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  • View in gallery

    (a) Time–height cross section of radar reflectivity (dB) after Capon processing and after correcting the range attenuation effects (i.e., power multiplied by r2) measured by the VHF MU radar in vertical incidence from 2200 LT 7 Jun 2008 to 0715 LT 8 Jun 2008 between 1.32- and 18-km altitudes. The nearly vertical thin white lines above 10 km are airplane echoes. The black line shows the balloon altitude vs time launched at 2357 LT 7 Jun 2006. The letters A, B, and C denote the echoing layers discussed in the text. (b) (top) the corresponding time–height cross section of vertical velocity of air motions measured by the VHF MU radar. (bottom) The corresponding time–height cross section of variance of the Doppler spectrum peaks using the vertical beam and after correcting from the beam-broadening effect. Values are not given for SNR < −5 dB.

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    (a) Vertical profiles of air (black) and dewpoint (gray) temperature obtained from radiosonde data at 2357 LT 7 Jun 2006 released from Shigaraki MU Observatory. The inset shows a closeup of the temperature profile between 7.4 and 8.8 km. (b) (left) Vertical profiles of dry squared Brunt–Vaïsälä frequency (multiplied by 1000, gray line) and Richardson number (Ri; black line). The vertical dashed line shows the value of Ri = 0.25. (right) The corresponding profiles of wind speed (m s−1, dashed black line), wind shear (m s−1 km−1, solid black line), and wind direction (°, solid gray line) shifted by −90°.

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    (top) Time–height cross section of horizontal wind speed (m s−1), (middle) wind direction (°), and (bottom) wind shear (m s−1 km−1) from 0430 to 0630 LT 8 Jun 2006. The black lines show the 28-dB contour of radar reflectivity shown in Fig. 1a.

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    (top) Time–height cross section of RMR lidar backscatter ratio (dB) from 2206 LT 7 Jun 2006 to 0236 LT 8 Jun 2006 showing cirrus clouds (and layers of aerosols mainly below 3 km). Optical attenuation was strong enough to cause range-limiting effects at some periods and altitudes. (bottom) The inset in the top is expanded: white dashed box in the top indicates inset in the bottom.

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    Infrared (IR1, 10.3–11.3 μm) images of the geostationary meteorological satellites showing high-level clouds, and then likely cirrus, over the southwest of Japan at 0200 LT 8 Jun 2006. The dot shows the MU radar location.

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    Time–height cross section of radar reflectivity (dB) after correcting the range attenuation effects for the period of lidar observations and between 6- and 9-km altitudes using (top) the vertical beam and (bottom) the oblique (180°, 10°) beam. The black contour shows the cirrus cloud edges arbitrarily defined as the 9.54-dB level of the backscatter ratio shown in Fig. 4.

  • View in gallery

    (top) Time–height cross section of vertical velocity (m s−1) for the period of lidar observations and between 6- and 9-km altitudes MSL. (bottom) The corresponding variance of Doppler spectrum (m2 s−2) at the range resolution of 150 m and after correcting the beam-broadening effects. The black contours show the cirrus edges arbitrarily defined in the same way as in Fig. 6.

  • View in gallery

    Expanded views of the time–height cross section of (top) vertical wind (m s−1) and (bottom) from 0055 to 0200 LT 8 Jun 2006 between 6.0 and 9.0 km. The black line contour shows the cirrus edges arbitrarily defined as in Fig. 6.

  • View in gallery

    (a) Time series of vertical velocities averaged between 6.57 and 7.92 km in the altitude range of turbulent layer C. (b) The corresponding power spectral density. A −3 slope line is also shown as a reference (gray line).

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MU Radar and Lidar Observations of Clear-Air Turbulence underneath Cirrus

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  • 1 Laboratoire de Sondages Electromagnétiques de l’Environnement Terrestre, Université du Sud-Toulon Var, La Garde, France
  • | 2 Research Institute for Sustainable Humanosphere, Kyoto University, Uji, Japan
  • | 3 Department of Space Information Science and Engineering, Fukui University of Technology, Fukui, Japan
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Abstract

Turbulence generation mechanisms prevalent in the atmosphere are mainly shear instabilities, breaking of internal buoyancy waves, and convective instabilities such as thermal convection due to heating of the ground. In the present work, clear-air turbulence underneath a cirrus cloud base is described owing to coincident observations from the VHF (46.5 MHz) middle and upper atmosphere (MU) radar, a Rayleigh–Mie–Raman (RMR) lidar, and a balloon radiosonde on 7–8 June 2006 (at Shigaraki, Japan; 34.85°N, 136.10°E). Time–height cross section of lidar backscatter ratio obtained at 2206 LT 7 June 2006 showed the presence of a cirrus layer between 8.0 and 12.5 km MSL. Downward-penetrating structures of ice crystals with horizontal and vertical extents of 1.0–4.0 km and 200–800 m, respectively, have been detected at the cirrus cloud base for about 35 min. At the same time, the MU radar data revealed clear-air turbulence layers developing downward from the cloud base in the environment of the protuberances detected by the RMR lidar. Their maximum depth was about 2.0 km for about 1.5 h. They were associated with oscillatory vertical wind perturbations of up to ±1.5 m s−1 and variances of Doppler spectrum of 0.2–1.5 m−2 s−2. Analysis of the data suggests that the turbulence and the downward penetration of cloudy air were possibly the consequence of a convective instability (rather than a dynamical shear instability) that was likely due to sublimation of ice crystals in the subcloud region. Downward clear-air motions measured by the MU radar were associated with the descending protuberances, and updrafts were observed between them. These observations suggest that the cloudy air might have been pushed down by the downdrafts of the convective instability and pushed up by the updrafts to form the observed protuberances at the cloud base. These structures may be virga or perhaps more likely mamma as reported by recent observations of cirrus mamma with similar instruments and by numerical simulations.

* Current affiliation: National Institute of Polar Research, Tokyo, Japan.

Corresponding author address: Hubert Luce, Laboratoire de Sondages Electromagnétiques de l’Environnement Terrestre, Université du Sud-Toulon Var, La Garde, 83957, France. Email: luce@lseet.univ-tln.fr

Abstract

Turbulence generation mechanisms prevalent in the atmosphere are mainly shear instabilities, breaking of internal buoyancy waves, and convective instabilities such as thermal convection due to heating of the ground. In the present work, clear-air turbulence underneath a cirrus cloud base is described owing to coincident observations from the VHF (46.5 MHz) middle and upper atmosphere (MU) radar, a Rayleigh–Mie–Raman (RMR) lidar, and a balloon radiosonde on 7–8 June 2006 (at Shigaraki, Japan; 34.85°N, 136.10°E). Time–height cross section of lidar backscatter ratio obtained at 2206 LT 7 June 2006 showed the presence of a cirrus layer between 8.0 and 12.5 km MSL. Downward-penetrating structures of ice crystals with horizontal and vertical extents of 1.0–4.0 km and 200–800 m, respectively, have been detected at the cirrus cloud base for about 35 min. At the same time, the MU radar data revealed clear-air turbulence layers developing downward from the cloud base in the environment of the protuberances detected by the RMR lidar. Their maximum depth was about 2.0 km for about 1.5 h. They were associated with oscillatory vertical wind perturbations of up to ±1.5 m s−1 and variances of Doppler spectrum of 0.2–1.5 m−2 s−2. Analysis of the data suggests that the turbulence and the downward penetration of cloudy air were possibly the consequence of a convective instability (rather than a dynamical shear instability) that was likely due to sublimation of ice crystals in the subcloud region. Downward clear-air motions measured by the MU radar were associated with the descending protuberances, and updrafts were observed between them. These observations suggest that the cloudy air might have been pushed down by the downdrafts of the convective instability and pushed up by the updrafts to form the observed protuberances at the cloud base. These structures may be virga or perhaps more likely mamma as reported by recent observations of cirrus mamma with similar instruments and by numerical simulations.

* Current affiliation: National Institute of Polar Research, Tokyo, Japan.

Corresponding author address: Hubert Luce, Laboratoire de Sondages Electromagnétiques de l’Environnement Terrestre, Université du Sud-Toulon Var, La Garde, 83957, France. Email: luce@lseet.univ-tln.fr

1. Introduction

The impact of clear-air turbulence on energy budget and vertical transport of heat and constituents is now well recognized even though still difficult to estimate. In the stably stratified atmosphere, turbulence is often confined within layers as thin as a few tens of a meter or less (e.g., Sato and Woodman 1982; Barat 1982). The main sources in the lower atmosphere are convective and dynamic instabilities. Some of these mechanisms have been reviewed by Fritts and Alexander (2003), for example. Thermal convection due to heating of the ground is also a very important source of turbulence. It mainly affects the planetary boundary layer and the troposphere and can also be an important source of gravity waves through various mechanisms (e.g., Kuettner et al. 1987).

In the present paper, we describe clear-air turbulence observations with the middle and upper atmosphere (MU) radar operating at very high frequencies (VHFs; Fukao et al. 1990). The MU radar is a mesosphere–stratosphere–troposphere/incoherent scatter (MST/IS) radar mainly sensitive to humidity and temperature irregularities in the neutral atmosphere and to electron density fluctuations in the ionized atmosphere. During observations performed in 7–8 June 2006, the MU radar revealed 0.5–2-km-deep turbulent layers with roll-like appearance developing downward below 8.0 km above mean sea level (MSL). Coincident observations with a Rayleigh–Mie–Raman (RMR) lidar (Behrendt et al. 2004) showed a 4-km-deep layer of cirrus above 8.0 km. The observed turbulence was likely the result of convective currents due to cooling by sublimation of ice crystals beneath the cirrus cloud base.

The RMR lidar also detected downward-protruding structures at the cloud base. The coincident radar observations suggest that downward and upward air motions generated by the convective instability very likely contributed to their formation. The protuberances observed by the RMR lidar bear striking similarity to cirrus mamma observed with a cloud polarization lidar (CPL; Sassen et al. 2001; Wang and Sassen 2006, Schultz et al. 2006). Mamma are “smooth lobe-like hanging protuberances extending downward from the base of cumulonimbus anvils and stratiform clouds like cirrus” (Glickman 2000). CPL observations revealed that cirrus mamma, which are more ragged than anvil mamma, are apparently quite common (K. Sassen 2008, personal communication). There are thus strong presumptions that mamma were indeed detected by the RMR lidar. However, because of the lack of visual identification of these structures, we will broaden the discussion to downward protuberances on the underside of cirrus.

The spatial resolution of a classical VHF stratosphere–troposphere (ST) radar is rather poor compared to the resolution of lidars and millimeter wavelength radars used for cloud dynamic studies. However, a range imaging technique based on frequency diversity [i.e., the Frequency Radar Interferometric Imaging (FII)] enables us to improve the range resolution of the MU radar (Luce et al. 2001a; Hassenpflug et al. 2008). The range resolution is typically improved by a factor of about 5 for high signal-to-ratios (SNRs) so that a few tens of meters of range resolution is obtained from an initial 150-m range resolution. Thin stable inversions and Kelvin–Helmholtz (KH) instabilities could be nicely delineated with this technique (e.g., Palmer et al. 2001; Luce et al. 2006, 2008). Also, the use of a Doppler VHF radar can be an advantage with respect to millimeter radars for measuring vertical air motions. A VHF (∼50 MHz) radar is sensitive to clear-air refractive index irregularities both in clear and precipitation conditions and thus can measure the clear-air motions without being affected by vertical motions of hydrometeors as is the case for millimeter wavelength radars (e.g., Fukao et al. 1985; Gage 1990). By comparing vertical velocities measured by the 47.0-MHz equatorial atmosphere radar, a Doppler-pulsed radar similar to the MU radar, and a W-band (95 MHz) radar, Yamamoto et al. (2008) confirmed that VHF radars can provide measurements of the vertical velocity of the clear-air motions in cloudy air at cirrus levels. This property is thus useful for studying clear-air dynamics within or around cirrus (e.g., Nishi et al. 2007; Yamamoto et al. 2008).

In section 2, the instruments, dataset, analysis methods used are briefly described. The results of observations with the MU radar, the RMR lidar and a GPS radiosonde are described in section 3. Possible interpretations and nature of the observed cloudy protuberances are proposed in section 4. Conclusions are given in section 5.

2. Instruments, analysis method, and dataset

The MU radar is a flexible and fast beam steering Doppler-pulsed radar operating between 46.0 and 47.0 MHz (3.5-MHz bandwidth and 1-MW peak output power; e.g., Fukao et al. 1990). It is located at Shigaraki MU Observatory (34.85°N, 136.10°E) in Japan and has been developed for studying the variability of the earth’s atmosphere (from the lower troposphere up to the ionosphere). The antenna area consists of 475 Yagi antennas arranged in a circular array of 103-m diameter with a resulting beamwidth of 3.6°. In 2004, the MU radar system was upgraded for being operated with a digital receiver system and for improving its spatial resolution (Hassenpflug et al. 2008). The FII observational mode is one of its new capabilities. Basically, the FII technique consists in transmitting several closely spaced frequencies switched pulse to pulse. The collected data at the various frequencies are processed using the adaptive Capon processing method (Palmer et al. 1999; Luce et al. 2001a). A more detailed description of the range imaging processing with the upgraded MU radar can be found in Luce et al. (2006). Also, the radar parameters applied for collecting the data discussed in the present paper are similar to those given by Luce et al. (2008). In particular, the MU radar was operated with five equally spaced frequencies from 46.0 to 47.0 MHz (i.e., with a frequency spacing of 0.25 MHz). Range sampling was performed from 1.32 up to 20.37 km MSL with a step of 5 m (the initial range resolution was 150 m). The number of coherent integrations was set equal to 32 times and one profile is calculated every ∼10 s and the time resolution is ∼20 s. The radar beam was steered into five positions: one vertical and four oblique directions at a zenith angle of 10° off zenith toward north, east, south, and west). In the present paper, high resolution time–height cross sections of echo power (or reflectivity) and vertical wind velocity are described. The other parameters are shown at the initial range resolution of 150 m.

The RMR lidar was originally installed at Shigaraki MU Observatory in 2000. The system has been upgraded to have pure rotational Raman scatter detection in order to lower the observational height for temperature down to 1–2 km (Behrendt et al. 2004). The transmitter is an injection-seeded-pulsed Nd:YAG laser with 532.25 nm output of 600 mJ pulse−1 at 50-Hz repetition frequency. The receiving telescope is a Cassegrainian telescope with a diameter of 82 cm. Received signal is divided into five channels, two-elastic channels for Rayleigh and Mie scatter detection, water vapor vibrational Raman channel, and two pure rotational Raman channels, and recorded by both photon counting and A/D recording. In the current study, observations were carried out with an integration time of 15 s and a height resolution of 18 m. A Raman reference signal (Behrendt et al. 2004; i.e., a temperature-independent signal proportional to molecular density) was obtained from two rotational channels, and then backscatter ratio was calculated with signal intensities of low sensitivity elastic channel and Raman reference signal. The backscatter ratio is used as an index of clouds in the present work.

The concurrent MU radar and RMR lidar experiment was carried out for about 3 days from 1623 LT 6 June 2006 to 1556 LT 9 June 2006. The lidar was operated during night periods. Because of precipitations occurring the third night, data could only be collected during the first two nights. In the present paper, we will show the RMR lidar observation results from 2200 LT 7 June 2006 to 0230 LT 8 June 2006.

3. Results of observations

a. Radar observations during the night of 7–8 June 2006

For a better understanding of the figures showing radar reflectivity, the main backscattering mechanisms operating at VHF (∼50 MHz) are briefly reviewed. The reader can refer to Röttger and Larsen (1990), for instance, for more details. VHF radar returns are caused by refractive index irregularities due to variations in temperature and humidity. Specular reflection from both thin humidity- and temperature-gradient sheets and Bragg scattering from clear-air isotropic turbulence are the main backscattering mechanisms (e.g., Röttger and Liu 1978; Gage 1990; Luce et al. 2001b). The VHF radars are thus complementary to weather radars that make use of Rayleigh scattering from hydrometeors. Because the thin sheets are generally horizontally spread out, radar echoes produced by these layers are aspect sensitive (i.e., strongly attenuated when using an oblique beam tilted sufficiently far from the zenith; e.g., Tsuda et al. 1986). On the contrary, radar echoes of similar intensities in vertical and oblique beam directions are usually interpreted in terms of scattering from isotropic turbulence (e.g., Gage 1990). At 10°–15° off zenith, scattering from isotropic turbulence is often the dominant backscatter mechanism in the troposphere (e.g., Tsuda et al. 1988) but not in the lower stratosphere (e.g., Worthington et al. 1999; Hirono et al. 2004).

Figure 1 shows time–height cross sections of radar reflectivity [more precisely, the radar echo power (dB) in arbitrary units after Capon processing and after correcting the range attenuation effects] using the vertical beam from 2205 LT 7 June 2006 to 0730 LT 8 June 2006 and for the altitude range of 1.32–18.0 km. The thermal tropopause was around 12.5 km according to the in situ measurements made by a Vaisala RS90G balloon launched at 2357 LT 7 June 2006 from the MU radar site. The enhancement of radar reflectivity above 12.5 km is consistent with the capability of a VHF ST radar to detect the tropopause (e.g., Röttger and Larsen 1990). Multiple thin and long lasting layers of radar backscatter exhibiting vertical oscillations of various scales and amplitudes due to buoyancy waves are observed below 6.0 km and above the thermal tropopause.

The vertical distribution of the radar echoes are very different between 6.0 km and the tropopause. About 3-km-deep vortexlike structures (probably due to convection in clouds) can be seen after 0200 LT in the upper troposphere (above 8.0 km). Unless not so clearly delineated, similar structures were also monitored in the upper troposphere with the UHF Arecibo radar (Woodman and Rastogi 1984). Between 6.5 and 8.0 km, intermittent layers with slanted lines of intense echoes, like paint trickling down a wall, appear at different times. We will focus on the layers seen from 2310 to 0010 LT (layer A), from 0105 to 0230 LT (layer B) and from 0300 LT (layer C). The selection is quite arbitrary and is given for an easy reference in the following description. The pattern of these layers show no resemblance whatsoever to the pattern of KH instabilities (see e.g., Gossard 1990 and references therein for typical radar images of KH billows and waves). From 0105 LT, the layer B starts to grow downward from the altitude of 8.0 km at a rate of 65 m min−1. Its depth reached a maximum of ∼1.2 km at ∼0125 LT. The downward penetration is even faster for the layer C around 0500 LT. The layer depth then levels out (about 2.0 km) for about 90 min before decreasing again.

The corresponding time–height cross sections of Doppler velocity and variance of the Doppler spectrum measured with the vertical beam are shown in Fig. 1b. Positive (negative) vertical velocities indicate upward (downward) vertical component of the air motions. Nearly monochromatic oscillations of vertical wind with periods close to the Brunt–Vaïsälä period (typically 10 min) can be identified below 6.0 km. These oscillations are commonly observed during quiet horizontal winds (say less than 5–10 m s−1 when Doppler shift effects and mountain wave contamination are minimal) in the lower troposphere (e.g., Ecklund et al. 1985). They are also consistent with the wavy oscillations of the multiple layers of radar backscatter revealed by Fig. 1a. More intense and smaller period fluctuations are noted within and below the layers A, B, and C. These fluctuations exceed ±1.5 m s−1 after 0510 LT for layer C and ±0.7 m s−1 between 0105 and 0230 LT for layer B.

The width of the Doppler spectrum basically arises from fluctuating velocity of turbulent scatterers in the radar sampling volume (e.g., Doviak and Zrnic 1984). Since turbulence enhances the variance of the velocity fluctuations, enhanced spectral width indicates the presence of turbulence. However, nonturbulent processes also contribute to the measured spectral width. Beam broadening, shear broadening, specular reflection, and (gravity wave) transient effects must be corrected (or avoided) before retrieving turbulence parameters from the Doppler spectral width (e.g., Hocking 1985). Because measurements are made with a vertical beam at a high time resolution, beam broadening is expected to be the main nonturbulent contribution to the spectral width enhancement in the (nonaspect sensitive) turbulent layers (e.g., Hocking 1985). Because the horizontal wind speed is also estimated from the radar measurements, the variance of the Doppler spectrum corrected from the beam-broadening effects can be obtained. It is approximately given by
i1520-0493-138-2-438-e1
where σobs is the half-width half power of the Doppler peaks, Vh is the horizontal wind speed, and θ0 (∼1.3° = 0.023 rad) is the half power half-width of the effective (two way) radar beam (e.g., Fukao et al. 1994). The bottom panel of Fig. 1b displays time–height cross sections of from the vertical beam at a range resolution of 150 m. Negative values of (1), due to statistical errors when turbulence is weak and Vh is large, have been put to zero and values for SNR < −5 dB removed. The variance is about 0.25 m2 s−2 just above the thermal tropopause (between 12.5 and 15.0 km approximately), about 0.2 m2 s−2 in the upper troposphere where deep roll-like structures are observed and can exceed 0.6 m2 s−2 in the layers A, B, and C (up to about 1.5 m2 s−2 around 0500 LT). These estimates are similar to those reported by Naström and Eaton (1997) with a 50-MHz radar and by VanZandt et al. (2002) with the MU radar in the same altitude range. Local turbulent “bursts” of a few minutes (for which exceeds 1.5 m2 s−2) can also be noted at ∼0200 and ∼0500 LT around 12.0 km. The presence of (at least 2) braided structures of a KH instability is suggested in Fig. 1a (and in the bottom panel of Fig. 1b) around 12.0 km and 0200 LT.

Assuming isotropic turbulence, it is possible to assess the total turbulent kinetic energy (TKE) from (e.g., VanZandt et al. 2002). It is also possible to infer the energy dissipation rate ε and the vertical eddy diffusivity K if the Brunt–Vaïsälä frequency N is known (e.g., Naström and Eaton 2005 and references therein). However, these derivations are only relevant for isotropic and inertial turbulence generated in a stably stratified atmosphere.

b. Dynamical conditions associated with the layers A, B, and C

Figure 2 shows vertical profiles of parameters measured or estimated from data collected by the GPS balloon launched at Shigaraki MU Observatory site from 2357 LT 7 June 2008. Figure 2a displays temperature (K), dewpoint temperature (K) for an altitude range of 1.0–10.0 km. Figure 2b shows wind speed (m s−1), direction (°), vertical shear S of horizontal wind (m s−1 km−1), dry squared Brunt–Vaïsälä frequency N2 = g/θdθ/dz (rad s−1) (θ is the potential temperature) and Richardson number Ri = N2/S2 for an altitude range of 6.0–9.0 km at a vertical sampling of 50 m.

A strong positive temperature gradient (1.8 K over a depth of 80 m) is observed at an altitude of 8.2 km. The dewpoint temperature profile suggests that the gradient coincides with the base of a cirrus cloud (saturation with respect to ice). The profiles also confirm the presence of a very dry region (TdT) underneath the cloud base. The relative humidity was about 20% between 7.7 and 8.2 km and as small as ∼5% between 7.0 and 7.7 km.

The cloud base is associated with a veering wind and a constant wind speed. The dry Richardson number has a maximum (about 3) at the cloud base despite the wind shear exceeding 20 m s−1 km−1 produced by the rotation of wind direction. Two maxima of wind shear (>30 m s−1 km−1) due to a wind speed increase and decrease are observed at 7.6 km and 8.4 km, respectively. At these altitudes, Ri < 0.25. The wind shears are thus conducive to KH instabilities and then dynamical turbulence. However, it is difficult to relate these characteristics to the radar echoes. The altitude ranges where Ri < 0.25 are not associated with enhanced radar echoes. Also, the altitudes where thin echoing layers are present do not coincide with the altitude of the stable gradient. Errors in altitude are not the cause of these differences. Previous comparisons between balloon and similar radar data showed excellent agreements in altitude (e.g., Luce et al. 2007). Disagreements could rather be explained by the horizontal inhomogeneity of the atmospheric parameters in the environment of the cloud base. The balloon altitude versus time is shown by the black line in Fig. 1a. The balloon drifted at a horizontal distance of about 10.0 km southeastward from the Shigaraki MU Observatory at the altitude of about 8.0 km. It is thus difficult to know if the conditions met by the balloon between 6.0 and 8.0 km are fully representative (or not) of those occurring above the MU radar.

The dynamical conditions associated with the turbulent layers can be better identified from the analysis of the horizontal wind and wind shear measured by the MU radar. Figure 3 shows the height–time cross sections of radar-estimated wind speed (top panel), wind direction (middle panel), and vertical shear of horizontal wind (bottom panel) between 5.0 and 10.0 km and from 0430 to 0630 LT. The black lines display the contour of the layer C shown in Fig. 1a. The sporadic maxima of wind shear before 0457 LT and below 7.8 km are artifacts due to low SNR and should not be considered. The wind shear, due to both an increase of the wind speed and oscillations of the wind direction, is maximum and exceeds 30 m s−1 km−1 about 200–400 m above the turbulent layer. The wind shear does not exceed 5–10 m s−1 km−1 inside the layer and below. Similar results are obtained for layers A and C (not shown). Shear-driven turbulence is usually associated with a horizontal wind shear maximum at the altitude of the turbulent layer (e.g., Browning et al. 1973; Sato and Woodman 1982; Fritts et al. 2003). Consequently, this observation result is also a strong indication that a dynamical shear instability is not the cause of the turbulent layers A, B, and C.

c. Lidar observations

The top panel of Fig. 4 shows the time–height cross section of RMR lidar backscatter ratio (dB) from 2206 LT 7 June 2006 to 0236 LT 8 June 2006. A cirrus layer was present above 8.0 km. Despite the optical attenuation, it can be seen that the cirrus cloud extended up to the altitude of 12.5 km (i.e., the tropopause, around 2205 and 0030 LT). Infrared (IR1) satellite images at 0200 LT 8 June 2006 over Japan confirm the presence of cirrus (Fig. 5).

The bottom panel of Fig. 4 shows an expanded view of the backscatter ratio between 0036 and 0206 LT from 7.0 to 9.0 km. Protuberances at the cloud base are particularly noticeable between 0100 and 0140 LT. The vertical extent of the protuberances is about 0.2 to 0.8 km around 0136 LT. Assuming that these structures are advected by the wind at the same speed as the cloud base (i.e., of the order of 20 m s−1, see Fig. 2b), a horizontal extent of 1.0–4.0 km is found.

d. Combined radar–lidar observations

The top panel of Fig. 6 shows an expanded view of Fig. 1a between 2200 and 0230 LT (i.e., during the RMR lidar observations when layers A and B were observed) and between 6.0 and 9.0 km. To identify isotropic and aspect sensitive echoes in layers, the bottom panel shows the corresponding results using the southward-oriented oblique beam. Similar plots are obtained when using the three other oblique beam directions (not shown). Superimposed is the cloud-base contour arbitrarily defined as the 9.54-dB level (9 in linear scale) of the lidar backscatter ratio. This level enables us to nicely delineate the cloud-base irregularities. The contour also shows that the presence of cirrus affects the radar returns. Since a VHF radar is not sensitive to (small) hydrometeors but to clear-air refractive index fluctuations, the effects of cirrus on radar backscattering only result from the refractive index perturbations associated with these clouds (without being able to separate the temperature and humidity contributions, however).

Figure 6 shows that the echoing layers A and B are observed at the cirrus cloud base, within and underneath the protruding structures and are not aspect sensitive. Therefore, the regions where the protuberances are observed are likely turbulent. The turbulent nature of these layers will be confirmed later from the analysis of . Their morphology suggests that they develop near the cloud base around 8 km.

Aspect sensitive echoing layers, sometimes thinner than ∼100 m, can be noted between 7.0 and 8.5 km before 0020 LT and around 8.3 km after 0140 LT in the surroundings of the cirrus. The reflectivity at vertical incidence is enhanced by 10–15 dB with respect to the oblique direction. The echoing layer above 8.0 km appearing from 2240 LT until the stop of the radar operation (∼0020 LT) in the top panel of Fig. 6 is attenuated in the vertical beam by about 10–20 dB when “passing through” the cirrus likely due to air saturation effects (a more thorough interpretation of the phenomenon is beyond the scope of the present work). Two thin aspect sensitive echoing layers can be noted just underneath the cloud base without deep penetrating lobes around 2310 LT indicating that the subcloud layer was quite stable (and maybe humid) around that time. The downward growth of the turbulent layers came after periods during which the radar echo power was minimum underneath the cloud base (around 2345 LT for layer A and 0100 LT for layer B). The downward extension of layer C also occurs after a period of weak reflectivity between 7.0 and 8.0 km (see Fig. 1a). Weak radar echoes can be caused by a weak background stability and a dry air (as revealed by the balloon data in Fig. 2b below the saturated layer). Therefore, the downward development of both the cloud protuberances and the turbulent layers underneath cirrus was likely allowed by the presence of a dry and weakly stable air in the subcloud layer while a similar development was possibly inhibited around 2310 LT due to the presence of stable and/or humid layers at that time.

Both the cloud base and a thin echoing layer show a subsiding motion from ∼0050 until ∼0105 LT. At 0130 LT, a minimum of reflectivity is observed in layer B between 7.2 and 8.0 km corresponding to the absence of protruding lobes and a higher altitude of the cloud base. The most prominent protuberances are observed around 0140 LT and coincide very well with the position of echo power maxima (a good correspondence is also observed around 2352 LT).

The corresponding expanded views of height–time cross sections of vertical air velocity and are shown in Fig. 7. The absence of detectable vertical velocity variations through the cloud interface confirms that the measurements of vertical air velocities are not affected by the fall speed of ice particles. The rapid oscillatory vertical wind disturbances pointed out in section 3a are clearly found in the cloudy protuberances and from the cloud base down to the altitude of 6 km (Fig. 7a). On the contrary, these disturbances do not penetrate deeply within the cirrus cloud (see also Fig. 8). Values of are also enhanced in the cloud-base environment and maximum when the protuberances are observed.

The correspondence between the cloud protuberances and the oscillations of the vertical wind can be more easily identified from the expanded view shown in Fig. 8a. In particular, from ∼0130 LT, the lobes deeply penetrated into the turbulent layer and suddenly disappeared after ∼0140 LT. The downward-penetrating lobes are associated with downward air motions while updrafts of similar intensity as the downdrafts are observed between the lobes and might be associated with dry air intrusions into the cloud base. In particular, the minimum of reflectivity between 7.2 and 8.0 km around 0130 LT (Fig. 6) is associated with updrafts indicating a less turbulent and drier flow from the subcloud layer. The bottom panel of Fig. 8 shows the expanded view of height–time cross section of . Maxima of are observed within and in the very close environment of the protuberances (about 1 m2 s−2 around 0133 LT, 0.8 m2 s−2 around 0137 LT, and 0.6 m2 s−2 around 0139 LT).

These observations thus indicate that 1) the vertical wind oscillations observed from the cloud base down to about 2 km below can find their origin at or just above the cloud base, and 2) that the region including the cloud base and the protuberances were turbulent from 2300 to 0000 LT and from 0110 to 0150 LT.

4. Discussion

a. Convective overturning and cloud-base protuberances

Radar and balloon data yielded an assessment of the dynamical conditions met in the environment of protuberances detected at the cirrus cloud base by the RMR lidar. In particular, we observed 1) a weak wind shear in the altitude range of turbulent layers A, B, and C; 2) a cirrus layer at their top (only assumed for layer C due to the lack of RMR data); and 3) a dry and weakly stable air in the subcloud layer. These observations are conducive to an interpretation of the turbulent layers in terms of convective instability generated by cooling due to sublimation of ice at the cloud base. The coherent structures in the reflectivity patterns can be due to some sort of convective cells (the reflectivity patterns differ from those classically produced by gravity waves and KH instabilities) and the observed downdrafts and updrafts are likely to be a signature of convective currents spawn by the convective instability. The downdrafts may have pushed down the cloudy air and the updrafts may have pushed up the clear and dry air to form the cloud-base protuberances.

b. Possible mechanisms for convective currents underneath clouds

1) Sublimation at cloud base

Several mechanisms are likely to produce convective instability at a cloud base. Cooling by sublimation can be one of these mechanisms. Sublimation can occur when ice crystals or snow aggregates precipitate into the underlying unsaturated air [subcloud evaporation–sublimation mechanism; Schultz et al. (2006) and references therein]. When the particles begin to sublime, they cool the unsaturated layer, steepening the lapse rate. A statically unstable layer can be generated if the cooling is enough so that the clear air becomes negatively buoyant. It can easily be generated if the dry subcloud layer is weakly stable. The drier the air parcels are, the stronger the evaporative cooling and thus the deeper the local instability (e.g., Harris 1977). As noted by Schultz et al. (2006), assuming a fall speed of 1 m s−1 into a layer with 60% relative humidity (with respect to water saturation), the evaporative cooling is of the order of 1 K h−1. Much smaller relative humidity would give rise to faster cooling. The presence of a very dry air underneath the cloud (Fig. 2a) and, to some extent, the reflectivity minima observed underneath the cloud base prior to the downward growth of the turbulent layers are strong indications that sublimation was possible. Even though the temperature profile does not reveal a superadiabatic gradient below the temperature inversion, which could infer the generation of a convective instability, the temperature gradient is neutral just underneath the cloud base between 7.9 and 8.2 km (insert of Fig. 2a). The persistence of turbulent layer C for about 90 min is also an indication that the cooling of the air just underneath the cloud base was likely sufficient for sustaining the superadiabatic lapse rate and then the convective instability.

Precipitation may not be the sole mechanism responsible for sublimation. Mixing of the dry and cloudy air at the cloud base is an alternative mechanism. Mixing can result from turbulence spawn by KH instability or by cloud-base detrainment instability (CDI). Proposed by Emanuel (1981) in analogy with the cloud-top entrainment instability (CEI; Deardorff, 1980), CDI differs from the subcloud evaporation–sublimation mechanism in that the ice crystals or droplets are introduced into the dry layer by spontaneous mixing at the cloud base. By mixing subsaturated and cloudy airs, sublimation takes place, air parcels become negatively buoyant and a convective instability is generated much the same way as the subcloud evaporation mechanism. CDI is relevant when cloudy air is advected over clear air because this condition is favorable to the criterion for CDI [i.e., a decrease of the liquid (ice) water static energy across the cloud base]. From a practical point of view, the vertical gradient of ice (liquid) water virtual potential temperature θ should be negative at the cloudy interface. Ice–liquid water potential temperature was introduced by Tripoli and Cotton (1981). Expressions for θ can be found in Emanuel (1994) or in Kanak et al. [2008, their expression (2)]. It is mainly function of virtual temperature, water vapor, ice and snow aggregate mixing ratios, and latent heat of sublimation. Using an idealized distribution of ice crystal ratio of 1 g kg−1 and snow aggregate of 0.75 g kg−1 relevant in the saturated layer from the bottom of cirrus clouds (Kanak et al. 2008), we found that the CDI criterion (/dz < 0) can be satisfied at the cloud base where /dz > 0. CDI can thus be an alternative to subcloud sublimation mechanism and KH instability. Kanak et al. (2008) showed CDI was a necessary condition for the mammatus formation.

2) Others

The cloud-base subsidence mechanism can also be a source of convective instability underneath clouds (e.g., Ludlam and Scorer 1953). When the air draws down, the saturated air warms at the pseudoadiabatic lapse rate Γs while the underlying unsaturated air warms at the dry-adiabatic lapse rate (Γd > Γs). Thus, this mechanism is able to transform a statically stable layer into an unstable layer so that a convective instability is generated. Therefore, this convective instability mechanism is also likely to produce turbulence as observed by the MU radar. However, this mechanism does not seem to be viable because cloud subsidence was not observed in the present case.

Radiative cooling at cloud top can generate instabilities and could result in deep vertical motions within the cloud. The observed protuberances, if mamma, would be a visible signature of the penetration of downdrafts at the cloud base. Schultz et al. (2006) showed that radiative cooling can have a similar or even stronger effect than evaporative cooling. Such a mechanism is likely not relevant in the present case because there is no signature of deep vertical motions within the cloud. However, radiative destabilization can also occur near the cloud bottom when terrestrial longwave radiation cannot penetrate deeply within optically thick clouds. Garrett et al. (2005) showed that radiative heating and destabilization can occur within a layer of 100 m of the cloud base in an optically dense thunderstorm anvil cirrus. Because it is concentrated in a too thin layer, this source of instability cannot be discounted in our case.

c. Did the lidar detect cirrus mamma?

The protuberances detected by the RMR lidar might be either mamma, virga, or wave clouds. According to the Glossary of Meteorology, virga (also called fall streaks) are defined as streaks of rain or ice crystals that evaporate or sublimate before reaching the ground as precipitation (Glickman 2000). They can exhibit a hooked form underneath precipitation source (mainly due to wind shear) or nearly vertical if the wind shear underneath the cloud is weak. Schultz et al. (2006) discussed the possible relationships between virga and mamma. These structures are sometimes observed together and virga could form from mamma (e.g., Clarke 1962). Therefore, in some occasions, mechanisms giving the impetus for virga and mamma formation might be the same.

The structures detected by the RMR lidar resemble those reported by Wang and Sassen (2006, their Fig. 11), with a cloud polarization lidar. Wang and Sassen (2006) reported vertical and horizontal scales of protuberances (interpreted as cirrus mamma) from 0.3 to 1.1 km and from 0.5 to 8.0 km, respectively. The scales of the structures observed by the RMR lidar are about 0.2–0.8 km vertically and 1.0–4.0 km horizontally (see section 3c) and are thus consistent with those given by Wang and Sassen (2006) for cirrus mamma.

The MU radar observations clearly revealed that the protuberances at the cloud base are associated with downward air motions and upward air motions between them. A similar correspondence between mamma and vertical motions was reported using millimeter wavelength radars (e.g., the review by Schultz et al. 2006). Kanak et al. (2008) reviewed estimates of vertical velocities (of the hydrometeors with respect to the air motions) in mammatus lobe areas published until then. A negative correlation is usually found between vertical velocity and reflectivity: large downdrafts (−2.0 to −6.0 m s−1) in mammatus lobe areas (associated with high reflectivities) and weak updrafts (0–1.5 m s−1) in the drier air penetrating within the clouds (usually associated with low reflectivities). Because few estimates for cirrus mamma are available, the values reviewed by Kanak et al. (2008) are not necessarily representative of the vertical velocities in cirrus mammatus areas. Also, because of the small hydrometeor size in cirrus, fall velocities cannot likely exceed a few tens of centimeters per second (e.g., Schultz et al. 2006). Only Wang and Sassen (2006) reported vertical velocities for cirrus mamma in the same altitude range as in the present work. Their Fig. 6 indicates the presence of downward and upward motions of the order of +/−0.7 m s−1. Vertical velocities estimated from the MU radar are of the same order (Fig. 7) suggesting that the ice particles could fall at the same speed as the air draws down. Kollias et al. (2005) reported high-resolution measurements of cirrus mamma with a 94 GHz radar. For one mammatus lobe at least, they measured mean vertical Doppler velocities from −1.0 down to −5.0 m s−1 within 0.7 km of descent without noticeable change in reflectivity. Such (negative) velocities are larger than the expected terminal fall velocities of ice crystals and thus suggest the presence of strong clear-air downdrafts. As a matter of fact, if mamma were indeed observed in the present work, clear-air motions should play a fundamental role for cirrus mamma formation and a nearly quiescent air around the mamma, as suggested by Martner (1995), cannot be assumed.

As proven by the variance enhancements of Doppler spectrum, the RMR lidar protuberances are embedded in turbulence. Lynch et al. (2001) showed an impressive image of a mammatus lobe at a cirrus fibratus cloud base with a cloud polarization lidar at ultrahigh resolution (1.5 m by 0.1 s). Cirrus mamma are ragged by turbulent eddies of various scales, consistent with our observation of turbulence. Wang and Sassen (2006) made a spectral analysis of Doppler velocities measured by a W-band (94 GHz) radar in the cloud-base region with mamma and showed the presence of Kolmogorov-type turbulence at least in the lower part of the cirrus. Figure 9b shows the power spectral density of the mean vertical Doppler velocity between 6.57 and 7.92 km (Fig. 9a) in layer C. The spectral slope is about −3 at time periods of 50–300 s (i.e., at horizontal ranges of 1.0–6.0 km for an average horizontal wind of 20 m s−1). Kollias et al. (2005) also showed power density spectra of vertical velocities measured with a 94-GHz radar in a cirrus where mamma were also present (their Fig. 3). The authors also observed a −3 spectral slope, not only in the surrounding of the mamma but also in all heights within the cirrus. They interpreted this observation in terms of gravity waves and argued that initiating processes conducive to mamma are deep in the cloud and can be related to gravity waves (even if the scales of the gravity waves were not consistent with the size of the mammatus). Schultz et al. (2006) suggested that gravity waves can contribute to the distribution in bands of the mamma rather than to their formation. The presence of the gravity wave could suggest that the protuberances are rather signatures of wave clouds. However, as already pointed out, the morphology of the radar echoes is not compatible with the interpretation in terms of gravity or KH waves (see examples in Gossard 1990).

5. Conclusions

In the present work, we described observations of clear-air turbulence with the MU radar in the environment of a cirrus cloud base at the altitude of 8.0 km. Turbulent layers up to 2.0 km in depth associated with downdrafts and updrafts up to +/−1.5 m s−1 have been monitored for about 10, 30 (when the RMR lidar was operated), and about 90 min. These vertical wind oscillations were associated with reflectivity and spectral width (or variance of the Doppler spectrum) enhancements. Owing to the high time and range resolutions of the MU radar, it could be observed that the clear-air echo layers were composed of coherent structures (especially during the third episode) deepening in time from the cirrus cloud base. They are very likely a signature of convective currents resulting from a convective instability. To our knowledge, such a convective instability underneath clouds had never been investigated from VHF ST radar observations. Further studies with the MU radar in high-resolution mode and complementary instruments such as instrumented balloons, lidar, and weather radars will be carried out for a better understanding of this process and a better knowledge of its occurrence and impacts in the troposphere.

In addition, the RMR lidar disclosed downward cloudy protuberances at the base of the cirrus. These structures can be vertical fall streaks or perhaps more likely mamma as already reported in the literature from lidar observations by Wang and Sassen (2006). Their occurrence coincides with the observed downward air motions measured by the MU radar.

Cooling by sublimation of ice crystals and/or snow aggregates into the subcloud bayer was likely the cause of the convective instability. As a result of the cooling, the air parcel becomes negatively buoyant giving the impetus to convective overturns that, in turn, could be responsible for the protuberances observed by the RMR lidar. The downward air motions push down the cloudy air and the upward air motions push up the clear and dry air to form the downward-protruding lobes of cloudy air. It is not clear, however, if the proposed mechanism is enough for a deep penetration of the protuberances or if an additional mechanism helps to their penetration. Wave clouds are likely not to be the cause of the observed cloudy protuberances. Wave clouds develop in stable conditions and would produce typical signatures in the radar echoes that are different from those observed by the MU radar.

Acknowledgments

We are grateful to Professor Henri Sauvageot and Doctor Bernard Campistron for their useful comments on this work. Special thanks are given to Prof. Schultz who also provided us many suggestions on this manuscript. We also thank Prof. Straka and two other reviewers for their comments. The MU radar is operated by and belongs to Kyoto University. The work provided by the main author was partly supported by Kyoto University.

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Fig. 1.
Fig. 1.

(a) Time–height cross section of radar reflectivity (dB) after Capon processing and after correcting the range attenuation effects (i.e., power multiplied by r2) measured by the VHF MU radar in vertical incidence from 2200 LT 7 Jun 2008 to 0715 LT 8 Jun 2008 between 1.32- and 18-km altitudes. The nearly vertical thin white lines above 10 km are airplane echoes. The black line shows the balloon altitude vs time launched at 2357 LT 7 Jun 2006. The letters A, B, and C denote the echoing layers discussed in the text. (b) (top) the corresponding time–height cross section of vertical velocity of air motions measured by the VHF MU radar. (bottom) The corresponding time–height cross section of variance of the Doppler spectrum peaks using the vertical beam and after correcting from the beam-broadening effect. Values are not given for SNR < −5 dB.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 2.
Fig. 2.

(a) Vertical profiles of air (black) and dewpoint (gray) temperature obtained from radiosonde data at 2357 LT 7 Jun 2006 released from Shigaraki MU Observatory. The inset shows a closeup of the temperature profile between 7.4 and 8.8 km. (b) (left) Vertical profiles of dry squared Brunt–Vaïsälä frequency (multiplied by 1000, gray line) and Richardson number (Ri; black line). The vertical dashed line shows the value of Ri = 0.25. (right) The corresponding profiles of wind speed (m s−1, dashed black line), wind shear (m s−1 km−1, solid black line), and wind direction (°, solid gray line) shifted by −90°.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 3.
Fig. 3.

(top) Time–height cross section of horizontal wind speed (m s−1), (middle) wind direction (°), and (bottom) wind shear (m s−1 km−1) from 0430 to 0630 LT 8 Jun 2006. The black lines show the 28-dB contour of radar reflectivity shown in Fig. 1a.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 4.
Fig. 4.

(top) Time–height cross section of RMR lidar backscatter ratio (dB) from 2206 LT 7 Jun 2006 to 0236 LT 8 Jun 2006 showing cirrus clouds (and layers of aerosols mainly below 3 km). Optical attenuation was strong enough to cause range-limiting effects at some periods and altitudes. (bottom) The inset in the top is expanded: white dashed box in the top indicates inset in the bottom.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 5.
Fig. 5.

Infrared (IR1, 10.3–11.3 μm) images of the geostationary meteorological satellites showing high-level clouds, and then likely cirrus, over the southwest of Japan at 0200 LT 8 Jun 2006. The dot shows the MU radar location.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 6.
Fig. 6.

Time–height cross section of radar reflectivity (dB) after correcting the range attenuation effects for the period of lidar observations and between 6- and 9-km altitudes using (top) the vertical beam and (bottom) the oblique (180°, 10°) beam. The black contour shows the cirrus cloud edges arbitrarily defined as the 9.54-dB level of the backscatter ratio shown in Fig. 4.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 7.
Fig. 7.

(top) Time–height cross section of vertical velocity (m s−1) for the period of lidar observations and between 6- and 9-km altitudes MSL. (bottom) The corresponding variance of Doppler spectrum (m2 s−2) at the range resolution of 150 m and after correcting the beam-broadening effects. The black contours show the cirrus edges arbitrarily defined in the same way as in Fig. 6.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 8.
Fig. 8.

Expanded views of the time–height cross section of (top) vertical wind (m s−1) and (bottom) from 0055 to 0200 LT 8 Jun 2006 between 6.0 and 9.0 km. The black line contour shows the cirrus edges arbitrarily defined as in Fig. 6.

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

Fig. 9.
Fig. 9.

(a) Time series of vertical velocities averaged between 6.57 and 7.92 km in the altitude range of turbulent layer C. (b) The corresponding power spectral density. A −3 slope line is also shown as a reference (gray line).

Citation: Monthly Weather Review 138, 2; 10.1175/2009MWR2927.1

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