Numerical simulation of the equatorial wind jet in the thermosphere

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 117,, doi: /2011ja017373, 2012 Numerical simulation of the equatorial wind jet in the thermosphere Yasunobu Miyoshi, 1 Hitoshi Fujiwara, 2 Hidekatsu Jin, 3 Hiroyuki Shinagawa, 3 and Huixin Liu 1 Received 13 November 2011; revised 10 January 2012; accepted 20 January 2012; published 8 March [1] We have examined excitation mechanism of the fast jet of the neutral atmosphere along the dip equator in the upper thermosphere. The zonal momentum balance of the neutral atmosphere is estimated using an atmosphere-ionosphere coupled model. The coupled model used in this study is a self-consistent global model of the atmosphere and ionosphere covering the height range from the ground surface to the exobase. It can reproduce the observed equatorial fast jet above 250 km heights. The analysis of the zonal momentum balance reveals that the pressure gradient and ion drag play an important role on the formation of the fast jet near the dip equator. In particular, the fast jet near the equator is closely related with the latitudinal difference of the ion drag force. We also investigate the zonal momentum balance of the longitudinal wave-4 structure of the zonal wind in the fixed local time frame. Furthermore, significant day-to-day variations in the neutral zonal wind and the ion drift near the dip equator are obtained although the solar UV/EUV fluxes and the energy input from the magnetosphere are assumed to be constant during the numerical simulation. This result indicates the importance of the lower atmospheric variability on day-to-day variations in the thermosphere/ionosphere. Citation: Miyoshi, Y., H. Fujiwara, H. Jin, H. Shinagawa, and H. Liu (2012), Numerical simulation of the equatorial wind jet in the thermosphere, J. Geophys. Res., 117,, doi: /2011ja Introduction [2] Behaviors of the equatorial thermosphere are unique and interesting. The latitudinal distributions of the neutral density and wind are geomagnetically controlled. From observations of the orbits of satellites, King-Hele [1973] showed that the thermosphere rotates 10 20% faster than the solid Earth (10 20% superrotation). Using DE-2 (Dynamics Explorer) satellite, Wharton et al. [1984] reported that the diurnal mean zonal wind at the equator in the km height region is about 20 m/s (about 5% superrotation). Measurements by the DE-2 have also revealed behaviors of the diurnal variation of the neutral zonal wind at low latitudes. For example, the zonal wind is westward from about 0500 LT to 1600 LT. The westward wind maximum appears at around noon, while the eastward wind maximum occurs at 2000 LT with a value of 230 m/s. These eastward and westward winds have maxima along the dip equator. In addition, the minimum of the eastward wind around midnight followed by the secondary eastward maximum around 0300 LT occurs. 1 Department of Earth and Planetary Sciences, Kyushu University, Fukuoka, Japan. 2 Faculty of Science and Technology, Seikei University, Musashino, Japan. 3 National Institute of Information and Communication Technology, Koganei, Japan. Copyright 2012 by the American Geophysical Union /12/2011JA [3] Taking advantage of the long-period CHAMP observation, Liu et al. [2006] have revealed detailed behaviors of the diurnal variation of the zonal wind at an altitude of 400 km. The eastward maximum in the evening and the superrotation increase with increasing solar activity. The occurring time of the secondary maximum after midnight has seasonal variations. The secondary maximum during December solstice appears around 0100 LT instead of 0300 LT during equinoxes. [4] Herrero et al. [1985] estimated the zonal momentum balance at the equatorial thermosphere and examined excitation mechanism of the eastward wind during night. The maximum of the eastward wind around 2000 LT is driven by the pressure gradient force. On the other hand, the pressure gradient force during the period from 2100 LT to 2400 LT is negligibly small, and the deceleration of the eastward wind between 2100 and 2400 LT is caused by the ion drag force. Based on these results, Rishbeth [2002] indicated that the F- layer dynamo plays an important role on the generation of the superrotation. However, they did not make use of the measurements of ion drift and the neutral pressure obtained by the DE-2 satellite. Namely, the diurnal variation of ion drift observed at Jicamarca is used, while the neutral pressure is presumed by the neutral density and temperature using the AE-E satellite. Furthermore, effects of the molecular viscosity on the momentum balance are omitted. Therefore, these assumptions provide an indication of the uncertainty in the estimation of the zonal momentum balance. The purpose of this study is to investigate the 1of10

2 excitation mechanism of the zonal wind jet along the dip equator using an ionosphere-atmosphere coupled model. [5] Jin et al. [2011] have recently developed a new whole atmosphere-ionosphere coupled model, which is called Ground-to-topside model of Atmosphere and Ionosphere for Aeronomy (GAIA). This model solves the ionospherethermosphere interaction self-consistently, including the electrodynamics. This means that we can estimate effects of the E-layer and F-layer dynamo processes on the diurnal variation of ion-drift. Using this model, we study the zonal momentum balance of the neutral atmosphere and discuss the mechanism for the fast jet formation. Kondo et al. [2011] examined the momentum balance of the fast thermospheric zonal wind at the dip equator using the NCAR TIEGCM. However, only a few numerical studies concerning the wind jet formation in the thermosphere have been performed. Recent observations [e.g., Lühr et al., 2007] have revealed the wave-4 structure of the zonal wind in a fixed local-time frame due to the upward penetration of the eastward moving diurnal tide with zonal wave number 3 (DE3). Thus, we focus our attention on the latitudinal and longitudinal variability of the zonal momentum balance and its influence on the latitudinal and longitudinal structures of the zonal wind. The descriptions of the GAIA model used in this study and numerical simulation are presented in section 2. Results and discussion are presented in sections 3 and 4, respectively. A summary follows in section Descriptions of the GCM and Numerical Simulation [6] The whole atmosphere-ionosphere coupled model (GAIA) has been recently developed by coupling three models; a whole atmosphere general circulation model (GCM) [Miyoshi and Fujiwara, 2003], an ionosphere model [Shinagawa, 2009], and an electrodynamics model [Jin et al., 2008]. The GAIA model solves the ionosphere- thermosphere interaction self-consistently, including the electrodynamics. The electrodynamics model treats the closure of global ionospheric currents induced both by the neutral wind and by the setup of a polarization electric field under the assumption of equipotential geomagnetic field lines. The present model assumes a tilted dipole for the geomagnetic field configuration. The detailed description of the GAIA model is found in the work of Jin et al. [2011]. [7] For neutral atmospheric part the GAIA model covers the whole range from the ground surface to the exobase and contains a full set of physical processes appropriate for investigating vertical coupling processes in the troposphere, stratosphere, mesosphere and thermosphere as well as the interaction with the ionosphere [Miyoshi and Fujiwara, 2003, 2006, 2008; Fujiwara and Miyoshi, 2006, 2010]. The horizontal resolution of neutral atmospheric part is T42 (maximum horizontal wave number is equal to 42), corresponding to a grid spacing 2.8 latitude by 2.8 longitude, and the vertical resolution is 0.4 scale height above the tropopause. A more detailed description of these physical processes is given by Miyoshi and Fujiwara [2003]. Using this model, we can simulate the excitation of atmospheric waves (such as tides, gravity waves, and planetary waves) in the whole atmosphere including their upward propagation into the upper atmosphere self-consistently without posing any boundaries between the lower and upper regimes. [8] The zonal wind along the equator and the supperrotaion are significant during higher solar activity levels [Liu et al., 2006]. Therefore, in this study, the solar F10.7 cm flux is fixed at W/m 2 /Hz, and geomagnetically quiet condition is assumed. The data are sampled at 10-min intervals from 16 September to 25 September. 3. Results 3.1. Latitudinal Structure of the Zonal Wind [9] The latitudinal structure of the neutral zonal wind obtained by the GAIA is examined first. Longitudinal variations of the neutral zonal wind are also observed at low and middle latitudes [Liu et al., 2009; Lühr et al., 2007]. In order to focus our attention on the latitudinal structure, Figure 1 is averaged over all longitudes. Thus, effects of non-migrating tides on the zonal wind distribution are excluded in Figure 1. The effects of non-migrating tides will be examined in section 3.4. Figure 1a shows the local time-geomagnetic latitude distribution of the zonal wind at 400 km height. A diurnal variation is clearly dominating. The zonal wind at low latitudes is westward between 0500 LT and 1500 LT, and eastward during the remaining hours of the day. The wind magnitude at the dip equator is higher than that around 20 latitudes. This weaker wind occurs at the latitudes where the electron density is the largest. The distribution of the electron density simulated by the GAIA was found in the work of Jin et al. [2011]. Other peaks of the eastward and westward winds are found around latitudes, and these peaks appear earlier than the peaks over the dip equator. The maximum of the eastward wind at the dip equator occurs at LT with a peak value of 220 m/ s. During 1930 LT to 2300 LT, the eastward wind decreases with time, and the secondary peak of the eastward wind appears at LT. On the other hand, the maximum of the westward wind at low latitudes is located around 1000 LT, and the wind reversal from the westward to the eastward at low latitudes occurs around LT. These results are in good agreement with the DE-2 observation [e.g., Herrero et al., 1985, Figure 1], hence the GAIA model can reliably reproduce the observed zonal wind structure. However, the observed maximum of the eastward wind at the dip equator is 170 m/s, which is somewhat smaller than the present result. The average solar flux value (F10.7) during the DE-2 observation is 166, while that in this study is 200. Thus, the difference in the maximum eastward wind is probably due to the difference of the solar activity level. The effect of the solar activity on the zonal wind magnitude will be discussed in section 4.2. [10] Liu et al. [2009] examined the features of the zonal wind in two sets of independent observations. One is from the DE-2 satellite, and the other is from the CHAMP satellite. Besides the similar latitudinal structure obtained by DE-2 and CHAMP, an apparent difference is seen in the occurring time of westward-to-eastward wind reversal and the second wind maximum after midnight. The wind reversal from westward to eastward occurs around LT for CHAMP, while that is around 1600 LT for DE-2. The second wind maximum occurs around 0100 LT for CHAMP, while that appears near 0300 LT for DE-2. Liu 2of10

3 Figure 1. (a) Local time-geomagnetic latitude distribution of the zonal wind at a height of 400 km averaged from 16 September to 25 September. Units are m s 1. Data are averaged over all longitudes in order to exclude the effects of non-migrating tides on the zonal wind distribution. Positive and negative values indicate eastward and westward winds, respectively. (b) Same as Figure 1a except for a height of 200 km. et al. [2009] discussed plausible candidates for these discrepancies, however the reasons for these discrepancies remain unclear. The occurring time of the wind reversal and the second wind maximum simulated in the present study is similar to that from DE-2. The reason why our simulated result resembles the result obtained from DE-2 is also unclear. We need further investigation to clarify the reason for these differences. [11] The zonal wind distribution at an altitude of 200 km is shown in Figure 1b. The westward wind is dominant in the morning hours, and the wind at low latitudes blows eastward at night. The westward-to-eastward wind reversal occurs around 1500 LT. These features are quite similar to those at 400 km height. However, the pronounced maximum of the eastward wind around 1930 LT disappears, and the minimum of the eastward wind occurs at LT. The eastward wind at the dip equator reaches its maximum around LT. The eastward-to-westward turning in the morning occurs near 0800 LT instead of 0600 LT at 400 km. It is interesting to note that the westward wind maximum around 1100 LT is located at 10 latitudes, and the faint minimum of the westward wind is found at the dip equator. There are no observations concerning the equatorial minimum of the westward wind in the morning and the local eastward minimum around LT near the equator. Further observational studies to support this result are required. The diurnal mean zonal wind is 48.4 m/s (10% superrotation) at 400 km height, while that is 27.8 m/s (6% superrotation) at 200 km. Thus the features of the zonal wind distribution at 200 km height are different from those at 400 km height. [12] In order to see the vertical structure of the zonal wind in more detail, the vertical distribution of the zonal wind near the dip equator (averaged over 5 geomagnetic latitude to 5 geomagnetic latitude) is investigated. Figure 2 shows the local time-height distribution of the zonal wind near the dip equator. Above an altitude of 250 km, the zonal wind indicates the eastward wind maximum at LT, the secondary eastward wind maximum around 0300 LT, and the westward wind maximum around 1000 LT. The occurring time of the eastward wind maximum changes with height at the height range from 160 to 250 km. It is noteworthy that the vertical shear of the zonal wind in the km height region is enhanced during the period from 1800 to 2400 LT. The semidiurnal variation is significant in the km height region, and the phase of the semidiurnal variation descends with time, indicating the upward propagating mode of the semidiurnal tide from the lower atmosphere. The terdirunal tide from the lower atmosphere is also evident. This means that the temporal variation of the zonal wind in the km height is determined by superposition of the zonal wind components due to the diurnal, semidiurnal and terdiurnal tides Zonal Momentum Balance [13] In order to investigate the excitation mechanism of the fast zonal wind along the dip equator, we will estimate zonal momentum balance. Using the NCAR TGCM, Killeen and Roble [1986] studied the zonal and meridional momentum balance over the Alaska at 300 km height. We adopt a method similar to Killeen and Roble [1986]. Namely, effects of the advection term, the Coriolis term, the viscous (molecular viscosity) term, the ion drag term and the pressure gradient term of the equation of zonal momentum on the zonal wind acceleration [see Killeen and Roble, 1986, Figure 9]. Figure 3 depicts the diurnal variations of these forcing terms near the dip equator (averaged over 5 geomagnetic latitude to 5 geomagnetic latitude) at 400 km height. [14] The ion drag and pressure gradient terms are the largest for the zonal momentum acceleration. The viscous term ranges from 0.02 to 0.01 m/s 2, which is smaller than Figure 2. Local time-height distribution of the zonal wind near the dip equator (averaged over 5 N geomagnetic latitude to 5 S geomagnetic latitude). Units are m s 1. 3of10

4 Figure 3. (a) Local time distribution of the effect of the advection term on the zonal momentum balance at the dip equator at 400 km height. Units are m s 2. Data are averaged over all the longitudes. (b) Same as Figure 3a except for the effect of Coriolis term. (c) Same as Figure 3a except for the effect of the molecular viscosity. (d) Same as Figure 3a except for the effect of the ion drag force. (e) Same as Figure 3a except for the pressure gradient force. (f) Local time distribution of the zonal wind acceleration due to all the terms. the ion drag by a factor of 2 4. Effects of the advection and Coriolis terms are negligibly small. The total acceleration (Figure 3f) reaches its maximum at 1800 LT with a value of 0.02 m/s 2. The pressure gradient term, ion drag and viscous terms at 1800 LT are 0.085, and m/s 2, respectively. The pressure gradient term drives the zonal wind eastward against the westward ion drag and viscous terms. The pressure gradient term is westward (eastward) during LT (2300 to 0400 LT). This means that the pressure has a weak minimum (maximum) at 2030 LT Figure 4. The broken line shows the local time distribution of the neutral zonal wind near the dip equator (averaged over 5 N geomagnetic latitude to 5 S geomagnetic latitude) at 400 km height averaged from 16 September to 25 September. The solid line is the local time distribution of the eastward ion drift near the dip equator (averaged over 5 N geomagnetic latitude to 5 S geomagnetic latitude) at 400 km. height. Units are m s 1. Figure 5. Same as in Figure 3 except for at 20 N geomagnetic latitude. (2300 LT). The pressure maximum at 2300 LT corresponds to the Midnight Temperature Maximum (MTM), and this will be mentioned in section 3.3. The total acceleration is westward (eastward) during 1900 to 2300 LT (2300 to 0330 LT). This westward acceleration during LT is caused by both the pressure gradient and ion drag terms, while the eastward acceleration during 2300 to 0330 LT is driven by the pressure gradient term against the westward ion drag term. The comparison of the zonal momentum balance between the previous study and this result will be made in section 4.2. [15] The analysis of the momentum balance reveals that the pressure gradient and ion drag terms play an important role on the zonal wind acceleration. In order to examine the behavior of the ion drag term in more detail, the diurnal variation of zonal ion-drift near the dip equator at 400 km height is shown (Figure 4). The diurnal variation of ion drift is quite similar to that of the neutral zonal wind. Namely, the ion drift is westward between 0630 LT and 1530 LT, and eastward during the remaining hours of the day. However, the magnitude of ion drift is smaller than that of the neutral zonal wind. These features of the diurnal variation of ion drift are quite similar to those observed at Jicamarca during the solar maximum periods [Fejer et al., 1985; Fejer, 1991]. [16] The latitudinal difference of the zonal momentum balance is examined here. Figure 5 shows the momentum balance at 20 N geomagnetic latitude where the zonal wind is the weakest. The pressure gradient and ion drag terms are dominant, while the advection, Coriolis and viscous terms are negligibly small. The diurnal variation of the pressure gradient force at 20 N is almost the same with that at the dip equator. On the other hand, the magnitude of the ion drag term at 20 N is larger than that at the dip equator. This larger westward ion drag force at 20 N attenuates the magnitudes of the total eastward acceleration and the zonal wind at 20 N. For example, the maximum of the total acceleration around 1800 LT at 20 N is m/s 2, which is 75% of that at the dip equator. [17] In order to examine the latitudinal structure of the zonal momentum balance in more detail, Figure 6 shows the 4of10

5 Figure 6. Latitudinal distribution of the pressure gradient (dotted line), ion drag (broken line), viscous (chain line) and total acceleration (solid line) terms averaged from 1700 to 1800 LT. latitudinal distributions of the pressure gradient, ion drag, viscous and total acceleration terms during 1700 to 1800 LT when the total eastward acceleration reaches its maximum. The latitudinal difference of the pressure gradient term is negligibly small, while the ion drag term is the smallest at the dip equator. It is clearly seen that the smallest ion drag force at the dip equator induces the maximum of the total acceleration at the dip equator. It is noteworthy that the ion drag force has its maxima around 20 25, and the ion drag force at is smaller than that at 20. [18] The zonal wind balance at various heights is investigated. Figures 7a 7c show the balance near the dip equator at 300 km height. The effects of the advection and Coriolis terms are small, so that these terms are omitted in Figure 7. The zonal wind acceleration during day is determined by the balance between the pressure gradient and ion drag terms. Figure 7. (a) The solid (dotted) line shows local time distribution of the effect of the ion drag force (molecular viscosity) on the zonal momentum balance at the dip equator at 300 km height. Units are m s 2. Data are averaged over longitudes. (b) Same as Figure 7a except for the effect of the pressure gradient. (c) Local time distribution of the zonal wind acceleration due to all the terms. (d) Same as Figure 7a except for 200 km height. (e) Same as Figure 7b except for 200 km height. (f) Same as Figure 7c except for 200 km height. Figure 8. Local time-geomagnetic latitude distribution of the temperature at a height of 400 km. Units are K. Data are averaged over all longitudes. During the period from 1800 to 2200 LT, the viscous term is larger than the ion drag term. This larger viscous term is closely related with the fact that the vertical shear of the zonal wind around 300 km is enhanced during LT (Figure 2). During 1400 to 2100 LT, the pressure gradient drives the wind eastward against the westward force due to the ion drag and viscous terms. The pressure gradient term is westward (eastward) during LT (2300 to 0400 LT), and this feature is similar to that at 400 km. [19] The momentum balance around the dip equator at 200 km height is shown in Figures 7d 7f. The pressure gradient term is significant, and the magnitude of the pressure gradient term is larger than that of the ion drag and viscous terms by a factor of 3 4. The temporal variation of the total acceleration is mainly determined by the temporal variation of the pressure gradient term. The ion drag force is enhanced during day, while the viscous term is enhanced during the period from 2000 to 2300 LT. The pressure gradient term and the total acceleration have clear semidiurnal variations. As mentioned above, the semidiurnal tide from the lower atmosphere is significant. The minima of the pressure gradient force at 0600 LT and 1800 LT are caused by the semidiurnal tide. Therefore, the minimum of the eastward wind around LT are generated by the upward propagating semidiurnal tide. Because the amplitude of the semidiurnal tide poleward of 40 latitude is negligible, the minima of the pressure gradient force poleward of 40 latitude at 0600 LT and 1800 LT disappear. This means that the pressure gradient force poleward of 40 latitude reaches its maximum around LT and produces the maximum of the eastward wind around 1800 LT Role of the Semidiurnal and Terdiurnal Tides [20] We consider the second maximum of the eastward wind after midnight. Figure 8 presents the local timegeomagnetic latitude distribution of the temperature at a height of 400 km. The temperature maximum at low latitudes is located at LT, while the temperature minimum appears around 0400 LT. The secondary maximum near the equator is evident at LT, and corresponds to the MTM. Using the NCAR TGCM, Fesen [1996] showed that the MTM was generated by the upward propagation of the migrating semidiurnal tide from the lower atmosphere. Miyoshi et al. [2009] has recently showed that not only the migrating semidiurnal tide but also the migrating terdiurnal tide played an important role on the generation of the MTM. As mentioned in section 3.2, the reversal of the pressure gradient force from westward to 5of10

6 Figure 9. (a) Local time distribution of the migrating semidiurnal zonal wind component at the dip equator at 400 km height. Units are m s 1. (b) Same as Figure 9a except for the migrating terdiurnal tide. (c) Same as Figure 9a except for superposition of the semidiurnal and terdiurnal tides. eastward occurs at 2300 LT. These results indicate that the behaviors of the migrating semidiurnal and terdiurnal tides are important for investigating the temporal variations of the pressure gradient force and the zonal wind. Thus, the behaviors of the migrating semidiurnal and terdiurnal tides are investigated. Figures 9a and 9b show the semidiurnal and terdiurnal variations of the zonal wind component at the dip equator at 400 km height, respectively. The amplitude of the semidiurnal (terdiurnal) tide is 43 m/s (39 m/s). The zonal Figure 10. (a) Longitude-geographic latitude distribution of the zonal wind at a height of 400 km at 1000 LT. Units are m s 1. (b) Same as Figure 10a except for 1900 LT. Figure 11. (a) The solid (dotted) line shows longitude distribution of the ion drag force (molecular viscosity) on the zonal momentum at the dip equator at 400 km height averaged over the period from 0400 LT to 1000LT. Units are m s 2. (b) Same as Figure 11a except for the effect of the pressure gradient force. (c) Longitude distribution of the zonal wind acceleration due to all the terms averaged over 0400 LT to 1000 LT. (d) Longitudinal distribution of the zonal wind at 1000 LT. Units are m s 1. wind due to the semidiurnal tide maximizes at 0530 and 1730 LTs, while that of the terdiurnal tide maximizes at 0300, 1100 and 1900 LTs. Figure 9c depicts the zonal wind variation due to the semidiurnal and terdiurnal tides. The eastward wind peaks are located around 0330 and 1830 LTs, and the westward wind is enhanced at 2300 LT. This result indicates that the second maximum of the eastward wind around LT is caused by superposition of the semidiurnal and terdiurnal tides Longitudinal Variation of the Zonal Wind [21] In this section, we investigate the longitudinal structure of the zonal wind at 400 km height. Figure 10a shows the longitude-geographic latitude section of the zonal wind at 1000 LT, when the westward wind at low latitudes reaches its maximum. The strong westward jet is aligned with the dip equator. This result clearly demonstrates that the behavior of the zonal wind is magnetically controlled. The longitudinal variation of the zonal wind is also evident. For example, the westward wind near the dip equator has peaks at 125, 220, and 320 E. Another weak peak of the westward wind is located around 20 E, indicating a wave-4 longitudinal structure (Figure 11d). Further analysis shows that this wave-4 structure is mainly generated by the DE3. The amplitude of the DE3 near the equator at 400 km height is 6 m/s, which is consistent with the previous result [Oberheide et al., 2009]. It is noteworthy that the longitudinal variation of the zonal wind is evident not only at low latitudes but also at middle latitudes. Using the GAIA, Miyoshi et al. [2012] has recently investigated the main source of the wave-4 longitudinal structure at middle latitudes. They showed that the eastward moving semidiurnal 6of10

7 Figure 12. (a) Diurnal variations of the neutral zonal wind at the dip equator and 0 E of 400 km height during September. Thin lines indicate diurnal variations in individual days. Units are m s 1. (b) Same as Figure 12a except for the eastward ion drift. tide with zonal wave number 2 is essential for the generation of the wave-4 longitudinal structure at middle latitudes. [22] In order to study the excitation mechanism of the wave-4 structure of the zonal wind at the dip equator, the longitudinal variation of the zonal momentum balance at 400 km is examined. Figures 11a and 11b show the longitudinal distributions of the ion drag, viscous and pressure gradient terms averaged over 0400 to 1000 LT. The average from 0400 to 1000 LT is chosen because the westward acceleration of the zonal wind is significant during this period. The magnitude of the viscosity term is smaller than that of the ion drag term by a factor of 4 5. The longitudinal variation of the total acceleration (Figure 11c) is mainly determined by the balance between the westward acceleration of the pressure gradient term and the eastward acceleration of the ion drag term. Both the pressure gradient and ion drag terms have the wave-4 structure. However, the longitudinal variability in the ion drag term is smaller than that in the pressure gradient term. For example, the ion drag term ranges from to m/s 2, while the pressure gradient term ranges from to m/s 2. This means that the shape of the longitudinal variability in the total acceleration is similar to that in the pressure gradient term. The westward acceleration and the westward wind are strong at the longitudes where the westward pressure gradient term has peaks. Therefore, the wave-4 structure of the pressure gradient term, which is mainly generated by the upward penetration of the DE3, produces the wave-4 structure of the zonal wind around 1000 LT. [23] Figure 10b is the longitude-geographic latitude section of the zonal wind at 1900 LT, when the eastward wind at low latitudes reaches its maximum. The strong eastward jet is aligned with the dip equator. The longitudinal variation of the zonal wind is also significant, and the peak eastward winds near the dip equator occurs around E. Other peaks of the eastward wind are located at 20 E, 90 E and 270 E. The features of the longitudinal variation of the momentum balance in the evening are quite similar to those in the morning (not shown). Namely, the longitudinal variation of the total acceleration (Figure 11c) is mainly determined by the balance between the eastward acceleration of the pressure gradient term and the westward acceleration of the ion drag term. The eastward acceleration and the eastward wind are strong at the longitudes where the eastward pressure gradient term has peaks Day-to-Day Variation of the Neutral Wind and Ion Drift [24] We examine day-to-day variations of the zonal neutral wind and ion drift. Figure 12a shows the diurnal variation of the neutral zonal wind at the dip equator and 0 E at an altitude of 400 km. Significant day-to-day variations are visible. These day-to-day variations obtained in this study are generated by the lower atmospheric variability because the solar UV/EUV fluxes and the energy input from the magnetosphere are assumed to be constant during the numerical simulation. Furthermore, fluctuations with periods of 1 3 h are also evident. Using GCM simulation, Miyoshi and Fujiwara [2008] showed that gravity waves with a shorter period could penetrate into higher altitudes to overcome dissipation processes in the thermosphere, such as the molecular viscosity and ion drag force. Thus, shortperiod fluctuations in Figure 12a is likely generated by the penetration of gravity waves from the lower atmosphere. [25] Figure 12b depicts the diurnal variation of the zonal ion drift at the same grid point. Day-to-day variations of the ion drift are also significant. The ion drift at 1200 LT ranges from 40 to 100 m/s, while that at 2000 LT ranges from 150 to 195 m/s. These day-to-day variations of the ion drift are generated by the variations of the neutral wind through the E-layer and F-layer dynamo processes. Thus, these results indicate the importance of the lower atmospheric variability on the variations in the ion drift of the F region height. 4. Discussions 4.1. Relation to the Lower Atmospheric Variability [26] In this section we discuss day-to-day variations of the neutral zonal wind and zonal ion drift as shown in section 3.5. We demonstrated that the zonal wind component due to the semidiurnal and terdiurnal tides play an important role on the generation of the eastward wind minimum before midnight and the secondary eastward maximum after midnight. Using GCM simulations, day-to-day variations of the MTM is closely related with the variability of the upward propagating tides [e.g., Akmaev et al., 2010; Fujiwara and Miyoshi, 2010]. Miyoshi et al. [2011] suggested that day-to-day variations of the terdiurnal tide of the lower atmospheric origin produce the variations in the MTM. The amplitudes of the semidiurnal and terdiurnal tides are closely related to the distribution of water vapor, cloud and ozone in the lower atmosphere. For example, Hagan and Forbes [2002] indicated that the seasonal variation of the convective activity in the tropics affects seasonal variation of the tidal amplitude in the mesosphere and lower thermosphere using a global Scale wave Model (GSWM). 7of10

8 Figure 13. (a) Local time-geomagnetic latitude distribution of the zonal wind at a height of 400 km under solar moderate (F10.7 = 135) condition. Units are m s 1. Data are averaged over all longitudes. (b) Same as Figure 13a except for solar minimum (F10.7 = 70) condition. Using a whole atmosphere GCM, Miyoshi and Fujiwara [2003] showed that day-to-day variations in tropical convective activity produces day-to-day variations in the intensity of diurnal tides which are propagating into the upper atmosphere. Thus, these results imply that day-to-day variations of the semidiurnal and terdiurnal tides associated with the lower atmospheric variability can cause day-to-day variations of the zonal wind in the upper thermosphere. [27] Takahashi et al. [2005] studied the variation of the minimum virtual height h F with 3 4 days period, and suggested that the 3 4 days period must be related to the Ultra Fast Kelvin (UFK) wave in the mesosphere. Pancheva et al. [2006] suggested that the 2-day planetary wave from the lower atmosphere produces the 2-day variations in the ionospheric electric currents and F-layer electron densities. The UFK and the 2-day planetary wave cannot penetrate into the upper thermosphere, so that these oscillations in the ionsopheric parameters are probably due to the modulation of the upward E B drift through the E region dynamo process. However, excitation mechanism of the day-to-day variations in the ionospheric parameters induced by the atmospheric waves is not well examined. In order to identify the relation between the mesospheric wind fluctuations and the oscillations in the ionospheric parameters, further investigation concerning wind fluctuations in the E region heights and the day-to-day variability of the upward E B drift through the E region dynamo process are required. In the next step, using an atmosphere-ionosphere coupled model, we will have to estimate the day-to-day variability of the upward E B drift caused by the atmospheric waves Comparison With Previous Studies [28] In this section, we make a comparison of the zonal momentum balance between the present result and the previous studies. Using the DE-2 satellite, Herrero et al. [1985] studied the zonal momentum balance at the equatorial thermosphere. They showed that the enhancement of the eastward wind around 2000 LT and after midnight is caused by the pressure gradient force. This result is consistent with the present result. During LT, the pressure gradient force is weak eastward, and the attenuation of the eastward wind is driven by the ion drag force. In the present result, the ion drag force during LT is westward, and contributes to the attenuation of the eastward wind. However, the pressure gradient force during LT in the present study is also westward, and plays an important role on the zonal momentum balance. This result is not consistent with their result. However, they estimated the diurnal variation of the pressure using the AE-E satellite data and the MSIS empirical model. This causes the uncertainty in the estimation of the pressure gradient force. In the present study, the pressure gradient force is evaluated using an atmosphere-ionosphere coupled model (GAIA). Thus, we can conclude that our estimation of the pressure gradient force is reliable. [29] Kondo et al. [2011] investigated the zonal momentum balance at the dip equator using the NCAR TIEGCM. They showed that the pressure gradient and ion drag force plays an important role on the zonal momentum balance in the dip equator. The pressure gradient term drives the zonal wind eastward against the westward ion drag force during 1700 to 2100 LT. This feature is consistent with the present result. However, the viscous term in their simulation is eastward during 1700 to 2000 LT, while that in the present study is westward. In our simulation, the vertical shear of the zonal wind near the dip equator, which is mainly generated by the upward propagating modes of the semidiurnal and terdiurnal tides, is enhanced in the evening, and this large vertical shear produces the westward acceleration of the viscous term above a height of 250 km. On the other hand, in the work of Kondo et al. [2011], the strong eastward winds at 2000 LT exist around 250 km height, and the vertical shear of the zonal wind at 2000 LT disappears. The lower boundary of the NCAR TIEGCM is set at 97 km height, and effects of the upward propagating terdiurnal tide are not taken into account in their simulation. Therefore, these results indicate that the upward propagating tides affect the vertical shear of the zonal wind in the km height region. We suggest that the difference of the diurnal variation of viscous term at 400 km height between the present result and their result is caused by the upward propagating tides. [30] Using the CHAMP satellite, Liu et al. [2006] examined effects of the solar activity on the diurnal variation of the zonal wind at an altitude of 400 km. They indicated that the magnitude of the eastward wind in the evening and the superrotation increases with increasing solar activity, while the magnitude of the westward wind in the morning decreases with increasing solar activity. Using the GAIA, the zonal wind distributions at various solar activities are studied. Figures 13a and 13b show the local time-geomagnetic latitude distribution of the zonal wind at 400 km height during solar moderate (F10.7 = 135) and solar minimum (F10.7 = 70) conditions, respectively. The magnitude of the eastward wind around 2000 LT is 180 m/s (130 m/s) at solar moderate (minimum) condition. The diurnal mean zonal winds at the dip equator during solar moderate and minimum conditions are 32 m/s and 15 m/s, respectively. This means 8of10

9 that the superrotaion is enhanced during solar maximum condition. The occurring time of westward-to-eastward wind reversal is insensitive to solar activity. These results are in good agreement with the CHAMP observation. [31] The behavior of the secondary maximum of the eastward wind after midnight during solar minimum condition is different from that during solar maximum condition. Namely, the occurring time of the secondary eastward maximum during solar minimum condition ( LT) is about 30 min later than that during solar maximum ( LT). A faint maximum of the eastward wind around LT is discernible during solar minimum condition. These differences are explained by the fact that the phase structures of the semidiurnal and terdiurnal tides during solar minimum condition are different from those during solar maximum condition. Mayr et al. [1979] showed that the ion drag force in the thermosphere modulates the phase structure of the semidiurnal tide. The present result also indicates that the solar activity level affects the phase structures of the semidiurnal and terdiurnal tides in the thermosphere. Moreover, in the present result, the magnitude of the westward wind in the morning also increases with increasing solar activity. This result is inconsistent with the result obtained by the CHAMP satellite. In order to clarify the effects of the solar activity on the zonal wind distribution, we have to investigate the zonal wind balance during solar moderate and minimum conditions. This will be a subject of the future study. 5. Summary [32] Using the GAIA model, the excitation mechanism of the equatorial fast jet at the dip equator during September equinox condition has been investigated. The zonal wind distributions obtained in the present study are in good agreement with the DE-2 and CHAMP observations. The analysis of the zonal momentum in the km height region indicates that the pressure gradient force and the ion drag force play important roles on the zonal wind balance. The pressure gradient force in the evening (morning) drives the zonal wind eastward (westward) against the westward (eastward) ion drag and viscous forces. The fastest eastward wind at the dip equator is generated by the fact that the westward ion drag force has its minima at the dip equator. On the other hand, the second maximum of the eastward wind after midnight is closely related with the semidiurnal and terdiurnal tides from the lower atmosphere. The magnitude and occurring time of the secondary maximum after midnight may be influenced by the seasonal variations of the semidiurnal and terdiurnal tides, because the semidiurnal and terdiurnal tides from the lower atmosphere have significant seasonal variations [Miyoshi et al., 2009]. To investigate the seasonal variations of the semidiurnal and terdiurnal tides and their impacts on the zonal momentum balance is a subject of the future study. [33] Significant day-to-day variations in the neutral zonal wind and the eastward ion drift are obtained although the solar UV/EUV fluxes and the energy input from the magnetosphere are assumed to be constant during the numerical simulation. Thus, day-to-day variations in the thermosphere/ ionosphere are closely related with the lower atmospheric variability. Day-to-day variations in the ion drift are considered to be generated by the variations in the neutral wind through the E-layer and F-layer dynamo processes. In the next step, using the GAIA, detailed analysis concerning the effects of the upward propagating atmospheric waves on day-to-day variations in the ion drift through the dynamo processes is desirable. [34] Acknowledgments. This work is supported by MEXT Grantin-Aid for Scientific Research ( ). 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