Airborne Doppler Radar and Sounding Analysis of an Oceanic Cold Front

Roger M. Wakimoto National Center for Atmospheric Research,* Boulder, Colorado

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Hanne V. Murphey Department of Atmospheric and Oceanic Sciences, University of California, Los Angeles, Los Angeles, California

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Abstract

An analysis of a cold front over the eastern Atlantic Ocean based on airborne Doppler wind syntheses and dropsonde data is presented. The focus and unique aspect of this study is a segment of the front that was near the center of the cyclone. The dual-Doppler wind synthesis of the frontal zone combined with an average dropsonde spacing of ∼30 km covers a total distance of >450 km in the cross-frontal direction. The finescale resolution and areal coverage of the dataset are believed to be unprecedented. The cold front was characterized by a distinct wind shift and a strong horizontal temperature gradient. The latter was most intense aloft and not at the surface, in contrast to the classical paradigm of surface cold fronts. The shear of the alongfront component of the wind was relatively uniform as a function of height within the frontal zone. This observation is contrary to studies suggesting that frontal zones decrease in intensity above the surface. The surface convergence within the frontal zone was weak. This may have been related to the closeness of the analysis region to the surface low pressure. The prefrontal low-level jet and the upper-level polar jet were both shown to be supergeostrophic based on the analysis of the geopotential height field. It is believed that a major contributing factor to the former was the isallobaric wind from the large pressure tendencies associated with the moving cyclone. A dry pocket accompanied by descending air was noted out ahead of the low-level jet. This pocket produced a region of potential instability that could have supported deep convection, although none was observed on this day. The vertical structure of the front revealed couplets of potential vorticity that appeared to be the result of diabatic heat sources from condensation. The diabatic effect in the frontogenesis equation was the dominant term, exceeding the combined effects of the confluence and tilting terms. As a result, an alternating pattern of frontogenesis–frontolysis developed along the flanks of the maxima of diabatic heating. This study highlights the importance of taking diabatic heating into account even in the absence of deep convection.

* The National Center for Atmospheric Research is sponsored by the National Science Foundation

Corresponding author address: Roger M. Wakimoto, NCAR/EOL, P.O. Box 3000, Boulder, CO 80307. Email: wakimoto@ucar.edu

Abstract

An analysis of a cold front over the eastern Atlantic Ocean based on airborne Doppler wind syntheses and dropsonde data is presented. The focus and unique aspect of this study is a segment of the front that was near the center of the cyclone. The dual-Doppler wind synthesis of the frontal zone combined with an average dropsonde spacing of ∼30 km covers a total distance of >450 km in the cross-frontal direction. The finescale resolution and areal coverage of the dataset are believed to be unprecedented. The cold front was characterized by a distinct wind shift and a strong horizontal temperature gradient. The latter was most intense aloft and not at the surface, in contrast to the classical paradigm of surface cold fronts. The shear of the alongfront component of the wind was relatively uniform as a function of height within the frontal zone. This observation is contrary to studies suggesting that frontal zones decrease in intensity above the surface. The surface convergence within the frontal zone was weak. This may have been related to the closeness of the analysis region to the surface low pressure. The prefrontal low-level jet and the upper-level polar jet were both shown to be supergeostrophic based on the analysis of the geopotential height field. It is believed that a major contributing factor to the former was the isallobaric wind from the large pressure tendencies associated with the moving cyclone. A dry pocket accompanied by descending air was noted out ahead of the low-level jet. This pocket produced a region of potential instability that could have supported deep convection, although none was observed on this day. The vertical structure of the front revealed couplets of potential vorticity that appeared to be the result of diabatic heat sources from condensation. The diabatic effect in the frontogenesis equation was the dominant term, exceeding the combined effects of the confluence and tilting terms. As a result, an alternating pattern of frontogenesis–frontolysis developed along the flanks of the maxima of diabatic heating. This study highlights the importance of taking diabatic heating into account even in the absence of deep convection.

* The National Center for Atmospheric Research is sponsored by the National Science Foundation

Corresponding author address: Roger M. Wakimoto, NCAR/EOL, P.O. Box 3000, Boulder, CO 80307. Email: wakimoto@ucar.edu

1. Introduction

The Bergen School in Norway provided the first conceptualization of the evolution of fronts within an extratropical cyclone. Subsequently, numerous studies have endeavored to understand the dynamics and observational structure of these discontinuities since they are the focus of a significant fraction of the cyclone’s precipitation and rapid changes in weather. Fronts also provide an environment for instabilities that may be responsible for generating smaller-scale weather phenomena (e.g., Carbone 1983; Emanuel 1983; Keyser 1986). Most of the literature has focused on cold fronts mainly as a result of large horizontal temperature gradients, abrupt wind shifts, and temperature changes, and their frequent association with severe weather (e.g., Browning and Harrold 1970; Hobbs and Biswas 1979; James and Browning 1979; Matejka et al. 1980; Carbone 1982; Hobbs and Persson 1982; Ogura and Portis 1982; Young and Johnson 1984; Shapiro et al. 1985; Bond and Shapiro 1991; Locatelli et al. 1995; Wakimoto and Bosart 2000).

Detailed and near-simultaneous observations of cold fronts have been difficult to obtain because of their long cross-frontal dimension spanning hundreds of kilometers and their depth that can encompass most of the troposphere. Accordingly, most past studies have focused on a small segment of the front or have made restrictive, steady-state assumptions during the data collection period. In addition, it is challenging to collect comprehensive data on the kinematic structure while also accurately defining the thermodynamic structure of the front. The wind field associated with a cold front with sparse information on its thermodynamic structure has been obtained using single-Doppler (e.g., Browning and Harrold 1970; Testud et al. 1980; Hobbs and Persson 1982) and dual-Doppler techniques (e.g., Carbone 1982; Lemaître et al. 1989; Roux et al. 1993; Miller et al. 1996; Wakimoto and Bosart 2000; Jorgensen et al. 2003). Bond and Fleagle (1985), Bond and Shapiro (1991), and Neiman et al. (1993) used dropsonde observations and in situ data collected at flight level to reproduce the two-dimensional structure of oceanic cold fronts. Ogura and Portis (1982) used rawinsonde and dense surface mesonetwork observations to examine a strong cold front associated with severe weather. Thorpe and Clough (1991) recorded one of the most comprehensive datasets by rapidly deploying dropsondes over a distance of ∼500 km in the cross-frontal direction. The average sounding spacing in their study was ∼60 km.

This paper presents an analysis of an oceanic cold front based on data collected by an airborne Doppler radar during the Fronts and Atlantic Storm-Track Experiment (FASTEX; Joly et al. 1997) intensive observation period (IOP) 11 on 5–6 February 1997. A research aircraft equipped with a Doppler radar flew several legs (∼200 km in length) in the cross-frontal direction, providing a detailed vertical cross section of its kinematic structure. This synthesized wind field was significantly enhanced by a series of dropsondes released by a second aircraft flying a parallel track at a higher altitude. Mesoscale structures such as frontal rainbands require observational spacing of 20–30 km (Thorpe and Clough 1991). The average spacing between drops in the present case was ∼30 km. This combination of kinematic and thermodynamic data, both in areal extent and finescale resolution, collected on a cold front is believed to be unprecedented. In addition, the total elapsed time for the dropsonde flight leg was 1 h while covering a distance of >450 km. This is not believed to be a restrictive time interval to assume it being stationary. The horizontal extent of the dual-Doppler wind synthesis relative to the aircraft track also allows for an assessment of the alongfrontal variability of the kinematic structure. Past studies using dropsonde and in situ data collected at flight level have assumed uniformity in the alongfront direction. Radar analyses have shown that some cold fronts can be associated with substantial alongfront variability (e.g., precipitation core and gap structure). FASTEX and the aircraft platforms are described in section 2. Section 3 provides an overview of IOP 11 and presents the tracks of the research aircraft. The detailed analysis of the cold front including examining the frontogenesis equation and the potential vorticity (PV) field is shown in section 4. A summary and discussion are presented in section 5.

2. FASTEX and the aircraft platforms

The field phase of FASTEX occurred in January and February 1997. The dataset collected during the experiment was intended to help improve forecasts of end-of-storm-track cyclogenesis over the northeastern Atlantic Ocean (Joly et al. 1997). While there were a number of scientific objectives, the analysis in this paper focuses on the goal to document the mesoscale organization of cyclone cloud systems. Major observing systems for FASTEX included several research ships, six research aircraft, and a number of buoys deployed over the Atlantic Ocean. For additional information on FASTEX, the reader is referred to Joly et al. (1997).

Two primary platforms used in the present study are the National Oceanic and Atmospheric Administration (NOAA) P-3 and the C-130 operated by the Met Office. The P-3 and the C-130 were deployed from Shannon, Ireland, and Lyneham, United Kingdom, respectively. The typical flight track required both aircraft to fly west over the Atlantic Ocean to intercept extratropical cyclones as they were approaching Europe. The C-130 was equipped to release dropsondes at regularly spaced intervals from altitudes between 6 and 8 km above mean sea level (MSL). The dropsondes were rapidly deployed and led to horizontal spacing as small as 19 km with an average spacing of 30 km. The flight leg that is the focus of this study included a series of 14 dropsondes. The P-3 is equipped with an X-band tail radar and a suite of probes capable of recording in situ measurements (Jorgensen et al. 1983). The scanning parameters for the airborne Doppler radar are shown in Table 1. Discussion of the wind synthesis technique is presented in the appendix.

3. Overview of IOP 11 and the P-3/C-130 flight tracks

The incipient upper-level trough and surface low associated with IOP 11 organized on 3–4 February 1997 over the western Atlantic Ocean (not shown). Subsequently, the entire system rapidly intensified as it moved eastward into the FASTEX domain. A strong baroclinic wave is apparent in the midtropospheric analyses at the 500- and 700-mb (hPa) levels at 0000 UTC 6 February 1997 in Fig. 1 (hereinafter, all times are UTC). The temperature analysis in the figure depicts prominent regions of cold- and warm-air advection that clearly delineate the location of the fronts at upper levels. The cold (warm) advection into the trough (ridge) is consistent with an amplifying system based on the quasigeostrophic height tendency equation. The surface cyclone was associated with its lowest central pressure (<984 mb) at 0000 UTC 6 February 1997 (Fig. 2a) and was moving to the northeast at ∼20 m s−1.

An extensive cirrus cloud shield was apparent in the infrared satellite image and a hooklike cloud pattern can be identified at ∼46°N and 32°W in Fig. 2. This cloud pattern was collocated with a positive potential vorticity anomaly on the 330-K isentrope (not shown). This anomaly was associated with a weak pressure trough at the surface. In contrast to the limited horizontal dimension of the warm front, the surface cold front extends for distances greater than 2000 km from the cyclone center to the south and southwest. The difference in the horizontal extent of these two fronts has been discussed in past studies (e.g., Hoskins and West 1979; Bluestein 1993, 255–270; Schultz et al. 1998).

Flight tracks for the two aircraft were designed during the pre–IOP 11 planning meetings. These tracks were based on the expected position and movement of the cold front. The P-3 flew primarily at an altitude between 2 and 3 km and executed 4 approximately perpendicular penetrations through the frontal surface (Figs. 2b and 3). The tracks plotted on the figures are ground relative and were near the center of the cyclone. The flight plans for the C-130 were designed such that the dropsondes would be deployed from an altitude of ∼8 km along a path that was parallel to the P-3 tracks. The time–space-adjusted aircraft tracks are shown in Fig. 4 using a frontal motion 9.6 m s−1 from 300°. The estimate of the speed and direction of the front is discussed in the appendix.

The data collected during the execution of the two tracks shown in Fig. 4 are the focus of the current study. The tracks were chosen based on the quality of the Doppler radar data and the completeness of the thermodynamic and kinematic information collected by the dropsonde deployments. The time synchronization of the flight tracks shown in Fig. 4 was good with the P-3 trailing the C-130 by a time difference that did not exceed 13 min. The P-3 and C-130 completed their flight legs in ∼40 and 60 min, respectively. The key assumptions for these analyses are that the front remains steady during the sampling period and that the P-3 and C-130 are sampling the same air mass even though the tracks are separated by ∼35 km. The detailed wind syntheses from the airborne Doppler radar were merged with the thermodynamic and kinematic information from the dropsondes based on these assumptions. These analyses are presented in the next section.

4. Structure of the cold front

a. Horizontal structure

The dual-Doppler synthesis of the ground-relative winds at 400 m from 0245–0320 UTC is presented in Fig. 5 (the flight track in this figure was also shown in Fig. 4). The position of the front is located on the western boundary of a strong southwesterly jet (speeds >40 m s−1) and at the leading edge of the strongest horizontal temperature gradient (shown in the next section). A prominent wind shift is evident as the wind direction changes from a southwesterly to a strong northerly flow ∼100 km behind the surface location of the front.

The precipitation echoes associated with the front were weak and scattered throughout the region. A linear band of echo is collocated with the front with maximum reflectivity values <30 dBZ. An analysis of the vertical vorticity is shown in an enlargement of the analysis region (Fig. 5b). The front is collocated with the peak cyclonic vertical vorticity >10−3 s−1. The maximum values are substantially weaker than other strong fronts documented in the literature (e.g., Hobbs and Persson 1982; Carbone 1982, 1983). However, the peak cyclonic vorticity is an order of magnitude larger than the Coriolis parameter at this latitude.

Since the peak magnitude of the vertical vorticity is scale dependent, the cross- and alongfront wind components through the frontal zone were compared with the aforementioned studies. Hobbs and Persson (1982) document a shear of ∼7 m s−1 in 2 km of the alongfront component of the horizontal wind through the wind-shift zone of a cold front. In comparison, the Doppler wind synthesis in the present study reveals a ∼25 m s−1 shear of the wind in 60 km (or ∼0.8 m s−1 in 2 km). No comparison of the across-front component of the wind between these two cases could be made. Carbone (1982) showed an ∼15 m s−1 change in the across-front component of the wind in 2 km within the wind-shift zone. The current study revealed ∼15 m s−1 over a distance of 60 km (or ∼0.5 m s−1 in 2 km). No comparison of the alongfrontal component of the wind between these two cases could be made. The comparisons discussed above highlight the fact that the FASTEX cold front was associated with a substantially weaker wind discontinuity than the fronts analyzed by Hobbs and Persson (1982) and Carbone (1982, 1983).

b. Vertical structure

A vertical cross section was created that was oriented approximately perpendicular to the cold front. The dual-Doppler wind synthesis was rotated such that the abscissa was parallel to the front and the ordinate was pointed into the postfrontal cold air. The cross sections of the horizontal and vertical components of the wind for each grid point along the abscissa were averaged (i.e., a total of 27 cross sections over a distance of 40 km). This mean field was merged with the wind and thermodynamic data recorded from the dropsondes deployed from 0231 to 0330 UTC by the C-130 along the track shown in Fig. 4. The merged analysis is presented in Fig. 6. A vertical cross section of standard deviations for the synthesized dual-Doppler winds was created (not shown) in order to assess the impact of horizontal alongfrontal variability. The maximum standard deviations were <2 m s−1 and 3° in wind speed and direction, respectively, in regions located away from the frontal zone. These numbers increased to 4 m s−1 and 20° within the frontal zone. The latter values are not surprising in a region characterized by strong horizontal and vertical wind shears.

The gray wind vectors in Fig. 6 highlight the dual-Doppler wind synthesis. Note the close agreement in the wind speed and direction between the dropsonde- and Doppler-derived wind fields, suggesting that the separation in flight tracks is not a major limitation of the dataset. The cross section covers a distance of >450 km, which is comparable to the results shown by Thorpe and Clough (1991). However, the average spacing between soundings in the present case is smaller and the dual-Doppler wind field provides a level of detail that has not been available in prior studies of cold fronts.

The kinematic boundary of the cold front is well defined by the shift in wind direction from a northerly flow within the postfrontal cold air to a southwesterly flow in the warm sector (Fig. 6a). The front is also apparent in the sloping isopleths of virtual potential temperature θυ over the polar air mass. The mixing ratio analysis reveals the relatively moist and dry air ahead of and behind the front, respectively.

Also apparent in the mixing ratio plot is the elevated pocket of dry air located 100–200 km ahead of the front. The location of this dry air is east of the low-level jet (Fig. 6c). Browning and Pardoe (1973) were the first to suggest the existence of dry descent in advance of the low-level jet. Ogura and Portis (1982) documented this drying in a region of descent in their study of a cold front (see their Figs. 16 and 19). More recently, Schultz (2005) has emphasized that prefrontal descent has been documented in a number of studies without addressing its cause (e.g., Hobbs et al. 1980; Hsie et al. 1984; Crook 1987; Bénard et al. 1992; Schultz et al. 1997; Thompson and Williams 1997; Chen and Bishop 1999). In a few cases, this descent appears to be related to compensating flow from deep, moist convection (e.g., Ogura and Portis 1982; Crook 1987; Bénard et al. 1992), although it is unclear what other physical processes may be contributing to the development of this feature.

The dry pocket in Fig. 6a is located in a region of slight warming (indicated by the dashed line), suggesting that the dry air is a result of descending air.1 The absence of strong convection in the analysis region effectively discounts forced descent as a causal mechanism on this day. Low equivalent potential temperatures are also apparent in this region, producing a region of potential instability (Fig. 6b). It is possible that the low values near the 700-mb level are a consequence of cloud-top radiational cooling. Descent in this region could have resulted in clearing at midtropospheric levels required for cooling of the cloud layer below. This type of thermodynamic structure could support prefrontal convective activity although none was observed on this day. A layer characterized by high values of θE (>310 K) is seen above the frontal surface. This layer does not extend continuously into the warm sector because of the pocket of low-θE air. It should be noted that the pocket of dry air does not appear to be related to a forward-tilting cold front associated with the split-cold front model (e.g., Browning and Monk 1982; Browning 1986; Hobbs et al. 1990).

There were weak positive vertical velocities near the leading edge of the cold front between the dropsondes deployed at 0253 and 0257 UTC(Fig. 6b). The maximum updraft speeds (>2 m s−1) were noted at ∼75 km in postfrontal air mass and were positioned near a wavelike pattern in isopleths in θE (note the 308- and 310-K isopleths near the dropsonde deployed at 0249). A similar feature in the θE field can be identified along the 310-K isopleth near the 0241 dropsonde location. Interestingly, the latter is also near an updraft region. Such features have been noted before and attributed to precipitation bands forming as a result of slantwise convection or gravity waves generated during frontogenesis (e.g., Emanuel 1983; Parsons and Hobbs 1983; Wolfsberg et al. 1986; Gall et al. 1988). Snyder et al. (1993) have cautioned, however, that the gravity waves discussed by Gall et al. (1988) were numerical artifacts resulting from insufficient vertical resolution and truncation errors. In the present case, three bands associated with weak radar reflectivities are noted in the echo pattern and are indicated by the gray arrows in Fig. 6c. Each band is near the regions of positive vertical velocity. The band spacing is ∼50 km and shows a multicellular structure that is typical within a mesoscale convective system with the stronger cells at the leading edge and progressively weaker echoes toward the rear of the system (e.g., Leary and Houze 1979).

The component of motion parallel to the front (Fig. 6c) illustrates the strong vertical wind shear that exists between the northerly and southwesterly flow in the polar- and warm-sector air masses, respectively. The wind shear is relatively uniform as a function of height within the frontal zone. This observation is contrary to studies suggesting that frontal zones decrease in intensity above the surface (also shown by Bond and Shapiro 1991). A low-level jet (speeds >34 m s−1) is apparent just ahead of the cold front between 2.5–3 km. Low-level jets developing in advance of cold fronts have been well documented in the literature (e.g., Browning and Harrold 1970; Browning and Pardoe 1973). Also apparent in the figure is a portion of the polar jet stream (located in the upper left-hand side of the figure).

The temperature and pressure data recorded by the dropsondes deployed during this flight leg were used to determine the geopotential height field in the cross-frontal direction. The alongfront geostrophic wind (ug) was subsequently calculated (not shown). The degree of geostrophic balance of the flow parallel to the front was estimated by calculating the ageostrophic wind (uag = uug, see Fig. 6d). A large region of supergeostrophic flow is evident in the warm sector. Browning and Pardoe (1973) proposed that low-level jets were in approximate geostrophic balance. More recently, Keyser and Anthes (1982), Dudhia (1993), and Lafore et al. (1994) used numerical simulations to show that the low-level jet was supergeostrophic. The magnitude of the ageostrophic wind shown in Fig. 6d is larger than previous examinations of cold fronts. A unique aspect of this case is the close proximity of the analysis region to the surface low pressure center. The isallobaric wind can be estimated assuming no change in the isobar pattern shown in Fig. 2a and using the propagation speed of the cyclone of ∼20 m s−1 to the northeast. The latitude is assumed to be 55° and the density is considered constant. These approximations yield a maximum isallobaric wind of ∼19 m s−1. This rough calculation suggests that the ageostrophic winds shown in the figure may not be inherent to the front.

The comparison of Figs. 6c,d also reveals the supergeostrophic nature of the upper-level jet stream located at levels >4 km near the soundings deployed at 0231 and 0237 UTC. The slight ridging aloft (Fig. 1) is estimated to contribute a few meters per second to the supergeostrophic flow at upper levels. The postfrontal flow in Fig. 6d is also supergeostrophic. Dudhia (1993) proposed that there was little ageostrophic flow to the rear of the front. In addition, Lafore et al. (1994) suggested that the polar airstream was in geostrophic balance but for the lowest few hundred meters. Numerical simulations by Keyser and Anthes (1982), however, do suggest that the polar air is slightly supergeostrophic above the surface boundary layer.

The horizontal gradient of virtual potential temperature across the front is >10 K per 100 km with the maximum located aloft at ∼2 km (Fig. 6d). The virtual potential temperature gradient at ∼2 km is nearly twice the value near the surface. Other observational studies of maritime cold fronts have documented the strongest temperature gradients near the surface (e.g., Wakimoto and Cai 2002) and above the boundary layer (e.g., Bond and Fleagle 1985; Thorpe and Clough 1991). It should be noted that an estimate of the temperature gradient based on an objective analysis of the surface data (not shown) was 3–4 K per 100 km. This is less than the near-surface values shown in Fig. 6d (6–7 K per 100 km). The difference is not surprising given the relatively coarse resolution of the surface observations compared to the analysis of the dropsonde data.

An enlarged view of the precipitation echoes and θυ analysis superimposed on the dual-Doppler and dropsonde wind field reveals the detailed kinematic structure of the cold front (Fig. 7a). The flight track of the NOAA P-3 is also plotted on the figure. The distinct wind shift within the frontal zone is illustrated, and the frontal slope over the polar air is approximately 1:60. As previously mentioned, there are distinct updraft maxima that appear to be related to the locations of the precipitation bands (Fig. 7b). The strongest region of cyclonic vertical vorticity (>10−3 s−1) is near the surface location of the dropsonde deployed at 0257 (Fig. 7b). The peak vorticity associated with cold fronts documented in the literature with comparable spatial resolution spans a wide range from 10−4 s−1 (e.g., Miller et al. 1996) to 10−2 s−1 (e.g., Carbone 1983). The maximum vertical vorticity associated with the leading edge of the surface frontal zone weakened substantially with increasing height similar to other documented cases (e.g., Carbone 1983; Wakimoto and Bosart 2000). There exists another zone of positive vertical vorticity (>0.5 × 10−3 s−1) that does not weaken with increasing height (note the profile of vorticity between the 0231 and 0249 UTC soundings), which is associated with the trailing portion of the front.

The relationship between θυ and potential vorticity is shown in Fig. 7c. There are several regions within the frontal zone where PV exceeds 2 potential vorticity units (PVU) with the maximum values (>4 PVU) positioned at the leading edge of the front and approximately 50 km to the west within the postfrontal cold air. The latter is a result of the increased static stability in this region (Fig. 7c) even though the vertical vorticity does not attain its peak values. It should be noted that PV is defined in terms of dry potential temperature (θ) instead of virtual potential temperature used in the present study. Accordingly, another calculation of PV was performed using θ (not shown). There were no substantive differences between the two PV plots that would have altered the conclusions presented in this paper.

Examination of the collocated moist PV values (not shown) revealed numbers near zero. This observation leads to the conclusion that frictional forces may not have contributed substantially to the positive values shown in Fig. 7c. There is a suggestion of two PV couplets (highlighted by the black crosses) in the vertical that are positioned near the two updrafts depicted in Fig. 7b. This may be an example of positive- and negative-PV sources in response to diabatic heating released via condensation in air parcels rising in updrafts. The position and orientation of these couplets in relation to a diabatic heat source are consistent with theory (e.g., Hoskins 1990). In addition, the PV couplets are aligned with the absolute vorticity vector (not shown) as discussed by Hoskins (1990, 1997). The other PV maximum, located at the leading edge of the front, is near the echo with radar reflectivities greater than 18 dBZ.

Although there appears to be a relationship between the PV couplet located between the dropsondes deployed from 0245 to 0249 UTC and the reflectivity maximum (Fig. 7a), a similar relationship is not readily apparent with the couplet located farther to the west. The airborne radar, however, is able to detect only precipitation-size hydrometeors. Condensation in the weak updrafts could be associated with cloud droplets that would not be represented in the plots of radar reflectivity. It is also possible that evaporational cooling in the lowest few kilometers is occurring (near the 0237 UTC dropsonde) based on the mixing ratio analysis shown in Fig. 6a.

The relationship between the horizontal gradient of virtual potential temperature and vertical wind shear of the component of the wind parallel to the cold front is shown in Fig. 7d. The low-level jet positioned above the leading edge of the cold front can also be seen. The two-dimensional front-relative flow (across-front and vertical wind components), based on the dual-Doppler wind syntheses from the surface to just above the frontal zone, is in the opposite direction of the frontal motion. The lack of front-relative flow of the polar air toward the leading edge of the surface front has been documented before (e.g., Browning and Harrold 1970; Testud et al. 1980). Smith and Reeder (1988) have argued that this is a characteristic difference between some cold fronts and density currents.

The thermodynamic analysis combined with the detailed kinematic wind synthesis provided an opportunity to calculate the frontogenesis in two dimensions with greater accuracy than has been previously attempted. For example, a common methodology to determine the vertical velocity field in past studies was to assume no alongfront variation of the wind component parallel to the front when calculating the horizontal divergence (e.g., Sanders 1955; Neiman et al. 1993). This assumption was not needed in the present case because the dual-Doppler coverage extended ∼20 km in both directions perpendicular to the aircraft track. Indeed, the alongfront component of horizontal divergence was nonnegligible in the present case and was equal and opposite to the magnitude of the cross-front component in several regions within the frontal zone. This analysis suggests using caution when interpreting derived vertical velocity fields that assume two-dimensionality of fronts also noted in other Doppler radar studies (e.g., Hobbs and Persson 1982). In contrast, a similar analysis of the vorticity field revealed that the alongfront variation of the wind shear contributed only 10%–20% to the total vertical vorticity. Accordingly, estimates of the vorticity assuming two-dimensionality of the front would not have significantly altered the conclusions presented in this study.

The frontogenesis equation in two dimensions was used as follows:
i1520-0493-136-4-1475-e1
where the term (∂u/∂y)(∂θυ/∂x) has been assumed to be small and is therefore neglected. These terms represent the frontogenetical effects of the confluence, tilting, and diabatic components, respectively. There have been few studies that examine the diabatic term in the frontogenesis equation (e.g., Koch 1984; Keyser 1986; Neiman et al. 1993; Segal et al. 1993; Miller et al. 1996). Turbulent fluxes of sensible heat within the boundary layer are ignored in the present study. Bond and Fleagle (1985) and Wakimoto and Cai (2002) have shown that fluxes of heat were small for their case studies of oceanic cold fronts.

The tilting and confluence components of frontogenesis are shown in Fig. 7e. The confluence term is primarily frontolytic with the minimum values located aloft and to the rear of the surface position of the front. It is not known why strong convergence was absent in the present case, although detailed observational studies of the kinematic structure of cold fronts near the center of the cyclone have not been previously shown. It is possible that the intense isallobaric winds are contributing to the weak convergence, but the physical processes involved are unknown. In addition, the weak convergence is likely responsible for the low values of vertical vorticity observed (Fig. 7b).

The tilting effect is dominant because of the updrafts and the strong static stability within the frontal zone. The combined effects of tilting and confluence (not shown) reveal weak frontogenesis at the surface location of the front. The diabatic heating can be estimated if it is assumed that the tendency term is negligible and that the alongfront variations of θυ are small (Fig. 7f). The latter assumption is partially supported by the surface analyses of θυ (not shown). The cold frontal region is largely associated with diabatic heating primarily through the vertical advective term (i.e., updrafts located where the vertical temperature gradients are strong). The two strongest regions of diabatic heating overlap with the PV couplets (Fig. 7c), which is consistent with the hypothesis that the couplets are a result of a heat source as predicted by theory. This relationship has rarely been shown in observational studies (e.g., Neiman et al. 1993). Weak diabatic heating is associated with the echo located at the leading edge of the front. It is possible that evaporational cooling is occurring near the surface (not resolved in the analysis shown in Fig. 7f) and is contributing to the positive-PV region located close to the 0257 UTC sounding.

The combined effects of the tilting, confluence, and diabatic terms are shown in Fig. 7f. In the present case, the diabatic term in the frontogenesis equation was significantly larger than confluence and tilting effects. Accordingly, the analysis is dominated by an alternating pattern of frontogenesis–frontolysis (Fig. 7f) that is located along the flanks of the maxima of diabatic heating.

5. Summary and discussion

Analysis of an oceanic cold front based on airborne Doppler wind syntheses and dropsonde data was presented. The finescale resolution and areal extent of the kinematic and thermodynamic data used in this study are believed to be unprecedented. Indeed, the average spacing of the dropsondes was ∼30 km while covering a horizontal distance of >450 km. These data were merged with an airborne dual-Doppler wind synthesis. The total elapsed time to collect both datasets was only ∼1 h. A unique aspect of this study was the location of the segment of the front under investigation. The position was very close to the surface low pressure center unlike previous studies. The front was characterized by a distinct wind shift and a horizontal gradient of virtual potential temperature. The latter was most intense aloft, which is different from the classical paradigm of cold fronts. The shear of the alongfront component of the wind was relatively uniform as a function to height within the frontal zone. This observation is contrary to studies suggesting that frontal zones decrease in intensity above the surface.

A dry pocket of descending air was noted out ahead of the front and low-level jet. Such features have been documented before, but the present study appears to eliminate compensating flow as a result of deep convection as a forcing mechanism. This dry area produced a region of potential instability that would have supported prefrontal deep convection, although none was observed on this day. It is not known whether this unstable region is related to the commonly observed prefrontal squall lines that form out ahead of cold fronts (Newton 1950). Improved understanding of the descending flow awaits future studies as suggested by Schultz (2005). Wavelike patterns were noted in the θE analysis of the frontal zone. The waves appeared to be associated with updraft patterns and weak radar reflectivity bands. These bands could have been a result of slantwise convection or gravity waves.

The supergeostrophic nature of the upper- and low-level jets was shown based on the observed geopotential height field. A major contributing factor to the ageostrophic flow associated with the low-level jet was the strong isallobaric wind due to the closeness of the analysis region to the center of the cyclone. The isallobaric winds may have contributed to the weak convergence observed along the frontal zone. Couplets of PV in the vertical were resolved, which appeared to be a result of diabatic heat sources in updrafts. This relationship has rarely been shown in observational studies. The importance of diabatic effects in the frontogenesis equation was presented. This diabatic term was dominant, in the present study, exceeding the combined effects of confluence and tilting. Accordingly, the analysis was dominated by an alternating pattern of frontogenesis–frontolysis located along the flanks of the maxima of diabatic heating. This case highlights the importance of taking diabatic effects into account when calculating frontogenesis even in the absence of deep convection.

Acknowledgments

Research results presented in this paper were partially supported by the National Science Foundation under Grant ATM-021048. Any opinions, findings, and conclusions or recommendations expressed in this publication are those of the authors and do not necessarily reflect the views of the National Science Foundation. Comments from Lance Bosart and Chris Snyder substantially improved an earlier version of this manuscript. Helpful comments from two anonymous reviewers also improved and clarified a number of aspects of this study.

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APPENDIX

Doppler Radar Methodology

The NOAA P-3 radar uses a beam-scanning methodology known as fore–aft scanning (FAST) as described by Jorgensen et al. (1996). The along-track and sweep-angle resolution of the Doppler radar data collected by the P-3 radar were 1.3 km and 1.2°, respectively, as shown in Table 1. The sweep-angle resolution resulted in an effective sampling in the vertical of ∼500 m in the analysis domain. These numbers determined the Cartesian grid with a horizontal and vertical grid spacing of 1.5 km and 500 m, respectively. The aircraft motion was removed from the velocity data by using “SOLO” software (Oye et al. 1995). Both the radar reflectivity and Doppler velocity data were edited and subsequently interpolated onto this grid using a Cressman filter (Cressman 1959) with a radius of influence of 1.5 km and 500 m in the horizontal and vertical directions, respectively. The lowest grid level was located at 400 m MSL.

The movement of the cold front was used to time–space adjust each radar scan. This velocity was 9.6 m s−1 from 300°. This vector was determined based on an analysis of sequential surface analyses of the frontal positions combined with the movement of radar echoes located at the leading edge of the front using the lower fuselage radar onboard the P-3 (Jorgensen et al. 1983). Both methods of estimating frontal movement were consistent in speed and direction. The wind synthesis of the radar data used the Custom Editing and Display of Reduced Information in Cartesian Space software (CEDRIC; Mohr et al. 1986). The hydrometeor fall speeds were estimated from the reflectivity–terminal fall speed relationships established by Joss and Waldvogel (1970) with a correction for the effects of air density (Foote and du Toit 1969). The effect of hydrometeor fall speeds was small because of the scattered and weak radar reflectivities documented during this event.

Vertical air velocities were obtained by vertical integration of the horizontal divergence using a variational adjustment scheme discussed by O’Brien (1970). A four-step Leise filter (Leise 1982) was applied to the wind synthesis that removes wavelengths of less than 24 km in the horizontal direction. This distance is close to the average spacing of the dropsondes.

The standard deviation in vertical velocity was estimated to be 1.0 and 3.0 m s−1 at 2 and 5 km MSL, respectively, following the discussion by Doviak et al. (1976). However, the vertical cross sections of the kinematic wind field shown in this paper were calculated by averaging 27 cross sections. This averaging will reduce the uncertainties of the vertical velocities from these values.

Fig. 1.
Fig. 1.

Midtropospheric analyses for the 500- and 700-mb levels at 0000 UTC 6 Feb 1997. Black lines are contours and dashed lines are isotherms. Wind vectors are plotted with the following notation: flag = 25, barb = 5, and half barb = 2.5 m s−1. Locations of upper-level cold and warm fronts are indicated at the 700-mb level.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 2.
Fig. 2.

(a) Surface analyses superimposed on top of the infrared satellite image at 0000 UTC 6 Feb 1997. Surface reporting stations are shown as black dots. Wind vectors are plotted when available (one barb = 5 m s−1, half barb = 2.5 m s−1). (b) Infrared satellite image at 0000 UTC 6 February 1997 with the superimposed flight tracks of the NOAA P-3 (black line) and the C-130 (dashed line).

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 3.
Fig. 3.

Enlargement of the flight tracks of the NOAA P-3 (black line) and C-130 (dashed line) on 5–6 Feb 1997 shown in Fig. 2b. The star indicates the location of Shannon Airport.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 4.
Fig. 4.

Time–space adjusted aircraft tracks for the NOAA P-3 (black line) and C-130 (dashed line). The locations of the dropsondes deployed by the C-130 are shown.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 5.
Fig. 5.

(a) Dual-Doppler wind syntheses from 0245:00 to 0320:00 UTC 6 Feb 1997 at 400 m MSL. Ground-relative winds are superimposed onto radar reflectivity. (b) Enlargement of the boxed-in area shown in (a). Ground-relative winds are superimposed onto radar reflectivity and vertical vorticity. Vertical vorticity in (b) is shown by the black lines. Radar reflectivity is shown by the gray lines with values >20 dBZ shaded gray. The dashed line is the NOAA P-3 flight track.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 6.
Fig. 6.

Vertical cross section of dropsonde data through the cold front from 0231 to 0328 UTC on 6 Feb 1997. Ground-relative winds from the dropsonde and dual-Doppler synthesis (plotted as gray wind vectors) are superimposed onto (a) virtual potential temperature (contour interval every 2 K, black lines) and mixing ratio (contour interval every 1 g kg−1, gray lines). (b) Equivalent potential temperature (contour interval every 2 K, black lines) and vertical velocities (contour interval every 0.5 m s−1, gray lines) derived from the Doppler synthesis (the dual-Doppler analysis region is shown by the arrows at the top of the figure). Solid and dashed gray lines represent positive and negative vertical velocities, respectively.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 6.
Fig. 6.

(Continued) (c) Component of the wind parallel to the front (u, black lines) superimposed onto radar reflectivity (contour interval every 4 dBZ, gray lines). The rotated coordinate system is shown in the inset. Positive values of u are into the figure. Radar reflectivity values >14 dBZ are shaded gray. The gray arrows at the bottom of the figure denote the location of three reflectivity bands. (d) Ageostrophic component of the wind parallel to the cold front (contour interval every 8 m s−1, black lines) and the gradient of the virtual potential temperature (contour interval every 2 K per 100 km, gray lines). Positive values of uag are into the figure. Values of ageostrophic wind >24 m s−1 are shaded gray. The large gray box is the dual-Doppler analysis region and is enlarged in Fig. 7. The plotted Doppler winds were based on mean vertical cross sections of the synthesized data. The 282 and 292 K isopleths of virtual potential temperature are highlighted by the dashed gray line in (b), (c), and (d) in order to approximately delineate the frontal zone.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 7.
Fig. 7.

Enlargement of the boxed-in area shown in Fig. 6 of dropsonde data through the cold front from 0231 to 0311 UTC on 6 Feb 1997. Ground-relative winds from the dropsondes and Doppler wind synthesis are superimposed onto (a) virtual potential temperature (contour interval every 2 K, black lines) and radar reflectivity (contour interval every 4 dBZ, gray lines). Values of reflectivity >14 dBZ are shaded gray. (b) Vertical vorticity (contour interval every 0.5 × 10−3 s−1, black lines) and vertical velocity (contour interval every 0.5 m s−1, gray lines). Positive and negative values of vertical vorticity are drawn as solid and dashed lines, respectively (the dashed dotted line is the 0.25 × 10−3 s−1 isopleth of vorticity).

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 7.
Fig. 7.

(Continued) (c) Virtual potential temperature (contour interval every 2 K, black lines) and cross-frontal potential vorticity (contour interval every 2 PVU, gray lines) defined as −g[ f + (−∂u/∂y)θ](∂θυ/∂p), where (−∂u/∂y)θ is the cross-frontal vorticity on an isentropic surface and (∂θυ/∂p) is the static stability. Black crosses embedded within a circle represent the center of PV couplets. PVU = 10−6 m2 s−1 K kg−1. Values of potential vorticity >4 × 10−6 m2 s−1 K kg−1 are shaded gray. (d) Component of wind parallel to the cold front (contour interval every 4 m s−1, black lines) and the virtual potential temperature gradient (contour interval every 2 K per 100 km, gray lines). Positive values of u are into the figure. Black arrows are the front-relative flow in the plane of the cross section (across-front and vertical wind components). Values of the virtual potential temperature gradient >8 K per 100 km are shaded gray.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Fig. 7.
Fig. 7.

(Continued) (e) Tilting (K per 100 km h−1, black lines) and confluence frontogenesis (K per 100 km h−1, gray lines). Positive and negative values are shown by the solid and dashed lines, respectively. (f) Total frontogenesis (K per 100 km h−1, black lines) and the diabatic term (K h−1, gray lines). Positive and negative values are shown by the solid and dashed lines, respectively. The track of the P-3 is shown by the dotted line. The 282 and 292 K isopleths of virtual potential temperature are highlighted by the dashed gray line in (b), (d), (e), and (f) in order to approximately delineate the frontal zone.

Citation: Monthly Weather Review 136, 4; 10.1175/2007MWR2241.1

Table 1.

Characteristics of the P-3 Doppler radar.

Table 1.

1

The reader should be cautioned that a single analysis cannot be used to infer local warming. In addition, warm-air advection from southerly flow cannot be ruled out.

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  • Bénard, P., J-L. Redelsperger, and J-P. Lafore, 1992: Nonhydrostatic simulation of frontogenesis in a moist atmosphere. Part I: General description and narrow rainbands. J. Atmos. Sci., 49 , 22002217.

    • Search Google Scholar
    • Export Citation
  • Bluestein, H. B., 1993: Observations and Theory of Weather Systems. Vol. 2, Synoptic-Dynamic Meteorology in Midlatitudes, Oxford University Press, 594 pp.

    • Search Google Scholar
    • Export Citation
  • Bond, N. A., and R. G. Fleagle, 1985: Structure of a cold front over the ocean. Quart. J. Roy. Meteor. Soc., 111 , 739759.

  • Bond, N. A., and M. A. Shapiro, 1991: Research aircraft observations of the mesoscale and microscale structure of a cold front over the eastern Pacific Ocean. Mon. Wea. Rev., 119 , 30803094.

    • Search Google Scholar
    • Export Citation
  • Browning, K. A., 1986: Conceptual models of precipitation systems. Wea. Forecasting, 1 , 2341.

  • Browning, K. A., and T. W. Harrold, 1970: Air motion and precipitation growth at a cold front. Quart. J. Roy. Meteor. Soc., 96 , 369389.

    • Search Google Scholar
    • Export Citation
  • Browning, K. A., and C. W. Pardoe, 1973: Structures of low-level jet streams ahead of mid-latitude cold fronts. Quart. J. Roy. Meteor. Soc., 99 , 619638.

    • Search Google Scholar
    • Export Citation
  • Browning, K. A., and G. A. Monk, 1982: A simple model for the synoptic analysis of cold fronts. Quart. J. Roy. Meteor. Soc., 108 , 435452.

    • Search Google Scholar
    • Export Citation
  • Carbone, R. E., 1982: A severe frontal rainband. Part I: Stormwide hydrodynamic structure. J. Atmos. Sci., 39 , 258279.

  • Carbone, R. E., 1983: A severe frontal rainband. Part II: Tornado parent vortex circulation. J. Atmos. Sci., 40 , 26392654.

  • Chen, C., and C. Bishop, 1999: Reply. Mon. Wea. Rev., 127 , 258263.

  • Cressman, G. P., 1959: An operational objective analysis scheme. Mon. Wea. Rev., 87 , 367374.

  • Crook, N. A., 1987: Moist convection at a surface cold front. J. Atmos. Sci., 44 , 34693494.

  • Doviak, R. J., P. S. Ray, R. G. Strauch, and L. J. Miller, 1976: Error estimation in wind fields derived from dual-Doppler radar measurements. J. Appl. Meteor., 15 , 868878.

    • Search Google Scholar
    • Export Citation
  • Dudhia, J., 1993: A nonhydrostatic version of the Penn State–NCAR mesoscale model: Validation tests and simulation of an Atlantic cyclone and cold front. Mon. Wea. Rev., 121 , 14931513.

    • Search Google Scholar
    • Export Citation
  • Emanuel, K. A., 1983: On assessing local conditional symmetric instability from atmospheric soundings. Mon. Wea. Rev., 111 , 20162033.

    • Search Google Scholar
    • Export Citation
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