1300 1600 1900 2200 0100 0400 0700 1000
hhmm (LST)
1300 -200,00
0,00 200,00 400,00 600,00 800,00
Rn H LE G
W·m-2
Figure 5.11: The heat energy fluxes in W·m−2 with the net radiation (Rn), the sensible heat flux (H), the latent heat flux (LE) and the ground heat flux (G) as a function of time between 12 October 1300 LST and 13 October 1300 LST taken at -78.5◦ x -5.2◦ the associated begin of radiation loss. After a maximum peak of 100 W·m−2 at 1800 LST, the radiation loss achieves nearly constant values of 70 W·m−2 during the night. The surface fluxes change their direction correspondingly. As a result of the heating effect during the day, the sensible heat flux (H) has positive values but is reduced to almost zero at 1900 LST. The latent heat flux (LE) reduces as well, but keeps clearly positive until 0200 LST indicating a flux from the surface to the atmosphere due to evaporation. The ground heat flux (G) reach negative values as a result of the temperature deficit between the atmosphere and the surface layer. Therefore, an energy flux from the PBL to the surface layer exists, which is caused by a constantly cooling atmospheric boundary layer as a result of a negative radiation budget which is typical for nighttime. This can lead to the development of a surface inversion, as already indicated by the amount of negative buoyancy, which induces a positive temperature gradient. The surface heat fluxes indicate the development of thermally-induced katabatic flows, which are associated with the characteristic features of the downslope flow (figure 5.10).
5.5 Discussion
As expected, the nocturnal formation of the presented convective cloud south-east of Ecuador in the Peruvian Amazon basin is affected by well defined and strong downslope winds from the Andean slopes and valleys. Those katabatic flows induce a surface cold front at the foothills of the Andes in the Amazon basin. The special terrain configuration in the target area can cause the formation of a quasi-concave
5 Nocturnal Convective Clouds at the Eastern Andes of South Ecuador
terrain line with strong convergences. The underlying processes resulting in the formation of a cold cloud are examined using a case study (12 October to 13 October 2009) diagnosed by GOES-E satellite images. This outlines the impact of the terrain and, consequently, of the katabatic flows on the cloud formation: the 4 km domain (D3) reveals the absence of any cloudiness in the target area at nighttime, but in the inner domain (D4) a convective cloud cluster was generated (figure 5.2 and 5.3).
The atmospheric stability of both domains regarding the frontal zone and the thermodynamic structure was examined. The development of the katabatic induced surface front was described by the horizontal distribution of the equivalent-potential temperature (see figure 5.5). They show similar structures regarding the evolution of a horizontal gradient ofθe in the cloud formation area. But even here, the impact of terrain and, subsequently, the downslope winds was visible. The cold drainage air in D4 propagated further into the basin, indicating a stronger flow.
A cross section through the front revealed first distinctions (see figure 5.6). While D4 demonstrated the vertical displacement of warm-moist air in the target area in one time step while the brightness temperature showed cold cloud patterns, D3 did not offer such structures. Although D3 has a potential instability and areas of high θe (354 K), which are sufficient for a destabilisation of the atmosphere (Machado et al., 2002), convective activities are completely absent.
An inspection of the environmental parameters gave more information concerning the reasons for the initiation and lacking convection, respectively (see figure 5.7).
Both domains are marginally unstable, with a great amount of negative buoyancy in the lower levels. However, they also show a good potential for the development of a thunderstorm on the basis of the LI and K-Index (see table 5.3). The LFC in D3 and D4 are significantly different. In D3, the LFC for an ascending air parcel is reached at 3359 m, which is 1000 m higher than in D4. In association with the great amount of convective inhibition, which must be overcome first, a strong trigger function must be present. This trigger mechanism is assumed to be the convergence of the katabatic flows from the Andean mountains in connection with the Amazonian warm-moist air. As well, D4 demonstrated stronger convergence patterns indicating the initiation of sufficient vertical motions (see figure 5.8). In this process, the most accurate representation of the terrain, the surface fluxes as well and the resulting katabatic flows are important features.
Laurent et al. (2004) described that MCS initiation in South America is mostly driven by topography. Calbo and Millan (1998) also show the role of the terrain and highlight its influence on the model grid size. The authors demonstrated the significant impact of horizontal resolution on the vertical structure of the model do-main. They ascertain that high resolved simulations increase the ability to develop stronger vertical velocities and that the vertical fluxes are more strongly developed.
The vertical fluxes can result in modifications in the surface temperature and humid-ity field, which are sensitive to thermodynamic activities (Mapes and Houze, 1992).
These surface fluxes are also the main driver for the development of thermally
in-5.5 Discussion
duced katabatic flows. However, the vertical resolution essential for the evolution of the flows of both domains (D3 and D4) is identical, but the horizontal grid size differs. By increasing the horizontal resolution, as in domain D4, the complexity of the digital elevation model (DEM) increases as well. Thus, the terrain is more structured and the concave shape of the Andes in the target area is better revealed.
The katabatic flows are stronger and the convergence in the target area increases.
The result is a compressional lifting with markedly high vertical velocities due to a larger increase in pressure. In D4 sufficiently dynamic vertical motions occur, whereby the ascending air parcels reach the LFC. In contrast, the 4 km simulation generates weaker horizontal convergences with consequently ascending air parcels, which fail to reach theirLFC (see figure 5.9). Therefore, the reason for the missing nocturnal cloudiness in domain D3 is that, in the current situation, the lower atmo-sphere requires a strong trigger to induce deep convection and the formation of a mesoscale cloud cluster. The chosen resolution of D3 is insufficient for this initiation mechanism, since it less represents the terrain, and thus, the katabatic flows and the increase in mass, which are the main driver.
In domain D4 ARPS underestimates the development of the generated convective cloud cluster, since the simulated cluster is smaller than the observed one (see fig-ures 5.2, 5.3 and 5.4). The simulated brightness temperatfig-ures at the beginning of the night (2115 LST -0115 LST) showed two distinguishable systems with low IR temperatures (220 K) emerging at two separated regions different from the GOES-E data. But in these areas the convergence patterns, as a direct responds to the terrain (figure 5.8b), are strongest causing the convective initiation. Unlike the observed data, the simulated data do not merge to one convective system. The reason is the lack of a stronger upper level wind field (see figure 5.10) reaching values of approx-imately 1.5 m·s−1. However, new cells emerged at 0215 LST, which are more equal to the observed ones. They consisted of several organised cells within a cloud shield.
These different occurring cells suggests two different initiation phases: one being the convergence due to the katabatic flows and the other in the second half of the night a cell regeneration due to the circulation dynamics of the convective cluster.
Two different mechanisms of storm initiation in South America was also mentioned by Valesco and Fritsch (1987), although the authors describe one in the afternoon and one nocturnal.
Aside from this deviation, there are several conformances. As in the observed satellite data, the cloud cluster develops in the evening, grows during the night and dissipates in the early morning hours. The cells gain similar brightness temperatures and exhibit typical interior areas with lower values (below 220 K) as well. Even if the location is not exactly collocated with the observed data, the orientation of the cells are similar to the GOES-E data and, the error is marginal in comparison to the cloud physics. So for these conditions there is a good agreement between the satellite images and the modeled data.
5 Nocturnal Convective Clouds at the Eastern Andes of South Ecuador
A closer look at the development of the cloud cluster demonstrates the influence of the surface convergence by an intermittent influx of cold moist air, indicating drainage flow (figure 5.9). An analysis of an atmospheric profile regarding wind velocity, theTKE and heat exchange (figure 5.10 and 5.11) contribute to this result.
Although the inversion layer is deeper than in a previous study concerning the confluence of katabatic flows (Trachte et al., 2010), it shows the specific jet-like profile andTKE characteristics. The reason for the deviation of the katabatic layer is its dependency on the vertical resolution; in this study, a coarser one was used.
The idea that a katabatic flow develops is further supported by the surface fluxes representing diabatic heating. The results in figure 5.11 disclose a typical nocturnal course with net radiation loss and a corresponding energy flux from the PBL to the surface layer. It induces a buoyancy deficit, which is the driving force of the downslope flow.
Based on these results, and due to the fact that the 1 km domain (D4) used no cumulus parametrisation scheme (see table 5.1), a scale dependency for the initiation of these convective activities is assumed. Considering that katabatic flows are small-scale features, the demonstrated simulation results lead to the conclusion that they play a major role in the activation of convection. However, the thermally-induced downslope winds act only on the lower levels of the atmosphere in the PBL. They cause the formation of a surface inversion, which produces negative buoyancy (-146 J·kg−1 for D4). CIN, thus obtained in the PBL, has to be overcome first to reach theLFC with subsequent formation of cold clouds. Therefore, an additional trigger mechanism must be available for strong and sufficient lifting of air. In the light of the concave Andean ridge line south-east of Ecuador, which forms an extensive drainage system directed into the Amazon basin, this specific terrain geometry is extremely important. Particularly the confluence of the cold drainage of air due to the shape of the terrain results in a compressional lifting, which effectively removes the negative buoyancy. In this process, a destabilisation of the atmosphere is caused by the interaction with the warm-moist air of the Amazon, causing deep convection as shown in figure 5.9 and a previous idealised study (Trachte et al. submitted).