Variations in lower thermosphere dynamics at midlatitudes during intense geomagnetic storms

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 109,, doi:10.1029/2003ja010244, 2004 Variations in lower thermosphere dynamics at midlatitudes during intense geomagnetic storms Larisa P. Goncharenko, Joseph E. Salah, John C. Foster, and Chaosong Huang Haystack Observatory, Massachusetts Institute of Technology, Westford, Massachusetts, USA Received 19 September 2003; revised 23 December 2003; accepted 5 February 2004; published 9 April 2004. [1] High-resolution observations of the dynamics of the ionospheric E region (100 150 km) at midlatitudes have been made with the incoherent scatter radar at Millstone Hill (42.6 N, 71.5 W) during several intense geomagnetic storms (K p 8 9). These included the storms of 25 September 1998, 15 July 2000, and 31 March 2001. The observations reveal electric fields of up to 100 mv/m during these events and indicate E region zonal plasma drifts of 300 1000 m/s, westward in the evening sector and eastward in the morning sector, with smaller meridional drifts. The tidal pattern of neutral winds in the lower thermosphere is heavily disrupted. The zonal component of the neutral wind follows convection-driven ion flow, while the meridional component shows a more complex response to geomagnetic forcing. INDEX TERMS: 3369 Meteorology and Atmospheric Dynamics: Thermospheric dynamics (0358); 2788 Magnetospheric Physics: Storms and substorms; 3384 Meteorology and Atmospheric Dynamics: Waves and tides; 2427 Ionosphere: Ionosphere/atmosphere interactions (0335); KEYWORDS: geomagnetic storms, lower thermosphere, neutral wind Citation: Goncharenko, L. P., J. E. Salah, J. C. Foster, and C. Huang (2004), Variations in lower thermosphere dynamics at midlatitudes during intense geomagnetic storms, J. Geophys. Res., 109,, doi:10.1029/2003ja010244. 1. Introduction [2] Assessing the extent of the geomagnetic influence on the Earth s thermosphere and ionosphere has been a challenge for the research community for a number of years. Although effects of geomagnetic storms are relatively well studied in the ionospheric F region and upper thermosphere, the response at lower altitudes is not well understood, primarily for two major reasons. First, the response at mesospheric and lower thermospheric altitudes is much weaker than at F region heights, is highly dependent on geographic location and local time, and can often be masked by other processes that usually dominate in these regions. Second, the dearth of observational data, particularly at lower thermospheric altitudes (100 150 km), combined with high variability of the wind field, make it difficult to trace the source responsible for a particular change in the dynamics. [3] In spite of these difficulties, the research community has accumulated to date a substantial dataset during geomagnetic storms. Available observations include continuous sets of wind measurements by meteor and medium frequency radars at 80 110 km [e.g., Hook, 1970; Manson and Meek, 1986; Price et al., 1991; Singer et al., 1994; Ma et al., 2001; Fahrutdinova et al., 2001]. At altitudes above 110 km, significant evidence of geomagnetic forcing was found in observations made by incoherent scatter radars, especially at high latitudes [e.g., Johnson et al., 1987; Kunitake and Schlegel, 1991; Nozawa and Brekke, 1995; Salah et al., 1996]. Large changes in the Copyright 2004 by the American Geophysical Union. 0148-0227/04/2003JA010244 wind pattern due to geomagnetic disturbances were found in Wind Imaging Interferometer (WINDII) data in both statistical [Emmert et al., 2001, 2002] and case studies [Zhang and Shepherd, 2000, 2002]. [4] In the upper mesosphere/lower thermosphere (80 105 km), the evidence for any storm-related effects is highly contradictory, perhaps because of different geographical locations and local times used in the various studies, as well as sampling errors arising from a variety of mechanisms. An early study of the neutral wind by Hook [1970] at College, Alaska (65 N, 147 W) using meteor wind radar found evidence of enhanced southward wind flow at altitudes of 75 110 km at nighttime. A recent study of mesospheric winds at midlatitude locations Yamagawa (31.2 N, 130.6 E) and Wakkanai (45.4 N, 141.7 E) by Ma et al. [2001] reported enhancement in the eastward wind several days after the storm onset, with the magnitude of the enhancement depending on the season. As tidal motion is the most prominent feature of mesospheric and lower thermospheric dynamics, different researchers characterized effects of geomagnetic disturbance in terms of changes in tidal components. Fahrutdinova et al. [2001] report a reduction in the mean zonal wind and appearance of the northward meridional component, i.e., rotation of the mean wind vector, together with an increase in the amplitude of the diurnal and semidiurnal tides. In contrast, Singer et al. [1994] found no reaction in the meridional component and a weak reduction of the semidiurnal tidal amplitude due to geomagnetic storms. [5] Several studies have suggested that an increase in geomagnetic forcing leads to changes in tidal modes observed at lower thermosphere altitudes, and theoretical 1of13

simulations imply that such changes could be measurable at high midlatitudes [Müller-Wodarg et al., 2001]. From multiday measurements over Millstone Hill, Wand [1983] found that during disturbed intervals the semidiurnal tidal amplitude decreases by 20 50% at heights of 105 115 km and suggested in situ generation of semidiurnal oscillations caused by auroral heating, while Salah et al. [1996] report increase in the amplitude of the semidiurnal tide around 110 km. Kunitake and Schlegel [1991] found that in EISCAT observations the amplitude of the diurnal tide at 117/120 km increases with rises in geomagnetic activity, and the correlation is better for the zonal component of the neutral wind than for the meridional component. However, Nozawa and Brekke [1995] showed that though the diurnal component increases during disturbed periods, it is likely to be a result of enhanced diurnal oscillations in the electric field (i.e., enhanced two-cell ion convection), and not tidal forcing. [6] Recent statistical studies using daytime winds measured by the WINDII instrument show that disturbance winds in the 100 260 km altitude region, defined as difference between winds during geomagnetically active and geomagnetically quiet periods, in general follow the two-cell ion convection pattern [Emmert et al., 2002] and are similar to the upper thermospheric winds [Emmert et al., 2001]. The strongest disturbance winds are observed at high midlatitudes, with eastward wind in the morning hours and westward in the afternoon, while the meridional disturbance wind is equatorward. The Richmond et al. [2003] study of high-latitude lower thermospheric winds using WINDII data in the summer season shows that the wind pattern is consistently similar to the ion convection pattern above 125 km, and the influences of ion convection can be detected at altitudes as low as 105 km. [7] The objective of this study is to assess the penetration extent of the storm effects in latitude and altitude. In this work, we present observations made by the Millstone Hill incoherent scatter radar (43 N, 72 W, magnetic latitude 54 ) during three intense geomagnetic storms: 25 September 1998, 15 July 2002, and 31 March 2001. Though incoherent scatter radars are unable to make continuous observations and collect data mostly on campaign basis, they provide simultaneous measurements of several parameters, which help to elucidate the details of particular mechanisms at work during storm events. In general, the effects of individual geomagnetic storms are not detected at midlatitudes in the E region unless the disturbances are intense and the K p index exceeds values of about 6 [Salah et al., 1996; Salah and Goncharenko, 2001]. The present study concentrates on three recent geomagnetic storms with the maximum K p index exceeding 8, and the intensities of these storms are therefore expected to introduce major perturbations to the ionospheric E region at midlatitudes. Although we report measurements of plasma drifts in the F region that are used to derive the electric fields in the ionosphere, we emphasize here the study of the variations of the plasma drifts and neutral winds at altitudes from 100 to 150 km. We will also address the coupling between the ion and neutral components. We concentrate on E region data only; other results obtained during these three storms are not discussed in this paper and the interested reader can find them elsewhere [Pavlov and Foster, 2001; Foster et al., 2002; Mishin et al., 2002; Huang et al., 2003]. 2. Solar and Geophysical Conditions and Experiment Description [8] Observations of the ionospheric E region were made by the Millstone Hill incoherent radar during three intense geomagnetic storms of the current solar cycle: 25 September 1998, 15 July 2000, and 31 March 2001. Although all three events are very intense, there is a significant difference between them as they occur during different levels of solar activity, different seasons, and during different local times. Figure 1 presents the variations of geomagnetic indices K p (upper panels) and D st (lower panels) for these three events. The range of coverage by the radar for each of the events is indicated in Figure 1 with the dark bars at the bottom parts of the panels. It is important to note that due to the heavy particle precipitation at midlatitudes during the storms, observations at Millstone Hill in the E region during the nighttime were possible and allowed a near-continuous description of the time variations of the ionospheric E region during the storm. This is in contrast to measurements made in this region during geomagnetically quiet times, which are limited only to daytime periods when the signalto-noise ratio from the backscattered signal is sufficiently strong. [9] For the first storm, 25 September 1998, the solar activity was moderate, with F 10.7 = 138 on 25 September, and 81-day average F 10.7 = 130. A severe geomagnetic disturbance began close to 0000 UT on 25 September (see Figure 1a) and was indicated by an increase in D st to 4 nt at 0100 UT. The main phase of the storm was represented by a rapid decrease in D st to 207 nt at 0800 UT. The D st index stayed at the 200 nt level until 1100 UT, and during the recovery phase D st increased to 66 nt by 2000 UT. The K p index reached a maximum value of 8+ at 0600 0900 UT, decreased to 7 and 6+ at 0900 1500 UT, and dropped to 3 at 1800 2100 UT. [10] The Millstone Hill radar was operated continuously from 1030 UT on 21 September 1998 to 0330 UT on 27 September 1998, which allows us to compare daytime storm-induced effects with the quiet conditions preceding the storm. The first incoherent scatter radar echoes from the E region on 25 September were obtained at 1200 UT, i.e., after the peak of the storm. The observational mode for the radar included a measurement of the full ion velocity vector with a 36 min. cycle time, as well as electron density and plasma temperatures with an altitude resolution of 3 km in the E region and 20 30 km in the F region, depending on the pulse length. [11] The second storm, 15 16 July 2000 (the so-called Bastille day event), occurred during maximum solar activity, with F 10.7 = 213 on 15 July and F 10.7 = 219 on 16 July, and a 81-day average F 10.7 = 185. The history of the magnetic activity for this storm is presented in Figure 1b. After a sudden storm commencement with D st increasing to 7 nt at 1600 UT on 15 July 2000, D st decreased to 301 at 0100 UT, and the recovery phase began near 0300 UT on 16 July. The radar collected data from 1304 UT on 15 July 2000 to 2004 UT on 16 July 2000. This experiment included wide-coverage azimuth and elevation scans, which 2of13

Figure 1. K p and D st indices during (a) 21 27 September 1998, (b) 15 16 July 2000, and (c) 31 March to 3 April 2001. The horizontal bar at the bottom of the K p plots indicates time periods of E region coverage. provided data over 20 of latitude, as well as local E region and F region measurements. In order to accommodate other scientific objectives, the cycle time for this experiment increased to 54 min. However, modes providing local E region and F region coverage used in this study were spaced out within 15 min in each cycle, thus allowing us to reduce errors introduced by the assumption of temporal homogeneity when deriving vector drifts. [12] Finally, the third event considered for this study occurred at solar maximum with F 10.7 = 245 on 31 March 2001 and 81-day average F 10.7 = 170. The sudden storm commencement was observed at 0400 UT with D st of 30 nt (see Figure 1c), and the D st index dropped to 358 nt at 0900 UT during the main phase of the storm. After the initial recovery phase, D st experienced a second decrease to 283 nt at 2000 UT, which was followed by a slow, long recovery over several days. This double-peak nature of the event is also reflected in the K p index, with the first maximum reaching 9- from 0300 UT to 0900 UT and the second weaker maximum of 8+ at 1800 2100 UT. The Millstone Hill radar operated from 2028 UT on 30 March 2001 to 1941 UT on 4 April 2001, thus allowing us to study both the disturbed and recovery periods. The radar operational mode for this experiment was the same as for 15 16 July 2000, i.e., combined wide-coverage scans and local E and F region data, resulting in a 54-min cycle time. 3. Data Analysis [13] To obtain an ion velocity vector, we used line-ofsight ion velocity measurements from at least three fixed antenna positions, assuming spatial and temporal homogeneity. As the elevation angle for all radar beam positions was 45, the maximum spatial separation covered by the beams is 250 km at E region heights and 400 km at F region heights. These relatively large separations in space and time filter out variations on small spatial and temporal scale, and therefore our results capture only the most prominent features of the disturbed thermosphere. To derive neutral wind velocity in the E region, we follow the commonly used approach, previously described by Salah et al. [1991] and Wand and Evans [1981]. In this analysis we use the updated mass spectrometer/incoherent scatter model of the neutral atmosphere NRLMSISE-00 [Hedin, 1987; Picone et al., 2002] and collision cross sections given by Banks and Kockarts [1973]. The geomagnetic field strength is determined from the International Geomagnetic Reference Field (IGRF-95) model. The electric field is derived from ion velocity measurements in the F region. Since the ion-neutral collision frequency becomes much smaller than ion gyrofrequency at the altitudes above 160 km, the electric field is calculated assuming that ion motion across the magnetic field line is caused by E B drift. Although the electric fields during disturbed periods can change in both direction and magnitude on the timescale of the order of minutes, we expect that such rapid variations do not contribute significantly to the ion-neutral momentum transfer. We will discuss this in greater detail in a later section of this paper. The uncertainty of neutral wind estimates is estimated to be 20 25 m/s [Salah et al., 1991], and analysis of error sources in computing wind from incoherent scatter radar data is provided by Johnson et al. [1987] and Rino et al. [1977]. During geomagnetic disturbances the error in the wind estimates can increase to 200 m/s. Uncertainties in neutral composition (and hence in ion-neutral collision frequency) create larger errors 3of13

Figure 2. The northward (E y ) and the eastward (E x ) components of the electric field measured above Millstone Hill for (a) 24 25 September 1998, (b) 15 16 July 2000, and (c) 31 March to 1 April 2001. in neutral wind determination at higher altitudes, where the ratio of collision frequency to gyrofrequency becomes too small. For this reason we limit the upper altitude of wind calculation to 130 km. 4. Results 4.1. Electric Field Observations [14] Electric field measurements are an essential part of the neutral wind calculation in the lower thermosphere and are a convenient indicator of the general level of disturbance in the ionosphere due to the storm. Figure 2 shows two components of the electric field measured above Millstone Hill for the three geomagnetic storms. The meridional (northward positive) component of the electric field is shown in dark circles, and the zonal (eastward positive) component of the electric field is shown in open circles. For each storm period, we show the electric field components for two consecutive days. Figure 2a presents electric field variations on 24 25 September 1998, Figure 2b on 15 16 July 2000, and Figure 2c on 31 March to 1 April 2001. [15] Since ions during storms are driven primarily by electric fields, these fields can be considered a direct presentation of ion convection in the E B direction, i.e., a northward electric field represents the westward component of the ion drift vector. Figure 2a shows that during quiet time on 24 September 1998 (1000 2000 UT), both components of the electric field do not exceed 3 mv/m. As the storm develops, increased ion convection changes from westward motion in the afternoon and premidnight sector (2100 0300 UT or 1600 2200 LT) to predominantly eastward motion in the morning sector (0800 1300 UT or 0300 0800 LT). The electric field magnitude increases to 20 30 mv/m during the most disturbed time but rapidly decreases after 1300 UT and soon after 1500 UT drops to the level observed during quiet time. This increase in the electric field represents the expansion and equatorward shift of the convection pattern, which typically can be observed at subauroral latitudes and was reported in numerous studies. The simultaneous colocated observations by the DMSP satellite (Figure 3a) confirm such expansion and show horizontal ion velocities of 600 m/s above the Millstone Hill location, consistent with the electric field observations. [16] In the case of 15 16 July 2000 storm (Figure 2b), the northward electric field reaches 120 mv/m (>2400 m/s westward plasma drift) in the dusk sector (2200 UT). It is noteworthy that the electric fields recorded during the July 2000 storm were the largest ever recorded by the Millstone Hill radar. An overflight of DMSP F13 satellite along the Millstone Hill meridian at 2230 UT observed increases in electron and ion particle precipitation at latitudes as low as 52 MLat (invariant latitude), clearly demonstrating that, at the time, the auroral oval extended equatorward beyond the Millstone Hill location. Large (800 2000 m/s) horizontal ion flow was observed by DMSP at latitudes as low as 45 MLat. [17] Figure 2c shows the electric field components for the 31 March to 1 April 2001 period. During the most disturbed time the spectrum of the scattered signal was so distorted that it made derivation of plasma drift and electric field unreliable, leading to a data gap between 0100 UT and 0500 UT. We would expect that the electric field was very high, but as the data are missing, we will concentrate in the later parts of the paper on the afternoon events of 31 March. Figure 2c clearly shows the same pattern in the ion convection as was described for two previous storms. As the storm develops, the plasma drifts in the westward direction in the afternoon and dusk sector (1800 2400 UT) and in the eastward direction after the local midnight (0500 0900 UT). The DMSP F13 pass close to Millstone Hill at 2216 UT on 31 March 2001 (Figure 3c) shows a double-peaked distribution of west- 4of13

Figure 3. Horizontal ion drift data (positive sunward) as observed by DMSP satellite on (a) 25 September 1998, (b) 15 July 2000, and (c) 31 March 2001. ward convection, with a secondary peak several degrees equatorward from the auroral precipitation, as discussed by Foster et al. [2002]. 4.2. E Region Plasma Drift and Neutral Wind Observations 4.2.1. 21 25 September 1998 [18] In the E region, where the atmosphere becomes more dense compared with the F region, ions lose momentum due to collisions with neutral particles. Figure 4 shows the mapping of ion convection down to E region altitudes for the 25 September 1998 storm. The left panel of Figure 4 shows the zonal (top) and meridional (bottom) components of the E region ion drifts in the altitude range 100 150 km, while the right panel of Figure 4 shows horizontal neutral wind components in the 100 130 km altitude range. A large enhancement in the eastward drift component is observed close to 1200 UT on 25 September, when radar measurements for the E region are first obtained. Eastward ion drifts exceed 300 m/s above 120 km but quickly decrease at lower altitudes, change direction, and become westward with speed of 25 m/s below 110 km. This results in the enhanced eastward neutral wind above 120 km in the morning hours (1200 1400 UT) but in westward winds at the same time below 115 km. The northward component of the ion drift at E region heights does not reveal large enhancements, except for an increase in the southward flow between 118 and 125 km. Owing to the increase in ion-neutral collision frequency at altitudes below 140 km, the ion drift vector rotates by an angle that depends on the ratio of gyrofrequency to collision frequency (Pedersen effect). This results in a region of strong meridional drift at altitudes around 120 km, directed to the south in the morning sector and to the north in the evening sector, as observed in meridional drift data for other storms in this study. [19] To fully appreciate the significance of the observed wind disturbance, it is important to establish a reference for the wind pattern during undisturbed times. A quantitative comparison of the wind disturbance is presented in Figure 5, where the top panel depicts the eastward component of the neutral wind, and the bottom panel shows the northward component. The disturbed wind at 1318 UT on 25 September 1998 (triangles) is compared with the average wind observed between 1248 and 1348 UT during 4 quiet days before the storm, 21 24 September 1998 (circles). The error bars for the quiet time represent the standard deviation of the mean for 21 24 September, while the error bars for the disturbed time correspond to the error of a single measurement. During quiet time, the zonal wind (top panel) in the morning hours is primarily westward, with wind speeds approximately 40 90 m/s, as typically seen for tidal oscillations [Goncharenko and Salah, 1998; Zhang et al., 2003]. The meridional wind (bottom panel) is primarily northward, with a smaller magnitude of 0 25 m/s. The storm-time enhancement in the wind field is well pronounced for the zonal component, where above 110 km, the direction of the wind is reversed from westward to eastward, and with changes in speeds of the order of 100 250 m/s. The storm-time disturbance penetrates down to the lowest altitude of 100 km, although the differences in winds are within the measurement 5of13

Figure 4. Zonal and meridional components of the ion drift vector (left panel) and the neutral wind (right panel) on 25 September 1998. uncertainty. A small cell of southward wind is observed at 1300 UT close to 120 km. 4.2.2. 15 16 July 2000 [20] The observations from the July 2000 storm also show a similar pattern, though in this case the increase in the convection results in large westward drift, as should be expected for the evening sector. Figure 6 (left panel) displays the plasma drifts velocities with very large speeds, exceeding 1000 m/s in the E region with a clear reversal from eastward to westward directions after the storm onset at 1900 UT. The results also show that the afternoon convection cell expanded to the radar location, and the westward plasma return flow in the afternoon convection cell extends from the F region to the E region. A region of northward drifts with peaks of 300 400 m/s is seen at 110 125 km after 1900 UT, while the meridional drifts at other times and altitudes are small. [21] The large plasma drifts introduce uncertainties in the simplified method used for deriving neutral dynamics from radar observations and limit the ability to reliably determine the neutral winds during such events. Nonetheless, a derivation of the neutral winds was attempted and the results are shown in Figure 6, right panel. Before discussing these results, it is useful to show the expected winds during quiet conditions. Climatological observations (Figure 7) compiled from many years of E region data [Goncharenko and Salah, 1998; Zhang et al., 2003] show that at Millstone Hill s location during geomagnetically quiet summer time, the winds are primarily driven by the semidiurnal tide, with a maximum wind magnitude of about 100 m/s. Before the storm onset, the zonal neutral wind (Figure 6, top right panel) appears to follow the pattern expected for quiet summer conditions (see Figure 7, top panel), with westward flow in the morning hours, reversing to an eastward direction after 1700 1800 UT. Following the storm onset, an increase in ion convection is observed above Millstone Hill shortly after 1900 UT, and the zonal wind reverses direction to westward. A large cell of westward neutral wind dominates the lower thermosphere above 110 km after 2000 UT, with values in excess of 300 500 m/s. The meridional component is seen to reverse toward large southward velocities, exceeding 300 m/s. The meridional wind is unidirectional at 110 130 km and is similar to equatorward surges commonly reported during magnetic storms in the F region. Such surges result from increased atmospheric pressure, created by precipitating particles and energy dissipation at high latitudes. The wind magnitude is expected to be smaller at lower altitudes. Observations of exceptionally large equatorward wind in 110 130 km range at subauroral latitudes point out once again the extreme nature of this storm and the extension of auroral zone to Millstone Hill s latitude. [22] The large zonal wind cells observed during the storm and their similarity to the ion drifts which drive them indicates that the normal pattern of relatively small amplitudes in the lower thermosphere induced by tidal propagation from the lower atmosphere are disrupted and become dominated by the ion-driven convection during intense storms. Comparison with winds obtained from atmospheric general circulation models [e.g., Fuller-Rowell, 1995] supports the expected large enhancements of neutral dynamics driven by the ion plasma flows as observed with the radar. Westward motion reflecting the expansion of the polar convection cell is predicted from the model, together with 6of13

Figure 5. Highly disturbed wind at 13.3 UT on 25 September 1998 is compared with the average wind observed between 12.8 and 13.8 UT during four quiet days on 21 24 September 1998. a circulation pattern characterized by a return northward flow in the lower thermosphere that completes the southward circulation cell generated in the upper thermosphere. 4.2.3. 31 March 2001 [23] In the case of the 31 March 2001 storm (Figure 8), increased ion speeds are observed in both morning and evening sectors due to the double-peaked nature of the storm. Ion drifts show large zonal speeds, reversing from a southeast direction in the morning hours to a northwest direction in the afternoon (LT = UT 5). In the morning sector, ion drifts reach their maximum observed values close to 0800 UT. The eastward ion drift exceeds 550 m/s at 150 km, 450 m/s at 130 km, and decreases to 80 m/s at 112 km. The afternoon sector displays larger ion speeds, with westward drift close to 1000 m/s at 150 km, 700 m/s at 130 km, and 250 300 m/s at 112 km. [24] It is important to note that high ion drift speeds observed by the radar in the afternoon sector are not really part of the storm-enhanced two-cell ion convection flow, which at the time lies more poleward (see Figure 3c), but are a signature of a subauroral polarization stream (SAPS), as discussed by Foster et al. [2002]. SAPS convection is often observed equatorward from a two-cell auroral convection, and ion drift peak velocities can be comparable to velocities in the auroral pattern [Foster and Vo, 2002]. According to statistical studies of SAPS [Foster and Vo, 2002], this feature spans 3 6 in latitude, and its separation from the auroral convection region increases with increased K p [Yeh et al., 1991]. Though it is most frequently observed in the nighttime sector, it shifts with increasing K p to earlier local times and lower latitudes, thus providing a strong influence on neutral wind dynamics at subauroral and middle latitudes. [25] Another interesting feature in Figure 8 is the large variation of midlatitude E region ion velocity observed around 2200 2300 UT. Zonal ion speeds around 2200 UT are reduced from 600 m/s to 200 400 m/s, while meridional ion drifts above 120 km change direction from northward to southward. At the same time, the reaction of the F region ionospheric plasma manifests as a 10 mv/m drop in the northward component of the electric field (see Figure 2c), while the electron density briefly increases by 50 100%. The GPS maps of total electron content based on data from >120 GPS receiving sites shows that the Storm Enhanced Density (SED) plume discussed by Foster et al. [2002] in the time period between 2130 UT and 2300 UT weakens and moves in the westward direction, and a brief highly localized enhancement in the electron density is observed over New England. Though the discussion of the motions of SED plumes and mechanisms driving the changes in the ionospheric electron density is beyond the scope of this study, we merely point out here that the variations of E region ion drifts around 2200 2300 UT are a real feature and are an inseparable part of the highly dynamic state of the ionosphere during the storm. [26] The complicated picture of the ion drift motion on 31 March 2002 produces an even more complex neutral wind pattern, as shown on the right side of the Figure 8. The eastward component of the neutral wind increases to maximum speeds of 150 200 m/s in the early morning hours, between 0900 and 1200 UT, and forms a strong cell with winds speeds exceeding 200 m/s in the afternoon (1700 2000 UT). Westward winds become dominant at altitudes below 130 km after 1900 UT, penetrating later to lower altitudes, and reach speeds up to 400 600 m/s at 130 km and up to 150 200 m/s at 112 km. The meridional wind magnitude and direction vary. A strong cell of northward wind is observed in the morning hours (1100 1700 UT), when K p decreased to 6+ and 7. The cause of this wind is not well understood. We suggest that this northward flow represents a return flow in the heating-induced circulation, coupled with the effects of tidal variation, which results in a northward wind reaching its maximum values during this local time. The pattern of tidal variation in the neutral wind during spring is similar to the one observed in summer (see Figure 7), but the magnitude of the wind in spring varies significantly from day to day [Goncharenko and Salah, 1998; Zhang et al., 2003]. While it is difficult to separate contributions from different processes, the northward flow observed at 1100 1700 UT is higher than could be expected due to the usual tidal pattern. Therefore we suggest that this flow results from tidal variation enhanced by a northward return flow. Without additional simultaneous data from other locations it is difficult to verify this suggestion. 7of13

Figure 6. Zonal (top) and meridional (bottom) components of the ion drift vector (left panel) and the neutral wind (right panel) on 15 July 2000. [27] Later, as the strength of the storm increases, the meridional wind changes direction and remains southward (300 600 m/s at 130 km) between 1800 and 2300 UT. Similar to the case of the July 2000 storm, it could result from large pressure gradients as the auroral zone moves closer to the latitude of observations. [28] The large difference between ion drift and neutral wind observed on 31 March 2001 can produce extreme frictional heating, as the ion temperature depends on the magnitude of relative drift between ions and neutrals [e.g., Schunk, 1975; St.-Maurice and Hanson, 1982] T i ¼ T n þ m n ðv i V n Þ 2 3K b where T i and T n are the ion and neutral temperatures, m n is an average neutral mass, K b is Boltzmann s constant, and V i and V n are the ion and neutral speeds, respectively. The ion temperature measurements, presented in Figure 9, confirm strong ion heating and support the case of very large neutral wind. In Figure 9, the circles illustrate the time sequence of T i observed on 31 March (3-point weighted running mean), and the specified altitude represents all data within ±5 km. The solid line shows the mean ion temperature observed on quiet days, 1 April to 4 April 2001. Although the averaged ion temperature data show a lot of variance, it is clear that during quiet time the ion temperatures at 130 km altitude vary between 450 K and 550 K, typical for values over Millstone Hill during equinox conditions and high solar activity [e.g., Holt et al., 2002] (also available at http://www.haystack. edu/madrigal/models/). In contrast, on 31 March the ion Figure 7. Climatological observations of neutral wind during summer quiet time at E region altitudes above Millstone Hill. Note that color scale is different from Figure 6. 8of13

Figure 8. Zonal (top) and meridional (bottom) components of the ion drift vector (left panel) and the neutral wind (right panel) on 31 March 2001. temperature at 130 km increased by 200 300 K between 0300 and 0900 UT and up to 500 K at 2100 UT. This agrees well with the amount of Joule heating expected from the difference between ion drifts and neutral winds described earlier. 5. Discussion [29] The zonal wind observations described above are generally consistent with the ion drag mechanism, which can modify the neutral air motion through momentum transfer from ions to neutrals. If the given pattern of ion motion persists for long enough time, momentum will be transferred from ions to neutrals by collisions. The timescale, necessary for the neutral species to approach constant ion drift velocity, is given by the interplay of decreasing ion drift as the storm subsided and more efficient ion momentum transfer as the electron density increased. At 1200 UT, eastward ion drifts reach 300 350 m/s at altitudes above 120 km, and eastward neutral wind speeds are only 90 130 m/s. Close to 1400 UT, eastward neutral winds increase to 150 t ¼ n n n i n in where t is the ion drag time constant, n n is neutral density, n i is ion density, and n in is ion-neutral collision frequency. Since the ion-neutral collision frequency in the lower thermosphere is roughly proportional to the total neutral density, the ion drag time constant is then dependent mainly on ion density. In the 25 September 1998 case, for an altitude h = 115 km the ion drag time constant decreases from 10 hours at UT = 12 h (n i =6.8 10 4 cm 3 )to 6 hours by UT = 13.3 (n i =1.1 10 5 cm 3 ), since electron density rapidly rises after the local sunrise. The behavior of the zonal wind in the 25 September 1998 case (Figure 4, right top panel) clearly shows these changes and reflects Figure 9. Time series of T i, averaged within 5 km from specified altitude. Solid line shows averaged data for the quiet time 1 April to 4 April 2001, while line with circles illustrates extreme Joule heating observed during 31 March 2001 storm. 9of13

180 m/s at altitudes above 120 km, while ion drift speeds slow down to 200 220 m/s. [30] In the case of the 15 July 2002 storm, strong particle precipitation increased the electron density in the E region after 2000 UT to the levels usually observed in the daytime F region. It resulted in the decrease in the ion drag time constant from 4.2 hours before 2000 UT to 1 1.5 hours at 2000 2200 UT. As seen in Figure 6, the zonal neutral winds (top right panel) tend to catch up with the zonal ion drift (top left panel) rather rapidly. As both convectiondriven ion drifts and electron density increase, wind disturbances penetrate deeper to lower altitudes, with the largest disturbances observed at 2100 UT. [31] The penetration of convection-driven wind disturbances to lower altitudes is strongly controlled by two major factors. First, due to increased collisions, ion drift speeds quickly subside at lower altitudes. Second, for any given case, the efficiency of ion-neutral momentum transfer depends on the ion density. For average daytime conditions, the E region layer has a maximum ionization at 110 115 km, with a slight minimum above this altitude and subsequent slow rise with altitude. Below 110 km, the electron density decreases, with the most pronounced drop starting at 100 km. The typical ion drag time constant at 105 km rises to over 10 hours for the daytime and is significantly larger at night. Similar time constants are found for the Chatanika location [Johnson et al., 1987] and EISCAT [Nozawa and Brekke, 1995], where below 105 km the ion drag time constant is of the order of 12 hours. [32] The expected efficiency of ion-neutral momentum transfer, and hence the scope of E region neutral wind disturbances can be estimated based on climatological studies of E region density. A comprehensive study by Ivanov-Kholodny and Nusinov [1979] based on ionosonde data shows that the maximum E region electron density can be described by the following expression N e ¼ N 0 e ðcos cþp where c is the solar zenith angle, and parameters N 0 e and p vary with latitude and season. For middle latitudes, p ffi 0.6 for equinoxes and p ffi 0.54 for solstices, while for a fixed zenith angle, N 0 e in winter is 5 10% higher than in summer and has additional maxima up to 10% at the equinoxes. Note that notwithstanding seasonal variations in N 0 e and p, the dependence of maximum E region electron density on solar zenith angle is dominant. Thus for the same local time, the typical E region electron density in winter will be much lower than in summer due to the higher solar zenith angle. Buonsanto s [1990] study of the maximum E region critical frequency f o E for the whole solar cycle based on data from Boulder (40.0 N, 254.7 E) and Wallops Island (37.8 N, 284.5 E) describes the observations based on local time instead of solar zenith angle. It shows that, for noontime, f o E in summer is 1.15 1.25 times larger than in winter (i.e., 1.3 1.6 factor in electron density). [33] Another source of variation in f o E is the dependence on solar activity. According to Buonsanto [1990], f o E for high solar activity is 1.2 higher than for the low solar activity for the same season. In more general terms, the correlation between maximum E region critical frequency f o E and solar flux index F10.7 can be expressed [Ivanov- Kholodny and Nusinov, 1979] as f o E ½1 þ 0:094ðF 10:7 66ÞŠ 1=4 [34] Thus one could expect that in the absence of an increase in electron density due to particle precipitation, neutral wind disturbances in the lower thermosphere induced by a geomagnetic storm would be higher at maximum solar activity, especially during summer and equinoxes. These general expectations are in good agreement with observations reported here, where the storms under study occurred during moderate and high solar activity. Two of the storms occurred during equinox, and one occurred during the summer with a particle precipitation event. [35] The observations reported here are consistent with the finding by Johnson et al. [1987] that periods of enhanced activity result in an increased eastward flow (by 80 m/s at 115 km) in the local morning sector and increased westward flow in the evening sector, while changes in the meridional wind component are more difficult to discern. Similar conclusions were reached in studies performed for other altitudes and latitudes. A recent study of F region nighttime wind pattern from Fabry-Perot measurements over Millstone Hill by Fejer et al. [2002] shows 130 m/s early-night westward and 40 80 m/s late-night eastward disturbance winds. The same local time pattern, i.e., westward disturbance wind in the afternoon sector and eastward wind in the morning sector, was found in the average daytime F region disturbance winds measured by WINDII at high midlatitudes [Emmert et al., 2001]. Extending the study of WINDII winds down to 90 km, Emmert et al. [2002] found that at high midlatitudes (55 magnetic latitude) disturbance winds follow this local time pattern for altitudes above 120 km, with maximum disturbance winds observed at 120 150 km. The magnitude of the wind perturbations in these statistical studies are much smaller than reported in our observations, primarily due to averaging and the inclusion of smaller storms in the studies. As Fejer et al. [2002] point out, westward disturbance winds increase sharply for K p > 5 and reach 300 m/s for K p 8, while their variability also increases with magnetic activity. Large zonal wind speeds (400 650 m/s) were seen at 42 magnetic latitude in the Dynamic Explorer-2 data [Reddy and Mayr, 1998], and even higher winds speeds, reaching 700 m/s at 180 km, were found in case studies of winds from WINDII by Zhang and Shepherd [2002], while observed features in these case studies clearly indicate that neutrals are primarily driven by ions. [36] While various observations agree that zonal winds during disturbed conditions are driven by ions, with a westward direction before local midnight and an eastward direction after local midnight, the behavior of meridional winds is more complicated. According to studies of average nighttime disturbance winds from the Millstone Hill FPI data by Fejer et al. [2002], meridional perturbation winds show large seasonal and solar activity dependence, and the nighttime F region meridional wind is mostly southward in direction, with maximum magnitude of 60 m/s at 0300 LT (0800 UT). The equatorward surges in the meridional wind were found in numerous case studies of Millstone Hill 10 of 13

incoherent scatter radar F region data, with wind magnitudes reaching 300 400 m/s [e.g., Buonsanto et al., 1990; Buonsanto, 1995]. They are commonly observed in the F region in the postmidnight periods, while daytime surges are relatively rare. Such surges are usually interpreted as a result of changes in global thermospheric circulation due to highlatitude heating and expansion of the neutral atmosphere. At lower altitudes (115 km), high-latitude meridional winds at Chatanika [Johnson et al., 1987] are directed to the south in the morning hours and turn northward later in the afternoon. Above EISCAT [Nozawa and Brekke, 1995], the northward turning of the neutral winds occurs earlier, before local noon. [37] As equatorward wind surges arise due to pressure gradients set up by high-latitude Joule heating, their magnitude is expected to be smaller at lower altitudes. According to simulations from a general circulation model [Fuller- Rowell, 1995], the equatorward circulation cell generated in the upper thermosphere is completed by a return northward flow in the lower thermosphere at locations to the south of regions of high-latitude Joule heating. Thus the meridional component of the wind pattern would be highly dependent on the altitude, ion drift (hence ion drag characteristics), and distance from the high-latitude heating regions. When these effects are added to the general circulation induced by the solar tide, a highly variable wind field is obtained, as illustrated in the various results shown in this paper. [38] Our results indicate that large thermospheric winds in the E region are driven by large ion drifts through ion drag. However, the ion drag mechanism cannot explain all observed features in the neutral wind field. It is important to remember that other hydrodynamic forces, such as Coriolis, pressure gradients, viscosity and Joule heating can also affect the wind pattern. Thayer et al. [1995] studied the effect of these forces at high latitudes and found that they can be significant and that during particular periods, the pressure gradients can even exceed the ion drag force. [39] Effects of other (non-ion drag) forces can be evident in the altitude variations of neutral wind. Statistical analysis of WINDII winds by Richmond et al. [2003] shows high correlation of winds with the ionospheric convection pattern, with a lag of several hours below 125 km as can be expected for an ion drag-driven wind. However, at altitudes above 125 km, the wind pattern is rotated 1.5 hours earlier with respect to the convection pattern, which is not expected from the Larsen and Mikkelsen [1987] model. Richmond et al. [2003] suggest several possible causes for differences between Larsen and Mikkelsen [1987] model and data, including neglected Joule heating in the model. St.-Maurice et al. [1999] report exceptionally high neutral wind at high latitude, which reached 350 400 m/s around 120 km, increased to 850 m/s at 160 km, started diminishing above 175 km, and then decreased to 100 m/s at 278 km. They attribute this result to ion drag forces in the lower ionosphere and the dominance of pressure gradient forces in the F region. Emmert et al. [2002] present additional evidence of the influence of other forces, pointing out that altitudional variations in wind magnitude are most pronounced below 150 km and winds have maxima at 120 140 km. This agrees well with the observations reported in this study, as our results indicate lower thermospheric winds with higher speeds than usually reported in the F region. To understand the extent of effects from other forces at subauroral and middle latitudes as well as in both the E region and F region, it would be extremely beneficial to simulate severe storm conditions using models based on first principles. [40] Disturbances in the zonal component of the neutral wind in the lower thermosphere at subauroral and midlatitudes are primarily driven by increased ion convection. We identify two fundamentally different sources of increased ion convection during geomagnetic disturbances: equatorward expansion of the auroral convection and SAPS. As the auroral region expands equatorward during storms, it still (mostly) retains a standard two-cell pattern, which reaches subauroral and middle latitudes mostly in the nighttime sector. As the nighttime electron density in the E region is low, the ion-neutral coupling is not efficient. The lower thermospheric winds at subauroral latitudes can be disturbed due to expansion of the auroral convection pattern only during extremely strong events, as happened during the 15 July 2000 storm. However, SAPS, the secondary, subauroral plasma convection peak, which is often associated with a deep F region ionization trough, is a persistent feature of the disturbed convection pattern and is observed several degrees equatorward from the main convection pattern. Though it is most frequently observed in the nighttime sector, it shifts with increasing K p to earlier local times and lower latitudes, thus providing a strong source of influence on neutral wind dynamics at subauroral and middle latitudes. [41] It is important to note that though numerous results suggest westward disturbance winds in the evening sector and eastward disturbance winds in the morning sector during geomagnetic storms, the wind pattern for any particular event will reflect a high-latitude convection geometry [i.e., Heelis and Hanson, 1980] and might have a variety of shapes depending on the configuration of interplanetary magnetic field. The interesting challenge for the next few years lies in the understanding of details of ionospherethermosphere response to magnetospheric forcing. [42] It also should be noted that the uncertainty in neutral density could cause large and systematic errors for the extreme events studied here. The NRLMSISE-00 model, used in this paper, contains more data covering extremes of geophysical condition (i.e., solar flux, geomagnetic parameters) compared with earlier models and is expected to provide more accurate neutral density predictions. Still, one clearly cannot expect statistical model to capture local structure associated with a particular storm. Multiple sensitivity analysis, not included in this paper, indicates that if the neutral density is underestimated, the true wind would penetrate less deeply than indicated by our wind estimates. Similarly, if the neutral density is overestimated, the true wind would penetrate deeper in thermosphere. In the altitude range 115 125 km and for electric field strength 30 50 mv/m, a 25% deviation from the NRLMSISE-00 model leads to a difference in wind magnitude of 100 200 m/s. Although large errors in neutral density can cause a difference in penetration altitude of the disturbed wind of the order of several kilometers, they do not change main interpretation of the results. 6. Summary [43] We have used incoherent scatter radar observations from Millstone Hill to study changes in the winds patterns 11 of 13

in the lower thermosphere during three recent geomagnetic storms. The main results of this study are as follows: [44] 1. The radar observations of drifts in the F region at subauroral latitude during intense geomagnetic storms reveal electric fields as high as 100 mv/m. Strong electric fields drive E region plasma drifts primarily in the two-cell convection pattern direction, with magnitudes of 300 500 m/s, eastward in the morning and westward in the evening. During the most intense storm in July 2000, westward ion drifts as large as 1000 m/s were observed. Smaller drift velocities in the meridional direction are measured, and a common feature amongst the storms is the development (due to the Pedersen effect) of a southward cell of plasma drift in the 120 km altitude range in the morning sector, turning to a northward cell in the afternoon sector. [45] 2. Derivation of neutral winds in the lower thermosphere during the storms is subject to larger errors due to uncertainty in the neutral density and increased significance of such uncertainty for periods with large electric field. Our estimates suggest the dominance of a westward component in the afternoon sector, indicating a strong coupling to the polar convection pattern that expands towards middle latitudes. The meridional component of the neutral wind turns southward, similar to equatorward wind often reported in the F region during storms. The wind magnitude exceeds 700 800 m/s at 120 130 km, decreasing to 500 m/s at 110 km and 100 m/s at 100 km during the 15 July 2000 storm. In the case of 31 March 2001 storm winds speeds reached 300 700 m/s at 120 130 km. This sharply contrasts with semidiurnal wind pattern having speeds of the order of 100 m/s, which usually observed over Millstone Hill during geomagnetically quiet time. [46] 3. The observations presented here show general agreement in the local time dependence of the wind pattern with previously reported results for higher latitudes or altitudes but reveal lower thermospheric wind speeds much higher than earlier studies. High wind speeds, reported in this paper, could result from both extremely disturbed conditions and more efficient ion-neutral coupling in the E region as compared with the F1 and F regions. [47] 4. The ion drag mechanism is shown to be a primary mechanism driving storm-induced zonal wind disturbances. We also identify the major factors responsible for the penetration of convection-driven ion flows to unusually low altitudes as well as the efficient momentum transfer from ions to neutrals. These factors, besides high ion speeds, include the sustainment of ion drift flow for a long period of time, the relatively high E region density typical for summer and equinox seasons under high solar activity and, in the case of the 15 July 2000 storm, the increased electron density due to particle precipitation. [48] 5. The meridional component of the neutral wind shows a complex response to geomagnetic forcing. We identify southward flow at 120 km as response to ion drag, as observed during September 1998 storm. We report northward flow at 1100 1700 UT on 31 March 2001 and interpret it as northward return flow in heating-induced Hadley circulation, coupled with northward winds resulting from tidal variation. Finally, the most pronounced feature during 15 July 2000 and 31 March 2001 storms is a strong southward wind, similar to equatorward winds commonly reported in the F region. [49] 6. During the March 2001 storm, we find a good agreement between storm-time increase in ion temperature and frictional heating expected from the difference in the neutral winds and plasma drifts. The observed rise in the ion temperature is an independent confirmation of the strength of the estimated neutral wind. [50] 7. Results of this study illustrate the important coupling that takes place between the polar and middle latitudes during intense storms and reveal the influence of the storm effects to altitudes as low as 100 km. [51] 8. Increased ion convection, necessary to affect lower thermospheric dynamics, is a result of equatorward expansion of the auroral convection pattern during geomagnetic storms or (more likely) Subauroral Polarization Streams (SAPS). This study directly links SAPS, a phenomena associated with the erosion of the outer plasmasphere, and dynamics in the lower thermosphere. The fundamental difference between two sources of large electric fields and their link to a variety of plasmaspheric and magnetospheric phenomena provide an opportunity for subsequent studies of the magnetosphere-ionosphere-thermosphere as a coupled system. [52] Acknowledgments. The analysis of the observations during geomagnetic storms was supported under the TIMED/CEDAR grant from NSF (ATM-0000958). Millstone Hill radar observations and analysis are supported by a NSF cooperative agreement with Massachusetts Institute of Technology. The authors are grateful to the Millstone Hill radar staff for the effort involved in the radar operation and data processing. We thank F. J. Rich for providing access to DMSP data and H. G. Mayr for useful discussions. [53] Arthur Richmond thanks John Emmert and another reviewer for their assistance in evaluating this paper. References Banks, P., and G. Kockarts (1973), Aeronomy, Academic, San Diego, Calif. Buonsanto, M. (1990), A study of the daytime E-F 1 region ionosphere at mid-latitudes, J. Geophys. Res., 95, 7735 7747. Buonsanto, M. J. (1995), Millstone Hill incoherent scatter F region observations during the disturbances of June 1991, J. Geophys. Res., 100, 5743 5755. Buonsanto, M. J., J. C. Foster, A. D. Galasso, D. P. Sipler, and J. M. Holt (1990), Neutral winds and thermosphere/ionosphere coupling and energetics during the geomagnetic disturbances of March 6 10, 1989, J. Geophys. Res., 95, 21,033 21,050. Emmert, J. T., B. G. Fejer, C. G. Fesen, G. G. Shepherd, and B. H. Solheim (2001), Climatology of middle- and low-latitude daytime F region disturbance neutral winds measured by Wind Imaging Interferometer (WINDII), J. Geophys. Res., 106, 24,701 24,712. Emmert, J. T., B. G. Fejer, G. G. Shepherd, and B. H. Solheim (2002), Altitude dependence of middle and low- latitude daytime thermospheric disturbance winds measured by WINDII, J. Geophys. Res., 107(A12), 1483, doi:10.1029/2002ja009646. Fahrutdinova, A. N., O. N. Sherstyukov, and S. V. Maksyutin (2001), Geomagnetic activity influence on the dynamics of the upper mesosphere and lower thermosphere, Int. J. Geomagn. Aeron., 2, 201 208. Fejer, B. G., J. T. Emmert, and D. P. Sipler (2002), Climatology and storm time dependence of nighttime thermospheric neutral winds over Millstone Hill, J. Geophys. Res., 107(A5), 1052, doi:10.1029/2001ja000300. Foster, J. C., and H. B. Vo (2002), Activity dependence of subauroral electric fields, J. Geophys. Res., 107(A12), 1475, doi:10.1029/ 2002JA009409. Foster, J. C., P. J. Erickson, A. J. Coster, J. Goldstein, and F. J. Rich (2002), Ionospheric signatures of plasmaspheric tails, Geophys. Res. Lett., 29(13), 1623, doi:10.1029/2002gl015067. Fuller-Rowell, T. J. (1995), The dynamics of the lower thermosphere, in The Upper Mesosphere and Lower Thermosphere: A Review of Experiment and Theory, Geophys. Monogr. Ser., vol. 87, edited by R. M. Johnson and T. L. Killeen, pp. 23 36, AGU, Washington, D. C. Goncharenko, L. P., and J. E. Salah (1998), Climatology and variability of the semidiurnal tide in the lower thermosphere over Millstone Hill, J. Geophys. Res., 103, 20,715 20,726. 12 of 13