Tropical Transition of an Unnamed, High-Latitude, Tropical Cyclone over the Eastern North Pacific
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Tropical Transition of an Unnamed, High-Latitude, Tropical Cyclone over the Eastern North Pacific Alicia M. Bentley1 Department of Atmospheric and Environmental Sciences, University at Albany, State University of New York, Albany, New York Nicholas D. Metz Department of Geoscience, Hobart and William Smith Colleges, Geneva, New York Submitted to Monthly Weather Review 5 June 2015 ABSTRACT In early November 2006, an unnamed tropical cyclone (TC) formed via the tropical transition (TT) process at ~40°N over the eastern North Pacific. An extratropical cyclone (EC), developing downstream of a thinning upper-tropospheric trough over the Gulf of Alaska, served as the precursor disturbance that would ultimately undergo TT. The TT of the unnamed TC was extremely unusual—occurring over ~16°C sea surface temperatures in a region historically devoid of TC activity. This paper (1) identifies the upper- and lower-tropospheric features linked to the formation of the EC that transitions into the unnamed TC, (2) provides a synoptic overview of the features and processes associated with the unnamed TC’s TT, and (3) discusses the landfall of the weakening cyclone along the west coast of North America. As observed in previous studies of the TT process in the North Atlantic basin, the precursor EC progresses through the life cycle of a marine extratropical frontal cyclone, developing a bent-back warm front on its northern and western sides prior to the isolation of its central circulation. Vertical cross sections taken through the center of the cyclone during its life cycle reveal its transformation from an asymmetric, cold-core, EC into an axisymmetric, warm-core, TC during the TT process.
Tropical cyclones (TCs) are not exclusive to the tropics. While the environmental conditions deemed favorable for tropical cyclogenesis in the seminal work of Palmén (1948), Gray (1968), and DeMaria et al. (2001) are typically observed at tropical latitudes, environmental conditions can become favorable for tropical cyclogenesis in locations removed from the tropics. The global climatology of tropical cyclogenesis events constructed by McTaggart-Cowan et al. (2013) reveals that the majority of TCs forming poleward of 25°N (25°S) in the Northern (Southern) Hemisphere during 1948–2010 developed in the presence of an upper-tropospheric disturbance in a baroclinic environment (their Fig. 7). TCs developing in the presence of an upper-tropospheric disturbance in a baroclinic environment typically form via the tropical transition (TT) process (Davis and Bosart 2003, 2004), during which an asymmetric, cold-core, extratropical cyclone (EC) transitions into an axisymmetric, warm-core, TC. In the initial stages of the TT process, vertical wind shear in a baroclinic environment produces a region of upward motion that focuses deep convection and diabatic heating (Sutcliffe 1947). Vertical wind shear values are subsequently reduced by the diabatic redistribution of potential vorticity (PV) in the vertical (e.g., Hoskins et al. 1985; Davis and Emanuel 1991; Raymond 1992; Stoelinga 1996; Grams et al. 2011; Čampa and Wernli 2012) and by divergent outflow in the upper troposphere, allowing the surface cyclone to intensify via air–sea interaction processes [i.e., wind-induced surface heat exchange (Emanuel 1986)]. Deep convection and diabatic heating upshear of the surface cyclone are also associated with the generation of lower-tropospheric vorticity that enhances cold air advection (CAA) on the cyclone’s northern and western sides, helping to isolate the cyclone’s developing warm core (e.g., Hulme and Martin 2009b; Galarneau et al. 2015). TCs forming via the TT process have been documented in many basins where tropical cyclogenesis events occur annually, including the western North Atlantic (e.g., Moore and Davis 1951; Bosart and Bartlo 1991; Bracken and Bosart 2000; Davis and Bosart 2001; McTaggart-Cowan et al. 2006; Evans and Guishard 2009; Guishard et al. 2009; Hulme and Martin 2009a,b), western North Pacific (e.g., Wang et al. 2008), and western South Pacific (e.g., Garde et al. 2010; Pezza et al. 2013). TCs forming via the TT process have also been documented in basins where tropical cyclogenesis events are extremely rare, including the eastern North Atlantic (e.g., Case 1990; Franklin 2006; Beven 2006), western South Atlantic (e.g., Pezza and Simmonds 2005; McTaggart-Cowan et al. 2006; Evans and Braun 2012; Gozzo et al. 2014), and Mediterranean Sea (e.g., Ernst and Matson 1983; Pytharoulis et al. 1999; Reale and Atlas 2001; Emanuel 2005; McTaggart-Cowan et al. 2010). In early November 2006, an unnamed TC [designated “Invest 91C” by the Central Pacific Hurricane Center (CPHC)] developed at ~40°N in the eastern North Pacific basin (Fig. 1a). An EC, forming downstream of a thinning upper-tropospheric trough over the Gulf of Alaska, served as the precursor disturbance that would ultimately undergo TT. The TT of Invest 91C, which took place between 0000 UTC 28 October 2006 and 0000 UTC 2 November 2006, was extremely unusual—occurring over ~16°C sea surface temperatures (SSTs; Fig 1b) in a region historically devoid of TC activity. The remnants of Invest 91C would ultimately make landfall along the northwest (southwest) coast of Washington (British Columbia) at approximately 1600 UTC 3 November 2006, with measured wind gusts at Destruction Island, WA, in excess of 29 m s−1. The remainder of this paper is organized as follows. The data and methodology used to analyze the development, TT, and landfall of Invest 91C are described in section 2. The upper- and lower-tropospheric features linked to the formation of the EC that transitioned into Invest 91C are identified and discussed in section 3. A synoptic overview of the life cycle of Invest 91C is presented in section 3 in order to document the features and processes associated with its TT and landfall. Key findings and conclusions are contained in section 5.
Gridded datasets of the National Centers of Environmental Prediction (NCEP) Climate Forecast System Reanalysis (CFSR; Saha et al. 2010), with 0.5° × 0.5° horizontal grid spacing and 6-h temporal resolution, constitute the primary data source for analyzing the development, TT, and landfall of Invest 91C. The upper- and lower-tropospheric features linked to the formation of the EC that transitioned into Invest 91C are examined every 24 h between 0000 UTC 24 October 2006 and 0000 UTC 28 October 2006 (hereafter, all dates in 2006) in section 3. The upper-tropospheric features linked to the formation of the EC that transitioned into Invest 91C are summarized using a hovmöller diagram (Hovmöller 1949) of 250-hPa meridional winds obtained every 6 h between 0000 UTC 22 October and 0000 UTC 30 October. Standardized anomalies, used to assess the characteristics of the 250-hPa meridional wind field, are computed from the long-term (1979–2009) climatology derived from the 0.5° NCEP CFSR dataset. A synoptic overview of the life cycle of Invest 91C will be presented in section 4 using four-panel diagrams plotted every 24 h between 0000 UTC 28 October and 0000 UTC 3 November. A four-panel diagram depicting Invest 91C at 1800 UTC 3 November (i.e., shortly after landfall) will also be presented for completeness. All four-panel diagrams are centered on position of the minimum mean sea level pressure (MSLP) value associated with Invest 91C in the 0.5° NCEP CFSR dataset. A track map indicating the position of Invest 91C every 6 h between 0000 UTC 28 October and 1800 UTC 3 November is shown in Fig. 2. The colored dots correspond to the time periods analyzed using four-panel diagrams in section 4. Vertical cross sections, taken through the center of Invest 91C at 0000 UTC 29 October and 0000 UTC 2 November, are presented in section 4 to illustrate the cyclone’s transition from an asymmetric, cold-core, EC into an axisymmetric, warm-core, TC. The coupling index, a measure of troposphere-deep instability, is defined as the difference between potential temperature (θ) on the dynamic tropopause (DT) and equivalent potential temperature (θe) at 850 hPa (Bosart and Lackmann 1995). Infrared (IR) brightness temperature, used to determine the location of deep convection in the four-panel diagrams presented in section 4, was obtained from the NCEP/Climate Prediction Center 4 km global IR dataset. The SSTs displayed in Fig. 1b were obtained from analyses of 0.25° Advanced Very High Resolution Radiometer data (Reynolds et al. 2007), available once daily from the National Climatic Data Center.
The upper- and lower-tropospheric features linked to the formation of the EC that ultimately transitions into Invest 91C began to interact over the extratropical North Pacific on 24 October. A meridionally confined 1000–500-hPa thickness gradient, located between a 996-hPa surface cyclone ~350 km south of the Aleutian Islands and a 1036-hPa surface anticyclone ~1400 km west of California, spans the North Pacific at 0000 UTC 24 October (Fig. 2a). This zonally elongated baroclinic zone is associated with a >70 m s−1 upper-tropospheric jet that extends from northern Japan to western Canada (Figs. 2a,b). A broad surface low begins to develop in the equatorward entrance region of the upper-tropospheric jet at 0000 UTC 24 October (EC1), downstream of an upper-tropospheric disturbance located over the Sea of Japan (Fig. 2b). Differential cyclonic vorticity advection by the thermal wind, implied in Figs. 2a,b, helps focus upward motion and diabatic heating over the western portion of EC1. This focused diabatic heating reduces MSLP values over the western portion of EC1 and shifts the MSLP minimum toward eastern Japan over the following 24 h (Figs. 2a,c). A second EC over southeastern Russia (EC2), located downstream of an upper-tropospheric trough in the 300–200-hPa layer-averaged PV field, begins to perturb the western portion of the North Pacific waveguide by 0000 UTC 25 October (Figs. 2c,d). Warm air advection (WAA) to the east of EC2 is associated with the amplification of the downstream ridge in the upper and lower troposphere (Figs. 2c,d). Upper-tropospheric divergent outflow, emanating from a region of 600–400-hPa layer-averaged ascent collocated with the center of EC2, results in negative PV advection in the upper troposphere that slows the eastward progression of the upstream trough, contributes to rapid downstream ridge amplification, and enhances northwesterly flow downstream of the ridge axis between 0000 UTC 24 October and 0000 UTC 25 October (Figs. 2b,d). EC1 and EC2 deepen and move northeastward between 0000 UTC 25 October and 0000 UTC 26 October (Figs. 2c,e). A corridor of WAA to the east of EC1 and EC2 extends from ~40°N to ~60°N by 0000 UTC 26 October (Fig. 2e), further amplifying the downstream ridge. Negative PV advection by the 300–200-hPa layer-averaged irrotational wind continues to contribute to rapid ridge amplification and enhances northwesterly flow downstream of the ridge axis between 0000 UTC 25 October and 0000 UTC 26 October (Figs. 2d,f). Enhanced northwesterly flow downstream of the ridge axis is associated with the formation and amplification of a positively tilted upper-tropospheric trough over the central North Pacific by 0000 UTC 26 October, inferred from the orientation of PV contours in the 300–200-hPa layer averaged PV field (Fig. 2f). A narrow corridor of relatively low MSLP values develops downstream of the upper-tropospheric trough over the eastern North Pacific, in the equatorward entrance region of the North Pacific jet, at 0000 UTC 26 October (Fig. 2e). This narrow corridor of relatively low MSLP values amalgamates into a ~1012-hPa closed low over the following 24 h as the positively tilted upper-tropospheric trough amplifies upstream (Figs. 2e,g). Negative PV advection by the 300–200-hPa layer-averaged irrotational wind on the east side of the positively tilted upper-tropospheric trough causes the trough to slow, stretch, and thin between 0000 UTC 26 October and 0000 UTC 27 October (Figs. 2f,h). Continued stretching and thinning occurs over the following 24 h, resulting in the equatorward movement of the closed low to 40.5°N, 147.5°W by 0000 UTC 28 October (Figs. 2g–j). The amplification of the North Pacific waveguide that results in the formation of the closed low located at 40.5°N, 147.5°W by 0000 UTC 28 October is summarized in a hovmöller diagram of 250-hPa meridional winds averaged between 40°N and 60°N (Fig. 4). Figure 3 highlights the eastward propagation of a Rossby wave train (Namias and Clapp 1944; Chang 1993; Orlanski and Sheldon 1993, 1995; Hakim 2003; Danielson et al. 2004) across the North Pacific basin between 0000 UTC 24 October and 0000 UTC 28 October. The upper-tropospheric trough associated with EC2 (Figs. 2a,b) is evident over eastern Asia at 0000 UTC 24 October, with the trough axis denoted by a “T” and a shift from northerly to southerly meridional winds at 250 hPa. WAA and divergent outflow emanating from regions of deep convection associated with EC1 and EC2 (Figs. 2a,b) help to enhance the downstream ridge over the following 24 h, with the ridge axis denoted by an “R” and a shift from southerly to northerly meridional winds at 250 hPa. Enhanced northerly flow downstream of the ridge axis between 0000 UTC 25 October and 0000 UTC 26 October is associated with the formation and amplification of an upper-tropospheric trough over the central North Pacific, with the trough axis denoted by a “T” and a shift from northerly to southerly meridional winds at 250 hPa. Negative PV advection by the 300–200-hPa layer-averaged irrotational wind on the western and eastern sides of the upper-tropospheric trough (Figs. 2f,h,j) causes the trough to slow, stretch, and thin between 0000 UTC 26 October and 0000 UTC 28 October. Figure 3 reveals northerly (southerly) 250-hPa meridional winds in excess of −2 σ (+2 σ) upstream (downstream) of the trough axis, suggesting that the upper-tropospheric trough associated with the formation of the closed low located at 40.5°N, 147.5°W is an anomalous feature in the eastern North Pacific basin in late October.
The ~1012-hPa elongated closed low, located at 40.5°N, 147.5°W at 0000 UTC 28 October, exists on the warm-side of a 925–500-hPa thickness gradient in a region of negative 925–500-hPa thermal vorticity (ζT) (Fig. 5a). The strip of 925–850-hPa layer-averaged relative vorticity associated with this ζT minimum is located downstream of the thinning upper-tropospheric trough identified in section 2 (Fig. 5b). A meridionally elongated corridor of >40 mm precipitable water (PW) values extends poleward downstream of the thinning upper-tropospheric trough, coinciding with a region of 600–400-hPa layer-averaged ascent (Fig. 5c). Upper-tropospheric divergent outflow to the west of the region of 600–400-hPa layer-averaged ascent tightens the upper-tropospheric PV gradient and slows the eastward propagation of the southern portion of the thinning upper-tropospheric trough. The cloud field at 0000 UTC 28 October aligns nicely with the region of 600–400-hPa layer-averaged ascent (Fig. 5d). Differential cyclonic vorticity advection over the center of the closed low is associated with the formation of an ~1004-hPa EC over the following 24 h (Figs. 5a and 6a). A region of WAA (CAA) begins to wrap around northern (southern) edge of the EC by 0000 UTC 29 October, resulting in a discernable warm (cold) front in the 925–500-hPa thickness and ζT fields (Fig. 6a). The southern portion of the thinning upper-tropospheric trough has become cut off from the northern portion at this time (Fig. 6b), allowing the corridor of warm, moist, air and 600–400-hPa layer-averaged ascent in Figs. 5a,c to begin to wrap cyclonically around the northern edge of the EC (Figs. 6a,c). A vertical cross section, taken along line A–A’ in Fig. 6d, highlights the different air masses wrapping cyclonically around the ~1004-hPa EC (Fig. 7). The corridor of warm, moist, air to the north of the EC in Figs. 6a,c is indicated by >314 K θe values and by relatively high θ values throughout the troposphere to the immediate north of the MSLP minimum. A region of cold, dry, air to the south of the EC in Figs. 5a,c is indicated by relatively low midtropospheric θe and θ values in the to the south of the MSLP minimum. PV values >2 PVU, associated with the upper-tropospheric cutoff in Fig. 5b, extend from the stratosphere to ~475 hPa to the south of the EC, associated with a region of highly unstable air in the midtroposphere. The corridor of warm, moist, air wraps around the western edge of the EC by 0000 UTC 30 October, resulting in the bent-back warm front visible in the 925–500-hPa ζT field (Fig. 8a). Negative ζT values, which coincide with the 925–850-hPa layer-averaged relative vorticity maximum, have become vertically aligned with the upper-tropospheric cutoff (Figs. 8a,b). Values of the coupling index are <−2.5 K over the EC center at this time (not shown), suggesting the continued presence of unstable air in the midtroposphere beneath the upper-tropospheric cutoff. Midtropospheric ascent and PW values >35 mm are collocated with the upper-tropospheric cutoff (Fig. 8c), creating a favorable environment for the development of deep convection near the EC center (Fig. 8d). Diabatic heating, associated with weak convection near the EC center, allows the cyclone to begin to gain TC-like characteristics by 0000 UTC 31 October. Negative ζT values associated with the cyclone center have broken away from the bent-back warm front (Fig. 9a) and remain collocated with the upper-tropospheric cutoff at this time (Fig. 9b). An increase in DT θ over the cyclone center (Fig. 9b), corresponding to a reduction in 300–200-hPa layer-averaged PV values (Fig. 9c), suggests that the upper-tropospheric cutoff is being eroded by the diabatic redistribution of PV in the vertical. Midtropospheric ascent continues to occur to the north and east of the center of circulation in a region of >35 mm PW values (Fig. 9c), coinciding with the position of a spiral band seen in the IR satellite imagery (Fig. 9d). The central pressure of the transitioning cyclone has dropped below 998 hPa by 0000 UTC 1 November (Fig. 10a). ζT values have decreased to <−14 × 10−5 s−1 near the cyclone center, with a reduction in ζT values observed throughout the majority of the analyzed domain (Figs. 8a and 9a). The 925–850-hPa layer-averaged relative vorticity maximum has moved northward in response to a progressive upper-tropospheric trough that wrapped around the southern portion of the transitioning cyclone during the previous 24 h (Figs. 8b and 9b). Midtropospheric ascent is observed surrounding, but not collocated with, the center of the transitioning cyclone at this time (Fig. 10c). Contrary to the plotted analysis, a region of deep convection (indicated by cold cloud tops in the IR satellite imagery in Fig. 10d) is collocated with the cyclone center. This discrepancy is likely due to the relatively coarse resolution of the 1° NCEP Global Forecast System FNL analysis, the lack of surface observations over the eastern North Pacific, and the relatively small size of the transitioning cyclone. The TC-like characteristics of the transitioning cyclone at 0000 UTC 1 November (Figs. 9a–d) caused the Central Pacific Hurricane Center to label the storm “Invest 91C” at 1200 UTC 1 November. The central pressure of Invest 91C decreases to <992 hPa by 0000 UTC 2 November (Figs. 9a and 10a), while maintaining 925–850-hPa layer-averaged relative vorticity values of >2.0 × 10−4 s−1 near the cyclone center (Figs. 9b and 10b). Midtropospheric ascent continues to be poorly represented near the region of deep convection surrounding the cyclone center at this time (Figs. 10c,d), likely due to the reasons previously discussed. A vertical cross section, taken along line B–B’ in Fig. 11d, indicates that Invest 91C has completely transitioned into an axisymmetric, warm-core, TC by 0000 UTC 2 November (Fig. 12). Relatively high θ values located over the center of Invest 91C confirm the warm-core structure of the cyclone suggested in Fig. 11a, which is shown to extend throughout the depth of the troposphere. Vertically orientated θe contours surrounding the MSLP minimum suggest that deep convection has been occurring in this region. The upper-tropospheric PV maximum that extended from the stratosphere to ~475 hPa at 0000 UTC 29 October (Fig. 7) has been eroded by deep convection (Fig. 12). A PV tower, also indicative of the diabatic redistribution of PV in the vertical, is located over the MSLP minimum between the surface and ~400 hPa. DISCUSSION OF 0000 UTC 3 November and LANDFALL!
The TT of Invest 91C occurred at ~40°N in the eastern North Pacific between 0000 UTC 28 October 2006 and 0000 UTC 2 November 2006. Despite occurring in a basin where tropical cyclogenesis events are extremely rare, the features and processes associated with Invest 91C’s transition from an asymmetric, cold-core, EC into an axisymmetric, warm-core, TC are consistent with those found in previous TT studies. This study supports the findings of Davis and Bosart (2003, 2004) and Hulme and Martin (2009a,b) that an EC, developing in association with an upper-tropospheric trough approaching a lower-tropospheric baroclinic zone, serves as the precursor disturbance to TT. As suggested by Hulme and Martin (2009b), the EC progresses through the life cycle of a marine extratropical frontal cyclone (Shapiro and Keyser 1990), developing a bent-back warm front on its northern and western sides during the TT process. Deep convection along the bent-back warm front diabatically redistributes PV in the vertical over the cyclone center, eroding upper-tropospheric PV over the lower-tropospheric relative vorticity maximum. The lower-tropospheric relative vorticity maximum ultimately separates from the bent-back warm front, allowing the cyclone to completely transition into an axisymmetric, warm-core, TC. It is important to note that Invest 91C, though a TC, is not included in the International Best Track Archive for Climate Stewardship (IBTrACS) dataset (Knapp et al. 2010) or in the global climatology of tropical cyclogenesis events constructed by McTaggart-Cowan et al. (2013). The authors suggests that Invest 91C, as well as TCs forming via the TT process in other unusual basins (i.e., the Mediterranean Sea), should be included in these datasets in order to obtain a comprehensive understanding of the frequency and location of tropical cyclogenesis events around the globe. Acknowledgments. We thank Drs. Ryan Torn (University at Albany) and Kristen Corbosiero (University at Albany) for helpful discussions and research assistance. This research was funded by the Hobart and William Smith Colleges Provost’s Office and NSF Grant AGS-1355960. REFERENCES Beven, J. L., 2006: Tropical cyclone report: Tropical Storm Delta, 22–28 November 2005. National Hurricane Center. [Available online at http://www.nhc.noaa.gov/pdf/TCR-AL292005_Delta.pdf.] Bosart, L. F., and J. A. Bartlo, 1991: Tropical storm formation in a baroclinic environment. Mon. Wea. Rev., 191, 1979–2013. Bosart, L. F., and G. M. Lackmann, 1995: Postlandfall tropical cyclone reintensification in a weakly baroclinic environment: A case study of Hurricane David (September 1979). Mon. Wea. Rev., 123, 3268–3291. Bracken, W. E., and L. F. Bosart, 2000: The role of synoptic-scale flow during tropical cyclogenesis over the North Atlantic Ocean. Mon. Wea. Rev., 128, 353–376. Čampa, J., and H. Wernli, 2012: A PV perspective on the vertical structure of mature midlatitude cyclones in the Northern Hemisphere. J. Atmos. Sci., 69, 725–740. Case, R. A., 1990: Preliminary tropical cyclone report: Tropical Storm Edouard, 02–11 August 1990. National Hurricane Center. [Available online at http://www.nhc.noaa.gov/archive/storm_wallets/atlantic/atl1990-prelim/edouard/.] Chang, E. K. M., 1993: Downstream development of baroclinic waves as inferred from regression analysis. J. Atmos. Sci., 50, 2038–2053. Danielson, R. E., J. R. Gyakum, and D. N. Straub, 2004: Downstream baroclinic development among forty-one cold-season eastern North Pacific cyclones. Atmos.–Ocean, 42, 235–250. Davis, C. A., and L. F. Bosart, 2001: Numerical simulations of the genesis of Hurricane Diana (1984). Part I: Control simulation. Mon. Wea. Rev., 129, 1859–1881. Davis, C. A., and L. F. Bosart, 2003: Baroclinically induced tropical cyclogenesis. Mon. Wea. Rev., 131, 2730–2747. Davis, C. A., and L. F. Bosart, 2004: The TT problem. Bull. Amer. Meteor. Soc., 85, 1657–1662. Davis, C. A., and K. A. Emanuel, 1991: Potential vorticity diagnostics of cyclogenesis. Mon.
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Extratropical Cyclones: The Erik Palmén Memorial Volume, C. W. Newton and E. Holopainen, Eds., Amer. Meteor. Soc., 167–191. Stoelinga, M. T., 1996: A potential vorticity-based study of the role of diabatic heating and friction in a numerically simulated baroclinic cyclone. Mon. Wea. Rev., 124, 849–874. Sutcliffe, R. C., 1947: A contribution to the problem of development. Quart. J. Roy. Meteor. Soc., 73, 370–383. Wang, W., W. Zhu, Y. Duan, H. Yu, L. Jiang, and X. Wang, 2008: Statistical analysis of large- scale backgrounds classification associated with tropical cyclone formations in the western North Pacific. J. Nanjing Inst. Meteor., 31, 277–286. FIGURE CAPTIONS FIG. 1. (a) GOES-10 visible satellite image, valid at 0045 UTC 2 November 2006, showing Invest 91C centered at ~42°N, 145°W. (b) Mean sea level pressure (black contours, hPa) from the 0.5° CFSR dataset, valid at 0000 UTC 2 November 2006, overlaid on infrared satellite sea surface temperatures (shaded, °C), obtained from analyses of 0.25° Advanced Very High Resolution Radiometer (AVHRR) data, valid on 2 November 2006. FIG. 2. Fig. 2. Track map indicating the position of the center of Invest 91C every 6 h between 0000 UTC 28 October 2006 (genesis) and 1800 UTC 3 November 2006 (shortly after landfall), identified using the minimum mean sea level pressure value in the 0.5° CFSR dataset. The positions plotted in color correspond to the time periods analyzed in section 3. The minimum mean sea level pressure value associated with the center of Invest 91C is shown for the time periods analyzed in section 3. FIG. 3. Synoptic analysis of mean sea level pressure (black contours, hPa), 1000–500-hPa thickness (red dashed contours, dam), and 250-hPa wind speed (shaded, m s−1) at (a) 0000 UTC 24 October, (c) 0000 UTC 25 October, (e) 0000 UTC 26 October, (g) 0000 UTC 27 October, and (i) 0000 UTC 28 October 2006; Precipitable water (grayscale, mm), 300–200-hPa layer-averaged potential vorticity (gray contours, PVU), 300–200-hPa layer-averaged irrotational wind (vectors, starting at 3 m s−1), 250-hPa wind speed (shaded, m s−1), and 600–400-hPa layer-averaged ascent (red contours every 2.5 × 10−3 hPa s−1), at (b) 0000 UTC 24 October, (d) 0000 UTC 25 October, (f) 0000 UTC 26 October, (h) 0000 UTC 27 October, and (j) 0000 UTC 28 October 2006. The “EC1” and “EC2” labels in (a), (c), (e), (g), and (i) denote the location of surface cyclone centers. FIG. 4. Hovmöller diagram of 250-hPa meridional wind (black contours, m s−1) and 250-hPa meridional wind standardized anomalies (shaded, σ), averaged between 40°N and 60°N, from 0000 UTC 22 October 2006 to 0000 UTC 30 October 2006. The “T” and “R” labels denote the position of 250-hPa trough and ridge axes, respectively. FIG. 5. Synoptic analysis of (a) mean sea level pressure (black contours, hPa), 925–500-hPa thickness (red dashed contours, dam), and 925–500-hPa thermal vorticity (shaded, 10−5 s−1), (b) dynamic tropopause (DT, 2-PVU surface) potential temperature (shaded, K), DT winds (barbs, kts), and 925–850-hPa layer-averaged relative vorticity (black contours every 0.5 × 10−4 s−1), (c) precipitable water (grayscale, mm), 300–200-hPa layer-averaged potential vorticity (gray contours, PVU), 300–200-hPa layer-averaged irrotational wind (vectors, starting at 3 m s−1), 250-hPa wind speed (shaded, m s−1), and 600–400-hPa layer-averaged ascent (red contours every 2 × 10−3 hPa s−1), and (d) infrared (IR) brightness temperature (shaded, °C) at 0000 UTC 28 October 2006. FIG. 6. As in Fig. 5, but at 0000 UTC 29 October 2006. Line A–A’ in (d) denotes the position of a vertical cross section shown in Fig. 7. FIG. 7. Vertical cross section along line A–A’ in Fig. 6d depicting potential vorticity (shaded, PVU), equivalent potential temperature (green contours, K), and potential temperature (gray contours, K) at 0000 UTC 29 October 2006. “L” represents the position of the mean sea level pressure minimum along the cross section. FIG. 8. As in Fig. 5, but at 0000 UTC 30 October 2006. FIG. 9. As in Fig. 5, but at 0000 UTC 31 October 2006.
FIG. 14. As in Fig. 5, but at 1800 UTC 3 November 2006. The magenta dot in (d) denotes the location of the Destruction Island, WA, lighthouse. FIG. 15. Meteogram depicting wind speed (blue line, m s−1), wind gusts (red line, m s−1), and pressure (green line, hPa), measured at the Destruction Island, WA, lighthouse (DESW1), from 0000 UTC 31 October 2006 to 1200 UTC 4 November 2006. DESW1 data obtained from the National Data Buoy Center (NDBC) data archive (available online at http://www.ndbc.noaa.gov/ station_history.php?station=desw1). 1 Corresponding author address: Alicia M. Bentley, Department of Atmospheric and Environmental Sciences, University at Albany, State University of New York, ES-234, 1400 Washington Ave., Albany, NY 12222. Email: ambentley@albany.edu Directory: student -> abentley -> research images -> ne pac tc -> cfsr -> writing student -> Comparing Hurricanes and Hurricane Seasons student -> Storm Surge: a rising Danger student -> Dr. Jeff Masters' WunderBlog student -> April 9, 2008 Hurricanes (Continued) student -> Instructions for Writing Hurricane Term Paper writing -> Tropical Transition of an Unnamed, High-Latitude, Tropical Cyclone over the Eastern North Pacific ne pac tc -> References List (Updated: 8 January 2015) Bosart, L. F., and J. A. Bartlo, 1991: Tropical storm formation in a baroclinic environment. Mon. Wea. Rev., 191 ne pac tc -> Tropical Transition in the Eastern North Pacific: Sensitivity to Microphysics research images -> AL192011 Sean pdf.] research images -> 1. Introduction The National Hurricane Center (nhc) online glossary defines a subtropical cyclone (stc) as a “non-frontal low-pressure system that has characteristics of both tropical and extratropical cyclones… Download 88.68 Kb. Share with your friends:
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