Strong Cross-Barrier Flow under Stable Conditions ... - AMS Journals

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Strong Cross-Barrier Flow under Stable Conditions Producing Intense Winter Orographic Precipitation: A Case Study over the Subtropical Central Andes MAXIMILIANO VIALE AND FEDERICO A. NORTE Programa Regional de Meteorologıá, Instituto Argentino de Nivologıá, Glaciologıá y Ciencias Ambientales (IANIGLA), CCT–CONICET, Mendoza, Argentina (Manuscript received 2 June 2008, in final form 6 February 2009) ABSTRACT The most intense orographic precipitation event over the subtropical central Andes (368–308S) during winter 2005 was examined using observational data and a regional model simulation. The Eta-Programa Regional de Meteorologıá (PRM) model forecast was evaluated and used to explore the airflow structure that generated this heavy precipitation event, with a focus on orographic influences. Even though the model did not realistically reproduce any near-surface variables, nor the precipitation shadow in the leeside lowlands, its reliable forecast of heavy precipitation over the windward side and the wind fields suggests that it can be used as a valuable forecasting tool for such events in the region. The synoptic flow of the 26–29 August 2005 storm responded to a well-defined dipole from low to upper levels with anomalous low (high) geopotential heights at midlatitudes (subtropical) latitudes located off the southeast Pacific coast, resulting in a large meridional geopotential height gradient that drove a strong anomalous cross-barrier flow. Precipitation enhancement in the Andes was observed during the entire event; however, the highest rates were in the prefrontal sector under the low-level stable stratification and crossbarrier winds exceeding 2.5 standard deviations (s) from the climatological monthly mean. The combination of strong cross-mountain winds with the stable stratification in the air mass of a frontal system, impinging on the high Andes range, appears to be the major factor in determining the flow structure that produced the pattern of precipitation enhancement, with uplift maximized near mountaintops and low-level blocking upwindleading to the formation of a low-level along-barrier jet. Additionally, only the upstream wind anomalies for the 15 heaviest events over a 10-yr (1967–76) period were investigated. They exhibited strong anomalous northwesterly winds for 14 of the 15 events, whereas for the remaining event there were no available observations to evaluate. Thus, these anomalies may also be exploited for forecasting capabilities.

1. Introduction Winter orographic precipitation events over the subtropical central Andes (368–308S) of South America have a large impact on the regional climate system. The principal water source in the surrounding lowlands is the snow that accumulates in the Andes during winter storm events (Araneo and Compagnucci 2008; Viale et al. 2008). When these events have a high intensity, however, they create large negative impacts, including flooding and road blockages. The winter events result from a complex interaction between the general circu-

Corresponding author address: Maximiliano Viale, Instituto Argentino de Nivologıá, Glaciologıá y Ciencias Ambientales, Av. Adriań Ruiz Leal s/n, Parque Gral. San Martıń, CC 330, 5500 Mendoza, Argentina. E-mail: [email protected] DOI: 10.1175/2009WAF2222168.1 Ó 2009 American Meteorological Society

lation of the atmosphere and steep orography. The midlatitude westerly and storm tracks have seasonal variability, traveling more northward and being more intense in winter than in summer (Trenberth 1991; Hoskins and Hodges 2005). Consequently, this mountainous region and its windward lowlands in central Chile have wintertime precipitation events almost exclusively (Erenõ and Hoffman 1978; Montecinos et al. 2000; Falvey and Garreaud 2007). Recently, Falvey and Garreaud (2007) showed that daily wintertime precipitation in the central Chile (328–358S) is strongly associated with the cross-mountain moisture flux, which is mostly determined by synoptic conditions and favors orographic lifting by the Andes. Pandey et al. (1999) remarked that precipitation in mountainous regions of the Sierra Nevada in California is maximized by an optimal combination of moisture transport and orographic lifting, inferred from a preferred

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FIG. 1. Map of the region under study, including orography (m, elevations higher than 500 m are shaded), and the stations listed in Table 1.

perpendicular wind direction to the mountain range. In addition, large anomalies in cross-barrier winds are useful tools for anticipating heavy precipitation in the northern Sierra Nevada (Junker et al. 2008). Orographic lifting and available moisture are responsible for a great deal of precipitation in mountainous terrain. Orographic lifting is a response to the interaction of synoptic airflow with mountains and depends on multiple factors, including static stability, the strength of cross-barrier winds, moisture content, and the topography. The Andes are roughly an ideal two-dimensional obstacle because of their perpendicular orientation to the westerlies, their narrow width (;150 km), and their high mean altitude (;5000 m), which, together with their proximity to the South Pacific Ocean (SPO; ;200 km from the coast to the Andes crest), make them a great barrier for moisture advection and baroclinic precipitation systems coming from the west (see Fig. 1). Hence, there is a significant climatic difference in the winter precipitation between the windward (about 350 mm) and leeward

sides (about 70 mm). However, there has been little treatment of the possible combinations of factors that favor orographic precipitation in the Andes. Smith and Evans (2007) examined orographic precipitation in the southern Andes (408–488S) where the topography is significantly lower than in the central Andes. Medina and Houze (2003) studied the differences in airflow patterns and orographic modifications of two baroclinic precipitation cases over the European Alps with different static stability conditions. Several studies have documented that under stable conditions, low-level upstream flow is profoundly blocked and separated from upper-level strong cross-barrier flow by strong vertical cross-barrier shear layer and along-barrier winds. This type of airflow pattern in baroclinic storms passing over different mountain ranges is associated with ensuing enhanced precipitation on windward slopes [e.g., over the Sierra Nevada in California (Marwitz 1987), the Cascades in Washington and Oregon (Medina et al. 2005; Garvert et al. 2007), and the European Alps (Bousquet and Smull 2003; Medina and Houze 2003)].

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In the Andean region of South America, the difficulty in characterizing local features in orographic precipitation mostly resides in the scarcity of meteorological stations. Falvey and Garreaud (2007) could not establish a clear rainfall spatial distribution in the cross-mountain direction of central Chile because of the poor rain gauge network in use on upslope zones. Colle and Mass (2000), for example, noted from a dense station network in the relatively high Cascades in Washington that the maximum precipitation was some kilometers upstream of the crest with a significant precipitation shadow in the lee. However, in relatively narrow and low barriers, this maximum is observed closer to the crest [e.g., over the Oregon coastal mountains (Colle and Mass 2000) and the New Zealand Southern Alps (Sinclair et al. 1997)]. The limitations due to the scarcity of stations in the Andes can be partially solved using numerical simulations with regional models. In addition to their precipitation forecasting capability, these simulations can be used to explore the spatial structures in precipitation that are associated with orographic influences. The regional Eta Model (Mesinger et al. 1988) has performed well in forecasting precipitation across high mountains [e.g., the Cascades in the northwestern United States; Black (1994); Colle et al. (1999)], as well as downslope leeward winds in central western Argentina during orographic precipitation events (Seluchi et al. 2003), locally named the zonda wind (Norte 1988). During the 2005 winter season, the most significant event took place on 26–29 August. The surrounding high areas of Santiago, Chile, experienced an intense rainfall of 150 mm in 2 days. It caused flooding and damage in the city and forced approximately 1500 residents to leave their homes. In the high mountains, the intense snowfall and winds caused the blockage of the principal route between Chile and Argentina for 6 days, interrupting international commerce and leaving many truck drivers and ski resort tourists stranded in bad weather conditions. The few precipitation measuring stations in the region revealed that the event’s precipitation totals ranged from 50–150 mm over the lowlands of central Chile to 60–300 mm over the Andes. In addition, the downslope zonda wind was observed in some Argentinean surface stations along the lee side. The goal of this study is to provide a complete description of the intense orographic precipitation episode of 26–29 August 2005 over the subtropical central Andes. This includes the synoptic and thermodynamic evolution patterns that accompanied the event, as well as a detailed examination of the precipitation distribution across the mountain range. We evaluate the effectiveness of the Eta-Programa Regional de Meteo-

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rologıá (PRM) regional model as a numerical forecasting tool. We examine the anomalous cross-barrier winds of the 26–29 August 2005 storm and the 15 heaviest storms over a 10-yr period (1967–76), as potential indicators of intense rainfall. In addition, we use the simulation to explore the features of the airflow during the heaviest event of 2005 that led to significant precipitation with a focus on identifying orographic influences.

2. Data and methods a. Data and geographic setting The study area is plotted in Fig. 1, and the station data used here are listed in Table 1. Of the 25 surface stations, 8 correspond to those of the Direccioń Meteorolo´gica de Chile (DMC), 9 to the Servicio Meteorolo´gico Nacional de Argentina (SMN), 1 to Centro Nacional del Medio Ambiente (CENMA) of Chile, and 7 are snow telemetry (SNOTEL) stations. The SNOTEL stations were recently installed in remote high mountainous regions by the Departamento General de Irrigacioń (DGI) of Mendoza, Argentina (MZA). They have automated snowpillow sensors that measure the daily snow water equivalent precipitation. The synoptic evolution was analyzed with data from the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis (Kalnay et al. 1996). These grid data are available every 6 h, with a horizontal resolution of 2.58 3 2.58 latitude–longitude and 17 vertical levels. We examined the atmospheric vertical profile, together with the Eta-PRM model evaluation, using rawinsonde observations at Santo Domingo, Chile, station. The rawinsonde observations are made twice a day (0000 and 1200 UTC) along the central Chilean coast. The rawinsonde launching was moved to less than 80 km from Quintero, Chile (QUI), to Santo Domingo (SDM) station in 1998 (see Fig. 1), and then, the 1961–90 climatology of the variables correspond to Quintero station. According to Falvey and Garreaud (2007), the shift in location had a minimal effect on the values of most variables for the winter baroclinic precipitation systems that have synoptic temporal and spatial scales, so following their conclusions, we adopt the term ‘‘coastal rawinsonde’’ without distinction between locations. For the anomalies explorations, the 1961–90 August monthly mean and its standard deviations (s) of geopotential height and u and y winds were calculated using the daily coastal rawinsonde and the 4-hourly NCEP– NCAR reanalysis data. Then, the normalized anomalies

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TABLE 1. Coordinates and heights of the stations used in the study (see Fig. 1). The horizontal line separates groups between the windward- (upward), and leeside (downward) stations. The leeside stations in italic correspond to SNOTEL stations. ID

Lat (S)

Lon (W)

Elev (m)

La Serena Valparaiso Pudahuel Aero Quinteros Santo Domingo Quinta Normal Curico Chillan Lagunitas

Station name

LSE VAL PUD QUI SDM QTN CUR CHI LAG

298549S 338019S 338239S 328509S 338399S 338309S 348589S 368359S 338089S

718129 718309 708479 718209 718379 708409 718139 728029 708259

50 120 475 8 0 520 228 124 2765

Jachal San Juan Aero Punta de Vacas Mendoza Aero San Martin Chacras de Coria San Carlos San Rafael Malargue Aero Horcones Toscas Palomares Salinillas Lag. Diamante Lag. Atuel Valle Hermoso

JAC SJA PVA MZA SMA CHA SCA SRA MAL HOR TOS PAL SAL LGD LGA VHE

308149S 318349S 328519S 32859S 338059S 328599S 338469S 338359S 358039S 328479S 338099S 338379S 338519S 348079S 348289S 358089S

688459 688259 698459 688479 688259 688529 698029 698359 698359 698569 698539 698539 698479 698429 708019 708129

1175 598 2225 704 653 921 940 748 1425 3038 3000 2900 2616 3300 3600 2250

are defined as those variables at the specific time of the storm minus the August monthly mean, divided by s.

b. Model setup and description The Eta Model is a hydrostatic model, which uses the eta (h) vertical coordinate defined by Mesinger (1984) in order to reduce to a large extent the errors generated when computing the pressure gradient force, as well as the advection and horizontal diffusion, along a steeply sloped coordinate surface. The eta coordinate is defined as h5

p ps

pt pt

pref (zs ) pref (0)

pt , pt

(1)

where pt is the pressure at the top of the domain; ps and zs are the pressure and height of the model’s surface, respectively; and pref is the pressure reference state that is a function of the distance above sea level. The first term in Eq. (1) corresponds to the standard sigma coordinate, while the second is a factor that depends only on the horizontal coordinates and converts the sigma into the eta coordinate. For a more complete description of the model’s structure and dynamical core, see Mesinger et al. (1988), Janjic (1990, 1994), and Black (1994).

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The Eta-PRM regional model used here is a version of the Eta-Centro de Previsaõ de Tempo e Estudos Clima´ticos (CPTEC-Brazil) regional model. The EtaCPTEC model has been used at CPTEC for weather forecasting over South America (SA) since late 1996 (Seluchi and Chou 2001). Its domain covers most of SA (108N–458S, 838–308W) and has a horizontal resolution of 40 km with 38 vertical levels, from the surface to the stratosphere. The initial and boundary conditions for Eta-CPTEC are NCEP operational analyses every 6 h with a horizontal resolution of nearly 93 km (T126 triangular truncation) and 28 vertical levels. The topography is represented as discrete steps whose tops coincide with the model-layer interfaces. It has a complete physics package, with schemes for grid-scale precipitation represented in explicit form (Zhao and Carr 1997), and the Betts–Miller scheme for convective precipitation (Betts and Miller 1986), as modified by Janjic (1994). The turbulence exchanges in the free atmosphere are based on the Mellor and Yamada (1974) level 2.5 scheme. The Eta-PRM had a horizontal resolution of 15 km, covering central Chile and central Argentina (;298–388S, 758–578W), and uses as initial and boundary conditions the 40-km Eta-CPTEC output. The model was recently implemented at PRM in Mendoza. The model integrations began at 1200 UTC 25 August 2005, approximately 36–48 h before maximum precipitation intensity, and ended at 1200 UTC 29 August 2005 (96-h forecast). To verify the Eta-PRM model output at the observation sites, the grid-point values surrounding the observation were bilinearly (linearly) interpolated in the horizontal (vertical) to each of the observation locations.

3. Brief synoptic overview of the 26–29 August 2005 storm Figures 2 and 3 show the synoptic situation during the precipitation event. An intense baroclinic zone located in the southeastern Pacific Ocean at 1200 UTC 25 August (not shown) approached the Chilean coast at 1200 UTC 26 August. This baroclinic zone accompanied a surface pressure depression also located near the coast of central Chile at 1200 UTC 26 August (Fig. 2a). This low strengthened and moved inland at 0000 UTC 27 August (Fig. 2b), when the maximum precipitation rates over the mountains began. At 500 hPa, the strong flow was orientated perpendicularly to the mountain (Figs. 3a and 3b). By the end of the event, another surface low pressure system approached and deepened around 458S over the Chilean coast (Figs. 2c and 2d), while the rear section of the strong westerly belt moved northward at midlevels (Figs. 3c and 3d). The end of the event was associated

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FIG. 2. The surface NCEP–NCAR reanalysis data at (a) 1200 UTC 26 Aug, (b) 0000 UTC 27 Aug, (c) 1200 UTC 28 Aug, and (d) 0000 UTC 29 Aug 2005 showing SLP (solid black line every 5 hPa) and 1000–500-hPa thickness (dotted–dashed gray line every 6 dam).

with a cold outbreak, appreciable by strong southerly flow between 858 and 908W at 0000 UTC 29 August (Figs. 2d and 3d). As we will show in section 5, the South Pacific anticyclone at subtropical latitudes along with the extended area of low pressure at midlatitudes were stronger than normal during the event (compared to the monthly long-term 1961–90 average obtained from the NCEP– NCAR reanalysis). These anomalous systems in most of the troposphere over the southeast Pacific produced a large meridional pressure gradient, resulting in a strong westerly flow. The satellite image in Fig. 4 shows the cloudiness during prefrontal conditions at 1745 UTC 26 August. The extratropical cyclone position around 408S along the Chilean coast can be identified by an inverted comma cloud (Southern Hemisphere) in the warm front part of

the surface cyclone (Browning 1990). Remarkably, this inverted comma cloud was abruptly disrupted in central western Argentina (north of 378S) by the high Andes forming a foehn wall cloud, with clear sky immediately downstream of the highest peaks (Fig. 4). This satellite image clearly shows the major influence of the Andean cordillera on cloud patterns, which have stratiform clouds upstream and clear sky downstream of the crest, suggesting upward and downward motions, respectively.

4. Synoptic and mesoscale observational features and model evaluation of the 26–29 August 2005 storm a. Surface time series The surface temporal evolution of the meteorological conditions is analyzed at stations located on the lowlands

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FIG. 3. The 500-mb NCEP–NCAR reanalysis data at (a) 1200 UTC 26 Aug, (b) 0000 UTC 27 Aug, (c) 1200 UTC 28 Aug, and (d) 0000 UTC 29 Aug 2005 showing geopotential heights (solid line every 5 dam) and winds (one pennant, full barb, and half-barb indicate 50, 10, and 5 m s21, respectively).

windward and leeward of the Andean cordillera, where 3-hourly data are available. Figures 5 and 6 show the evolution of the observed and model-forecasted 10-m winds, sea level pressure (SLP), 2-m temperature, 2-m dewpoint temperature, and precipitation at two windward-side and three leeside stations.

1) WINDWARD

SIDE

The windward-side stations registered two distinct rainfall periods, where the first was more intense than the second (Figs. 5a and 5b). These rainfall periods were associated with the passage of two cold fronts, the principal cold front (PCF) and the second cold front (SCF), respectively. The PCF passage at Pudahuel station (PUD, site of Santiago’s international airport) was recognized at 0900 UTC 27 August by an SLP minimum, a wind shift to the southwest, and a principal

maximum rainfall (Fig. 5b). The SCF arrived at 1800 UTC 28 August and was also denoted by an SLP minimum, a drastic decrease in surface temperature, a secondary maximum rainfall, and winds veering from the north to the southwest (Fig. 5b). The polar airmass entrance at PUD was marked by an almost continuous temperature decrease from the PCF passage to 1.38C at 1200 UTC 29 August (Fig. 5b). The SDM station had no observations at dawn (i.e., after 0000 UTC and before 1200 UTC); however, the two rainfall times and some traces of veering winds with frontal passages were recorded (Fig. 5a). A slight difference in the observations from PUD was that the continuously decreasing temperature was not observed at SDM because it is located on the coast and was influenced more strongly the marine layer than was PUD, which is located more inland in a closed valley.

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FIG. 4. Satellite image in the visible channel at 1745 UTC 26 Aug 2005 (approximately 6 h before maximum rainfall rates in central Chile) from SMN.

Many of the observed features during the event were well forecast by the model, such as two distinct rainy periods and minimums of SLP, as well as a generally decreasing 2-m temperature associated with the cold fronts (Figs. 5c and 5d). However, the model underestimated the 2-m dewpoint temperature by around 128C (28C) before (after) the PCF passage at PUD, and by 48C before PCF at a coastal station (SDM). Also, it overestimated the daily diurnal 2-m temperature range, and overestimated the 10-m wind speed by 5 m s21 at SDM (Fig. 5c). At PUD (Fig. 5d), the forecast wind directions did not agree with the observations, likely due to fact that the horizontal resolution is not high enough to reproduce the discontinuous wind directions across the closed valley, where this station is located (Fig. 1). The precipitation amounts at PUD and SDM were realistically simulated by the Eta-PRM, but the initial times were forecast nearly 6 h later than the actual onset (Figs. 5c and 5d). The model captured 88% and 110% of the observed precipitation totals at PUD and SDM, respectively. Most of the underprediction (overprediction) occurred during the principal (secondary) rainfall rates at both Chilean (only at SDM) station locations (Figs. 5c and 5d).

2) LEEWARD SIDE The most important signals of the PCF passage in the northernmost leeside region are the onset and duration of the downslope (zonda) winds. At San Juan, Argentina (SJA), a weak downslope wind event was identified by the wind direction (west component) and a quick increase of the dewpoint depression between 0000 and 1200 UTC 27 August (Fig. 6a). This abrupt increase in air temperature during the night is also a clear signal of the onset of a zonda event. Even though it is frequently difficult to determine the onset of a zonda wind episode in the city of San Juan, there is usually a transition period (about 3–4 h on average) where there are gradational interruptions of westerly air in the prevailing north-northeast flow around the Atlantic anticyclone. Atlantic-sourced air circulation progressively loses influence until the station is definitively dominated by the zonda [i.e., northwest or west winds; Norte (1988)]. The zonda was also observed at about the same time at San Jose´ de Jaćhal, Argentina (JAC, not shown), whereas it was not observed at MZA (Fig. 6b). MZA is located in a relatively low-lying area,

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FIG. 5. (a) SDM and (b) PUD observations plotted every 3 h from 1200 UTC 25 Aug to 1200 UTC 29 Aug 2005 showing temperature (8C, black solid line), dewpoint temperature (8C, dotted–dashed black line), SLP (mb, black dots), winds (full barb 5 10 m s21), and 6-h precipitation (shaded, mm). The SMD and PUD locations are shown in Fig. 1. Gaps in temperature, dewpoint temperature, SLP, and winds at SMD are present in the night time. (c),(d) Same as in (a),(b), respectively, but for the Eta-PRM simulation.

and sometimes the zondas do not extend fully to the surface because they cannot break the thermal inversion, due to the cold static air over the urban area (Norte 1988). The principal and secondary rainfall times at Malarguë, Argentina (MAL), agreed with the PCF and SCF passages in the central Chile region, while the zonda wind first occurred at 0000 UTC 26 August and then between 1500 UTC 27 August and 0900 UTC 28 August (Fig. 6c). The mean mountain heights close to MAL are lower than those surrounding SJA and MZA. Thus, diminished orographic blocking effects would result in more frequent downslope winds and increased spillover of precipitation onto the lee side of the barrier close to MAL, as suggested in the meteorograms in Fig. 6. The last zonda

wind period at MAL began nearly 9 h earlier than at SJA and JAC (cf. Figs. 6a and 6c), showing the northward advance of the downslope wind occurring with the northward cold front displacement (Norte 1988). The Eta-PRM underpredicted the rainfall at MAL and MZA (Figs. 6e and 6f), which is in agreement with findings that the cold season 10-km Eta Model does not generate enough precipitation on the lee of major barriers in the northwestern United States (Colle et al. 1999). In addition, the model outputs for the leeward side had errors similar to those observed on the windward side. For example, the model erroneously predicted 10-m wind directions, underpredicted the 2-m dewpoint temperature, and overpredicted the magnitude

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FIG. 6. Same as in Fig. 5 but for leeside stations in Argentina: (a) San Juan (SJA), (b) Mendoza (MZA), and (c) Malarguë (MAL). (d)–(f) As in (a)–(c) but for the Eta-PRM simulation. SJA, MZA, and MAL are shown in Fig. 1. The SLP at MAL is replaced by the 850-hPa geopotential height (GH) because its altitude is 1425 m (higher than 800 m).

of the diurnal 2-m temperature ranges, as well as delaying by nearly 6 h the PCF passage (Figs. 6d and 6f). The zonda wind at SJA was discernible in the model output by the veering and strengthening of the winds from the north to the west, and a slight drop in 2-m dewpoint temperature at 0000 UTC 27 August (Fig. 6d). However, the model did not reproduce the observed weak north-northwesterly winds during this time at SJA (Fig. 6d), nor did it reproduce the observed southerly winds from 1200 UTC 27 August onward at SJA and MZA (Figs. 6d and 6e). At MAL (Fig. 6f), the simulated winds were mostly from the west, and during the occurrence of observed downslope winds, their intensity was underestimated. Gallus (2000) also detected an underestimation of the near-surface winds using the eta vertical coordinate of the 10-km Eta Model to simulate

a downslope windstorm along the lee of the central Rocky Mountains. Gallus pointed out that erroneous wind forecasts are due to flow blocking downstream from high topography, which appeared to worsen as the horizontal resolution increased (,32 km). These wind forecasts could be improved by setting the Eta-PRM model with the recently developed sloping-steps eta discretization, which showed better downslope winds for a zonda case than did the standard step-mountain approach (Mesinger et al. 2006). Nevertheless, the model was able to forecast several aspects of the zonda wind effects, such as wind strength, temperature, and dewpoint changes, but a delay in their ending time was detected (Figs. 6d and 6f). Seluchi et al. (2003) found similar satisfactory results for the 40-km Eta-CPTEC simulated zonda wind cases.

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b. Vertical upstream profiles During the wintertime, the typical stable vertical atmospheric profile in central Chile, associated with the subtropical South Pacific anticyclone circulation (Satyamurty et al. 1999), is frequently replaced by a less stable profile with the advance of a frontal system from the SPO. The vertical profile evolution was observed during the event using the coastal rawinsonde (Fig. 7). At 1200 UTC 25 August, the profile exhibited a subsidence inversion around the 850-hPa level with dry air aloft, which then was replaced by a quasi-saturated 600–200-hPa layer at 0000 UTC 26 August (not shown). About the time of the PCF passage (Figs. 7a and 7b), the environment was saturated below 650 hPa and subsaturated above it, while a strong cross-barrier flow in the whole troposphere was observed. An isothermal layer below 900 hPa was present at 0000 UTC 27 August (Fig. 7a), and above it to the 700-hPa level the sounding was slightly stable. The sounding remained relatively stable below the 700-hPa level at 1200 UTC 27 August (Fig. 7b). At two times in Figs. 7a and 7b, the sounding exhibited a layer between 700 and 600 hPa that was less stable than the rest of the air column. About the time of SCF passage (Figs. 7c and 7d), a slight reduction in humidity and a rotation in wind directions from westerly to northwesterly was noted at low levels. The soundings were also stable at two times and the westerly upper winds reached 60–65 m s21 above 400 hPa (Figs. 7c and 7d). Above 700 hPa, the environment was very dry at 0000 UTC 29 August (Fig. 7d). The main features in the vertical profile evolution were mostly reproduced by the Eta-PRM model. However, the model had some difficulty in capturing the saturated conditions and simulated slightly less stable conditions than were observed below 700 hPa during the PCF (Figs. 7a and 7b). The model also reproduced drier conditions before the SCF passage (Fig. 7c), while after the SCF the only difference was a slightly deeper moist layer in the model (saturated to nearly 600 hPa compared with observations to only 700 hPa Fig. 7d). The aforementioned underprediction (overprediction) of rainfall in the model during PCF (SCF) at PUD and SDM is likely associated with a model forecast of shallower (deeper) saturation in a layer around 850 hPa (650 hPa) than was found in the observations (Figs. 7a, 7b, and 7d). Nevertheless, the wind profiles were well solved by the Eta-PRM model. The only noticeable difference in Fig. 7a was a shallow near-surface layer with easterly–northerly winds in the model, while the observations showed northerly to northwesterly winds. The dry Brunt–Vaïsa¨la¨ frequency (Nd2) calculated from the sounding indicates that profiles were stable through-

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out the event (Fig. 8). Since the atmosphere was saturated below 650 hPa at about the time of the PCF (Figs. 7a and 7b), the moist Brunt–Vaïsa¨la¨ frequency was also calculated (Nm2 ; Durran and Klemp 1982), which accounts for the latent heat of the condensation in a saturated environment. Below 700 hPa, Nm2 remained weakly positive (Figs. 8a and 8b), indicating neutral to stable flow. Above 700–600 hPa, Nm2 was slightly negative at 0000 UTC 27 August (Fig. 8a), while this negative value was reduced to a shallower layer around 700 hPa at 1200 UTC 27 August (Fig. 8b). These layers with neutral to unstable flow are consistent with the less stable 700–600-hPa layer mentioned previously in Figs. 7a and 7b. At the same time, for those layers in Figs. 7c and 7d (i.e., when the atmosphere was not saturated), the Nd2 remained positive, indicating stable stratifications (Figs. 8c and 8d). The Eta-PRM simulation reproduced slightly lower Nm2 than were observed at 0000 and 1200 UTC 27 August, but it closely reproduced the observed Nd2 at the four times shown in Fig. 8, including its sharp variations with height. The cross-barrier (u wind) and along-barrier (y wind) profiles for different times of the storm obtained from the coastal soundings are shown in more detail in Fig. 9. The early stage of the storm featured a significant increase with height in the cross-barrier profile at the lowest levels, producing a strong vertical shear below 900 hPa (Fig. 9a). This strong cross-barrier shear was coincident with a near-surface along-barrier jet, showing peak speeds of 15 m s21, and with a very stable nearsurface layer (i.e., an isothermal layer in Fig. 7a and the largest values of Nm2 in Fig. 8a). After the early stage of the storm, the wind profiles remained nearly constant, as characterized in an intense cross-barrier shear and an along-barrier jet around the 700-hPa level, and above it a strong cross-barrier flow (Figs. 9b and 9d). The results of Medina et al. (2005) could explain the development of the strong u-wind shear (cf. the observed values for Nm2 and Nd2 in our Fig. 8 with their Fig. 5). Their idealized 2D simulations of Mesoscale Alpine Programme (MAP) and Improvement of Microphysical Parameterization through Observational Verification Experiment (IMPROVE) cases indicated that, in the presence of sufficient upstream static stability, orographic effects support the development of a strong shear layer on the windward side of the mountain range. The observed wind profiles upstream of the Andes also bear similarities to those flow structures of cases of profoundly blocked flow associated with orographic storms over the Sierra Nevada (Marwitz 1987) and the Alps (Bousquet and Smull 2003). The model did not reproduce the near-surface shear and the along-barrier jet during the beginning of the storm (Fig. 9a). However, it successfully simulated the

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FIG. 7. The skew T–log p diagram from the coastal rawinsonde station at (a) 0000 UTC 27 Aug, (b) 1200 UTC 27 Aug, (c) 1200 UTC 28 Aug, and (d) 0000 UTC 29 Aug 2005 showing observed profiles (black solid line) and the respective 12-, 36-, 50-, and 84-h Eta-PRM forecasts (gray dashed line). The profiles correspond to temperature (8C, right line), dewpoint temperature (8C, left line), and wind (half-barb, 5 m s21; full barb, 10 m s21; pennant, 50 m s21).

near-constant observed wind profiles during the remaining stages of the storm (Figs. 9b and 9d). The strong cross-barrier shear above 700 hPa and the along-barrier jet around 700 hPa were well simulated, although in a smoother way than in the observations.

c. Observed and simulated orographic precipitation amounts Figure 10 shows the observed and forecasted precipitation totals, and the percentage of observed precipitation

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FIG. 8. Vertical profiles of squared dry (gray) and moist [black, according to definition of Durran and Klemp (1982)] Brunt–Vaïsa¨la¨ frequency (1024 s22) calculated using the coastal rawinsonde observations (solid line) and the Eta-PRM model output (dashed line) at (a) 0000 UTC 27 Aug, (b) 1200 UTC 27 Aug, (c) 1200 UTC 28 Aug, and (d) 0000 UTC 29 Aug 2005.

(POP) at the station locations (numbers alongside circle). The POP is defined as the surrounding grid points of the model output—bilinearly interpolated to the observation location—divided by the observed precipitation (F/O), and then multiplied by 100. Most observations recorded an increase in precipitation with height above the lowlands in central Chile (between 50 and 100 mm), reaching a maximum of 286 mm upslope at Lagunitas, Chile (2765 m), from where it decreased for SNOTEL stations immediately downstream of the highest peaks (i.e., precipitation between 100 and 200 mm immediately east of the Argentina–Chile border in Fig. 10). A notable decrease in precipitation in the lowlands of the lee side was also recorded, with rainfall between 0 and 10 mm in northernmost sector and 10–50 mm in the southernmost sector. According to Falvey and Garreaud (2007), the enhanced precipitation of the upwind effect on the Andes is tested by examining the ratio between the precipitation amounts in upslope and immediate leeside

highland areas (SNOTEL stations) with those of lowland areas in central Chile. The observations from the 26–29 August 2005 storm depicted an enhancement ratio around 338S of close to 3 and 2 for areas upslope and immediately in the lee, respectively. These values were close to 3.5 for the cluster of events with major precipitation enhancement rates on upslope areas reported by Falvey and Garreaud (2007) (cf. their Fig. 11g). Our results are also consistent with the precipitation regime across the central Andes documented by Erenõ and Hoffman (1978) based on a slightly better rain gauge network around 338S, especially at high altitudes. The Eta-PRM precipitation forecast had acceptable POP scores (70%–130%) over the lowlands of central Chile and north of 348S along the immediate lee side. West of the Andes, there was over- (under-) estimation in the southern (northern) sector. The model prediction alternated between over- and underestimation at SNOTEL stations near the crest line on the lee side,

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FIG. 9. Vertical profiles of u wind (cross-barrier) and y wind (along-barrier) from coastal sounding (black line) and the Eta-PRM model output (gray line) at (a) 0000 UTC 27 Aug, (b) 1200 UTC 27 Aug, (c) 0000 UTC 28 Aug, and (d) 0000 UTC 29 Aug 2005.

suggesting that the horizontal resolution is not high enough to reproduce in detail the local circulation patterns across valleys and complex terrain. East of the Andes, the model predicted less than 70% of the observed storm-total precipitation at most of the lowlying stations. The bias of zero, in some isolated locations of the northern part, corresponds to stations that reported no or little precipitation, while the model forecasted no precipitation. The overestimation (underestimation) in the southern lowlands of the windward (leeward) side agrees with the results of Colle et al. (1999), who attributed it to possible problems with the microphysical scheme and/or excessive flow blocking upwind of the barrier with the 10-km Eta Model. Even though the scarcity of meteorological stations in the region does not allow us to accurately establish the precipitation distribution, the relatively good performance of the model in precipitation forecasts makes it a useful diagnostic tool. In addition, sensitivity studies have documented that although a horizontal resolution of 4 km or less in a regional model results in more realistic precipitation structures over the Cascades range, the main orographic contributions to the precipitation distribution are also well simulated at 12-km resolution [i.e., with increased (decreased) precipitation on steep windward (leeward) slopes; Colle et al. (2000); Colle and Mass (2000)]. As expected the Eta-PRM simulated a relatively coarse pattern of orographic precipitation

in accord its smooth step-topography terrain with a resolution of 15 km. However, the main features can be estimated in this Andean region for a heavy winter storm (Fig. 10). The Eta-PRM model presented a marked northward decay pattern in central Chile and the Andean region (Fig. 10), consistent with long-term wintertime patterns (Hoffman 1975; Falvey and Garreaud 2007). Moreover, the precipitation maximum was concentrated to the south of Santiago (;33.58S), which is in agreement with the precipitable water pattern over the west coast of central Chile for the topographically enhanced events of Falvey and Garreaud (2007) (cf. their Fig. 12c). The simulation also reproduced a very large zonal precipitation gradient across the mountains with a rain shadow east of the crest. The spillover rainfall extended a few kilometers eastward onto the lee side, with the precipitation amount almost reduced to zero about 150 km east of the highest terrain (Fig. 10). The large orographic effect of the Andes in precipitation at this latitude can also be noted by the simulation of the zone of maximum precipitation (250–275 mm) upslope of the crest (i.e., before the Chile–Argentina border). The Eta-PRM simulated the upslope precipitation maximum between 2500 and 3000 m in central part of the study area (33.58–35.58S), representing a precipitation enhancement of the order of 3.5, a value very close to the available observations. It is clear from Fig. 10 that the subtropical central Andes are high enough to produce the

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FIG. 10. Total 96-h precipitation forecasted by the Eta-PRM model from 1200 UTC 25 Aug through 1200 UTC 29 Aug 2005 (shaded, mm). The numbers at the filled circle locations indicate the percent of observed storm-total precipitation, the observed storm-total precipitation (fill-sized circle), and the model topography above 10 hm (contour every 5 hm).

precipitation maximum on the upwind slope, as is generally the case for a larger barrier (Rauber 1992; Colle and Mass 2000; Roe 2005).

5. Wind anomalies observed during the 26–29 August 2005 storm and during other heavy precipitation events Significant links between winter rainfall in central Chile and the zonal flow were found by Falvey and Garreaud

(2007), while anomalous winds, especially in the crossmountain direction, are expected to be associated with significant orographic lifting upstream and heavy precipitation rates on windward slopes and over high mountains, as was noted over the similar Sierra Nevada mountain range (Junker et al. 2008; Pandey et al. 1999). Based on these studies, we also examine the existence of significant anomalies that could be used by a forecaster as an extra tool in predicting intense orographic precipitation events, as was found useful by Junker et al. (2008).

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FIG. 11. Temporal evolution of the zonal (cross mountain) wind at the 850-, 700-, 500-, and 250-hPa levels for the coastal rawinsonde observations (dotted) and the Eta-PRM model output (solid line). The black solid straight line represents the 1961–90 August monthly mean and the gray dashed line represents the 12.5 std dev.

The anomalous cross-mountain flow at the coastal rawinsonde location for 26–29 August 2005 storm is evident in Fig. 11. At 850 and 700 hPa (Figs. 11a and 11b), the u wind has two higher values corresponding to PCF and SCF passage; in the first, the climatological departures surpassed 12.5s. At mid- and upper levels (Figs. 11c and 11d), the zonal flow exhibited a continuous increase until 1200 UTC 28 August (SCF) and later a marked decrease. The Eta-PRM closely forecasted the cross-barrier wind evolution at lower levels, with slightly lower values at the time of maximum speeds (Figs. 11a and 11b). At upper levels (Figs. 11c and 11d), the maximum u winds were forecast about 6 h before, but the model output was close to the observations, with a 12.5s departure at 500 hPa. The spatial distribution of the normalized wind and height anomalies using the NCEP–NCAR reanalysis data is shown in Fig. 12. The largest westerly (easterly) wind anomalies were found between 308 and 408S (508–608S and 108–208S) throughout the storm stages and at multiple vertical levels. These anomalies were associated with a region of deep and anomalously low pressure centered at 408S. At the 850- and 700-hPa levels, the anomalous west and north winds around the extratropical cyclone exceeded 13s (Figs. 12a and 12b) and 21.5s (not shown), respectively, on the Chilean coast during prefrontal conditions (i.e., 15 min after the satellite image time in Fig. 4). A similar wind pattern at

low levels was noted by Junker et al. (2008) from anomalous cyclonic circulation for heavy precipitation events in the Sierra Nevada. Nevertheless, its wind and u-component anomaly magnitudes were much less than in our case, but similar to the anomalies for the three major flooding events (cf. Fig. 12a with Figs. 8 and 14 in Junker et al. 2008). At 300 hPa (Fig. 12c), the u-wind anomalies were between 11.5s and 12.5s in the exit region of the jet stream with diffluent flow over central Chile and the Andean region, at the time of the highest rainfall rates. In the jet stream core (358–408S, 958–808W), the anomalies surpassed 13s, exhibiting a peak speed of 90 m s21. Thirty-six hours later (Fig. 12d), the maximum wind speed and westerly anomalies slightly decreased at 300 hPa (;60 m s21 and 12s, respectively) and were located over central Chile and west of Argentina (cf. the highest u winds of the rawinsonde in Fig. 11d). At the end of the event (not shown), the largest westerly wind anomalies at low and high levels were located over western Argentina, while a remarkably anomalous southerly wind was over the SPO (i.e., polar air incursion). As we introduced in section 3, the synoptic flow during the 26–29 August 2005 storm responded to intense low and high geopotential heights located off the southeast Pacific coast at mid- and subtropical latitudes, respectively. It was associated with strong negative (approximately south of 358S) and positive (approximately

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FIG. 12. The u-wind and height anomalies (shaded and contours, respectively), and the wind fields at 1800 UTC 26 Aug at (a) 850 and (b) 700, and at (c) 0600 UTC 27 Aug and (d) 1800 UTC 28 Aug 2005 at 300 hPa. The u-wind and height anomalies are normalized by the std dev according to the 1961–90 August monthly mean; wind barbs have the following units: half-barb, 5 m s21; full barb, 10 m s21; and pennant, 50 m s21.

north of 208S) geopotential height anomalies from the lower to upper levels (Fig. 12), resulting in a large meridional geopotential height gradient that drove a strong anomalous westerly flow. These geopotential height anomalies were configured into a well-defined dipole with center reaching 2s over the southeast Pacific coast (Fig. 12). It is important to note that the positive anomalies of the semipermanent anticyclone had the same or larger magnitude compared to the negative anomalies of the extratropical cyclone associated with frontal systems, so it could be an additional tool for anticipating potentially heavy rainfall. A similar dipole pattern was found for extreme cases over the Sierra Nevada in California, favoring strong anomalous southwesterly (cross barrier) winds (Junker et al. 2008; Pandey et al. 1999). The existence of anomalously strong northwest winds at low levels was confirmed in 14 of the 15 highest precipitation events in the subtropical central Andes for the period 1967–76, as recorded by the coastal rawinsondes (Table 2); the one unconfirmed event did not have a rawinsonde launch. During this period, there were at least four stations in the highlands areas that measured daily snowfall by an observer in situ. This

allowed for identification of the 15 heaviest episodes, which had a daily snowfall of 35 mm or greater. Table 2 shows that 13 of the 15 heaviest events had maximum speeds in the cross-mountain direction at 850 or 700 hPa that varied by 12.5s from the 1961–90 August climatological mean, while 8 of these 13 varied by 13s. The remaining 14th event had no rawinsonde observations on its maximum rainfall days; however, the u winds varied by 12s for the 2 days before it. For the alongbarrier winds, anomalous northerly winds were observed at 850 or 700 hPa during 10 of 14 events. Anomalous northwesterly winds were also observed at midlevels in these events, but less frequently than at lower levels (Table 2).

6. Model diagnosis of the 26–29 August 2005 storm The Eta-CPTEC 40-km regional model has already been used effectively as a diagnostic tool to explore physical mechanisms linked with subtropical Andean influences on the disruption of cold fronts that cross South America (Seluchi et al. 2006), and on downslope wind cases (Seluchi et al. 2003). These precedents of the Eta-CPTEC 40-km model, the current initial conditions

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TABLE 2. Maximum winds (m s21) observed at 850, 700 and 500 hPa at the coastal rawinsonde station (windward side of the Andes cordillera) for the 14 of 15 most intense wintertime (April–September) events for 1967–76. Wind values that vary by 63s, between 62.5 and 63s, or between 62s and 62.5s from the 1961–90 climatological August monthly mean are shown in boldface, underlining, and italics, respectively. Max 850-hPa wind

Max 700-hPa wind

Max 500-hPa wind

Event date

u

y

u

y

u

5–11 Jun 1969 19–24 Jun 1971 26–30 Jun 1972* 12–16 Aug 1972 24–29 Jun 1974 5–11 Jun 1969 13–17 Jul 1970 24–29 Jul 1970 5–9 May 1972 18–20 May 1973 19–21 May 1974 7–9 Jun 1974 2–4 Jul 1975 8–13 Jul 1975

10 15

2 18 222 29 221 238**

17 22 15 25 21 17 15 18 19 15 21 17 21 20

223 2 24 220 231 2 24

45 35

13 11 11 12 8 10 12 11 7 13 13

2 20 223 216 216 2 18

2 24 235

33 38 35 33 34

221

44

y 2 31 226 225

248 227

39 40

* Does not include rawinsonde observations during the maximum daily precipitation on 28 Jun 1972. ** The highest y-wind value at 700 hPa for 1961–90.

of the Eta-PRM model, indicate that the Eta-PRM should be reliable for use as a diagnostic tool. The Froude number (Fr) is widely used to check the flow regime of winter storm moving over mountain ranges. This number is defined as Fr 5 u/hN, where u is the crossmountain component wind, N the dry or moist Brunt– Vaïsa¨la¨ frequency, and h the height of the mountain ‘‘seen’’ by each pressure level (distance between the level and the top of the terrain). The Eta-PRM model output depicted a blocked or flow-around regime (i.e., Fr , 1) throughout the event at 925 and 850 hPa for three different latitudes on the windward side (not shown), mainly due to the stable profile and high altitude of the Andes. A vertical cross section of Eta-PRM model output at 338S shows the simulated kinematic structure’s evolution across the Andes in Figs. 13 and 14. At the time of PCF passage (Figs. 13a and 13b), the ascending air was evidenced almost throughout most of the column with the maximum upward motions reaching 28 Pa s21 on the windward slope near the mountaintop. On the lee side, the sinking air also had a large vertical extension, consistent with the presence of a vertically propagating mountain wave noticed by upward motion above 200 hPa (Fig. 13a). The westerly winds were relatively low, below 900 hPa, while the jet stream was noticed directly above the Andes crest, denoting the strong momentum transport upward (Fig. 13a). The meridional winds depicted a broad along-barrier maximum below crest level over the windward slope (Fig. 13b), and veering winds from northerly to southerly near the surface west of 728W, showing the approach of the PCF at the Chilean

coast. On the lee side, the isentropes abruptly descended nearer to the surface (Fig. 13b), consistent with strong downward motions, and a northerly flow at lower levels related with prefrontal conditions was present in central Argentina. Twelve hours later (Figs. 13c and 13d), the cold-air incursion arrived over central western Argentina, distinguished by southeasterly winds at lower levels (Fig. 13d). The weak near-surface easterly winds were associated with slight upward motions below 650 hPa over eastern Andean and Sierras de San Luis (;668W; see Fig. 1) slopes, while above them, there were downward motions associated with strong cross-barrier winds across the Andes (Fig. 13c). It can also be seen in Figs. 13c and 13d that the flow exhibited a strong crossbarrier vertical shear, separating the cross-barrier flow above from the weak flow below it, and there was also a well-defined along-barrier jet below crest level. This flow pattern on the windward side appreciably compares to cases of profoundly blocked flow patterns adjacent to mountain ranges during stable orographic storms, such as the Sierra Nevada (Marwitz 1987), the European Alps (Medina and Houze 2003; Bousquet and Smull 2003), and the Oregon Cascades (Medina et al. 2005). This simulated flow pattern is also consistent with the abovementioned blocked or flow-around regimes on upstream slopes on the western side of the Andes cordillera. During the next day (Fig. 14), the flow structure across the Andes remained relatively constant, consistent with the stationary wind profiles observed along coastal of central Chile (Figs. 9b and 9d). About the time of SCF passage (Fig. 14a), the cross-barrier flow

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FIG. 13. Vertical cross-mountain section at 338S showing (left) the vertical wind velocity v (Pa s21, solid line) and the u wind (cross barrier, m s21, shaded) and (right) the potential temperature (K, dashed line) and the y wind (along barrier, m s21, shaded) at (a),(b) 0600 UTC 27 Aug and (c),(d) 1800 UTC 27 Aug 2005; the thick black lines correspond to the 08C isotherm.

below 750 hPa was very weak, and above it was markedly increased, resulting in a strong cross-barrier shear layer (600–500 hPa). The along-barrier jet was also well defined between 700 and 600 hPa, reaching greater velocities of up to 14 m s21 (Fig. 14b). At 0000 UTC 29 August (Figs. 14c and 14d), the downward motion was reduced in magnitude along the lee side of the mountaintop, but the flow structure (along-barrier jet and the strong cross-barrier winds above) still remained. The polar air incursion in the ending stage of the storm is indicated in Fig. 14d by the overall drop of the isentropic values and the 08C isotherm height, particularly west of the Andes. The presence of the Andes also influenced the blocking of the eastward advance of the polar air at lower levels west of the mountains, denoted by the fairly colder airmass characteristics on the windward rather than leeward side (Fig. 14d). To provide a complement to the instantaneous snapshot of the vertical structure of the flow across the Andes shown above, we present in Figs. 15 and 16 time–

height plots of the simulated winds near and far upstream from the crest, allowing for the diagnosis of the vertical structure and the orographic effects. The main synoptic features of the storm become visible in Fig. 15: the strengthening of the westerly and northerly winds at lower levels around the PCF (;0000 UTC 27 August) and SCF (; midday on 28 August) time passage, and the jet streak around the 200-hPa level. The averaged flow plots over a four-latitude band (328–368S) also demonstrated the influence of the Andes decelerating the cross-barrier flow near the windward slope at nearsurface levels (cf. Figs. 15a and 15c), and increasing the along-barrier winds (cf. Figs. 15b and 15d) at low and middle levels. Also distinguishable is a slight increase in the cross-barrier winds into jet stream core above crest level. The difference between the near and far winds plots emphasizes these orographic effects in the northern (328–348S) and southern (348–368S) parts of the study area (Fig. 16). The blocking of the west wind and its southward turning of the flow below 800 hPa around

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FIG. 14. Same as in Fig. 13 but at (a),(b) 1800 UTC 28 Aug and (c),(d) 0000 UTC 29 Aug 2005.

the two cold front passages are more intense in the southern than northern part (Fig. 16), where the model topography is approximately 1 km lower (see Fig. 10). Alternatively, above 800 hPa, but still below crest level, this same orographic effect is clearer in the northerly parts. The simulation exhibited a near-constant alongbarrier jet at about 700 hPa from 1200 UTC 27 August to 0000 UTC 29 August, along the northerly part that is consistent with the upstream rawinsondes (cf. Figs. 16b and 9b–d). The observed along-barrier jet near the surface at 0000 UTC 27 August (Fig. 9a) is not well represented by the Eta simulation on the northern part (Figs. 9a and 16b), but appears to be present in the southern part (Fig. 16d), where there was a great difference in the cross- and along-barrier winds below 900 hPa. Unfortunately, there were no upstream rawinsonde observations in the southern part that verified the nearsurface along-barrier jet.

7. Discussion and concluding remarks An intense orographic precipitation event over the subtropical central Andes (308–368S) has been examined

in detail; only the wind anomalies for this heavy rainfall event and the 15 most intense events over a 10-yr period are investigated in this study. The comprehensively studied event occurred on 26–29 August 2005 and produced major flooding and landslides in low-lying areas of Santiago, Chile; heavy snowfall in the mountains; and downslope (zonda) winds in leeside lowlands. We simulated the event with the 15-km Eta (PRM) Model, which showed some deficiencies in forecasting lower precipitation rates and zonda near-surface winds to the lee of the Andes. The deficiencies observed were similar to those of previous studies in the mountainous region of the western United States (Colle and Mass 2000; Colle et al. 2000; Gallus 2000). The downslope wind forecasts could be improved using the sloping discretization of the eta coordinate in the Eta-PRM, as was noted for a zonda wind case (Mesinger et al. 2006). Despite its limitations, the Eta-PRM model produced a reliable forecast of heavy precipitations on windward-side lowlands and over the mountains, and offered the potential to be a useful tool for diagnosing orographic influences. The synoptic pattern accompanying the heaviest precipitation event in 2005 is characterized by not only the

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FIG. 15. Time–height (pressure level) plots of simulated upstream (left) cross- and (right) along-barrier flows at (a),(b) ;150 km and (c),(d) ;300 km distance from the crest of the Andes. The simulated upstream winds shown were averaged over 328–368S (i.e., where the maximum precipitation was observed and simulated in Fig. 10).

strong negative anomalies in the extratropical cyclone moving over the midlatitudes, but also by the same or larger magnitude of the positive anomalies in the South Pacific anticyclone at subtropical latitudes. This determined a well-shaped dipole pattern of pressure anomalies that produced a large meridional pressure gradient, resulting in strong anomalous cross-barrier flows. These synoptic characteristics could be useful in signaling a possible heavy precipitation event. Previous studies have documented synoptic patterns with a dipole trough– ridge pattern enhancing the cross-barrier flow for heavy composite events over the Sierra Nevada range (Pandey et al. 1999; Junker et al. 2008) and in the northwestern United States along the windward side of the Cascades (Lackmann and Gyakum 1999). A more intense orographic lifting and moisture crossbarrier flux producing heavy orographic precipitation is expected as a result of larger departures from climatology in the cross-barrier flow. This is indirectly supported by the fact that 14 of the 15 highest precipitation events over a 10-yr period, and also the 26–29 August 2005

event, exhibited cross-mountain winds exceeding 2s. The along-barrier flow also exhibited anomalous northerly winds in 10 of the 15 heaviest events and in the 2005 event, denoting the anomalous cyclonic circulation and/or the generation of an along-barrier jet as a consequence of the majority of the blocked cases. These findings agree with those of Falvey and Garreaud (2007), who showed that heavy precipitation enhancement in the Andes occurred in the warmer and prefrontal sector of the strong frontal episodes (with northwesterly winds), and had stronger midlevel zonal wind speeds, and more low-level stability than nonenhanced events. The previous work of Junker et al. (2008) also found similar strong anomalous cross-barrier flows leading to heavy precipitation episodes over the Sierra Nevada. On the other hand, the Eta-PRM model closely forecasted the wind speed observed during the heaviest 2005 event, so one might conclude that at least the model output with such anomalous values could anticipate a possible extreme precipitation event. Hence, this would be an additional valuable tool to be taken into account by the forecasters.

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FIG. 16. Same as in Fig. 15 but for the difference for the near (;150 km) minus far (;300 km) upstream (left) cross- and (right) along-barrier flows over (a),(b) 328–348S and (c),(d) 348–368S.

Despite the fact that the Eta-PRM model (with a resolution of 15 km) has a smooth step topography and simulated relatively coarse patterns of airflow and precipitation across the mountain, its diagnosis allowed the inference of the main topographic influence on the flow structure that led to the significant orographic precipitation on 26–29 August 2005. The strong flow that impinged on the Andes was profoundly blocked and then diverted into a strong along-barrier jet below crest level that was clearly separated from the strong cross-barrier winds above. The upslope motion reached its maximum near the mountaintop and downslope motion reached it immediately at the lee side. These topographically induced vertical motions would explain the enhanced precipitation pattern over the upslope and the rapid decrease in the precipitation on the lee side. The stable conditions during the storm and the topography features (i.e., a mean height between 5 and 3.5 km and a north–south disposition) appear to be the main reasons of this flow structure, as well as the ensuing precipitation pattern. It is likely that the stable vertical profile reduced the forced ascent of the strong cross-barrier flow

over the lower slope when compared with unstable events, and then would play a role in the reduced accumulated precipitation. We cannot support this with one stable case study from the Andes; however, it was observed by Medina and Houze (2003) over the European Alps during two heavy events under distinct stability conditions. Even though the precipitation in our case study occurred under stable conditions, it was significant and produced considerable damage that deserves a forecaster’s attention. During other field projects of similar midlatitude mountain ranges around the world, such as the Sierra Cooperative Pilot Project (SCPP) of the Sierra Nevada, IMPROVE of the Cascades, and MAP of the Alps, many winter storms have been observed in detail using advanced dual-Doppler and airborne radar, and dense networks. Some studies from these projects have documented finescale airflow structures during winter storms under stable flow around regimes that compare to the coarse patterns of the flow captured by the Eta-PRM model and the upstream rawinsondes (e.g., Marwitz 1987; Medina and Houze 2003; Garvert et al. 2007).

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These similarities suggest that the finescale precipitation enhancement mechanisms proposed by Medina et al. (2005), Houze and Medina (2005), and Garvert et al. (2007), among others, may have been operating during the 2005 storm over the Andes. However, in our case, given the limited observations available, and the limitations in the Eta-PRM model for representing flow structures (with a resolution of 15 km), we can only suppose that these finescale mechanisms operated during the event. Future detailed radar and in situ observations are needed to gain insights into the precipitation growth mechanisms over the steep slopes of the central Andes. The results of this work allow us to infer some initial suggestions corresponding to environmental conditions and flow structures that lead to heavy orographic precipitation over the subtropical central Andes that may be useful for forecasters. Little is known about this orographic phenomenon in this region, which has a significant impact on the regional climate system. Therefore, comparing cases under different flow regimes using higher resolution in numerical models could be the natural next step, as well as the climatologic studies from many events, which would provide a more comprehensive understanding of the orographic influences in precipitation of the central Andean region. Acknowledgments. The authors thank Dr. M. Seluchi from CPTEC-Brazil for helping in the Eta–PRM model implementation. We are also grateful to Eng. G. Pereyra from the DGI of Mendoza, to SMN-Argentina, and to R. Leoń Gallardo and C. Belmar from CENMA, for the provision of the SNOTEL, SYNOP, and snowfall at Lagunitas data, respectively. Dr. D. Araneo and the anonymous reviewers made detailed and insightful comments. M. Silva provided the model computational support and J. Cristaldo helped with the data management. The authors wish to thank Dr. Gregory Hoke for his help in revising the English in our manuscript.

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