3.3.1. W and VZ
 Differences in W, VZ, P, and δ between cases A and B are described. Upward W above 6.2 km MSL (greater than 0.2 m s−1) in case B was greater than that in case A, which was less than 0.1 m s−1 (Figures 4a and 8a). The disturbances of W above 6.2 km MSL in case B (dominantly greater 0.2 m s−1) was greater than that in case A (∼0.05 m s−1). Latent heat release resulting from deposition growth is the most plausible cause that produced the moderate upward motion in case A [Nishi et al., 2007]. On the other hand, the greater upward W and W disturbance in case B indicates that all the mechanisms that are able to produce upward W in stratiform region (i.e., latent heat release resulting from deposition growth, gravity waves, old cells which had origin in the convective region, and latent heat release resulting from riming) are candidates that produced the observed upward W. It is noted that riming was limited to the altitude below ∼7.6 km MSL, where temperature was greater than 258 K (i.e., −15°C) (see Figure 1) and hence riming can be observed [Woods et al., 2005].
 VZ at 7.0 km MSL in case B (1.3 m s−1) was greater than in case A (1.0 m s−1). Because D0 estimated at the bottom of ML in case B was ∼2.8 times greater than in case A (described later), the greater VZ at 7.0 km MSL in case B indicates that hydrometeors with larger size existed in case B, though fall velocity of ice hydrometeors depends on their shape and density [Heymsfield et al., 2002]. Because both vapor condensation occurring in stratiform region and riming in convective cells which have typical updrafts of several m s−1or more significantly contribute to the size growth of ice-phased hydrometeors in the upper part of stratiform region [Houze, 1993], they are candidates that produced larger hydrometeors size in case B. Let us mention again that because the values of VZ+air above 7.0 km MSL in cases A and B were nearly the same (see section 3.2), detailed considerations on hydrometeor fall velocity cannot be made if accurate W measurement was not available.
 The increase of VZ with decreasing altitude was observed in the altitude range above 4.9 km MSL, where temperature was below 0°C (Figures 4b and 8b). Above ∼6.0 km MSL, where upward W was observed, both vapor condensation and aggregation are candidates which produced the increase in hydrometeor size. Though it is possible that riming played a role in the size growth of hydrometeors, the contribution of riming was limited to the altitude below 7.6 km MSL, where temperature was greater than 258 K (i.e., −15°C).
 The increase of VZ was seen even in the altitude range 4.9–6.0 km MSL, where upward W was not observed. In case A, VZ was ∼1.2 m s−1 around 6.0 km MSL and ∼1.3 m s−1 around 5.1 km MSL (Figure 4b). In case B, VZ was ∼1.4 m s−1 around 6.2 km MSL and ∼1.8 m s−1 around 5.1 km MSL (Figure 8b). Hereafter it is explained that aggregation is the most plausible cause that produced the increase of VZ in the altitude range 4.9–6.0 km MSL. In a 1.0–1.5 km thick layer lying just above the 0°C level, riming is an important process that contributes to size growth of hydrometeors [Houze, 1993]. Riming is closely related to upward W because freezing of liquid drops produces upward W through latent heat release and upward W is necessary for maintaining water saturation [Houze, 1993]. However, because upward W in the altitude range ∼4.9–6.0 km MSL was not observed both in cases A and B (Figures 4a and 8a), it is concluded that riming was not active or absent. Though aggregation is also the process that contributes to size growth of hydrometeors, it does not produce heat. Temperature in the altitude range ∼4.9–6.0 km MSL was 267–273 K (i.e., −6°C–0°C) (Figure 1), and occurrence frequency of aggregation becomes much greater for temperature above ∼−5°C [see Houze, 1993, section 3.2.4; Pruppacher and Klett, 1997]. In situ measurements of hydrometeors using aircrafts have confirmed the hydrometeor size growth by aggregation in the ∼1 km thick layer just above the 0°C level [McFarquhar et al., 2007; Stewart et al., 1984; Willis and Heymsfield, 1989].
 The increase of VZ in the altitude range 4.9–6.0 km MSL was greater in case B than in case A (Figures 4b and 8b). In addition to the temperature above ∼−5°C, the presence of planar dendritic crystals or spatial dendrites is necessary for producing aggregation [Rauber, 1987]. The growth of arms of dendritic crystals becomes entangled between −10 and −16°C [Houze, 1993, p. 86], and the altitudes with temperature between −10 and −16°C (i.e., between 263 and 257 K) was 6.7–7.7 km MSL (Figure 1). Because the greater upward W in the altitude range 6.7–7.7 km MSL in case B provided more preferable condition for entangled arm growth of dendritic crystals than in case A, it is concluded that the greater upward W in case B contributed to the greater increase of VZin the altitude range 4.9–6.0 km MSL through enhanced occurrence of aggregation. For further investigation, simultaneous in situ measurement of hydrometeors and 50 MHz wind-profiling radar are necessary. However, it is stressed that accurate measurement ofW provides the opportunity to discuss the microphysical processes which relate to stratiform precipitation in detail. In case A, δ at 7.0 km was 0.015, increased with decreasing altitude, and was 0.06 at 5.0 km MSL (Figure 4d). Though the disturbance observed in δ was significant probably due to returns which produced small δ (i.e., specular reflection from crystal's plane face and/or horizontally oriented plates), this increase in δ indicates the increase of nonsphericity by vapor condensation and aggregation.
 By using the model of raindrop fall velocity and assuming DSD, median volume diameter of raindrops (hereafter D0) just below the ML was estimated. VZ was calculated using the relation
where D is the raindrop diameter in m, N(D) is DSD, σ(D) is backscattering cross section of raindrops, and Vt(D) is the terminal fall velocity in m s−1. In the measurement, W and VZ+air are spatially weighted by the antenna beam pattern [Fang and Doviak, 2008]. Because the two-way half-power full width of a radar beam and vertical resolution were 2.4° and 150 m, respectively, the EAR observed ∼260 m × 150 m volume at 7.0 km MSL. However, the horizontal extent of radar beam was small compared to the advection scale. For the sampling time length of 131 s, advection scale of a scatterer is greater than the horizontal extent of radar beam even for the small horizontal wind velocity of 2 m s−1. Inhomogeneity of clear air reflectivity can occur by shear instability, and the reflectivity inhomogeneity can produce misestimation of W by contaminating Doppler velocity measured by the vertical radar beam with horizontal wind velocity [e.g., Yamamoto et al., 2003]. However, the possible misestimation of W is negligible in the study because horizontal wind velocity was small (i.e., mostly less than 10 m s−1). Under the condition that water drop is spherical and D is small compared to the radar wavelength λ (i.e., D ≤ λ/16), σ(D) is expressed by
where Km = (m2 − 1)/(m2 + 2) and m is the complex refractive index of water. Equation (2) is called the Rayleigh approximation. We used the Rayleigh approximation because the condition D ≤ λ/16 is satisfied for wind-profiling radars which are operated below 3 GHz and hence the Rayleigh approximation has been widely used for raindrop measurements by the wind-profiling radars [Renggono et al., 2006; Schafer et al., 2002, Williams, 2002].
 The raindrop fall velocity is given by
where Vt(D) is the terminal fall velocity in m s−1, ρ is the air density, and ρ0 is the air density at a pressure of 1013 hPa and temperature of 293.15 K. For details of modeling for raindrop fall velocity, see section 8.2 of Doviak and Zrnić . ρwas determined by pressure and temperature measured by the radiosonde soundings at the observatory. The Marshall-Palmer distribution is given by
where N0 = 8 × 103 m−3 mm−1, Λ = 4100 × R−0.21 m−1, and R is the rain rate in mm h−1. For details of the Marshall-Palmer distribution, see section 8.1.2 ofDoviak and Zrnić . Because only Λ (i.e., R) is a variable and the relation between Λ and D0 is given by , D0 is able to be related to VZ by combining equations (1) to (4). The rainfall amount measured by the surface rain gauge from 19:40 to 22:00 LST on 16 December was 6.8 mm, and that estimated at the lowest altitude that the EAR measured (i.e., 2.4 km MSL) were 6.5 mm. The agreement in rainfall amount between the surface rain gauge and the EAR indicates that the Marshall-Palmer distribution was useful for estimatingD0. For VZ of 3.7 m s−1 at 4.5 km MSL (i.e., case A), D0 was 0.4 mm. For VZ of 7.6 m s−1 at 3.9 km MSL (i.e., case B), D0 was 1.1 mm. Though VZ of raindrops had disturbances and the assumption was used for DSD, the estimation result suggests that D0 at the bottom of ML in case B was ∼2.8 times greater than in case A. The greater raindrop size in case B is consistent with the greater VZ above the ML in case B. The uncertainty of D0 was evaluated by assuming the modified gamma distribution (i.e., N(D) = N0Dμ exp(−ΛD) where μ is the shape parameter; see Ulbrich  for details). μ was varied from 0 to 3 because Tokay and Short  showed that μ was generally less than 3 for very light, light, and moderate rains which have rainfall rate less than 5 mm h−1. D0 is given by . Note that the value of N0 did not affect the computation (see equation (1)). For μ of 0–3, D0 for cases A and B varied 0.4–0.5 mm and 1.1–1.4 mm, respectively.
3.3.2. Melting Layer Thickness
 As described in sections 3.1 and 3.2, the thickness of ML in case B (900 m) was greater than in case A (300 m). Most of ice-phased hydrometeors melt rapidly as they reach to the 0°C level and only large-sized aggregates determine the thickness of melting layer [Willis and Heymsfield, 1989]. Further, the estimation using VZ at the bottom of ML suggests that D0 at the bottom of ML in case B was ∼2.8 times greater than in case A. Both the thickness of ML and the VZ value at the bottom of ML indicate that hydrometeors with greater size existed in case B. The results support the idea that the size growth of hydrometeors in convective cells by riming and that in stratiform region by condensation above ∼6.0 km MSL and aggregation at 4.9–6.0 km MSL (see section 3.3.1) contributed to the formation of thicker ML in case B. Though δ above the ML was not able to be measured in case B, the peak value of δ in the ML in case B (0.25) was greater than in case A (0.18) and δ in case B was greater than 0.2 over the 400 m altitude range (see Figures 4d and 8d). This result indicates that the nonsphericity of ice hydrometeors was greater in case B, and supports the conclusion that the greater upward W in the altitude range 6.7–7.7 km MSL in case B provided more preferable condition for the entangled arm growth of dendritic crystals than in case A, and contributed to the greater size increase of hydrometeors in the altitude range 4.9–6.0 km MSL through enhanced aggregation.
3.3.3. Volume δ of Raindrops
 Volume δ of raindrops (i.e., below 4.0 km MSL) was 0.08–0.10 in case B, though it was close to zero in case A (see Figures 4d and 8d). Using a three-dimensional polarization-dependent ray-tracing algorithm,Roy and Bissonnette  demonstrated that static and dynamical raindrop deformation from the sphere (i.e., steady state flattening and oscillation as a resonant response to eddy shedding) can produce strong increase of δfor single raindrop at off-zenith angles. However, their calculation results also showed that the increase ofδ at the vertical direction (i.e., the direction of the lidar laser beam) does not show significant increase for the raindrop models that are consistent with the measurement results of the scanning lidar [see Roy and Bissonnette, 2001, Figures 16 and 17]. Therefore δ for single raindrop is not likely a cause for the increase of δ in case B. Multiple scattering can generate the increase of volume δunder the presence of large-sized raindrops [Tatarov and Kolev, 2001]. Because scattering by single raindrop is not limited to the backward direction, multiple scattering can increase volume δ even when the lidar laser beam points to the vertical direction. Other lidar measurements showed that the volume δ (i.e., δ including the effects by multiple scattering) for raindrops at the vertical direction can increase ∼0.1 or greater for the FOV as small as 1 mrad [see Roy and Bissonnette, 2001, Figures 6–9]. Because our lidar measurement used the FOV similar to the work of Roy and Bissonnette (i.e., 0.5 mrad), stronger multiple scattering triggered by large-sized raindrops is likely a cause that generated the increase of volumeδ in case B.
 Effects of turbulence on raindrop oscillations need to be further investigated [Szakáll et al., 2010]. Therefore it is worth noting that wind disturbance is able to increase δ of single raindrop at the vertical incidence by distorting the orientation of raindrops from the horizontal. Standard deviation of W below the ML altitude was greater than 0.1 m s−1 in case B, while it was less than 0.1 m s−1 in case A. Altitude variation of W in case B was also greater than in case A (Figures 4a and 8a), and hence it is speculated that the greater disturbance of W in case B caused the greater δ in case B. Further theoretical end experimental studies are necessary to prove our speculation.
 Another interesting feature observed in δ is a dip around 4.0 km MSL (i.e., the bottom of ML) in case B (see Figure 8d). Most of ice-phased hydrometeors melt rapidly as they reach to the 0°C level and only large-sized aggregates contribute to the ML formation [Willis and Heymsfield, 1989]. The large-sized aggregate disrupts into small-sized particles at the final stage of melting, and the disrupted small-sized particles display round, plate, or lens-like shapes [Fujiyoshi, 1986; Knight, 1979]. The decrease of nonsphericity at the final stage of melting explains the decrease of volume δ at the bottom of ML than that at the altitudes above 4.0 km MSL. We speculate that the stronger multiple scattering effect for raindrops explains the greater volume δ below 4.0 km MSL. By producing stronger scattering to various directions, surface wave of raindrops can cause stronger multiple scattering than melting particles. The presence of lidar dark band, which indicates the stronger raindrop backscattering than melting particles, supports the speculation (see Figures 4c and 8c). Unfortunately there are no theoretical and experimental studies that treated the effects of multiple scattering on melting particles. Further theoretical and experimental studies are required for thorough understanding of δ both for raindrops and melting particles.