Effect of Topography on the Propagation of Extratropical Cyclones over the Rocky Mountains

Size: px

Start display at page:

Download "Effect of Topography on the Propagation of Extratropical Cyclones over the Rocky Mountains"

Mercy Porter
5 years ago
Views:

1 University of Alabama in Huntsville ATS740 Final Project Effect of Topography on the Propagation of Extratropical Cyclones over the Rocky Mountains Author: Brian Freitag December 9, 2013

2 1. Introduction In addition to monitoring changes in global temperature and sea level, shifts in precipitation patterns are important for understanding and preparing for global climate change in the future. In the mid-latitude belt, shifts in precipitation patterns are typically understood by modeling the variation of extratropical cyclone tracks through time. Previous modeling studies for extratropical cyclone tracks have removed model results over regions where topography exceeds 1000 m and these models have grids in excess of 1.5 resolution (Wang et al., 2006; Ulbrich et al., 2008; Gulev, 2001). At this resolution, mesoscale processses affecting the propagation of cyclones over topography cannot be fully resolved. Topographical influences on extratropical cyclone propagation include synoptic, mesoscale, and microscale forcings that intiate a spin down on the windward side of the topography and cyclogenesis on the leeward side. On the synoptic scale, vertical motion is initiated by an increase in the surface roughness length and the drag coefficient which alters the horizontal and vertical momentum fluxes. Enhanced vertical motion on the windward side initiates negative vertical vorticity as the flow is deflected to the right (in the Northern Hemisphere) by the Coriolis Force. The initiation of a secondary circulation opposite the mean flow is called Ekman pumping; which, transports the opposing momentum flux vertically through the column (Holton and Hakim, 2013). The spin-down effect generated by Ekman pumping reduces the surface moisture flux, thereby reducing the vapor content in the atmosphere and precipitation within the storm. On the leeward side, there is divergent flow and a stretching of vorticity as the mean flow descends the topography. This results in a positive vorticity anomaly on the leeward side, initiating cyclogenesis. Additionally, the sinking flow accelerates as it descends the mountain because the descending air is colder and more dense than the surrounding air. The accelerated flow descending the topography is

3 known as the katabatic winds. The katabatic winds feed into the stretched vorticity column, thereby enhancing the cyclogenesis. The forcings observed on the synoptic scale for the evolution of extratropical cyclones are a result of mesoscale and microscale processes occurring at the surface (Holton and Hakim, 2013 ; Cotton et al., 2011). As the cyclone passes over the topography it goes through several stages as defined by Cotton et al. (2011). The most important of these stages with respect to mesoscale processes is the late sector; which, is the cold unstable zone behind the cold front. This region is where the enhanced turbulence alters the horizontal momentum flux and increases vertical motion leading to multiple, smaller scale, orographically forced cores of enhanced condensate loading. According to Cotton et al. (2011), the cells observed in the late sector are typically broadened on the windward side because of orographic lifting. The lee cyclogenesis is driven by baroclinic vorticity generation associated with the descending cold air off the terrain. Along the Rockies, ahead of the cold front, the cyclonic flow initiates slantwise instability as warm air associated with enhanced moisture from the Gulf of Mexico is loftedat an incline and wrapped around the center of rotation (Cotton et al., 2011; Markowski and Richardson, 2010). For this project, the microphysical processes associated with small scale turbulent forcings that affect extratropical cyclone propagation are not discussed or considered because of the limitations of model resolution. 2. Previous Research Because many of the features that alter extratropical cyclone flow are not provided from in situ meteorological data, atmospheric models are used to gain a greater understanding of the physical processes. There have been several studies done on the external forcings for cyclones crossing the Alps in Europe. In these

4 cases, similar to the Rocky Mountains, flow splitting occurs because of the vertical and horizontal scales of the obstruction are such that the main flow dissipates into multiple separate flow features as discussed above. Previous modeling studies have shown that flow splitting occurs on the windward side and rapid intensification associated with strong vertical motion on the leeward side before the cold front reaches the lee cyclone ( Mattocks and Bleck, 1986 ; Aebischer and Schar, 1998). In addition, baroclinic vorticity generated from down-slope winds strengthens the existing lee cyclone further promoting cyclogenesis (Mattocks and Bleck, 1986). Enhanced katabatic winds are generated by breaking gravity waves associated with the topographical features. This leads to the development of perturbation potential vorticity fields which are advected down-slope promoting lee cyclogenesis. Modeling studies have also focused on the physical processes initiating spindown on the windward side of topography. Beare et al. (2007) modeled the potential vorticity and velocity fields with changing boundary layer conditions to analyze extratropical cyclone characteristics. The study found that Ekman pumping led to vortex squashing and synoptic scale cyclolysis, but potenetial vorticity fields needed to be inverted in the warm sector to see a deepening of the negative vorticity anomaly. These boundary layer forcings shift the track of the cyclone to the South as it traverses the topography as seen in Brayshaw et al. (2009) as well. This study will analyze the shift in the cyclone track to some degree, but will mostly focus on the physical processes affecting the cylcone propagation. 3. Methods a. Ocean Land Atmosphere Model The Ocean Land Atmosphere Model (OLAM) is a non-hydrostatic, unstructured grid model that allows for specification of local grid refinement constrained to a user defined region of arbitrary shape. OLAM utilizes a shaved cell method

in the vertical, allowing for grid cells that intersect terrain. In this project, regional grid refinement will be applied over the Rocky Mountains in Colorado.

5 in the vertical, allowing for grid cells that intersect terrain. In this project, regional grid refinement will be applied over the Rocky Mountains in Colorado. The OLAM grid is centered over the midway point between the windward cyclolysis and the leeward cyclogenesis, more specifically, N and W near Aspen, Colorado. The OLAM gridmesh used for this project plotted over the local topography, can be seen in Figure 1. Figure 1: OLAM gridmesh overplotting topography for Western Colorado. The grid structure for OLAM consists of eight horizontal, hexagonal grids scaling from a global grid down to a 100 x 100 km domain with 1 km resolution centered over the latitude and longitude defined above. This grid structure will allow for the model to identify and resolve mesoscale processes that lead to the

6 windward cyclolysis and lee cyclogenesis to be analyzed in this project. While OLAM has a global domain and thus, does not require nudging, it does have the capability for nudging using external analysis. One of the reasons for utilizing OLAM in this study is to eliminate uncertainties associated with lateral boundaries in limited area models; therefore, analysis nudging will not be utilized. OLAM calculates surface data at a 5-minute interval and outputs a history file for plotting once every hour for the duration of the model run. OLAM was run for 72 hours for each of the three events and was initialized using reanalysis data. b. Data Collection Global reanalysis data from the NCEP Climate Forecast System (CFSR) is used to initialize OLAM at 6 hour intervals. The CFSR is a coupled ocean-landatmosphere surface-sea ice system that uses assimilation of satellite radiances to provide an instantaneous atmospheric state. According the the NOMADS website (N.D.), the CFSR global atmosphere resolution is 38 km with 64 vertical levels extending from the surface to 0.26 hp a. Analysis data was used for a three-day period for the duration of the model runs, further discussed in the next paragraph. c. Events For this study, three events of extratropical cyclones crossing the Rocky Mountains were chosen using UNISYS surface and upper air weather charts. The three events analyzed were an event from January, 1996, one from April, 2003, and another from December, In these cases the surface weather charts show that the extratropical cyclone enters the Pacific Northwest and approaches the Rocky Mountains near the Colorado/Wyoming Border. As the center of circulation interacts with the terrain there is a synoptic scale spin down, followed by cyclogenesis to the south of this intersection point, east of the topography several hours later. The goal of this project is to provide an analysis of the physical forcings acting on the extratropical cyclone and to compare the modeled results

7 to previous research. Surface plots for the April 2003 and December 2010 event are presented in Figure 2. The UNISYS plot for the January 1996 event did not accurately portray the synoptic setup and therefore, is not presented in this figure. Figure 2: Surface plots for the April 2003 and the December 2010 events (UNISYS, 2013). 4. Results The model results for the extratropical cyclone tracks in each of the three cases was fairly consistent with the UNISYS weather plots for two of the three cases used. When comparing each of the events, the model showed that each event provided different results when looking at the vertical velocity, cloudwater spectral width, and perturbation density fields. The vertical velocity and cloudwater spectral width were analyzed on the windward side of the topography and the perturbation density field was analyzed on both the windward and leeward sides. Plots for the January 1996 event as it traversed the Rocky Mountains can be seen in Figure 3.

8 Figure 3: Horizontal Cross sections of vertical velocity (left), condensate spectral density (center) and perturbation density (right) for the 09 UTC on 17 January In this case, there is evidence of condensate loading on the windward side of the highest peaks on the left hand side of the center plot as was seen in Cotton et al. (2011). This is driven by the upward vertical motion plumes shaded in blue on the far left plot. When comparing these two images, it is pretty clear that the enhanced cloudwater spectral density cores occur in the regions where there is upward vertical motion. The enhanced vertical motion and condensate loading is evidence of the initiation of cyclolysis forced by increased surface roughness. The perturbation density field also provides insight into the forcing mechanisms for lee cyclogenesis and windward Ekman pumping. The perturbation density field for this event as the extratropical cyclone crosses the topography shows negative density perturbations on either side of the topography with a positive density perturbation in over the highest peaks. This scenario promotes air parcels on the windward side to rise vertically, which, when coupled with the Coriolis effect is the trigger for Ekman pumping. On the leeward side a tongue of higher density

9 perturbations can be seen riding down the sloping topography. This is representative of the baroclinic forcing that initiates lee cyclogenesis as discussed above. In this case, the higher density air is carried by the katabatic winds down the slope of the terrain; which, lifts less dense parcels vertically upward as is typically seen along a cold frontal boundary. The results from the 1996 study are very well aligned with the results from previous research discussed above. The other two cases provided less common representations of topographical influences on these three fields. The plots for the April 2003 display these values at the approximate time the extratropical cyclone passed over the Rocky Mountains and can be seen in Figure 4. Figure 4: Horizontal Cross sections of vertical velocity (left), condensate spectral density (center) and perturbation density (right) for the 02 UTC on 03 April At first glance, it is fairly obvious that this event produced a much weaker signal at the surface compared to the 1996 event discussed above. While the surface winds still show updraft plumes along the windward side, the amount of condensate loading is moderately reduced on the windward side; however, condensate loading

10 also occurs in conjunction with the initiation of downslope katabatic wind flows. This can be explained by two possible scenarios. The first explanation is that the storm could have moved to the North or South of the grid center reducing the surface signatures seen in the January 1996 event. The second explanation is that the atmospheric low pressure system was a stronger 500-mb feature compared to the surface low suggesting an upper level disturbance. The perturbation pressure density field was analyzed in this case as well to look for the presence of surface forcings leading to lee cyclogenesis. The perturbation density plot shows no signature of the katabatic winds pulling higher density air down the front of the mountain slope initiating vertical motion. In fact, the entire domain below 6 km is a positive perturbation density which is expected given the vertical profile of density through the troposphere. The only evidence of topographically influenced flow on this plot is the dip in the negative perturbation density above the highest peak ( 3.3 km). Speculatively, this strong perturbation density gradient between 5 7 km could be evidence of the upper level disturbance being enhanced by the elevated peaks. The 2010 event provided a completely different result than both the 1996 and 2003 events discussed above. The main difference between this case compared to the others is the perturbation density field. The plots for this event were selected based on the time when the greastest condensate spectral density was observed and they can be seen in Figure 5 below. The cloudwater spectral density signature for this event was even weaker than the 2003 event shown by the two very small and narrow plumes of condensate in the center of the image. This is also represented in the profile of vertical wind with height. The plumes of upward vertical motion are not as defined and tend to have greater curvature than the plumes in the other two cases. The curved vertical wind plumes do not promote the lofting of condensate higher into the atmosphere. The

Figure 5: Horizontal Cross sections of vertical velocity (left), condensate spectral density (center) and perturbation density (right) for the 06 UTC on 30 December 2010.

11 Figure 5: Horizontal Cross sections of vertical velocity (left), condensate spectral density (center) and perturbation density (right) for the 06 UTC on 30 December katabatic wind flow is still evident in the velocity plot and even shows the initiation of a linear vertical plume at the base of the topography; however, the perturbation density field does not appear to be the driving factor. The perturbation density field does not resemble either of the previous plots and really doesn t resemble what we would expect for the passage of an extratropical cyclone. The perturbation density field shows two positive perturbations on the far east and west side of the plot with no terrain influences present. The katabatic winds appear to be forcing negative perturbation density (less dense air) under a region with relatively higher density, which would induce a downward vertical velocity and subsidence. Thus, there is no evidence in this plot of lee cyclogenesis being forced by baroclinic vorticity generation from the advection of perturbation by katabatic winds.

12 5. Conclusions The different synoptic set up and propagation properties for each of these events provided three unique and very different case studies. Previous research suggested condensate loading and enhanced vertical motion on the windward side of mountains and katabatic wind driven lee cyclogenesis. The January 1996 event aligned very well with the findings of the previous research in almost every aspect. There was evidence of the condensate loading forced by enhanced vertical motion on the windward side and density driven vertical motion on the leeward side forced by descending katabatic winds. The other events did not present nearly as well as this case. The April 2003 case consisted of a weak extratropical cyclone interacting with the topography north of the center model grid point. Once the cyclone encounters the topography, the initial cyclone completely dissipates as seen in the left hand side of Figure 2. Since this cyclone had a weaker response in the model output, the model actually handled this case well because it suggested weaker vertical motions and less condensate loading. While the surface low pressure system dissipated, the upper level disturbance maintained and actually spun up a weak lee cyclone that also dissipated once the upper level disturbance moved away from the topography. While surface plots suggest the presence of a surface low pressure system for the December 2010 case, the model output did not seem to include much evidence supporting its presence. Whether it s an inaccuracy with the model or one with the UNISYS surface data, the synoptic set up in the right hand side of Figure 2 was not well represented by the model. There was very little condensate loading, vertical velocity plumes were non-linear and actually curved back westward with height. The density pertubations on the leeward side of the topography actually suggested subsidence, thereby inhibiting the development of the lee cyclone.

13 To gain a better understanding of the small scale physical processes that took place in the last two cases, it would be beneficial to float the grid mesh to the intersection point. The use of higher temporal and spatial resolution surface plots to identify the time and location the cyclone passage would help to determine where to center the model grid. Model runs with higher temporal resolution plots would also provide insight into the tendencies of these turbulent forcings and their progression through time with respect to the cyclone s entire path over the topographical feature as opposed to hourly snapshots. The use of additional cases would also provide a larger data set with which to understand the primary forcings since the use of only a few events leads to a greater probability of anomalous features.

14 References 1. Aebischer, U., C. Schar, 1998: Low-Level Potential Vorticity and Cyclogenesis to the Lee of the Alps. Journal of the Atmospheric Sciences. 55, Gulev, S.K., O. Zolina, S.Grigoriev, 2001: Extratropical cyclone variability in the Northern Hemisphere winter from NCEP/NCAR reanalysis data. Climate Dynamics. 17, Holton, J.R., G.J. Hakim, 2013: An Introduction to Dynamic Meteorology., 5th, Waltham: Academic Press. Print. 4. Markowski, P., Y. Richardson, 2010, Mesoscale Meteorology in Midlatitudes, 1st, Hoboken: Wiley-Blackwell Print 5. Mattocks, C., R. Bleck: Jet Streak Dynamics and Geostrophic Adjustments Processes during the Initial Stages of Lee Cyclogenesis. Monthly Weather Review. 114, NOAA Satellite and Information Service, (N.D.). Retrieved from 7. Ulbrich, U., J.G. Pinto, H. Kupfer, G.C. Leckebusch, T. Spangehl, M. Reyers, 2008: Changing Northern Hemisphere storm tracks in an ensemble of IPCC climate change simulations. Journal of Climate. 21, Unisys Weather, Retrieved from 9. Wang, X.L., V.L. Swain, F.W. Zwiers, 2006: Climatology and changes of extratropical cyclone activity: comparison of ERA-40 with NCEP-NCAR reanalysis for Journal of Climate. 19,

The Planetary Circulation System

12 The Planetary Circulation System Learning Goals After studying this chapter, students should be able to: 1. describe and account for the global patterns of pressure, wind patterns and ocean currents