3.1. Event Overview

Typhoon Soudelor (2015) made landfall in Putian City, Fujian province at 1400 UTC on August 8, 2015, as a severe tropical typhoon (Table 1). The spiral rainband, distributed in a southeast-to-northwest direction, remained stationary at the northeast quadrant of the typhoon from east Wenzhou to Taizhou and the Siming Mountains. This phenomenon, as referred to the “echo training” feature, persisted for ~11 h. During this period, over 108 meteorological stations in Eastern Zhejiang province recorded accumulated rainfall exceeding 100 mm. The rainfall center manifested in Taizhou City, where the total rainfall exceeded 300 mm, and the highest 1-h (3-h) rainfall was 84.9 mm and 205.9 mm, respectively (Figure 2a).

Typhoon Fitow (2013) made landfall in Fuding City, Fujian province, as a severe typhoon at 1700 on October 6, 2013. Concurrently, Typhoon Danas (2013) reached the eastern East China Sea. The linear mesoscale convective rainband between these two typhoons exhibited a strengthening trend after making landfall, resulting in persistent rainfall for ~12 h. The rainfall amounts at 117 stations exceeded 100 mm, with Taizhou experiencing the highest amount of 252.0 mm (Figure 2b).

Although both “echo training” processes have produced extreme rainfall in East Zhejiang, there were notable differences in the organizational and developmental modes between them. Figure 3 shows the distribution of composite reflectivity and the direction of convection cell movement during the two processes. During Process 1, the convective rainband exhibited an SSE‒NNW trend of movement and was comprised of convective cells arranged in rows, with new convective cell motion parallel to the rainband. In contrast, during Process 2, the convective rainband exhibited an south–north trend as new convective cells merged into the eastern boundary of the rainband. The moving direction of the convective cells was nearly perpendicular to the stagnant rainband.

3.2. Comparison of Largescale Circulations and Synoptic Conditions

Figure 4 illustrates that the long axis of the “echo training” rainbands was approximately aligned with the ambient 500 hPa wind. The rainband of Process 1 developed within the expansive southerly flow situated between the 580 isobar of Typhoon Soudelor and the 588 isobar of the subtropical high. The wind speed over the area of extreme rainfall reached 22 m s⁻¹. In contrast, the rainband during Process 2 was located between Typhoon Fitow and Typhoon Danas, with a relatively weak 500 hPa wind (ranging from 4 to 6 m s⁻¹).

Additionally, discrepancies are observed in the water vapor transport pathway (Figure 5a,b). The strong water vapor flux at low latitudes reaches ~28 × 10⁻³ g s⁻¹ cm⁻¹ hPa⁻¹ during Process 1, while the westward transport of moisture was lower than 20 × 10⁻³ g s⁻¹ cm⁻¹ hPa⁻¹ during Process 2. However, the uplift and condensation of water vapor during these two processes occurred between the coastline and the windward slope of inland hills (Figure 5c,d). Meanwhile, the convergence of the water vapor flux during both processes was concentrated in the boundary layer below 925 hPa, with the maximum value exceeding 3.0 × 10⁻⁶ g cm⁻² hPa⁻¹ for Process 1 and 1.5 × 10⁻⁶ g cm⁻² hPa⁻¹ for Process 2, respectively (Figure 5e,f).

The 72-h backward trajectory (every 6-h) of water vapor sources during the two processes was further investigated using Lagrange tracking with the HYSPLIT4.8 model (Figure 6). The initial altitudes were 500 m, 1000 m, and 1500 m, respectively. As shown in Figure 6a, the water vapor in Process 1 originated from the South China Sea and the East Philippine Ocean, whereas that of Process 2 was mostly transported from the northwest Pacific Ocean.

The soundings observed at Hongjia station indicated the presence of a conditionally unstable boundary during the two processes (Figure 7). In particular, the troposphere was nearly fully saturated during Process 1. The southeasterly jet extended from 925–500 hPa, with wind speeds exceeding 20 m s⁻¹ between 850 and 400 hPa. The wind speed difference between the ground and 850 hPa was 12 m s⁻¹, indicating substantial vertical wind shear during this process. Moreover, the convective available potential energy (CAPE) value was 1698 J kg⁻¹, with a negligible CIN (8.4 J kg⁻¹). In Process 2, the near-saturated layer was thinner (below 600 hPa), and the difference between temperature and dew point exceeded 20 K above 600 hPa. The wind speed below 850 hPa did not surpass 8 m s⁻¹, with a moderate vertical wind shear between the ground level and 700 hPa. In this instance, the CAPE was 742 J kg⁻¹.

3.3. Comparison of Meso- and Micro-Scales Thermal and Dynamical Characteristics

With regard to Process 1, Typhoon Soudelor resulted in rainfall over 100 mm within a 24-h period in the hilly areas situated to the west of 121.2°E. A cold pool was formed in this area, with surface temperatures remaining below 26°C. In contrast, surface temperatures to the east of 121.2°E increased to ~30°C due to the sustained influence of warm and moist air from the southeast. The temperature distribution indicates the emergence of a mesoscale temperature front zone situated over the cold pool region (Figure 8a). The scattered weak convection exhibited gradual linear organization characteristics, culminating in the formation of a linear mesoscale rainband trending in a southeast–northwest direction.

Subsequently, the mesoscale frontal zone was quasistationary and the temperature anomalies near the surface intensified (Figure 8b). The cold pool rapidly intensified as the temperature differential between the two flanks of the convective rainband reached 4°C. The southeasterly flow in the coastal region reached 15.6 ms⁻¹. Consequently, a meso-γ scale convergence center was formed at the east edge of the rainfall area (Figure 8e), exhibiting a convergence rate in excess of 10 × 10⁻⁴ s⁻¹. The rainband was well-organized to form an “echo training” process. Convective cells with reflectivity factors exceeding 50 dBZ proceeded in a northward in sequence, resulting in heavy rainfall exceeding 200 mm in Taizhou City. By 1800 UTC, the linear structure of the mesoscale rainband had dissipated (Figure 8c).

The aforementioned observations illustrate that the interaction between the southeasterly flow and the inland cold pool (caused by prolonged heavy rainfall) is conducive to the formation of a mesoscale front zone. This further promotes the lifting of warm and moist air by the solenoidal term. The positive and negative divergence pairs generated by lateral friction under the hilly area of Eastern Zhejiang serve to reinforce the organization of the convective rainband. In this regard, the southeasterly jet represents a fundamental condition for the formation of the mesoscale frontal zone and convergence.

For Process 2, the surface temperature in the coastal region of Zhejiang increased to 28°C−30°C, coinciding with a cessation of rainfall as Typhoon Fitow weakened. However, the hilly area (west of 121.3°E) experienced rainfall for approximately 3 h longer than the surrounding regions, and the near-surface temperature remained at around 24°C throughout this period. From 1100 UTC, scattered weak cells over the sea approached the land amidst the strong easterly flow. Following the landing, the scope and intensity of convection developed gradually to form an south–north stripe rainband (Figure 9a). The rainband was located in a coastal area where the disturbance temperature is higher, with a meso-γ scale convergence center was present to the east of the rainband (Figure 9d).

While the rainband remained in a quasistationary state with a relatively weak easterly air flow (6 m s⁻¹) in the eastern sector, two meso-γ scale convergence centers have developed in the rainband (Figure 9e), with a temperature anomaly of 2°C (Figure 9b). The surface convergence and mesoscale barotropism exhibited slight enhancement, though not to the same degree as in Process 1. Subsequently, the rainband retreated to the east of 121.3°E, accompanied by a reduction of convergence and temperature anomalies (Figure 9c–f). In general, Process 2 indicates that persistent rainfall can occur in the absence of a strong mesoscale barocline and convergence, due to the coordination between a continuously weak easterly flow and a weak but stagnant cold pool, resulting in strong and continuous rainfall.

Previous research suggests that the conjunction of LLJ with a warm tongue would augment the θ_se at lower levels and curtail the convection inhibition energy, thereby creating a conducive milieu for the genesis and evolution of convection [22]. The VWP production of Doppler radar is an effective method for accurately identifying the wind field over the radar coverage area [23]. Here, the vertical structure of convection during the two processes was compared using radar reflectivity factor, radial velocity structure, and VWP products (Figures 10 and 11).

At 06 : 59 UTC on August 9, 2015, the southeasterly wind was dominated below 5 km height, with a 20 m s⁻¹ jet observed at 1.2–3.7 km height. The 30 (40) dBZ echo top exceeded 5 (7) km in altitude, indicative of the involvement of ice-phase processes. At about 10 : 00 UTC, the southeasterly jet (over 20 m s⁻¹) expanded upwards to 6.1 km, in conjunction with the southwesterly flow of the cold pool, forming a deep convergence zone that extended from the ground to ~5 km in height. In these circumstances, the maximum reflectivity factor increased rapidly to 55 dBZ at 1.5–2 km. This low centroid and high-intensity feature was conducive to the formation of heavy rainfall [24]. At 18 : 05 UTC, the wind direction at 5 km altitude changed from south to southwest while the rainband dissipated.

During Process 2, a deep ESE air flow (4–8 m s⁻¹) occurred at ~5.5 km above the coastal region at 1205 UTC on October 7, 2013. Figure 11 shows that the 30 dBZ echo top can reach as high as 5 km, while the echo centroid over 45 dBZ was located at a height of 1–2 km, implying that warm rain process are dominant in the formation of heavy rainfall. Between 1600 and 1700 UTC, a nearly 50 km long north–south convergence zone was formed below 1.5 km altitude. The 30 dBZ echo top was mainly distributed at elevations below 3.2 km. The comparative wind speed on both sides of the convergence zone (ranging from 5 to 10 m s⁻¹) resulted in stagnant convergence and persistent rainfall in that area. After 2100 UTC, the ambient wind changed to southwest at 3.4 km altitude and to northwest near the surface, causing the dissipation of the convergent flow and the eastward retreat of the rainband. The analysis indicates that the convergence of the westerly wind from the cold pool and the easterly wind from the open ocean represents a pivotal factor in the genesis and sustenance of the mesoscale convective typhoon rainband.