3.1 Conceptual Models of Extratropical Transition

Both Evans and Prater-Mayes (2004) and Agusti-Panareda et al. (2004) have studies the explosive reintensification of Hurricane Irene (1999) as an extratropical storm. The former approaches the ET problem through a model simulation and a largely quasigeostrophic diagnosis while the latter develops a potential vorticity (PV) conceptual model. Each study provides a unique framework with which to understand, model, and diagnose ET. Their approaches will be summarized, combined, and augmented with a vector frontogenesis function analysis (Harr and Elsberry 2001) and remote sensing products including QuikScat derived winds and satellite imagery to provide a coherent observational model of marine ET. The troposphere will be scrutinized in terms of surface winds and frontogenesis, mid-level frontogenesis and trough interaction, and upper-level PV anomalies and jet interactions with the TC. It is commonly accepted that the midlatitude circulation into which the TC moves determines the final outcome of the ET (Jones et al. 2003). In this analysis, the midlatitude circulation not only influences the intensity and outcome of ET, it also determines the structure and evolution of low-level cyclones and fronts. First, it is hypothesized that the large-scale circulation interaction during marine ET often spawns a Shapiro-Keyser (1990) type extratropical cyclone with the observed properties of frontal fracture and warm core seclusion. This extratropical cyclone may deepen significantly after ET as in the case of many storms including Irene (1999) and Kate (2003). A second observed extratropical transition does not mimic the Shapiro-Keyser lifecycle but maintains a hybrid tropical structure as evidenced by Felix (1995) (Thorncroft and Jones, 2000) and Fabian (2003) exemplified by weak frontal structure and baroclinic interaction.

3.2 Introduction to Structural Changes

As a northern hemisphere tropical cyclone moves northward and begins to interact with the baroclinic environment (westerly winds, increased vertical shear, lower sea surface temperatures (SST) or strong SST gradients), it evolves from an axisymmetric warm core tropical cyclone into an asymmetric cold-core extratropical cyclone. In addition, the translation speed of the cyclone increases dramatically under the effects of the midlatitude steering currents. Precipitation and strong winds near the center of the tropical cyclone circulation become distinctly asymmetric and often expand greatly in area due to interaction with synoptic features including midlatitude cyclones and upper-level features such as troughs and jet streaks. Interaction with these features, especially in the presence of vertical wind shear, contributes to loss of symmetry in a tropical storm and frontogenesis (Evans and Prater-Mayes 2004). The tropical cyclone may continue to weaken under the effects of the baroclinic environment, merge with an existing extratropical cyclone, or reintensify into its own extratropical cyclone.

Figure 3.1: Klein et al. (2000) conceptual model of transformation stage of ET in the Western Pacific: (1) environmental advection of cold, dry air equatorward creating open cell cumulus clouds; (2) decreased TC convection on the western semicircle with dry slot formation in step 1, which extends southward during steps 2 and 3; (3) environmental advection of warm, moist air poleward maintains convection in the eastern half creating an asymmetric distribution of clouds and precipitation in steps 1 and 2; steps 2 & 3 feature a southerly jet that ascend over the tilted baroclinic zone (4); ascent undercut by dry-adiabatic descent produces cloudbands that wrap westward and equatorward around the storm center maintaining convection (5); dry-adiabatic descent occurs near circulation center to erode eyewall convection in step 3; (6) cirrus shield with a sharp cloud edge if confluent with polar jet. (Adapted from Klein et al. (2000) their Figure 5.)

Klein et al. (2000) develop a conceptual model of the ET transformation stage from a composite study of 30 North Pacific typhoons. Evans and Hart (2003) prefer the terminology “onset of transition� rather than “start of transformation� but essentially reference the same process of asymmetric TC evolution. The transformation stage begins as the tropical cyclone moves over lower SSTs with its outer circulation impinging upon a preexisting baroclinic zone (Fig. 3.1). A dipole of (cold/warm) advection of (dry equatorward / moist poleward) moving air develops (west/east) of the TC. Typically a dry-slot forms in the southwestern quadrant of the storm as deep convection decreases over the western quadrant. However, warm moist air advected poleward maintains the deep convection over the northern and eastern quadrants of the storm. The poleward advected flow turns cyclonically and ascends over the tilted isentropic surfaces inherent with the preexisting baroclinic zone producing a region conducive for warm frontogenesis (Harr and Elsberry, 2000). This asymmetry in TC cloud patterns is acknowledged as an early indicator of transformation (Klein et al, 2000).

During step 2 of the Klein et al. (2000) conceptual model (Fig 3.1), the lower tropospheric temperature advection dipole accentuates the baroclinic zone as well as expands the dry slot over the southern quadrant. The ascending poleward flow of warm moist air turns cyclonically and subsides into the western quadrant. A vertical motion dipole exists with dry adiabatic descent west of the storm center with ascent east of the center. The increased vertical wind shear south of the upper-tropospheric jet begins to ventilate the storm by advecting the top of the upper-tropospheric warm core downstream. Deep convection may still persist in the inner-core even as mid-level westerlies envelop the weakening midtropospheric warm core.

During step 3 of the conceptual model (Fig. 3.1), the storm becomes embedded in the baroclinic zone with an increase in vertical wind shear, decrease in SSTs, and a temperature advection dipole. The dry-adiabatic descent west of the storm center continues to erode the inner-core convection with only a weak warm core at low-levels remaining. A broad multilayer cloud mass exists on the poleward side of a warm front with a weaker cloud band to the southeast that resembles a cold front (Fig. 3.2). The broad cirrus shield with a sharp edge implies confluence between the TC outflow and the upper-level jet (Bader et al. 1995). Harr and Elsberry (2000) explain the continuing convection and cirrus shield to the north and northeast as the commencement of warm frontogenesis, a more vigorous process than cold frontogenesis in most ET cases. Ascent northwest of the storm center is undercut by dry adiabatic descent originating from the poleward advected environmental flow as part of the eastern vertical motion dipole branch. This poleward flow of warm, most air ascends over the tilted isentropes and joins a strong southwesterly jet aloft that resembles the warm conveyor belt of an extratropical storm (Carlson 1991).

It is the interaction of the TC circulation with the baroclinic zone and its associated vertical wind shear that initiates the transformation stage of ET. However, not all TCs that enter the transformation stage and may dissipate due to extreme vertical shear or vortex spin-down over cooler SSTs (Eastern Pacific hurricanes commonly dissipate quite quickly over cooler waters without experiencing unfavorable vertical wind shear). Furthermore, TCs that enter the transformation stage do not always complete transition. The definition of a “transitioned� tropical cyclone implies that it is indistinguishable from an extratropical cyclone; the system specifically develops a lower troposphere cold-core (Evans and Hart, 2003). Consequently, differences in TC transformation and ultimate completion of ET may depend upon the angle of entry of the TC into the baroclinic zone and the specific characteristics of the midlatitude circulation into which the TC translates (Hart and Evans 2004).

After completion of ET, an extratropical cyclone develops with an extensive warm frontal region but with an ill-defined cold front. The cold front is often suppressed due to a direct thermal circulation that includes the descent of cold air from upstream of the reintensifying cyclone (Jones et al. 2003). A constructive or destructive interaction with a preexisting midlatitude system may either reintensify or decay the transitioned TC. Without interaction, the transitioned TC may continue to decay in terms of MSLP but continue to have strong winds in excess of hurricane force.

Figure 3.2: Schematic of midlatitude trough/TC interaction during ET. A multilayered cloud shield in the warm frontogenetic region poleward along the trough is caused by ascent of warm, moist air over the tilted isentropes of the baroclinic zone. Cool, dry air from behind the midlatitude cold front entrains into the southwest quadrant of the storm creating a dry slot. The strongest surface winds are associated with the cool, dry air along the southern side added to the storm translation speed. (Adapted from Fogarty 2002).

3.3 Trough Interaction and Vertical Shear

As the tropical cyclone translates poleward, it often increases its forward motion, encounters lower SSTs, and vertical shear due to interaction with a midlatitude mid-level trough. The effects of vertical shear on tropical cyclones have been studied by many (Frank and Ritchie 2001; Ritchie and Elsberry 2001; Reasor et al. 2000; Hanley et al. 2001) with considerable attention paid to ET. Briefly, the TC core will weaken under strong environmental vertical wind shear because the upper-level warm core cannot be sufficiently maintained upright to support the surface low pressure. Due to the minimum in inertial stability at the top of the TC, the warm core aloft is easily advected downstream. The ventilation of the warm core under the effects of strong environmental wind shear will reduce the strength of the warm core aloft. The height of the maximum warm core is thus reduced while the lower level warm core is enhanced. Subsidence due to convergence between the environmental winds and the cyclonic circulation of the TC are responsible for the warm core enhancement at low levels along with a rise in sea-level pressure. The cyclonic circulation is weakened further aloft, which entails even lower inertial stability. This negative feedback may weaken the upper-level warm core dramatically and erode deep convection. However, it is not well understood how midlatitude trough interactions weaken or strengthen a tropical cyclone’s circulation, yet numerical models attempt to resolve the issue (Ritchie and Elsberry 2001; 2003).

Hanley et al (2001) performed a composite analysis of TC-trough interactions and identified two instances where the trough favorably interacted with the TC and resulted in intensification. Most applicable to ET is the favorable interaction between an upper-level jet streak and the TC outflow. As the temperature gradient between the warm tropical cyclone outflow and the cold trough well to the west of the TC narrows, the intervening jet streak is enhanced, which in turn improves the upper-level TC outflow channels. Linear jet dynamics implies that ascent is favored beneath the right-entrance region of the jet streak where a conveniently located TC center would experience even greater divergence aloft and hence improved outflow. As a result of the small inertial stability aloft, the improved outflow will contribute to low-level convergence and hence cyclonic spinup of the TC (Hanley et al. 2001). This favorable interaction usually results in an increase in cyclone upper-level cirrus in satellite imagery due to the upper-level outflow enhancement. This situation is occurs when the trough is well away or ‘distant’ from the TC. As the trough nears, vertical wind shear weakens the TC by ventilating its warm core and creating an asymmetry in the cloud/precipitation patterns. Ascent is favored in the downshear left quadrant of the tropical cyclone coupled with the formation of a dry slot upshear in response to forced subsidence (Ritchie and Elsberry 2001). Weakening of the vortex usually results.

3.4 Paradigms of Cyclone Development

Harr and Elsberry (2000) note that many studies (Thorncroft et al. 1993; Evans et al. 1994; Shultz et al. 1998) have defined variability associated with the structures of maritime cyclones that depend on the dynamics of the large-scale circulation. Their examination of two transitions of Western Pacific Typhoons discusses the evolution of structural characteristics of a TC during ET with respect to two different baroclinic environmental influences. The use of vector frontogenesis functions (Chapter 5) describes the frontal evolution during ET with emphasis upon the interaction between the midlatitude baroclinic zone and the TC circulation. However, their analysis focuses on the mid-level tropospheric interactions (500hPa) where they found the largest magnitude of structural changes. Schultz et al. (1998) consider the effects of large-scale flow on low-level frontal structures in marine midlatitude cyclones through both observational and numerical approaches. Different upper-level flow regimes are found to produce variable distinct cyclone/frontal structures and evolutions, including the Norwegian and Shapiro-Keyser (Fig. 3.3) cyclone models. It is the latter of a large spectrum of different cases that is closest related and observed most often during marine ET.

The Norwegian cyclone model (Bjerknes 1919; Bjerknes and Solberg 1922) has proven to be a limited and inadequate explanation of possible cyclone/frontal structures and evolutions (Browning 1990; Shapiro and Keyser 1990; Evans et al. 1994; Bosart 1999). The model consists of a disturbance on the polar front that forms a warm/cold air advection dipole and creates warm and cold fronts. The occlusion occurs when the cold front catches up to the warm front, a fundamental difference with the Shapiro-Keyser model. A Norwegian model cold front is typically meridionally oriented and much stronger than the warm front. This is based upon the zonal index cycle of jet stream oscillation from weak westerlies in a high-amplitude planetary wave pattern (low zonal index) to strong westerlies in a low-amplitude planetary wave pattern (high zonal index) and back again (Rossby and Collaborators 1939; Rossby and Willett 1948; Namais 1950). Rossby and Willet (1948) characterized low zonal index by deep occlusions with north-south orientation of frontal systems with maximum east-west temperature contrasts.

The opposite configuration holds for periods of high zonal index with east-west oriented frontal systems and latitudinal temperature gradients. The former case usually occurs within diffluent blocking patterns downstream of the cyclone whereas the latter is associated with strongly confluent flow (Saucier 1955). Fundamentally, it is the deformation pattern associated with each pattern that favors either meridionally or zonally oriented fronts with their own respective thermodynamically forced secondary circulations. Sawyer (1950) noted that warm occlusions tend to occur within the jet stream entrance region where confluence deformation and thermal fields exist typically in high zonal index flow.

Figure 3.3: Shapiro-Keyser (1990) frontal-cyclone evolution: incipient broad-baroclinic phase (I), frontal fracture (II), bent-back front and frontal T-bone (III), and warm-core frontal seclusion (IV), Upper: sea level pressure (solid), fronts (bold), and cloud signature (shaded). Lower: temperature (solid), and cold and warm air currents (solid and dashed arrows, respectively).

3.5 Shapiro-Keyser (1990) Cyclone Model

The Shapiro-Keyser (1990) cyclone model can be partially utilized to describe marine ET (Fig. 3.3). Instead of a small-amplitude disturbance over a broad low-level baroclinic zone, ET deals with a tropical cyclone vortex impinging upon the pre-existing baroclinic zone. Warm frontogenesis occurs when warm, moist air is advected equatorward and ascended over the baroclinic zone. A weak cold front with little convection is created by the poleward advection and descent of cold-dry air around the southwest periphery of the TC. The weak cold front and strong warm front move nearly perpendicular to each other and form a frontal T-bone (II). A frontal fracture appears in the horizontal temperature gradient along the poleward portion of the cold front near the low center first associated with subsidence and frontolysis by Godske et al. (1957). Schultz et al. (1998) note that differential adiabatic warming may weaken the low-level temperature gradient or differential rotation of the isentropes in absence of vertical motion may create frontal fracturing. The baroclinicity in the warm-front region is transported westward relative to the cyclone center forming a bent-back warm front (III). The bent-back warm front wraps around the low center enclosing a pool of relatively warmer air and forms a warm seclusion (IV) (Schultz et al. 1998). A noticeable difference between the Norwegian and Shapiro-Keyser models concerns the maintenance of a nearly perpendicular orientation of the weak cold front and strong warm front as discussed in the latter model. The end result of each observational paradigm is very different in that the Norwegian produces a cold-core occlusion as opposed to the warm-core seclusion of the Shapiro-Keyser lifecycle.

3.6 LC1 and LC2 Baroclinic Lifecycles

A complimentary view of the relationship between cyclone/frontal structure and the upper-level flow concerns the introduction of barotropic shear to the basic-state zonal flow. Thorncroft et al. (1993) (THM) simulates two distinct cyclone lifecycles LC1 (basic state with no barotropic shear) and LC2 (basic state with cyclonic barotropic shear) that roughly parallel the Norwegian and Shapiro-Keyser model lifecycle paradigm. Schultz et al. (1998) mention limitations to applying this barotropic shear model to the real atmosphere yet acknowledge its usefulness to broadly conceptualize upper-level flow and cyclone-frontal structure. This is just one of the many observational and theoretical studies that identify distinct large-scale flow patterns that lead to different modes of cyclone development (Davies et al. 1991; Sinclair and Revell 2000; Schultz et al. 1998). [For a complete analysis of cyclone life cycle characteristics discussed below, the reader is referred to Thorncroft et al. (1993).]

3.61 LC1 and LC2 Types

The LC1 is more comparable to the Norwegian lifecycle with strong temperature gradients in the cold frontal region and bent-back baroclinicity surrounding the low-pressure region (THM). The cold front eventually pinches off the warm sector, which decreases in area reminiscent of a Norwegian occlusion. Eventually, the baroclinicity in the midlatitudes is destroyed and the surface pressure elongates zonally. From a PV point of view, before the occlusion, the wave tilts westward with height in the NW-SE direction on the cyclonic side of the mean jet (THM). A pronounced cyclonic-wrap of the system continues until anticyclonic behavior of the poleward isentropic flow develops. Mean anticyclonic shear untilts the wave to a NE-SW direction and barotropic decay ensues as poleward momentum fluxes are returned to the jet (THM). This process, described as “trough thinning� as cross θ-contour flow destroys the gradient, is analogous to equatorward “Rossby-wave breaking� (McIntyre and Palmer 1984).

The NW-SE tilt of LC2 is stronger than in LC1 due to the effects of stronger cyclonic mean shear. The surface pressure tilts more definitively in the NW-SE direction with the strongest temperature gradients in the warm frontal zone (Hoskins and West 1979). Similarly, a warm-core seclusion occurs as baroclinicity associated with the extended bent-back warm front encircles the low-pressure center. THM notes that surface cyclone is elongated more zonally than the LC1 and confined more meridionally. The strongest pressure gradients exist on the northern and western sides of the storm causing strong winds. However, instead of anticyclonic shear dominating south of the mean jet, cyclonic shear continues to wrap up the wave with uniform θ inside and strong θ-gradients outside the mesoscale vortex. Cyclonic wrap-up on the poleward side of the jet expands to a larger degree than LC1 with the PV-contours near the jet core remaining largely undular.

3.7 Jet Dynamics

Sinclair (2004) uses EOF analysis to identify significant midlatitude circulation patterns associated with ET and assess the impact upon the intensity and structure of the system. He hypothesized that a period of coupling with the divergent quadrant of an upper-level jet is a required condition for extratropical redevelopment of a tropical cyclone. McTaggert-Cowan et al. (2003) also examined the ET of Hurricane Earl (1998) and Hurricane Danielle (1998) and differentiated between two different modes of development based upon the TC remnant’s position relative to the upper-level jet. It is well known that the secondary or vertical circulation causes the typical distribution of convergence and divergence associated with a jet (Fig. 3.4). Anticyclonic (cyclonic) maximum relative vorticity is found on the anticyclonic (cyclonic) side of the jet. A maximum of positive (negative) vorticity advection is found in the left (right) exit and right (left) entrance regions. Along with the attendant ageostrophic circulation, convergence (divergence) is found in the right (left) exit and left (right) entrance regions of the jet.


Figure 3.4: Idealized jet streak model showing convergence associated with low-level horizontal ageostrophic circulation (arrows) and positive (negative) vorticity advection PVA (AVA). Shaded areas show cold (warm) temperature advection in the entrance (exit) regions. (Adapted from McTaggart-Cowan et al. 2003).





3.8 Potential Vorticity (PV) Point of View

The use of potential vorticity provides important insight into the interactions between the tropical cyclone and midlatitude circulation during ET. Agusti-Panareda et al. (2004) model and analyze the ET of Hurricane Irene (1999) and create a PV framework/conceptual model that illustrates many important points. In order for extratropical cyclogenesis to initiate after ET, especially reintensification, the surface low and upper-level trough must tilt westward with height necessary for extraction of potential energy from baroclinic instability. The approach of the tropical cyclone to the midlatitude environment can precondition the upper-level trough and enhance the upper-level jet and thus create a favorable cyclogenetic environment. Three important anomalies associated with the moist convective processes of a TC are critical to the ET process: (1) negative PV anomaly associated with TC upper-level outflow and (2) positive PV anomaly and (3) moisture anomaly co-located with PV tower cyclonic flow and positive thermal anomaly (Fig. 3.5). Emphasis will be placed upon the features of upper-level PV anomalies responsible for the structure and evolution of a post-transition TC extratropical storm.

3.9 Upper-Level Preconditioning of Midlatitude Environment by Transitioning Tropical Cyclone

The TC outflow at upper-levels (5, Fig. 3.5) causes a large-scale negative PV anomaly that can be advected very easily into the extratropical environment where it can be identified as a region of tropopause lift (Agusti-Panareda et al. 2004). In their analysis of Supertyphoon Flo (1990), Merrill and Velden (1996) indicate that PV decreases within the isentropic layers associated with the upper-level outflow of the TC. Following Hoskins et al (1985) and Haynes and McIntyre (1987), the effect of heating due to condensation above the environmental tropopause is to produce a PV source below the heating and a sink above. Due to a decrease of heating with height, tropical cyclones are considered an upper-tropospheric sink of PV (Schubert and Altworth 1987). PV is approximately conserved outside of the convective region for individual parcels so the outflow from tropical cyclones would therefore have relatively low PV values as the effects of the sink are spread by advection (Wu and Emanuel 1994). The effects of vertical wind shear on the TC upper-level warm core discussed in Section 3.3 provided a description from the point of view of trough interactions, which is consistent with this explanation. When this negative PV anomaly is advected into the vicinity of a positive PV anomaly associated with an upper-level trough, it can enhance the PV gradient by steepening the tropopause and hence increasing the intensity of the upper-level jet (Agusti-Panareda et al. 2004). This favorable trough/TC interaction is fundamental to determining if/when/where extratropical cyclogenesis will occur.

As the TC nears an upper-level jet, it enters into a vertical shear environment. The top of the PV tower and co-located moisture tower is advected downwind analogous to the discussion in Section 3.3. At the surface, the PV remains relatively intact allowing for intermittent bursts of convection to be maintained along the warm frontal region. Agusti-Panareda et al. (2004) indicate that the burst of convection generated a new diabatic PV tower comparable to observations of extratropical systems by Rossa et al. (2000). This reinvigorated PV tower maintained and enhanced the negative PV anomaly aloft despite the decay of the tropical cyclone PV tower primarily due to the advection of very moist air poleward near the warm frontal region. As the TC moves poleward, the negative PV anomaly approaches the positive PV anomaly of the upper-level trough and creates a meridionally oriented PV dipole, which strengthened the upper-level jet by steepening the tropopause. As the latent heat release wanes with the decay of the PV tower, the production of negative PV ends and the upper-level negative PV anomaly is advected downstream and sheared, which in turn flattens the tropopause and reduces the upper-level jet strength (Thorncroft and Jones, 2000; Agusti-Panareda et al. 2004).

Rapid cyclogenesis may occur if the upper-level trough untilts sufficiently from the upper-level jet consistent with Petterssen-Smybe type B development (Petterssen and Smybe 1971). Thus for baroclinic instability to develop, the upper-level PV anomaly associated with the trough and the surface thermal anomaly must tilt against the shear. This is essentially a favorable trough interaction described by Hanley et al. (2001) where the surface cyclone lies immediately to the east of the upper-level trough. The upper-level outflow of a TC is highly divergent, which amplifies the upper-level jet and engineers a cyclogenetic feedback process.

Figure 3.5: (Agusti-Panareda et al. 2004) Vertical cross section schematic with potential vorticity (PV) anomalies and other anomalies associated with extratropical transition: (1) surface thermal anomaly on baroclinic zone, (2) diabatically-generated positive PV anomalies along the baroclinic zone, (3) positive PV anomaly associated with a midlatitude upper-level trough, (4) TC positive PV tower, (5) negative PV anomaly associated with the tropical-cyclone’s outflow. The funny arrow represents the upper-level jet, which is modulated by the horizontal gradient of PV at upper-levels, i.e. the steepness of the tropopause.

3.10 Cyclone Phase Space Diagnostics

The cyclone phase space (CPS, Hart 2003) is a three-dimensional continuum that describes the frontal and thermal structure of synoptic scale systems. Symmetry is defined in the phase diagnostics by the difference in 900-600 hPa thickness across the storm relative to storm motion quantified by


where a thickness difference of 10 m was determined empirically to define the onset of ET (Evans and Hart 2003). ET completion is characterized by the first point in time after the storm has become both asymmetric and cold-cored; the sign of the thermal wind in the 900-600 hPa layer defines the core structure, where a negative slope indicates a warm-cored system and a positive slope indicates a cold core (Evans and Hart 2003). Determination of the thermal structure of the storm is related to the vertical structure of the cyclone’s height perturbation, which follows from the hypsometric relationship discussed more fully by Hirschberg and Fritsch (1993). If the magnitude of the cyclone isobaric height gradient increases (decreases) with height above the surface, by thermal wind relations, the cyclone is defined to be cold (warm) core. To differentiate between tropical and extratropical systems in the continuum of possibilities, two layers of equal mass (900-600 hPa and 600-300 hPa) are used to calculate the height perturbation and thus the thermal wind. The strength and depth of the thermal structure is easily estimated from the CPS. See example below of the Ocean Ranger storm (Fig. 3.6).

Figure 3.6: Cyclone Phase Space (Hart 2003; Evans and Hart 2003) for the Ocean Ranger storm of 1982. The left plot illustrates the frontal asymmetry (B) and thermal structure (-VTL) in the 900-600 hPa layer. The right plot corresponds to the overall thermal structure of the storm from deep cold core to moderate warm core and back again. Figure obtained from

3.11 Précis

The observational and theoretical cyclone paradigms described above are antipode discrete examples of a vast variety of cyclone structures and evolutions. The differences between the Norwegian lifecycle and the Shapiro-Keyser cyclone evolution are largely a function of observation location. Bjerknes and Solberg (1922) observed cyclones in the diffluent region of the upper-level storm-track over Western Europe. Thus, their conclusions relied heavily upon understanding the circulations inherent with primarily low zonal index flows and often associated blocking patterns. The Experiment on Rapidly Intensifying Cyclones over the Atlantic (ERICA, December 1988 – January 1989) examined many cyclones with various frontal structures and evolutionary characteristics. The Shapiro-Keyser (1990) model resulted from a perceived inadequacy in the Norwegian model to explain the observed patterns of cyclone lifecycles explained in Section 3.4. The theoretical paradigms of Thorncroft et al. (1993) model antipode cyclone lifecycles also based upon upper-level flow and found many similarities between the previously mentioned paradigms. However, as a collection, all of them still fail to define the vast continuum of cyclone lifecycle evolutions. Extratropical transition is a primary example of where the paradigms fail for a couple major reasons to be addressed in the forthcoming chapters. Primarily, the presence of a powerful tropical cyclone vortex impinging upon the baroclinic zone and its ability to precondition the midlatitude environment into which it is moving is not described in either the observational or theoretical cyclone paradigms. The Ocean Ranger storm (Fig. 3.6) is characteristic of a strong wintertime Nor’easter yet has a very similar lifecycle to the post-transition TC: a warm seclusion. Conversely, the final outcome of the extratropical transition of Floyd (1999) is a cold-core occlusion reminiscent of the Norwegian lifecycle.

Note: Schultz et al. (1998) studied the effects of different upper-level flow upon low-level frontal structure, which highlight the differences between Norwegian and Shapiro-Keyser type systems. Idealized simulations using a nondivergent barotropic model illustrate the different lifecycles to a good approximation.