Numerical simulations were conducted using the Cloud Model 1 (CM1) version 21.1 (McGowan et al., 1995; Barros et al., 2000; Singh and Nakamura 2009), a non-hydrostatic, fully compressible cloud model designed for idealized studies of atmospheric phenomena, such as tropical cyclones, thunderstorms, and topographic gravity waves. In this study, we conducted a series of idealized simulations along a cross-section through the Nam Co Basin, to further understand the physical processes involved in the generation of mesoscale circulation and convection initiation driven by the synergistic effects of the lake and surrounding mountains (Fig. 1a). Land cover characteristics were derived from the MCD12C1 Production (Sulla-Menashe and Friedl, 2018), with grassland assigned to lowland areas and snow cover specified over mountainous terrain in summer (Fig. 1b). When large-scale atmospheric forcing is weak or absent, circulation within the basin is primarily driven by the thermally induced lake breeze and mountain-valley circulation. Hence, a cross-sectional transect over smoothed terrain, passing through Nam Co Lake and the Nyenchen Thanglha Mountain, was selected as the reference for the experimental design to investigate the associated thermal and dynamical characteristics (Fig. 1c and d).

The model was configured to solve the fully compressible governing equations on a three-dimensional domain of size 275km×20km×8km (x, y and z dimensions, respectively) with a horizontal grid spacing of 100×100m2 and an adaptive timestep initialized suitable value of 1 s (Fig. 2). The terrain-following vertical grids were stretched (Wilhelmson and Chen, 1982) with constant vertical spacing of 100 m below 4 km and 400 m from 6–8 km, resulting in a total of 53 levels. In the default configuration, the model operates in large eddy simulation (LES), and a fifth-order WENO scheme was employed for the advection of both scalars and momentum (Jiang and Shu, 1996). Since the simulations span a few hours and a relatively small domain, the Coriolis force has a negligible impact on cloud and precipitation processes, consistent with previous idealized terrain studies (Feldmann et al., 2024; Katona and Markowski, 2021). Physical parameterizations included the 1.5-order turbulent kinetic energy (TKE) closure scheme for subgrid-scale turbulence (Deardorff, 1980), with vertical turbulence tendencies calculated implicitly using the Crank–Nicholson scheme. The Klemp–Wilhelmson compressible pressure solver (Klemp and Wilhelmson, 1978) and the Morrison double-moment moisture scheme (Morrison et al., 2005) were employed. The revised land surface scheme developed for the Weather Research and Forecasting (WRF) model is also applied (Jiménez et al., 2012).

Five sensitivity experiments were conducted based on the model configuration (detailed in Table 1) to examine the interactions between the lake and mountains of varying heights, assess the contributions of lake breeze and mountain-valley circulation, and investigate cloud-precipitation processes associated with the lake and surrounding topography.

In the model configuration, simulations incorporating both lake and topography are oriented along the x-direction of the domain (see the cross-sections in Fig. 2d). ‌Periodic boundary conditions are applied in the lateral boundary (y-direction), while an open-radiative boundary condition is imposed along the (x-direction) horizontal boundary to minimize wave reflection (Klemp and Wilhelmson, 1978). Consequently, the geometry of the open lake (x=70-130km) is effectively infinitely long in the y-direction because of the use of periodic flow-parallel lateral boundary conditions. This configuration has been shown to support the initiation of open-cell convection and lake-effect precipitation induced by the open lake (Gowan et al., 2021, 2022). Based on the observed water temperature at a depth of 6 m in Nam Co, which was used as a proxy for surface temperature by Wang et al. (2023), the lake surface temperature was set to 284 K (10.85 °C) and kept constant throughout the simulation. This assumption is justified by the relatively large heat capacity of the lake, which limits substantial lake temperature variations over the simulated period. Although the lake temperature is prescribed, the associated surface turbulent fluxes allow the lake to continuously influence atmospheric moisture availability through evaporation. This treatment is intended to represent the persistent thermal and moisture effects of the lake during the development of lake-mountain convection.

The terrain configuration is intended to provide insight into downstream processes that occur where lake-effect systems interact with lower terrain features. The “bell-cosine” terrain profile, which is used to investigate hill effects in large-eddy simulations (Xue and Giorgetta, 2021), is incorporated into the idealized model and defined by the following equation: 1zt=H21+cos⁡2π4L(x-xc), where H is the maximum hill height (1500 m) (set for control experiment: LM-1500), L is the half-width of the terrain (15 km) and xc represents the ridge center location (175 km). An idealized bell-shaped mountain defined by the cosine profile was employed to ensure smooth curvature and continuous derivatives across the terrain. This configuration minimizes numerical artifacts, such as wave reflection or artificial convergence near sharp gradients, while providing a controlled, symmetric topographic forcing suitable for isolating orographic effects on atmospheric processes. We initialized the land surface with a temperature of 300 K (26.85 °C), and the soil temperature to 285 K (11.85 °C). To represent the observed cooling of terrain surfaces with altitude, the surface temperature over the mountain was prescribed to decrease linearly with height at a fixed lapse rate of 6.5 Kkm-1.

Rather than optimizing parameterization schemes to reproduce real atmospheric conditions, our simulation intended to emphasize land–atmosphere interactions over the underlying surface of the lake and mountain, and to explore the change in convection initiation and cloud-precipitation processes over the TP. The initial profiles of potential temperature and water vapor mixing ratio used in this study were derived from radiosonde observations at the Naqu Station, whose atmospheric conditions are similar to those over Nam Co Lake, on 14 August 2023, at 00:00 UTC. It was a sunny day and the extensive low stratus cloud with a vertical extent of several hundred meters persisted over Nam Co Lake, and the initial sounding profile is shown in Fig. 3. The convective available potential energy (CAPE) was 392.94 Jkg-1, and the convective inhibition (CIN) was -44.01Jkg-1, indicating a favorable condition for convection initiation.

We initialized the simulations using a linearly increasing wind profile, in which the u-component of the wind increases from 0 ms-1 at the surface to 15 ms-1 at the model top. As noted earlier, such winds are maintained at the inflow boundary using Rayleigh damping. The simulations were integrated for 20 h, with model output saved every 5 min. Additionally, we included random 0.1 K potential-temperature perturbations in the initial conditions to seed turbulence.

The prescribed land–air thermal contrast represents the strong daytime surface heating over the TP and produces a positive sensible heat flux, which provides the thermal forcing for lake-mountain circulation development. Rather than reproducing the complete evolution of a specific convective event, this idealized configuration is designed to isolate the thermodynamic and dynamical mechanisms controlling convective initiation, organization, and precipitation development.

The simulations illustrate the pronounced lake and downstream terrain joint effect on the distribution and intensity of precipitation. Overall, during the whole simulation period, the accumulated precipitation exhibits a characteristic lake-mountain coupled pattern. (Fig. S2). Although largely disorganized, the striated nature of the precipitation pattern suggests weak organization by roll convection. These precipitation features resemble those observed in open-cell lake-effect convection (Sakakibara et al., 1988; Steenburgh and Campbell, 2017; Steiger et al., 2013).

The temporal pattern, quantified by interquartile ranges (25th–75th percentiles), reveals that lake-mountain interactions modulate the timing and variability of precipitation (Fig. 4). The lake-breeze circulation, together with the background flow, causes precipitation to develop preferentially over the lake and supports continued precipitation development beyond the shoreline in experiments involving the lake (Fig. 4). Precipitation appears firstly over the lake (Fig. 4a and b), and subsequently intensifies on the windward and leeward slopes (Fig. 4c–f), and finally shifts farther downstream (Fig. 4g and h), consistent with the direction of background flow. The LM_3000 configuration exhibits an early, mechanically driven precipitation, with peak values occurring within 2–4 h as topographic lifting rapidly enhances condensation and rainfall. In contrast, reducing the mountain height weakens the initial mechanical triggering, and the development of convection becomes more dependent on the gradual establishment of lake-induced moisture transport and circulation. The initiation of precipitation is postponed until 8–12 h, followed by a sharp and well-defined peak that is most pronounced over the lake and its downstream region, producing a much steeper temporal gradient in LM_1500. In OL, precipitation peaks even later (approximately 12 h) and is weaker, concentrated mainly over the lake and its downstream area. The LM_750 shares a similar onset time with LM_1500 but yields smaller precipitation amounts.

The differences in precipitation are broadly consistent with the relative timescales of thermally driven circulations under clear skies and weak synoptic forcing. Upslope wind can develop early as sloping terrain warms rapidly, promoting mechanically and thermally assisted ascent over the mountain (Barry, 2008; Serafin and Zardi, 2010; Whiteman, 2000). In contrast, lake breeze circulations require a more extended period of daytime heating to establish a sufficient land-water thermal contrast, delaying the initiation of convective activity (Arritt, 1993; Miller et al., 2003). Accordingly, we define the period of strongest overlap between the lake breeze and mountain circulation as the processes of synergistic interaction. Using this definition, we treat times after 7 h as the lake-mountain synergy regime, when cloud and precipitation processes are well developed and begin to influence the downstream.

During the lake-mountain synergy phase, precipitation fields exhibit distinct spatial distributions among the different configurations. In LM_1500, the combination effect of lake-induced moisture and the mechanically forced ascent on the windward slope produce stronger and more concentrated terrain-driven precipitation that extends upslope from the downwind foothills (Fig. 5a). Compared to LM_1500, increasing mountain height shifts the primary rainfall enhancement toward the lake area/windward side and limits precipitation expansion over the leeward, yielding a more confined precipitation footprint on the downstream side (Fig. 5b). By contrast, weaker terrain forcing reduces the vertical ascent, resulting in weaker and more spatially confined precipitation in LM_750 (Fig. 5c). When only the lake is present, precipitation remains largely confined to the lake and its immediate downwind vicinity and is more symmetric about lake-effect rainfall pattern (Fig. 5d). In the absence of the lake, precipitation appears as a narrow band locked to the terrain over the leeward, with limited downstream growth, indicating that background moisture is insufficient to sustain broader precipitation (Fig. 5e).

Stronger orographic ascent shifts and amplifies precipitation toward the lake area and the windward slope. In contrast, the lower mountain limits upslope lifting, leading to weaker and more spatially confined rainfall. Removing the mountain enhances precipitation over the lake while suppressing it over the terrain, whereas removing the lake produces a pronounced precipitation deficit over the former lake area and its downwind region. Overall, the combined effects of lake-supplied moisture and terrain ascent govern the intensity, timing, and spatial distribution of precipitation. Removing either component substantially reduces both the spatial coverage and the magnitude of rainfall.

To diagnose the moisture supply and convective instability in a high-altitude environment, we analyze vertical cross sections of water vapor mixing ratio and indicators of convection initiation for the experiments exhibiting distinct precipitation patterns (Fig. 6). In all configurations, near-surface humidity is generally low across the domain, consistent with the dry climatological background of the TP reported in previous studies (Maussion et al., 2011; You et al., 2015). In the NLM experiment, where both lake and mountain forcing are absent, the dry background environment is insufficient to support deep convection, and no measurable precipitation develops (Fig. S3). This implies that convective initiation is more sensitive to localized moisture enhancements and thermal perturbations. Against this background, surface heterogeneity (lake and adjacent terrain) produces pronounced contrasts in convection initiation.

In LM_1500, LM_3000, and LM_750, lake moistening combined with terrain-forced ascent produces comparatively favorable conditions for convective initiation. Over the lake, a deep moist layer develops between 90 and 150 km, where evaporation increases near-surface humidity and dewpoint, thereby lowering the lifting condensation level (LCL) and the level of free convection (LFC) over the lake relative to the surrounding land (Fig. 6a, c, and e). Over the TP, reduced air density often favors a larger kinematic buoyancy flux and a deeper mixed layer, which facilitates vertical moisture transport to support convection initiation (Wang et al., 2020b; Xu et al., 2023). Stronger mixing promotes parcel ascent to the LCL, and subsequent latent heat release helps further into the free convection layer (lower the LCL). Consistent with these changes, CAPE increases and CIN decreases, creating a less inhibited environment for deep convective initiation over the lake areas (Fig. 6b, d, and f). Notably, in these simulations, CAPE remains modest even when initiation thresholds are low. Although enhanced moisture supply over the lake lowers the LFC and weakens CIN, CAPE accumulation is still constrained by the thermodynamic background of the TP. Convection may initiate under relatively weak inhibition, but sustained deep development is less favorable than in warm and moist low-elevation regions. In plain environments, numerous studies have shown that substantial CAPE can accumulate beneath a strong capping inversion and be released quickly once the cap is breached (Chen et al., 2020a; Emanuel, 2023). Where the LCL-LFC separation remains large, parcels are less likely to reach free convection without additional lifting, implying a greater dependence on windward moisture advection and terrain-forced lifting around 150 km. It should be noted that as mountain height increases, CAPE over the lake region decreases and CIN strengthens (Fig. 6b, d, and f), whereas precipitation over the lake exhibits the opposite tendency (Fig. 5a–c).

In OL, the lake enhances low-level moisture directly, lowering LCL and LFC over the water surface (Fig. 6g and h). However, the absence of orographic lifting inhibits forcing ascent. Although CAPE is elevated over the lake, CIN remains substantial downwind (near 150 km) and the LFC stays relatively high, indicating suppressed convective development away from the lake. In OM, the near-surface layer is drier and the LFC is higher (Fig. 6i and j). CAPE is weak to nearly absent across much of the section, while CIN remains strong. Although terrain-forced ascent occurs near the crest, the conditions remain unfavorable for deep, sustained convection. On the leeward slope (around 180 km), CAPE decreases while CIN increases, indicating a more stable lee environment. Rolling ascent associated with trapped lee waves may nevertheless locally erode inhibition and support the development of convection and produce rain shadow regions (Bai et al., 2021; Zängl, 2005), consistent with the precipitation signal on the leeward slope (Fig. 5a and e). Over the TP, lake moistening alone favors lower LCL but provides limited triggering, whereas terrain forcing without lake moisture yields strong inhibition and weak instability. The interplay between lake-sourced moisture and the downstream orography is critical for modulating convective initiation.

Having established the thermodynamic preconditions for initiation, we next examine convective growth and cloud formation using the hydrometeor fields (Fig. 7). The LM_1500 exhibits the deepest convective development, with an enhanced total water column extending above 4 km and collocated with the strongest updraft core (Fig. 7a). Hydrometeors are mainly concentrated near and above the LCL, indicating that near-surface layer water vapor supplied by the lake plays a key role in promoting condensation (Fig. 6a). Notably, the isentropes (θ) rise sharply and fold over the mountain crest, indicating strong diabatic heating within the convective column. When topographic lifting is present, vertical motions reach sufficient magnitude and depth to overcome CIN and support deep convective updrafts capable of initiating precipitation (Figs. 7a–c and e).

Contrasts among configurations clarify the distinct roles of moisture supply and dynamical triggering. With the lake retained and a higher mountain in LM_3000, ascent intensifies over the windward slope, producing more hydrometeors, but the condensate column is narrower and less laterally extensive than in LM_1500 (Fig. 7b). Consistent with the thermodynamic differences diagnosed in the previous section, CAPE over the lake sector is reduced as mountain height increases, and condensate loading over the lake correspondingly decreases, whereas condensate becomes more concentrated near the crest where orographic ascent is strongest. With a lower mountain (LM_750), cloud development is shallower and hydrometeor mixing ratio is reduced (Fig. 7c), consistent with weaker mechanical forcing that makes it more difficult to reach the LFC. In OL, low-level moistening together with the lake-land thermal contrast forcing generates a broad and shallow convective response (Fig. 7d). Although relatively strong vertical motion develops over the lake, the absence of terrain-forced lifting limits further convective development. In contrast, hydrometeor in OM is only caused by the lee wave, indicating strong moisture limitation in the absence of the lake, with convection development supported only by the background flow (Fig. 7e). Overall, these configurations indicate that deep convection is favored when sufficient moisture supports is lifted by suitable orographic forcing, promoting efficient condensation and sustained vertical development.

To further characterize the evolution of convective growth and condensate production, we examine the temporal evolution of the vertical profiles of total non-rain hydrometeor content (qhc=qc+qi+qs+qg) and rain water mixing ratio qr, averaged over the lake, mountain, and downstream subregions (Fig. 8). Across the configurations that include the lake, low-level moisture promotes early hydrometeor formation over the lake, but the subsequent vertical deepening and conversion to rainwater vary markedly. The qhc occur primarily above the freezing level (1.4–1.5 km; not shown), indicating that cloud condensate mass accumulates through efficient ice-phase growth (e.g., riming and aggregation) within the mixed-phase layer. As these hydrometeors descend through the freezing level, melting increases qr, so rainwater maxima occur mainly below the freezing level. This vertical structure suggests that precipitation is primarily produced through ice-crystal processes of precipitation, with hydrometeors growing aloft and melting as they fall through the melting layer, rather than through warm-rain processes (Gupta et al., 2023).

In LM_1500, hydrometeor content develops rapidly over the lake, and qr increases sharply between 8–12 h (Fig. 8a). Over the mountain, the hydrometeor column is deeper and more persistent (Fig. 8b). Trapped lee waves promote enhanced production of qhc aloft and a more efficient conversion to qr. The downstream response subsequently strengthens (16 h; Fig. 8c), consistent with the timing of precipitation shown in Fig. 4b. In LM_3000, hydrometeor growth and rainwater production begin earlier, but the lake-region hydrometeor column is narrower and less vertically developed (Fig. 8d), consistent with stronger inhibition (higher LFC and larger CIN) limiting deepening. Over the mountain, both qhc and qr maximize at low levels, producing intense orographically forced precipitation (Fig. 8e). In LM_750, reduced terrain height weakens lifting; qhc increases over both the lake and mountain, but qr remains comparatively weak (Figs. 8g–i). In OL, hydrometeors persist over the lake and its downwind region for most of the simulation period, with their evolution lagging behind that in the other configurations (Figs. 8j–l). Lake-breeze forcing is able to maintain cloudiness, but in the absence of strong dynamical lifting it is inefficient at rapidly converting condensate into qr. Both qhc and qr remain low across all subregions (Figs. 8m–o), indicating that moisture limitation is pronounced in OM.

Downstream hydrometeor amounts are generally lower than those over the lake and mountain in all configurations, and the downstream precipitation signal primarily reflects the evolution of upstream systems rather than strong local production. Overall, earlier and stronger rainfall occurs when lake moistening coincides with terrain-induced lifting, consistent with the precipitation evolution shown in Fig. 4.

To diagnose how lake-terrain forcing regulates condensate production and its conversion to surface precipitation over the TP, Fig. 9 shows the distributions of hydrometeor mass growth (Qgrowth), hydrometeor mass loss (Qloss), and the surface precipitation mass rate (Pr).

Across all lake-present experiments, the first Qgrowth peak occurs on the downwind side of the lake area (80–110 km), further proving that near-surface moistening promotes convection initiation (Fig. 9a). In the same sector, the Qloss and Pr also exhibit peak (Fig. 9b and c), indicating efficient conversion of hydrometeors to precipitation. Only LM_1500 and LM_3000 exhibit a second, stronger paired maximum of Qgrowth and Qloss in the windward slope (145–175 km), collocated with a peak in Pr. However, in LM_750 and OL, the absence or weakness of terrain-induced lifting results in the suppression of precipitation formation. Along the leeward slope, the configurations with the same mountain height (LM_1500 and OM) exhibit pronounced differences in Qgrowth and Qloss, but their maximum Pr are comparable. This indicates that forced ascent associated with trapped lee wave can substantially enhance precipitation release even under limited moisture.

Over the TP, the distinctive thermodynamic background constrains the magnitude and distribution of hydrometeor and finally shapes the precipitation pattern (Fu et al., 2020; Xu et al., 2023). In previous studies over the low-elevation regions using an open water configuration comparable to OL, Qgrowth and Pr typically exceed Qloss by about one order of magnitude (Gowan et al., 2021), reflecting the more moist near-surface layer and lower saturation deficit. In contrast, over the TP, Qgrowth, Qloss, and Pr are of comparable magnitude and all smaller than the low-elevation regions, indicating the condensate budget that is limited by moisture and convection initiation. Specifically, the low moisture condition and elevated lifting condensation levels over the TP, together with stronger entrainment of dry air, limit condensate production, thereby delaying saturation and suppressing the development of deep convection (Li et al., 2022a; Wang et al., 2020b). As a result, even under comparable dynamical forcing, Qgrowth increases more slowly than over the low-elevation regions. Notably, despite the weaker Qgrowth, the efficiency of converting condensate into Pr does not decrease with altitude. Mixed-phase processes are prevalent over the TP (as shown in Fig. 8), and ice-phase growth processes can efficiently generate precipitation (Gao et al., 2018). These processes partly compensate for the relatively low Qgrowth and high Qloss, yielding Pr comparable to that over the low-elevation regions. In this setting, convective clouds tend to be short-lived and spatially limited, with frequent precipitation occurrence but typically weaker event intensity, consistent with the constrained moisture supply and inhibited deep maintenance over the TP (Chen et al., 2019; Fu et al., 2020; Singh and Nakamura, 2009; Ueno et al., 2007).

To further assess convective triggering under different surface configurations, Fig. 10 illustrates vertical cross-sections of potential temperature perturbation (θ′) and wind perturbation. These fields highlight localized circulation and thermodynamic anomalies and are used to relate dynamical forcing to the corresponding cloud development.

In LM_1500, the upstream of the mountain (60–150 km) exhibits a clear vertical contrast in θ′ (Fig. 10a). Near the windward foot, terrain blocking promotes low-level cold-air pooling and a thicker stable layer, yielding negative θ′ within low-level layer. Above this layer, positive θ′ develops as convection over the lake intensifies, consistent with diabatic warming associated with sustained condensation and deposition. At the same time, a deep ascending branch is collocated with the windward slope and extends above 6 km. Above the mountain, a cold core (θ′<-0.5K) associated with lifted isentropes is collocated with the condensate region (Fig. 7a), indicating that ascent and adiabatic cooling over the leeward support hydrometeor growth and downstream transport. In LM_3000, stronger orographic lifting leads to larger cold and warm θ′ perturbations above the mountain (Fig. 10b), collocated with enhanced hydrometeor mixing ratio and precipitation (Figs. 7b and 5b). The enhanced precipitation produces stronger diabatic cooling through the evaporation of precipitation and phase-change processes, contributing to the development of negative near-surface temperature perturbations. Meanwhile, stronger subsidence warming on the lee slope limits the downstream extension of low-level moist flow. The terrain-induced blocking effect further suppresses the horizontal ventilation of the cooled air, allowing the precipitation-induced cold pool to accumulate and persist near the windward slope. Some hydrometeors are transported back toward the lake region, where precipitation modifies the local thermodynamic structure and weaken the warm potential temperature anomaly above the lake (Fig. 8d). In LM_750, θ′ is weaker and ascent is shallower, consistent with a lower and more diffuse hydrometeor distribution (Fig. 7c).

Vertical transport is shallow over the lake, with weak θ′ confined below 3–4 km in OL (Fig. 10d). The absence of sharp vertical motion or strong θ′ limits the development of deep convective clouds. The shallow hydrometeors are transported to the downstream driven by the background flow (Fig. 7d). In OM, the intense windward updraft and strong warm core are evident above the mountain, yet hydrometeor remains limited in extent (Fig. 7e), indicating that strong lifting cannot sustain the development of deep convection cloud under moisture-limited conditions. Overall, the synergy between low-level instability supplied by the lake and terrain lifting favors the initiation and development of deep convection.

To connect the dynamic structure diagnosed in the different configurations with the downstream evolution of clouds and precipitation, we further quantify the moist dynamical regime using the moist flow-stability index (Imfs) (Fig. 11). For relatively small upstream values of Imfs, the background flow tends to be blocked by the topographic barrier, consistent with the low-level perturbation zonal wind structure shown in Fig. 10. When the lake is included, enhanced moisture availability modifies the moist stability through changes in the vertical distribution of water substances, while surface thermal effects also contribute to the adjustment of atmospheric stratification. These combined effects reduce Nm, and increase Imfs, favoring stronger moisture transport toward the mountain.

In LM_1500, the strongest and most vertically extensive upstream Imfsanomalies indicate that the combined lake and terrain act to weaken moist stability and to enhance the dynamical lifting (Fig. 11a). This pattern supports the development of convection and sustained hydrometeor production. However, increasing the terrain height cause dynamical competition (Fig. 11b). While the stronger mechanical lifting provides more kinetic energy to overcome the potential energy of the topographic barrier, the larger H in Imfs exhibits more blocked effect near the mountain. This moist dynamic response tends to concentrate convection and precipitation on the mountain while promoting the downstream drying tendency. In LM_750 (Fig. 11c), the near-surface moisture persists, but the lifting is weaker and less vertically extended. Accordingly, this configuration is less effective at transporting low-level moisture to the development of deep convection compared with the LM_1500. In OL, Imfs only serves as a compact diagnostic of U/Nm, the near-surface moisture reduces effective moist stability and promotes shallow lifting even in the absence of mechanical forcing (Fig. 11d). In OM, where the low-level moisture supply is absent, convective initiation is restricted and the development of deep convection is suppressed despite favorable orographic lifting (Fig. 11e).

To diagnose how lake-terrain forcing regulates the pathways by which near-surface moisture supply affects convection, Fig. 12 shows forward trajectories calculated using a Eulerian wind field trajectory approach for air parcels released in the upstream near-surface layer (z=50m, x=80km) for the different configurations. (Forward trajectories were computed from the model output using an open-source Python package developed by Gowan, 2019). Although water vapor can generate phase changes, our purpose is not to construct a closed moisture budget, but rather to diagnose the transport pathways and lifting of near-surface moisture from the lake toward the terrain and downstream, thereby isolating the thermodynamic and dynamical linkages between moisture transport and orographic triggering for convective development.

In LM_1500, lake-supplied tracer parcels lift and spread laterally to the downstream under the background flow (Fig. 12a). Terrain forcing organizes a coherent upslope ascent branch along the windward slope (Fig. 12b), transporting moisture from the mountain foot to levels more favorable for cloud growth and precipitation formation. On the leeward slope, the trapped lee wave produces two downstream branches: one pathway is lifted by the ascending phase of the lee wave and advected downstream aloft, while the other follows the descending phase and channels moisture downstream within the lower levels. In LM_3000, stronger orographic lifting causes more moisture to be retained near the mountain front rather than transported farther downstream (Fig. 12c and d). Trajectories released closer to the terrain also show a clearer blocking for the higher mountain (Fig. S4). Compared with LM_1500 and LM_3000, LM_750 still exhibits moisture transport, but only shallow lifting occurs, and fewer trajectories reach higher levels (Fig. 12e and f). In OL, a moist tongue develops over the lake and its downwind areas. Most trajectories remain within this humid zone and are transported primarily horizontally with modest lateral spreading (Fig. 12g). Vertical transport is weak, with trajectories largely confined to low levels and lacking sustained deep ascent (Fig. 12h). In OM, trajectories show a narrow ascent near the terrain, but moisture content is lower than in the experiments involving lake.