GARY M. LACKMANN,* DANIEL KEYSER, AND LANCE F. BOSART

Transcription

1 VOLUME 125 MONTHLY WEATHER REVIEW NOVEMBER 1997 A Characteristic Life Cycle of Upper-Tropospheric Cyclogenetic Precursors during the Experiment on Rapidly Intensifying Cyclones over the Atlantic (ERICA) GARY M. LACKMANN,* DANIEL KEYSER, AND LANCE F. BOSART Department of Earth and Atmospheric Sciences, State University of New York at Albany, Albany, New York (Manuscript received 12 August 1996, in final form 19 March 1997) ABSTRACT This paper documents a characteristic life cycle of upper-tropospheric precursors to surface cyclogenesis observed during the field phase of the Experiment on Rapidly Intensifying Cyclones over the Atlantic (ERICA, December 1988 February 1989). This life cycle begins with the development of an elongated region of lower dynamic tropopause that forms in association with an intensifying midtropospheric jet/front over central North America. The elongated disturbance subsequently compacts into a more circular configuration prior to crossing the east coast of North America and frequently is associated with rapid surface cyclogenesis offshore. A representative example of the life cycle outlined above is documented through a detailed case study of the upper precursor associated with the second ERICA intensive observation period (IOP 2) cyclone. Emphasis is placed upon (i) description of the tropopause structure and evolution during the upper-precursor life cycle, (ii) diagnosis of mechanisms leading to the development and intensification of a midtropospheric cyclonic vorticity maximum and frontal zone, (iii) analysis of the role of transverse jet front circulations in deforming the dynamic tropopause, (iv) documentation of the influence of the low- and high-frequency flow components on the upperprecursor life cycle, and (v) isolation of dynamic and thermodynamic factors that render this life cycle especially conducive to rapid surface cyclogenesis. Confluence downstream of the axis of a low-frequency (i.e., periods greater than 120 h), troposphere-deep ridge over western North America facilitates the organization of a midtropospheric jet/front over central North America. As this precursor disturbance approaches the inflection between the western ridge and a downstream trough, tilting, in the presence of cold advection along the midtropospheric frontal zone, becomes an important vorticity generation and frontogenesis mechanism in the upper precursor. Transverse circulations accompanying the jet/front steepen and lower the dynamic tropopause prior to surface cyclogenesis. Compaction of the initially elongated upper precursor is shown to involve deformation in the highfrequency component of the upper-tropospheric flow. The compacted upper-precursor configuration, lowered tropopause, and reduced static stability in the offshore environment lead to strong vertical coupling and vigorous surface cyclogenesis as the upper precursor passes offshore. The foregoing results suggest that the life cycle of a common class of cyclogenetic precursors is closely related to midtropospheric frontogenesis. A favored location for the development of midtropospheric jet/fronts is over central North America during northwesterly flow episodes. Production of vorticity in the midtropospheric jet/front and subsequent compaction of this vorticity feature suggest a link between midtropospheric frontogenesis and mobile upper-trough genesis. This link may explain the existence of a maximum in the upper-trough-genesis distribution over central North America documented by Sanders. 1. Introduction Since the advent of routine upper-air observations in the 1940s and 1950s, it has become evident that extratropical cyclogenesis usually involves an upper-tropospheric precursor disturbance in the form of a mobile upper trough or jet streak (e.g., Petterssen 1956, section 16.7; Palmén and Newton 1969, section 11.2). Although * Current affiliation: Department of the Earth Sciences, State University of New York College at Brockport, Brockport, New York. Corresponding author address: Prof. Gary M. Lackmann, Dept. of the Earth Sciences, SUNY College at Brockport, 350 New Campus Dr., Brockport, NY garyl@weather.brockport.edu a large number of studies have documented the role of such precursor disturbances during cyclogenesis, few studies have focused on the development and evolution of the upper-precursor disturbances themselves. The main objectives of this paper are to (i) document a characteristic life cycle of upper-tropospheric precursor disturbances, (ii) describe dynamical mechanisms that are important to their evolution, (iii) investigate the effect of the low-frequency flow on this life cycle, and (iv) identify aspects of this life cycle that are especially conducive to rapid surface cyclogenesis. The linkage between the upper-precursor life cycle and the low-frequency flow may explain the presence of a maximum in the upper-trough-genesis distribution over and east of the Rocky Mountains documented by Sanders (1988). In this paper, we use the general term upper disturbance to refer to both mobile upper troughs and jet 1997 American Meteorological Society 2729

2 2730 MONTHLY WEATHER REVIEW VOLUME 125 streaks. In those instances where an upper disturbance is known to play a major role in the dynamics of surface cyclogenesis, we apply the terms upper-tropospheric cyclogenetic precursor or upper precursor. Upper disturbances typically have wavelengths 1 of km, periods of days, and translational speeds of m s 1 (e.g., Sanders 1986, 1988; Lackmann 1995, section 4.3; Lackmann et al. 1996). These high-frequency features do not contain significant variance on timescales within the day range often used to encompass synoptic-scale disturbances (e.g., Blackmon et al. 1984). Several previous studies have identified mobile upper troughs using either curvature of geopotential height contours or isobaric vorticity, usually at the 500- hpa level (Sanders 1988; Lefevre and Nielsen-Gammon 1995; Dean and Bosart 1996). Mobile upper troughs attain maximum amplitude at the tropopause (see, e.g., Sanders 1988, section 5c) and are accompanied by significant downward extrusions of the dynamic tropopause (e.g., Hoskins et al. 1985, section 6e; Hoskins and Berrisford 1988), where the dynamic tropopause is defined in terms of potential vorticity (PV). Therefore, we define mobile upper troughs as local maxima of pressure and minima of potential temperature on the dynamic tropopause (p trop and trop, respectively). Significant downward extrusions of the tropopause invariably are associated with vorticity maxima at 500 hpa; thus our definition of upper troughs is consistent with those based on isobaric vorticity. Jet streaks (Bluestein 1993, section 2.8) appear as regions of steeply sloping dynamic tropopause (e.g., Shapiro 1981; Uccellini et al. 1985). Because both mobile upper troughs and jet streaks are identifiable on maps of p trop and trop, we utilize such maps to locate and track upper disturbances. Classic theoretical treatments of baroclinic instability (e.g., Charney 1947; Eady 1949) are not concerned with finite-amplitude cyclogenetic precursor disturbances (see, e.g., Palmén 1951). Initial-value treatments of the baroclinic development problem (Farrell 1982, 1984, 1985; Rotunno and Fantini 1989), although taking a significant step toward observations, do not attempt to account for the origin or initial evolution of cyclogenetic precursors. Farrell (1989a) isolates optimally growing perturbations in order to describe the initial growth of cyclone-scale disturbances. This approach yields a set of disturbances that initially exhibit a pronounced vertical phase tilt; nevertheless, a systematic comparison of these disturbances with observed cyclogenetic precursors has not, to our knowledge, been undertaken. A mechanism referred to as superposition can account for the intensification of cyclone-scale disturbances. This mechanism, considered theoretically by Farrell (1989b) 1 The use of terminology pertaining to wavelike disturbances does not necessarily imply that mobile upper troughs and jet streaks are spatially periodic. In fact, many of the upper disturbances examined in this study appear to be spatially localized and isolated. and observationally by Nielsen-Gammon (1995) and Nielsen-Gammon and Lefevre (1996), explains how perturbation energy can increase in response to changes in the proximity of two separate PV anomalies or to changes in the shape of a single PV anomaly. Additional theories that have been proposed to describe the development of disturbances with wavelengths in the km range include modal baroclinic instability (e.g., Blumen 1979, 1980; Weng and Barcilon 1987; Whitaker and Barcilon 1992a,b) and modal and local barotropic instability (e.g., Kuo 1949; Mak and Cai 1989). Downstream baroclinic development (e.g., Namias and Clapp 1944; Rossby 1945; Hovmöller 1949; Orlanski and Katzfey 1991; Orlanski and Chang 1993; Chang 1993; Orlanski and Sheldon 1993, 1995) can explain the formation of synoptic-scale troughs through the downstream propagation of extant Rossby wave energy. Nielsen-Gammon (1995) and Nielsen-Gammon and Lefevre (1996) present evidence for the role of downstream development in a trough-genesis event over North America. Nevertheless, it is not clear that this mechanism can account for the generation of smallerscale mobile troughs because of the similarity of the Rossby wave group and phase velocities for the km range of wavelengths presently under consideration. The relevance of downstream development to the behavior of smaller-scale mobile troughs also is questioned by Bosart et al. (1996, section 6b) in the context of the Superstorm 93 cyclone event over eastern North America. Uccellini et al. (1984), Uccellini et al. (1985), Uccellini (1986), and Whitaker et al. (1988) have demonstrated the role of upper-tropospheric jet streaks in rapid maritime cyclogenesis. Specifically, Uccellini et al. (1985) and Whitaker et al. (1988) use observational data and numerical simulations, respectively, to document the role of ageostrophic circulations in the development of an upper precursor prior to the rapid amplification of the Presidents Day cyclone of They show that jet front circulations contribute to the downward extrusion of air possessing stratospheric values of PV in advance of explosive cyclogenesis. Based on idealized simulations of vortices superimposed on a baroclinic westerly basic-state jet stream, Takayabu (1991) presents evidence that the thermally direct transverse ageostrophic circulation in a confluent jet entrance region acts to steepen and lower the tropopause. As discussed by Uccellini et al. (1985, section 5a), there have been numerous studies documenting the downward extrusion of stratospheric air during midtropospheric frontogenesis and cyclogenesis. Nevertheless, as noted by Uccellini et al. (1985) and Keyser and Shapiro (1986, section 5), additional research is needed concerning the link between midtropospheric frontogenesis and cyclogenesis. In a comprehensive observational study of cyclogenetic precursors, Sanders (1988) shows that mobile upper troughs form preferentially downstream of major

3 NOVEMBER 1997 LACKMANN ET AL TABLE 1. Summary of 18 upper disturbances during the ERICA field phase (December 1988 through February 1989). This 18-case sample represents all instances of 500-hPa relative vorticity maxima of at least s 1 in the ECMWF analysis, indicating the presence of upper disturbances, that pass into the ERICA domain (defined as the region bounded by 30 N and 50 N, and 60 W and the east coast of North America). The first column lists the time and date of coastal crossing of the 500-hPa relative vorticity maxima for each of the 18 cases, followed in the second column by the corresponding ERICA intensive observation period (IOP), limited observation period (LOP), or special observation period (SOP) number, where applicable. The third column indicates whether surface cyclogenesis (defined by the appearance of at least two closed isobars in the sea level pressure field contoured at a 4-hPa interval in the ECMWF analysis) occurs within 24 h of coastal crossing of the 500-hPa relative vorticity maximum. The fourth column provides a subjective classification describing the life cycle of the upper disturbance; see section 3 and Lackmann (1995, section 4.9) for details on classification terminology. In addition to the terminology introduced in section 3, we include the categories northern stream, southern stream, and noncompaction. The former two apply to upper disturbances that are isolated north and south of the polar-front jet stream, respectively; the latter applies to a particular event (time of coastal crossing: 1800 UTC 18 January 1989) in which the upper disturbance is elongated initially but does not become more isotropic with time. Time/date of coastal crossing ERICA No. Cyclogenesis Classification 0000 UTC 2 December 1988 No Merger/compaction 1800 UTC 4 December 1988 No Merger/compaction 1200 UTC 12 December 1988 Yes Northern stream 0600 UTC 13 December 1988 Yes Fracture/compaction 0000 UTC 14 December 1988 IOP 2 Yes Compaction 1800 UTC 17 December 1988 IOP 3 Yes Compaction 0600 UTC 4 January 1989 IOP 4 Yes Compaction 1800 UTC 13 January 1989 SOP 4B No Fracture 1800 UTC 18 January 1989 No Noncompaction 1200 UTC 19 January 1989 IOP 5 Yes Compaction 0000 UTC 21 January 1989 LOP 5A Yes Compaction 0600 UTC 23 January 1989 Yes Southern stream 0000 UTC 28 January 1989 LOP 6P No Compaction 1200 UTC 31 January 1989 No Merger 1800 UTC 10 February 1989 No Compaction 0000 UTC 13 February 1989 IOP 7 Yes Compaction 1200 UTC 23 February 1989 IOP 8 Yes Merger 0000 UTC 27 February 1989 Yes Compaction orographic barriers, such as the Rocky Mountains and the Himalayas, in the presence of a northerly component to the midtropospheric flow. More recently, the distribution of trough genesis obtained by Sanders (1988) has been reproduced using objective numerical procedures (Lefevre and Nielsen-Gammon 1995; Dean and Bosart 1996). Many earlier studies have emphasized that midtropospheric frontogenesis is favored in northwesterly flow (e.g., Reed and Sanders 1953; Reed 1955; Bosart 1970; Shapiro 1981; Sanders et al. 1991); these studies highlight the importance of frontogenetical tilting. Shapiro (1981) hypothesized that alongfront cold advection contributes to shifting the direct transverse ageostrophic circulation in a confluent jet entrance region to the anticyclonic-shear side of the jet axis such that subsidence is focused within, or on the warm side of, the midtropospheric baroclinic zone. Numerical studies by Keyser and Pecnick (1985a,b) and Reeder and Keyser (1988) confirm this hypothesis. The possibility exists of a positive feedback arising between this frontogenetical configuration of subsidence and the developing jet/front in its confluent entrance region; such a feedback is referred to by Rotunno et al. (1994) as the Shapiro effect. An investigation of the life cycles of 18 upper disturbances over North America during the field phase of the Experiment on Rapidly Intensifying Cyclones over the Atlantic [ERICA, December 1988 through February 1989; see Hadlock and Kreitzberg (1988)] by Lackmann (1995, chapters 4 and 5) reveals that substantial structural modifications occur as upper disturbances interact with larger-scale waves. Specifically, 12 of the 18 upper disturbances (summarized in Table 1) undergo an increase in isotropy during their evolution prior to coastal crossing (i.e., the length of the major axis of the disturbance decreases relative to that of the minor axis), a process referred to here as compaction. Upper disturbances typically take on the character of jet streaks as they traverse larger-scale ridges and are manifested as increasingly distinct troughs in the geopotential height field as they approach the axes of larger-scale troughs and experience compaction. This structural transformation has important ramifications for understanding mobile-trough genesis, which becomes, in part, a question of explaining how upper disturbances behave in the presence of larger-scale waves. The presence of planetary-scale flow anomalies during the formation of upper-tropospheric precursors to western North Atlantic cyclogenesis is documented in the composite study of Lackmann et al. (1996); their analysis reveals a troposphere-deep ridge over western North America throughout the composite evolution. The composite upper precursor becomes organized in association with a jet streak in a region of northwesterly flow downstream of the western North American ridge. The investigation of upper-precursor life cycles by Lackmann (1995, chapters 4 and 5) suggests that com-

4 2732 MONTHLY WEATHER REVIEW VOLUME 125 paction of an initially elongated upper disturbance in northwesterly flow over the North American continent may commonly precede explosive cyclogenesis over the western North Atlantic Ocean. Accordingly, the overall objective of this paper is to document the characteristic life cycle of compacting upper disturbances through a detailed case study of the upper precursor associated with the second ERICA intensive observation period (IOP 2) cyclone. Specific goals of this study are to (i) document the tropopause structure and evolution during the IOP 2 upper-precursor life cycle, (ii) determine the mechanisms responsible for the intensification of a cyclonic vorticity maximum and midtropospheric frontal zone associated with the IOP 2 upper precursor, (iii) examine the role of transverse jet front circulations in deforming the dynamic tropopause during this event, (iv) assess the influence of the low- and high-frequency components of the flow on the life cycle of this precursor disturbance, and (v) isolate the characteristics of the IOP 2 upper precursor that lead to explosive surface cyclogenesis. These goals represent extensions of the work of Uccellini et al. (1985) and Sanders (1988) in the following sense: the preference for mobile upper-trough formation in northwesterly flow over central North America found by Sanders (1988) is hypothesized to be related to the fact that this location is conducive to midtropospheric frontogenesis of the type documented by Uccellini et al. (1985). The format of the remainder of this paper is as follows. First a discussion of the dataset and methodology is presented in section 2. A summary of upper-precursor life cycles during ERICA follows in section 3. In section 4, the ERICA IOP 2 event, a representative example of the characteristic life cycle outlined in section 3, is diagnosed in greater detail, with emphasis on vorticity generation mechanisms, frontogenesis, jet front circulations, and the influence of the larger-scale flow on the precursor life cycle. A concluding discussion is given in section Data and methodology The sources for gridded data used in this study are global analyses generated at the European Centre for Medium-Range Weather Forecasts (ECMWF) and obtained from the National Center for Atmospheric Research (NCAR). A description of the data assimilation procedure used by ECMWF to develop these analyses is presented by Bengtsson et al. (1982), and modifications to this procedure are discussed by Trenberth and Olson (1988) and Trenberth (1992). The data are stored in spherical harmonic form with a horizontal resolution of in latitude and longitude and with 6-h time resolution, although the 0600 UTC and 1800 UTC analyses are derived largely from the short-term model forecast. The dataset comprises the 14 mandatory isobaric levels from 1000 to 10 hpa. The ECMWF grids are bilinearly interpolated onto a grid with 1 latitude longtitude resolution and are stored and displayed using the General Meteorological Package (GEMPAK; Koch et al. 1983). A concise depiction of midlatitude cyclogenesis is obtained through maps of trop, along with maps of surface potential temperature or surface equivalent potential temperature ( esfc ). This approach embodies an interpretation of cyclone dynamics that is consistent with the Eady model of baroclinic development, namely, that the relevant dynamics can be accounted for using maps of the thermal field at the upper and lower boundaries (i.e., the dynamic tropopause and the surface of the earth, respectively). An important distinction between these displays and the classic Eady problem is that unlike in the idealized model, the observed dynamic tropopause is deformable (i.e., a function of position and time). An aspect of cyclogenesis that is not included explicitly in this approach is diabatically generated PV features within the troposphere, such as those typically found in warm-frontal regions of maritime cyclones (e.g., Davis and Emanuel 1991; Davis 1992; Reed et al. 1992; Stoelinga 1996). Nevertheless, diabatic processes at the tropopause and at the surface of the earth are manifested as nonconservation of potential temperature on these respective surfaces. Furthermore, diabatically induced modifications to the PV field at the level of the tropopause may result in nonconservation of potential temperature on the tropopause by changing the altitude of this surface. For additional discussion of the dynamic tropopause and for examples of tropopause maps from other cases, refer to, for example, Danielsen (1968), Hoskins and Berrisford (1988), Hakim et al. (1995), and Bosart and Lackmann (1995). The choice of an appropriate value of PV to define the dynamic tropopause has been discussed by Danielsen and Hipskind (1980). They point out that values ranging from 1 to 2 PV units [PVU, defined as 10 6 m 2 s 1 Kkg 1 following Hoskins et al. (1985, section 2a)] can be found in the literature. Here we choose the 1.5 PVU surface to represent the dynamic tropopause. Inspection by the first and third authors of many cross sections during various synoptic situations indicates that interpretations based on tropopause maps are not highly sensitive to this choice. The tropopause is located through a downward search of PV as a function of pressure starting from the lower stratosphere for the first occurrence of PV values less than 1.5 PVU. In those regions where the tropopause is multivalued, the possibility exists either of a tropopause fold or of isolated regions of lower-tropospheric PV in excess of 1.5 PVU (e.g., in warm-frontal regions or in shallow, stable arctic air masses). The adoption of a downward search procedure avoids the need to distinguish objectively between tropopause folds and lower-tropospheric PV maxima; in the present study, these features are considered to be part of the troposphere. Once the pressure of the tropopause is determined, potential temperature, horizontal vector wind, and pressure-coordinate vertical ve-

5 NOVEMBER 1997 LACKMANN ET AL locity are linearly interpolated with respect to pressure to this surface; the latter two quantities will be denoted as V trop and trop, respectively. The degree to which flow induced by tropopausebased disturbances penetrates to the surface of the earth is hypothesized to be directly proportional to the ratio H R /H trop, where H R is the Rossby depth and H trop is the depth of the troposphere (Hoskins et al. 1985, section 3). At the level of quasigeostrophic theory, H R f 0 L/N, where f 0 is the Coriolis parameter, L is a length scale representative of the upper disturbance, and N is the Brunt Väisälä frequency. If L and N are constant, then H R is constant and H R /H trop 1/H trop. Therefore, the ability of flow induced by tropopause-based disturbances to reach the surface of the earth is anticipated to be inversely proportional to the depth of the troposphere, which in the absence of topography reduces to the tropopause height. Because of the close correspondence between tropopause height and p trop, we can infer the ability of upper disturbances to interact with surface disturbances by examining maps of p trop. In addition, p trop also reveals changes in the structure and the intensity of tropopause-based disturbances. Regarding the latter, if p trop is increasing (decreasing) at the center of an upper trough, we infer that it is intensifying (weakening) even if trop remains relatively constant. An example of this behavior is provided in section 4a. 3. Summary of upper-precursor life cycles during ERICA We begin by reviewing briefly the results of an observational study of 18 upper disturbances during the ERICA field phase undertaken by Lackmann (1995, chapters 4 and 5) and summarized here in Table 1. Despite a variety of documented disturbance evolutions, certain characteristic behaviors are evident. Examples include (i) initially elongated upper disturbances that develop in northwesterly flow and subsequently become more isotropic (i.e., compaction); (ii) the formation of an upper disturbance through the detachment of the equatorward extremity of a trough from the stratospheric reservoir of PV associated with the circumpolar vortex (i.e., trough fracture); and (iii) the amalgamation or phasing of two or more upper disturbances, often involving surface cyclogenesis (i.e., trough merger). As disclosed in section 1, scenario (i) is the most common of the above evolutions, having occurred to varying degrees in 12 of the 18 cases (Table 1), and, as such, will be the focus of investigation in this paper. This scenario is characterized by the formation or intensification of an elongated upper disturbance in association with a midtropospheric jet/front, manifested as an elongated tongue of higher p trop (i.e., lower-tropopause height) forming east of the axis of a troposphere-deep ridge located over western North America or over the Gulf of Alaska. In many instances, the elongated upper disturbance undergoes compaction prior to crossing the east coast of North America, with rapid surface cyclogenesis following coastal crossing. Diffluent flow typically is present immediately downstream of the elongated upper disturbance when compaction is occurring. During compaction, which occurs when the upper disturbance is well upstream of the surface cyclogenesis region, the tropopause steadily lowers and the absolute vorticity increases on midtropospheric isobaric surfaces. In several instances of this scenario, there is a separate disturbance downstream of the compacting disturbance; this leading disturbance is referred to as a predecessor trough or predecessor disturbance. To expose similarities between cases and to provide an overview of the tropopause structure during events of the type described above, we now summarize three examples from the ERICA period: the IOP 2, IOP 3, and limited observation period (LOP) 5A events (Table 1). Each of these examples exhibits a well-defined upper precursor that undergoes compaction and that is associated with an intense maritime cyclone after passing into the ERICA domain. Figure 1 shows p trop at 12-h intervals from 1200 UTC 12 December through 0000 UTC 14 December 1988, a period corresponding to the initial development of the upper precursor associated with the powerful IOP 2 cyclone. At 1200 UTC 12 December, the IOP 2 upper precursor is evident as a region of locally higher p trop (i.e., lower tropopause height) oriented approximately from northwest to southeast across southwestern Canada (Fig. 1a). An expansive area of lower p trop (i.e., higher tropopause height) and anticyclonic flow is evident over the western United States. A predecessor trough is centered over southwestern Arkansas at this time (Fig. 1a). By 0000 UTC 13 December, the IOP 2 precursor has moved southeastward and extends in an elongated band from northern Alberta to Iowa, while the ridge persists over the western United States (Fig. 1b). There is a strong jet located in the region of steeply sloping tropopause between the IOP 2 precursor and the western ridge, with V trop in excess of 80 m s 1 between southern Saskatchewan and northwestern South Dakota. The predecessor trough now is an elongated band extending from the Florida panhandle to Virginia. By 1200 UTC 13 December, p trop in the IOP 2 upper precursor has increased to over 450 hpa, and the gradient of p trop has further increased between the center of the disturbance and a region of p trop less than 150 hpa that is centered over Mississippi (Fig. 1c). The predecessor trough has dissipated rapidly between 0000 UTC and 1200 UTC 13 December (Figs. 1b,c). By 0000 UTC 14 December, the IOP 2 upper precursor is characterized by a nearly circular region of p trop in excess of 500 hpa centered immediately southeast of Cape Hatteras, North Carolina (Fig. 1d). A tropopause fold exists at this time imme-

6 2734 MONTHLY WEATHER REVIEW VOLUME 125 FIG. 1. Pressure (solid, contour interval 50 hpa, shaded as indicated in legend at bottom of panel) and wind (every fourth grid point plotted using standard convention) on the dynamic tropopause (defined as the 1.5-PVU surface) for IOP 2 upper-precursor event: (a) 1200 UTC 12 December 1988, (b) 0000 UTC 13 December 1988, (c) 1200 UTC 13 December 1988, (d) 0000 UTC 14 December diately south and southeast of the upper precursor (not shown). The life cycle of the upper precursor associated with the IOP 3 cyclone is marked by a sequence of events similar to the IOP 2 case. At 0000 UTC 16 December, the IOP 3 upper precursor appears as an elongated band of higher p trop extending from the Northwest Territories into Manitoba (Fig. 2a). A complex predecessor disturbance, which consists of three separate maxima in the p trop field, lies nearly perpendicular to the major axis of the IOP 3 upper precursor. This predecessor disturbance extends from a cutoff cyclone immediately west of California, across the continental United States, and northeastward to the Maritime Provinces. A region of lower p trop and associated anticyclonic flow is present over western Canada and the adjacent eastern North Pacific Ocean. The IOP 3 precursor advances southeastward into northern Minnesota by 1200 UTC 16 December as the trailing portion of the predecessor trough over Kentucky weakens (Fig. 2b). The IOP 3 precursor is marked by lowering of the tropopause over southern Minnesota and northern Iowa at 0000 UTC 17 December, with values of p trop increasing to over 450 hpa (Fig. 2c). The large gradient of p trop along the southwestern edge of the precursor disturbance coincides with ECMWF-analyzed V trop of over 70 m s 1. The portion of the predecessor trough located southeast of the IOP 3 upper precursor has weakened dramatically by 0000 UTC 17 December and is barely detectable as a region of higher p trop over eastern North Carolina, while the anticyclonic region of lower p trop remains over western Canada. By 1200 UTC 17 December, the IOP 3 upper precursor is nearing the East Coast, and p trop in the disturbance center has increased to over 500 hpa in an elliptical region centered over southern West Virginia (Fig. 2d). The upper trough associated with the ERICA LOP

7 NOVEMBER 1997 LACKMANN ET AL FIG. 2. As in Fig. 1 except for ERICA IOP 3 event: (a) 0000 UTC 16 December 1988, (b) 1200 UTC 16 December 1988, (c) 0000 UTC 17 December 1988, (d) 1200 UTC 17 December A cyclone follows an evolution similar to that in the previous two examples, with the exception that a dissipating predecessor disturbance is absent in the present case. 2 At 0000 UTC 19 January 1989, there is a wellorganized upper disturbance nearing the East Coast that is associated with the IOP 5 cyclone (Fig. 3a). Meanwhile, the LOP 5A upper precursor is becoming organized over central British Columbia, Alberta, and Saskatchewan. This region of higher p trop extends into northeastern Montana, North Dakota, and northern Minnesota by 1200 UTC 19 January (Fig. 3b). A region of lower p trop and anticyclonic flow is evident over western Canada and the Pacific Northwest at this time. Values of 2 Although we do not address the significance of the absent predecessor disturbance in relation to the tropopause evolution and surface cyclogenesis for the LOP 5A event, we speculate that the presence of a predecessor disturbance downstream of a precursor disturbance may facilitate the compaction of the precursor. The possible role of the predecessor in the compaction process will be considered for the IOP 2 event in section 4e. p trop in the center of the LOP 5A upper precursor have increased to over 500 hpa by 0000 UTC 20 January, at which time the disturbance is characterized by a band of lower tropopause and strong p trop gradient extending northward into central Canada from a local maximum in p trop over northern Iowa (Fig. 3c). This maximum advances eastward during the subsequent 12 h (Fig. 3d) and eventually is associated with surface cyclogenesis on 21 January (not shown). Each of the cases presented above exhibits (i) an initially elongated region of higher p trop corresponding to the upper precursor, (ii) a strong p trop gradient and associated jet streak along the southern or southwestern flank of the upper precursor, (iii) steadily rising p trop in the disturbance center as it traverses North America, (iv) an increasingly circular region of locally higher p trop near the eastern edge of the upper precursor, and (v) a region of lower p trop and anticyclonic flow centered west of the developing upper precursor. Each upper precursor is also associated with rapid surface cyclogenesis after crossing the east coast of North America (not shown).

8 2736 MONTHLY WEATHER REVIEW VOLUME 125 FIG. 3. As in Fig. 1 except for ERICA LOP 5A event: (a) 0000 UTC 19 January 1989, (b) 1200 UTC 19 January 1989, (c) 0000 UTC 20 January 1989, (d) 1200 UTC 20 January The similarities between these three cases and between these cases and the other compaction events listed in Table 1 suggest that common mechanisms are at work in this life cycle. In the following section, we examine such mechanisms for a representative example of this life cycle the ERICA IOP 2 event. This event is selected for diagnosis because it presents the clearest signature of the compacting upper-precursor life cycle from the 18-case sample summarized in Table ERICA IOP 2: A representative example Following a detailed documentation of the evolution of upper- and lower-tropospheric disturbances during the IOP 2 event in section 4a, we will address specific questions that arise from the foregoing discussion of Figs These include (i) what is the signature of the upper-precursor evolution in terms of midtropospheric vorticity, and what mechanisms are responsible for midtropospheric vorticity generation? (ii) What processes are responsible for the initiation and intensification of the midtropospheric frontal zone? (iii) What role do transverse jet front circulations play in organizing the initial disturbance structure? (iv) What are the relative contributions of deformation associated with the lowand high-frequency components of the flow to the formation and compaction of the upper precursor? (v) What aspects of this life cycle render it particularly conducive to surface cyclogenesis? We apply vorticity and frontogenesis budgets in sections 4b and 4c, respectively, to examine mechanisms leading to the initial amplification of the precursor disturbance. In section 4d, the issue of the initial precursor organization will be addressed through diagnosis of the divergent and total flows transverse to the upper precursor and their role in deforming the dynamic tropopause. Digitally filtered flow fields are analyzed in section 4e to assess the degree to which the formation of the upper precursor and its subsequent compaction are related to deformation in the low- and high-frequency components of the flow. Consideration of what makes this life cycle particularly con-

9 NOVEMBER 1997 LACKMANN ET AL FIG. 4. Summary plot for IOP 2 event: (a) potential temperature (solid, contour interval 10 K, shaded as indicated in legend at bottom of panel) and wind (every fourth grid point plotted using standard convention) on the dynamic tropopause, and surface equivalent potential temperature (dashed, contour interval 10 K, 310-K isentrope denoted by thick dashed lines) at 0000 UTC 12 December 1988; (b) pressure on the dynamic tropopause (solid, contour interval 50 hpa, shaded as indicated in legend at bottom of panel) and sea level pressure (dashed, contour interval 4 hpa) at 0000 UTC 12 December 1988; (c) as in (a) except for 1200 UTC 12 December 1988; (d) as in (b) except for 1200 UTC 12 December ducive to surface cyclogenesis will be deferred to section 5. a. The life cycle of the ERICA IOP 2 upper precursor We now provide a more complete summary of the life cycle of the IOP 2 upper precursor than is presented in Fig. 1. Detailed analyses of the ERICA IOP 2 cyclogenesis event may be found elsewhere (e.g., Sanders 1990; Roebber et al. 1994). The IOP 2 life cycle is presented using two maps, the first exhibiting trop, V trop, and esfc ; the second, p trop and sea level pressure (SLP). At 0000 UTC 12 December 1988, a jet with V trop approaching 75 m s 1 is located over central British Columbia (Fig. 4a). This jet is oriented parallel to the southern flank of a region of lower trop and higher p trop, which extends from northern Alberta westward to the Gulf of Alaska (Figs. 4a,b). This region marks the location of the incipient upper precursor for the IOP 2 event. Two predecessor disturbances are located over eastern North America; the northern one is centered over eastern Ontario and Quebec, and the southern one is centered over Kansas. A broad region of higher trop, lower p trop, and accompanying anticyclonic flow is evident over the eastern Pacific and western United States (Figs. 4a,b). The SLP field exhibits a lee-trough signature immediately east of the Rocky Mountains over the northern plains of the United States (Fig. 4b). At 1200 UTC 12 December, the downstream edge of the aforementioned jet has advanced southeastward to a position over northern Montana and North Dakota (Fig. 4c), and the region of lower trop and higher p trop over southwestern Canada has moved southeastward as well (Figs. 4c,d). The region of strongest flow coincides with large gradients of trop and p trop over southern Can-

10 2738 MONTHLY WEATHER REVIEW VOLUME 125 FIG. 5. As in Fig. 4 except for 0000 UTC 13 December 1988 [(a) and (b)] and 1200 UTC 13 December 1988 [(c) and (d)]. ada and the north-central United States. The southern predecessor disturbance is now centered over Arkansas (Figs. 4c,d). The northern portion of the SLP trough that had been located along the lee slopes of the Rocky Mountains 12 h earlier has moved eastward to the Dakotas (Fig. 4d). The IOP 2 upper precursor has developed into an elongated band of lower trop and higher p trop extending from northern Alberta into Iowa by 0000 UTC 13 December (Figs. 5a,b). The southern predecessor trough is thinning and is oriented southwest northeast at this time. There are no strong gradients of esfc near the IOP 2 upper precursor, although an area of enhanced esfc gradient lies east of the southern predecessor disturbance along the Georgia coast (Fig. 5a). The trough in the SLP field that originally developed along the lee slopes of the Rocky Mountains is now located immediately east of the leading edge of the IOP 2 upper precursor (Fig. 5b). By 1200 UTC 13 December, the upper precursor has buckled from its straight-line configuration 12 h earlier (cf. Figs. 5c,d with Figs. 5a,b). Although trop within the upper precursor has remained nearly constant during the preceding 12 h (with values slightly less than 300 K), p trop has increased, with values exceeding 450 hpa over a large region extending from Wisconsin to Kentucky (Fig. 5d). The southern predecessor disturbance has nearly disappeared by this time (Figs. 5c,d). The trop gradient along the southwestern edge of the upper precursor has strengthened during the past 12 h (Figs. 5a,c). The most intense trop gradient is now located over southeastern Missouri, western Tennessee, and northeastern Alabama between the upper precursor and an area of locally higher trop that is centered over Mississippi (Fig. 5c). The upper precursor is nearing the East Coast at this time and is approaching a region of relatively strong esfc gradient evident along the coast of the southeastern United States. There is an inverted trough in the SLP field extending northward from a cyclone that has formed offshore of the southeastern United States in association with the now defunct predecessor trough (Fig. 5d). A small region of higher trop and lower p trop has developed east of the upper precursor, offshore of Cape Hatteras, North Carolina (Figs. 5c,d). Although

11 NOVEMBER 1997 LACKMANN ET AL FIG. 6. As in Fig. 4 except for 0000 UTC 14 December 1988 [(a) and (b)] and 1200 UTC 14 December 1988 [(c) and (d)]. this feature is possibly spurious, its location corresponds to an area of deep convection (not shown) that is erupting offshore immediately west of the inverted trough axis at this time. The presence of convection suggests that the redistribution of PV due to convective latent heat release could be responsible for this feature. Although trop in the upper precursor has remained quasi-constant from 0000 UTC 12 December through 0000 UTC 14 December, p trop continues to increase, reaching values in excess of 500 hpa by 0000 UTC 14 December (Figs. 6a,b). The region of maximum p trop is now nearly circular (Fig. 6b), and a tropopause fold has developed along its southern and southeastern portions (not shown). During the preceding 12 h, the northern extremity of the 310-K esfc contour has advanced along the East Coast from east of the Carolinas (Fig. 5c) to east of the Delmarva Peninsula (Fig. 6a). The similarity between values of trop and esfc east of North Carolina at 0000 UTC 14 December indicates that the tropospheric stratification is conditionally neutral in this region. That such a stratification is highly conducive to vertical coupling between upper and lower disturbances is evident from the discussion in section 2, where the Rossby depth becomes very large relative to the depth of the troposphere in the limit of vanishing stratification parameter, N, modified appropriately to account for moist processes (e.g., Whitaker and Davis 1994). The inverted trough extending northward along the East Coast that was beginning to form at 1200 UTC 13 December has become more pronounced, and a separate cyclone center has developed east of Virginia north of the initial center (Fig. 6b). The 310-K trop and esfc contours in the vicinity of the IOP 2 upper disturbance intersect over the eastern portion of the upper precursor at 1200 UTC 14 December, indicative of a conditionally unstable stratification, strong vertical coupling, and cyclogenesis (Fig. 6c). Values of trop in the center of the IOP 2 upper precursor have increased to between 300 and 310 K, presumably due to nonconservative processes such as latent heat release. There is a strong esfc gradient in the vicinity of a warm front northeast of the upper precursor and

12 2740 MONTHLY WEATHER REVIEW VOLUME 125 FIG. 7. Sequence of plots of 500-hPa absolute vorticity (thin solid, contour interval s 1, shaded as indicated in legend at bottom of panel), wind speed (thick dashed, contour interval 10 m s 1, isotachs less than 40 m s 1 omitted), and geopotential height (medium solid, contour interval 6 dam) for IOP 2 event: (a) 0000 UTC 12 December 1988, (b) 1200 UTC 12 December 1988, (c) 0000 UTC 13 December 1988, (d) 1200 UTC 13 December an anticyclonically curved jet at the tropopause level north of this front (Fig. 6c). At the center of the upper disturbance, p trop now exceeds 550 hpa, and rapid surface cyclogenesis is under way with a SLP minimum of 975 hpa in the ECMWF analysis (Fig. 6d, minimum value not shown). Manual analyses performed by Sanders (1990) for this case indicate a central SLP of 964 hpa at this time. This 11-hPa discrepancy may be due to the fact that the inner core of the cyclone lies below the resolution of the ECMWF analysis. b. Vorticity diagnosis Many published papers concerning the behavior of mobile upper troughs have demonstrated that the 500- hpa vorticity field provides an adequate description of these disturbances (e.g., Sanders 1986, 1987, 1988; Lefevre and Nielsen-Gammon 1995). Therefore, we focus on the vorticity evolution at this level throughout this subsection. The evolution of the 500-hPa absolute vorticity a and geopotential height fields during the initial intensification of the IOP 2 upper precursor (0000 UTC 12 December through 1200 UTC 13 December) is summarized in Fig. 7. The IOP 2 upper precursor appears as a disorganized region of relatively large a over British Columbia and Alberta at 0000 UTC 12 December (Fig. 7a). This a maximum is located within a region of cyclonic wind shear north of a jet streak that extends across southern British Columbia into Alberta. There is no apparent trough in the geopotential height field corresponding to the IOP 2 upper precursor at this time, although very slight cyclonic curvature is evident over northern Alberta (Fig. 7a). At 1200 UTC 12 December, there is an indication of cyclonic curvature in the geopotential height contours over Manitoba and Saskatchewan (Fig. 7b). The jet streak has advanced southeastward, with analyzed wind speeds in excess of 40 m s 1 over Montana at this time. By 0000 UTC 13 December,

13 NOVEMBER 1997 LACKMANN ET AL the upper precursor exhibits a coherent signature in both the a and the geopotential height fields (Fig. 7c). The precursor now is located in the left exit of the jet streak, which has intensified to over 50 m s 1 and is advancing southeastward into Nebraska. The upper precursor is amplifying rapidly by 1200 UTC 13 December, with significant increases in a and in the curvature of the geopotential height contours (Fig. 7d). The continuing intensification of the jet streak is evident from the expansion in the areal coverage of the 50 m s 1 contour. The a maximum now is centered upstream of the trough axis in the geopotential height field (Fig. 7d). Comparison of Figs. 7a d with Figs. 4b,d and 5b,d demonstrates that the evolution of a at 500 hpa closely parallels that of p trop, with both quantities increasing simultaneously in the IOP 2 upper precursor. Several published studies have defined and tracked mobile upper troughs using some measure of 500-hPa curvature vorticity, either deduced subjectively from the geopotential height field (e.g., Sanders 1988) or calculated explicitly (e.g., Lefevre and Nielsen-Gammon 1995). Nevertheless, a jet streak often accompanies a mobile upper trough and is an integral component of the class of compacting upper-precursor disturbances investigated in the present study. Because jet streaks exhibit the signature of a cyclonic anticyclonic dipole of shear vorticity elongated along the jet axis, we propose that the 500-hPa shear vorticity is a useful complement to the curvature vorticity in defining and tracking mobile upper troughs. To assess this proposition, we examine maps of 500-hPa curvature and shear vorticity to quantify the extent to which the IOP 2 upper precursor is manifested as a trough or a jet streak at various stages of its life cycle. For background on the properties of curvature and shear vorticity, the reader is referred to Bell and Keyser (1993). The curvature and the shear vorticity ( cu and sh, respectively) may be expressed in spherical coordinates as and [ ] 1 u u 2 2 cu u u 2 V x y x y u tan f, (1) a [ ] 1 u u 2 2 sh u u, (2) 2 V x y x y where is latitude, a is the radius of the earth, and other symbols have their usual meaning. Horizontal derivatives are given by / x (a cos ) 1 / and / y a 1 /, where denotes longitude. Centered finite differences are used in evaluating (1) and (2), with and equal to 1. We compute cu and sh at 500 hpa and present their distributions in Figs. 8 and 9 for the period between 0000 UTC 12 December and 1200 UTC 13 December. At 0000 UTC 12 December, there are two lobes of relatively large cu that overlap a zonally elongated band of large sh over north-central British Columbia and Alberta (Figs. 8a,b). This elongated band of large sh is associated with the jet streak extending from off the Pacific Northwest coast eastward over the southern portions of British Columbia and Alberta. An area of negative sh is located off the Pacific Northwest coast south of the jet streak (Fig. 8b). At 1200 UTC 12 December, a region of enhanced cu extends southward into North Dakota near the incipient IOP 2 upper precursor (Fig. 8c). Two maxima in sh are present over southwestern Canada at this time, with negative sh centers located over Oregon and southwestern Montana; these regions of positive and negative sh bracket the jet streak (Fig. 8d). The tongue of large cu that extends southward into North Dakota at 1200 UTC 12 December penetrates as far south as Missouri by 0000 UTC 13 December (Fig. 9a). A well-organized dipole in the sh field is evident over central North America (Fig. 9b). At 0000 UTC 13 December, the sh distribution in the IOP 2 upper precursor resembles the a distribution (Figs. 9b and 7c). Despite this resemblance, increases in cu and in the curvature of geopotential height contours relative to 12 h earlier are becoming apparent at the downstream edge of the precursor (Figs. 9a and 8c). The cu maximum in the vicinity of the IOP 2 upper precursor extends southward between 0000 UTC and 1200 UTC 13 December (Figs. 9a,c). A marked intensification of the sh extrema, especially the cyclonic member of the dipole, occurs during the same time period (Figs. 9b,d). The sh contribution largely determines the structure of the a field at 1200 UTC 13 December (Figs. 9d and 7d), which is consistent with the displacement of the a maximum upstream of the trough axis in the geopotential height field (Fig. 7d). Although the inclusion of planetary vorticity f in cu (1) increases the magnitude of the cu contribution to a value similar to that of sh at 1200 UTC 13 December, the planetary vorticity does not have a significant impact on the structure of the upper precursor because of its relative uniformity over the spatial extent of this feature. Furthermore, when the contribution of f is subtracted from cu, the resulting relative curvature vorticity, which reflects streamline curvature, is subordinate in magnitude to the shear vorticity. For example, at the location of the cu maximum associated with the upper precursor (near the Tennessee North Carolina border, Fig. 9c), f is approximately s 1 (based on a latitude of 36 N). Subtracting this value of f from the corresponding value of cu at this location ( s 1 ) yields a relative curvature vorticity of s 1. This value is approximately one-half that of the maximum sh associated with the upper precursor of s 1 (located in western Illinois, Fig. 9d).