1. Introduction
The rainfall distributions for heavy rainfall events over Taiwan under favorable large-scale settings differ from case to case due to terrain, local winds, convective feedbacks, orientation of the front, and depth of the postfrontal cold air, etc. (Li et al. 1997; Teng et al. 2000; Yeh and Chen 2002; Lin et al. 2011; Tu et al. 2014, 2017; Chen et al. 2018; Ke et al. 2019, and others). This presents a significant challenge when predicting the timing, duration, and location of heavy rainfall over the Taiwan area during the mei-yu season. The unique aspects of the 2 June 2017 case studied in this research include persistent torrential rain on the northern side of the Yang-Ming Mountains, with a rainfall accumulation of >600 mm in less than 8 h (0200–1000 LST) along the northern coast. Furthermore, around noontime (1100–1200 LST 2 June), an east–west-oriented rainband occurred over the Taipei basin and the northwestern coast, with a rainfall rate of >70 mm h−1. Torrential rains also occurred over the southwestern slopes of the Snow Mountains and the Central Mountain Range, with a rainfall accumulation of >500 mm day−1.
There are a few recent studies on the heavy rainfall event of 2 June 2017. Tu et al. (2020) studied the long-duration widespread mei-yu heavy rainfall event over the Taiwan area during 1–4 June 2017. For this event, the southwesterly marine boundary layer jet (MBLJ) (Tu et al. 2020) was present prior to the arrival of the mei-yu jet/front system. The MBLJ was associated with horizontal moisture transport from the northern South China Sea to the Taiwan area mainly within the marine boundary layer. Wang et al. (2021) applied an ensemble-based sensitivity analysis (ESA) based on 45 forecast members with a grid size of 2.5–5 km. They suggested that there are several factors important in the prediction of the frontal rainfall in their models: (i) position and speed of the surface mei-yu front and the low-level wind-shift line; (ii) moisture amount in the environment; (iii) location and timing of the mesoscale low pressure disturbances along the front; and (iv) frontal intensity. Arakane et al. (2019) studied the remote effect of a tropical cyclone (TC Mora) in the Bay of Bengal on the mei-yu front by including and excluding TC Mora in the model simulations. Chung et al. (2020) combined the Weather Research and Forecasting (WRF) Model with the Community Radiative Transfer Model (CRTM) to simulate the multichannel brightness temperatures (BTs) from the Advanced Himawari Imager (AHI) on board the Himawari-8 geostationary satellite for the mei-yu frontal system.
Our main foci are the front-terrain interactions over the northern Taiwan coast, Taipei basin, and northwestern Taiwan, as well as the impacts of these interactions on the timing, duration, and location of heavy rainfall occurrences from observational analyses and high-resolution model simulations. Past studies of orographic precipitation associated with landfalling cold fronts in different parts of the world (Bosart et al. 1973; Bond et al. 1997; Bougeault et al. 2001; Chien et al. 2001; Colle et al. 2002; Stoelinga et al. 2003; Neiman et al. 2004; Medina et al. 2005; Garvert et al. 2005a,b; Kingsmill et al. 2006; Rotunno and Houze 2007; Barrett et al. 2009; and others) were conducted during the winter season in midlatitudes in a relatively stable environment. The orographic precipitation associated with the mei-yu front over the Taiwan area occurs in a subtropical, warm, and potentially unstable monsoon flow in the early summer under the presence of steep terrain.
In this work, we would like to investigate the heavy-rain-producing mei-yu front (>70 mm h−1) along the northern Taiwan coast, northwestern Taiwan, and the subsequent rainfall maximum over the Taipei basin as the shallow (<1 km) east–west-oriented front makes landfall. We will investigate the interplay of the moisture supply, lifting mechanisms (e.g., frontal lifting and orographic lifting), depth of the postfrontal cold air, and frontal movement for the development of torrential rains. In addition to the National Centers for Environmental Prediction (NCEP) Final Operational Global Analysis (FNL) data, soundings, and surface and rain gauge observations, we also conduct high-resolution modeling to further assess the interactions between the mei-yu jet/front system and the terrain. Our investigations include the roles of the barrier jet (BJ) (Chen and Li 1995a; Li et al. 1997; Li and Chen 1998), marine boundary layer jet (MBLJ) (Tu et al. 2019), and terrain effects (Fig. 1c) on the development of heavy rainfall over northern Taiwan.
(a) The four nested domains of the WRF Model simulation with horizontal grid sizes of 27, 9, 3, and 1 km, respectively. (b) Terrain height (m). The dark red line marks the location of the northeast–southwest cross section across the South China Sea (SCS) and Taiwan Strait. The purple line marks the location of the NE–SW cross section across the SCS and the Central Mountain Range. The blue dot marks Dongsha Island. (c) Terrain height (m) of northern Taiwan. The dark red line marks the location of the northwest–southeast cross section along the Tamsui River Valley. The purple line marks the location of the northeast–southwest cross section along the Keelung River Valley. The pink line marks the location of the north–south cross section across the Yang-Ming Mountains and Snow Mountains. The Anbu (25.18°N, 121.53°E), Zhuzihu (25.16°N, 121.54°E), Tamsui (25.16°N, 121.45°E), Keelung (25.13°N, 121.74°E), Wenshan (25°N, 121.58°E), and Guishan (25.03°N, 121.39°E) stations are denoted by a red (Central Weather Bureau manned weather station) or blue (CWB automatic weather station) dot. Banciao (25°N, 121.44°E) rawinsonde site is denoted by a red triangle. Peng-Chia Yu Island (25.63°N, 122.07°E) and Linkou Plateau are marked.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
2. Literature review
During late spring and early summer, the large-scale circulations undergo two distinct changes over East Asia (He et al. 1987). The first transition in May corresponds to the onset of the Southeast Asian monsoon when rains spread over Assam, the west coast of the Malay Peninsula, Indochina, and south China. The second transition occurs about one month later, which corresponds to the onset of the Indian summer monsoon. During the second transition, the mei-yu season starts over the Yangtze River Valley, China plain, and the adjacent waters. The onset of the early summer rainy season over Taiwan (15 May–15 June) occurs during the first transition and ends once the mei-yu season begins over the China plain (Chen 1983; Chen et al. 1989; Kuo and Chen 1990).
In a prolonged mei-yu frontal system during 10–15 June 1975, the apparent temperature gradients across the surface front and the associated upper-level jet/front were not revealed by the filtered and subjectively analyzed 240-km grid data using only four vertical levels (e.g., 850-, 700-, 500-, and 400-hPa level) (Chen and Chang 1980). Hsiao and Chen (2014) revisited this case using the European Centre for Medium-Range Weather Forecast (ECMWF) Re-Analysis (ERA-40) data with a 2.5° grid spacing and 23 vertical levels (from the 1000-hPa level to the 1-hPa level). Similar to other frontal systems found during the Taiwan Area Mesoscale Experiment (TAMEX) (Chen et al. 1989; Trier et al. 1990; Chen 1993; Chen et al. 1994, 1997; and others) and recent studies (Tu et al. 2014; Chen et al. 2018; and others), this mei-yu front exhibited baroclinic characteristics: (i) appreciable temperature/moisture gradients in the lower troposphere; (ii) a marked vertical tilt; (iii) an upper-level jet/front with tropopause folding; (iv) a thermally direct circulation across the frontal zone (Hsiao and Chen 2014). A midlatitude omega blocking pattern on 10 June 1975 transformed into a Rex blocking pattern by 12 June 1975. During this period, the midlatitude blocking patterns and associated trough axis were almost stagnant. As a result, this mei-yu front was a slow mover and affected the Taiwan area for more than four days (Hsiao and Chen 2014), unlike propagating fronts observed during TAMEX (Chen 1993).
Over southern China, the mei-yu fronts have a marked northward vertical tilt (Chen et al. 1989; Chen and Hui 1990, 1992). The frontal slope is determined by the balance between Coriolis acceleration and the density difference (neglecting friction) across the front (Margules 1906). The frontal slope is maintained under an equilibrium condition between the wind and density fields (Palmen and Newton 1969). However, as the mei-yu front penetrates into the subtropics, the depth of the postfrontal cold air decreases due to a decrease in the Coriolis parameter (Chen et al. 1989). As the mei-yu fronts from southeastern China move southward, they interact with the hilly terrain along the southeastern China coast (Chen and Hui 1990) and the terrain over Taiwan (Wang 1986; Chen et al. 1989; Trier et al. 1990). The postfrontal cold air is ducted around the southeastern China coast and moves rapidly southward, a result of a pressure ridge along the southeastern China coast and strong northeasterlies blowing down the pressure gradients (Chen and Hui 1990, 1992). This is similar to the ducted coast ridge observed over southeastern Australia (Baines 1980; Colquhoun et al. 1985; Holland and Leslie 1986; and others) and the backdoor cold fronts along the northeastern coast of the United States (Bosart et al. 1973).
In the vicinity of Taiwan, the mei-yu front is shallow with a depth of ∼1 km. It is cut into two branches by the mountainous terrain of the Central Mountain Range, with a local pressure ridge along the northeastern coast of Taiwan (Wang 1986; Chen et al. 1989; Chen and Hui 1990, 1992; Trier et al. 1990). The microscale structure of the leading edge of the cold front resembles a density current as observed by instrumented aircraft (Chen and Hui 1990; their Fig. 19). It is important to note that the microscale structure of the cold front cannot be resolved by mesoscale models with a grid spacing on the order of 15 km (Sun and Chern 2006). Because of the existence of the frontal slope, the passage of the 850-hPa trough over the Taiwan area frequently occurs up to 1–2 days after the passage of the surface front (e.g., Chen and Li 1995a; Tu et al. 2014; and others).
During frontal passage, mesoscale convective systems are frequently embedded in the mei-yu frontal zone (Wang et al. 1990; Lin et al. 1990; Tao et al. 1991; Li et al. 1997; Teng et al. 2000; Wang et al. 2005; Xu et al. 2012; Tu et al. 2014; and others), with localized heavy precipitation as they move onshore. After the passage of the mei-yu fronts over Taiwan, the low-level northeasterlies are deflected by the Central Mountain Range (Fig. 1b), leaving warm, moist air over the southwestern plain (Chen et al. 1989; Chen and Hui 1990, 1992).
In addition to front–terrain interaction, the prefrontal flow is affected by orographic blocking of the southwesterly flow by the Central Mountain Range. Three types of low-level jets (LLJs) may occur during the mei-yu season over Taiwan. (i) The subsynoptic low-level jet (SLLJ; Chen and Yu 1988; Y.-L. Chen et al. 1994; G. Chen et al. 2005; Du et al. 2014) occurs between the 900- and 600-hPa levels (∼1–4 km). The 850–700-hPa SLLJ in the frontal zone ahead of the mei-yu trough is associated with the secondary circulation of the moist baroclinic jet/front system (Chen et al. 1994, 1997; Chen and Tseng 2000; Chen and Chen 2002; Hsiao and Chen 2014). (ii) The marine boundary layer jet (MBLJ; Tu et al. 2019, 2020) occurs in the planetary boundary layer (PBL) below the 900-hPa level (≤1 km). The MBLJ develops when the mei-yu trough is deeper and/or the western Pacific subtropical high (WPSH) is stronger, with the ridge axis extending farther westward than the climatological mean. In the MBLJ core, the vertical wind profile resembles an Ekman spiral in response to the three-way balance among pressure gradients, Coriolis force, and frictional force, and has a wind speed maximum of more than 10 m s−1 in the PBL (Tu et al. 2019). (iii) The barrier jet (BJ; Chen and Li 1995a,b; Li and Chen 1998; Yeh and Chen 2003) off or along the northwest Taiwan coast is a result of orographic blocking of the SLLJ by the island obstacle. The BJ is a local-scale feature (Chen and Li 1995a,b; Li et al. 1997) with a wind speed maximum at the top of the PBL, at approximately ∼1 km (Li and Chen 1998). The barrier jet is the strongest when the surface windward ridge–leeside trough pressure gradient is the strongest. It has an early morning maximum when the land surface is the coldest (∼3 m s−1 stronger than the afternoon minimum) (Lin et al. 2011). The convergence between the BJ and southwest monsoon flow and/or frontal wind shift line leads to the development of localized heavy rainfall over northwestern Taiwan (Li et al. 1997; Yeh and Chen 2002; Chen et al. 2018; Ke et al. 2019). Moreover, the BJ transports abundant moisture locally to the convergence area from the south, which contributes to enhanced heavy coastal precipitation (Li et al. 1997; Kerns et al. 2010). In the absence of a SLLJ, the BJ could develop due to orographic blocking of the MBLJ. A notable example occurred during the Terrain-influenced Monsoon Rainfall Experiment (TiMREX) Intensive Observing Period (IOP) 8 (Tu et al. 2017, their Figs. 4a,c).
Tu et al. (2019) studied the MBLJs during the later period of the early summer rainy season of Taiwan in June for 5 years (2008–12) and examined their relationship with mei-yu heavy rainfall. The moisture transport by the MBLJ supplies moisture whereas the SLLJ provides the subsynoptic lifting needed to produce heavy mei-yu rainfall in the frontal zone (Tu et al. 2019, 2020). Furthermore, under favorable large-scale settings, orographic lifting of the southwesterly MBLJs by the Snow Mountains and/or the Central Mountain Range (Fig. 1b) contributes to heavy rainfall over the southwestern slopes of these mountains and over southwestern Taiwan (Tu et al. 2019, 2020). Du and Chen (2018, 2019) studied the impacts of the coupling of double low-level jets, e.g., a SLLJ on the south side of an inland cold front and the boundary layer jet from the South China Sea, on convective initiation of warm-sector heavy rainfall over the southern China coastal area.
Chen et al. (2018) studied the interaction between a mei-yu jet/front system and topography for a heavy rainfall event over northern Taiwan, with ∼435 mm over the northwestern Taiwan coast and ∼477 mm within the Taipei basin, during 11–12 June 2012. From 2200 LST 11 June to 0200 LST 12 June, the rainfall maximum along the northwestern coast is closely related to the arrival of scattered prefrontal echoes, followed by the east-northeast–west-southwest-oriented convective line associated with the mei-yu front. The convective activity over the northwestern coast is enhanced by the localized convergence between the southerly BJ and the postfrontal west-northwesterly flow. From 0200 to 0800 LST 12 June, the relatively deep postfrontal cold air (∼1.5 km) moves over the Yang-Ming Mountains, with peaks ∼1120 m (Fig. 1c), into the Taipei basin as the mei-yu front/850-hPa trough arrives over northern Taiwan with cold northeasterlies near the surface and northwesterlies behind the 850-hPa trough. In this case, the mei-yu front stalls over the Taipei basin for 6 h. The 850-hPa postfrontal northwesterlies impinge on the northwestern slopes of the Snow Mountains (Fig. 1c), which results in heavy rainfall. In addition, during 10–12 June 2012, the prefrontal southwesterly SLLJ and strong low-level winds associated with the MBLJ with large horizontal moisture fluxes (>360 g kg−1 m s−1 at the 950-hPa level) were lifted by the southwestern slopes of the Snow Mountains and the Central Mountain Range, and resulted in a 3-day heavy orographic precipitation accumulation of >1000 and >1500 mm, respectively.
For the coastal heavy rainfall event during TiMREX IOP 8, the warm, moist south-southwesterly MBLJ is lifted by a ∼500-m-deep cold pool over the southwestern coast. The cold pool was caused by rain-evaporative cooling resulting from antecedent rains (Xu et al. 2012; Tu et al. 2014, 2017). Above the cold pool, the south-southwesterly flow deflected by orographic blocking is nearly parallel to the Central Mountain Range (Fig. 1b). Thus, the heavy rainfall maximum occurs over the southwest coast of Taiwan, where the MBLJ converges with the offshore flow associated with the cold pool. The rain showers are enhanced by localized low-level convergence, drift inland, and then diminish.
3. Data and methodology
a. Data
The National Centers for Environmental Prediction (NCEP) Final Operational Global Analysis (FNL) data with 0.25° × 0.25° grids at 6-h intervals and 31 pressure levels are used to delineate the subsynoptic weather patterns. The rainfall and surface wind observations are from conventional surface weather stations and the Automatic Rainfall and Meteorological Telemetry System (ARMTS; Kerns et al. 2010) (512 stations with hourly rainfall data). The time series of station-observed temperature, dewpoint temperature, surface winds, and rainfall at Anbu, Zhuzihu, Tamsui, Keelung, Wenshan, and Guishan (Fig. 1c) are used to analyze rain-evaporative cooling, the arrival of postfrontal cold northerly flow at the surface, and the interaction between the cold northerlies and the terrain of northern Taiwan. Banciao (Fig. 1c) sounding data, available from the Central Weather Bureau, are used to delineate the depth of the postfrontal cold air behind the mei-yu front.
b. Model description
The Weather Research and Forecasting (WRF) Model (Skamarock et al. 2008) uses the sigma (terrain following) hydrostatic-pressure vertical coordinate (Laprise 1992). Four model domains with two-way nesting are used with horizontal grids of 27 (domain 1), 9 (domain 2), 3 (domain 3), and 1 km (domain 4), respectively (Fig. 1a). There are 65 sigma levels from the surface to the 30-hPa level. There are 22 and 34 sigma levels below 1 km and through the depth of Taiwan’s terrain, respectively. The WRF model physics schemes used in this study include the Rapid Radiative Transfer Model for GCMs (RRTMG; Iacono et al. 2008) longwave and shortwave schemes; Noah land surface model (LSM) (Chen and Dudhia 2001); Yonsei University (YSU) planetary boundary layer scheme (Hong et al. 2006); modified Kain–Fritsch cumulus parameterization scheme (Kain 2004); and Goddard microphysics (Tao and Simpson 1993) with three classes of ice (including ice, snow, and graupel). The cumulus parameterization scheme is turned off for 3- and 1-km domains. The surface layer uses the fifth-generation Pennsylvania State University–NCAR Mesoscale Model (MM5) similarity scheme with stability functions from Paulson (1970), Dyer and Hicks (1970), and Webb (1970). A 0.083° daily real-time global sea surface temperature (RTG_SST_HR) analysis is used as the lower boundary condition over the ocean (Gemmill et al. 2007). The land use data are from the Moderate Resolution Imaging Spectroradiometer (MODIS) dataset with a 500-m resolution.
In this study, we analyze the IVT within the boundary layer (defined for convenience as the surface to 900 hPa). The interaction between the mei-yu jet/front system and the terrain of Taiwan is presented using the model results from the 3- and 1-km grid domains. The Yang-Ming Mountains are removed in a model sensitivity test (RmY run) to discuss the impacts of the terrain over northern Taiwan on the mei-yu jet/front system and the distribution of rainfall. In the RmY run, the model terrain height is set to 150 m (flat terrain) where the Yang-Ming Mountains’ terrain (with peaks ∼1120 m) is higher than 150 m.
4. Weather conditions
a. Synoptic conditions
At 1200 UTC 31 May 2017 (Fig. 2a), a SLLJ prevails at the 850-hPa level between a lee cyclone (L), which originates in the leeside of the Yun-Guei Plateau over southwestern China, and the WPSH extending westward to the South China Sea. At 1200 UTC 1 June 2017 (Fig. 2b), the lee cyclone L is stationary and the other two mei-yu frontal cyclones (L1 and L2) develop over southeastern China with a stronger SLLJ over the southeastern flank of the cyclones. At 0000 UTC 2 June 2017 (Fig. 2c), the mei-yu frontal cyclone L1 (L2) moves eastward off southeastern China (over southeastern China) and another mei-yu frontal cyclone L3 develops over southern China. At 1200 UTC 2 June 2017 (Fig. 2d), as the mei-yu frontal cyclones (L1, L2, and L3) weaken, the SLLJ also weakens.
The NCEP FNL 850-hPa winds (m s−1; one full barb represents 5 m s−1) and geopotential height (gpm; contours) at (a) 1200 UTC (2000 LST) 31 May 2017, (b) 1200 UTC (2000 LST) 1 Jun 2017, (c) 0000 UTC (0800 LST) 2 Jun 2017, and (d) 1200 UTC (2000 LST) 2 Jun 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
At 0000 UTC 2 June 2017 (Fig. 3a), the north-northeast–south-southwest upper-level trough at the 500-hPa level extends from the East China Sea to east of Taiwan. An east–west-oriented westerly upper-level jet passes through Taiwan (Fig. 3a). Another upper-level jet is present in the southwestern flank of the midlatitude cyclones (La and Lb; Fig. 3a). At the 300-hPa level (Fig. 3b), the north-northeast–south-southwest upper-level trough extends from the Yellow Sea to the eastern China coast showing the westward tilt of the upper-level trough.
(a) The NCEP FNL 500-hPa winds (m s−1; one full barb represents 5 m s−1) and geopotential height (gpm; contours) at 0000 UTC (0800 LST) 2 Jun 2017. The dashed line represents the upper-level trough axis. (b) As in (a), but for 300-hPa winds (m s−1; one full barb represents 5 m s−1), geopotential height (gpm; contours), and relative vorticity (10−4 s−1; color shaded).
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
b. Rainfall and the mei-yu front over the Taiwan area
At 0000 LST 2 June 2017, an east-northeast–west-southwest-oriented convective line (radar echoes > 50 dBZ) associated with the mei-yu front is off the northern coast of Taiwan (not shown). There is very little rainfall over northern Taiwan before 0200 LST 2 June (Fig. 4a). From the station-observed hourly rainfall, the heavy rainfall occurs over the northern (windward) slopes of the Yang-Ming Mountains during 0200–1000 LST 2 June 2017 (∼8 h) (Figs. 4a–f). During 1100–1200 LST 2 June (Fig. 4g), the east–west-oriented mei-yu front moves over the Yang-Ming Mountains and arrives at the Taipei basin, resulting in an east–west-oriented rainband with the maximum accumulated rainfall over the northwestern coast and the Taipei basin (>70 mm h−1) (Fig. 4g). During 1300–1400 LST 2 June, the mei-yu frontal rainband continues to move southward with rainfall over the northern slopes of the Snow Mountains (Fig. 4h) (Fig. 1c). There are persistent radar echoes on the southwestern windward side of the Snow Mountains and the Central Mountain Range as the prefrontal southwesterly flow impinges on these mountains (not shown).
The observed station surface winds (m s−1; one full barb represents 5 m s−1) and hourly rainfall accumulation (mm; color shaded) from rain gauges ending at (a) 0200 LST 2 Jun 2017, (b) 0300 LST 2 Jun 2017, (c) 0400 LST 2 Jun 2017, (d) 0600 LST 2 Jun 2017, (e) 0800 LST 2 Jun 2017, (f) 1000 LST 2 Jun 2017, (g) 1200 LST 2 Jun 2017, and (h) 1400 LST 2 Jun 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
Figures 5 and 6 show the time series of temperature, dewpoint temperature, surface winds, and rainfall at Anbu, Zhuzihu, Tamsui, Keelung, Wenshan, and Guishan stations (Fig. 1c) from 0100 to 1500 LST 2 June. During 0200–1000 LST 2 June (Fig. 4), the shallow mei-yu front cannot move over the Yang-Ming Mountains (with peaks ∼1120 m) when it first arrives. The postfrontal cold air at the surface is deflected by the Yang-Ming Mountains and moves through two river valleys, Keelung River and Tamsui River (Figs. 4 and 5c,d) with shallow northerly winds anchored over the northern slopes of the Yang-Ming Mountains (Fig. 4). As the depth of the postfrontal cold air increases, the shallow front is finally able to pass over the Yang-Ming Mountains (Figs. 5a,b) and move into the Taipei basin (Fig. 5e) around 1100 LST 2 June.
The time series of observed temperature (black solid line), dewpoint temperature (blue dashed line), and surface winds (m s−1; one full barb represents 1 m s−1) from 0100 to 1500 LST 2 Jun 2017 at various stations shown in Fig. 1c: (a) Anbu, (b) Zhuzihu, (c) Tamsui, (d) Keelung, (e) Wenshan, and (f) Guishan. Red line represents the sunrise time.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
The time series of observed hourly rainfall (mm) from 0100 to 1500 LST 2 Jun 2017 at various stations given in Fig. 1c: (a) Anbu, (b) Zhuzihu, (c) Tamsui, (d) Keelung, (e) Wenshan, and (f) Guishan. Orange triangle represents a trace (rainfall < 0.1 mm).
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
At Tamsui and Keelung, the temperature drop is evident at 0300 LST 2 June as the surface cold air arrives at the entrances of the Tamsui River Valley and the Keelung River Valley (Figs. 5c,d), having been modified by rain evaporative cooling (Figs. 6c,d). Furthermore, the frontal convection is enhanced over the northern slopes of the Yang-Ming Mountains with the most significant rainfall along the northern coast (Figs. 4b,c). The entire Yang-Ming Mountains (including Zhuzihu and Anbu) receive notable rainfall from the mei-yu frontal system (Figs. 4b,c and 6a,b). However, before 1100 LST, the shallow postfrontal cold air does not reach Zhuzihu (607 m) and Anbu (838 m) on the leeside slopes of the Yang-Ming Mountains as the cold air behind the front is unable to move over the hilltop (∼1120 m).
At Anbu and Zhuzihu, there are two temperature drops on 2 June (Fig. 5). The first temperature drop around 0200 LST is due to rain evaporative cooling as heavy rains are anchored along the windward coast (Figs. 4 and 6), whereas the second temperature drop (at around 1100 LST) is due to the arrival of the deepening postfrontal air that is eventually able to climb over the Yang-Ming Mountains and has been modified by rain evaporative cooling (Figs. 4, 5, and 6). From 0200 to 1000 LST 2 June (Figs. 4a–f), the cold pool from intense convection over the northern side of the Yang-Ming Mountains further enhances the strength and depth of the postfrontal air. Note that nocturnal cooling will result in a gradual decrease in temperature, rather than a sudden drop in temperature.
At Wenshan in the foothills of the Snow Mountains south of the Taipei basin, there are also two temperature drops around 0400–0500 LST 2 June and 1100 LST 2 June (Fig. 5e). The first temperature drop (Fig. 5e) is due to the arrival of the surface cold air through the Keelung River Valley modified by rain evaporative cooling (Figs. 4c and 6e), whereas the second temperature drop (Fig. 5e) is due to deepening of the postfrontal air that moves over the Yang-Ming Mountains into Taipei basin (Fig. 6e). Note that the easterly flow at the Wenshan Station at around 1100 LST 2 June (Fig. 5) could be modified by the terrain and the frontal convection (Figs. 4 and 6).
Over the southwestern end of the Taipei basin (Guishan) (Fig. 5f), the southwesterly flow prevails during 0100 LST–1000 LST 2 June. Afterward, the deepened postfrontal cold northerly flow moves over the Yang-Ming Mountains and arrives at Guishan at 1100 LST 2 June (Fig. 5f) with notable rainfall (Fig. 6f). There is a slight temperature drop before 0400 LST 2 June (Fig. 5f), perhaps due to the arrival of the shallow cold air through the Tamsui River Valley. When the shallow surface cold air arrives at the Taipei basin by way of the Tamsui River Valley and the Keelung River Valley, only light rainfall occurs within the Taipei basin before the cold front moves over the Yang-Ming Mountains at 1100 LST 2 June (Fig. 4).
Figure 7 shows the time series of potential temperature and wind profiles constructed from rawinsonde data collected at the Banciao station, located on the southwestern corner of the Taipei basin (Fig. 1c). Prior to 0000 UTC (0800 LST) 2 June, the southwesterly flow prevails at the Banciao station. The shallow (<1 km) cold front arrives before 1200 UTC (2000 LST) 2 June.
The time series of wind (m s−1; one full barb represents 5 m s−1) and potential temperature (K; color shaded) profiles from 1200 UTC (2000 LST) 31 May to 1200 UTC (2000 LST) 3 Jun 2017 constructed from rawinsonde data collected at Banciao Weather Station (Fig. 1c).
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
5. Simulations of structure of the mei-yu system and moisture transport
At 1200 UTC (2000 LST) 1 June (Fig. 8a), at the 950-hP level, the mei-yu frontal cyclone (L1) is located over the southeast coast of China and the WPSH extends westward to the southern South China Sea with large pressure gradients and a MBLJ in between. The axis of maximum wind speed (>20 m s−1) of the MBLJ extends from the northern South China Sea to the Taiwan Strait. The mei-yu front is located north of Taiwan. The prefrontal southwesterly flow is blocked by the Central Mountain Range, resulting in the development of the BJ along the northwestern Taiwan coast (>25 m s−1) and a lee vortex off the southeastern coast as found by Chen and Tsay (1978), Sun et al. (1991), and Sun and Chern (1993). Note that as the MBLJ intensifies and interacts with the Central Mountain Range, the BJ intensifies simultaneously (Fig. 8a). The horizontal moisture transport by the southwesterly MBLJ along the Taiwan Strait is evident (Figs. 8a and 9a) with a maximum associated with the BJ (>430 kg m−1 s−1) off northwestern Taiwan (Fig. 9a). Additionally, strong orographically induced winds (>25 m s−1) are present in the southeastern flank of the lee vortex, a result of the large pressure gradients between the lee vortex (Lt) and the WPSH (Fig. 8a).
Simulated 950-hPa winds (m s−1; one full barb represents 5 m s−1), and geopotential height (gpm; contours) from the 9-km gridded domain (domain 2): (a) 1200 UTC (2000 LST) 1 Jun 2017, (b) 1600 UTC 1 Jun (0000 LST 2 Jun) 2017, (c) 2200 UTC 1 Jun (0600 LST 2 Jun) 2017, and (d) 0600 UTC (1400 LST) 2 Jun 2017. The light yellow shading represents an area with no data where the station pressure is below 950 hPa; L1, L2, and L3 represent mei-yu frontal cyclones; and Lt represents a lee vortex.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
Simulated integrated vapor transport (IVT; kg m−1 s−1) in the boundary layer (below 900 hPa) from the 9-km gridded domain (domain 2): (a) 1200 UTC (2000 LST) 1 Jun 2017, (b) 2200 UTC 1 Jun (0600 LST 2 Jun) 2017, and (c) 0600 UTC (1400 LST) 2 Jun 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
At 1600 UTC 1 June (0000 LST 2 June) (Fig. 8b), the mei-yu frontal cyclone (L1) moves eastward off southeastern China. At 2200 UTC 1 June (0600 LST 2 June) (Fig. 8c), the mei-yu frontal cyclone (L1) moves eastward and is located northeast of Taiwan. Within the northern Taiwan Strait, the northwesterly flow in the southwestern flank of L1 converges with southwesterly flow along the Taiwan Strait and the BJ off the northwest coast (Fig. 8c). The local maximum in moisture transport associated with the BJ off the northwestern coast is greater (>460 kg m−1 s−1) than before (Fig. 9b).
Figure 10 shows a northeast–southwest cross section along the Taiwan Strait (∼2350-km-long dark red line in Fig. 1b). The SLLJ wind maximum (>20 m s−1) extends vertically from 900 to 600 hPa, mainly in the frontal zone (Fig. 10a), with moisture fluxes > 420 g kg−1 m s−1 (Fig. 11). In addition, the MBLJ with a wind maximum of 20 m s−1 at the 925-hPa level around 20.5°N is also simulated (Fig. 10a). Within the MBLJ, the simulated maximum axis of the moisture fluxes (>360 g kg−1 m s−1) occurs at the 950-hPa level (Figs. 10a and 11) in agreement with the analyses by Tu et al. (2020). The horizontal moisture fluxes are locally enhanced off northwestern Taiwan (∼24.4°N; >450 g kg−1 m s−1 at the 950-hPa level) (Fig. 11) due to the presence of the BJ (Figs. 8c and 10a).
The vertical cross section along the dark red line in Fig. 1b (across the SCS and Taiwan Strait) at 2200 UTC 1 Jun (0600 LST 2 Jun) 2017 from the 9-km gridded domain (domain 2): (a) winds (m s−1; one full barb represents 5 m s−1) and potential temperature (K; contours), (b) relative humidity (%; color shaded) and equivalent potential temperature (K; contours), (c) radar reflectivities (dBZ; color shaded), and (d) potential vorticity (PVU; color shaded) and specific humidity (g kg−1; contours).
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
The vertical cross section of the simulated horizontal moisture fluxes (vectors with magnitude color shaded every 30 g kg−1 m s−1) along the dark red line in Fig. 1b (across the SCS and Taiwan Strait) at 2200 UTC 1 Jun (0600 LST 2 Jun) 2017 from the 9-km gridded domain (domain 2).
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
In the frontal region, the moisture tongue extends upward along the sloping frontal surface to the 350-hPa level (Figs. 10a,b,d; RH > 80%). In the postfrontal region, the cold dry northerly flow prevails (Figs. 10a,b). Similar to cases found in previous studies, the mei-yu front has a vertical tilt with a slope approximately 1/200 (Fig. 10a). The depth of the postfrontal cold air is deeper farther to the north behind the leading edge. The latent heat release associated with frontal convection, having radar echoes > 30 dBZ up to the 250-hPa level, feeds back to the environmental flow and high potential vorticity extends vertically upward from the 700-hPa level to the 350-hPa level (Figs. 10c,d). The tropopause folding associated with the high-PV air (>2 PVU; 1 PVU = 10−6 K kg−1 m2 s−1) (Lamarque and Hess 1994) occurs in the midlatitudes (∼34°N) where the upper-level jet prevails (>60 m s−1 at the 200-hPa level) (Figs. 10a,d) as found in Chen and Hui (1990, 1992) and Hsiao and Chen (2014).
Figure 12 shows the 1109-km-long southwest–northeast cross section across the Central Mountain Range (purple line in Fig. 1b). The upstream southwesterly flow impinges on the Central Mountain Range with a wind speed maximum in the 950–900-hPa layer (Fig. 12a). To the southwest, high moisture (RH > 90%) is confined below the 900-hPa level (Fig. 12b). The horizontal moisture transport from the northern South China Sea to southwestern Taiwan occurs mainly in the marine boundary layer (Figs. 12a,c) in agreement with Tu et al. (2020). Orographic lifting (Fig. 12d) of the moist low-level southwesterly flow results in the development of heavy rainfall on the windward side of the Central Mountain Range and the Snow Mountains.
The vertical cross section along the purple line in Fig. 1b (across the SCS and the Central Mountain Range) at 2200 UTC 1 Jun (0600 LST 2 Jun) 2017 from the 9-km gridded domain (domain 2): (a) winds (m s−1; one full barb represents 5 m s−1) and potential temperature (K; contours), (b) relative humidity (%; color shaded) and equivalent potential temperature (K; contours), (c) horizontal moisture flux (vectors with magnitude color shaded every 30 g kg−1 m s−1), and (d) vertical velocity (cm s−1; color shaded).
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
At 0600 UTC (1400 LST) 2 June (Figs. 8d and 9c), the mei-yu frontal cyclone (L1) continues to move eastward with the strong northeasterly flow in the western flank of the cyclone (L1) penetrating farther southward and converging with the moisture-laden prefrontal southwesterly flow. Note that during 1–4 June 2017, the southwesterly monsoon flow over the northern South China Sea satisfies the criteria of MBLJ days (Tu et al. 2020).
6. Simulations over the Taiwan area with a grid spacing of less than 3 km
At 0600 LST 2 June (Figs. 13a,b), the mei-yu frontal cyclone (L1) northeast of Taiwan is characterized by the cold northerly flow in the northwestern flank of the cyclone. The cold west/northwesterly flow is deflected by the Yang-Ming Mountains with weak winds passing through the Tamsui River Valley and the Keelung River Valley (Fig. 14a). Note that the observed surface winds along the northern slopes of the Yang-Ming Mountains exhibit a more northerly wind component (Fig. 4) when compared to the model simulations (Fig. 14a). At the 900-hPa level (Fig. 13b), the southwesterly BJ converges with the northwesterly flow in the southwestern flank of the mei-yu frontal cyclone (L1) slightly off the northern tip of Taiwan, resulting in the enhancement of the frontal convection.
(a) Simulated 2-m temperature (°C; color shaded) and 10-m winds (m s−1; one full barb represents 5 m s−1) from the 3-km gridded domain (domain 3) at 0600 LST 2 Jun (2200 UTC 1 Jun) 2017. (b) As in (a), but for 900-hPa vertical velocity (cm s−1; color shaded) and winds (m s−1; one full barb represents 5 m s−1). The light yellow shading represents an area with no data where the station pressure is below 900 hPa. (c),(d) As in (a) and (b), but for 1200 LST 2 Jun (0400 UTC 2 Jun) 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
Simulated 2-m temperature (°C; color shaded) and 10-m winds (m s−1; one full barb represents 5 m s−1) from the 1-km gridded domain (domain 4): (a) 0600 LST 2 Jun (2200 UTC 1 Jun) 2017 and (b) 1200 LST 2 Jun (0400 UTC 2 Jun) 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
At 1200 LST 2 June, the surface front moves southward (Figs. 13c and 14b) with an intense frontal rainband arriving at the Taipei basin (Fig. 4g). At the 900-hPa level (Fig. 13d), the localized convergence between the northwesterly flow in the southwestern flank of the mei-yu trough and the southwesterly BJ enhances the frontal convection over the Taipei basin and the northwestern coast. Over northwestern Taiwan, at the surface (Fig. 14b), the postfrontal northerly flow advances southward and moves to south of the Linkou plateau (Fig. 1c). Over northeastern Taiwan, the northerly flow passes through the Keelung River Valley and arrives at the Taipei basin (Fig. 14b).
Figure 15 shows the north–south cross section through the Yang-Ming Mountains and the Snow Mountains (133-km-long pink line in Fig. 1c). From the figure, it is evident that in the early morning of 2 June 2017, the shallow stable cold northerly flow is anchored on the north side of the Yang-Ming Mountains at the lowest levels (Fig. 15a) with a relatively warm westerly flow above. The shallow northerly flow deflected by the Yang-Ming Mountains passes through the Keelung River Valley and the Tamsui River Valley, and arrives at the Taipei basin (Figs. 15a and 16a,c). Around noontime, the cold northerly flow with increasing depth (∼1 km) advances southward and arrives at the Taipei basin (Figs. 15b and 16b,d). Within the Taipei basin, south of the Yang-Ming Mountains, the cold air becomes deeper as time progresses.
Model results from the 1-km gridded domain (domain 4). (a) Vertical cross section of winds (m s−1; one full barb represents 5 m s−1) and potential temperature (K; color shaded) along the pink line in Fig. 1c (across the Yang-Ming Mountains and Snow Mountains) at 0400 LST 2 Jun (2000 UTC 1 Jun) 2017. (b) As in (a), but for 1300 LST 2 Jun (0500 UTC 2 Jun) 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
Model results from the 1-km gridded domain (domain 4). (a) Vertical cross section of winds (m s−1; one full barb represents 5 m s−1) and potential temperature (K; color shaded) along the purple line in Fig. 1c (along the Keelung River Valley) at 0400 LST 2 Jun (2000 UTC 1 Jun) 2017. Keelung Weather Station is denoted by a red triangle with a letter “K.” (b) As in (a), but for 1300 LST 2 Jun (0500 UTC 2 Jun) 2017. (c) Vertical cross section of winds (m s−1; one full barb represents 5 m s−1) and potential temperature (K; color shaded) along the dark red line in Fig. 1c (along the Tamsui River Valley) at 0600 LST 2 Jun (2200 UTC 1 Jun) 2017. Tamsui Weather Station is denoted by a red triangle with a letter “T.” (d) As in (c), but for 1200 LST 2 Jun (0400 UTC 2 Jun) 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
From the Hovmöller diagrams of the simulated winds and potential temperature at the 990-hPa level (∼100 m) along a 49-km-long purple line in Fig. 1c through the Keelung River Valley, and a 54-km-long dark red line in Fig. 1c through the Tamsui River Valley (Fig. 17), it is evident that the simulated cold northerly flow starts to advance from the surrounding ocean to the Keelung River Valley and the Tamsui River Valley at 0300 LST 2 June and 0400 LST 2 June, respectively, and then into the Taipei basin.
Simulated Hovmöller diagrams at the 990-hPa level from the 1-km gridded domain (domain 4). (a) Winds (m s−1; one full barb represents 5 m s−1) and potential temperature (K; color shaded) along the dark red line in Fig. 1c (along the Tamsui River Valley) during 0100–1500 LST 2 Jun 2017. Tamsui Weather Station is denoted by a red semicircle with a letter “T.” (b) Winds and potential temperature along the purple line in Fig. 1c (along the Keelung River Valley). Keelung Weather Station is denoted by a red semicircle with a letter “K.”
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
When the shallow cold air reaches the northern side of the Yang-Ming Mountains in the early morning of 2 June 2017, Fr0 = 0.4–0.7 and M = 3.3 (θ0 = 299 K, Δθ ∼ 3 K, g* = 0.098 m s−2, U0 ∼ 2–4 m s−1, H0 = 300 m, hm = 1000 m). Under these conditions, the cold airflow is unable to climb over the top of the Yang-Ming Mountains. This results in flow separation on the windward side (Regime III) with given Fr0, Ms ∼ 1.1–1.2 and H0s ∼ 850–900 m. Note that, H0s is defined as hm/Ms from Eq. (3). Therefore, based on the M × Fr0 regime diagram, when the cold air depth exceeds ∼850 m, the cold airflow can pass over the Yang-Ming Mountains.
From the evening of 1 June to the early morning of 2 June (LST), before the mei-yu front reaches Taiwan, the front demonstrates features of propagation (Figs. 8a–c) as found by Chen et al. (1989), Trier et al. (1990), Chen (1993) and others. Because the leading edge of the front is very shallow in this case, the front is unable to pass over the Yang-Ming Mountains when it first arrives. Instead, it stays anchored over the northern side of the Yang-Ming Mountains for almost 8 h (0200–1000 LST 2 June), until the depth of the cold air exceeds ∼850 m. In the CTRL run (Fig. 18b), in addition to a rainfall maximum along the northern coast, the simulated daily rainfall accumulation exhibits a rainfall maximum over the southwestern windward slopes of the Snow Mountains and the Central Mountain Range, in agreement with rain gauge observations (Fig. 18a).
(a) Daily (0000–2400 LST) rainfall accumulation (mm; color shaded) for 2 Jun 2017 from rain gauge observations. The simulated daily rainfall accumulation (mm; color shaded) for 2 Jun 2017 (LST) from the 3-km gridded domain (domain 3): (b) the CTRL run and (c) the RmY run.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
In the RmY run, at 0600 LST 2 June (Figs. 19b and 20a), the BJ off the northwestern coast and/or over northern Taiwan is stronger and extends farther northward when compared to the CTRL run (Figs. 13b and 14a). As a result, the surface front in the RmY run does not reach the northern tip of Taiwan at this time (Figs. 19a and 20a). Unlike the CTRL run (Fig. 14a), the southerly flow at the surface prevails in the Taipei basin in the RmY run (Fig. 20a). Furthermore, at the 900-hPa level, the southwesterly BJ converges with the northwesterly flow in the southwestern flank of the mei-yu frontal cyclone (L1) northeast of Taiwan, producing frontal convection off the northern tip of Taiwan (Fig. 19b). Without the Yang-Ming Mountains in the model, the surface front continues to move southward (Figs. 19 and 20). At 1200 LST 2 June (Fig. 20b), the surface front has already passed through northern Taiwan and is located farther south than the CTRL run (Fig. 14b). As a result, on 2 June 2017 (LST) (Fig. 18c), there is no daily rainfall maximum over the northern tip of Taiwan in the RmY run. Instead, a rainfall maximum is simulated over northwestern Taiwan (Fig. 18c) where the BJ converges with the postfrontal northwesterly flow.
Model results from the 3-km gridded domain (domain 3) for the RmY run. (a) 2-m temperature (°C; color shaded) and 10-m winds (m s−1; one full barb represents 5 m s−1) at 0600 LST 2 Jun (2200 UTC 1 Jun) 2017. (b) As in (a), but for 900-hPa vertical velocity (cm s−1; color shaded) and winds (m s−1; one full barb represents 5 m s−1). The light yellow shading represents an area with no data where the station pressure is below 900 hPa. (c),(d) As in (a) and (b), but for 1200 LST 2 Jun (0400 UTC 2 Jun) 2017.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
Simulated 2-m temperature (°C; color shaded) and 10-m winds (m s−1; one full barb represents 5 m s−1) from the 1-km gridded domain (domain 4) for the RmY run: (a) 0600 LST 2 Jun (2200 UTC 1 Jun) 2017 and (b) 1200 LST 2 Jun (0400 UTC 2 Jun) 2017. The black (gray) frontal symbol marks the location of the surface front for the RmY (CTRL) run.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
7. Summary and conclusions
On 2 June 2017, the marine boundary layer jet (MBLJ) over the northern South China Sea prevails between the mei-yu trough over southeastern China and the western Pacific subtropical high (WPSH), extending westward to the southern South China Sea. Orographic blocking of the prefrontal southwesterly flow by the Central Mountain Range results in a barrier jet (BJ) off and along northwestern Taiwan. Prior to the arrival of the mei-yu front, orographic lifting of the prefrontal moisture-laden southwesterly flow contributes to heavy rainfall on the windward side of the Central Mountain Range and the Snow Mountains. The horizontal moisture transport from the northern South China Sea to the Taiwan area mainly occurs below the 900-hPa level (Tu et al. 2020). The localized enhanced horizontal moisture transport within the boundary layer by the BJ brings in abundant moisture to the east–west-oriented mei-yu front over northern Taiwan.
The depth of the mei-yu front impacts the rainfall pattern as the front interacts with the terrain over northern Taiwan. The unusually shallow east–west-oriented mei-yu front (<850 m) is anchored over the northern side of the Yang-Ming Mountains (with peaks ∼1120 m) for 8 h during the early morning of 2 June 2017 (0200–1000 LST), producing torrential rain along the northern coast (>600 mm). In contrast, for the 12 June 2012 case, the relatively deep mei-yu front (∼1.5 km) moves quickly over the Yang-Ming Mountains and stalls over the Taipei basin for 6 h (0200–0800 LST), producing rainfall maxima within the Taipei basin and along the northwestern coast (Chen et al. 2018). In the 2 June 2017 case, the postfrontal stable cold air at the surface is shallower than it is in the 12 June 2012 case, causing it to stay anchored along the northern side of the Yang-Ming Mountains during 0200–1000 LST (Fig. 21a). At the surface, the postfrontal cold air is deflected by the Yang-Ming Mountains and moves through the Keelung River Valley and Tamsui River Valley into the Taipei basin (Fig. 21a). At the 900-hPa level, the southwesterly BJ prevails over northern/northwestern Taiwan and converges with the northwesterly flow in the southwestern flank of the mei-yu frontal cyclone around the northern tip of Taiwan (Fig. 21a).
A schematic diagram shows the surface (black solid arrows) and 900-hPa (red solid arrows) flow patterns with the barrier jet (thick red solid arrows). (a) The shallow (<1 km) mei-yu front is anchored over the northern side of the Yang-Ming Mountains for almost 8 h during the morning of 2 Jun 2017 (0200–1000 LST 2 Jun). At the 900-hPa level, the southwesterly BJ converges with the northwesterly flow in the southwestern flank of the mei-yu frontal cyclone around the northern tip of Taiwan. (b) Around noontime, the mei-yu front finally moves southward over the Yang-Ming Mountains into the Taipei basin and converges with the southwesterly BJ over the Taipei basin.
Citation: Monthly Weather Review 150, 3; 10.1175/MWR-D-21-0113.1
Around noontime, with an increase in the depth of the postfrontal cold air to >850 m, the mei-yu front is finally able to pass over the Yang-Ming Mountains and moves into the Taipei basin (Fig. 21b). At this time, the mei-yu frontal rainband develops within the Taipei basin. The southwesterly BJ converges with the postfrontal northwesterly flow at the 900-hPa level (Fig. 21b). A temperature drop is evident at the Tamsui and Keelung stations, as the cold air at the surface arrives at 0300 LST 2 June. The Taipei basin also experienced two temperature drops in the early morning (∼0400–0500 LST 2 June) and around noontime (∼1100 LST 2 June), corresponding to the arrival of the surface cold air through two river valleys (Tamsui River and Keelung River valleys), followed by the arrival of the mei-yu front that moved over the Yang-Ming Mountains. Furthermore, two temperature drops are recorded at Zhuzihu and Anbu stations in the Yang-Ming Mountains at 0200 LST 2 June and 1100 LST 2 June, corresponding to rain evaporative cooling and to the deepening of the postfrontal air that is also modified by rain evaporative cooling. Note that before 1100 LST 2 June, the shallow front anchored over the northern slopes of the Yang-Ming Mountains does not reach Zhuzihu and Anbu.
The BJ and rainfall maxima over the northern tip of Taiwan, Taipei basin, and the southwestern slopes of the Snow Mountains and the Central Mountain Range are well simulated using a high-resolution (<3 km) numerical model. With the terrain of the Yang-Ming Mountains removed in the RmY run, as the mei-yu front makes landfall over northern Taiwan, it quickly moves southward without producing a rainfall maximum over the northern side of the Yang-Ming Mountains on 2 June 2017. Instead, in addition to heavy rainfall over the southwestern slopes of the Snow Mountains and the Central Mountain Range, a rainfall maximum is simulated over northwestern Taiwan.
Based on studies of heavy rainfall events over northern Taiwan on 11–12 June 2012 (Chen et al. 2018) and 2 June 2017, we found that under favorable large-scale settings with abundant low-level moisture, the depth of the postfrontal cold air is crucial for frontal movement and heavy rainfall distribution over northern Taiwan (e.g., a rainfall maximum occurs over the northern coast of Taiwan versus within Taipei basin). In the future, we would like to measure the depth of the postfrontal cold air from soundings along the northern coast of Taiwan and off the northern coast at Peng-Chia Yu (Fig. 1c), or use radio occultation (RO) data from the Constellation Observing System for Meteorology, Ionosphere and Climate-2 (COSMIC-2) global positioning system (GPS). Sounding observations from future field experiments, e.g., Prediction of Rainfall Extremes Campaign In the Pacific (PRECIP)/Taiwan-Area Heavy rain Observation and Prediction Experiment (TAHOPE), will be analyzed for this scientific goal. Additionally, the Storm Tracker radiosondes (Hwang et al. 2020) will be used to measure the depth of the postfrontal cold northerly flow at the entrances of the Tamsui River Valley and Keelung River Valley, and along the northern coast of Taiwan. Note that the Storm Tracker is a newly developed, smaller, lighter, cheaper, and multichannel simultaneous capable radiosonde that is beneficial for high spatiotemporal resolution observations (Hwang et al. 2020). Further, we would also like to study the diurnal variations of the MBLJ and BJ under the presence of a MBLJ, but without a SLLJ, using a high-resolution long-term climatological dataset, e.g., ERA5 analysis, wind profiler observations (at Dongsha Island and over northwestern Taiwan; Fig. 1b), and WRF model simulations, as well as the impacts of the MBLJ on heavy rainfall over southwestern Taiwan and the windward slopes of the Central Mountain Range and the Snow Mountains.
Acknowledgments.
This work is jointly funded by the Ministry of Science and Technology (MOST) under Grants MOST 104-2923-M-008-003-MY5, MOST 109-2111-M-008-017, and MOST 110-2111-M-008-021, and the Featured Areas Research Center Program within the framework of the Higher Education Sprout Project by the Ministry of Education (MOE) to the National Central University; and by the National Science Foundation under Grants AGS-1142558 and AGS-1854443 to the University of Hawai‘i at Mānoa. We thank two anonymous reviewers for their comments, and May Izumi, David Hitzl, and Rachael Eichelberger-Iga for editing the text.
REFERENCES
Arakane, S., H.-H. Hsu, C.-Y. Tu, H.-C. Liang, Z.-Y. Yan, and S.-J. Lin, 2019: Remote effect of a tropical cyclone in the Bay of Bengal on a heavy-rainfall event in subtropical East Asia. npj Climate Atmos. Sci., 2, 25, https://doi.org/10.1038/s41612-019-0082-8.
Baines, P. G., 1980: The dynamics of the southerly buster. Aust. Meteor. Mag., 28, 175–200.
Barrett, B. S., R. Garreaud, and M. Falvey, 2009: Effect of the Andes Cordillera on precipitation from a midlatitude cold front. Mon. Wea. Rev., 137, 3092–3109, https://doi.org/10.1175/2009MWR2881.1.
Bond, N. A., and Coauthors, 1997: The Coastal Observation and Simulation with Topography (COAST) experiment. Bull. Amer. Meteor. Soc., 78, 1941–1955, https://doi.org/10.1175/1520-0477(1997)078<1941:TCOASW>2.0.CO;2.
Bosart, L. F., V. Pagnotti, and B. Lettau, 1973: Climatological aspects of eastern United States back-door cold frontal passages. Mon. Wea. Rev., 101, 627–635, https://doi.org/10.1175/1520-0493(1973)101<0627:CAOEUS>2.3.CO;2.
Bougeault, P., and Coauthors, 2001: The MAP special observing period. Bull. Amer. Meteor. Soc., 82, 433–462, https://doi.org/10.1175/1520-0477(2001)082<0433:TMSOP>2.3.CO;2.
Chen, F., and J. Dudhia, 2001: Coupling an advanced land surface–hydrology model with the Penn State–NCAR MM5 modeling system. Part I: Model implementation and sensitivity. Mon. Wea. Rev., 129, 569–585, https://doi.org/10.1175/1520-0493(2001)129<0569:CAALSH>2.0.CO;2.
Chen, G. T.-J., 1983: Observational aspects of the Mei-Yu phenomenon in subtropical China. J. Meteor. Soc. Japan, 61, 306–312, https://doi.org/10.2151/jmsj1965.61.2_306.
Chen, G. T.-J., and C.-Y. Tsay, 1978: A synoptic case study of Mei-Yu near Taiwan. Papers Meteor. Res., 1, 25–36.
Chen, G. T.-J., and C.-P. Chang, 1980: The structure and vorticity budget of an early summer monsoon trough (mei-yu) over southeastern China and Japan. Mon. Wea. Rev., 108, 942–953, https://doi.org/10.1175/1520-0493(1980)108<0942:TSAVBO>2.0.CO;2.
Chen, G. T.-J., and C.-C. Yu, 1988: Study of low-level jet and extremely heavy rainfall over northern Taiwan in the mei-yu season. Mon. Wea. Rev., 116, 884–891, https://doi.org/10.1175/1520-0493(1988)116<0884:SOLLJA>2.0.CO;2.
Chen, G. T.-J., C.-C. Wang, and D. T.-W. Lin, 2005: Characteristics of low-level jets over northern Taiwan in mei-yu season and their relationship to heavy rain events. Mon. Wea. Rev., 133, 20–43, https://doi.org/10.1175/MWR-2813.1.
Chen, X. A., and Y.-L. Chen, 2002: Kinetic energy budgets of the low-level jet during TAMEX IOP 5. J. Meteor. Soc. Japan, 80, 1–19, https://doi.org/10.2151/jmsj.80.1.
Chen, Y.-L., 1993: Some synoptic-scale aspects of the surface fronts over southern China during TAMEX. Mon. Wea. Rev., 121, 50–64, https://doi.org/10.1175/1520-0493(1993)121<0050:SSSAOT>2.0.CO;2.
Chen, Y.-L., and N. B.-F. Hui, 1990: Analysis of a shallow front during the Taiwan Area Mesoscale Experiment. Mon. Wea. Rev., 118, 2649–2667, https://doi.org/10.1175/1520-0493(1990)118<2649:AOASFD>2.0.CO;2.
Chen, Y.-L., and N. B.-F. Hui, 1992: Analysis of a relatively dry front during the Taiwan Area Mesoscale Experiment. Mon. Wea. Rev., 120, 2442–2468, https://doi.org/10.1175/1520-0493(1992)120<2442:AOARDF>2.0.CO;2.
Chen, Y.-L., and J. Li, 1995a: Large-scale conditions favorable for the development of heavy rainfall during TAMEX IOP3. Mon. Wea. Rev., 123, 2978–3002, https://doi.org/10.1175/1520-0493(1995)123<2978:LSCFFT>2.0.CO;2.
Chen, Y.-L., and J. Li, 1995b: Characteristics of surface airflow and pressure patterns over the island of Taiwan during TAMEX. Mon. Wea. Rev., 123, 695–716, https://doi.org/10.1175/1520-0493(1995)123<0695:COSAAP>2.0.CO;2.
Chen, Y.-L., and S.-F. Tseng, 2000: Comments on “The intensification of the low-level jet during the development of mesoscale convective systems on a mei-yu front”. Mon. Wea. Rev., 128, 495–506, https://doi.org/10.1175/1520-0493(2000)128<0495:COTIOT>2.0.CO;2.
Chen, Y.-L., Y.-X. Zhang, and N. B.-F. Hui, 1989: Analysis of a surface front during the early summer rainy season over Taiwan. Mon. Wea. Rev., 117, 909–931, https://doi.org/10.1175/1520-0493(1989)117<0909:AOASFD>2.0.CO;2.
Chen, Y.-L., X. A. Chen, and Y.-X. Zhang, 1994: A diagnostic study of the low-level jet during TAMEX IOP 5. Mon. Wea. Rev., 122, 2257–2284, https://doi.org/10.1175/1520-0493(1994)122<2257:ADSOTL>2.0.CO;2.
Chen, Y.-L., X. A. Chen, S. Chen, and Y.-H. Kuo, 1997: A numerical study of the low-level jet during TAMEX IOP 5. Mon. Wea. Rev., 125, 2583–2604, https://doi.org/10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2.
Chen, Y.-L., Y.-J. Chu, C.-S. Chen, C.-C. Tu, J.-H. Teng, and P.-L. Lin, 2018: Analysis and simulations of a heavy rainfall event over northern Taiwan during 11–12 June 2012. Mon. Wea. Rev., 146, 2697–2715, https://doi.org/10.1175/MWR-D-18-0001.1.
Chien, F.-C., C. F. Mass, and P. J. Neiman, 2001: An observational and numerical study of an intense landfalling cold front along the northwest coast of the United States during COAST IOP 2. Mon. Wea. Rev., 129, 934–955, https://doi.org/10.1175/1520-0493(2001)129<0934:AOANSO>2.0.CO;2.
Chung, K.-S., H.-J. Chiu, C.-Y. Liu, and M.-Y. Lin, 2020: Satellite observation for evaluating cloud properties of the microphysical schemes in Weather Research and Forecasting simulation: A case study of the Mei-Yu front precipitation system. Remote Sens., 12, 3060, https://doi.org/10.3390/rs12183060.
Colle, B. A., B. F. Smull, and M. J. Yang, 2002: Numerical simulations of a landfalling cold front observed during COAST: Rapid evolution and responsible mechanisms. Mon. Wea. Rev., 130, 1945–1966, https://doi.org/10.1175/1520-0493(2002)130<1945:NSOALC>2.0.CO;2.
Colquhoun, J. R., D. J. Shepherd, C. E. Coulman, R. K. Smith, and K. McInnes, 1985: The southerly burster of south eastern Australia: An orographically forced cold front. Mon. Wea. Rev., 113, 2090–2107, https://doi.org/10.1175/1520-0493(1985)113<2090:TSBOSE>2.0.CO;2.
Du, Y., and G. Chen, 2018: Heavy rainfall associated with double low-level jets over southern China. Part I: Ensemble-based analysis. Mon. Wea. Rev., 146, 3827–3844, https://doi.org/10.1175/MWR-D-18-0101.1.
Du, Y., and G. Chen, 2019: Heavy rainfall associated with double low-level jets over southern China. Part II: Convection initiation. Mon. Wea. Rev., 147, 543–565, https://doi.org/10.1175/MWR-D-18-0102.1.
Du, Y., Q. Zhang, Y.-L. Chen, Y. Zhao, and X. Wang, 2014: Numerical simulations of spatial distributions and diurnal variations of low-level jets in China during early summer. J. Climate, 27, 5747–5767, https://doi.org/10.1175/JCLI-D-13-00571.1.
Dyer, A. J., and B. B. Hicks, 1970: Flux–gradient relationships in the constant flux layer. Quart. J. Roy. Meteor. Soc., 96, 715–721, https://doi.org/10.1002/qj.49709641012.
Garvert, M. F., B. A. Colle, and C. F. Mass, 2005a: The 13–14 December 2001 IMPROVE-2 event. Part I: Synoptic and mesoscale evolution and comparison with a mesoscale model simulation. J. Atmos. Sci., 62, 3474–3492, https://doi.org/10.1175/JAS3549.1.
Garvert, M. F., C. P. Woods, B. A. Colle, C. F. Mass, P. V. Hobbs, M. T. Stoelinga, and J. B. Wolfe, 2005b: The 13–14 December 2001 IMPROVE-2 event. Part II: Comparisons of MM5 model simulations of clouds and precipitation with observations. J. Atmos. Sci., 62, 3520–3534, https://doi.org/10.1175/JAS3551.1.
Gemmill, W., B. Katz, and X. Li, 2007: Daily real-time, global sea surface temperature—High-resolution analysis: RTG_SST_HR. NOAA/NWS/NCEP/EMC/MMAB, Science Application International Corporation, and Joint Center for Satellite Data Assimilation Tech. Note 260, 39 pp., http://polar.ncep.noaa.gov/mmab/papers/tn260/MMAB260.pdf.
He, H., J. W. McGinnis, Z. Song, and M. Yanai, 1987: Onset of the Asian summer monsoon in 1979 and the effect of the Tibetan Plateau. Mon. Wea. Rev., 115, 1966–1995, https://doi.org/10.1175/1520-0493(1987)115<1966:OOTASM>2.0.CO;2.
Holland, G. J., and L. M. Leslie, 1986: Ducted coastal ridging over S.E. Australia. Quart. J. Roy. Meteor. Soc., 112, 731–748, https://doi.org/10.1002/qj.49711247310.
Hong, S.-Y., Y. Noh, and J. Dudhia, 2006: A new vertical diffusion package with an explicit treatment of entrainment processes. Mon. Wea. Rev., 134, 2318–2341, https://doi.org/10.1175/MWR3199.1.
Hsiao, F., and Y.-L. Chen, 2014: Revisiting the structure and characteristics of an early summer monsoon trough over South China in 1975. SOLA, 10, 194–198, https://doi.org/10.2151/sola.2014-041.
Hwang, W.-C., P.-H. Lin, and H. Yu, 2020: The development of the “Storm Tracker” and its applications for atmospheric high-resolution upper-air observations. Atmos. Meas. Tech., 13, 5395–5406, https://doi.org/10.5194/amt-13-5395-2020.
Iacono, M. J., J. S. Delamere, E. J. Mlawer, M. W. Shephard, S. A. Clough, and W. D. Collins, 2008: Radiative forcing by long-lived greenhouse gases: Calculations with the AER radiative transfer models. J. Geophys. Res., 113, D13103, https://doi.org/10.1029/2008JD009944.
Kain, J. S., 2004: The Kain–Fritsch convective parameterization: An update. J. Appl. Meteor. Climatol., 43, 170–181, https://doi.org/10.1175/1520-0450(2004)043<0170:TKCPAU>2.0.CO;2.
Ke, C.-Y., K.-S. Chung, T.-C. Chen Wang, and Y.-C. Liou, 2019: Analysis of heavy rainfall and barrier-jet evolution during Mei-Yu season using multiple Doppler radar retrievals: A case study on 11 June 2012. Tellus, 71A, 1571369, https://doi.org/10.1080/16000870.2019.1571369.
Kerns, B. W. J., Y.-L. Chen, and M.-Y. Chang, 2010: The diurnal cycle of winds, rain, and clouds over Taiwan during the mei-yu, summer, and autumn rainfall regimes. Mon. Wea. Rev., 138, 497–516, https://doi.org/10.1175/2009MWR3031.1.
Kingsmill, D. E., P. J. Neiman, F. M. Ralph, and A. B. White, 2006: Synoptic and topographic variability of northern California precipitation characteristics in landfalling winter storms observed during CALJET. Mon. Wea. Rev., 134, 2072–2094, https://doi.org/10.1175/MWR3166.1.
Kuo, Y.-H., and G. T.-J. Chen, 1990: The Taiwan Area Mesoscale Experiment (TAMEX): An overview. Bull. Amer. Meteor. Soc., 71, 488–503, https://doi.org/10.1175/1520-0477(1990)071<0488:TTAMEA>2.0.CO;2.
Lamarque, J.-F., and P. G. Hess, 1994: Cross-tropopause mass exchange and potential vorticity budget in a simulated tropopause folding. J. Atmos. Sci., 51, 2246–2269, https://doi.org/10.1175/1520-0469(1994)051<2246:CTMEAP>2.0.CO;2.
Laprise, R., 1992: The Euler equations of motion with hydrostatic pressure as an independent variable. Mon. Wea. Rev., 120, 197–207, https://doi.org/10.1175/1520-0493(1992)120<0197:TEEOMW>2.0.CO;2.
Lavers, D. A., G. Villarini, R. P. Allan, E. F. Wood, and A. J. Wade, 2012: The detection of atmospheric rivers in atmospheric reanalyses and their links to British winter floods and the large-scale climatic circulation. J. Geophys. Res., 117, D20106, https://doi.org/10.1029/2012JD018027.
Li, J., and Y.-L. Chen, 1998: Barrier jets during TAMEX. Mon. Wea. Rev., 126, 959–971, https://doi.org/10.1175/1520-0493(1998)126<0959:BJDT>2.0.CO;2.
Li, J., Y.-L. Chen, and W.-C. Lee, 1997: Analysis of a heavy rainfall event during TAMEX. Mon. Wea. Rev., 125, 1060–1082, https://doi.org/10.1175/1520-0493(1997)125<1060:AOAHRE>2.0.CO;2.
Lin, P.-L., Y.-L. Chen, C.-S. Chen, C.-L. Liu, and C.-Y. Chen, 2011: Numerical experiments investigating the orographic effects on a heavy rainfall event over the northwestern coast of Taiwan during TAMEX IOP 13. Meteor. Atmos. Phys., 114, 35–50, https://doi.org/10.1007/s00703-011-0155-7.
Lin, Y.-J., T.-C. C. Wang, R. W. Pasken, H. Shen, and Z.-S. Deng, 1990: Characteristics of a subtropical squall line determined from TAMEX dual-Doppler data. Part II: Dynamic and thermodynamic structures and momentum budgets. J. Atmos. Sci., 47, 2382–2399, https://doi.org/10.1175/1520-0469(1990)047<2382:COASSL>2.0.CO;2.
Margules, M., 1906: Über Temperaturschichtung in stationär bewegter und in ruhender Luft. Hann-Band. Meteor. Z., 23, 243–254. [Translated and edited by E. Volken, A. Giesche, and S. Brönnimann, 2016: On temperature stratification in resting and non-accelerated moving air. Meteor. Z., 25, 795–804, https://doi.org/10.1127/metz/2016/0833.]
Medina, S., B. F. Smull, R. A. Houze, and M. Steiner, 2005: Cross-barrier flow during orographic precipitation events: Results from MAP and IMPROVE. J. Atmos. Sci., 62, 3580–3598, https://doi.org/10.1175/JAS3554.1.
Neiman, P. J., P. O. G. Persson, F. M. Ralph, D. P. Jorgensen, A. B. White, and D. E. Kingsmill, 2004: Modification of fronts and precipitation by coastal blocking during an intense landfalling winter storm in southern California: Observations during CALJET. Mon. Wea. Rev., 132, 242–273, https://doi.org/10.1175/1520-0493(2004)132<0242:MOFAPB>2.0.CO;2.
Palmen, E., and C. W. Newton, 1969: Atmospheric Circulation Systems. Academic Press, 603 pp.
Paulson, C. A., 1970: The mathematical representation of wind speed and temperature profiles in the unstable atmospheric surface layer. J. Appl. Meteor. Climatol., 9, 857–861, https://doi.org/10.1175/1520-0450(1970)009<0857:TMROWS>2.0.CO;2.
Rotunno, R., and R. A. Houze, 2007: Lessons on orographic precipitation from the Mesoscale Alpine Programme. Quart. J. Roy. Meteor. Soc., 133, 811–830, https://doi.org/10.1002/qj.67.
Schär, C., and R. B. Smith, 1993: Shallow-water flow past isolated topography. Part I: Vorticity production and wake formation. J. Atmos. Sci., 50, 1373–1400, https://doi.org/10.1175/1520-0469(1993)050<1373:SWFPIT>2.0.CO;2.