• Barnes, G. M., and K. Sieckman, 1984: The environment of fast- and slow-moving tropical mesoscale convective cloud lines. Mon. Wea. Rev.,112, 1782–1794.

  • Bartels, D. L., and R. A. Maddox, 1991: Midlevel cyclonic vortices generated by mesoscale convective systems. Mon. Wea. Rev.,119, 104–118.

  • Battan, L. J., 1973: Radar Observations of the Atmosphere. University of Chicago Press, 324 pp.

  • Biggerstaff, M. L., and R. A. Houze Jr., 1991: Kinematics and precipitation structure of the 10–11 June 1985 squall line. Mon. Wea. Rev.,119, 3034–3065.

  • ——, and ——, 1993: Kinematics and microphysics of the transition zone of a midlatitude squall-line system. J. Atmos. Sci.,50, 3091–3110.

  • Braun, S. A., and R. A. Houze Jr., 1994: The transition zone and secondary maximum of radar reflectivity behind a midlatitude squall line: Results retrieved from Doppler radar data. J. Atmos. Sci.,51, 2733–2755.

  • Chong, M., P. Amayenc, G. Scialom, and J. Testud, 1987: A tropical squall line observed during the COPT 81 experiment in West Africa. Part I: Kinematic structure inferred from dual-Doppler data. Mon. Wea. Rev.,115, 670–694.

  • Davis, C. A., and M. L. Weisman, 1994: Balanced dynamics of simulated, long-lived mesoscale convective systems. J. Atmos. Sci.,51, 2005–2030.

  • Fairall, C. W., E. F. Bradley, D. P. Rogers, J. B. Edson, and G. S. Young, 1996: Bulk parameterization of air–sea fluxes for TOGA COARE. J. Geophys. Res.,101, 3747–3765.

  • Fovell, R. G., and Y. Ogura, 1988: Numerical simulation of a midlatitude squall line in two dimensions. J. Atmos. Sci.,45, 3846–3879.

  • Gal-Chen, T., 1978: A method for the initialization of the anelastic equations: Implications for matching models with observations. Mon. Wea. Rev.,106, 587–696.

  • ——, and R. A. Kropfli, 1984: Buoyancy and pressure perturbations derived from dual-Doppler radar observations of the planetary boundary layer: Applications for matching models with observations. J. Atmos. Sci., 41, 3007–3020.

  • Hane, C. E., 1993: Storm motion estimates derived from dynamic retrieval calculations. Mon. Wea. Rev.,121, 431–443.

  • ——, and D. P. Jorgensen, 1995: Dynamic aspects of a distinctly three-dimensional mesoscale convective system. Mon. Wea. Rev.,123, 3194–3214.

  • ——, R. B. Wilhelmson, and T. Gal-Chen, 1981: Retrieval of thermodynamic variables within deep convective clouds:Experiments in three dimensions. Mon. Wea. Rev.,109, 564–576.

  • Hauser, D., and P. Amayenc, 1986: Retrieval of cloud water and water vapor contents from Doppler radar data in a tropical squall line. J. Atmos. Sci.,43, 823–838.

  • ——, ——, and M. Chong, 1984: A new optical instrument for simultaneous measurement of raindrop diameter and fallspeed distributions. J. Atmos. Oceanic Technol.,1, 256–269.

  • Houze, R. A., Jr., 1977: Structure and dynamics of a tropical squall-line system. Mon. Wea. Rev.,105, 1540–1567.

  • Jabouille, P., J. L. Redelsperger, and J. P. LaFore, 1996: Modification of surface fluxes by atmospheric convection in the TOGA COARE region. Mon. Wea. Rev.,124, 816–837.

  • Johnson, R. H., and M. E. Nicholls, 1983: A composite analysis of the boundary layer accompanying a tropical squall line. Mon. Wea. Rev.,111, 308–319.

  • Jorgensen, D. P., and M. A. LeMone, 1989: Vertical velocity characteristics of oceanic convection. J. Atmos. Sci.,46, 621–640.

  • ——, and B. F. Smull, 1993: Mesovortex circulations seen by airborne Doppler radar within a bow-echo mesoscale convective system. Bull. Amer. Meteor. Soc.,74, 2146–2157.

  • ——, P. H. Hildebrand, and C. L. Frush, 1983: Feasibility test of an airborne pulse-Doppler meteorological radar. J. Climate Appl. Meteor.,22, 744–757.

  • ——, M. A. LeMone, and B. J.-D. Jou, 1991: Precipitation and kinematic structure of an oceanic mesoscale convective system. Part I: Analysis of airborne Doppler radar data. Mon. Wea. Rev.,119, 2608–2637.

  • ——, T. Matejka, and J. D. DuGranrut, 1996: Multi-beam techniques for deriving wind fields from airborne Doppler radars. J. Meteor. Atmos. Phys.,59, 83–104.

  • Keenan, T. D., and R. E. Carbone, 1992: A preliminary morphology of precipitation systems in northern Australia. Quart. J. Roy. Meteor. Soc.,118, 283–326.

  • Leise, J. A., 1981: A multidimensional scale-telescoped filter and data extension package. NOAA Tech. Memo. ERL WPL-82, 18 pp. [NTIS PB82-164104.].

  • LeMone, M. A., and D. P. Jorgensen, 1991: Precipitation and kinematic structure of an oceanic mesoscale convective system. Part II: Momentum transport and generation. Mon. Wea. Rev.,119, 2638–2653.

  • ——, G. M. Barnes, E. J. Szoke, and E. J. Zipser, 1984a: The tilt of the leading edge of mesoscale tropical convective lines. Mon. Wea. Rev.,112, 510–519.

  • ——, ——, and E. J. Zipser, 1984b: Momentum flux by lines of cumulonimbus over the tropical oceans. J. Atmos. Sci.,41, 1914–1932.

  • Matejka, T., and S. A. Lewis, 1997: Improving research aircraft navigation by incorporating INS and GPS information in a variational solution. J. Atmos. Oceanic Technol.,14, 495–511.

  • O’Brien, J. J., 1970: Alternative solutions to the classical vertical velocity problem. J. Appl. Meteor.,9, 197–203.

  • Ogura, Y., and M.-T. Liou, 1980: The structure of a midlatitude squall line. A case study. J. Atmos. Sci.,37, 553–567.

  • Parsons,D., and Coauthors, 1994: The integrated sounding system: Description and preliminary observations from TOGA COARE. Bull. Amer. Meteor. Soc.,75, 553–567.

  • Rotunno, R., and J. B. Klemp, 1982: The influence of the shear-induced pressure gradient on thunderstorm motion. Mon. Wea. Rev., 110, 136–151.

  • Roux, F., 1988: The West African squall line observed on 23 June 1981 during COPT 81: Kinematics and thermodynamics of the convective region. J. Atmos. Sci.,45, 406–426.

  • Rutledge, S. A., R. A. Houze Jr., M. I. Biggerstaff, and T. J. Matejka, 1988: The Oklahoma–Kansas mesoscale convective system of 10–11 June 1985: Precipitation structure and single-Doppler radar analysis. Mon. Wea. Rev.,116, 1409–1430.

  • Schmidt, J. M., and W. R. Cotton, 1989: A high plains squall line associated with severe surface winds. J. Atmos. Sci.,46, 281–302.

  • Scott, J. D., and S. A. Rutledge, 1995: Doppler radar observations of an asymmetric mesoscale convective system and associated vortex couplet. Mon. Wea. Rev.,123, 3437–3457.

  • Skamarock, W. C., M. L. Weisman, and J. B. Klemp, 1994: Three-dimensional evolution of simulated long-lived squall lines. J. Atmos. Sci.,51, 2563–2584.

  • Smull, B. F., and R. A. Houze Jr., 1987: Dual-Doppler radar analysis of a midlatitude squall line with a trailing region of stratiform rain. J. Atmos. Sci.,44, 2128–2148.

  • Sun, J., and R. A. Houze Jr., 1992: Validation of a thermodynamic retrieval technique by application to a simulated squall line with trailing stratiform precipitation. Mon. Wea. Rev.,120, 1003–1018.

  • Szoke, E. J., E. J. Zipser, and D. P. Jorgensen, 1986: A radar study of convective cells in mesoscale systems in GATE. Part I: Vertical profile statistics and comparison with hurricanes. J. Atmos. Sci.,43, 182–197.

  • Trier, S. B., W. C. Skamarock, M. A. LeMone, D. B. Parsons, and D. P. Jorgensen, 1996: Structure and evolution of the 22 February 1993 TOGA COARE squall line: Numerical simulations. J. Atmos. Sci.,53, 2861–2886.

  • ——, ——, and ——, 1997: Structure and evolution of the 22 February 1993 TOGA COARE squall line: Organization mechanisms inferred from numerical simulation. J. Atmos. Sci.,54, 386–407.

  • Verlinde, J., and W. R. Cotton, 1990: A mesoscale vortex couplet observed in the trailing anvil of a multicellular convective complex. Mon. Wea. Rev.,118, 993–1010.

  • Webster, P. J., and R. Lukas, 1992: TOGA COARE: The Coupled Ocean–Atmosphere Response Experiment. Bull. Amer. Meteor. Soc.,73, 1377–1416.

  • Weisman, M. L., 1993: The genesis of severe, long-lived bow echoes. J. Atmos. Sci.,50, 645–670.

  • Yuter, S. E., and R. A. Houze Jr., 1995: Three-dimensional kinematic and microphysical evolution of Florida cumulonimbus. Part II: Frequency distributions of vertical velocity, reflectivity, and differential reflectivity. Mon. Wea. Rev.,123, 1941–1963.

  • ——, ——, B. F. Smull, F. D. Marks, and J. R. Daugherty, 1995: TOGA COARE aircraft mission summary images: An electronic atlas. Bull. Amer. Meteor. Soc.,76, 319–328.

  • Zhang, D.-L., K. Gao, and D. B. Parsons, 1989: Numerical simulations of anintense squall line during 10–11 June 1985 PRE-STORM. Part I: Model verification. Mon. Wea. Rev.,117, 960–994.

  • Zipser, E. J., 1977: Mesoscale and convective-scale downdrafts as distinct components of squall-line circulation. Mon. Wea. Rev.,105, 1568–1589.

  • ——, and M. A. LeMone, 1980: Cumulonimbus vertical velocity events in GATE. Part II: Synthesis and model core structure. J. Atmos. Sci.,37, 2458–2469.

  • ——, and K. R. Lutz, 1994: The vertical profile of radar reflectivity of convective cells: A strong indicator of storm intensity and lightning probability? Mon. Wea. Rev.,122, 1751–1759.

  • View in gallery

    GMS infrared satellite photograph from 2132 UTC 22 February 1993. Cloud top temperatures (K) are contoured with the color scale at the bottom of the figure. The island of Guadalcanal in the Solomon Islands is shown as the dark line with the sounding site at Honiara (HIR) indicated as the cross. Aircraft tracks of the two National Oceanic and Atmospheric Administration (NOAA) aircraft for the period near 2130 UTC are show as the green (N42) and blue (N43) parallel lines near the eastern edge of the cloud shield.

  • View in gallery

    Composite radar reflectivity (dBZ) over the time periods (a) 2057–2103 UTC, (b) 2141–2147 UTC, (c) 2219–2234 UTC, and (d) 2302–2305 UTC from the P-3’s lower fuselage C-band radar. Each of the four panels is approximately 45 min apart. The flight tracks of the aircraft during the composite period are indicated by the lines with arrows. The composites were constructed by maximizing the reflectivity from each sweep at each grid point. Color scale for radar reflectivity is shown to the right of each panel. The domain of analysis is 240 × 240 km2. Islands are shown by the bold lines, with the largest island extending from the right side of each figure being Guadalcanal, Solomon Islands. Small boxes with tick marks in all but the 2300 UTC map are the Doppler analysis domains. Points labeled “A” and “B” refer to transverse precipitation bands.

  • View in gallery

    Composite sounding (top panel) and hodograph (bottom panel) for the environment of the 22 February squall line. Data is from both P-3 aircraft and the balloon ascent at Honiara, Solomon Islands. See text for a description of the compositing procedure. The system motion vector (eastward at 12 m s−1) is shown as the arrow starting from u = 0, υ = 0.

  • View in gallery

    (a) Horizontal analysis of radar reflectivity and system-relative wind flow and (b) vertical air motion (m s−1) at 1.5 km MSL valid at 2116 UTC. Domain is 75 km × 75 km, the location of which is shown in Fig. 2a. Reflectivity contour gray scale and the 10 m s−1 scaling vector for winds are shown to the right and above (a) respectively. The box located near the center of both panels is the domain over which a line average cross section was constructed.

  • View in gallery

    As in Fig. 4, except for 9.5 km MSL.

  • View in gallery

    As in Fig. 4, except for (a) perturbation pressure (in 10−1 mb) at 1.5 km MSL and (b) perturbation buoyancy (in 10−1 K) at 1.0 km MSL retrieved from the airborneDoppler-derived wind fields.

  • View in gallery

    Vertical cross section of average reflectivity and average line-normal system relative wind in (a) the plane of the cross section and (b) vertical motion (m s−1) averaged within the boxed domain of Fig. 4. In the construction of the line average, the coordinate system was rotated so that u was aligned in the direction of x′ and υ was aligned in the direction of y′ as shown in Fig. 4b. The symbols A, B, C, and D refer to the locations of various maxima in the vertical velocity, pressure perturbation, and buoyancy fields.

  • View in gallery

    As in Fig. 7, except for line average vertical cross sections of (a) perturbation pressure (in 10−1 mb) and (b) perturbation buoyancy (in 10−1 K).

  • View in gallery

    Horizontal analysis of radar reflectivity and system-relative wind flow at 1.5 km MSL valid at (a) 2100 UTC, (b) 2116 UTC, (c) 2135 UTC, (d) 2154 UTC, (e) 2203 UTC, (f) 2227 UTC, and (g) 2245 UTC. The analysis domains of (a) and (f) are shown on the larger scale radar view of Figs. 2a and 2c.

  • View in gallery

    Horizontal analysis of radar reflectivity and system-relative wind flow at 4.5 km MSL at 2227 UTC. Thin black line oriented approximately NNE to SSW near the center of the figure is the location of the cross section shown in Fig. 11.

  • View in gallery

    Vertical cross section of reflectivity and (a) system-relative wind (dBZ) and (b) vertical air velocity (m s−1) in the plane of the cross through the two transverse bands along the cross-section line shown in Fig. 10. Horizontal distance of the cross section is 53.7 km.

  • View in gallery

    Vertical cross section of perturbation D-value (m, the difference between radar altitude and pressure altitude determined from an environmental sounding) obtained by (a) N43 and (b) N42 during several east–west transects normal to the convective line (bottom panel of each figure). The environmental D value at each height was subtracted to yield perturbation values. Starting and ending times (UTC) of each flight leg segment are indicated at the end points of each leg. Arrows denote direction of flight. The top panel of each figure shows the system-relative horizontal winds from the topmost flight leg (at 940 m MSL) segment. The aircraft were separated by about 0.3° of latitude (∼35 km).

  • View in gallery

    Time series from 2148 to 2156 UTC of in situ data collected by N42 (the easternmost flight track in Fig. 9d) as it proceeded parallel to the convective leading edge at 150 m MSL. Total distance covered is about 60 km. From top to bottom the time series are east–west ground-relative wind speed (u, m s−1), and north–south ground-relative wind speed (υ, m s−1), equivalent potential temperature (θe, K), potential temperature (θ, K), mixing ratio (q, g kg−1), and cloud water mixing ratio (ql, g kg−1) observed by the Johnson–Williams (JW) hot-wire instrument. For the θe, θ,and q plots the dashed lines are raw data; the solid lines are adjusted for the effects of sensor wetting by liquid water by replacing temperature with bias-adjusted temperature from the side-looking CO2 radiometer; dewpoints in excess of this temperature are set equal to it. The delay in warming and cooling of “adjusted” θ results from the slow response of the radiometer. Similarly, the warm temperature spike in the cold air is probably more accurate than the radiometric value. The raw values should be taken as “correct” where ql = 0. The aircraft penetrated the gust front during the time approximated by the arrow on the θ plot. During the penetration, the liquid water content ql rose from zero to about 0.4–0.5 g kg−1.

  • View in gallery

    Time series of in situ data obtained by N43 as it flew at 150 m MSL from south to north and parallel to and ∼50 km to the rear (west) of the convective line from 2129 to 2134 UTC. Total distance is about 40 km. From top to bottom, the time series are equivalent potential temperature (θe, K), potential temperature (θ, K), mixing ratio (q, g kg−1), and east–west wind velocity (u, m s−1). The aircraft penetrated the rearward moving (in a relative sense) convective-scale downdraft air between the arrows indicated on the θ time series. The track is indicated as the westernmost track in Fig. 9d.

  • View in gallery

    Scatter plots of earth-relative east–west wind velocity (u, m s−1) versus potential temperature (θ, K, top panel), mixing ratio (q, g kg−1, middle panel), and equivalent potential temperature (θe, K, bottom panel) from flight legs behind the convective line by N43. The vertical line at u = 12 m s−1 marks the line motion and hence the division between rear-to-front and front-to-rear moving air.

  • View in gallery

    Horizontal depiction of radar reflectivity derived from a composite of N42 lower fuselage radar scans over the time period 2336 to 2356 UTC. The flight tracks of the two P-3 aircraft are shown as the lines with arrows and wind barbs. Wind barbs depict ground-relative flow and are plotted every 60 s with one full barb equal to 5 m s−1. Domain size is 240 km × 240 km. The box with tick marks is the domain of the dual-Doppler analysis. Color scale for reflectivity (dBZ) is shown to the right.

  • View in gallery

    Time series of in situ data obtained by N43 at 150 m MSL between 2327 and 2356 UTC along the southernmost flight track shown in Fig. 16. From top to bottom are the time series of mixing ratio (q, g kg−1), potential temperature (θ, K), equivalent potential temperature (θe;t2, K), and cloud water content (g kg−1) measured by the Johnson–Williams hot-wire instrument, the line-relative wind normal to the leading edge and along the flight track (us, m s−1), storm-relative wind parallel to leading edge and normal to the flight track (υs, m s−1), and total wind speed relative to the ground (m s−1). Vertical dashed lines denote relative distance (km) from the leading edge. Thin vertical lines indicate the locations of the relative minima inθe. Line motion at this time was toward 55° at 12 m s−1. No bias adjustment has been applied to these time series.

  • View in gallery

    Schematic diagram of air flow contributing to the minimum in θe seen in the flight-level data at 150 m MSL. Convective line is toward the right of the schematic. The effects of surface flux and entrainment are included in the deepening of the boundary layer toward the rear of the system.

  • View in gallery

    Domain-averaged vertical velocity (cm s−1) as a function of height over the Doppler analysis area shown in Fig. 16.

  • View in gallery

    Vertical profiles of mean mixing ratio (q g kg−1, top panel) and mean potential temperature (θ, K, bottom panel) collected by N43. The solid lines (labeled “a”) are values characteristic of the environment ahead of the line sampled by N42. Dashed lines (labeled “b”) are values taken at the θe minimum point ∼120 km behind the convective line while the dash-dotted lines (labeled “c”) are values ∼20 km behind the line along N43’s flight track. The shaded line with boxes is the environmental q profile taken by N42 ahead of the line and “adjusted” for possible biases in temperature and dewpoint between aircraft that intercomparisons indicate may be 0.5°C for dewpoint and 0.1°C for temperature (N43 lower than N42). To be consistent with Fig. 21, values of q should be raised by 0.6 g kg−1.

  • View in gallery

    Bulk surface fluxes computed from the flight-level data obtained by N43 at 150 m MSL (Fig. 17). Solid line represents latent heat fluxes (W m−2), while the dashed line is sensible heat fluxes. Fluxes were computed using the COARE 1.0 bulk flux routine (Fairall et al. 1996) with sea surface temperatures and temperatures from N43 taken as “correct,” and N43 dewpoint temperatures raised by 0.5 K. For this version of the aircraft dataset (available from NOAA/National Severe Storms Laboratory, Boulder, the dewpoint correction is consistent with the Burns–Friehe “best guess” values (S. Burns 1996, personal communication).

All Time Past Year Past 30 Days
Abstract Views 0 0 0
Full Text Views 21 21 1
PDF Downloads 12 12 1

Structure and Evolution of the 22 February 1993 TOGA COARE Squall Line: Aircraft Observations of Precipitation, Circulation, and Surface Energy Fluxes

View More View Less
  • 1 NOAA/NSSL/Mesoscale Research Division, Boulder, Colorado
  • | 2 National Center for Atmospheric Research, Boulder, Colorado
© Get Permissions
Full access

Abstract

This study documents the precipitation and kinematic structure of a mature, eastward propagating, oceanic squall line system observed by instrumented aircraft during the Tropical Ocean Global Atmosphere Coupled Ocean–Atmosphere Response Experiment (TOGA COARE). Doppler radar and low-level in situ observations are used to show the evolution of the convection from an initially linear NNW–SSE-oriented convective line to a highly bow-shaped structure with an embedded low- to midlevel counterclockwise rotating vortex on its northern flank. In addition to previously documented features of squall lines such as highly upshear-tilted convection on its leading edge, a channel of strong front-to-rear flow that ascended with height over a “rear-inflow” that descended toward the convective line, and a pronounced low-level cold pool apparently fed from convective and mesoscale downdrafts from the convective line; rearward, the observations of this system showed distinct multiple maxima in updraft strength with height and reflectivity bands extending rearward transverse to the principal convective line. Vertical motions within the active convective region of the squall line system were determined using a new approach that utilized near-simultaneous observations by the Doppler radars on two aircraft with up to four Doppler radial velocity estimates at echo top. Echo-top vertical motion can then be derived directly, which obviates the traditional dual-Doppler assumption of no vertical velocity at the top boundary and results in a more accurate estimate of tropospheric vertical velocity through downward integration of horizontal divergence.

Low-level flight-level observations of temperature, wind speed, and dew point collected rearward of the squall line are used to estimate bulk fluxes of dry and moist static energy. The strong near-surface fluxes, due to the warm sea and high winds, combined with estimates of mesoscale advection, are used to estimate boundary layer recovery time; they indicate that the boundary layer could recover from the effects of the cold dome within about 3 h of first cold air injection if the observed near-surface winds were maintained. However, the injection and spreading of air from above leads to cooling at a fixed spot ∼20 km rearward of the convective line (surface θe minimum point), suggesting that the cold pool could be still intensifying at the time of observation. Recovery time at a point is probably similar to that measured in previous studies.

Corresponding author address: Dr. David P. Jorgensen, NOAA/NSSL/Mesoscale Research Division, Mail Code: N/C/MRD, 325 Broadway, Boulder, CO 80303.

Email: davej@ncar.ucar.edu

Abstract

This study documents the precipitation and kinematic structure of a mature, eastward propagating, oceanic squall line system observed by instrumented aircraft during the Tropical Ocean Global Atmosphere Coupled Ocean–Atmosphere Response Experiment (TOGA COARE). Doppler radar and low-level in situ observations are used to show the evolution of the convection from an initially linear NNW–SSE-oriented convective line to a highly bow-shaped structure with an embedded low- to midlevel counterclockwise rotating vortex on its northern flank. In addition to previously documented features of squall lines such as highly upshear-tilted convection on its leading edge, a channel of strong front-to-rear flow that ascended with height over a “rear-inflow” that descended toward the convective line, and a pronounced low-level cold pool apparently fed from convective and mesoscale downdrafts from the convective line; rearward, the observations of this system showed distinct multiple maxima in updraft strength with height and reflectivity bands extending rearward transverse to the principal convective line. Vertical motions within the active convective region of the squall line system were determined using a new approach that utilized near-simultaneous observations by the Doppler radars on two aircraft with up to four Doppler radial velocity estimates at echo top. Echo-top vertical motion can then be derived directly, which obviates the traditional dual-Doppler assumption of no vertical velocity at the top boundary and results in a more accurate estimate of tropospheric vertical velocity through downward integration of horizontal divergence.

Low-level flight-level observations of temperature, wind speed, and dew point collected rearward of the squall line are used to estimate bulk fluxes of dry and moist static energy. The strong near-surface fluxes, due to the warm sea and high winds, combined with estimates of mesoscale advection, are used to estimate boundary layer recovery time; they indicate that the boundary layer could recover from the effects of the cold dome within about 3 h of first cold air injection if the observed near-surface winds were maintained. However, the injection and spreading of air from above leads to cooling at a fixed spot ∼20 km rearward of the convective line (surface θe minimum point), suggesting that the cold pool could be still intensifying at the time of observation. Recovery time at a point is probably similar to that measured in previous studies.

Corresponding author address: Dr. David P. Jorgensen, NOAA/NSSL/Mesoscale Research Division, Mail Code: N/C/MRD, 325 Broadway, Boulder, CO 80303.

Email: davej@ncar.ucar.edu

1. Introduction

One of the principal objectives of TOGA COARE (Tropical Ocean Global Atmosphere Coupled Ocean–Atmosphere Response Experiment; Webster and Lukas 1992) is to understand the processes that control and organize convection in the warm pool and the effect ofthe convection on larger scales of atmospheric circulation. A primary method of accomplishing this objective was with the in situ sensors and airborne Doppler radars installed on three turboprop aircraft, the two WP-3Ds operated by the National Oceanic and Atmospheric Administration (NOAA) and the Electra aircraft operated by the National Center for Atmospheric Research. These highly mobile platforms allow for the collection of Doppler datasets that describe the precipitation structure, its evolution, and basic horizontal and vertical circulations within the mesoscale convective systems (MCSs) that were observed during TOGA COARE.

In this study we describe the structural and airflow characteristics of an MCS that is from a class of systems described in previous studies as squall lines. The general characteristics of tropical squall lines were documented initially using aircraft and sounding data by Zipser (1977) and Houze (1977), and later from ground-based dual-Doppler radars by Chong et al. (1987). Structurally, tropical squall lines and their midlatitude counterparts are very similar, as documented by Ogura and Liou (1980), Rutledge et al. (1988), Biggerstaff and Houze (1991), and many others. Their relatively simple structure usually consists of a rapidly moving leading convective line that slopes rearward with height, 30–50 km behind which can be a broad secondary and more stratiform rainfall maximum. In between the two rainfall maxima can be a “transition zone” of weaker surface precipitation (Biggerstaff and Houze 1993; Braun and Houze 1994). The fastest moving squall lines are generally associated with an environmental low-level wind speed jet with the convective line oriented roughly perpendicular to the low-level wind shear (e.g., surface to 800 mb). Speed of propagation of the leading edge is generally near the speed of the low-level jet maximum (Keenan and Carbone 1992; Barnes and Sieckman 1984). The principal system-relative, two-dimensional flows at the mature stage of squall line evolution consist of a near-surface front-to-rear flow of low-θe (equivalent potential temperature) air originating at the convective line, a rear-to-front current of dry midlevel air that slopes downward toward the surface near the leading edge, and a front-to-rear flow of generally high θe air originating at low levels ahead of the leading edge and extending rearward at the upper levels in the stratiform rain region.

This basic description of the squall line circulation has been augmented greatly by recent three-dimensional nonhydrostatic simulations of midlatitude squall lines by Weisman (1993) and Skamarock et al. (1994), theoretical arguments by Davis and Weisman (1994), and the advent of airborne Doppler radars that can collect pseudo-dual-Doppler data (Jorgensen and Smull 1993). These three-dimensional studies emphasized the role played by the descending rear inflow and ascending front-to-rear flows in twisting and converging environmental and MCS-generated shear to produce counterrotating circulations at low to mid levels, at the ends of the convective line. In extreme conditions of line-normal shear, short (∼50–100 km) convective line segments can produce strong vortices and lead to strong rear-inflow jets that contribute to the production of severe surface winds as well as the characteristic convex shape of the convective line [leading to the term “bow echoes”; see Schmidt and Cotton(1989)]. To date, however, very little Doppler radar data have been collected in tropical squall lines to verify simulations of vortex generation in such systems. Such cyclonic vortices, induced by convection and sustained and grown to larger scales by mesoscale circulations and/or vertical gradients of microphysically induced heating within the stratiform region, may be key to the production of incipient cyclonic vorticity for tropical cyclone genesis. Indeed, several days after the aircraft’s investigation of this system, satellite images showed the development of a strong tropical cyclone (Tropical Cyclone Polly) from the region of cloudiness that had been this system, which caused considerable damage to the island of Fiji. Unfortunately, it is not possible to track circulation centers back to the vortices seen at the time of the aircraft’s investigation. Indeed, the strongest circulation seen by the aircraft was anticyclonic, although there were indications of a companion cyclonic vortex at the south end of the line.

In this study analyses of airborne Doppler radar data have been performed over an extended period of a squall line life cycle. These analyses reveal the development of counterrotating vortices (particularly the anticyclonic, or counterclockwise in the Southern Hemisphere, member on the northern or equatorward end of the convective line) and the coincident development of a eastward extending “bow” in the leading convective line from an initially linear structure. Many of the structural features of this tropical squall line system were successfully replicated in numerical simulations performed by Trier et al. (1996), including the observations of the line-end vortices, the secondary updraft maxima in the upper troposphere located rearward of the surface position of the leading edge, and westward extending transverse bands of enhanced rainfall, the most pronounced of which was near the southern end of the bow.

The system studied here is representative of a class of mesoscale convective systems that are generally oriented approximately normal to the low-level (i.e., surface to about 800 mb) shear vector and propagate near the speed of a low-level “jet” in the environmental wind profile. This system was, admitedly, one of the fastest moving sytems seen in COARE with the most pronounced 800-mb westerly jet maximum in the environmental wind profile. However, a close examination of the composite wind soundings and radar data for all COARE convective events revealed that about half of all convective lines observed by aircraft or ship-based radar in COARE had an orientation relative to the low-level flow like that of the 22 February MCS. Therefore, it is felt that the 22 February system is representive of COARE convective lines.

Section 2 presents an overview of the system and its evolution. Section 3 presents the Doppler analysis procedures that are believed to be a significant improvement in estimating vertical motions over the traditional dual-Doppler approach of assuming zero vertical motion at echo top. Section 4 presents the precipitation and air flow structure and their evolution. Because squall lines are known to produce extensive cooling and drying of the boundary layer air from both mesoscale and convective-scale downdrafts, we calculate surface fluxes and boundary layer recovery time in section 5 to quantify how rapidly the boundary layer can recover from gust front passage over the very warm waters of the warm pool region.

2. Description of the convective system, its evolution, and its environment

On 22–23 February 1993 the two NOAA P-3 aircraft, designated N42 and N43, (also the DC-8 and ER-2 aircraft operated by the National Aeronautics and Space Administration) intercepted a rapidly eastward propagating north–south orientedconvective line near Guadalcanal Island in the Solomon Islands. Figure 1 presents an infrared satellite view at the beginning of the aircraft investigation. This convective system was located at the leading edge (east side) of a disturbed region associated with pronounced divergent anticyclonic flow at cloud top in a region with strong low-level westerly winds (K. Boyd 1994, personal communication). For over 5 h the two airborne Doppler-equipped aircraft flew highly coordinated, low-level patterns to examine the convective line and its immediate environment. Summaries of the flight tracks can be found in the electronic atlas cited by Yuter et al. (1995).

The squall line persisted from at least ∼1950 UTC when it was first penetrated by N43 en route to another target farther west, to after 0100 UTC the following day after it passed over Guadalcanal. Geostationary Meteorological Satellite (GMS) imagery also shows cold cloud tops indicative of deep convection at 2132 UTC. The line at 1950 UTC was linear and vigorous [strongest 1 s or ∼100 m resolved vertical velocities seen in the in situ data observed during the mission were seen on this first pass at 1920 above mean sea level (MSL) and were a 13.5 m s−1 updraft and an 8.0 m s−1 downdraft], shallow with radar tops at only about 8 km, and very narrow (<5 km), the last two characteristics suggesting recent formation. Radar depiction of the squall line system from 2057 UTC to 2305 UTC taken from the horizontally scanning lower fuselage (LF) C-band radar from N43 are shown in Fig. 2. Although the squall line was still linear (top left panel at 2057 UTC), it later developed a pronounced “bow” in the leading convective line (Fig. 2b) in the next 1.5 h. The initial focus of the aircraft investigation was on the segment of the convective zone that bowed rapidly toward the east. The two P-3 aircraft executed flight patterns that enabled near-simultaneous measurements of dual-beam radar data (called “quad-Doppler” patterns because up to four radial velocity estimates per grid point could be obtained) from two parallel straight line flight tracks spaced about 30–50 km apart. Line motion during this period was estimated from the sequence of radar reflectivity maps at approximately 12 m s−1 toward the east (090°), and it was that motion that was used as the reference frame motion for the Doppler analysis that focused on the convective line. Later, as the southern part of the convective line neared and then passed over the island of Guadalcanal focus shifted farther north to a portion of the line moving 55° at 12 m s−1. The aircraft then reoriented their patterns to fly normal to the convective line.

An interesting aspect of this convective system is the appearance of one or more bands of reflectivity maxima oriented approximately transverse to the convective line. The first band of echo (labeled “A” in Fig. 2b) in the southern part of the pseudo-dual-Doppler domain initially appeared relatively narrow on horizontal radar depictions and exhibited vertical characteristics associated with stratiform precipitation (e.g., pronounced reflectivity “bright band” near the melting level and weak ascent above). As the vortex intensified and the bowing part of the line surged eastward, an east–west kink in the line developed along the northern flank of the vortex (labeled “B” in Fig. 2c). In spite of the bow development, the forward motion of the main body of the line south of the vortex remained about the same as before the bow. The bow was largely manifested as the sudden development of an east–west band north of the vortex core.

The squall line system studied hereexhibited many of the same characteristics seen in previous studies (e.g., the conceptual model of Smull and Houze 1987). At the system’s leading edge (the eastern side), a line of deep convective storms was organized in a multicellular fashion that resulted from discrete propagation; that is, new cells developed ahead of the line of mature cells because of low-level convergence from the outflowing gust front. We will later show the character of this outflow from in situ penetration data. During the linear phase of the system, far to the rear (>50 km from the leading edge) a region of weak stratiform precipitation was seen. This region of stratiform precipitation (seen as a bright band of enhanced reflectivity near the melting level) was not as extensive or intense as in previously observed squall lines, which may be due to the rapid decay of the convective line as it passed over Guadalcanal Island after only about 2 h from its initial development.

The environmental sounding, shown in Fig. 3, shows moderate instability [convective available potential energy (CAPE) was about 1440 J kg −1] and low-level wind shear (800 mb to surface shear of east–west wind u was about 13 m s−1 or about 6.7 m s−1 km−1). The u maximum at 800 mb is nearly identical to the convective line propagation speed. The sounding is a composite designed to apply to 2100 UTC. The upper-level thermodynamic parameters (above about 6 km) are an average of the rawinsonde ascent from Honiara, Guadalcanal, (HIR) at 1800 and 2400 UTC, but winds are from 1800 UTC because of their close resemblance to those from the upper levels of the last N43 sounding, taken just ahead of the line from 0105 to 0130 UTC 23 February From 1.3 to 6 km, a time interpolation was performed from the N43 aircraft sounding and the HIR 1800 UTC sounding for wind, and the 1800 UTC and 2400 UTC HIR sounding for thermodynamic data. The N43 dewpoints were adjusted upward by 0.5°C to correct for a bias between N42 and N43 (see section 5 for a description of the method for bias determination). The lowest 1.3 km of the sounding derives from a sounding by N42 just ahead of the line around 2100 UTC and straight and level leg segments flown by N42 ahead of the system at 150 and 300 m.

3. Data and processing methodology

a. Description of the radar and antenna scanning methodology

The Doppler radar installed on N42 and N43 is described in detail in Jorgensen et al. (1983). The radar is a vertically scanning X-band (3.2-cm wavelength) Doppler radar mounted in the aircraft’s tail. The basic scanning methodology is termed the Fore–Aft Scanning Technique (FAST). This scanning procedure consists of collecting two scans of data, one pointing ∼25° forward from a plane normal to the flight track and one scan pointing aft ∼25°. As the aircraft moves in a quasi-straight flight path, FAST scans sweep out a three-dimensional region of space surrounding the aircraft’s track. If the antenna is rotated at its maximum rate of 10 rpm (60° s−1), at typical P-3 ground speeds (∼120 m s−1) adjacent beams intersect with horizontal spacing of ∼1.4 km. An improvement in the horizontal data spacing to ∼750 m can be obtained if the antenna is scanned in a sector confined to one side of the track. The increase in horizontal data density comes at the expense of the areal coverage on both sides of the track.1 Since the start times of most flight legs were synchronized to within 2 min (a few of the legs were within 30 s) most of the time delay between beams arises because of the time delay between intersecting forward- and aft-looking beams (Jorgensen et al.1996). For the P-3 antenna this delay is roughly 1 min for each 10-km range from the aircraft’s track. For two flight tracks separated by 40 km, the range to a midpoint would imply a time delay of 2 min. Whatever system evolution occurs during this time delay (plus any delay arising because of different flight leg start times) would produce an error inherent in the analysis. System propagation is handled through advecting the grid with the system motion vector (12 m s−1 to the east).

b. Determination of horizontal and vertical velocities

Where two radar beams intersect, an estimate of the horizontal wind velocity can be obtained using the traditional “dual-Doppler” overdetermined synthesis methods described in Ray et al. (1985). Vertical velocity can then obtained by downward (or upward) integration of the equation of continuity subject to boundary conditions near the surface and echo top, using an O’Brien (1970) correction to the divergence profile. This traditional method of vertical velocity determination using vertical divergence integration uses a zero vertical velocity assumption at echo top. This procedure is known to be in error in regions of active convection because often what is detected at echo top is not really updraft top due to radar insensitivity (Jorgensen et al. 1991), and therefore the echo-top vertical motion could be substantially different from zero. Should echo-top vertical motion be near zero, such as within the stratiform region of an MCS, the dual-Doppler approach should work well. To alleviate this problem in TOGA COARE, where we expected to see deep convective verical velocities, however, a special flight strategy was designed to utilize two Doppler-equipped aircraft to gather two sets of FAST data nearly simultaneously. This strategy (called “quad-Doppler” since up to four Doppler observations are made at each point) required that the aircraft fly parallel flight legs bracketing the area of interest, spaced by 30–50 km. The echo-top vertical air velocity can then be estimated from a “triple-Doppler” solution for the wind field (Jorgensen et al. 1996) using an empirical relationship to remove ice particle terminal velocity from radar reflectivity. The echo-top vertical velocity is then used as a boundary condition to start the downward integration of horizontal divergence to get the column vertical velocity, followed by the O’Brien (1970) correction to zero the surface vertical motion. This method of vertical integration is called a hybrid analysis method since it combines the traditional dual-Doppler scheme with a triple-Doppler solution. The hybrid technique produces about 50% stronger updrafts in the convective line region in TOGA COARE compared to those found applying the traditional assumption of zero vertical velocity at echo top. A comparison with NASA DC-8 in situ vertical velocities obtained during a penetration of the 22 February squall line coincident with quad-Doppler measurements reveals much closer agreement between the in situ vertical velocities and the hybrid technique compared to velocities derived from the dual-Doppler methodology (Jorgensen et al. 1996), lending credence to the validity of using the quad-Doppler methodology in TOGA COARE. To filter noise in the vertical velocities, the horizontal wind fields and the vertical winds derived from the triple Doppler solution at echo top are lightly smoothed with a two-pass Leise filter (Leise 1981) prior to divergence calculation and vertical integration. This procedure dampens features with wavelengths less than about 3–4 times the horizontal grid spacing (1.5 km for this study), which implies that features with wavelengths smaller than 4.5–6.0 km are removed. Theanalysisregions are generally 75 × 75 km on a horizontal grid of 1.5 km and vertical grid of 500 m, centered on the reflectivity maximum of the squall line. The lowest analysis level is 500 m. To partially compensate for the effects of attenuation of the radar beam by intervening precipitation, radar reflectivity is calculated by maximizing the estimates from all the radar gates that pass through a given Cartesian grid point (i.e., up to four beams). Attenuation will mostly affect the estimation of water loading in the buoyancy calculation, with the consequence that the water loading would be underestimated and the buoyancy fields (e.g., Fig. 8) would be too weak. In the course of determining radar biases for TOGA COARE, extensive intercomparison of point values of maximum reflectivity were made for all COARE flights between the aircraft (tail and lower fuselage) and ship radars. The result of this intercomparison indicated that maximum reflectivity values generally differed by less than 2–3 dBZ and could generally be explained by target distances from each radar. If the attenuation were as large as 3 dBZ, the reduction in estimated water content and hence buoyancy, according to Eq. (1), would be 24/7 or about 50%. Reflectivity is also used in the wind synthesis to remove terminal fallspeeds from radial velocities. Attenuation is not felt to be a serious problem for wind determination since the aircraft were flying at very low altitudes and attenuation would be greatest at very flat radar elevation angles, where the fallspeed contamination is minimal anyway.

It generally took about 5–8 min to complete each set of flight legs. Over this time, the system is assumed to be stationary.

c. Aircraft positions and ground speeds

Accurate navigation of each aircraft (to an absolute position within ±250 m) is required to properly combine the four radial velocity estimates. In addition, removal of the aircraft motion from the Doppler radial velocities requires that the aircraft’s ground speed be known to an accuracy of better than 0.5 m s−1. The standard aircraft navigation (inertial navigation system, INS) in TOGA COARE was augmented by Global Positioning System (GPS) information. Although normally the GPS positions are sufficiently accurate to properly position the aircraft (the GPS does not suffer from accelerometer drift, or the Shuler oscillation, as does the INS), during TOGA COARE the GPS observations exhibited data gaps and slow updating that resulted in spikes in the ground speed time series. Dropouts most commonly occurred during aircraft turns; it would often take as much as 5 min after the turn was completed before the system locked on to the required number of satellites and navigation was restored. A procedure was developed to merge the accurate ground speed measurements of the INS with the accurate position information of the GPS using a variational adjustment procedure to interpolate over the GPS position gaps (Matejka and Lewis 1997). This procedure produced an accurate dataset of aircraft positions and ground speeds that are mutually consistent; that is, differentiation of the positions produces the ground speeds.

d. Calculation of pressure and buoyancy perturbations

Interpretation of dynamic circulation features of mesoscale convective systems is aided through the determination of perturbation pressure and buoyancy fields. Here the term “perturbation” refers to deviations from horizontal averages. The method used to calculate these fields is termed dynamic retrieval and is based originally on the work of Gal-Chen (1978) as applied to deep convection by Hane et al. (1981). The method has been used successfully to diagnose kinematic structure and inthe calculation of momentum budgets of mesoscale systems (e.g., Jorgensen et al. 1991; LeMone and Jorgensen 1991; Hane and Jorgensen 1995). Although the methodology allows for the use of local time derivatives of the three wind components, for this case steady state is assumed due to the lack of apparent evolution in mesoscale organization and line-averaged flow structure during the highly linear stage of the system during the early phase of the aircraft’s investigation. There could be, however, substantial time variability associated with individual convective cells (Sun and Houze 1992) that might impact retrieval results. To evaluate those impacts, an evaluation of temporal changes in the wind fields and their effect on the retrieved pressure fields was made using a least-squares approach in estimating time derivatives described in Hane and Jorgensen (1995) for the first three coordinated flight legs (scans roughly every 10 min). The differences in the deduced pressure values at each grid point between the steady state fields and those calculated with estimates of time derivatives for the linear part of the line were less than about 10% with very similar spatial patterns. Therefore the steady state assumption for the calculation of retrieved mesoscale fields is felt to be good and is not surprising given the degree of filtering employed and the coarseness of the time resolution. However, the system clearly evolved greatly following the appearance of the line-end vortices, and retrieval calculations will not be discussed here for those later times.

The “most optimal” frame of reference was calculated according to the method of Hane (1993) in order to move the reference frame with the most dynamically active features in the flow. Probably because of the highly organized and consistent motion of the updrafts associated with the leading convective line, the most optimal reference frame determined by the Hane method was virtually identical to the motion of the line determined by tracking the leading edge radar reflectivity pattern (eastward at 12 m s−1). A measure of the relative error in the retrieved pressure gradients is Er (Gal-Chen and Kropfli 1984). The 22 February retrievals had Er values of 0.331–0.358, which is in the same range as other Doppler radar studies of mesoscale convective systems (Hane and Jorgensen 1995).

The buoyancy deviation from horizontal averages is represented by
i1520-0469-54-15-1961-eq1
where θ is potential temperature, qυ the deviation from base-state water vapor mixing ratio, and qr the condensate mixing ratio (cloud and rain). Buoyancies are presented as perturbation virtual potential temperatures and do not include water loading, which has been removed with an empirical formula based on radar reflectivity derived by Hauser and Amayenc (1986) from data collected by an optical spectropluviometer (Hauser et al. 1984) in tropical convection:
i1520-0469-54-15-1961-e1
where ρ is air density and Z the radar reflectivity (mm6 m−3). Since water vapor and cloud water cannot be independently estimated from radar data, their effects are included in the retrieved temperatures. Although horizontal gradients of retrieved temperatures are accurately portrayed (to the extent the velocity field allows), vertical gradients of temperature should not be interpreted as departures from a base state because active convection covers asubstantial portion of the domain. This caution applies most readily to direct comparisons of retrieved buoyancies (temperatures) to model simulations. In addition, horizontal gradients of retrieved fields of perturbation pressure and perturbation buoyancy are considered more accurate in vertical cross sections than vertical gradients since the calculations contain an unknown integration constant at each vertical level.

e. Data collected

The flight patterns of the two P-3 aircraft consisted of seven quad-Doppler patterns parallel to and on either side of the convective leading edge from 2055 UTC to 2200 UTC at altitudes of 150 and 300 m MSL. Flight leg duration was generally 5–8 min, which equates to leg lengths of 40–64 km. The low-level flight altitudes and proximity to the line gain two advantages: accurate echo-top precipitation vertical motions can be derived because of the high antenna elevation angles, and in situ measurements in the boundary layer could be simultaneously made. Since it takes about 2 min for the aircraft to reverse its heading, a complete Doppler volume scan was repeated every 7–10 min. Up until 2200 UTC, the foci of the flight patterns was on the leading convective line and its change from a highly linear orientation to one that had a pronounced bow in the leading edge. After 2200 UTC both aircraft began flight tracks oriented perpendicular to the convective line (termed “whole system” patterns since the leg lengths were ∼100–150 km to encompass most of the stratiform rainfall region of the system).

4. Convective precipitation structure, airflow, and evolution

In this section we summarize the basic structure of the convective zone of the squall line system and its evolution over the period extending from its linear phase from about 2100 UTC to the point in which the convective line exhibited its most pronounced bow feature (2219 UTC).

a. Horizontal structure

At the lowest levels during the highly linear phase of the squall line the circulation was predominantly two-dimensional. Fairly uniform and strong convergence of air at 1.5 km MSL is indicated in Fig. 4, near the leading reflectivity maxima. At 1.5 km, the system-relative flow behind the line was from the rear to the front of the system. A slight augmentation of this rear inflow of about 4–5 m s−1 be seen near the location in the line where the bow as to take place (within the box region of Fig. 4a). This rear-to-front flow was stronger than the westerlies shown at 800 mb (about 1.5 m s−1 system relative) in the environmental sounding (Fig. 3) indicating some acceleration of the westerlies by the system. Approximately 12–15 km to the rear of the convective line was a relative minimum in radar reflectivity of 25–30 dBZ. The minimum in radar reflectivity is much closer to the convective line than the usual position of the “transition zone” seen in other squall line systems (e.g., Biggerstaff and Houze 1993; Smull and Houze 1987) and is associated with the updraft minimum (to be shown later) between the leading edge updraft and the strong updraft aloft ∼20 km to the rear of the convective line. The effects of attenuation of the X-band radar beam by intervening rainfall cannot by completely discarded since the reflectivity minimum is nearly equidistant from the two flight tracks.2 However, the double-banded reflectivity structure is also evident in the model results (Trier et al. 1997; Fig. 2).

Updrafts at 1.5 km MSL (Fig. 4b) were oriented in aribbon approximately 4–7 km wide with maxima of 2–3 m s−1 just ahead of the reflectivity maxima, although the entire width of the low-level updraft channel is not resolved due to the close proximity of the aircraft during this pass. These low-level Doppler-derived vertical motions are much lower than those simulated by Trier et al. (1997) and weaker than vertical air velocities measured by the aircraft during direct penetration of the line (up to 13.5 m s−1 updraft). Since the majority of the low-level inflow occurred below 1 km MSL, it is likely that the Doppler radar suffered from significant sea-clutter contamination and hence underestimation of this low-level inflow, which reduced the magnitudes of the low-level updrafts. Downward motion up to 1 m s−1 dominated the areas farther west, indicating a descending rear-to-front flow. At 9.5 km, the reflectivity maximum has shifted westward, indicative of a westward sloping updraft (Fig. 5) with the dominant flow being front to rear. The updraft character also changed from the narrow ribbon of updraft at low levels to a much more cellular structure at 9.5 km with maximum updraft velocities of 7–8 m s−1.

The retrieved pressure perturbation field at 1.5 km MSL and the retrieved perturbation thermal buoyancy (potential temperature) field at 1.0 km MSL (lowest information level) during the linear stage are shown in Fig. 6. The lowest perturbation pressures (approximately 0.7–0.8 mb) lies just behind (to the west) of the convective line near the north end of the domain near the center of counterclockwise curvature in the low-level flow shown in Fig. 4a. The region of coolest air is also located immediately behind the high reflectivity zone at the north end of the domain. The difference in retrieved temperatures between the air ahead of the system and the rain-cooled air in the wake of the convective line is only about 0.6°C. In situ data collected during a low-level penetration of the leading edge gust front (shown later) shows about a 3°C cooling as the gust front is penetrated at 150 m MSL. This discrepancy is probably explained by the shallowness of the cold pool, which sounding data (to be shown later) reveals is confined mostly below 0.5 km MSL. Such cold pool shallowness is not unexpected; previous studies by Parsons et al. (1994) and Jabouille et al. (1996) documented such structure for COARE convective systems.

b. Line-average vertical cross section through the convective line

The highly two-dimensional structure of the airflow and vertical circulation during the early, linear stage of the convective line can be better characterized by alongline averages that are approximately parallel to the principal part of the leading convective line, which were well observed by the two aircraft radars. The domain of the average is indicated by the rectangles in Figs. 4 through 6. Line-averaged cross sections of radar reflectivity and vertical velocity are shown in Fig. 7. A pronounced characteristic of this squall line is the rearward tilt (east to west, which is relative to the direction of propagation toward the east) with height of the leading edge and airflow. A tilt of roughly 35° from the horizontal is seen in both the vertical airflow and reflectivity contours in the cross section. This pronounced tilt was evident in the early, linear stage of the convective line as well as the bow stage and is very similar to that seen in earlier studies (LeMone et al. 1984a). This tilt was also replicated in the numerical simulations of this system by Trier et al.(1996) and Trier et al. (1997). The cross section depicts many of the common characteristics of tropical squall lines, including a relatively narrow leading edge reflectivity maximum of 10–15 km just to the rear of the leading edge updraft, a rapid reflectivity reduction with height above the melting level (∼4.5 km), and a rear-to-front flow below 4.5 km penetrating to the convective line. Pronounced reflectivity decreases above the melting level are commonly observed in tropical convection, even in convective regions, in distinct contrast to midlatitude convection (Jorgensen and LeMone 1989; Szoke et al. 1986; Zipser and Lutz 1994) and has been associated with much weaker updraft velocities in tropical convection (Zipser and LeMone 1980).

The vertical motion cross section (Fig. 7b) reveals distinct updraft maxima at successively higher heights from the leading edge rearward. Located at distance of 46 km (just ahead of the high reflectivity leading edge) was a 4–5 m s−1 updraft extending upward from about 2.5 km MSL. Since this low-level maximum was ahead of the leading edge it was probably associated with new convective development forced by lifting at the gust front. Just rearward of the leading edge (around point “D”) was a relative minimum in vertical velocity at midlevels (4–9 km MSL), followed farther rearward (at point “B”) by another relative maximum in updraft strength at very high levels (9–15 km MSL). This double-peaked updraft structure with height has been seen in previous Doppler-derived vertical motion fields of tropical squall lines that were embedded in strong lower-tropospheric sheared environments (Chong et al. 1987). The double-peaked structure has also been documented in several numerical simulation studies of multicellular squall lines over a broad range of idealized environmental wind profiles (e.g., Fovell and Ogura 1989) as well as the simulations of this system by Trier et al. (1997), although the secondary updraft maximum seemed to be somewhat transitory in nature compared with these observations. Associated with the double-peaked updraft structure were twin maxima in reflectivity at distances of 42 and 28 km from the left edge of the plot. Farther rearward (distances >20 km from the left edge) there is a faint bright band in reflectivity indicative of stratiform precipitation.

The line-averaged cross section also shows that the region rearward from the active leading edge is dominated by weak descent (between 0 and −1 m s−1) below the sloping updraft and below about 3 km MSL (point “A”). As in earlier squall line studies, this weak descent is associated with the system-relative rear-to-front flow, while the strong updraft zone extending rearward from the leading edge is predominantly front-to-rear flow. The shallow near-surface front-to-rear flow often seen in squall lines [and the simulations of this system by Trier et al. (1996)] just rearward from the leading edge was not seen by the airborne Doppler radar, probably due to sea clutter contamination. Near the echo top, a zone of relatively strong downdraft (up to about −4 m s−1) was seen rearward of the decelerating principal updraft above about 11 km MSL, although the equilibrium level of a undiluted parcel is ∼15 km. Roux (1988) diagnosed strong downdrafts in Doppler-derived vertical motion fields near echo top in a West African squall line. Yuter and Houze (1995) have also documented upper-level downdrafts near deep convective updrafts that are apparently “mechanically” forced by pressure gradientforces that arise near the top of buoyant updrafts. According to their hypothesis, near the tops of these rising updrafts air is being pushed laterally, producing local convergence and downward motion through mass continuity.

The pressure and buoyancy perturbation fields (from virtual temperature and cloud water only) for the same line-average cross section are shown in Fig. 8. The perturbation buoyancy (Fig. 8b) for the principal updraft was positive between 1.5 and 12 km MSL, consistent with active, and perhaps growing, convection. The largest positive buoyancies (∼0.8 K) of the principal updraft were located just above the freezing level between about 5 and 8 km MSL (point “D”) near the location of relative updraft minimum (Fig. 7b). These positive buoyancies in the updraft region are still low compared to an undiluted parcel (∼3°C), which indicates probable mixing of the updraft with environmental air.

The perturbation buoyancy calculated from Doppler-radar wind retrievals are lower by a factor of 2–3 from those calculated by Trier et al. (1997) from simulations, in part because the model outputs temperature (the major contributor to the buoyancy in this case) as deviations from the initial model base state rather than a departure from a horizontal average. In the Doppler retrievals convection covers a moderate percentage of the domain, which would reduce perturbation buoyancy compared to deviations from a base state. Although it may be tempting to attribute the vertical velocity minimum to water loading (especially since it occurs within the maximum of thermal buoyancy), Trier et al. (1997) showed in the simulation that it is the dynamic pressure gradient forces within a few kilometers of the leading edge that act to vertically decelerate upward-moving air parcels immediately rearward of the leading-edge updraft. In addition to the dynamic pressure force, there is a downward-directed pressure force associated with the buoyancy-generated mesolow located farther to the rear (point “A” in Fig. 8a) that nearly compensates the upward-directed buoyancy force (e.g., Trier et al. 1997). Once air moves upward and rearward of this adverse pressure gradient associated with the mesolow, its positive buoyancy results in upward accelerations leading to the secondary updraft maximum.

Near the echo top existed substantial deceleration and cooling, although the updraft region shows negative buoyancy as low as 11–13 km MSL, which is about 4 km below the tropopause. Such strong negative buoyancy this far below the tropopause in Fig. 8b may be partially an artifact of the line-averaging methodology as Fig. 5b shows only two strong updraft cores in the line-average domain. A cross section through one of those cores shows rapid vertical deceleration probably associated with the updraft penetrating beyond its equilibrium level.

Lack of scatterers prevents complete documentation of the buoyancy perturbation ahead of the line. Cool air is evident to the rear of the line below about 3 km MSL, although warming is suggested farther to the rear. The region above 2 km MSL <20 km rearward from the line was weakly descending and was also weakly cool relative to the surroundings, perhaps indicative of melting and evaporative cooling. The simulations by Trier et al. (1997) are consistent with these observations and show this region from 1 to 2 km MSL is warmer than initial conditions, indicative of subsidence-induced adiabatic warming exceeding cooling by evaporation. In all likelihood, the low- to midlevel warming or cooling of the rearward region of squall lines is a delicate balance between rearinflow–produced temperature and humidity advection, hydrometeor supply from the overhead anvil, and mesoscale downdraft strength. The “onion” sounding in the wake region of squall lines (Zipser 1977) is a common observation indicating descending air and subsidence warming exceeding cooling by microphysical processes.

The most striking feature of the perturbation pressure cross section is a mesolow of about −0.3 mb located beneath the sloping updraft in the reflectivity minimum region (distance of 36 km at point “A”). The mesolow is commonly observed in convective lines that exhibit considerable leading-edge tilt (LeMone et al. 1984b), which results in lowering of pressure due to the presence of warm updraft air above and rain-cooled near surface air below. The commonly observed shallow mesohigh pressure ridge beneath the mesolow is not seen in the Doppler-derived retrieval fields, probably either due to the shallow cold pool or because of inadequate resolution of near-surface winds or sea clutter contamination. In situ data from the low-level flight legs, however, show a well-defined mesohigh that is well replicated in the simulations of Trier et al. (1996). The ascending front-to-rear flow is dominated by high perturbation pressure, reminiscent of interaction of the vertical shear of the horizontal wind with the vertical motion (Rotunno and Klemp 1982). Another possible explanation for the high perturbation pressure seen in the ascending flow is that the buoyancy decreases upward between 6 and 9 km MSL.

c. Vortex structure associated with the bow in the line

The flight paths of the two aircraft from 2100 to 2245 UTC were focused on that part of the convective line exhibiting rapid eastward bowing motion. The N42 and N43 aircrafts flew seven consecutive coordinated quad-Doppler flight legs during this period. Low-level flow (1.5 km MSL) obtained from the quad-Doppler wind synthesis is shown in Fig. 9 to illustrate the rapid evolution in circulation as the bow developed. Coincident with the rapid eastward surge in the convective line was the development of a mesoscale vortex at the northern end of the bow. Vortex development (Figs. 9a and 9b) occurred at the northern end of the convective line, where low-level flow was able to turn southward behind the line. The apex of the bow (in Fig. 4a at y = ∼58 km) coincided with the location of strongest rear inflow. At 1.5 km MSL, both the horizontal scale and the magnitude of the relative vertical vorticity increased dramatically near the northern end of the bow within 1 h (from ∼1 × 10−3 s−1 in Fig. 9a to ∼2.5 × 10−3 s−1 in Fig. 9d). Although initially confined to the lowest levels, the vortex eventually extended to near 5 km MSL. At the later times (Figs. 9e and 9f) the strong low-level north winds produced convergence and new convective line development that was oriented approximately east–west. As the line passed over Guadalcanal Island (Fig. 9g), the convection rapidly decayed (see Fig. 2).

d. Structure of the “transverse bands”

As noted previously in the low-level radar depictions (Fig. 2), several bands of enhanced reflectivity were noted that were oriented approximately transverse to the principal convective line that, in its early stages, was oriented north–northwest to south–southeast. The band farthest south (labeled “A” inFig. 2b) resembles the structure simulated in Trier et al. (1997; their Fig. 10). Upper-level enhanced front-to-rear flow near the southern end of the line apparently transported hydrometeors rearward creating the appearance on radar of a transverse band of reflectivity, although in situ generation of precipitation may have also played a role as the model’s bands formed in moist ascending flow. Such enhanced flow at the extreme southern part of the Doppler analysis domain along y = 5 km is seen in Fig. 10, although the location is too high (4.5 km MSL) to show much enhanced reflectivity. At low levels (Fig. 9g) the band is much better defined.

A notable transformation in the orientation of the enhanced reflectivity region associated with the upper-level updraft and secondary reflectivity maximum to the rear of the convective line occurred as the convective line bowed eastward. During the highly linear stage of the line, this band was primarily oriented parallel to the leading line (Fig. 9a). As the central part of the leading line surged eastward, the portion of the line to the north of the bow became aligned east–west, its inflow being enhanced by southward flow on the west side of the northern vortex. A secondary east–west band appeared immediately to the south of this segment (Fig. 9f along y = 45 km), but it simply marks the location of the upper-level updrafts associated with the east–west portion of the leading edge (cf. Fig. 5). This conclusion is supported by the vertical cross section in Fig. 11, which is oriented approximately NNE to SSW through the line segment. In the figure (which resembles the average east–west cross sections in Fig. 7), the strong front-to-rear flow is provided by the southward flow on the west side of the vortex. Farther south the precipitation structure was more stratiform in nature, as indicated by the bright band near the melting level and the mesoscale downdraft below it.

e. Low-level structure from in situ data

A comparison of the dynamic retrieval values to in situ data is possible at the lowest levels, although the in situ data were collected at a different location and at later times than the retrieved pressures provided by the Doppler data. Following the quad-Doppler patterns parallel to the line, both aircraft performed east–west passes through the line and rearward 100–150 km at various altitudes from 2218 UTC to about 2319 UTC. Cross sections of perturbation (i.e., horizontal means removed) D value (the difference between pressure altitude and radar altitude, i.e., the departure from a “standard” atmosphere defined by the environment) for both aircraft are presented in Fig. 12. A D value of 1 m is about 0.1 mb at altitudes <500 m, so there is qualitatively good comparison between the retrieved mesolow strength (Fig. 8a) and N43’s cross section (Fig. 12a), which went approximately through the apex of the bow (the southernmost track in Figs. 9f and 9g), although by the time of the cross section the convective line was weakening rapidly, probably due to its proximity to Guadalcanal Island. The perturbation high pressure evident at the lowest levels from −20 km to about −90 km from the leading edge in the in situ data is not well represented in the Doppler-derived perturbation pressure, probably due to inadequate resolution of the winds due to sea cluttercontamination. The low-level high pressure is likely due to the cold near-surface outflow from the downdrafts in the high rainfall region of the convective line. The cross section by N42, taken along the northern boundary of the vortex (Fig. 12b) and about 35 km north of N43’s, shows a much narrower and slightly weaker mesolow than N43’s, probably because N43’s was taken underneath the broad, positively buoyant, front-to-rear flow. The N42 winds between 0 and 20 km are influenced by the mesolow; to the west of −20 km to about −90 km the aircraft passes into the cold pool, which is marked by southerly flow and low θe air (not shown).

5. Surface fluxes and boundary layer recovery

a. Characteristics of the air behind the convective line

In this section we quantify the rate at which the boundary layer can “recover” from squall line passage through enhanced surface fluxes of energy from the ocean. The character of the gust front air following passage of the 22 February squall line is shown in Fig. 13, which displays plots of equivalent potential temperature, potential temperature, mixing ratio, liquid water content, and ground-relative wind for the N42 flight leg segment flown at 150 m from 2148 to 2156 UTC. The flight track shown is the eastern track in Fig. 9d; the aircraft clearly penetrated the heavy rain region of the convective line roughly between the end points of the arrow on the θe plot. Note the increased westerlies, colder temperature, and lower mixing ratios; these values correspond to penetrations of the outflow air, while the higher values of θ and q and the northerly winds at the ends of the leg appear to be those of the environment. Although the cloud water content measured by the Johnson–Williams (JW) hot-wire sensor measured around 0.5 g kg −1 during the penetration, this amount seems high for 150 m MSL, which is below cloud base. Probably rain splashed on the JW sensor giving a false high reading. Within the outflowing air at 150 m, marked by the shift in wind to strong westerlies (recall the system motion was east at 12 m s−1), the potential temperature dropped about 3°C and the air became saturated as the aircraft penetrated rain. The abrupt drop in θ and q as the aircraft penetrated the gust front air corresponds to a drop in θe from about 356 K in the inflowing air to about 346 K behind the gust front. Associated with this θe drop was a downdraft (not shown) of ∼1.5 m s−1.

The data obtained at 150 m by N43 flying parallel to N42 but to the west of the convective line (western track in Fig. 9c, ∼50 km behind the leading edge) indicate that the aircraft was penetrating air that was a mixture of convective downdraft air and mesoscale downdraft air (Fig. 14). Depending on the convective downdraft intensity and the timing of the downdraft penetration, the demarcation between the two air masses is quite distinct. The penetration of the convective downdraft outflow is marked by a sharp increase in θe and q, and a sharp decrease in θ and u (shown in Fig. 14 between the arrow on the θ plot). At this time the back edge of the rain-cooled convective downdraft air was moving eastward at 10 m s−1. Thus, the cool and moist high θe air spreads rearward from below the convective band, and the warm, drier, lower θeair is associated with the mesoscale downdraft air coming in from the west. A penetration of the convective band about 0.5 h later indicated considerable low-level divergence normal to the band, confirming the association of such air with the convection. The time series in Fig. 14 also suggests that the aircraft was very near the top of the outflow layer; just after the θ drop between the arrows on the q time series, the aircraft passes through about 20 s of cooler air, followed by alternating warm-dry and moist-cool events.

The high degree of correlation between thermodynamic variables and east–west wind u from all quad-Doppler flight legs at 150 m MSL of N43 behind the convective line is shown in Fig. 15. As in the time series plot of Fig. 14, the air encroaching from the west has lower mixing ratios, higher values of θ, and lower values of θe. Both the time series and scatterplots suggest that considerable mixing is occurring as well as variation in the thermodynamic character of both convective-scale downdraft and frontward flowing air.

Following the quad-Doppler legs that focused on the convective line from 2100 to 2219 UTC, the two P-3 aircraft began flight legs oriented approximately normal to the convective line and extending >100 km toward the rear of the system. An example of these patterns for the time period 2336 to 2356 UTC is shown in Fig. 16. The southern aircraft (N43) flew at ∼150 m MSL while the northern track was at ∼325 m to avoid some small islands near the middle of the flight leg. Even though the squall line had locally reformed into a quasi-linear north–northwest to south–southeast line, there remained in the rear evidence of a large-scale counterclockwise vortex (anticyclonic in the Southern Hemisphere) in the in situ wind field. The line motion at this time was toward 55° at 12 m s −1.

Time series of thermodynamic quantities, liquid water, ground-relative wind velocity, and storm-relative wind components along and normal to the flight track (us and υs, in right-handed, storm-relative coordinates with x axis aligned toward 55° at 12 m s−1) from data obtained by N43 from (the southernmost flight track in Fig. 16) appears in Fig. 17. Two minima in q and θe are evident, at ∼20 and ∼80 km behind the leading edge. These minima correspond to divergence in us, probably representing maximum penetration of the low θe air. The minimum θe of 343 K along the flight leg near 80 km approaches but is significantly less than the minimum (338 K) in the environmental sounding, its higher value may indicate some mixing during descent as well as some modification from surface fluxes. The θe minimum near the leading edge (∼20 km) is probably associated with a convective-scale downdraft. Rearward of both θe minima, the boundary layer seems to be recovering fairly rapidly but not quite to the θe values (i.e., 355 K) seen in the environment ahead of the convective line. Cloud or rain, indicated by the JW liquid water meter, are not significant at or below flight level between 20 and 100 km behind the line. Above flight level, however, a deep stratiform cloud was evident on the radar display and in photographs from N43.

The hypothesized air motion associated with the minima is idealized in two dimensions in Fig. 18. At theminimum θe point the descending mesoscale downdraft depresses the top of the boundary layer, which grows in depth as the diverging air recovers. The radar-determined vertical velocity (Fig. 19), averaged over the 105 km × 105 km box in Fig. 16, indicates that the air below 2 km is descending on average.

b. Boundary layer recovery from convective line passage

The profiles of θ and q obtained by both P-3 aircraft based on “minisoundings” (quick ascents/descents through the boundary layer) and flight-level data characteristic of the environment ahead of the squall line (Fig. 20a), at the θe minimum point ∼20 km to the rear of the convective line (Fig. 20b), and in a more “recovered” region (∼120 km) farther to the rear (Fig. 20c). Also shown in Fig. 20 is an environmental q profile “adjusted” for bias between the two aircraft’s dewpoint sensors.3

We refer to the well-mixed lower layer of air as the boundary layer (BL). The modified BL is about 200 m deep, and the environmental BL is about 500 m deep. We estimate the BL recovery time from observed data using the simplified two-dimensional equation for 〈〉 and θ̄. The equation for the time rate of change of 〈〉 is written as
i1520-0469-54-15-1961-e2
where the first term on the right-hand side of (2) represents horizontal advection, the second term vertical advection, and the third term represents the contribution of water-vapor fluxes from the surface and through entrainment at the top of the BL. The equation for θ̄ takes similar form. Entrainment at the top of the air column is assumed zero, since vertical averaging (〈 〉) is over the depth of the fairweather BL (depth h). In other words, we assume that there is no turbulent mixing at h = 500 m in the modified air, something that is certainly valid until the BL gets close to h in depth. Overbars and primes represent local horizontal means and deviations, respectively. The BL is considered recovered when θ̄〉 and 〈 reach their environmental values. The effects of rainfall are assumed negligible.
To calculate how quickly the BL air recovers as it moves away from the cold air source, advection is set to zero and the time change of q is approximated by
i1520-0469-54-15-1961-eq2
where tr is the approximate recovery time assumed the same for q and θ. The two resulting recovery equations for θ̄ and 〈〉 have two unknowns, tr and w̄, and tr is given by
i1520-0469-54-15-1961-e3
where Δz denotes a change over a fixed depth (here, 100 m) and is found by substituting tr in the finite-difference form of (2).
The terms in (3) are estimated from the profiles of θ̄ and from the bias-adjusted N43dataand N42 data in Fig. 20. Estimates of latent and sensible heat fluxes are computed using the flight-level data using the TOGA COARE bulk aerodynamic technique (Fairall et al. 1996) in the recovering area (Fig. 21) as
i1520-0469-54-15-1961-eq3
With these values, solving the two simultaneous equations yields tr ∼ 3 h with a mean 500-m MSL subsidence of 0.02 m s−1. The flight-level data suggests that at the θe minimum point the air was receding from the line at about 4 m s−1, implying, if no other processes were acting, a BL recovery within about 45 km from this point. Figure 17 shows similar rates of recovery, and the profiles in Fig. 20 show that air that is about 120 km from the convective line is significantly warmer and moister, although not quite to environmental values. Figure 20 also shows warming and drying above the BL at this location, apparently due to subsidence, which produces an “onion” sounding (Zipser 1977). However, the inversion is not much stronger than that for the environmental sounding in Fig. 20.

To estimate the recovery of the air in squall line relative coordinates, the effects of advection must be added, here idealized as only along the flight leg (NE–SW). From Fig. 17, θ̄/∂x ≈ −2.6 × 10−5 K m−1. Adding this advection to the surface flux contribution produces a net cooling with time. Although these estimates are a gross simplification, the implication is that this squall line cold dome continued to cool or deepen, since advection of cooler and dryer air from the convective line was greater than the effects of surface fluxes near the θe minimum point, unless it is counteracted by the neglected y-advection term.

Since surface fluxes are not significantly different that those computed from earlier field projects (Johnson and Nicholls 1983), we suspect that Eulerian recovery times (at a point) are also similar (∼10 h).

6. Summary and conclusions

This study examined the three-dimensional structure and evolution of a tropical squall line system observed by instrumented aircraft using in situ and airborne Doppler radar data. The quad-Doppler scanning and data processing methodology was used to obtain as accurate an estimate of vertical velocity as has been obtained to date using airborne Doppler radar. The observations encompassed a time period during which the squall line underwent a change from a linear orientation (nearly perpendicular to the environmental low-level jet), to a three-dimensional stage exhibiting a pronounced “bow” in the convective line. Many aspects of the squall line structure have been documented in earlier observational and numerical studies of both midlatitude and tropical squall lines, including a strongly rearward tilted convective leading edge, and a mesolow pressure zone just rearward of the convective line and under the highly sloped front-to-rear flow.

Several characteristics (for tropical lines) were also documented in greater detail including the following.

  • The appearance of mesoscale line-end vortices coincident with the pronounced eastward bow in the middle of the convective line reminiscent of “bookend” vortices seen in midlatitude mesoscale convective systems with larger vertical shear and CAPE(Scott and Rutledge 1995; Verlinde and Cotton 1990; Bartels and Maddox 1991; Weisman 1993).
  • A double-peaked updraft structure within the leading edge convective line, that is, a maximum at low levels just ahead of the maximum rain region associated with lifting by the gust front and new convective cell development, a distinct vertical velocity minimum near the freezing level, and another maximum in the upper troposphere.
  • The growth of elongated transverse precipitation bands toward the systems’ rear.
  • Rapid recovery of the boundary layer in the wake of the squall line through intense fluxes of latent and sensible heat from the underlying warm sea. Continued injection and strong horizontal advection of gust front air sustained the cold pool. Following squall line demise, recovery would probably take much longer since fluxes would be reduced as winds dropped and the boundary layer became moister and warmer (i.e., smaller values of Δθ and Δq).

Previous observational studies (e.g., LeMone et al. 1984b) have documented the “countergradient” nature of the line-normal momentum transport associated with highly linearly oriented, and rapidly moving, squall lines. That was also the case for this system when it was highly linear as the two tiered flows (front-to-rear over rear-to-front flow) were accelerated and thus increasing the mean vertical shear (i.e., countergradient transport). The acceleration of the line-normal wind (over the environmental sounding) was due, presumably, to the generation of the rear mesolow that was itself tied to the strongly rearward sloped front-to-rear flow (LeMone et al. 1984b). The development of vortices associated with nonlinear convective lines will undoubtedly complicate the interpretation of momentum transport since the flow becomes much more three-dimensional. The parameterization of momentum transport in terms of larger scale variables is also more complicated in the case of vortex generation, because, as Trier et al. (1997) points out, the primary source for vertical vorticity generation in a numerical simulation of this case is not simply tilting and stretching of ambient horizontal vorticity arising from ambient shear, but the large contribution by system-induced vertical shear produced by horizontal buoyancy gradients. As convective lines become nonlinear, the local orientation of the environment shear relative to the cold pool also influences the production of system-induced vertical shear. This problem will be further explored by the construction of momentum flux budgets from the airborne Doppler-derived wind fields.

Acknowledgments

We wish to acknowledge the efforts of the flight crew members of the P-3 aircraft from NOAA’s Office of Aircraft Operations in Tampa, Florida. Their dedication and enthusiastic participation in the long 4-month TOGA COARE project from the remote base of operations at Honiara, Solomon Islands, made the collection of data in or near hazardous weather conditions possible. Mr. Robert Hueftle was instrumental in development of the Doppler editing and display software, Dr. Bradley Smull helped execute the dual P-3 flight patterns as Chief Scientist on N43, Mr. Thomas Shepherd and Ms. Diana Bartels operated the Doppler radars on the two P-3 aircraft, Dr. Thomas Matejka was instrumental in the design of the quad-Doppler flight strategies and also developed the methodology for synthesizing the horizontal and vertical winds from the radial velocity data, Dr. John Daugherty prepared the P-3 radar composites and managed most of the datasets, and Mr. David Johnson edited most of the P-3 Doppler radar data. We also acknowledge Carl Friehe and Shawn Burns of theUniversity of California, Irvine, for their assistance in determining biases in the P-3 in situ data. The NCAR Zebra software was used to construct the satellite photo and flight track plots in Fig. 1. This work was supported by NSF Grant 9215507 and NOAA/OGP Grant 8R1DA110. We lastly acknowledge the many stimulating discussions about the nature of squall lines, their structure, and momentum flux characteristics as well as the use and interpretation of airborne Doppler radar data with Peter Hildebrand (NCAR), Robert Houze (University of Washington), Dave Parsons (NCAR), Frank Roux (Laboratoire d’Aerologie, Toulouse, France), Bradley Smull (NOAA/NSSL), Wei-Kuo Tao (NASA/GSFC), Jacques Testud (CRPE/France), Morris Weisman (NCAR), and Edward Zipser (Texas A&M University).

REFERENCES

  • Barnes, G. M., and K. Sieckman, 1984: The environment of fast- and slow-moving tropical mesoscale convective cloud lines. Mon. Wea. Rev.,112, 1782–1794.

  • Bartels, D. L., and R. A. Maddox, 1991: Midlevel cyclonic vortices generated by mesoscale convective systems. Mon. Wea. Rev.,119, 104–118.

  • Battan, L. J., 1973: Radar Observations of the Atmosphere. University of Chicago Press, 324 pp.

  • Biggerstaff, M. L., and R. A. Houze Jr., 1991: Kinematics and precipitation structure of the 10–11 June 1985 squall line. Mon. Wea. Rev.,119, 3034–3065.

  • ——, and ——, 1993: Kinematics and microphysics of the transition zone of a midlatitude squall-line system. J. Atmos. Sci.,50, 3091–3110.

  • Braun, S. A., and R. A. Houze Jr., 1994: The transition zone and secondary maximum of radar reflectivity behind a midlatitude squall line: Results retrieved from Doppler radar data. J. Atmos. Sci.,51, 2733–2755.

  • Chong, M., P. Amayenc, G. Scialom, and J. Testud, 1987: A tropical squall line observed during the COPT 81 experiment in West Africa. Part I: Kinematic structure inferred from dual-Doppler data. Mon. Wea. Rev.,115, 670–694.

  • Davis, C. A., and M. L. Weisman, 1994: Balanced dynamics of simulated, long-lived mesoscale convective systems. J. Atmos. Sci.,51, 2005–2030.

  • Fairall, C. W., E. F. Bradley, D. P. Rogers, J. B. Edson, and G. S. Young, 1996: Bulk parameterization of air–sea fluxes for TOGA COARE. J. Geophys. Res.,101, 3747–3765.

  • Fovell, R. G., and Y. Ogura, 1988: Numerical simulation of a midlatitude squall line in two dimensions. J. Atmos. Sci.,45, 3846–3879.

  • Gal-Chen, T., 1978: A method for the initialization of the anelastic equations: Implications for matching models with observations. Mon. Wea. Rev.,106, 587–696.

  • ——, and R. A. Kropfli, 1984: Buoyancy and pressure perturbations derived from dual-Doppler radar observations of the planetary boundary layer: Applications for matching models with observations. J. Atmos. Sci., 41, 3007–3020.

  • Hane, C. E., 1993: Storm motion estimates derived from dynamic retrieval calculations. Mon. Wea. Rev.,121, 431–443.

  • ——, and D. P. Jorgensen, 1995: Dynamic aspects of a distinctly three-dimensional mesoscale convective system. Mon. Wea. Rev.,123, 3194–3214.

  • ——, R. B. Wilhelmson, and T. Gal-Chen, 1981: Retrieval of thermodynamic variables within deep convective clouds:Experiments in three dimensions. Mon. Wea. Rev.,109, 564–576.

  • Hauser, D., and P. Amayenc, 1986: Retrieval of cloud water and water vapor contents from Doppler radar data in a tropical squall line. J. Atmos. Sci.,43, 823–838.

  • ——, ——, and M. Chong, 1984: A new optical instrument for simultaneous measurement of raindrop diameter and fallspeed distributions. J. Atmos. Oceanic Technol.,1, 256–269.

  • Houze, R. A., Jr., 1977: Structure and dynamics of a tropical squall-line system. Mon. Wea. Rev.,105, 1540–1567.

  • Jabouille, P., J. L. Redelsperger, and J. P. LaFore, 1996: Modification of surface fluxes by atmospheric convection in the TOGA COARE region. Mon. Wea. Rev.,124, 816–837.

  • Johnson, R. H., and M. E. Nicholls, 1983: A composite analysis of the boundary layer accompanying a tropical squall line. Mon. Wea. Rev.,111, 308–319.

  • Jorgensen, D. P., and M. A. LeMone, 1989: Vertical velocity characteristics of oceanic convection. J. Atmos. Sci.,46, 621–640.

  • ——, and B. F. Smull, 1993: Mesovortex circulations seen by airborne Doppler radar within a bow-echo mesoscale convective system. Bull. Amer. Meteor. Soc.,74, 2146–2157.

  • ——, P. H. Hildebrand, and C. L. Frush, 1983: Feasibility test of an airborne pulse-Doppler meteorological radar. J. Climate Appl. Meteor.,22, 744–757.

  • ——, M. A. LeMone, and B. J.-D. Jou, 1991: Precipitation and kinematic structure of an oceanic mesoscale convective system. Part I: Analysis of airborne Doppler radar data. Mon. Wea. Rev.,119, 2608–2637.

  • ——, T. Matejka, and J. D. DuGranrut, 1996: Multi-beam techniques for deriving wind fields from airborne Doppler radars. J. Meteor. Atmos. Phys.,59, 83–104.

  • Keenan, T. D., and R. E. Carbone, 1992: A preliminary morphology of precipitation systems in northern Australia. Quart. J. Roy. Meteor. Soc.,118, 283–326.

  • Leise, J. A., 1981: A multidimensional scale-telescoped filter and data extension package. NOAA Tech. Memo. ERL WPL-82, 18 pp. [NTIS PB82-164104.].

  • LeMone, M. A., and D. P. Jorgensen, 1991: Precipitation and kinematic structure of an oceanic mesoscale convective system. Part II: Momentum transport and generation. Mon. Wea. Rev.,119, 2638–2653.

  • ——, G. M. Barnes, E. J. Szoke, and E. J. Zipser, 1984a: The tilt of the leading edge of mesoscale tropical convective lines. Mon. Wea. Rev.,112, 510–519.

  • ——, ——, and E. J. Zipser, 1984b: Momentum flux by lines of cumulonimbus over the tropical oceans. J. Atmos. Sci.,41, 1914–1932.

  • Matejka, T., and S. A. Lewis, 1997: Improving research aircraft navigation by incorporating INS and GPS information in a variational solution. J. Atmos. Oceanic Technol.,14, 495–511.

  • O’Brien, J. J., 1970: Alternative solutions to the classical vertical velocity problem. J. Appl. Meteor.,9, 197–203.

  • Ogura, Y., and M.-T. Liou, 1980: The structure of a midlatitude squall line. A case study. J. Atmos. Sci.,37, 553–567.

  • Parsons,D., and Coauthors, 1994: The integrated sounding system: Description and preliminary observations from TOGA COARE. Bull. Amer. Meteor. Soc.,75, 553–567.

  • Rotunno, R., and J. B. Klemp, 1982: The influence of the shear-induced pressure gradient on thunderstorm motion. Mon. Wea. Rev., 110, 136–151.

  • Roux, F., 1988: The West African squall line observed on 23 June 1981 during COPT 81: Kinematics and thermodynamics of the convective region. J. Atmos. Sci.,45, 406–426.

  • Rutledge, S. A., R. A. Houze Jr., M. I. Biggerstaff, and T. J. Matejka, 1988: The Oklahoma–Kansas mesoscale convective system of 10–11 June 1985: Precipitation structure and single-Doppler radar analysis. Mon. Wea. Rev.,116, 1409–1430.

  • Schmidt, J. M., and W. R. Cotton, 1989: A high plains squall line associated with severe surface winds. J. Atmos. Sci.,46, 281–302.

  • Scott, J. D., and S. A. Rutledge, 1995: Doppler radar observations of an asymmetric mesoscale convective system and associated vortex couplet. Mon. Wea. Rev.,123, 3437–3457.

  • Skamarock, W. C., M. L. Weisman, and J. B. Klemp, 1994: Three-dimensional evolution of simulated long-lived squall lines. J. Atmos. Sci.,51, 2563–2584.

  • Smull, B. F., and R. A. Houze Jr., 1987: Dual-Doppler radar analysis of a midlatitude squall line with a trailing region of stratiform rain. J. Atmos. Sci.,44, 2128–2148.

  • Sun, J., and R. A. Houze Jr., 1992: Validation of a thermodynamic retrieval technique by application to a simulated squall line with trailing stratiform precipitation. Mon. Wea. Rev.,120, 1003–1018.

  • Szoke, E. J., E. J. Zipser, and D. P. Jorgensen, 1986: A radar study of convective cells in mesoscale systems in GATE. Part I: Vertical profile statistics and comparison with hurricanes. J. Atmos. Sci.,43, 182–197.

  • Trier, S. B., W. C. Skamarock, M. A. LeMone, D. B. Parsons, and D. P. Jorgensen, 1996: Structure and evolution of the 22 February 1993 TOGA COARE squall line: Numerical simulations. J. Atmos. Sci.,53, 2861–2886.

  • ——, ——, and ——, 1997: Structure and evolution of the 22 February 1993 TOGA COARE squall line: Organization mechanisms inferred from numerical simulation. J. Atmos. Sci.,54, 386–407.

  • Verlinde, J., and W. R. Cotton, 1990: A mesoscale vortex couplet observed in the trailing anvil of a multicellular convective complex. Mon. Wea. Rev.,118, 993–1010.

  • Webster, P. J., and R. Lukas, 1992: TOGA COARE: The Coupled Ocean–Atmosphere Response Experiment. Bull. Amer. Meteor. Soc.,73, 1377–1416.

  • Weisman, M. L., 1993: The genesis of severe, long-lived bow echoes. J. Atmos. Sci.,50, 645–670.

  • Yuter, S. E., and R. A. Houze Jr., 1995: Three-dimensional kinematic and microphysical evolution of Florida cumulonimbus. Part II: Frequency distributions of vertical velocity, reflectivity, and differential reflectivity. Mon. Wea. Rev.,123, 1941–1963.

  • ——, ——, B. F. Smull, F. D. Marks, and J. R. Daugherty, 1995: TOGA COARE aircraft mission summary images: An electronic atlas. Bull. Amer. Meteor. Soc.,76, 319–328.

  • Zhang, D.-L., K. Gao, and D. B. Parsons, 1989: Numerical simulations of anintense squall line during 10–11 June 1985 PRE-STORM. Part I: Model verification. Mon. Wea. Rev.,117, 960–994.

  • Zipser, E. J., 1977: Mesoscale and convective-scale downdrafts as distinct components of squall-line circulation. Mon. Wea. Rev.,105, 1568–1589.

  • ——, and M. A. LeMone, 1980: Cumulonimbus vertical velocity events in GATE. Part II: Synthesis and model core structure. J. Atmos. Sci.,37, 2458–2469.

  • ——, and K. R. Lutz, 1994: The vertical profile of radar reflectivity of convective cells: A strong indicator of storm intensity and lightning probability? Mon. Wea. Rev.,122, 1751–1759.

Fig. 1.
Fig. 1.

GMS infrared satellite photograph from 2132 UTC 22 February 1993. Cloud top temperatures (K) are contoured with the color scale at the bottom of the figure. The island of Guadalcanal in the Solomon Islands is shown as the dark line with the sounding site at Honiara (HIR) indicated as the cross. Aircraft tracks of the two National Oceanic and Atmospheric Administration (NOAA) aircraft for the period near 2130 UTC are show as the green (N42) and blue (N43) parallel lines near the eastern edge of the cloud shield.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 2.
Fig. 2.

Composite radar reflectivity (dBZ) over the time periods (a) 2057–2103 UTC, (b) 2141–2147 UTC, (c) 2219–2234 UTC, and (d) 2302–2305 UTC from the P-3’s lower fuselage C-band radar. Each of the four panels is approximately 45 min apart. The flight tracks of the aircraft during the composite period are indicated by the lines with arrows. The composites were constructed by maximizing the reflectivity from each sweep at each grid point. Color scale for radar reflectivity is shown to the right of each panel. The domain of analysis is 240 × 240 km2. Islands are shown by the bold lines, with the largest island extending from the right side of each figure being Guadalcanal, Solomon Islands. Small boxes with tick marks in all but the 2300 UTC map are the Doppler analysis domains. Points labeled “A” and “B” refer to transverse precipitation bands.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 3.
Fig. 3.

Composite sounding (top panel) and hodograph (bottom panel) for the environment of the 22 February squall line. Data is from both P-3 aircraft and the balloon ascent at Honiara, Solomon Islands. See text for a description of the compositing procedure. The system motion vector (eastward at 12 m s−1) is shown as the arrow starting from u = 0, υ = 0.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 4.
Fig. 4.

(a) Horizontal analysis of radar reflectivity and system-relative wind flow and (b) vertical air motion (m s−1) at 1.5 km MSL valid at 2116 UTC. Domain is 75 km × 75 km, the location of which is shown in Fig. 2a. Reflectivity contour gray scale and the 10 m s−1 scaling vector for winds are shown to the right and above (a) respectively. The box located near the center of both panels is the domain over which a line average cross section was constructed.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 5.
Fig. 5.

As in Fig. 4, except for 9.5 km MSL.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 6.
Fig. 6.

As in Fig. 4, except for (a) perturbation pressure (in 10−1 mb) at 1.5 km MSL and (b) perturbation buoyancy (in 10−1 K) at 1.0 km MSL retrieved from the airborneDoppler-derived wind fields.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 7.
Fig. 7.

Vertical cross section of average reflectivity and average line-normal system relative wind in (a) the plane of the cross section and (b) vertical motion (m s−1) averaged within the boxed domain of Fig. 4. In the construction of the line average, the coordinate system was rotated so that u was aligned in the direction of x′ and υ was aligned in the direction of y′ as shown in Fig. 4b. The symbols A, B, C, and D refer to the locations of various maxima in the vertical velocity, pressure perturbation, and buoyancy fields.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 8.
Fig. 8.

As in Fig. 7, except for line average vertical cross sections of (a) perturbation pressure (in 10−1 mb) and (b) perturbation buoyancy (in 10−1 K).

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 9.
Fig. 9.

Horizontal analysis of radar reflectivity and system-relative wind flow at 1.5 km MSL valid at (a) 2100 UTC, (b) 2116 UTC, (c) 2135 UTC, (d) 2154 UTC, (e) 2203 UTC, (f) 2227 UTC, and (g) 2245 UTC. The analysis domains of (a) and (f) are shown on the larger scale radar view of Figs. 2a and 2c.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 10.
Fig. 10.

Horizontal analysis of radar reflectivity and system-relative wind flow at 4.5 km MSL at 2227 UTC. Thin black line oriented approximately NNE to SSW near the center of the figure is the location of the cross section shown in Fig. 11.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 11.
Fig. 11.

Vertical cross section of reflectivity and (a) system-relative wind (dBZ) and (b) vertical air velocity (m s−1) in the plane of the cross through the two transverse bands along the cross-section line shown in Fig. 10. Horizontal distance of the cross section is 53.7 km.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 12.
Fig. 12.

Vertical cross section of perturbation D-value (m, the difference between radar altitude and pressure altitude determined from an environmental sounding) obtained by (a) N43 and (b) N42 during several east–west transects normal to the convective line (bottom panel of each figure). The environmental D value at each height was subtracted to yield perturbation values. Starting and ending times (UTC) of each flight leg segment are indicated at the end points of each leg. Arrows denote direction of flight. The top panel of each figure shows the system-relative horizontal winds from the topmost flight leg (at 940 m MSL) segment. The aircraft were separated by about 0.3° of latitude (∼35 km).

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 13.
Fig. 13.

Time series from 2148 to 2156 UTC of in situ data collected by N42 (the easternmost flight track in Fig. 9d) as it proceeded parallel to the convective leading edge at 150 m MSL. Total distance covered is about 60 km. From top to bottom the time series are east–west ground-relative wind speed (u, m s−1), and north–south ground-relative wind speed (υ, m s−1), equivalent potential temperature (θe, K), potential temperature (θ, K), mixing ratio (q, g kg−1), and cloud water mixing ratio (ql, g kg−1) observed by the Johnson–Williams (JW) hot-wire instrument. For the θe, θ,and q plots the dashed lines are raw data; the solid lines are adjusted for the effects of sensor wetting by liquid water by replacing temperature with bias-adjusted temperature from the side-looking CO2 radiometer; dewpoints in excess of this temperature are set equal to it. The delay in warming and cooling of “adjusted” θ results from the slow response of the radiometer. Similarly, the warm temperature spike in the cold air is probably more accurate than the radiometric value. The raw values should be taken as “correct” where ql = 0. The aircraft penetrated the gust front during the time approximated by the arrow on the θ plot. During the penetration, the liquid water content ql rose from zero to about 0.4–0.5 g kg−1.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 14.
Fig. 14.

Time series of in situ data obtained by N43 as it flew at 150 m MSL from south to north and parallel to and ∼50 km to the rear (west) of the convective line from 2129 to 2134 UTC. Total distance is about 40 km. From top to bottom, the time series are equivalent potential temperature (θe, K), potential temperature (θ, K), mixing ratio (q, g kg−1), and east–west wind velocity (u, m s−1). The aircraft penetrated the rearward moving (in a relative sense) convective-scale downdraft air between the arrows indicated on the θ time series. The track is indicated as the westernmost track in Fig. 9d.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 15.
Fig. 15.

Scatter plots of earth-relative east–west wind velocity (u, m s−1) versus potential temperature (θ, K, top panel), mixing ratio (q, g kg−1, middle panel), and equivalent potential temperature (θe, K, bottom panel) from flight legs behind the convective line by N43. The vertical line at u = 12 m s−1 marks the line motion and hence the division between rear-to-front and front-to-rear moving air.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 16.
Fig. 16.

Horizontal depiction of radar reflectivity derived from a composite of N42 lower fuselage radar scans over the time period 2336 to 2356 UTC. The flight tracks of the two P-3 aircraft are shown as the lines with arrows and wind barbs. Wind barbs depict ground-relative flow and are plotted every 60 s with one full barb equal to 5 m s−1. Domain size is 240 km × 240 km. The box with tick marks is the domain of the dual-Doppler analysis. Color scale for reflectivity (dBZ) is shown to the right.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 17.
Fig. 17.

Time series of in situ data obtained by N43 at 150 m MSL between 2327 and 2356 UTC along the southernmost flight track shown in Fig. 16. From top to bottom are the time series of mixing ratio (q, g kg−1), potential temperature (θ, K), equivalent potential temperature (θe;t2, K), and cloud water content (g kg−1) measured by the Johnson–Williams hot-wire instrument, the line-relative wind normal to the leading edge and along the flight track (us, m s−1), storm-relative wind parallel to leading edge and normal to the flight track (υs, m s−1), and total wind speed relative to the ground (m s−1). Vertical dashed lines denote relative distance (km) from the leading edge. Thin vertical lines indicate the locations of the relative minima inθe. Line motion at this time was toward 55° at 12 m s−1. No bias adjustment has been applied to these time series.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 18.
Fig. 18.

Schematic diagram of air flow contributing to the minimum in θe seen in the flight-level data at 150 m MSL. Convective line is toward the right of the schematic. The effects of surface flux and entrainment are included in the deepening of the boundary layer toward the rear of the system.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 19.
Fig. 19.

Domain-averaged vertical velocity (cm s−1) as a function of height over the Doppler analysis area shown in Fig. 16.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 20.
Fig. 20.

Vertical profiles of mean mixing ratio (q g kg−1, top panel) and mean potential temperature (θ, K, bottom panel) collected by N43. The solid lines (labeled “a”) are values characteristic of the environment ahead of the line sampled by N42. Dashed lines (labeled “b”) are values taken at the θe minimum point ∼120 km behind the convective line while the dash-dotted lines (labeled “c”) are values ∼20 km behind the line along N43’s flight track. The shaded line with boxes is the environmental q profile taken by N42 ahead of the line and “adjusted” for possible biases in temperature and dewpoint between aircraft that intercomparisons indicate may be 0.5°C for dewpoint and 0.1°C for temperature (N43 lower than N42). To be consistent with Fig. 21, values of q should be raised by 0.6 g kg−1.

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

Fig. 21.
Fig. 21.

Bulk surface fluxes computed from the flight-level data obtained by N43 at 150 m MSL (Fig. 17). Solid line represents latent heat fluxes (W m−2), while the dashed line is sensible heat fluxes. Fluxes were computed using the COARE 1.0 bulk flux routine (Fairall et al. 1996) with sea surface temperatures and temperatures from N43 taken as “correct,” and N43 dewpoint temperatures raised by 0.5 K. For this version of the aircraft dataset (available from NOAA/National Severe Storms Laboratory, Boulder, the dewpoint correction is consistent with the Burns–Friehe “best guess” values (S. Burns 1996, personal communication).

Citation: Journal of the Atmospheric Sciences 54, 15; 10.1175/1520-0469(1997)054<1961:SAEOTF>2.0.CO;2

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

1

During TOGA COARE a mixture of sector and 360° rotation was used, depending on the precipitation extent. On 22 February, both radars were operated in the continuous rotation mode.

2

Calculation of attenuation using the X-band attenuation coefficients published by Battan (1973) yields an expected radar signal reduction of ∼1 dBZ due to the intervening rainfall from N43, the westernmost aircraft. The apparent reflectivity minimum evident in Fig. 4a is ∼5–7 dBZ.

3

S. Burns and C. Friehe (1996, personal communication) confirmed this dewpoint bias between the two P-3 aircraft from intercomparisons during formation flying to/from target regions as being ∼0.5° C with N43 being lower than N42.Forcalculations presented here, N42 is treated as the “correct” dewpoint.

Save