electric fields and neutral winds. The relative importance of the two depends on the ionneutral collision frequency which varies considerably with height within the E-region. Thus at altitudes above about 130 km, the ion-neutral collision frequency is smaller than the ion-gyro frequency and so the ions move readily under the influence of electric field while at lower altitudes the neutral wind is the main driving mechanism. To use ion velocity as a tracer for neutral wind the effect of electric field on the ions must be calculated. This was first done by Brekke et al. in 1974. They showed that the measured ion velocity at an appropriate altitude can be used to determine the velocity of the neutral wind providing the electric field and ion-neutral collision frequency are known. If the two components of ion velocity perpendicular to the magnetic field line are represented by v~*,~ and vln.E in the E-region and v~,~,~and vl+ in the Fregion, the corresponding components of neutral velocity, u, and u, are given by :
us == VI,E
+ K,
’ VL~,F -
K,
* vL.~,F
(14
and
n, = v,,,,fK,.v,,~,--K,.v,,,~
(lb)
where K, = wi~vi,/(w’+v~) and
where the number densities n[N,], n[O,] and n[O] are given in units of molecules m- ‘. This method of determining neutral winds should be limited to relatively quiet geomagnetic conditions. At high latitudes geomagnetic activity can play an important part in the dynamics of the lower thermosphere. In addition to the problem of measuring neutral winds in the presence of strong electric fields, during periods of high activity the neutral winds themselves are affected in two different ways. Firstly under such conditions the ion motion perpendicular to the magnetic field lines in the aurora1 zone is greatly enhanced due to the electric field and through collisions with neutral molecules the ion population transfers substantial momentum to the neutral atmosphere. Secondly the energy input from the magnetosphere heats the lower thermosphere via joule heating and this also perturbs the neutral wind. The overall effect on the neutral wind in the aurora1 zone will have strong diurnal and semi-diurnal components (MIKKELSEN et al., 1981). It follows that to study genuine tidal oscillations, which are generated by solar UV energy input at lower altitudes, incoherent-scatter data at high latitudes should only be used for periods of low to moderate geomagnetic activity. As a rule data were rejected if the electric field had values greater than 25 mV mm ‘. However, the nature of the measurements imposes another limitation. Under very quiet conditions, the ion density in the E-region at night is usually insufficient to produce scattering and large gaps occur in the data. In this study, therefore, suitable data sets were only available for days of moderate geomagnetic activity. 4. DATA
Kz = c$/(c$
+ v,‘,).
wi and v,, are, respectively, the ion-gyro and ionneutral collision frequencies. The value of oi used in this study was calculated for a mean ion mass of 30.5 a.m.u. The height profile of vin is difficult to measure directly using routine EISCAT observations but several authors have estimated this parameter using special experimental modes (KIRKWOOD, 1986; HUUSKONEN et al., 1986). Alternatively this parameter can be estimated by using empirical models of the neutral atmosphere, such as MSIS-86 (HEDIN, 1987), to indicate the number densities of N,, O2 and 0 at each height and then calculating vi,, from the equation defined by SCHUNK and WALKER (1973) :
Between 1987 and 1990 the CP-2-D experiment was run on 20 occasions for a total of over 33 days. On most occasions the experiment ran continuously for at least 36 h. Twice, in April 1988 and in August 1989, CP-2-D was run for four days continuously as part of the Worldwide Atmospheric Gravity-wave Study (WAGS). The days when EISCAT ran in CP-2-D mode between 1987 and 1990 are listed in Table 1 with the corresponding Z Kp values. During seven of the twenty runs the magnetic activity and the electric fields were too large to allow tidal and magnetospheric effects to be separated. On five occasions, in contrast, very low ion densities in the E-region, especially at night times, or technical problems resulted in large gaps in the data.
T. S. VIRDI and P. J. S. WILLIAMS
700
Table 1. Days between 1987 and 1990 when EISCAT ran CP-2-D. Days in italic typescript are those used in the present study. The right column gives C Kp for the corresponding day 1987 January (20,21) February (17, 18) March (If, 18) May (OS,06) June (23) October (20, 21, 22)
25,18+ 17-, 1517-, 19+ 9-, 19f 16 13,13+, 22-
1988 April (II, 12, 13, 14, 15)
June (13.14.15) Au& (09, 10)’ August (16, I?, 18)
November (15, 16, 17)
17+, 23,18+, 18, II+ 7, lo+, 27 22-, 1616, io-, 1722,25+, 19
1989 Febnravy (24, 15)
April (25, 26) August (28,29, 30,SI) September (01) October (23,24,25)
21,20+ 32-, 46 21+, 38-‘, 24’) 12+ 15+ 26-, 2.5+, 25
1990
March (20,21, 22) May (21, 22) October (24,25)
November (13, 14, 15) November (20, 21)
30-, 46,32-l 26+, 3429, 16 6,2+, 6 13+, 18
The remaining eight runs, corresponding to 16 days of data, were used in the present study. These days are listed in Table 1 in italic typeface. Unfortunately they were not spread evenly over the year and most are clustered around April and August. Therefore only a tentative study of the seasonal variations could be conducted with this data set. Ion velocities were determined between 100 and 135 km at 5 km intervals and at 279 km using the monostatic method ~WILLIAMS et al., 1984). As Tromsa antenna positions 1, 2 and 3 (pointing vertically, to the south and to the southeast) presented the best available geometry, line-of-sight velocities from these positions were used to calculate the ion velocities perpendicular to the field line. The velocities were then averaped over intervals of 10 min. The values derived at a height of 279 km were used to make the appropriate correction to the E-region ion velocities [equations (la) and (lb)] to calculate neutral wind (Section 3). Tristatic measurements of ion velocity, made simultaneously at 279 km, were used to monitor the electric field. When the electric field was large, subject to rapid changes, the corresponding data were rejected.
5. NEUTRAL WINDS
AND TIDAL
COMPONENTS
There are two ways in which the average tidal components can be determined from EISCAT measurements of neutral winds. The data from each day can be analysed separately and the amplitudes and phases determined in this way can be averaged over the whole data set (or an appropriate subset). Alternatively, the original measurements of neutral wind can be averaged first and the analysis into tidal components performed subsequently. In our previous studies only the first method was used. This did pose a problem on one or two occasions when large genuine variations in the phase of a component combined with large errors in measurement to give a value of the phase which was ambiguous (i.e. it was uncertain whether on a few exceptional days a particular component was early or late when compared with the tidal components indicated by the rest of the data, see VIRDI and WILLIAMS, 1989 ; WILLIAMS and VIRDI, 1989). This problem does not arise using the second method as the actual velocity measurements at a given local time are unambiguous. Both
methods were used in the present study. For all the days listed in Table 1 as suitable for studying tidal components, the measurements of neutral wind at a given local time and a given height were averaged to indicate the prevailing diurnal variation of the neutral wind in the lower thermosphere above Tromss. Figure I shows the zonal and the meridional winds in contour form where positive values correspond to winds towards the west and south, respectively. For clarity of presentation, the data used in this figure were averaged again, this time over intervals of 1 h in local time. Diurnal and semi-diurnal variations in the two wind components are clearly visible and the descending phase-fronts of a tidal mode are also prominent, especially for the semi-diurnal component. The process was then repeated after dividing the original data set into subsets according to season. Unfortunately, the data available were not sufficient to allow division into summer, winter and equinox subsets. Consequently, data from June, August and early September were grouped together to represent summer months and data from October, February and April were combined to represent winter and equinox. In each case, the average time variations of the wind at a given altitude were fitted to a mathematical function consisting of a mean wind A,, a diurnal variation with amplitude AZ4 and phase +24, a semi-diurnal variation with amplitude A, z and phase 4, 2 and a terdiurnal variation with amplitude A, and phase rt)*.
Tidal oscillations
Uns (m/s)
701
135 %25 5 2 115
67 2; -33 -67
at high latitude
,’
$105 d”
Uew (m/s)
1
I
I
k
bl
I
I
I
135 g125
?3 -3:
g 115
-67
@ 105 I”
I
0
;6
10
Fig. I. Contours of the meridional (top) and zonal (bottom) winds with respect to time and height. Positive values are towards south and west, respectively. For clarity. the data are averaged over one hour.
b
Tidal oscillations
Thus, in the meridional
J4,fit= &,+A,,,*
plane :
~0s wf
-cos (2n[r-$,,,1/12)+A,,~cos
- &24.71/24)+ A, 2r (27c[f--&J/8).
(3)
A similar expression was used for the zonal component u,,~,. A least-squares method was used in fitting. The results are shown in Figs 2(a-d), 3(ad) and 4(a-d) for the whole data set and for summer and winter/ equinox, respectively. In the meridional plane the mean wind below 115 km is very small in all cases, but above 120 km there is a pronounced southward wind which is especially strong in summer. In the zonal plane there is a clear reversal of direction with height, with an eastward wind below 115 km and a westward wind above. Again these winds appear to be stronger in summer. The diurnal variation also shows a similar pattern throughout the year. The meridional component has a maximum amplitude at 115 km and there is a slow variation of phase with height. For the zonal component the maximum amplitude is higher and there is a very sharp phase change between 110 and 120 km. This phase profile demonstrates one of the advantages of using CP-2. With the height-profile sampled every 5 km it was possible to follow the phase profile unambiguously. Using CP-1, in contrast, it was only possible to determine the phase at 110 and 120 km so that the phase profile between these heights was ambiguous. It now seems that in a previous study (VIRDI and WILLIAMS, 1989 ; WILLIAMS and VIRDI, 1989), of two alternative profiles linking the data at 110 and 120 km, the wrong one was chosen. In the case of the semi-diurnal tide there is evidence of a difference between summer and winter equinox. Below 120 km both show a mode which reaches a maximum amplitude at 105 km in the meridional plane and at 110 km in the zonal plane, but drops sharply in amplitude at greater heights. This mode shows a fairly rapid variation of phase with height, with a relatively short vertical wavelength, which is compatible with the (2,4) mode. Above 120 km, however, there is some evidence that in winter there is a second mode. HUUSKONEN et al. (1991) observed a similar pattern in an extended series of observations during March 1988 and pointed out that the phase variation of the semi-diurnal component above 120 km resembled that of a (2,2) tidal model. This mode, generated by solar heating at lower heights in the atmosphere, is attenuated sharply in the mesosphere, where it is evanescent and only a vestigial component of the (2,2) mode ‘tunnels’ through to the thermo-
at high latitude
703
sphere. Once in the lower thermosphere, however, this mode is again free to propagate with an amplitude which initially increases exponentially with height until at 130 km it becomes the dominant mode. There is no evidence of a similar mode in the summer observations which may reflect even greater attenuation of the (2, 2) mode in the summer mesosphere. The semi-diurnal component above 120 km in summer seems much closer to a tidal extension mode of the (2,4) component, as predicted theoretically by FORBESand HAGAN (1982). Finally, the CP-2 data, averaged over many days, were of a high enough quality to allow us to fit a terdiurnal component. The most important result is that the ter-diurnal mode is much weaker than the others (note that the scale in Figs 2d, 3d and 4d has been increased by a factor of 3). There is evidence that this
MO-
130 -
T S
120 -
5 f
110 -
100 -
go-100
140
-50 &,s
0 (m/s>
50
loo
-50 &w
0 (m/s)
50
loo
1
130 -
T c 2
.IF
120 -
110 -
P 100 -
go-100
Fig. 2a. Height profiles of the mean meridional (top) and zonal (bottom) winds averaged over all days.
704
T. S.
VIRDI
and P. J. S. WILLIAMS
MOX30-
g
l30-
‘2O-
.r;l 1102
g
‘20-
$
llo-
P
100 -
90 , 0
lOO, 30
, f 60 90 A24s (m/s)
1
120
ml
I
150
12 “$24~
140 -
WO-
130 -
l30-
‘-i 55
t20-
$5 P
110 100 -
g
‘xl-
z .P 3
no -
&)
6
t2
loo -
901 12 18fi24w
&) ’
t2
Fig. 2b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds averaged over all days.
Tidal oscillations at high latitude
l40l30-
Emz ?.S
120 -
g
*o-
5 z!
110 -
zi I”
llo-
lOO-
l0.i
90’ 0
g z .rr P
90
30
60 90 A72s (m/s)
l20
I
0
150
6p2s
140
l40-
130
UO-
120
g
fLo-
110
ZE .P 2
110 -
i
(E) l8
t
O
lOO-
100 I)
901 0
I
90’ 30
60 90 A12w (m/s)
120
150
0 6 $lZw (ET,
l8
O
Fig. 2c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds averaged over all days.
T. S. VIRDI and
P. J. S. WILLIAMS
l40130
l30-
T‘ &
120
s m-
g
110
5
llo-
2
L
i
x)0-
100
go6
9 Ps
z .P f
140
wo-
130
l30-
110
PO-
5
llo-
15
18
2
100 s
g
12 (LT)
120 1
XIO-
goI 6 g
Fig. 2d. Height profiles of the ter-diurnal amplitude and phase of the meridional winds averaged over all days.
(IF) l5
@3w
l8
(top) and zonal (bottom)
707
Tidal oscillations at high latitude
90
I I -50 A=s &,
-100
I 50
90
I loo
-100
I I -50 Aaow &‘s,
I 50
I 100
Fig. 3a. Height profiles of the mean meridional (left) and zonal (right) winds on summer days.
3 5 P
140-l
'40-j
130 -
l30-
120 nOloo-
‘::
901 0
90
30
60 90 A249 (m/s)
120
150 12
140,
1407
130
130 -
g
mJ-
2
no -
& 5 P T
110
.P
la@249 &)
6
l2
%$24w &I
6
l2
f .._i_-:-
100 120 1!
Qoh
90 0
30
60 A24w @I-$:]
120
150 12
Fig. 3b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonai (bottom) winds on summer days.
7.
708
S. VIRDI and P. J. S. WILLIAMS
MO130 -
g
?20-
g
m-
5 ii
110 -
,F
no-
100-
E
i\-
UYO-
90’ 306090
0
A12s (m/s)
120
140 1
MO-
130 -
l30-
120 -
2 f$
901 0
150
110 loo-90 -0
;'
g
d
p
no-
B$12s(~)
la
O
6 @2w
la
O
MO-
f 30
I I 60 (rn,:;) A12w
120
150
90’
0
(is)
Fig. 3c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on summer days.
Tidal oscillations at high latitude
709
130
loo
I
I
7
6
140 130
5 P
110
100
QOi 6 Fig. 3d. Height profiles of the ter-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on summer days.
710
T. S. VIRDI and P. J. S. WILLIAMS
:;g
{
110
90 -100
I '0
I -50
A&
)
1.
I
I
50
loo
90
I
-100
(m/s>
-50 A,w
1
I
(:/a,
50
1
100
Fig. 4a. Height profiles of the mean meridional (left) and zonal (right) winds on winter/equinox days.
g 5 f
140-l
WI-
130 -
l30-
120 -
g
PO-
110 -
&
tlo-
P
joo-
3
x)0-
goI
0
30
60
90
A24s (m/s)
120
go12
150
140
140
730 -
DO-
g
120 -
S 120
5 P
110 -
.P
2
P 100 -
0
(k)
6
l2
18qJ24w LT)
6
l2
no -.i
x)0-
901
"$249
901 30
60 A24w (m,$
I20
1.50
72
Fig. 4b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on winter/equinox days.
Tidal oscillations at high latitude
711
140 l30
130
3$ I”
no
1
120 ~
100
901 0
30
60 90 AlZs (m/s)
120
750
0
I 30
1 I 60 90 A12w (m/s)
I 120
1 150
90
9o”l----0
90 0
6$12s(~)l6
O
I 6
1
I 12 $12~ (LT)
I 18
0
Fig. 4c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on winter/equinox days.
T. S. VIRDI and P. J. S. WILLIAMS
712
z 5 P
'9
l40-
13
?30-
120
g
m-
110
5
llo-
100 ~
2
i
x)0-
go’
6
901 6
Fig. 4d. Height profiles of the ter-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds on winter/equinox days.
713
Tidal oscillations at high latitude mode reaches a maximum at about 120 km but the phase variation with height does not show any clear pattern and the errors in phase are always large.
6. DISCUSSION OF RESULTS To conclude our study we compared the results of the present work with our previous results using CP1 and with the results obtained at Chatanika, an incoherent-scatter radar located at a similar geographic latitude as EISCAT (JOHNSONet al., 1987). The amplitude and phase of the diurnal and semi-diurnal components were also compared with the theoretical values predicted by FORBES(1982a, b). The main comparisons, based on the analysis of the mean wind profile averaged over the whole data set, are summarised in Fig. 5(a-c). For the mean wind there is extremely good agreement between all three sets of data, except for the meridional component above 115 km where the Chatanika data indicate a northward wind while both sets of EISCAT data agree in showing a southward wind which increases with height. There is also a reassuring agreement in the case of the diurnal component-including the sharp phase change in the zonal component between 110 and 120 km. The new data from EISCAT agree well with the Chatanika result and suggest a re-interpretation of the CP-1 data which were published previously (VIRDI and WILLIAMS, 1989 ; WILLIAMS and VIRDI, 1989). The diurnal amplitude and phase values predicted by FORBES’(1982a) model compare well with the three data sets below about 110 km. Above this altitude the predicted amplitudes in both directions are smaller than the measured values. The predicted meridional phase agrees with observation but the zonal phase above 110 km shows large disagreement. Finally, there is very good agreement for the semidiurnal tidal mode, especially in the zonal plane, but two features deserve comment. The new results indicate a very similar profile of amplitude with height for both meridional and zonal components, with a maximum at 110 km. The previous results at EISCAT and the results from Chatanika, suggest that the amplitude profile in the meridional plane reaches a maximum at about 120 km. The difference may be partly due to seasonal effects combined with the uneven distribution of data throughout the year. It must be added, however, that the results quoted for heights above 120 km in the previous EISCAT study were based on very few data points and so the present results must be regarded as more reliable. This may
also explain the second feature: the phase profile of the meridional semi-diurnal mode in the present study agrees perfectly with the Chatanika results but not with the previous EISCAT results at 120 and 130 km. Once again, it is likely that the present results are more reliable. The semi-diurnal amplitude and phase are predicted reasonably well by FORBES’(1982b) model. Meridional and zonal phase show a similar variation with height although the times of the maxima are approximately 2 h later. The model predicts peaks in both meridional and zonal amplitudes at about 110 km. These peaks are observed in the measurements but with much smaller values. The final result is based on the analysis of the tidal components for each separate day. WILLIAMS and
-75
-50
-25 0 A,” (mls)
25
50
75
-75
-50
-25 0 A2 (m/s)
25
50
75
Fig. 5a. Height profiles of the mean meridional (top) and zonal (bottom) winds from the current study (open circles), from JOHNSONet al. (1957) (squares) and from WILLIAMS and VIRDI(1989) (crosses).
T. S. VIRDI and
714
P. J. S. WILLIAMS
140 -
120
120 -
-z
g
25 E
E
.a 2
.o,
2
100
0
25
0
25
50 A24s (m/s)
75
loo-
loo
140
120
E E
ca f
100
80
I
I 50 A24w (m/s)
1 75
I 100
80 12 324w(-$
6
l2
Fig. 5b. Height profiles of the diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds from the current study (open circles), from JOHNSON et al. (1987) (squares), from WILLIAMS and VIRDI (1989) (crosses) and from FORBES’ (1982a) theoretical model (stars).
Tidal oscillations
715
at high latitude
140
120
g
E .En f
lcm
80
I
0
25
0
25
80
I
I
50 Al 2s (m/s)
/
50 Al 2w (m/s)
I
75
I
75
I
I
I
I
I
100
I
100
00
0
Fig. 5c. Height profiles of the semi-diurnal amplitude and phase of the meridional (top) and zonal (bottom) winds from the current study (open circles), from JOHNSON et al. (1987) (squares), from WILLIAMS and VIRDI (1989) (crosses) and from FORBES’ (1982b) theoretical model (stars).
T. S. VIRDI and
716
P. J. S. WILLIAMS 120 -
Table 2. Correlation coefficient between the day-to-day variations in the measured amplitude of the meridional and zonal components of the different tidal modes at different altitudes
lGQ-
Height (56)
Mean
Diurnal
(%)
(%)
105 110 115 120 125 130
81 57 31 39 39 44
16 44 21 27 20 28
Semi-diurnal (X)
0
Ter-diurnal
12 60 50 39 36 36
%
:
(“/) 58 36 17 14 17 15
.
.
0
20-
VIRDI (1989) had examined
the correlation between the day-to-day variation in the amplitude of the meridional and zonal components of each tidal mode. As a result they tentatively concluded that there was very little correlation in the case of the diurnal variation but a much higher correlation in the case of the semidiurnal component, especially at heights below 120 km. This result is confirmed by the present set of results. Table 2 indicates the correlation coefficient between the amplitudes of the meridional and zonal components of each mode at different heights. There appears to be a relatively high correlation at lower heights for the mean wind and for the semi-diurnal component, but much lower correlation for the diurnal and ter-diurnal modes. In all cases the correlation coefficient tends to drop with height, but it is not yet certain whether this is an intrinsic property of the tidal modes or an effect due to the larger errors in the analysis of tidal modes at greater heights. The correlation in the case of the semi-diurnal mode at 105, 110 and 115 km is illustrated in Fig. 6. At these heights the phase of the two components of the semi-diurnal mode differs by about 3 h. To express this in very simple terms, we can say that the semi-diurnal mode below 120 km usually has a similar amplitude in the meridional and zonal planes and the two components are in quadrature ; in other words, the mode shows a There is no similar conclusion circular polarisation.
0
I
0
/ 20
I
I
loo
1
120
A12 W:'t (m/s?
Fig. 6. The correlation between daily values of the amplitude of the semi-diurnal tide in the meridional and zonal planes as measured by EISCAT CP-2 at 105 (squares), 110 (crosses) and 115 km (open circles).
to be drawn for -the diurnal mode where the day-today variations in the amplitude of the meridional and zonal components seem almost independent and the phase variations are much larger. It is hoped that in the near future EISCAT will introduce alternating codes into the operation of CP2 and this will provide E-region data of an even higher quality. There are also plans to organise even longer continuous runs which will allow a better study of the day-to-day variability of the tidal modes. In the meantime the present study removes some of the discrepancies noted in previous work and provides a reliable outline of the tidal modes in the lower thermosphere at high latitudes. Acknowledgenzentsm-The authors would like to thank the EISCAT Director and his staff for their help in measurement and analysis of the data. EISCAT is a scientific association
supporteh by the research councils of Finland, France, Germany. . , was ., Norwav. __Sweden and the U.K. One of us (TSV) supported by the SERC carried out.
during
REFERENCES BREKKE A., DOUPNIK J. R. and BANKS P. M. FORBES J. M. FORBES J. M. FORBES J. M. and HAGAN A. E. HEDIN A. E. HUUSKONEN A., NYGREN T., JALONEN T., TLIRUNENT. and GLEN J. HUUSKONENA., VIRDI T. S., JONES G. 0. L. and WILLIAMS P. J. S.
I
8
40
1974 1982a 1982b 1982 1987 1986
J. J. J. J. J. J.
geophys. Res. 78, 8235. geophys. Res. 87, 5222. geophys. Res. 87,5241. geophys. Res. 81, 5253. geophys. Ref. 92,4649. atmos. ierr. Phys. 48, 827.
1991
Ann. Geophys. 9,407.
the period
this study
was
Tidal oscillations at high latitude JOHNSON R. M. and VIRLX T. S. JOHNSON R. M., WICKWAR V. B., ROBLER. G. and LUHMANNJ. G. KIRKW~~D S. KUNITAKEM. and SCHLEGEL K. MIKKEUENI. S., JORGENSEN T. S., KELLEYM. C., LARSENM. E., PEIREIRA E. and VICICREY J. SCHUNKR. W. and WALKERJ. C. G. TURUNENT. VIRDIT. S., JONESG. 0. L. and WILLIAMSP. J. S. VIRDI T. S. and WILLIAMSP. J. S. WILLIAMSP. J. S., JONESG. 0. L. and JAINA. R. WILLIAMSP. J. S., JONEXG. 0. L., JONESB., OPGENNCKIRTH H. and HAGGSTR~MI. WILLIAMSP. J. S. and VIRDI T. S.
1991 1987
J. geophys. Res. %, 1099. Ann. Geophys. SA, 383.
1986 1991 1981
J. atmos. terr. Phys. 48,817. Ann. Geophys. 9, 143. J. geophys. Res. 86, 15 13.
1973 1986 1986 1989 1984 1990
Planet. Space Sci. 21, 1875. J. atmos. terr. Phys. 48, 777. Nature 324,354. Adv. Space Res. 9(5), 83. J. atmos. terr. Phys. 86, 521. J. atmos. terr. Phys. 52,439.
1989
J. atmos. terr. Phys. 51, 569.
717