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The Extratropical Transition of Tropical Cyclone Edisoana (1990) [Monthly Weather Review](Monthly Weather Review Via Acquire Media NewsEdge) ABSTRACT Documentation of southwest Indian Ocean (SWIO) tropical cyclones (TCs) and extratropical transition (ET) events is sparse in the refereed literature. The authors present a climatology of SWIO TC and ET events for 1989-2013. The SWIO averages ~9 tropical cyclones (TCs) per year in this modern era. Of these TCs, ~44% undergo extratropical transition (ET), or ~four per year. A case study of TC Edisoana (1990), the most rapidly intensifying SWIO post-ET TC between 1989 and 2013, shows that extratropical interactions began when an approaching trough embedded in the subtropical jet stream (STJ) induced ET on 7 March. As Edisoana underwent ET, a subtropical ridge downstream amplified in response to poleward-directed positive potential vorticity (PV) advection associated with diabatically (convectively) driven upper-level outflow from TC Edisoana. This amplifying lower-latitude ridge phased with a lower-amplitude higher-latitude ridge embedded in the polar front jet (PFJ), resulting in the merger of the two jets. This ridge phasing and jet merger, combined with the approach of an upstream trough embedded in the PFJ, resulted in a decrease in the half-wavelength between the approaching trough and the downstream phased ridges and provided extratropical cyclone Edisoana with a prime environment for rapid reintensification (RI). Poleward-directed positive PV advection into the phased ridge strengthened the upper-level jet downstream of Edisoana, which provided the primary baroclinic forcing throughout the RI phase. A backward trajectory analysis suggests that strong diabatic heating enhanced favorable synoptic-scale forcing for ascent from the upstream and downstream jet streaks and played a crucial role in the deepening of Edisoana through the ET and RI periods. 1. Introduction A subset of tropical cyclones (TCs) interact with the midlatitudes at some point in their life cycle. When this interaction leads to a change in the structure of the TC into a structure more comparable to that of an extra- tropical cyclone and, on occasion, a reintensification of the cyclone, the process is termed extratropical transi- tion (ET). The extratropical transition of tropical cy- clones worldwide has been well documented through climatologies (e.g., Sinclair 1993, 2002, 2004; Foley and Hanstrum 1994; Harr and Elsberry 2000; Klein et al. 2000, 2002; Hart and Evans 2001; Jones et al. 2003; Dare and Davidson 2004), numerous case studies (e.g., Sekioka 1956a,b; DiMego and Bosart 1982a,b; Thorncroft and Jones 2000; McTaggart-Cowan et al. 2001, 2003; Agust^i- Panareda et al. 2004, 2005; Evans and Prater-Mayes 2004; Kitabatake 2008; Harr and Dea 2009; Chen and Pan 2010; Cordeira and Bosart 2011; Grams et al. 2011, 2013; Pantillon et al. 2013), and idealized and ensemble modeling studies (e.g., Ritchie and Elsberry 2007; Riemer et al. 2008; Riemer and Jones 2010; Anwender et al. 2010; Davis 2010; Torn 2010; Pantillon et al. 2013). Jones et al. (2003) provided a comprehensive summary of the state of global ET knowledge and research. Absent from these publications is any comprehensive assessment of ET cli- matologies in the southwest Indian Ocean (SWIO) in general and case studies of TCs that undergo ET in this ocean basin in particular. A prior climatology of TC activity in the SWIO from 1960 to 1989 presented by Valadon (1992) focused on the relationship between TCs and El Nin~o-Southern Oscillation. More recently, Mavume et al. (2009) pre- sented a climatology of TC activity in the SWIO between November and April (1980-2007) that emphasized landfalling TCs as well as the relationship between sea surface temperatures and TC genesis. The purpose of this paper is to present the results of a compilation of TC and ET events in the SWIO and to describe variations in TC and ET activity on monthly and annual time scales. This climatology will serve to establish a set of long-term av- erages of TC and ET activity, as already established in other basins, and further provide a simple basis for comparing variations in TC and ET activity in the SWIO to other TC basins. In compiling the ET portion of the climatology, a number of deep (minimum central pressure , 950 hPa) post-ET extratropical cyclones (ECs) were noted, the strongest of which was post-TC1 Edisoana in early March 1990 (http://www.nws.noaa.gov/om/notification/ scn10-06trop_cyclone_terms.txt). As a post-TC, Edisoana deepened explosively (.40 hPa), reached a minimum central pressure of 938 hPa, and motivated us to con- duct a case study of a storm that, arguably, may be one of the first well-documented ET events in the SWIO. TCs undergoing ET in the SWIO typically do not have a direct impact on any highly populated regions since only isolated islands lie between 258S and Antarctica in the southern Indian Ocean. However, a subset of SWIO TCs that undergo ET experience explosive re- intensification as post-TCs (e.g., Edisoana) and can create downstream impacts near southern Australia and New Zealand via downstream development (e.g., Harr and Dea 2009; Orlanski and Chang 1993; Orlanski and Sheldon 1993; Simmons and Hoskins 1979). As we will show, strong TC-jet interactions played a signifi- cant role in the ET of TC Edisoana and its subsequent explosive reintensification as an EC, analogous to similar post-ET explosively deepening cyclones over the North Atlantic and western North Pacific Oceans (e.g., Thorncroft and Jones 2000; McTaggart-Cowan et al. 2003, 2004; Agust^i-Panareda et al. 2004; Evans and Prater-Mayes 2004; Harr and Dea 2009; Cordeira and Bosart 2011). A more recent example of a post-TC that deepened explosively over the SWIO was TC Jade in April 2009. Although the ET of TC Jade is only briefly discussed here, it helped to motivate our investigation of the TC and ET events in the SWIO that are examined in this paper. The paper is organized as follows. Section 2 presents a discussion of the data quality of both TC records and reanalyses in the SWIO as well as the general procedure used to develop the climatology. Section 3 discusses the climatology of TC and ET activity on both annual and intraseasonal time scales. Section 4 presents a detailed examination of the ET and subsequent rapid reintensi- fication [RI;definedbyHart(2003)to be24hPaday21]of TC Edisoana, focusing on the evolution of synoptic-scale features during the extratropical RI period. Section 5 summarizes and synthesizes the results of sections 3 and 4 and compares the evolution of TC Edisoana to other documented ET cases from other ocean basins. 2. Data and methodology a. Climatology data sources and methodology The TC portion of the SWIO climatology utilized 6-hourly track and intensity data points from the best track files of Me^t^eo-France La Reunion, France [the official Regional Specialized Meteorological Center (RSMC) for the southern Indian Ocean west of 908E], via the International Best Track Archive for Climate Stew- ardship (IBTrACS; Knapp et al. 2010) and software pro- vided by M^et^eo-France La Reunion. In compiling the ET climatology, inconsistencies were noted in the best track files where not all seasons contain the extratropical portion of storm tracks. A subset of seasons (many in the early 1990s) contain storms that recurve into mid- latitudes (suggesting the occurrence of ET), but include no official ET designation or track for an extratropical phase of the storm, while other more recent seasons do contain extratropical tracks. Because of these in- consistencies in records, the authors decided that a sub- jective analysis of ET events would allow for a more consistent and complete analysis to be used in the cli- matology while also providing the opportunity to be- come familiar with the synoptic patterns and typical TC activity within the basin. This subjective analysis utilized the European Centre for Medium-Range Weather Forecasts (ECMWF) Interim Re-Analysis (ERA-Interim) gridded dataset (1.58 horizontal resolu- tion; see Uppala et al. 2005; Dee et al. 2011) as this dataset appeared to be the best compromise between datasets with a finer spatial resolution covering a shorter time range and datasets that include older seasons, but with a coarser spatial resolution. The majority of the SWIO TC and ET climatology was performed prior to the availability of the higher-resolution (0.58 horizontal res- olution, ;25 hPa vertical resolution) National Centers for Environmental Prediction (NCEP) Climate Forecast System Reanalysis (CFSR; Saha et al. 2010)gridded datasets that extend from 1979 to present. A limited com- parison between the ERA-Interim and CFSR analyses show good agreement in the representation of synoptic- scale features examined in the climatology and suggest that similar results could have been obtained from either set of gridded analyses. ET events between 1989 and 2013 were subjectively determined based on the application of diagnostic analyses such as looking for increases in the 1000- 500-hPa thickness asymmetry and the development of pressure troughs in mean sea level pressure (MSLP) fields. Such an analysis approach utilized patterns identified in the climatology of TCs captured by frontal zones off the western coast of Australia (Foley and Hanstrum 1994) as well as variables similar to those used in the objective cyclone phase space methodology in- troduced by Evans and Hart (2003). To apply this analysis technique, a broad selection of prospective ET cases was chosen to include any system whose track reached into the subtropics or showed signs of in- teraction with a midlatitude trough. While many of these prospective cases of ET lead to clear-cut decisions as to whether ET occurred, a spectrum of transition types emerged during this subjective analysis and relatively few of the prospective ET cases remained ambiguous after applying this analysis. For each ET event, the date and time of the ET were noted along with whether the post-TC weakened or strengthened, and if the latter, the net decrease in MSLP during the reintensification pe- riod. Over 150 possible ET cases were examined, with 93 ET events identified. A vast majority (;90%) of these ET events agree well with (timed to be within 12 h of) the first extratropical point in the best track file or where a first extratropical point could be inferred in cases where the extratropical portion of the track is not available. Many of the discrepancies were present in seasons prior to the mid-1990s. A comparison of the subjectively derived and best track ET times suggests the subjective ET analysis is reasonably accurate with its timing of ET and leads to a reasonably accurate representation of ET events throughout the climatol- ogy period. Gaps in SWIO best track data serve to limit the his- torical extent of the TC and ET climatology. Incomplete track data appear as recently as the 1985-86 season. Prior to 1980-81, the best track data contains an in- tensity estimate only once daily with some seemingly incomplete TC tracks as well. Given these best track data quality issues and the questionable quality of gridded analyses in the SWIO region prior to the early 1990s, large uncertainties are present in the classification of cyclones as TCs, undergoing ET, or post-TCs. This period of apparent questionable quality of gridded analyses corresponds roughly with the time prior to the establishment of geostationary environmental satellite over the Indian Ocean. Because confidence in our ability to analyze the structure and evolution of TCs and whether the TCs had experienced ET was lessened prior to the late 1980s, we decided to start our climatological analysis with the 1988-89 TC season. Decreased confi- dence in the analysis of earlier TC seasons stems largely from the weak or nonexistent representation of TCs in the gridded reanalyses that were noted while initially performing this study and further described by Schenkel and Hart (2012). This choice also serves to continue the previous TC climatology as reported by Valadon (1992) for the period 1960-89 and update and expand the Mavume et al. (2009) climatology for the 1980-2007 period. Although the databases used by both Valadon (1992) and Mavume et al. (2009) to construct these previous TC climatologies likely contain TC omissions and incomplete TC tracks before the advent of the modern satellite era and (for Valadon) modern global gridded datasets, we will still compare our climatology with those of Valadon (1992) and Mavume et al. (2009). b. Case study data sources and methodology In first preparing the case study of TC Edisoana, a comparison was performed between the CFSR dataset with its increased spatial resolution and the ERA- Interim dataset. Synoptic-scale features associated with the Edisoana ET were comparable in both datasets as was the case for most of the subsynoptic-scale features. Fur- thermore, both analyses agreed reasonably well with a set of surface observations from the Kerguelen Islands (around 49.58S, 70.08E) that were extracted from opera- tional National Oceanic and Atmospheric Administra- tion (NOAA) synoptic observation charts at 0000, 0600, and 1200 UTC 9 March 1990 (not shown). EC Edisoana's closest point of approach to the Kerguelen Islands oc- curred just prior to 0000 UTC 9 March 1990 as the EC was explosively intensifying. The Kerguelen Island station reported an MSLP observation of 958.2 hPa at 0000 UTC 9 March, which is ;2 hPa deeper than the gridded analyses (both CFSR and ERA-Interim) suggest for the station at that time, but appears to be within 1- 2 hPa of the cyclone's true central MSLP at that time due to the proximity of the analyzed center to the island. By 0600 UTC 9 March, Kerguelen Island reported its minimum MSLP reading of 949.1 hPa, at which time the cyclone's analyzed center was located ;150 km southeast of the island. This MSLP value from Kerguelen Island compares favorably to MSLP values of ;954 hPa that were interpolated from the gridded analyses. The reported pressure tendency from Kerguelen Is- land at 0600 UTC 9 March indicated the pressure at the station had already started to rise by the time of the observation, suggesting the minimum MSLP value at the storm's closest approach was likely lower than ob- served from the 6-hourly observations. Furthermore, the entire set of the Kerguelen Island observations show lower MSLP values than do the gridded analyses interpolated to the position of Kerguelen Island, suggesting either the depth of the cyclone or the strength of its pressure gradi- ents is likely underestimated. Nevertheless, we argue that the CFSR gridded datasets provided a good quality overall representative analysis for the ET of TC Edisoana despite potential problems associated with analyses pre- dating the geostationary satellite era over the southwest Indian Ocean [from the National Climatic Data Center (NCDC; http://www.ncdc.noaa.gov/oa/rsad/gibbs/satellite- info.html, retrieved 2013]. With this argument in mind, the CFSR gridded dataset is utilized for the synoptic analysis portion of the case study in section 4. 3. Climatology The SWIO experiences TCs year-round, which pres- ents difficulties in establishing yearly TC statistics. Since the turn of the calendar year lies in the heart of the TC season, TC statistics are kept with a new season begin- ning either 1 July (as practiced by the U.S. Navy's Joint Typhoon Warning Center) or 1 August (as practiced by RSMC La Reunion). Because the entire climatology (1989-2013) only contains four TC events during July and August, the annual distributions are not significantly affected by a choice of a 1 July or a 1 August season starting date. An annual cycle of 1 July-30 June is used for both TC and ET events in this study for the conve- nience of having a 6-month offset with Northern Hemispheric (NH) records. References to a season with a single year hereafter refer to the season ending in that year (e.g., ''2009'' refers to 1 July 2008-30 June 2009). a. Annual variations in TC and ET activity TC frequency in the SWIO has remained steady since 1989, averaging 9.4 TCs of tropical storm strength (10-min sustained winds .34 kt, 1 kt 5 0.5144m s21)or higher per TC season (total number of TCs is 235). Of these TCs, 4 undergo ET every year, leading to an av- erage ET rate of 43.8% (Fig. 1a). This SWIO rate of ET is comparable to the rate of ET in the North Atlantic basin (46%) as found by Hart and Evans (2001) and is noticeably higher than the observed ET rates of 27% in the northwest Pacific from Klein et al. (2000) and 32.2% in the South Pacific from Sinclair (2002). The 9.4 count of annual TC events roughly agrees with the ;10 TC events per year derived from the previous SWIO TC climatology computed by Valadon (1992) over the 1960- 89 TC seasons. b. Monthly variations in TC and ET activity TCs occur in all months of the year in the SWIO as shown in the 25-yr monthly TC climatology presented in Fig. 1b. Although TC season in the SWIO is not as clearly defined as in the North Atlantic and east Pacific, 86% of TCs occur between November and April. Valadon (1992) also defines 1 November-30 April as a TC season based on activity records during the 1960- 89 TC seasons. Using these months to define the SWIO TC season is consistent with Mavume et al. (2009) who constructed a climatology of SWIO TCs between No- vember and April and found these months to contain 85% of TC activity. Because each of these 6 months averages at least 1.0 TCs per month, a TC season can be established utilizing November and April as the season boundaries. The peak of the SWIO TC season occurs in January and into February, with an average of 2.3 and 2.4 TC events per month, respectively (Fig. 1b). March, with 1.7 TC events, is the third most active month. No- vember, December, and April make up the rest of the SWIO TC season, all averaging ;1.0 TC events per season. The remaining months average no more than 0.3 TC events per season, providing a clear delineation be- tween periods of higher and lower TC activity. ET events in the SWIO occur almost exclusively be- tween November and April (Fig. 1b), with a higher frequency later in the TC season. The ET rates in the early portion of the TC season are relatively low, with November and December statistics yielding ET rates of 4% and 32%, respectively. As TC activity begins to peak, the ET rate also spikes to 57% in January before falling off slightly later in the season, with rates of 48%, 48%, and 42% for February, March, and April, re- spectively (Fig. 1b). In comparison, these late-season ET rates in the SWIO are similar to the rates calculated by Hart and Evans (2001) for the peak ET months of September and October in the North Atlantic and share the late-season bias found by Dare and Davidson (2004) in the Australian region. 4. TC Edisoana (28 February-10 March 1990) a. Overview The post-TC reintensification of TC Edisoana as an EC (;40-hPa SLP decrease) was the strongest such event during 1989-2013. TC Edisoana first attained tropical storm intensity at 0000 UTC 2 March 1990 while located 900 km northeast of Mauritius (Fig. 2a). TC Edisoana intensified slowly before attaining TC in- tensity (a sustained 10-min wind greater than 64 kt) on 5 March (Fig. 2b). Per RSMC La Reunion records, TC Edisoana maintained a 72-kt intensity (10-min wind) with a minimum central pressure of 953 hPa for the next 48 h, while Joint Typhoon Warning Center (JTWC) re- cords indicate that TC Edisoana reached TC strength a full 24 h earlier before peaking in intensity with 100 kt winds (1-min; 10-min equivalent is ;85 kt) at 0600 UTC 6 March. Furthermore, RSMC La Reunion declared that TC Edisoana had become extratropical at 0600 UTC 7 March, a full 24 h before the JTWC ended its tracking of the TC. Discrepancies in TC wind speeds and TC struc- ture (e.g., whether it has transitioned to an EC after completing ET) between RSMC La Reunion and the JTWC are not uncommon in the historical TC record of the late 1980s and early 1990s. Given these stated dis- crepancies and given the location of TC Edisoana in the SWIO, we opted to use the best track data from RSMC La Reunion for the purposes of this case study. However, best track data truncates 24 h after TC Edisoana is de- clared extratropical, resulting in a need to use an esti- mated track as obtained from following the minimum sea level pressure center within the ERA-Interim and CFSR gridded datasets for the extratropical portion of the track of Edisoana. Use of these reanalyses allow EC Edisoana to be tracked through its deepest minimum central pres- sure of 938 hPa at 1200 UTC 9 March before the storm began to decay near the Antarctic coast around 1108Eat 1200 UTC 10 March. To present an objective analysis of the ET of TC Edisoana, we first show a Hart (2003) cyclone phase space diagram for this storm (Fig. 3). A distinct increase in the 900-600-hPa thickness asymmetry beginning on 7 March is indicative of a TC that is taking on frontal characteristics (Fig. 3). It is consistent with the timing of ET as designated by the best track from RSMC La Reunion. Between 0600 UTC 7 and 8 March, the warm core of TC Edisona gradually weakens and becomes cold core. This timing of the cyclone's transition from warm to cold core is also coincident with the end of the JTWC track (and presumably a declaration that the TC is extratropical) at 0600 UTC 8 March. Regardless of the exact timing of ET onset, gridded reanalyses indicate that post-TC Edisoana starts to expe- rience a period of rapid deepening beginning 1200 UTC 8March (Fig. 2b). In the 36 h leading up to this time, the cyclone tracks through the cyclone phase space via a curving arc from symmetric with a warm core to asym- metric with a cold core, a track similar to other ET events that rapidly reintensify as extratropical cyclones (Hart 2003; Hart et al. 2006). After briefly reaching the asym- metric cold-core section of the cyclone phase space at 1200 UTC 8 March, the core of post-TC Edisoana begins warming and the low-level thickness asymmetry parame- ter begins to decrease, indicating that the cyclone was developing a warm seclusion (Fig. 3) during its period of rapid deepening through 1200 UTC 9 March (Fig. 2b). This evolution will be discussed further in the section 4d. b. Precursor environment and ET In examining the source of the synoptic-scale features that helped facilitate the ET of TC Edisoana, the im- portance of a high-amplitude Rossby wave train (RWT) becomes apparent. Initially noted on Hovmoeuroller dia- grams as early as 25 February (Fig. 4), a RWT is asso- ciated with trough amplification over the southeast Pacific Ocean (near 1008W) on 26 February (Fig. 5a). At the same time, a pool of warm potential tempera- ture air on the dynamic tropopause [DT, defined as 22.0-potential vorticity unit (PVU) surface (1 PVU 5 1026 Kkg21 m2 s21); labeled ''SWP'' in Figs. 5 and 6] bulges poleward along 808W, just off the west coast of South America. While the leading edge of the RWT moves rapidly downstream (;30 m s21 as indicated in Fig. 4), the SWP slowly progresses eastward and pole- ward over southern South America while accompanied by a progressive and sharpening trough at 0000 UTC 28 February (Fig. 5b). Simultaneously, an additional warm pool (labeled PWP in Fig. 5) in the polar front jet stream (PFJ) flow triggers a larger anticyclonic wave break (AWB; Thorncroft et al. 1993) event poleward and downstream of the SWP. While there does not appear to be a high-amplitude RWT associated with the PWP, the PWP does appear to amplify the polar trough located along 208W at 0000 UTC 28 February (labeled PV1). While the relatively slow movement of SWP (;11 m s21) and the accompanying upstream trough continues through 0000 UTC 2 March (Fig. 5c), the warm pool and associated upper-level ridge expands eastward and amplifies poleward over the South At- lantic (Figs. 5c and 6a) into the remnants of the AWB event associated with PWP. Between 0000 UTC 4 and 6 March, this ridge amplification leads to an AWB event associated with the SWP poleward of South Africa (yellow star in Fig. 4; Figs. 6a,b). In the 48-h period ending 0000 UTC 6 March, the SWP AWB event breaks equatorward while south of South Africa and traps the equatorward portion of trough PV1. This portion of PV1 thins and fractures from the main part of PV1 (labeled as PV2 in Fig. 6) and creates a PV streamer before cutting off and being left behind in weaker flow in the STJ along the southern coast of South Africa (Figs. 6a,b). Trough PV1 continues eastward across the southern Indian Ocean in the 48 h ending at 0000 UTC 6 March (Figs. 6a,b). Between 0000 UTC 6 and 7 March the northern portion of trough PV1 thins and weakens yet again as it is trapped beneath the continued AWB event associated with the eastward-progressing SWP, resulting in the creation of another PV streamer. This second streamer also becomes cut off from the PFJ flow and trough PV1 and is labeled as trough PV3 in Figs. 6b,c. While the PFJ trough (PV1) continues eastward across the southern Indian Ocean (Fig. 6b), trough PV3 be- comes embedded in the subtropical jet stream (STJ) and will be the first trough to interact with and stimulate the ET of TC Edisoana. Trough PV3, located in the STJ, helps to enhance the irrotational outflow associated with TC Edisoana (Figs. 7a,b), a sign of increasing interaction between the TC and the approaching subtropical trough (e.g., Archambault et al. 2013). Irrotational outflow increases to a localized maximum of 15 m s21 between TC Edisoana and the STJ at 0600 UTC 7 March (Fig. 7c),thetimeat which TC Edisoana is first declared extratropical. As the poleward portion of the outflow from TC Edisoana is enhanced, this also leads to strong poleward-directed positive PV advection by the irrotational wind toward PV3 (Fig. 7d). The resulting dipole of PV advection serves to strengthen the PV gradient on the eastern edge of the trough PV3, which in turn contributes to the local intensification of the STJ. Simultaneously, the component of the irrotational outflow directed per- pendicular to the axis of the STJ serves to strengthen the STJ as well, via the leftward-directed Coriolis torque in the Southern Hemisphere (SH). Intensification of the STJ by outflow from tropical convection has been observed in a number of individual and climatological studies (e.g., Bosart and Lackmann 1995; Agust^i-Panareda et al. 2004; Archambault et al. 2013) and is frequently di- agnosed via a tightening of the PV gradient by advection and an increase in the STJ wind speed. While the role of PV advection and irrotational outflow in accelerating the STJ is not quantized here, the overall increase in the magnitude of the irrotational wind with respect to the magnitude of the total wind is easily observed during the 18 h prior to TC Edisoana being declared extra- tropical (Fig. 8). The increasing importance of the irrotational wind can be noted in Fig. 8a with roughly 10% of the mag- nitude of the winds in the STJ streak (;50 m s21) im- mediately poleward of TC Edisoana contributed by the irrotational wind. By the time of ET (at 0600 UTC 7 Mar; Fig. 8b), the ratio of irrotational wind to total wind has increased to ;40% of the total wind magnitude poleward of the cyclone, representative of the ;10 m s21 increase in the magnitude of the irrotational wind in the region. Simultaneously, the total wind speed has roughly remained constant during this 18-h period prior to ET, implying an increasingly significant role of the irrotational wind in the dynamics associated with the STJ. While the enhanced STJ quickly moves down- stream, the local forcing for ascent in the equatorward entrance region of the STJ allows the central pressure of postTC [now extratropical cyclone (EC)] Edisoana to be maintained and even deepen slightly prior to beginning its interaction with a high-amplitude long-wave trough in the polar jet stream on 8 March. c. Post-ET ridge phasing The important role of ridge amplification during and after the ET of TC Edisoana is demonstrated in Fig. 9. The remnant of the AWB event (labeled SWP for con- sistency with the previous section) creates a significant region of anticyclonic curvature poleward of the tran- sitioning TC Edisoana at 1200 UTC 7 March (Fig. 9a) and allows for poleward amplification of the Edisoana- induced subtropical ridge in conjunction with the pole- ward movement of EC Edisoana (Figs. 9b,c). Although deformed, the elongated remnants of the subtropical PV streamers labeled PV2 and PV3 can each be seen be- tween the amplifying subtropical ridge and the ap- proaching upstream long-wave trough (Fig. 9a). While still significant anomalies both on the DT (Figs. 9a-d) and in the upper-level PV field (Figs. 10a and 11a), it does not initially appear that PV2 and PV3 played a significant role in the posttropical evolution of EC Edisoana. That said, the resolution of these gridded datasets might be insufficient to allow for a finer meso- scale analysis that would be necessary to account for the potential contributions from troughs PV2 and PV3. The subtropical ridge downstream of EC Edisoana continues to amplify through 8 March (Figs. 9b,c), driven by persistent upper-level outflow from Edisoana (represented by irrotational wind; Fig. 10a). As this ridge amplification continues, the STJ along the crest of the ridge begins to phase with the PFJ. Support for the upper-level outflow driving this ridge amplification is apparent in a meridional cross section through the center of EC Edisoana at 1200 UTC 7 March (Fig. 10b). An upright region of strong ascent is located on the poleward edge of the PV tower associated with the former TC. Above this region of ascent, pronounced poleward and equatorward outflow is evident. The poleward outflow impinges upon a thin body of PV and erodes trough PV3 near 458S(Fig. 10a). At 0000 UTC 8 March, a similar cross section (at 63 8 E, Fig. 11b) shows two distinct upper-level jet cores poleward of the tilting PV tower associated with the EC. The STJ, maximized at 200 hPa, and the PFJ, maximized at 300 hPa, are both characterized by distinct ''vertical steps'' on the DT. In addition, this cross section highlights strong ascent along the developing low-level frontal zone (as portrayed by steeply sloped isentropes below ;800 hPa), which further supports a region of small values of positive upper-level PV and associated irrotational outflow both poleward and equatorward (Fig. 11b), with the poleward branch of outflow impinging on the STJ. Strong vertical motion persists as EC Edisoana progresses poleward in the equatorward entrance region of the STJ (Figs. 11a and 12a). Persistence of this strong vertical motion helps to maintain a region of low- magnitude PV values equatorward of the STJ and a re- sultant poleward outflow channel across both the STJ and PFJ (implied by the irrotational wind analysis in Fig. 12a). This outflow continues to amplify the subtropical ridge and, by 1200 UTC 8 March, results in a phasing of the ridges in the subtropical and polar jet streams. While horizontal maps suggest that the two jet streams have merged at this point (Fig. 12a), a cross section taken through the center of EC Edisoana (along 668E; Fig. 12b) shows that the vertical steps associated with each jet on the DT remain distinct and that the two jets re- main separate features with semidistinct zonal wind maxima. On isobaric surfaces, the layer-mean 300-200-hPa wind speed associated with the two jets increases from over 50 m s21 with each jet (Fig. 11a) to a single region of wind speeds greater than 60 m s21 (Fig. 12a), suggesting that the jet intensification processes observed during ET are likely still playing a role in enhancing the upper-level support for EC Edisoana. Despite the STJ and PFJ remaining distinct in the cross sections (Figs. 11b and 12b), the ridges in each stream are sufficiently in phase to seemingly act as a cohesive, singular ridge over subsequent short time scales. Upstream, an amplifying trough in the PFJ begins interacting with EC Edisoana as it approaches the cy- clone from the west (Figs. 9c,d). The progress of this polar trough is slowed by the downstream phased ridges, resulting in the collapse of the half-wavelength between the ridge and the trough axes (Figs. 13a-c). The collapse of the half-wavelength downstream of a trough is a pro- cess previously identified to be associated with enhanced favorability for the deepening of midlatitude cyclones (e.g., Bosart and Lackmann 1995). The enhanced fa- vorability can be attributed to forcing from dynamical processes such as geostrophic adjustment and cyclonic vorticity advection occurring on shorter spatial scales and thus providing stronger forcing for ascent and the resultant deepening of the surface cyclone. In the case of Edisoana, the most readily observable function of this decrease in half-wavelength of the flow is an increase in mid to upper-level frontogenesis. Figures 13d-f show an approximate doubling of the 400-hPa frontogenesis be- tween 0000 UTC 8 March and 1200 UTC 9 March. While this increase in frontogenesis is occurring, the location of the frontogenesis with respect to the ap- proaching trough axis shifts dramatically westward. At 0000 UTC 8 March, the frontogenesis is associated with the leading edge of the approaching long-wave polar trough (Fig. 13d). By 1800 UTC 8 March, this fronto- genesis is located near the trough axis (Fig. 13e) and continues to shift to the upstream edge of the trough by 1200 UTC 9 March (Fig. 13f). The role of this fronto- genesis is discussed further with an additional PV ad- vection perspective in the following subsection. d. Rapid reintensification as an EC As the half-wavelength downstream of the polar trough collapses around 1200 UTC 8 March, EC Edisoana be- gins deepening rapidly. Beginning at a central pressure of ;976 hPa, the cyclone deepens to a minimum central pressure of 936 hPa at 1200 UTC 9 March, just 24 h later (Fig. 2b). At the beginning of this deepening period, EC Edisoana is located in the favorable equatorward en- trance region of the downstream jet streak (Fig. 14a), but the interaction between the cyclone and the nearby up- stream and downstream jet streaks becomes a two-way interaction during the RI process. The advection of upper-level PV by the irrotational wind serves as the primary process by which EC Edisoana continues interacting with and shaping the jet stream. During the RI period, positive differential PV advection by the irrotational wind (;15-20 PVU day21; Fig. 15a) erodes the poleward portion of the polar trough and allows the downstream jet to backbuild, helping EC Edisoana to maintain its position in the equatorward entrance region of this jet. Simultaneously, negative PV advection by the nondivergent wind in the base of the polar trough (Fig. 15c) serves to bring high- magnitude PV air east faster than the high-magnitude PV air to the south. With the base of the polar trough being advected eastward faster than its core, the ensuing evolution is reminiscent of a PV hook from the second life cycle (LC2) discussed in Thorncroft et al. (1993) and Thorncroft and Jones (2000). Development of a LC2- style PV hook in association with a former TC suggest that EC Edisoana may have deepened more than it would have in the absence of either feature (Thorncroft and Jones 2000). It also appears the interaction between EC Edisoana and the midlatitude flow serves to separate the merged STJ/PFJ into two distinct pieces upstream and down- stream of the cyclone. This process becomes apparent around 0000 UTC 9 March as the strong downstream merged jet spanning the crest of the phased ridge sep- arates from a weaker, upstream merged jet streak in the base of the polar trough (Fig. 14b). During this jet- separation process, the peak wind speed in the down- stream jet quickly shifts eastward from ;708 to 1008Ein 12 h. A simultaneous increase in the maximum wind speed observed in the upstream jet streak (;20 m s21 in 6 h) further enhances the favorable synoptic-scale forc- ing from the associated ageostrophic circulations over EC Edisoana, as the cyclone is clearly located in both the equatorward entrance region of the downstream jet and the poleward exit region of the upstream jet. Typically, when the above-described superposition of synoptic-scale forcing for ascent occurs, the cyclone found in such a favorable location is subsequently ob- served to deepen, sometimes rapidly. Unlike the para- digm of rapidly deepening midlatitude cyclones being located in a position that is favorable for ascent in both the equatorward entrance region of a downstream jet and the poleward exit region of an upstream jet as rapid deepening commences (e.g., Uccellini and Kocin 1987), EC Edisoana begins deepening in the presence of only a single midlatitude jet (Fig. 14a). While this is not necessarily uncommon, EC Edisoana then develops a secondary jet streak on its equatorward side around the base of the upstream trough such that the surface cyclone associated with EC Edisoana is in the poleward exit region of this new secondary upper-level jet streak (Figs. 14b,c). This development of the secondary jet streak upstream of EC Edisoana can be associated with the approach of the polar trough. As the trough ap- proaches the surface cyclone, the surface cyclone ''phase locks'' with the polar trough in a dynamical process described by Hoskins et al. (1985). This phase locking allows for a mutual intensification of both the surface and upper-level PV anomalies that, in turn, aids in the intensification of the upstream jet streak in the base of the polar trough. Once separated and intensified, the upstream jet streak is positioned such that its poleward exit region is superposed with the equatorward entrance region of the downstream jet and this coupled jet structure is such that strong synoptic-scale forcing for ascent can be inferred over EC Edisoana (Fig. 14b). As the upstream jet streak intensifies, the downstream jet streak weakens and progresses farther downstream and away from EC Edisoana near the end of the RI phase (Fig. 14c), leaving the EC with forcing for ascent from only the poleward exit region of the upstream jet streak. This appears to be an unusual sequence of events due to the existence of a deep preexisting cyclone moving into a region favorable for further deepening and, through phase locking with the polar trough and sub- sequent interactions with the upper-level jets during the period of rapid deepening, making the environment even more favorable for deepening. Physically, the in- tensification of the upstream jet streak could also be tied to an increase in vertical motion (Fig. 16) associated with the deepening of EC Edisoana and the resultant strengthening of the thermally indirect circulation lo- cated in the exit region of the upstream jet streak. With these strong positive feedbacks occurring on relatively short time scales, it is unclear whether such a favorable environment would have come to exist had TC Edisoana not maintained its intensity through the ET process and remained a potent EC at the beginning of its interaction with the polar trough. Gridded reanalyses also indicated the development of a short-lived low pressure center between the transitioning TC Edisoana and the ap- proaching long-wave trough (see MSLP contoured in Figs. 7, 11a, and 14a). While this MSLP minimum is weak and is quickly absorbed by Edisoana, it suggests that a significant extratropical cyclone may have been able to form in this environment without the complicating interaction of TC Edisoana, although such a hypothesis would need to be tested via modeling experiments. Harr and Elsberry (2000) present additional insight into the evolution of strengthening or weakening post-TCs down- stream of a midlatitude trough. e. Backward trajectory analysis To investigate the source regions of air parcels that were involved in the development of the PV hook documented in section 4d, a backward trajectory anal- ysis was performed for the 72 h prior to 0000 UTC 9 March. Calculations utilized spatial and temporal in- terpolation between the 6-hourly grids in the CFSR data. Trajectories were taken from 300 hPa with 28 spacing in the region where 300-200-hPa layer-average PV was greater than 22 PVU between 528 and 588S and 708 and 788E. These parcel trajectories can be separated into three groups based upon similar origins (both hor- izontal and vertical) and evolutions over the 72-h anal- ysis (Fig. 17a). The first group of parcels to interact with TC Edisoana is entrained into the midlevels of the circulation of the TC late on 6 March and is represented by the blue trajectory in Fig. 17b. Late on 7 March a second group of parcels is entrained into the circula- tion of Edisoana and this group is represented by the red trajectory in Fig. 17b. These parcels originate at very low levels in the boundary layer (Fig. 18a) on the equator- ward side of the cyclone (Fig. 17a) and possess relatively low initial potential temperatures. The third group of parcels originates in the upper troposphere in the vi- cinity of the subtropical jet around southern Africa on 6 March and is represented by the green trajectory in Fig. 17b. These subtropical jet parcels are not entrained into the circulation of the TC or EC, but instead wrap around the downstream side of the long-wave trough that in- duces the RI of EC Edisoana while generally remaining above 500 hPa and conserving their potential tempera- tures throughout the 72-h analysis window. The midlevel parcel group and the tropical boundary layer parcel group both experience significant vertical motion throughout the 72-h analysis window. Midlevel parcels originate ;100-200 hPa higher than the tropical boundary layer parcels, with the tropical boundary layer parcels vertically dispersed between 700 and 900 hPa at 0000 UTC 8 March (Figs. 17 and 18a). In the subsequent 24 h, these parcels rapidly ascend ;400-600 hPa to the termination of trajectories at 300 hPa at 0000 UTC 9 March (Fig. 18a). Although less rapid, the midlevel parcels also undergo a significant period of ascent be- tween 0000 UTC 8 March and 0000 UTC 9 March, with vertical displacements nearly as large as the parcels from the tropical boundary layer group (350-500 hPa). This rapid ascent occurs in the region of strong upward ver- tical motion discussed in the previous section. Non- conservation of potential temperature is observed along the parcel paths indicating that diabatic heat release, associated with strong rising motion, is at work. Heating along the paths of both parcel groups is observed to be as high as ;35 K in the final 24-h period prior to 0000 UTC 9 March, although most parcels average 20- 25 K over the period (Fig. 18b). The heating and ascent rates have been shown to be fairly stable over small variations in time and space. Further testing as to the sensitivity of these trajectories show that trajectories run with ERA-Interim data (not shown) provide similar results to these run with CFSR data. A similar backward trajectory analysis was performed on parcels originating from regions of high-magnitude negative PV. Trajectories were taken as previously described except from within the developing PV hook structure, defined as PV values less than 22 PVU. These trajectories demonstrate the complex nature of the SH flow patterns, with most parcels located within the polar trough at 0000 UTC 9 March and having tracked through at least one upstream trough over the previous 3 days (Fig. 19). In general, these parcels originate over oceanic regions of the South Atlantic between 308 and 708S. However, the trajectories in Fig. 19 indicate that three parcels originate farther south (;758S) over the Antarctic continent. Additional sampling of spatial and temporal perturbations in the parcel trajectories (not shown) demonstrate that this source region is more susceptible to changes in the initial conditions chosen for the trajectory calculations than others in both the high- and low-magnitude PV regions, but does appear on multiple trajectory samples (not shown). Assuming the air in the polar trough does contain some air of Antarctic origin, this could account for increased baro- clinicity associated with the trough. A distinct confluence of parcels occurs in the final ; 12 h prior to the trajectory termination points at 0000 UTC 9 March as parcels originating from the subtropics and parcels originating south of 408S become juxtaposed near the base of the polar trough. Furthermore, each grouping of parcels also experiences confluence in the vertical sense, with subtropical trajectories sinking from above 200 hPa and polar trajectories rising from below 300 hPa (Fig. 19). The confluence of these trajectories further strengthens the earlier discussion of the STJ-PFJ merger downstream of EC Edisoana. Some evidence for the collapse of the downstream half-wavelength is present as well, with the enhanced meridional motion with the trajectories in their final 12-24 h near 608E suggesting a local amplification of the upper-level flow pattern associated with the stronger upper-level jet between the progressive upstream trough and quasi-stationary downstream-phased ridge discussed in the previous section. 5. Discussion TCs in the SWIO remain an infrequently studied phenomenon. A brief climatology of both TC and ET events in the SWIO highlighted an average of nine TCs per year, of which four undergo ET. This results in an ET rate of 43.8%, a rate higher than all other TC basins with the exception of the North Atlantic. The majority of TC events (86%) and almost all ET events occur between the months of November and April, suggesting that these months form a rough basis for the boundaries of a SWIO TC season, consistent with previous sugges- tions from Valadon (1992) and Mavume et al. (2009). A case study of TC Edisoana, motivated by the cy- clone's rank as the deepest post-TC analyzed during the 1989-2013 period, was also constructed. Key compo- nents in the post-tropical RI of EC Edisoana included a phased ridge downstream of the EC, and an ap- proaching upstream polar trough. Significant and pro- longed diabatic heating aided in strengthening the subtropical ridge downstream of EC Edisoana by en- hancing upper-level outflow (represented by irrotational wind) and the resultant positive PV advection by that irrotational wind (Figs. 7a-c). This positive PV advection allowed the ridge to build poleward into the polar flow and enabled the two-flow stream to phase through the ridge (Figs. 10, 11, and 12). As a result of this strength- ening of the PV gradient and the poleward edge of the ridge (Figs. 15a,b), the upper-level winds poleward of the phased ridge intensified, creating a stronger upper- level jet and supporting a more favorable environment for cyclone development in the poleward entrance re- gion of the jet streak (Carlson 1998; Evans and Prater- Mayes 2004). The approaching polar trough remained progressive while the two ridges downstream of EC Edisoana phased, reducing the distance between the polar trough and the phased ridge and resulting in the collapse of the half-wavelength downstream of the polar trough and its associated surface cyclone (Figs. 13a-c). A downstream half-wavelength collapse has been shown to be condu- cive for the intensification of ECs (Rogers and Bosart 1991) as cyclonic vorticity advection and geostrophic adjustment occur on smaller spatial scales and necessitate a stronger ageostrophic circulation in response. While it would seem likely that a collapse of the half-wavelength downstream of a transitioning and reintensifying TC would occur relatively frequently, such a process does not appear to be widely documented for cases of ET. In the case of TC Edisoana, establishing a connection between the ridge amplification and phasing triggered by the upper- level outflow associated with the TC and the subsequent collapse of the downstream half-wavelength due to the phased quasi-stationary ridge suggests that such a process is tied to these relatively ubiquitous events throughout the ET process and could occur in any ET event, although the frequency of such a sequence remains unknown. During the interaction between the polar trough and the phased downstream ridge, the upper-level jet streaks associated with both features played a crucial role in the strengthening of Edisoana. The initial jet streak span- ning the crest of the downstream ridge provided support for an extended period of steady intensification during and immediately following ET as the cyclone main- tained its position in the equatorward entrance region to the jet streak, akin to the post-tropical intensification process described with TC Irene (1999) by Evans and Prater-Mayes (2004). In an evolution not documented in previous ET case studies, a secondary jet streak develops on the down- stream side of the base of the polar trough and intensifies through the RI of EC Edisoana. Simultaneously, the core of the jet streak across the crest of the downstream ridge shifts farther downstream. As the two jet streaks coevolve in this manner, the source of the forcing for the RI of the EC shifts from the equatorward entrance re- gion of the downstream jet to the intensifying jet streak in the base of the polar trough. This ''handoff'' of forcing from the downstream jet to the upstream is a variation on the coupled jet structure more commonly discussed in cases of RI with ordinary ECs. In such a situation, the downstream and upstream jet streaks position them- selves in a manner to allow the favorable superposition of their ageostrophic circulations and respective forcing for synoptic-scale ascent. Conversely, EC Edisoana be- gan RI while in the equatorward entrance region of the downstream jet streak and, while undergoing RI, ap- pears to have help intensified the upstream jet streak. Suggestions of a similar process, albeit less distinct, were presented by Atallah and Bosart (2003) for the case of TC Floyd (1999), where the ET of the TC appears to have been associated with a break in the jet stream poleward of the cyclone's location. Further work would be necessary to objectively attribute this jet streak de- velopment to the RI of EC Edisoana. From a PV perspective, the synoptic-scale evolution around EC Edisoana during the RI period is well de- scribed by the LC2 paradigm presented by Thorncroft et al. (1993). In general, cyclones following the LC2 life cycle paradigm develop well-formed PV hooks (see Figs. 15c,d) in association with strong PV anomalies at the dynamic tropopause and high potential temperature anomalies near the surface, usually in the warm seclu- sion (Fig. 14c). This positive surface potential tem- perature anomaly, in the warm seclusion, allows for a ''phase lock'' with the approaching upper-level trough as described by Hoskins et al. (1985) and is shown to be favorable for intensification of the cyclone. Addition- ally, Hart et al. (2006) documented North Atlantic TCs undergoing ET and showed those that develop a warm seclusion are more likely to intensify as an EC than those that experience cold-core post-tropical evolutions. Regardless of the synoptic perspective used to analyze this storm, diabatic heating played a significant role in amplifying the ridge downstream of EC Edisoana. In- tense and prolonged upward vertical motion (e.g., Figs. 10, 11, and 12; Fig. 14; Fig. 18) in the developing warm sector of the transitioning cyclone aided in the production of higher PV at upper levels near the dynamic tropo- pause. This resulted in a steepening of the tropopause as observed in the meridional cross sections through the cyclone and the jet streaks (Figs. 10, 11,and12) similar to the tropopause steepening observed by Bosart and Lackmann (1995) during the reintensification of TC David (1979). When this occurs in the presence of an upper-level disturbance on the tropopause, Bosart and Lackmann (1995) state the steepening of the tropo- pause makes the environment more favorable for as- cent and low-level cyclonic vorticity generation. The synoptic setup surrounding the intensification of EC Edisoana led to the initiation of RI prior to the coupling of the upper-level jet streaks. It seems likely this evolution is associated with the intensification of the downstream jet streak by diabatic heating and the re- sulting upper-level irrotational outflow from the TC, but a more complete examination of similar ET events would be necessary to be certain as such an evolution of upper-level jet streaks in relation to the timing of the RI of the associated surface cyclone has not been described often in the literature. Regardless of the frequency of this kind of cyclone-jet interactions, both the phased ridging and jet streak evolution observed with EC Edisoana have highlighted the importance of these synoptic-scale fea- tures in the post-tropical life cycle of TC Edisoana. Al- though not shown here, an RI between a high-amplitude upstream trough and a phased downstream ridge is similar to the evolution observed in the case of TC Jade from April 2009 that originally motivated this study. While many details in the case of Jade are more complex than the case of Edisoana, broad similarities between the two cases suggest the possibility of some common features among some cases of RI as a post-TC in the SWIO. It would be enlightening to see if future studies find similar evolutions in other cases of extreme post-TC reintensification. The Lagrangian perspective helps to identify the diverse source regions of the synoptic-scale factors that contrib- uted to the environment surrounding EC Edisoana during the period of RI. A backward trajectory analysis of the warm sector emphasized the role of TC Edisoana as it entrained air parcels from the tropical boundary layer prior to undergoing ET (Fig. 17). These parcels rapidly ascended shortly after the completion of ET and contrib- ute a significant portion of the area of the low-magnitude PV downstream of Edisoana (Figs. 18a,b). Simultaneously, the polar trough upstream of EC Edisoana brought to- gether upper-level air parcels originating from near the STJ, PFJ, and over the Antarctic continent (Fig. 19). The collapse of the downstream half-wavelength between the downstream ridge with its parcels lifted from the tropical boundary layer and the upstream trough with parcels brought in from the Antarctic allows for un- usually strong temperature gradients (and associated PV gradients) to contribute to the RI of EC Edisoana. Acknowledgments. The authors acknowledge the support of National Science Foundation Grant ATM- 0553017, especially for travel to the Fourth International Workshop on Extratropical Transition where this work was rekindled and refined. The authors extend our gratitude to Sebastian Langlade of M^et^eo-France La R^eunion, who provided software containing best track data and other historical records from RSMC La R^eunion. Finally, we appreciate the works of three anonymous reviewers whose comments greatly aided in the revision of this manuscript. 1 Post-tropical cyclone (post-TC) terminology was introduced by the National Weather Service in 2010 and describes the broad category of cyclone classifications that a TC can become after a TC, including a transitioning TC and an EC. While a number of other terms can be used for these transitioning and post-transition TCs (e.g., extratropically transitioned, post-ET, or former TC), the au- thors have chosen post-TC for its simplicity and all-encompassing nature. REFERENCES Agusti-Panareda, A., C. D. Thorncroft, G. C. Craig, and S. L. Gray, 2004: The extratropical transition of Hurricane Irene (1999): A potential-vorticity perspective. Quart. J. Roy. Meteor. Soc., 130, 1047-1074, doi:10.1256/qj.02.140. _____, S. L. Gray, G. C. Craig, and C. Thorncroft, 2005: The ex- tratropical transition of Tropical Cyclone Lili (1996) and its crucial contribution to a moderate extratropical de- velopment. Mon. Wea. Rev., 133, 1562-1573, doi:10.1175/ MWR2935.1. Anwender, D., S. C. Jones, M. Leutbecher, and P. A. Harr, 2010: Sensitivity experiments for ensemble forecasts of the extra- tropical transition of Typhoon Tokage (2004). Quart. J. Roy. Meteor. Soc., 136, 183-200, doi:10.1002/qj.527. Archambault, H. M., L. F. Bosart, D. Keyser, and J. M. Cordeira, 2013: A climatological analysis of the extratropical flow response to recurving western North Pacific tropical cyclones. Mon. Wea. Rev., 141, 2325-2346, doi:10.1175/MWR-D-12-00257.1. Atallah, E. H., and L. F. Bosart, 2003: The extratropical transition and precipitation distribution of Hurricane Floyd (1999). Mon. Wea.Rev.,131,1063-1081,doi:10.1175/1520-0493(2003)131,1063: TETAPD.2.0.CO;2. 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, doi:10.1175/1520-0493(1995)123,3268: PTCRIA.2.0.CO;2. Carlson, T. N., 1998: Mid-Latitude Weather Systems. Amer. Me- teor. Soc., 507 pp. Chen, H., and W. Pan, 2010: Targeting studies for the extratropical transition of Hurricane Fabian: Signal propagation, the interaction between Fabian and midlatitude flow, and an observation strategy. Mon. Wea. Rev., 138, 3224-3242, doi:10.1175/2010MWR2888.1. Cordeira, J. M., and L. F. Bosart, 2011: Cyclone interactions and evolutions during the ''perfect storms'' of late October and early November 1991. Mon. Wea. Rev., 139, 1683-1707, doi:10.1175/2010MWR3537.1. Dare, R. A., and N. E. Davidson, 2004: Characteristics of tropical cyclones in the Australian region. Mon. Wea. Rev., 132, 3049- 3065, doi:10.1175/MWR2834.1. Davis, C. A., 2010: Simulations of subtropical cyclones in a baro- clinic channel model. J. Atmos. Sci., 67, 2871-2892, doi:10.1175/2010JAS3411.1. Dee, D. P., and Coauthors, 2011: The ERA-Interim reanalysis: Configuration and performance of the data assimilation sys- tem. Quart. J. Roy. Meteor. Soc., 137, 553-597, doi:10.1002/ qj.828. DiMego, G. J., and L. F. Bosart, 1982a: The transformation of Tropical Storm Agnes into an extratropical cyclone. Part I: The observed fields and vertical motion computations. Mon. Wea. Rev., 110, 385-411, doi:10.1175/1520-0493(1982)110,0385: TTOTSA.2.0.CO;2. _____, and _____, 1982b: The transformation of Tropical Storm Agnes into an extratropical cyclone. Part II: Moisture, vorticity, and kinetic energy budgets. Mon. Wea. Rev., 110, 412-433, doi:10.1175/1520-0493(1982)110,0412:TTOTSA.2.0.CO;2. Evans, J. L., and R. E. Hart, 2003: Objective indicators of the life cycle evolution of extratropical transition for Atlantic trop- ical cyclones. Mon. Wea. Rev., 131, 909-925, doi:10.1175/ 1520-0493(2003)131,0909:OIOTLC.2.0.CO;2. _____, and B. E. Prater-Mayes, 2004: Factors affecting the post- transition intensification of Hurricane Irene (1999). Mon. Wea. Rev., 132, 1355-1368, doi:10.1175/1520-0493(2004)132,1355: FATPIO.2.0.CO;2. Foley, G. R., and B. N. Hanstrum, 1994: The capture of tropical cyclones by cold fronts off the west coast of Australia. Wea. Forecasting, 9, 577-592, doi:10.1175/1520-0434(1994)009,0577: TCOTCB.2.0.CO;2. Grams, C. M., and Coauthors, 2011: The key role of diabatic pro- cesses in modifying the upper-tropospheric wave guide: A North Atlantic case-study. Quart. J. Roy. Meteor. Soc., 137, 2174-2193, doi:10.1002/qj.891. _____,S.C.Jones,C.A.Davis,P.A.Harr,andM.Weissmann,2013: The impact of Typhoon Jangmi (2008) on the midlatitude flow. Part I: Upper-level ridge building and modification of the jet. Quart. J. Roy. Meteor. Soc., 139, 2148-2164, doi:10.1002/qj.2091. Harr, P. A., and R. L. Elsberry, 2000: Extratropical transition of tropical cyclones over the western North Pacific. Part I: Evolution of structural characteristics during the transi- tion process. Mon. Wea. Rev., 128, 2613-2633, doi:10.1175/ 1520-0493(2000)128,2613:ETOTCO.2.0.CO;2. _____, and J. M. Dea, 2009: Downstream development associated with the extratropical transition of tropical cyclones over the western North Pacific. Mon. Wea. Rev., 137, 1295-1319, doi:10.1175/2008MWR2558.1. Hart, R. E., 2003: A cyclone phase space derived from thermal wind and thermal asymmetry. Mon. Wea. Rev., 131, 585-616, doi:10.1175/1520-0493(2003)131,0585:ACPSDF.2.0.CO;2. _____, and J. L. Evans, 2001: A climatology of the extra- tropical transition of Atlantic tropical cyclones. J. Cli- mate, 14, 546-564, doi:10.1175/1520-0442(2001)014,0546: ACOTET.2.0.CO;2. _____, _____, and C. Evans, 2006: Synoptic composites of the extra- tropical transition life cycle of North Atlantic tropical cy- clones: Factors determining posttransition evolution. Mon. Wea. Rev., 134, 553-578, doi:10.1175/MWR3082.1. Hoskins, B. J., M. E. McIntyre, and A. W. Robertson, 1985: On the use and significance of isentropic potential vorticity maps. Quart. J. Roy. Meteor. Soc., 111, 877-946, doi:10.1002/ qj.49711147002. Jones, S. C., and Coauthors, 2003: The extratropical transition of tropical cyclones: Forecast challenges, current understanding, and future directions. Wea. Forecasting, 18, 1052-1092, doi:10.1175/1520-0434(2003)018,1052:TETOTC.2.0.CO;2. Kitabatake, N., 2008: Extratropical transition of tropical cyclones in the western North Pacific: Their frontal evolution. Mon. Wea. Rev., 136, 2066-2090, doi:10.1175/2007MWR1958.1. Klein, P. M., P. A. Harr, and R. L. Elsberry, 2000: Extratropical transition of western North Pacific tropical cyclones: An overview and conceptual model of the transformation stage. Wea. Fore- casting, 15, 373-395, doi:10.1175/1520-0434(2000)015,0373: ETOWNP.2.0.CO;2. _____, _____, and _____, 2002: Extratropical transition of western North Pacific tropical cyclones: Midlatitude and tropical cyclone contri- butions to reintensification. Mon. Wea. Rev., 130, 2240-2259, doi:10.1175/1520-0493(2002)130,2240:ETOWNP.2.0.CO;2. Knapp, K. R., M. C. Kruk, D. H. Levinson, H. J. Diamond, and C. J. Neumann, 2010: The International Best Track Archive for Climate Stewardship (IBTrACS). Bull. Amer. Meteor. Soc., 91, 363-376, doi:10.1175/2009BAMS2755.1. Mavume, A. F., L. Rydberg, M. Rouault, and J. R. E. Lutjeharms, 2009: Climatology and landfall of tropical cyclones in the south- west Indian Ocean. Western Indian Ocean J. Mar. Sci., 8, 15-36. McTaggart-Cowan, R., J. R. Gyakum, and M. K. Yau, 2001: Sen- sitivity testing of extratropical transitions using potential vorticity inversions to modify initial conditions: Hurricane Earl case study. Mon. Wea. Rev., 129, 1617-1636, doi:10.1175/ 1520-0493(2001)129,1617:STOETU.2.0.CO;2. _____, _____, and _____, 2003: The influence of the downstream state on extratropical transition: Hurricane Earl (1998) case study. Mon. Wea. Rev., 131, 1910-1929, doi:10.1175/2589.1. _____, _____, and _____, 2004: The impact of tropical remnants on extratropical cyclogenesis: Case study of Hurricanes Danielle and Earl (1998). Mon. Wea. Rev., 132, 1933-1951, doi:10.1175/ 1520-0493(2004)132,1933:TIOTRO.2.0.CO;2. Orlanski, I., and E. K. M. Chang, 1993: Ageostrophic geo- potential fluxes in downstream and upstream development of baroclinic waves. J. Atmos. Sci., 50, 212-225, doi:10.1175/ 1520-0469(1993)050,0212:AGFIDA.2.0.CO;2. _____, and J. Sheldon, 1993: A case of downstream baroclinic development over western North America. Mon. Wea. Rev., 121, 2929-2950, doi:10.1175/1520-0493(1993)121,2929: ACODBD.2.0.CO;2. Pantillon, F., J.-P. Chaboureau, C. Lac, and P. Mascart, 2013: On the role of a Rossby wave train during the extratropical transition of Hurricane Helene (2006). Quart. J. Roy. Meteor. Soc., 139, 370-386, doi:10.1002/qj.1974. Riemer, M., and S. C. Jones, 2010: The downstream impact of tropical cyclones on a developing baroclinic wave in idealized scenarios of extratropical transition. Quart. J. Roy. Meteor. Soc., 136, 617-637, doi:10.1002/qj.605. _____, _____, and C. A. Davis, 2008: The impact of extratropical transition on the downstream flow: An idealized modelling study with a straight jet. Quart. J. Roy. Meteor. Soc., 134, 69- 91, doi:10.1002/qj.189. Ritchie, E. A., and R. L. Elsberry, 2007: Simulations of the extra- tropical transition of tropical cyclones: Phasing between the upper-level trough and tropical cyclones. Mon. Wea. Rev., 135, 862-876, doi:10.1175/MWR3303.1. Rogers, E., and L. F. Bosart, 1991: A diagnostic study of two in- tense oceanic cyclones. Mon. Wea. Rev., 119, 965-996, doi:10.1175/1520-0493(1991)119,0965:ADSOTI.2.0.CO;2. Saha, S., and Coauthors, 2010: The NCEP Climate Forecast System Reanalysis. Bull. Amer. Meteor. Soc., 91, 1015-1057, doi:10.1175/ 2010BAMS3001.1. Schenkel, B. A., and R. E. Hart, 2012: An examination of tropical cyclone position, intensity, and intensity life cycle within at- mospheric reanalysis datasets. J. Climate, 25, 3453-3475, doi:10.1175/2011JCLI4208.1. Sekioka, M., 1956a: A hypothesis on complex of tropical and ex- tratropical cyclones for typhoon in the middle latitudes. I. Synoptic structure of Typhoon Marie over the Japan Sea. J. Meteor. Soc. Japan, 34, 276-287. _____, 1956b: A hypothesis on complex of tropical and extra- tropical cyclones for typhoon in the middle latitudes. II. Synoptic structure of Typhoons Louise, Kezia and Jane passing over the Japan Sea. J. Meteor. Soc. Japan, 34, 336- 345. Simmons, A. J., and B. J. Hoskins, 1979: The downstream and upstream development of unstable baroclinic waves. J. Atmos. Sci., 36, 1239-1254, doi:10.1175/1520-0469(1979)036,1239: TDAUDO.2.0.CO;2. Sinclair, M. R., 1993: Synoptic-scale diagnosis of the extratropical transition of a southwest Pacific tropical cyclone. Mon. Wea. Rev., 121, 941-960, doi:10.1175/1520-0493(1993)121,0941: SSDOTE.2.0.CO;2. _____, 2002: Extratropical transition of southwest Pacific tropical cy- clones. Part I: Climatology and mean structure changes. Mon. Wea. Rev., 130, 590-609, doi:10.1175/1520-0493(2002)130,0590: ETOSPT.2.0.CO;2. _____, 2004: Extratropical transition of southwest Pacific tropical cy- clones. Part II: Midlatitude circulation characteristics. Mon. Wea. Rev., 132, 2145-2168, doi:10.1175/1520-0493(2004)132,2145: ETOSPT.2.0.CO;2. Thorncroft, C. D., and S. C. Jones, 2000: The extratropical transitions of Hurricanes Felix and Iris in 1995. Mon. Wea. Rev., 128, 947-972, doi:10.1175/1520-0493(2000)128,0947:TETOHF.2.0.CO;2. _____, B. J. Hoskins, and M. E. McIntyre, 1993: Two paradigms of baroclinic-wave life-cycle behaviour. Quart. J. Roy. Meteor. Soc., 119, 17-55, doi:10.1002/qj.49711950903. Torn, R. D., 2010: Diagnosis of the downstream ridging associ- ated with extratropical transition using short-term ensem- ble forecasts. J. Atmos. Sci., 67, 817-833, doi:10.1175/ 2009JAS3093.1. Uccellini, L. W., and P. J. Kocin, 1987: The interaction of jet streak circulations during heavy snow events along the East Coast of the United States. Wea. Forecasting, 2, 289-308, doi:10.1175/ 1520-0434(1987)002,0289:TIOJSC.2.0.CO;2. Uppala, S. M., and Coauthors, 2005: The ERA-40 Re-Analysis. Quart. J. Roy. Meteor. Soc., 131, 2961-3012, doi:10.1256/ qj.04.176. Valadon, Y., 1992: Variability of tropical cyclones over the southwest Indian Ocean. J. Afr. Meteor. Soc., 1, 45-51. KYLE S. GRIFFIN AND LANCE F. BOSART Department of Atmospheric and Environmental Sciences, University at Albany, State University of New York, Albany, New York (Manuscript received 8 September 2013, in final form 12 March 2014) Corresponding author address: Kyle S. Griffin, University of Wisconsin-Madison, 1225 W. Dayton St., Madison, WI 53706. E-mail: [email protected] (c) 2014 American Meteorological Society |