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 unusualoccurring 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 near the United States/Canada border. 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.





  1. Introduction

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. During this TT process, the EC often exhibits characteristic features of an evolving marine EC (i.e., bent-back warm front; Shapiro and Keyser 1990; Hulme and Martin 2009a,b). 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 unusualoccurring 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 extreme 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. This paper documents the formation of the EC that eventually became Invest 91C, the TT of this EC, and the subsequent evolution of the TC over the eastern North Pacific Ocean.

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 4 in order to document the features and processes associated with its TT and landfall. Key findings and conclusions are contained in section 5.


  1. Data and methodology

Gridded datasets of the National Centers of Environmental Prediction Climate Forecast System Reanalysis (NCEP 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 from 0000 UTC 24 October 2006 through 0000 UTC 28 October 2006 (hereafter, all dates in 2006) in section 3. The upper-tropospheric features deemed important to the formation of the EC that transitioned into Invest 91C are summarized using a Hovmöller diagram (Hovmöller 1949) of 250-hPa meridional winds plotted every 6 h from 0000 UTC 22 October through 0000 UTC 30 October. The standardized anomalies presented in the Hovmöller diagram, used to assess the characteristics of the 250-hPa meridional wind field, are computed from a long-term (1979–2009) climatology derived from the 0.5° NCEP CFSR dataset.

Figure 2 indicates the position of the center of Invest 91C every 6 h from 0000 UTC 28 October (genesis as a weak EC) through 1800 UTC 3 November (shortly after landfall). The position of the center of Invest 91C was identified using the minimum mean sea level pressure value (MSLP) associated with the cyclone within the 0.5° NCEP CFSR dataset. The colored dots in Fig. 2 correspond to the time periods analyzed in the synoptic overview of the life cycle of Invest 91C presented in section 4. Vertical cross sections, taken through the center of Invest 91C at 0000 UTC 29 October (red dot in Fig. 2) and 0000 UTC 2 November (dark blue dot in Fig. 2), are also presented in section 4 to illustrate the cyclone’s transition from an asymmetric, cold-core, EC into an axisymmetric, warm-core, TC.

As mentioned in section 1, deep convection and diabatic heating are required for the vertical redistribution of PV to occur during the TT process (Davis and Bosart 2003, 2004). Previous observational studies of TCs have identified regions of deep convection using satellite-derived infrared (IR) brightness temperature data (e.g., Romps and Kuang 2009; Hulme and Martin 2010a,b; Monette et al. 2012). Similarly, this study utilizes IR brightness temperature data, obtained from the NCEP/Climate Prediction Center 4-km global (60°N–60°S) IR dataset (Janowiak et al. 2001; available online at http://www.cpc.ncep.noaa.gov/products/
global_precip/html/wpage.full_res.html), in section 4 to identify regions of deep convection during the life cycle of Invest 91C.

Near-surface wind speed (m s−1), wind gusts (m s−1), and pressure (hPa) values associated with the landfall of Invest 91C along the extreme northwest (southwest) coast of Washington (British Columbia) are presented in section 4, with data obtained from the National Data Buoy Center (NDBC) online data archive (available online at http://www.ndbc.noaa.gov/maps/


nw_straits_sound_hist.shtml). The SST values displayed in Fig. 1b were obtained from the analyses of daily 0.25° Optimum Interpolation SST (OISST) data (Reynolds et al. 2007), collected using the Advanced Very High Resolution Radiometer (AVHRR-Only) (available online at http://www.ncdc.noaa.gov/oisst).


  1. Upper- and lower-tropospheric precursors

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. 3a). 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. 3a,b). A broad surface low (EC1) begins to develop in the equatorward entrance region of the upper-tropospheric jet at 0000 UTC 24 October, downstream of an upper-tropospheric disturbance located over the Sea of Japan (Fig. 3b). Differential cyclonic vorticity advection by the thermal wind, implied in Figs. 3a,b, helps to focus upward motion and diabatic heating over the western portion of EC1. This implied forcing for ascent and 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. 3a,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. 3c,d). Warm air advection (WAA) to the east of EC2 is associated with the amplification of the downstream ridge in the upper troposphere (Fig. 3d). 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 (e.g., Archambault et al. 2013) 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. 3b,d).

EC1 and EC2 deepen and move northeastward between 0000 UTC 25 October and 0000 UTC 26 October (Figs. 3c,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. 3e), 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 meridional flow downstream of the ridge axis between 0000 UTC 25 October and 0000 UTC 26 October (Figs. 3d,f). Enhanced meridional 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. 3f).

The development of the positively tilted upper-tropospheric trough over the central North Pacific coincides with the formation of a narrow corridor of relatively low MSLP values over the eastern North Pacific at 0000 UTC 26 October, in the equatorward entrance region of the North Pacific jet (Fig. 3e). This narrow corridor of relatively low MSLP values amalgamates into a ~1012-hPa closed low (L) over the following 24 h as the positively tilted upper-tropospheric trough amplifies upstream (Figs. 3e,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. 3f,h). Continued stretching and thinning occurs over the following 24 h, coinciding with the equatorward movement of L to 40.5°N, 147.5°W by 0000 UTC 28 October (Figs. 3g–j).



The amplification of the North Pacific waveguide that results in the formation of L, 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 4 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. 3a,b) (T1) is evident over eastern Asia at 0000 UTC 24 October, with the trough axis denoted by 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. 3a,b) help to enhance the downstream ridge (R) over the following 24 h, with the ridge axis denoted by 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 (T2), with the trough axis again denoted by 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 T2 (Figs. 3f,h,j) cause T2 to slow and thin between 0000 UTC 26 October and 0000 UTC 28 October. Figure 4 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 (L) located at 40.5°N, 147.5°W, is an anomalous feature in the eastern North Pacific basin in late October.


  1. Tropical transition of Invest 91C

An elongated region of low MSLP values, centered 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) values (Fig. 5a). The strip of 925–850-hPa layer-averaged relative vorticity associated with this region of negative ζT is located downstream of the thinning upper-tropospheric trough identified in section 3 (Fig. 5b). A meridionally elongated corridor of >40 mm precipitable water (PW) values extends poleward downstream of this thinning upper-tropospheric trough, collocated 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 in Fig. 5d aligns nicely with the region of 600–400-hPa layer-averaged ascent in Fig. 5c.

Differential cyclonic vorticity advection occurring over the elongated region of low MSLP values (not shown) is associated with the formation of an 999-hPa EC over the following 24 h (Figs. 5a and 6a). A region of WAA (CAA) begins to wrap cyclonically around the eastern (western) 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 999-hPa EC (Fig. 7). The narrow corridor of warm, moist, air to the east of the EC center in Figs. 6a,c is indicated by >312 K equivalent potential temperature (θe) values and by relatively high potential temperature (θ) values in the lower and midtroposphere immediately to the east of the MSLP minimum. A region of cold, dry, air to the west of the EC center in Figs. 6a,c is indicated by <306 K θe values and by relatively low θ values in the lower and midtroposphere to the west of MSLP minimum. A region of PV values >2 PVU associated with the upper-tropospheric cutoff shown in Fig. 6b extends from the stratosphere to ~600 hPa to the west of the EC center. A region of >2 PVU values also exists directly over the EC center in the lower troposphere, collocated with a lower-tropospheric frontal boundary.

The corridor of warm, moist, air wraps around the northern and western edges 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, defined in Bosart and Lackmann (1995) as the difference between θ on the dynamic tropopause (DT) and θe at 850 hPa, are <−2.5 K over the EC center at this time (not shown), suggesting the 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 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 and begin to lose its frontal structure by 0000 UTC 31 October. Negative ζT values associated with the cyclone center have become removed from the rest of the bent-back warm front (Fig. 9a) and remain collocated with the upper-tropospheric cutoff at this time (Fig. 9b). An slight increase in DT θ values over the cyclone center (Fig. 9b), corresponding to a 1 PVU reduction in 300–200-hPa layer-averaged PV values over the cyclone center (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 observable in the IR satellite imagery (Fig. 9d).

The central pressure of the transitioning cyclone has reached 991 hPa by 0000 UTC 1 November (Figs. 2 and 10a). ζT values have decreased to <−16 × 10−5 s−1 near the cyclone center at this time, with a reduction in ζT values observed throughout the majority of the analyzed domain (Figs. 9a and 10a). The 925–850-hPa layer-averaged relative vorticity maximum has moved northward in response to a progressive upper-tropospheric trough that wrapped cyclonically around the southern portion of the transitioning cyclone during the previous 24 h (Figs. 8b and 9b). Midtropospheric ascent is observed over and surrounding the center of the transitioning cyclone at this time (Fig. 10c). A small region of deep convection (indicated by cold cloud tops in the IR satellite imagery in Fig. 10d) is collocated with the cyclone center at time, corresponding to the analyzed regions of midtropospheric ascent shown in Fig. 10c.

The TC-like characteristics of the transitioning cyclone at 0000 UTC 1 November (Figs. 10a–d) caused CPHC to label the storm “Invest 91C” at 1200 UTC 1 November. The central pressure of Invest 91C decreases from 991 hPa to 989 hPa between 0000 UTC 1 November and 0000 UTC 2 November (Figs. 10a and 11a), while 925–850-hPa layer-averaged relative vorticity values >2.0 × 10−4 s−1 are maintained near the cyclone center (Figs. 10b and 11b). Midtropospheric ascent persists near the near cyclone center at 0000 UTC 2 November (Fig. 11c), collocated with >40 mm PW values and a region of deep convection surrounding the cyclone center (Figs. 11c,d).

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, indicated by a bowing down of θ contours toward the surface throughout the depth of the troposphere, confirm the warm-core structure of the cyclone suggested in Fig. 11a. Vertically orientated θe contours surrounding the MSLP minimum suggest that deep convection has been occurring near the center of the cyclone. The upper-tropospheric PV maximum that extended from the stratosphere to ~600 hPa at 0000 UTC 29 October (Fig. 7) no longer exists at 0000 UTC 2 November (Fig. 12), likely due to the diabatic redistribution of PV in the vertical associated with persistent deep convection. 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.

Invest 91C begins to approach the west coast of North America between 0000 UTC 2 November and 0000 UTC 3 November, traveling 10° of longitude in 24 h (Fig. 2). The minimum MSLP value associated with the center of the cyclone has increased from 989 hPa to 993 hPa during this period as the MSLP field becomes broader and less organized (Figs. 11a and 13a). ζT values, however, remain robust (<−14 × 10−5 s−1) immediately to the south of the minimum in MSLP (Fig. 13a), corresponding to the position of the 925–850-hPa layer-averaged relative vorticity maximum (Fig. 13b). Figure 13c indicates that midtropospheric ascent is still occurring over Invest 91C at 0000 UTC 3 November, collocated with a region of >35 mm PW values. This midtropospheric ascent, however, is associated with minimal convective activity in the IR brightness temperature field (Fig. 13d). Vertical wind shear, inferred from DT wind speeds >60 kts (Fig. 13b), forces convective activity downshear of Invest 91C (to the northeast of 925–850-hPa layer-averaged relative vorticity maximum) (Figs. 13b,d) and likely prevents the cyclone from intensifying further.

CPHC issued the final advisory for Invest 91C at 1200 UTC 3 November, only a few hours prior to the landfall of the weakening disturbance. Figure 2 reveals that the remnants of Invest 91C traveled rapidly to the northeast between 0000 UTC 3 November and 1800 UTC 3 November, making landfall along the extreme northwest (southwest) coast of Washington (British Columbia) between 1600 UTC and 1800 UTC 3 November. A region of >486 dam 925–500-hPa thickness values is located over extreme northwest (southwest) coast of Washington (British Columbia) at 1800 UTC 3 November (Fig. 14a), collocated with a region of 925–850-hPa layer-averaged relative vorticity values in excess of 1.5 × 104 s1 (Fig. 14b). The strip of relatively high 925–850-hPa layer-averaged relative vorticity values along the Puget Sound is likely associated with surrounding orography, occurring between the Olympic Mountains and the Cascade Range at 0000 UTC 3 November and 1800 UTC 3 November (Figs. 13b and 14b). Midtropospheric ascent and convection associated with the remnants of Invest 91C are difficult to distinguish from surrounding features at 1800 UTC 3 November (Figs. 14c,d), making further synoptic-scale analysis difficult.

The sensible weather associated with the landfall of the remnants of Invest 91C can be analyzed using observations obtained from the NDBC online data archive (see section 2). Figure 15 reveals that pressure values at Destruction Island, WA, decrease steadily from 1024 hPa to 1000 hPa in the days leading up to the landfall of the remnants of Invest 91C (i.e., between 2000 UTC 31 October and 2200 UTC 2 November). Wind speed values remain relatively weak during the first portion of this period, but increase from 2.8 m s−1 to 14.0 m s−1 between 0200 UTC 2 November and 2200 UTC 2 November. Pressure (wind speed) values are shown to increase (decrease) gradually between 2200 UTC 2 November and 0800 UTC 3 November. However, at 0900 UTC 3 November, pressure (wind speed) values begin to fall (rise) rapidly at the Destruction Island, WA, lighthouse in association with the approaching remnants of Invest 91C. Wind speed and wind gust values of 26.2 m s−1 and 29.2 m s−1, respectively, are observed at 1600 UTC 3 November, corresponding to a pressure measurement of 998.4 hPa. Pressure (wind speed) values rise (fall) rapidly after the passage of the remnants of Invest 91C, indicating the relatively small areal extent of the disturbance.

The data recorded at the Destruction Island, WA, lighthouse is representative of wind speed, wind gust, and pressure measurements recorded at surrounding observation stations included in the NDBC online data archive. Table 1 indicates that the southern observation stations (Destruction Island, WA, and Cape Elizabeth, WA) recorded minimum pressure values at 1600 UTC 3 November, 2 h earlier than the northern observation stations (Neah Bay, WA, and Tatoosh Island, WA). The northern observation stations, however, recorded lower minimum pressure values than the southern observation stations, likely due to their alignment with the remnants of Invest 91C (Fig. 2). Finally, the observation stations closest to the coast of Washington (Destruction Island, WA, and Tatoosh Island, WA) recorded the fastest wind speeds and wind gusts of the observation stations included in Table 1, likely due to terrain channeling.




  1. Summary and conclusions

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 region where tropical cyclogenesis events have not been documented, 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. The TC analyzed in this study was able to form over a region of ~16°C SSTs, much lower than the typical SST value required for TC development (e.g., Palmén 1948; Gray 1968). Invest 91C formed in the presence of a cold upper-tropospheric cutoff that decreased the deep-layer atmospheric stability [e.g. coupling index values below 0 K (Bosart and Lackmann 1995)] near the cyclone center, facilitating the development of the deep convection necessary for the TT process to occur (e.g, McTaggert-Cowan et al. 2006). 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. The remnants of Invest 91C produced wind gusts in excess of 29.2 m s−1 in extreme northwestern Washington when they made landfall at approximately 1600 UTC 3 November. Observation stations along the western coast of Washington recorded pressure values of 995–999 hPa at the time of landfall, several hPa below the minimum MSLP value associated with the weakening cyclone in the 0.5° NCEP CFSR dataset.

It is important to note that Invest 91C, though shown to be 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 considered for inclusion in these datasets in order to obtain a comprehensive understanding of the frequency and location of tropical cyclogenesis events around the globe.


Acknowledgments. The authors would like to thank Drs. Ryan Torn (University at Albany), Lance Bosart (University at Albany), and Kristen Corbosiero (University at Albany) for helpful discussions and research assistance.  This study was initially motivated by the weekly Friday map discussion of this event at the University at Albany. This research was funded by the Hobart and William Smith Colleges Provost’s Office and NSF Grant AGS-1355960.  


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TABLES
TABLE 1. Measurements recorded by observation stations during the landfall of the remnants of Invest 91C on 3 November 2006. Measurements included in Table 1 correspond to the time when the lowest pressure value was recorded at each observation station. Observation stations are listed from north to south.


Station Name
[Station Number]

Location

Time

Pressure (hPa)

Wind Speed (m s−1)

Wind Gust (m s−1)

Neah Bay, WA

[46087]


48.5°N, 124.7°W

1800 UTC

995.5

15.9

21.1

Tatoosh Island, WA

[TTIW1]


48.4°N, 124.7°W

1800 UTC

994.8

20.0

24.6

Destruction Island, WA

[DESW1]


47.7°N, 124.5°W

1600 UTC

998.4

26.2

29.2

Cape Elizabeth, WA

[46041]


47.3°N, 124.7°W

1600 UTC

999.1

14.4

18.1


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° NCEP CFSR dataset, valid at 0000 UTC 2 November 2006, overlaid on infrared satellite sea surface temperatures (SST; shaded, °C), plotted using 0.25° Optimum Interpolation SST data, valid on 2 November 2006.


FIG. 2. Track map indicating the position of the center of Invest 91C every 6 h between 0000 UTC 28 October 2006 and 1800 UTC 3 November 2006, identified using the minimum mean sea level pressure value associated with the cyclone in the 0.5° NCEP CFSR dataset. The positions plotted in color and the minimum sea-level pressure values correspond to 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 s1) 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 s1), 250-hPa wind speed (shaded, m s1), and 600–400-hPa layer-averaged ascent (red contours every 2.5 × 10−3 hPa s1), 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 position of surface cyclone centers. The “L” labels in (g) and (i) denote the position of a closed low.


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 “T1” and “T2” labels denote the position of 250-hPa trough axes. The “R” label denotes the position of a 250-hPa ridge axis. The “L” label denotes the location of a closed low.
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, 105 s1), (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 × 104 s1), (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 s1), 250-hPa wind speed (shaded, m s1), and 600–400-hPa layer-averaged ascent (red contours every 2 × 10−3 hPa s1), 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. 10. As in Fig. 5, but at 0000 UTC 1 November 2006.
FIG. 11. As in Fig. 5, but at 0000 UTC 2 November 2006. Line B–B’ in (d) denotes the position of a vertical cross section shown in Fig. 12.
FIG. 12. As in Fig. 7, except along line B–B’ in Fig. 10d at 0000 UTC 2 November 2006.
FIG. 13. As in Fig. 5, but at 0000 UTC 3 November 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.

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


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