• Adang, T. C., and R. L. Gall, 1989: Structure and dynamics of the Arizona monsoon boundary. Mon. Wea. Rev.,117, 1423–1438.

  • Augustine, J. A., and F. Caracena, 1994: Lower-tropospheric precursors to nocturnal MCS development over the central United States. Wea. Forecasting,9, 116–135.

  • Badan-Dangon, A., C. E. Dorman, M. A. Merrifield, and C. D. Winant, 1991: The lower atmosphere over the Gulf of California. J. Geophys. Res.,96, 16877–16896.

  • Bonner, W. D., 1968: Climatology of the low-level jet. Mon. Wea. Rev.,96, 833–850.

  • ——, and J. Paegle, 1970: Diurnal variations in the boundary layer winds over the south-central United States in summer. Mon. Wea. Rev.,98, 735–744.

  • Brenner, I. S., 1974: A surge of maritime tropical air—Gulf of California to the southwestern United States. Mon. Wea. Rev.,102, 375–389.

  • Bryson, R. A., and W. P. Lowry, 1955: The synoptic climatology of the Arizona summer precipitation singularity. Bull. Amer. Meteor. Soc.,36, 329–339.

  • Carleton, A. M., 1986: Synoptic-dynamic character of “bursts” and “breaks” in the southwest U.S. summer precipitation singularity. J. Climatol.,6, 605–623.

  • ——, 1987: Summer circulation climate of the American Southwest, 1945–1984. Ann. Assoc. Amer. Geogr.,77, 619–634.

  • ——, D. A. Carpenter, and P. J. Weser, 1990: Mechanisms of interannual variability of the southwest United States summer rainfall maximum. J. Climate,3, 999–1015.

  • Douglas, M. W., 1995: The summertime low-level jet over the Gulf of California. Mon. Wea. Rev.,123, 2334–2347.

  • ——, R. A. Maddox, K. Howard, and S. Reyes, 1993: The Mexican monsoon. J. Climate,6, 1665–1677.

  • Hagemeyer, B. C., 1991: A lower-tropospheric climatology for March through September. Some implications for thunderstorm forecasting. Wea. Forecasting,6, 254–270.

  • Hales, J. E., Jr., 1972: Surges of maritime tropical air northward over the Gulf of California. Mon. Wea. Rev.,100, 298–306.

  • ——, 1974: Southwestern United States summer monsoon source—Gulf of Mexico or Pacific Ocean? J. Appl. Meteor.,13, 331–342.

  • Harman, J. R., 1991: Synoptic Climatology of the Westerlies: Process and Patterns. Association of American Geographers, 80 pp.

  • Helfand, H. M., and S. D. Schubert, 1995: Climatology of the Great Plains low-level jet and its contribution to the continental moisture budget of the United States. J. Climate,8, 784–806.

  • Higgins, R. W., J. E. Janowiak, and Y. Yao, 1996a: A gridded hourly precipitation data base for the United States (1963–1993). NCEP/Climate Prediction Center Atlas 1, 47 pp. [Available from Climate Prediction Center, NOAA/NWS/NCEP, Washington, DC 20233.].

  • ——, K. C. Mo, and S. D. Schubert, 1996b: The moisture budget of the central United States in spring as evaluated in the NCEP/NCAR and the NASA/DAO reanalyses. Mon. Wea. Rev.,124, 939–963.

  • ——, Y. Yao, E. S. Yarosh, J. E. Janowiak, and K. C. Mo, 1997: Influence of the Great Plains low-level jet on summertime precipitation and moisture transport over the central United States. J. Climate,10, 481–507.

  • Johnson, A. M., 1976: The climate of Peru, Bolivia and Ecuador. Climates of Central and South America, W. Schwerdtfeger and H. E. Landsberg, Eds., World Survey of Climatology, Vol. 12, Elsevier, 147–218.

  • Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-year reanalysis project. Bull. Amer. Meteor. Soc.,77, 437–471.

  • Mitchell, M. J., R. A. Arritt, and K. Labas, 1995: A climatology of the warm season Great Plains low-level jet using wind profiler observations. Wea. Forecasting,10, 576–591.

  • Mock, C. J., 1996: Climatic controls and spatial variations of precipitation in the western United States. J. Climate,9, 1111–1125.

  • Negri, A. J., R. F. Adler, E. J. Nelkin, and G. J. Huffman, 1994: Regional rainfall climatologies derived from Special Sensor Microwave Imager (SSM/I) data. Bull. Amer. Meteor. Soc.,75, 1165–1182.

  • Okabe, I. T., 1995: The North American monsoon. Ph.D. dissertation, University of British Columbia, 146 pp. [Available from Dept. of Geography, University of British Columbia, 2075, Wesbrook Place, Vancouver, BC V6T1W5, Canada.].

  • Oort, A. H., 1983: Global atmospheric circulation statistics 1958–1973. NOAA Professional Paper 14, U.S. Government Printing Office, Washington, DC, 180 pp.

  • Parker, S. S., J. T. Hawes, S. J. Colucci, and B. P. Hayden, 1989: Climatology of 500-mb cyclones and anticyclones, 1950–1985. Mon. Wea. Rev.,117, 558–570.

  • Parrish, D. F., and J. C. Derber, 1992: The National Meteorological Center’s spectral statistical interpolation analysis system. Mon. Wea. Rev.,120, 1747–1763.

  • Rasmusson, E. M., 1966: Atmospheric water vapor transport and the hydrology of North America. Planetary Circulations Project, Massachusetts Institute of Technology Rep. A-1, 170 pp.

  • ——, 1967: Atmospheric water vapor transport and the water balance of North America: Part I. Characteristics of the water vapor flux field. Mon. Wea. Rev.,95, 403–426.

  • Roads, J. O., S. Chen, A. K. Guetter, and K. P. Georgakakos, 1994: Large-scale aspects of the United States hydrologic cycle. Bull. Amer. Meteor. Soc.,75, 1589–1610.

  • Rowson, D. R., and S. J. Colucci, 1992: Synoptic climatology of thermal low-pressure systems over south-western North America. J. Climatol.,12, 529–545.

  • Schmitz, J. T., and S. Mullen, 1996: Water vapor transport associated with the summertime North American monsoon as depicted by ECMWF analyses. J. Climate,9, 1621–1634.

  • Schubert, S. D., J. Pfaendtner, and R. Rood, 1993: An assimilated dataset for earth science applications. Bull. Amer. Meteor. Soc.,74, 2331–2342.

  • Sellers, W. D., and R. H. Hill, 1974: Arizona Climate, 1931–1972. The University of Arizona Press, 616 pp.

  • Starr, V. P., J. P. Peixoto, and H. R. Crisi, 1965: Hemispheric water balance for the IGY. Tellus,17, 463–472.

  • Stensrud, D. J., R. L. Gall, S. L. Mullen, and K. W. Howard, 1995: Model climatology of the Mexican monsoon. J. Climate,8, 1775–1794.

  • Tang, M., and E. R. Reiter, 1984: Plateau monsoons of the Northern Hemisphere: A comparison between North America and Tibet. Mon. Wea. Rev.,112, 617–637.

  • Trenberth, K. E., and J. G. Olson, 1988: An evaluation and intercomparison of global analyses from the National Meteorological Center and the European Centre for Medium Range Weather Forecasts. Bull. Amer. Meteor. Soc.,69, 1047–1057.

  • Wallace, J. M., 1975: Diurnal variations in precipitation and thunderstorm frequency over the conterminous United States. Mon. Wea. Rev.,103, 406–419.

  • Whittaker, L. M., and L. H. Horn, 1981: Geographical and seasonal distribution of North American cyclogenesis, 1958–1977. Mon. Wea. Rev.,109, 2312–2322.

  • WMO, 1975: Climatic Atlas of North and Central America. Vol. I, Maps of Mean Temperature and Precipitation, World Meteorological Organization.

  • Xie, P., and P. A. Arkin, 1996: Analyses of global monthly precipitation using gauge observations, satellite estimates, and numerical model predictions. J. Climate,9, 840–858.

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    Topography of North America (excluding northern Canada). Data courtesy of the United States National Geophysical Data Center. The resolution of the data is 5 min (0.083°). The data are available on the World Wide Web at http://www.ngdc.noaa.gov/mgg/mggd.html. The elevation (in meters) is indicated by the bar at right. The box over Arizona and New Mexico indicates the region 112.5°–107.5°W, 32°–36°N used to define the precipitation index (see section 3).

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    Mean (July–September 1979–95) 925-hPa vector wind, 200-hPa streamlines, and merged satellite estimates and station observations of precipitation (shading). Circulation data are taken from the NCEP–NCAR reanalysis archive. The position of the North American monsoon anticyclone is indicated by “A.” The Bermuda and North Pacific subtropical high pressure centers are indicated by “H.” Precipitation amounts are in millimeters. The characteristic vector length is 10 m s−1. The approximate location of the Great Plains low-level jet is indicated by the heavy solid arrow.

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    Histograms of the mean (1963–94) daily and 5-day running mean precipitation (units: mm) during JJA at each grid point in the box 115°–105°W, 30°–38°N from the precipitation database of Higgins et al. (1996a).

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    Mean (1963–94) pentad (bars) and standard deviation (heavy black line) of observed precipitation (Higgins et al. 1996a) over (a) Arizona (112.5°–110°W, 32°–36°N) and (b) New Mexico (107.5°–105°W, 32°–36°N). (Units: mm.)

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    Evolution of the composite mean (1963–94) daily precipitation index (units: mm) over Arizona and New Mexico. The average date of monsoon onset is 7 July (defined as day 0 in the composite analysis).

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    Mean monthly precipitation (units: mm month−1) for July, August, and September from (a) the Xie and Arkin (1996) merged analysis for 1987–95 and (b) the Higgins et al. (1996a) analysis for 1963–94. Contribution of the precipitation during July, August, and September to the annual total, expressed in percent, from (c) the Xie and Arkin (1996) merged precipitation analysis and (d) the Higgins et al. (1996a) precipitation analysis. The contributions in (c) and (d) are for the same years used in (a) and (b), respectively. In (a) and (b), the contour interval is geometric and values greater than 80 mm are shaded. In (c) and (d), the contour interval is 5% and values greater than 40% are shaded.

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    Analysis of the contribution of the Xie and Arkin (1996) mean (1987–95) monthly precipitation to the annual mean (expressed in percent) over the conterminous United States and Mexico in (a) June, (b) July, (c) August, and (d) September. In each case, the contour interval is 5% and areas with values exceeding 15% are shaded.

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    Analysis of mean July minus June precipitation (units: mm month-1) from (a) the Xie and Arkin (1996) merged analysis for 1987–95 and (b) the Higgins et al. (1996a) analysis for 1963–94. The letters a–g in (b) indicate the locations of grid points for the histograms in Fig. 9. In each case, the contour intervals are . . . , −160, −80, −40, −30, −20, −10, 10, 20, 30, 40, 80, 160, . . . , and positive (negative) values are shaded dark (light).

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    Histograms of mean (1963–94) monthly precipitation (units: mm day−1) from selected grid points over the conterminous United States (indicated in Fig. 8b by the letters a–g). (a) Over Arizona (34°N, 110°W), (b) over Texas (30°N, 97.5°W), (c) over Oklahoma (36°N, 95°W), (d) over Montana (46°N, 110°W), (e) over North Carolina (36°N, 77.5°W), (f) over the Florida panhandle (30°N, 85°W), and (g) over Kentucky (38°N, 85°W).

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    NCEP time mean (JJA 1979–95) for (a) surface wind (units: m s−1) and specific humidity (units: g kg−1), (b) 500-hPa wind (units: m s−1) and specific humidity (units: g kg−1), (c) 500-hPa vertical velocity (units: μb s−1), and (d) precipitable water (units: mm). In (a) and (b), the standard vector length is 10 m s−1. The contour intervals are (a) 2 g kg−1, (b) 0.5 g kg−1, (c) 0.1 μb s−1, and (d) 5 mm. The shading denotes (a) specific humidity greater than 14 g kg−1, (b) specific humidity greater than 2 g kg−1, (c) upward motion, and (d) precipitable water greater than 25 mm.

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    NCEP time mean (JJA 1979–95) meridional moisture flux (shaded) and vector moisture flux (units: m s−1 g kg−1) for (a) the bottom 7 sigma levels (approximately 1000–850 hPa), (b) the next 5 sigma levels (approximately 850–550 hPa), (c) the next 6 sigma levels (approximately 550–200 hPa), and (d) the full vertical integral (sigma levels 1–28). For comparison with Rasmusson (1967), the standard vector length in (d) is approximately 15 × 102 g (cm s)−1. Values of meridional flux greater than 1 m s−1 g kg−1 (less than −1 m s−1 g kg−1) are shaded dark (light).

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    Map sequence of observed precipitation represented as the composite mean (1963–94) for 7-day periods centered at (a) day −10, (b) day −3, (c) day +4, and (d) day +11 relative to monsoon onset. The contour interval is 0.5 mm day−1. Values greater than 1.5 mm day−1 are shaded.

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    (a) Longitude–time diagram of the composite mean (1963–94) observed precipitation anomalies (departures from the JJA 1963–94 time mean) averaged between 34° and 38°N. Results are shown for a 3-day running mean. (b) Map of observed precipitation represented as the composite mean (1963–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1). In (a) and (b), the contour interval is 0.25 mm day−1, the zero contour is omitted for clarity, and values greater than 0.25 mm day−1 (less than −0.25 mm day−1) are shaded dark (light).

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    Maps of NCEP (a) 200-hPa wind (units: m s−1) and divergence (units: 10−6 s−1), and (b) 500-hPa vertical velocity (units: μb s−1) represented as the composite mean (1979–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1). In (a), the contour interval is 0.5 × 10−6 s−1, and in (b), the contour interval is 0.05 μb s−1.

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    Map sequence of the NCEP 500-hPa height (units: m) and 500-hPa specific humidity (units: g kg−1) represented as the composite mean (1979–94) for 3-day periods centered at (a) day −12, (b) day −9, (c) day −6, (d) day −3, (e) day 0, (f) day +3, (g) day +6, (h) day +9, and (i) day +12 relative to monsoon onset. The contour interval is 20 m, except for heights above 5880 m where it is 10 m. The light (dark) shading denotes values of specific humidity greater than 2 g kg−1 (greater than 2.5 g kg−1).

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    Maps of (a) surface specific humidity and (b) 500-hPa specific humidity represented as the composite mean (1979–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1) and normalized by the summer time mean (JJA 1979–95) value (units: percent). The contour interval is 5%.

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    Map sequence of the NCEP precipitable water (units: mm) represented as the composite mean (1979–94) for 7-day periods centered at (a) day −10, (b) day −3, (c) day +4, and (d) day +11 relative to monsoon onset. The contour interval is 3 mm, and values greater than 24 mm are shaded.

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    Analysis of the NCEP July minus June (a) percent frequency of criterion 1 jets (Bonner 1968) during the nighttime (0600 UTC and 1200 UTC combined) and (b) vertically integrated, nighttime moisture flux for criterion 1 jets. In each case, the analysis is for 1985–89. The units are (a) in percent and (b) m s−1 g kg−1. The standard vector length in (b) is 2 m s−1 g kg−1.

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    Maps of NCEP (a) precipitation, (b) evaporation, (c) evaporation minus precipitation, and (d) moisture flux divergence represented as the composite mean (1979–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1). In each case, the units are mm day−1 and the contour intervals are . . . , −4, −3, −2, −1, −0.5, −0.25, 0.25, 0.5, 1, 2, 3, 4, . . . . In (a) and (b), values greater than 0.25 mm day−1 (less than −0.25 mm day−1) are shaded dark (light); in (c) and (d), the shading is reversed.

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    Map sequence of the NCEP meridional moisture flux (shaded) and vector moisture flux (units: m s−1 g kg−1) for sigma levels 22–28 (approximately 1000–850 hPa) represented as the composite mean (1979–94) for 7-day periods centered at (a) day −10, (b) day −3 (c) day +4, and (d) day +11 relative to monsoon onset. The standard vector length is 15 m s−1 g kg−1. Values of meridional flux greater than 1 m s−1 g kg−1 (less than −1 m s−1 g kg−1) are shaded dark (light).

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    Same as Fig. 20 except for sigma levels 17–21 (approximately 850–550 hPa). The standard vector length is 10 m s−1 g kg−1. Values of meridional flux greater than 1 m s−1 g kg−1 (less than −1 m s−1 g kg−1) are shaded dark (light).

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    Same as Fig. 20 except for sigma levels 11–16 (approximately 550–200 hPa). The standard vector length is 3 m s−1 g kg−1. Values of meridional flux greater than 0.5 m s−1 g kg−1 (less than −0.5 m s−1 g kg−1) are shaded dark (light).

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Influence of the North American Monsoon System on the U.S. Summer Precipitation Regime

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  • 1 Climate Prediction Center, NOAA/NWS/NCEP, Washington, D.C.
  • | 2 Research and Data Systems Corporation, Greenbelt, Maryland
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Abstract

Key features of the U.S. summer precipitation regime are examined within the context of the evolving North American monsoon system. The focus is on the antecedent and subsequent atmospheric conditions over the conterminous United States relative to the onset of monsoon precipitation over the southwestern United States, which typically begins in early July. The onset of the monsoon in this region is determined using a precipitation index, based on daily observed precipitation for a 31-yr (1963–94) period. Lagged composites of the observed precipitation and various fields from the NCEP–NCAR reanalysis for the period 1979–94 provide a comprehensive picture of atmospheric conditions during the evolution of the U.S. warm season precipitation regime.

The summer precipitation regime is characterized by an out-of-phase relationship between precipitation over the Southwest and the Great Plains–northern tier and an in-phase relationship between precipitation over the Southwest and the East Coast. Changes in the upper-tropospheric wind and divergence fields (mean vertical motion) are broadly consistent with the evolution of this precipitation pattern. Enhanced upper-tropospheric divergence in the vicinity and south of the upper-tropospheric monsoon high coincides with enhanced upper-tropospheric easterlies and Mexican monsoon rainfall after onset. Over the Great Plains and along the northern tier, the middle- and upper-tropospheric flow is more convergent and rainfall diminishes after onset to the north and east of the monsoon high. The frequency of occurrence of the Great Plains low-level jet (LLJ) and southerly moisture transport change little during the evolution. However, LLJ-related precipitation is controlled by changes in the large-scale flow related to the North American monsoon system. There is increased upper-tropospheric divergence and precipitation after onset in the vicinity of an “induced” trough over the eastern United States. The pattern of evaporation minus precipitation from the NCEP–NCAR reanalysis shows broad consistency with the divergence of the vertically integrated flux of water vapor during the monsoon, although the resolution in the NCEP–NCAR reanalysis is inadequate to yield quantitatively accurate regional estimates of these fields. In agreement with earlier studies, the NCEP–NCAR reanalysis indicates that most of the moisture below 850 hPa over the desert Southwest comes from the northern Gulf of California, while most of the moisture at and above 850 hPa arrives from over the Gulf of Mexico.

Corresponding author address: Dr. R. W. Higgins, Analysis Branch, Climate Prediction Center, NOAA/NWS/NCEP, Washington, DC 20233.

Email: wd52wh@sgi85.wwb.noaa.gov

Abstract

Key features of the U.S. summer precipitation regime are examined within the context of the evolving North American monsoon system. The focus is on the antecedent and subsequent atmospheric conditions over the conterminous United States relative to the onset of monsoon precipitation over the southwestern United States, which typically begins in early July. The onset of the monsoon in this region is determined using a precipitation index, based on daily observed precipitation for a 31-yr (1963–94) period. Lagged composites of the observed precipitation and various fields from the NCEP–NCAR reanalysis for the period 1979–94 provide a comprehensive picture of atmospheric conditions during the evolution of the U.S. warm season precipitation regime.

The summer precipitation regime is characterized by an out-of-phase relationship between precipitation over the Southwest and the Great Plains–northern tier and an in-phase relationship between precipitation over the Southwest and the East Coast. Changes in the upper-tropospheric wind and divergence fields (mean vertical motion) are broadly consistent with the evolution of this precipitation pattern. Enhanced upper-tropospheric divergence in the vicinity and south of the upper-tropospheric monsoon high coincides with enhanced upper-tropospheric easterlies and Mexican monsoon rainfall after onset. Over the Great Plains and along the northern tier, the middle- and upper-tropospheric flow is more convergent and rainfall diminishes after onset to the north and east of the monsoon high. The frequency of occurrence of the Great Plains low-level jet (LLJ) and southerly moisture transport change little during the evolution. However, LLJ-related precipitation is controlled by changes in the large-scale flow related to the North American monsoon system. There is increased upper-tropospheric divergence and precipitation after onset in the vicinity of an “induced” trough over the eastern United States. The pattern of evaporation minus precipitation from the NCEP–NCAR reanalysis shows broad consistency with the divergence of the vertically integrated flux of water vapor during the monsoon, although the resolution in the NCEP–NCAR reanalysis is inadequate to yield quantitatively accurate regional estimates of these fields. In agreement with earlier studies, the NCEP–NCAR reanalysis indicates that most of the moisture below 850 hPa over the desert Southwest comes from the northern Gulf of California, while most of the moisture at and above 850 hPa arrives from over the Gulf of Mexico.

Corresponding author address: Dr. R. W. Higgins, Analysis Branch, Climate Prediction Center, NOAA/NWS/NCEP, Washington, DC 20233.

Email: wd52wh@sgi85.wwb.noaa.gov

1. Introduction

A fundamental and necessary first step toward understanding warm season precipitation variability over North America is the clear documentation of the major elements of the warm season precipitation regime within the context of the evolving atmosphere–ocean–land annual cycle. Monsoon circulation systems, which develop over low-latitude continental regions in response to thermal contrast between the continent and adjacent oceanic regions, are a major component of continental warm season precipitation regimes. The North American warm season is characterized by such a monsoon system [hereafter referred to as the North American monsoon system (NAMS)]. This system provides a useful framework for describing and diagnosing warm season climate controls and the nature and causes of year-to-year variability. This system displays many similarities (as well as differences) with its Asian counterpart (e.g., Tang and Reiter 1984). While the NAMS is less impressive than its Asian sister, it still has a tremendous impact on local climate. It is a specific goal of this study to improve our understanding of the large-scale features of the warm season precipitation regime over the United States as it relates to the evolution of the NAMS.

The life cycle and large-scale features of the NAMS can be described using terms typically reserved for the much larger Asian monsoon system; that is, we can characterize the life cycle in terms of development, mature, and decay phases (see section 2 for a description of the life cycle of the NAMS). Attempts to describe and understand the major elements of the NAMS and their evolution have been hampered to a large extent by the lack of adequate in situ and satellite data. Over the last two decades operational data-assimilation systems have provided global gridded atmospheric analyses necessary for climate studies of phenomena such as the monsoon. However, operationally mandated changes in the assimilation systems have introduced inhomogeneities in the data record that have led to difficulties in defining anomalies (e.g., Trenberth and Olson 1988). In addition, computer limitations have generally restricted the range of variables and the number of levels that can be archived. Clearly, advances in describing, understanding, and modeling the warm season precipitation regime over North America will require homogeneous datasets.

The National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP–NCAR reanalysis project currently under way at NCEP in cooperation with NCAR; Kalnay et al. 1996) will provide at least 40 yr (1957–96+) of global gridded fields produced with a fixed state-of-the-art analysis system and a large input database (including data available after the operational cutoff time). This dataset provides many new opportunities for studies of the NAMS. A gridded hourly precipitation database for the United States (1963–95) developed by Higgins et al. (1996a, hereafter HJY96) provides the basis for daily (and hourly) precipitation analyses over regions influenced by the NAMS.

In this study, we exploit the reanalysis data (for the period 1979–94) and the observed precipitation data (for the period 1963–94) to diagnose the large-scale features of the United States warm season precipitation regime. The focus is on the antecedent and subsequent atmospheric conditions over the conterminous United States relative to the onset of monsoon precipitation over the southwestern United States, which typically begins in early July. This is usually the period when the NAMS enters its mature phase, as described in section 2. The onset of monsoon rainfall over the southwestern United States is determined using a precipitation index, based on the daily observed precipitation for a 31-yr (1963–94) period (HJY96), and a threshold crossing procedure. The strategy will be to use lagged composites of various atmospheric circulation fields relative to monsoon onset in the southwestern United States to identify major features of the U.S. summer precipitation regime. Emphasis will be on regional aspects of the warm season precipitation regime and on relationships between precipitation over the southwestern United States and elsewhere over the United States. The overall intent is to provide a comprehensive picture of the warm season precipitation regime of the United States within the context of the evolving NAMS.

Section 2 presents a brief overview of the life cycle of the North American monsoon system. Section 3 describes the datasets and the methodology. Section 4 provides an overview of the climatology of the U.S. summer precipitation regime. The composite evolution of the U.S. summer precipitation regime is discussed in section 5. A summary and discussion are given in section 6.

2. Life cycle of the North American monsoon system

As noted above, the evolution of the North American monsoon system can be characterized in terms of development, mature, and decay phases. The development (May–June phase) is characterized by a period of transition from the cold season circulation regime to the warm season regime. This is accompanied by a decrease in midlatitude synoptic-scale transient activity over the conterminous United States and northern Mexico as the extratropical storm track weakens and migrates poleward to a position near the Canadian border by late June (e.g., Whittaker and Horn 1981; Parker et al. 1989). During this time, there are increases in the amplitude of the diurnal cycle of precipitation (e.g., Wallace 1975; Higgins et al. 1996a) and in the frequency of occurrence of the Great Plains low-level jet (LLJ) (e.g., Bonner 1968; Bonner and Paegle 1970; Augustine and Caracena 1996; Mitchell et al. 1995; Helfand and Schubert 1995; Higgins et al. 1997). The onset of the Mexican monsoon (Douglas et al. 1993; Stensrud et al. 1995) is characterized by heavy rainfall over southern Mexico, which quickly spreads northward along the western slopes of the Sierra Madre Occidental (hereafter SMO) and into Arizona and New Mexico by early July. Precipitation increases over northwestern Mexico coincide with increased vertical transport of moisture by convection (Douglas et al. 1993) and southerly winds flowing up the Gulf of California (Badan-Dangan et al. 1991). Increases in precipitation over the southwestern United States coincide with the development of a pronounced anticyclone at the jet stream level (e.g., Okabe 1995), the development of a thermally induced trough in the desert Southwest (Tang and Reiter 1984; Rowson and Colucci 1992), northward displacements of the Pacific and Bermuda highs (Carleton 1986, 1987), the formation of southerly low-level jets over the Gulf of California (Carleton 1986; Douglas 1995), the formation of the Arizona monsoon boundary (Adang and Gall 1989), and increases in eastern Pacific sea surface temperature gradients (Carleton et al. 1990). From June to July, there is also an increase in sea level pressure over the southwestern United States (Okabe 1995) and a general height increase in midlatitudes associated with the seasonal heating of the troposphere. The largest increases in height occur over the western and southern United States and are likely related to enhanced atmospheric heating over the elevated terrain of the western United States and Mexico (Fig. 1), and increased latent heating associated with the development of the North American monsoon. The resulting middle- and upper-tropospheric “monsoon high” is analogous to the Tibetan high over Asia (e.g., Tang and Reiter 1984) and the warm season Bolivian high over South America (e.g., Johnson 1976).

During the mature (July–August) phase, the NAMS is fully developed and can be related to the seasonal evolution of the continental precipitation regime; a schematic illustrating the key elements of the NAMS during the mature phase is shown in Fig. 2. The monsoon high is associated with enhanced upper-tropospheric divergence in its vicinity and to the south, and with enhanced easterlies (or weaker westerlies) and enhanced Mexican monsoon rainfall (Douglas et al. 1993). To the north and east of the monsoon high, the atmospheric flow is more convergent at middle- and upper-tropospheric levels, and rainfall diminishes from June to July (e.g., Harman 1991). Surges of maritime tropical air northward over the Gulf of California are linked to active and break periods of the monsoon rains over the deserts of Arizona and California (Hales 1972). The mature phase has also been linked with increased upper-level divergence and precipitation in the vicinity of an “induced” trough over the eastern United States (e.g., M. Barlow et al. 1997, manuscript submitted to J. Climate).

The decay (September–October) phase of the NAMS can be characterized as the reverse of the onset phase, although the changes tend to proceed at a slower rate. During this phase, the ridge over the western United States weakens as the monsoon high retreats southward and Mexican monsoon precipitation diminishes. The decay phase is also accompanied by an increase in rainfall over much of the surrounding region (Okabe 1995).

Numerous authors have attempted to identify the primary source of moisture for the summer rains over the southwestern United States. Bryson and Lowery (1955) suggested that horizontal advection of moist air at middle levels from the east or southeast around a westward extension of the Bermuda high might explain the onset of summer rainfall over the southwest; this was later corroborated by Sellers and Hill (1974). Several authors (Hales 1972, 1974; Brenner 1974; Douglas et al. 1993) expressed skepticism of this type of explanation since moisture from the Gulf of Mexico would first have to traverse the Mexican Plateau and the SMO (see Fig. 1) before contributing to Arizona rainfall. Rasmusson (1967) was among the first to show a clear separation between water vapor east of the continental divide, which clearly originates from the Gulf of Mexico–Caribbean Sea, and moisture over the Sonoran Desert, which appears to originate from the Gulf of California. Schmitz and Mullen (1996) examined the relative importance of the Gulf of Mexico, the Gulf of California, and the eastern tropical Pacific as moisture sources for the Sonoran Desert using European Centre for Medium-Range Weather Forecasts (ECMWF) analyses. They found that most of the moisture at upper levels over the Sonoran Desert arrives from over the Gulf of Mexico, while most of the moisture at lower levels comes from the northern Gulf of California.

3. Data analysis

The NCEP–NCAR assimilation system consists of the NCEP Medium Range Forecast spectral model and the operational NCEP Spectral Statistical Interpolation (Parrish and Derber 1992) with the latest improvements (Kalnay et al. 1996). The assimilation is performed at a horizontal resolution of T62 and 28 sigma levels in the vertical, with 7 levels below 850-hPa. In this study, we utilize the reanalysis heights, winds, and specific humidity, which are instantaneous fields available every 6 h, and several diagnostic fields (e.g., precipitation and evaporation) that are generated by the GCM’s physical parameterizations. In the NCEP system, the precipitation and evaporation fields are based on a 6-h forecast valid at the initial synoptic time. Moisture transport was computed directly from winds and specific humidity on the respective model sigma levels. These calculations were performed in the spherical domain to minimize truncation errors. All reanalysis fields are averaged to daily values prior to compositing as described below.

In order to study the summer precipitation regime over the United States, relative to the onset of monsoon rainfall over the southwestern United States, we use a monsoon index based on hourly, gridded precipitation analyses over the conterminous United States (HJY96); these analyses were developed from station observations obtained from the National Weather Service (NWS)/Techniques Development Laboratory. The time domain covers the period 1 January 1963 through 31 December 1994. The analyses were gridded to a horizontal resolution of 2° latitude × 2.5° longitude. We note that in this study the term “rainfall” is equivalent to measurable precipitation.

The procedure for identifying the onset of the summertime rains over Arizona and New Mexico is briefly summarized below. A precipitation index (PI) was obtained by averaging daily accumulations of observed precipitation at each grid point of the rectangular region (112.5°–107.5°W, 32°–36°N) over Arizona and western New Mexico (see the box in Fig. 1). Histograms of the mean (1963–94) daily rainfall (and the 5-day running mean) during summer at each grid point over the southwestern United States (Fig. 3) show that all of the grid points used for the PI exhibit rapid increases in rainfall by early July. Care was used in choosing the grid points for the PI because Arizona exhibits a “pure” monsoon signal—that is, a sudden onset of monsoon rains—while New Mexico has a more gradual increase due to mixed influences (Fig. 4); western New Mexico has a pure monsoon signal, but eastern New Mexico is influenced by the dryline and the Great Plains low-level jet. Monsoon onset dates were determined using the resulting time series and a threshold-crossing procedure. The PI magnitude and duration criteria used were +0.5 mm day−1 and 3 days, respectively; the monsoon onset date for each year occurs when the selection criteria are first satisfied after 1 June. Composite evolution fields for 1963–94 (observed precipitation) and 1979–94 (reanalysis fields) were obtained by averaging over all of the monsoons relative to the day when the PI first satisfies the threshold criteria; this day is designated as the onset day, or day 0. Note that by realigning the time series in this way we are not performing a simple average based on calendar day. Table 1 shows the onset date of the monsoon over the southwestern United States for each year based on this index; the average date of the monsoon onset for the period 1963–94 is 7 July.

The composite evolution of the PI (Fig. 5) shows the onset of the southwestern rains. It is important to note that the compositing scheme based on the PI makes the monsoon onset appear to be abrupt because it is keyed to synoptic as well as climate variability, as evidenced by the pronounced overshoot in Fig. 5. Composites of observed precipitation and various fields from the NCEP–NCAR reanalysis (including heights, winds, moisture, moisture flux, precipitation, evaporation, and sea level pressure) were constructed each day for the 90-day period from 45 days prior to onset (day −45) to 44 days after onset (day +44).

In section 4a the precipitation of HJY96 is compared to the monthly mean global merged analysis of Xie and Arkin (1996, hereafter XA96). Input data for the merged analysis, which is gridded at a resolution of 2.5° latitude × 2.5° longitude, include gauge observations over land, atoll gauge data, satellite estimates (including IR data from the GOES Precipitation Index, microwave scattering estimates, and microwave emission estimates), and numerical model predictions (based on 12–36-h ECMWF forecasts). This analysis was available for an 8-yr period from July 1987 to June 1995.

4. Features of the U.S. summer precipitation regime

In this section, we examine the large-scale features of the U.S. summer precipitation regime. A primary focus is the characteristics of the monsoon rainfall over the southwestern United States, including the geographical extent, timing, and local characteristics. Relationships between monsoon rainfall in southwestern North America and rainfall elsewhere over the conterminous United States are discussed. The mean distributions of the NCEP wind, vertical velocity, moisture, and moisture transport fields are examined and compared to previous studies. Our purpose is not to present a detailed discussion of the North American mean summer precipitation regime, but rather to provide the proper background for interpreting changes in these fields or anomalies (departures from mean summertime values) during the evolution of the monsoon (see section 5).

a. Precipitation

Rainfall associated with the Mexican monsoon is clearly evident in Fig. 6a, which shows mean monthly precipitation for July–September (JAS), usually the 3 rainiest summer months, from the XA96 merged analysis based on monthly (1987–95) data. The XA96 dataset shows heavy precipitation west of the SMO along the west coast of Mexico (Fig. 6a), with local mounts in excess of 160 mm. Atlases of monthly mean precipitation (e.g., WMO 1975) show that the greatest July precipitation occurs west of the SMO, with amounts ranging from more than 250 mm in Sinaloa to 50 mm in southern Arizona. Both the XA96 analysis and the HJY96 analysis, based on hourly (1963–94) gauge data (Fig. 6b), show that New Mexico experiences slightly more rainfall than Arizona. Rainfall in excess of 80 mm is found in both datasets from the Great Plains to the East Coast (except for the Great Lakes region in Fig. 6b). We note that when the same period is used in both datasets (i.e., 1987–95), the basic characteristics of the comparison are much the same.

The extension of the Mexican monsoon rainfall into the southwestern United States is clearly evident in Figs. 6c and 6d, which show the ratio (expressed in percent) of rain falling during the 3-month period July–September to the annual-mean precipitation; Douglas et al. (1993) used this ratio as a “monsoonal index.” The highest monsoon index values (exceeding 70%) are found along the west coast of Mexico; similar results were found by Douglas et al. (1993). The maximum monsoon index values extend northward along the axis of the SMO, northeastward along the Rio Grande Valley in New Mexico, and into the high plains of southeastern Colorado and western Kansas (Figs. 6c and 6d). As found by Douglas et al. (1993), New Mexico appears to be the U.S. state most affected by the monsoon.

The contribution of the summer monsoon rainfall during JAS to the annual total does not reveal the intraseasonal variations in rainfall over the United States. To show these changes, the contributions to the annual total precipitation for each month between June and September were calculated using the XA96 and the HJY96 datasets; results are similar for each dataset, so only the XA96 result is shown (Fig 7). Striking features include 1) the rapid increase in values along the west coast of Mexico and over southeastern Arizona and New Mexico from June to July, 2) the decrease and eastward migration of the maximum values along the northern tier between June and August, 3) the rapid decrease in contributions along the west coast of Mexico from August to September, and 4) the southeastward movement of the maximum values from the monsoon region during July and August to northeastern Mexico by September. The increases in September over northeastern Mexico may be due to increases in the frequency of land-falling tropical storms during this month.

Climatological aspects of the onset of the summer rains over the southwestern United States and western Mexico can be viewed from maps of the mean rainfall difference between June and July (Fig. 8), when the largest monthly variation in rainfall occurs for the southwestern United States. The XA96 analysis (Fig. 8a) shows that the most dramatic increases in rainfall over the continent occur along the west slopes of the SMO, where local increases greater than 160 mm are found. Douglas et al. (1993) reported increases in rainfall of more than 200 mm along the western slopes and foothills of the SMO. Over the southwestern United States, both the XA96 and the HJY96 data (Fig. 8b) show increases exceeding 20 mm over much of Arizona and New Mexico; there is a rapid decrease in the rainfall change to the east and north of this region and an increase in rainfall along the East Coast.

Figure 8 suggests that the summer precipitation regime over the conterminous United States is characterized by coherent phase relationships between rainfall over the southwest and rainfall elsewhere. Over the Great Plains and along the northern tier, the rainfall decreases between June and July by roughly the same amount as the rainfall increases in the southwestern United States. Along the eastern seaboard and portions of the Gulf Coast, there are rainfall increases between June and July. Previous studies have linked the onset of summer rains over northern Mexico and the southwestern United States to a decrease of rainfall over the Great Plains (e.g., Mock 1996; Tang and Reiter 1994; Douglas et al. 1993; A. Douglas 1996, personal communication) and to an increase of rainfall along the East Coast (Tang and Reiter 1984). In section 5a, lagged composites of the observed precipitation field are used to explore how these phase relationships change during the evolution.

Histograms of the mean monthly precipitation at various locations around the conterminous United States reveal other aspects of these regional relationships (Fig. 9). Over Arizona (Fig. 9a), the maximum precipitation occurs in August during the peak of the monsoon. Over Texas (Fig. 9b) and Oklahoma (Fig. 9c), there are two peaks (May and September), with a relative minimum in rainfall during July and August. Similar behavior is found over Montana (Fig. 9d), although the September maximum is much weaker. This out-of-phase relationship is consistent with changes in the large-scale circulation and moisture transport over the western and central United States (sections 5b and 5e). Rainfall over North Carolina (Fig. 9e), the Florida panhandle (Fig. 9f), and Kentucky (Fig. 9g) exhibits a single peak during July. This in-phase relationship is likely related to an induced trough over the eastern United States, which appears after monsoon onset (see section 5b).

b. Wind and moisture fields

The large-scale, low-level flow over the southern United States and Mexico is strongly influenced by the subtropical highs (Fig. 10a), with brisk southerlies over the southern Great Plains (reflecting the Great Plains low-level jet) and northwesterlies west of Baja California (reflecting the Baja jet). Weak evidence of the thermal low can be seen in the cyclonic winds over the desert Southwest. The NCEP–NCAR reanalysis has difficulty capturing the southerly component in the surface winds over the northern Gulf of California and southern Arizona found in observational studies (e.g., Badan-Dangon et al. 1991; Douglas et al. 1993). Some of the problems with the strength of the southerly component found in the NCEP–NCAR reanalysis surface winds may be a ramification of the model’s horizontal resolution, which is too coarse to resolve the Gulf of California and the surrounding terrain (e.g., Schmitz and Mullen 1996). The winds at 500 hPa are characterized by easterlies over the Tropics, widespread westerlies poleward of roughly 35°N, and an anticyclonic circulation centered over the Rio Grande Valley. We note that at 500 hPa there is a strong southerly component in the winds around the west side of the monsoon high.

The 500-hPa vertical velocity (Fig. 10c) indicates localized ascent over the Great Basin, the southern Rockies, the SMO, and the southeastern United States. Descent is situated over the plains states, extending from the Gulf Coast to the Canadian border, the Great Lakes, the east Pacific Ocean, and the northern Gulf of California. The region of upward motion is consistent with precipitation (see sections 4a and 5a), while the overall distribution of vertical velocity is in qualitative agreement with the June–August (JJA) mean of Oort (1983). The ascent in the ECMWF analyses (Schmitz and Mullen 1996) is much stronger than that found here.

The distribution of specific humidity at the surface (Fig. 10a) strongly reflects the underlying terrain, with high values flanking the southern Rockies and the Mexican Plateau. Surface humidity increases eastward into the southeastern United States, and a moist tongue extends up the Gulf of California and along the western coast of Mexico. In the middle troposphere (Fig. 10b), a band of enhanced moisture extends from the eastern tropical Pacific into Mexico and then New Mexico. The precipitable water (Fig. 10d) indicates abundant moisture over the eastern tropical Pacific, the southern Gulf of California and Baja California, western Mexico, and the eastern half of the United States; similar moisture distributions have been discussed by numerous authors (Starr et al. 1965; Hales 1974; Hagemeyer 1991; Douglas et al. 1993; Negri et al. 1994; Schmitz and Mullen 1996).

The vertical structure of the moisture flux is highlighted in Fig. 11, which presents the fluxes as vertical integrals over three different layers, as well as over the full vertical extent of the atmosphere; this presentation highlights differences in the lower-, middle-, and upper-tropospheric levels. On the whole, the vector flux for the full vertical integral (Fig. 11d) bears close resemblance to the surface wind field (Fig. 10a), due to the large specific humidity at low levels. The strongest flux onto the continent occurs at low levels, below 850 hPa, over the south-central United States and northeastern Mexico in the vicinity of the Great Plains low-level jet. Strong southward flux associated with the large-scale circulation of the east Pacific anticyclone occurs off the west coast of the United States and Mexico. The flux vectors over the Mexican Plateau tend to be easterly and are noticeably smaller than in neighboring regions to the east and west, indicating little transport of moisture across Mexico at low levels. Along the west coast, the low-level flow associated with the Baja jet is primarily parallel to the continent with little influx of moisture (Fig. 11a). The onshore transport of moisture from the Gulf of California into southwestern Arizona is relatively weak, with little evidence of a time mean transport of moisture at low levels from the tropical Pacific into the desert Southwest; difficulties in explaining the moisture transport in this region remain, in part, because Baja California and the Gulf of California are not properly resolved.

The flux at middle levels (Fig. 11b) can be characterized as a large-scale rotation about the subtropical highs, yielding easterly transport over the Tropics and western Mexico, southerly transport over the northern Gulf of California and the desert Southwest, and westerly transport over much of the United States poleward of 35°N. Note that there is more meridional transport into the Sonoran Desert region at midlevels and more easterly transport across the Mexican Plateau. As in Schmitz and Mullen (1996), we find that there is a relative minimum in the meridional moisture transport near 25°N, 110°W and that most of the moisture transport out of the eastern tropical Pacific veers to the west.

In summary, the mean distributions of the NCEP wind, vertical velocity, moisture, and moisture transport are qualitatively consistent with results from prior studies. With the noted exception of surface winds over the northern Gulf of California, the NCEP–NCAR reanalysis seems to generate quantitatively accurate time mean fields.

5. Evolution of the U.S. summer precipitation regime

The precipitation index discussed in section 3 is used to examine the evolution of the summertime precipitation regime over the United States. In particular, we examine lagged composites of the observed precipitation, as well as the analyzed tropospheric circulation and moisture fields relative to the onset of monsoon precipitation in Arizona and New Mexico. Composites of observed precipitation in section 5a are based on daily accumulations from HJY96 for the period 1963–94. Composites in sections 5b–5e are obtained from the NCEP–NCAR reanalysis for the period 1979–94.

a. Precipitation

A map sequence of composite mean (1963–94) observed precipitation for 7-day periods (Fig. 12) provides some details of the spatial variation of precipitation over the conterminous United States during the evolution of the monsoon. Prior to onset, relatively heavy rainfall is observed over the Great Plains region, while relatively light rainfall is observed in the Southwest and along the East Coast (Fig. 12a). A longitude–time diagram of composite mean observed precipitation anomalies (departures from the JJA 1963–94 mean) averaged between 34° and 38°N shows that this precipitation pattern persists for several weeks prior to monsoon onset (Fig. 13a).

During and after onset the rainfall increases in the Southwest and along the East Coast (roughly 1 week after onset) and diminishes over the Great Plains (Figs. 12c and 12d). Increases in Arizona, New Mexico, southern Utah, southern Colorado, and southeastern California are associated with the extension of the summer monsoon rains into the southwestern United States. The precipitation anomaly pattern (Fig. 13a) takes on the opposite phase after onset; this phase of the anomaly pattern also persists for several weeks. The eastward spread of negative anomalies and the transition to the summer precipitation regime occurs very rapidly around the time of onset (Fig. 13a). Figure 13a also shows a decrease in rainfall over the central United States roughly 5 days prior to onset. It is evident that the summer rainfall over the central and particularly the eastern United States is much more episodic than that in the Southwest monsoon region.

A map of the change in composite mean observed precipitation, represented as the difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1) (Fig. 13b), yields a pattern quite similar to the intermonthly (July minus June) precipitation map discussed earlier; this result is consistent with the fact that the southwestern rains typically appear near the beginning of July. After onset, the Great Plains and the northern tier experience a drying trend as a result of the strengthened and expanded middle- and upper-tropospheric monsoon high (see section 4b), but mesoscale convective activity, occasional upper-level disturbances, localized mountain valley circulations, and nocturnal LLJ-related precipitation keep the region relatively wet compared to the drier winter months (Tang and Reiter 1984; Augustine and Caracena 1996; Higgins et al. 1997). Some large-scale climatic controls on the spatial variation of precipitation during the monsoon are discussed in the following sections.

b. Tropospheric circulation and moisture

Changes in the tropospheric circulation and divergence (mean vertical motion) fields during the evolution can be related to changes in the continental precipitation regime. A region of enhanced upper-tropospheric divergence in the vicinity and south of the monsoon high coincides with enhanced upper-tropospheric easterlies (or weaker westerlies) (Fig. 14a), enhanced midtropospheric vertical motion (Fig. 14b), and enhanced monsoon rainfall (Fig. 13b). In contrast, the flow is more convergent, and rainfall diminishes to the north and east of the monsoon high (Fig. 14a), particularly over the Great Plains in the vicinity of strong midtropospheric subsidence (Fig. 14b); the monsoon high organizes the flow at middle and upper levels over the Great Plains, thereby acting as an effective control on LLJ-related rainfall (see section 5c).

Near the surface, the divergence field (not shown) is broadly consistent with the strongest low-level convergence located in the intertropical convergence zone, along the west coast of Mexico, and over southeastern Arizona. A compensating pattern of low-level divergence is found over the Great Plains, along the northern tier, and in the Pacific Northwest. There is some evidence of increased upper-tropospheric divergence (Fig. 14a), vertical motion (Fig. 14b), and precipitation (Fig. 13b) in the vicinity of an induced trough over the eastern United States, which is also evident in composites of the mean sea level pressure field after onset (not shown). Typically, enhanced rainfall appears along the East Coast roughly 10 days after onset (Fig. 13a), which may be a characteristic lag time for the development of the eastern trough.

A map sequence of the composite mean 500-hPa height field (Fig. 15) shows clear evidence of the northward shift of the monsoon ridge center from north-central Mexico to the Arizona–New Mexico region. Typically, the monsoon ridge extends eastward a bit after onset, where it meets a westward extension of the Bermuda high, resulting in an east–west axis of maximum heights across the southern United States–northern Gulf Coast region.

The large-scale, low-level reanalysis flow changes little during the evolution, but the middle- and upper-level flows change considerably. Prior to onset, the winds at 500 hPa (inferred from the height field on Fig. 15) show a large anticyclonic rotation around the monsoon ridge centered over northwestern Mexico. Southwesterly flow predominates in the southwestern United States, while easterly flow is generally confined south of 25°N. As the monsoon ridge shifts northward, deep easterlies become firmly established over all of Mexico after onset except for northwestern Baja California, where the winds are southerly. West of the monsoon ridge, the winds become southerly over Arizona and southern California, while near the center of the monsoon ridge, the winds remain light and variable. While the southwesterly flow over much of the western United States persists throughout the evolution, the westerlies over the eastern half of the United States turn more northwesterly as the monsoon ridge shifts northward and the induced trough develops along the East Coast.

During the evolution, the moisture at 500 hPa expands northward along the west coast of Mexico into Arizona and New Mexico on the southern and western flanks of the monsoon ridge (Fig. 15). Examination of the vertical structure of the moisture transport (see section 5e) shows that much of the moisture at and above 850 hPa over the desert Southwest originates over the Gulf of Mexico. Although moisture is concentrated in the lower troposphere over southwestern North America, the relative change in specific humidity during the evolution is much larger and more widespread in the middle troposphere (Fig. 16), where local increases of more than 50% (over summertime mean values) are found over a large portion of southwestern North America, the eastern Pacific, and the Gulf of California. At the head of the Gulf of California, the increases in moisture and southerly flow lead to larger meridional transports of moisture at middle levels than at lower levels (see Figs. 20–22, section 5e). Elsewhere over North America the distribution of specific humidity shows relatively minor changes.

A map sequence of composite mean 7-day-averaged precipitable water (Fig. 17) shows clear evidence of a “surge” of moisture from the Gulf of California into Arizona and from western Texas–northeastern Mexico into New Mexico. These low-level surges of moisture are identical to “bursts” of monsoonal activity (e.g., Carleton 1986). For the 4-week period shown, precipitable water increases of more than 50% (when normalized by summer mean values) are found over Arizona, New Mexico, southeastern California, northwestern Mexico, northern Baja California, and the northern Gulf of California. Increases of more than 25% extend over Colorado, Utah, southern Wyoming, southern Nevada, southeastern California, and the remainder of Baja California. Over the eastern United States, the changes in the precipitable water are quite small except over eastern New England, where local increases of roughly 25% are found.

c. Impact on LLJ-related rainfall

In a study of the influence of the Great Plains low-level jet on summertime precipitation over the United States, Higgins et al. (1997) found that enhanced LLJ-related rainfall over the Great Plains was associated with suppressed rainfall over the southwestern United States and along the East Coast. Their precipitation anomaly pattern (see Fig. 14 of that study) is strikingly similar to the one shown in Fig. 13b, except that it is for the period prior to monsoon onset. Recall from Fig. 13a, however, that there is a marked decrease in LLJ-related rainfall over the Great Plains after onset. To determine whether this decrease in rainfall after onset is due to some fundamental change in the LLJ, we examined the average frequency of occurrence of criterion 1 jets (Bonner 1968) during the nighttime hours for May–August 1985–89 using the NCEP–NCAR reanalysis and the approach outlined in Higgins et al. (1997). Interestingly, we find very little intermonthly change in the LLJ frequency maxima over the Great Plains, with a relatively small increase in LLJ frequency from June to July over the southern plains (Fig. 18a). Intermonthly changes in moisture transport for criterion 1 jets (Fig. 18b) show some evidence of the impact of the middle- and upper-tropospheric ridge over western North America, but there is little change in the moisture transport over the Great Plains (also see Fig. 20 in section 5e). This implies that LLJ-related rainfall is controlled by changes in the large-scale flow related to the evolution of the North American monsoon system. In particular, the northward migration of the upper-tropospheric monsoon high (Fig. 15) leads to upper-tropospheric convergence (Fig. 14a) and midtropospheric subsidence (Fig. 14b) over the Great Plains, thereby limiting LLJ-related rainfall after onset.

d. Atmospheric balance of water vapor

Some idea of the atmospheric balance of water vapor is given in Fig. 19, which shows the change in precipitation, evaporation, evaporation minus precipitation, and moisture flux divergence during the evolution; all composites are from the NCEP–NCAR reanalysis. The region of enhanced moisture flux convergence in the vicinity of and south of the upper-tropospheric monsoon high (Fig. 19d) coincides with enhanced monsoon rainfall (Fig. 19a). The largest values of enhanced convergence are situated along the western slopes of the SMO, where P exceeds E by as much as 5 mm day−1 (Fig. 19c). The band of enhanced convergence extends northward into the Sonoran Desert region and New Mexico, and southward into the tropical east Pacific and along the intertropical convergence zone. Another region of enhanced convergence is located along the East Coast, where enhanced precipitation also occurs (Fig. 19a). Enhanced moisture flux divergence prevails over the Great Plains and along the northern tier, where rainfall diminishes during the summer (Fig. 19a). We also find enhanced moisture flux divergence over the southern Rio Grande valley, in agreement with earlier studies (Rasmusson 1966; Roads et al. 1994; Schmitz and Mullen 1996). This may be quite realistic because the observed May–June rainfall in this region is considerably larger than the July–August rainfall (e.g., Douglas et al. 1993); the heavier spring rainfall could provide the potential for evaporation to exceed precipitation during the summer. Overall, the pattern of evaporation minus precipitation (Fig. 19c) shows broad consistency with the moisture flux divergence pattern (Fig. 19d).

e. Atmospheric flux of water vapor

For the southern United States, Rasmusson (1967) demonstrated a clear separation between southerly water vapor transport east of the continental divide, which clearly originates from the Gulf of Mexico–Caribbean Sea, and southerly transport over the desert Southwest (roughly west of 110°W) that appears to originate from the Gulf of California; these “moisture pipelines” play an integral role in the warm season precipitation regime, especially over the Great Plains and the desert Southwest, respectively.

The vertical structure of the moisture flux during the evolution is highlighted in Figs. 20–22, which present fluxes over three different layers to highlight the lower-(∼1000–850 hPa), middle- (∼850–550 hPa), and upper-(∼550–200 hPa) tropospheric levels, respectively (please note that the standard vector lengths are different in each figure).

On the whole, the flux in the lower troposphere (Fig. 20) resembles the surface wind field (Fig. 10a) throughout the evolution. Weak meridional transport of moisture into the desert Southwest from the northern Gulf of California appears to be separated from a second feature off the southwest coast of Mexico. This second feature, which emanates from the tropical eastern Pacific, shows a tendency to strengthen after onset, but flows no farther than about 25°N. Based solely on lower levels, it appears that moisture over the desert Southwest region comes predominantly from the northern Gulf of California without a direct link to the tropical eastern Pacific. During the evolution, there seems to be very little impact on the southerly transports at low levels associated with the Great Plains low-level jet (see section 5c).

At middle-tropospheric levels (Fig. 21), the flux can be characterized as a large-scale rotation about the monsoon high. After onset, there are increases in the easterly transport over northern Mexico and southerly transport over the southwestern United States, which appear to be a direct consequence of the northward shift of the monsoon ridge. After onset, a relative minimum in the meridional transport near 25°N over southern Baja California appears, with a local maximum in meridional transport to the south along the west coast of Mexico (directly above the maximum in lower-tropospheric meridional transport). The transport over northern Mexico is primarily from the east after onset. Overall, the pattern suggests that much of the moisture at midlevels over the southwestern United States originates over the Gulf of Mexico. At upper-tropospheric levels (Fig. 22), moisture is transported into the western United States on the western flank of the upper-tropospheric monsoon ridge.

6. Summary and discussion

The North American warm season is characterized by a monsoon system that provides a useful framework for describing and diagnosing the large-scale features of the summertime precipitation regime over the United States. This regime is characterized by an out-of-phase relationship between precipitation over the Southwest and the Great Plains–northern tier and an in-phase relationship between precipitation over the Southwest and the East Coast. Prior to onset, enhanced LLJ-related rainfall over the Great Plains is associated with suppressed rainfall over the desert Southwest and along the East Coast. As the Mexican monsoon extends into the desert Southwest, typically during early July, the opposite phase of this pattern appears. In most years, this transition occurs quite rapidly.

Changes in the continental precipitation regime during the monsoon are related to changes in the tropospheric circulation and divergence. Enhanced upper-tropospheric divergence in the vicinity and south of the monsoon high coincides with enhanced upper-tropospheric easterlies, vertical motion, and monsoon rainfall. A shift in the middle- and upper-tropospheric winds from westerly to easterly over northwestern Mexico and from southwesterly to southerly over Arizona was noted during the onset of the southwestern rains. Over the Great Plains and along the northern tier, the upper-tropospheric flow is convergent, and rainfall diminishes as a result of the strengthened and expanded middle- and upper-tropospheric monsoon high. The East Coast experiences wetter conditions after onset in response to a weak trough, which typically lags the onset of the southwestern rains by roughly 10 days.

Our earlier studies have shown that the Great Plains LLJ has a considerable impact on the distribution and the intensity of nighttime precipitation over the central United States during the summer season (Higgins et al. 1997). During the monsoon, however, there is little change in the intensity of the LLJ or in the southerly moisture transport from the Gulf of Mexico, despite a considerable reduction in LLJ-related rainfall after onset. LLJ-related precipitation is limited by changes in the large-scale flow related to the North American monsoon—that is, by the northward migration of the monsoon high, which helps organize the tropospheric flow over the Great Plains.

Changes in the pattern of evaporation minus precipitation from the NCEP–NCAR reanalysis show broad consistency with changes in the divergence of the vertically integrated flux of water vapor during the evolution. While the difference patterns appear to be quite realistic, the magnitude of the flux divergence field is questionable. A recent study of the regional balance of water vapor over the central United States (Higgins et al. 1996b) showed that the model bias is as large as the divergence itself in both the NCEP–NCAR reanalysis and the National Aeronautics and Space Administration/Data Assimilation Office reanalysis (Schubert et al. 1993). The extent to which such imbalances are related to topography, resolution (space and time truncation errors), and model forecast errors is still an open question. We note that the source region over the lower Rio Grande Valley found in earlier studies (e.g., Schmitz and Mullen 1996) appears to be real, though higher resolution is required for regional features; in our calculations, the moisture transport and its divergence were computed directly from winds and specific humidity on the NCEP–NCAR reanalysis model sigma levels.

The NCEP–NCAR reanalysis indicates that most of the moisture below 850 hPa over the desert Southwest comes from the northern Gulf of California, while most of the moisture at and above 850 hPa arrives from over the Gulf of Mexico. Similar results were recently reported by Schmitz and Mullen (1996), who also concluded (using ECMWF analyses) that the primary upper-level stream of moisture over the desert Southwest originates over the Gulf of Mexico. The reanalysis shows little evidence of a time mean transport of moisture at low levels from the tropical Pacific into the desert Southwest. However, difficulties in explaining the observed precipitation distribution and its timing from available upper-air wind and moisture analyses remain, in part because Baja California and the Gulf of California are not properly resolved. In addition, the proximate source(s) of moisture for the Southwest monsoon rains cannot be determined until we understand how precipitable moisture over the Southwest is linked to the vertical structure of the moisture transport. In general, finer-resolution models and long-term observations will be required to understand regional-scale impacts of the monsoon.

It is hoped that our analyses will provide a useful foundation for follow-on studies of the variability of the warm season precipitation regime over the United States. For example, warm season precipitation needs to be examined further for possible relationships with internal atmospheric dynamics, ocean–atmosphere interactions, and land–atmosphere interactions. Although these processes operate together, it is also necessary to isolate them in order to determine their relative contributions to the development, maintenance, and decay of warm season hydrologic anomaly regimes. Our analyses also highlight the critical need for adequate in situ precipitation data and upper-air analyses at sufficiently high temporal resolution and in a form convenient for analysis. Recent efforts to collect daily precipitation, temperature, and radiosonde data from Mexico with good spatial and temporal coverage (A. Douglas 1996, personal communication) should provide new opportunities for the development of convenient datasets covering both Mexico and the United States at sufficient temporal and spatial resolution for comprehensive studies of the warm season precipitation regime over North America.

Acknowledgments

We wish to thank Jess Charba of NOAA/Techniques Development Laboratory for the precipitation data and the NCEP–NCAR reanalysis team for the reanalysis data. Special thanks go to John Janowiak and Kingtse Mo for thorough reviews of an early version of this manuscript, and to Muthuvel Chelliah for Fig. 4. The authors are also indebted to Hugo Berbery, Vernon Kousky, Gene Rasmusson, Chester Ropelewski, and Jae Schemm for insightful discussions. This work was partially supported by the NOAA Office of Global Programs under the Pan American Climate Studies project, by the GEWEX Continental-Scale International Project, and by Interagency Agreement S-41376-F under the authority of NASA/GSFC.

REFERENCES

  • Adang, T. C., and R. L. Gall, 1989: Structure and dynamics of the Arizona monsoon boundary. Mon. Wea. Rev.,117, 1423–1438.

  • Augustine, J. A., and F. Caracena, 1994: Lower-tropospheric precursors to nocturnal MCS development over the central United States. Wea. Forecasting,9, 116–135.

  • Badan-Dangon, A., C. E. Dorman, M. A. Merrifield, and C. D. Winant, 1991: The lower atmosphere over the Gulf of California. J. Geophys. Res.,96, 16877–16896.

  • Bonner, W. D., 1968: Climatology of the low-level jet. Mon. Wea. Rev.,96, 833–850.

  • ——, and J. Paegle, 1970: Diurnal variations in the boundary layer winds over the south-central United States in summer. Mon. Wea. Rev.,98, 735–744.

  • Brenner, I. S., 1974: A surge of maritime tropical air—Gulf of California to the southwestern United States. Mon. Wea. Rev.,102, 375–389.

  • Bryson, R. A., and W. P. Lowry, 1955: The synoptic climatology of the Arizona summer precipitation singularity. Bull. Amer. Meteor. Soc.,36, 329–339.

  • Carleton, A. M., 1986: Synoptic-dynamic character of “bursts” and “breaks” in the southwest U.S. summer precipitation singularity. J. Climatol.,6, 605–623.

  • ——, 1987: Summer circulation climate of the American Southwest, 1945–1984. Ann. Assoc. Amer. Geogr.,77, 619–634.

  • ——, D. A. Carpenter, and P. J. Weser, 1990: Mechanisms of interannual variability of the southwest United States summer rainfall maximum. J. Climate,3, 999–1015.

  • Douglas, M. W., 1995: The summertime low-level jet over the Gulf of California. Mon. Wea. Rev.,123, 2334–2347.

  • ——, R. A. Maddox, K. Howard, and S. Reyes, 1993: The Mexican monsoon. J. Climate,6, 1665–1677.

  • Hagemeyer, B. C., 1991: A lower-tropospheric climatology for March through September. Some implications for thunderstorm forecasting. Wea. Forecasting,6, 254–270.

  • Hales, J. E., Jr., 1972: Surges of maritime tropical air northward over the Gulf of California. Mon. Wea. Rev.,100, 298–306.

  • ——, 1974: Southwestern United States summer monsoon source—Gulf of Mexico or Pacific Ocean? J. Appl. Meteor.,13, 331–342.

  • Harman, J. R., 1991: Synoptic Climatology of the Westerlies: Process and Patterns. Association of American Geographers, 80 pp.

  • Helfand, H. M., and S. D. Schubert, 1995: Climatology of the Great Plains low-level jet and its contribution to the continental moisture budget of the United States. J. Climate,8, 784–806.

  • Higgins, R. W., J. E. Janowiak, and Y. Yao, 1996a: A gridded hourly precipitation data base for the United States (1963–1993). NCEP/Climate Prediction Center Atlas 1, 47 pp. [Available from Climate Prediction Center, NOAA/NWS/NCEP, Washington, DC 20233.].

  • ——, K. C. Mo, and S. D. Schubert, 1996b: The moisture budget of the central United States in spring as evaluated in the NCEP/NCAR and the NASA/DAO reanalyses. Mon. Wea. Rev.,124, 939–963.

  • ——, Y. Yao, E. S. Yarosh, J. E. Janowiak, and K. C. Mo, 1997: Influence of the Great Plains low-level jet on summertime precipitation and moisture transport over the central United States. J. Climate,10, 481–507.

  • Johnson, A. M., 1976: The climate of Peru, Bolivia and Ecuador. Climates of Central and South America, W. Schwerdtfeger and H. E. Landsberg, Eds., World Survey of Climatology, Vol. 12, Elsevier, 147–218.

  • Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-year reanalysis project. Bull. Amer. Meteor. Soc.,77, 437–471.

  • Mitchell, M. J., R. A. Arritt, and K. Labas, 1995: A climatology of the warm season Great Plains low-level jet using wind profiler observations. Wea. Forecasting,10, 576–591.

  • Mock, C. J., 1996: Climatic controls and spatial variations of precipitation in the western United States. J. Climate,9, 1111–1125.

  • Negri, A. J., R. F. Adler, E. J. Nelkin, and G. J. Huffman, 1994: Regional rainfall climatologies derived from Special Sensor Microwave Imager (SSM/I) data. Bull. Amer. Meteor. Soc.,75, 1165–1182.

  • Okabe, I. T., 1995: The North American monsoon. Ph.D. dissertation, University of British Columbia, 146 pp. [Available from Dept. of Geography, University of British Columbia, 2075, Wesbrook Place, Vancouver, BC V6T1W5, Canada.].

  • Oort, A. H., 1983: Global atmospheric circulation statistics 1958–1973. NOAA Professional Paper 14, U.S. Government Printing Office, Washington, DC, 180 pp.

  • Parker, S. S., J. T. Hawes, S. J. Colucci, and B. P. Hayden, 1989: Climatology of 500-mb cyclones and anticyclones, 1950–1985. Mon. Wea. Rev.,117, 558–570.

  • Parrish, D. F., and J. C. Derber, 1992: The National Meteorological Center’s spectral statistical interpolation analysis system. Mon. Wea. Rev.,120, 1747–1763.

  • Rasmusson, E. M., 1966: Atmospheric water vapor transport and the hydrology of North America. Planetary Circulations Project, Massachusetts Institute of Technology Rep. A-1, 170 pp.

  • ——, 1967: Atmospheric water vapor transport and the water balance of North America: Part I. Characteristics of the water vapor flux field. Mon. Wea. Rev.,95, 403–426.

  • Roads, J. O., S. Chen, A. K. Guetter, and K. P. Georgakakos, 1994: Large-scale aspects of the United States hydrologic cycle. Bull. Amer. Meteor. Soc.,75, 1589–1610.

  • Rowson, D. R., and S. J. Colucci, 1992: Synoptic climatology of thermal low-pressure systems over south-western North America. J. Climatol.,12, 529–545.

  • Schmitz, J. T., and S. Mullen, 1996: Water vapor transport associated with the summertime North American monsoon as depicted by ECMWF analyses. J. Climate,9, 1621–1634.

  • Schubert, S. D., J. Pfaendtner, and R. Rood, 1993: An assimilated dataset for earth science applications. Bull. Amer. Meteor. Soc.,74, 2331–2342.

  • Sellers, W. D., and R. H. Hill, 1974: Arizona Climate, 1931–1972. The University of Arizona Press, 616 pp.

  • Starr, V. P., J. P. Peixoto, and H. R. Crisi, 1965: Hemispheric water balance for the IGY. Tellus,17, 463–472.

  • Stensrud, D. J., R. L. Gall, S. L. Mullen, and K. W. Howard, 1995: Model climatology of the Mexican monsoon. J. Climate,8, 1775–1794.

  • Tang, M., and E. R. Reiter, 1984: Plateau monsoons of the Northern Hemisphere: A comparison between North America and Tibet. Mon. Wea. Rev.,112, 617–637.

  • Trenberth, K. E., and J. G. Olson, 1988: An evaluation and intercomparison of global analyses from the National Meteorological Center and the European Centre for Medium Range Weather Forecasts. Bull. Amer. Meteor. Soc.,69, 1047–1057.

  • Wallace, J. M., 1975: Diurnal variations in precipitation and thunderstorm frequency over the conterminous United States. Mon. Wea. Rev.,103, 406–419.

  • Whittaker, L. M., and L. H. Horn, 1981: Geographical and seasonal distribution of North American cyclogenesis, 1958–1977. Mon. Wea. Rev.,109, 2312–2322.

  • WMO, 1975: Climatic Atlas of North and Central America. Vol. I, Maps of Mean Temperature and Precipitation, World Meteorological Organization.

  • Xie, P., and P. A. Arkin, 1996: Analyses of global monthly precipitation using gauge observations, satellite estimates, and numerical model predictions. J. Climate,9, 840–858.

Fig. 1.
Fig. 1.

Topography of North America (excluding northern Canada). Data courtesy of the United States National Geophysical Data Center. The resolution of the data is 5 min (0.083°). The data are available on the World Wide Web at http://www.ngdc.noaa.gov/mgg/mggd.html. The elevation (in meters) is indicated by the bar at right. The box over Arizona and New Mexico indicates the region 112.5°–107.5°W, 32°–36°N used to define the precipitation index (see section 3).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 2.
Fig. 2.

Mean (July–September 1979–95) 925-hPa vector wind, 200-hPa streamlines, and merged satellite estimates and station observations of precipitation (shading). Circulation data are taken from the NCEP–NCAR reanalysis archive. The position of the North American monsoon anticyclone is indicated by “A.” The Bermuda and North Pacific subtropical high pressure centers are indicated by “H.” Precipitation amounts are in millimeters. The characteristic vector length is 10 m s−1. The approximate location of the Great Plains low-level jet is indicated by the heavy solid arrow.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 3.
Fig. 3.

Histograms of the mean (1963–94) daily and 5-day running mean precipitation (units: mm) during JJA at each grid point in the box 115°–105°W, 30°–38°N from the precipitation database of Higgins et al. (1996a).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 4.
Fig. 4.

Mean (1963–94) pentad (bars) and standard deviation (heavy black line) of observed precipitation (Higgins et al. 1996a) over (a) Arizona (112.5°–110°W, 32°–36°N) and (b) New Mexico (107.5°–105°W, 32°–36°N). (Units: mm.)

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 5.
Fig. 5.

Evolution of the composite mean (1963–94) daily precipitation index (units: mm) over Arizona and New Mexico. The average date of monsoon onset is 7 July (defined as day 0 in the composite analysis).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 6.
Fig. 6.

Mean monthly precipitation (units: mm month−1) for July, August, and September from (a) the Xie and Arkin (1996) merged analysis for 1987–95 and (b) the Higgins et al. (1996a) analysis for 1963–94. Contribution of the precipitation during July, August, and September to the annual total, expressed in percent, from (c) the Xie and Arkin (1996) merged precipitation analysis and (d) the Higgins et al. (1996a) precipitation analysis. The contributions in (c) and (d) are for the same years used in (a) and (b), respectively. In (a) and (b), the contour interval is geometric and values greater than 80 mm are shaded. In (c) and (d), the contour interval is 5% and values greater than 40% are shaded.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 7.
Fig. 7.

Analysis of the contribution of the Xie and Arkin (1996) mean (1987–95) monthly precipitation to the annual mean (expressed in percent) over the conterminous United States and Mexico in (a) June, (b) July, (c) August, and (d) September. In each case, the contour interval is 5% and areas with values exceeding 15% are shaded.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 8.
Fig. 8.

Analysis of mean July minus June precipitation (units: mm month-1) from (a) the Xie and Arkin (1996) merged analysis for 1987–95 and (b) the Higgins et al. (1996a) analysis for 1963–94. The letters a–g in (b) indicate the locations of grid points for the histograms in Fig. 9. In each case, the contour intervals are . . . , −160, −80, −40, −30, −20, −10, 10, 20, 30, 40, 80, 160, . . . , and positive (negative) values are shaded dark (light).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 9.
Fig. 9.

Histograms of mean (1963–94) monthly precipitation (units: mm day−1) from selected grid points over the conterminous United States (indicated in Fig. 8b by the letters a–g). (a) Over Arizona (34°N, 110°W), (b) over Texas (30°N, 97.5°W), (c) over Oklahoma (36°N, 95°W), (d) over Montana (46°N, 110°W), (e) over North Carolina (36°N, 77.5°W), (f) over the Florida panhandle (30°N, 85°W), and (g) over Kentucky (38°N, 85°W).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 10.
Fig. 10.

NCEP time mean (JJA 1979–95) for (a) surface wind (units: m s−1) and specific humidity (units: g kg−1), (b) 500-hPa wind (units: m s−1) and specific humidity (units: g kg−1), (c) 500-hPa vertical velocity (units: μb s−1), and (d) precipitable water (units: mm). In (a) and (b), the standard vector length is 10 m s−1. The contour intervals are (a) 2 g kg−1, (b) 0.5 g kg−1, (c) 0.1 μb s−1, and (d) 5 mm. The shading denotes (a) specific humidity greater than 14 g kg−1, (b) specific humidity greater than 2 g kg−1, (c) upward motion, and (d) precipitable water greater than 25 mm.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 11.
Fig. 11.

NCEP time mean (JJA 1979–95) meridional moisture flux (shaded) and vector moisture flux (units: m s−1 g kg−1) for (a) the bottom 7 sigma levels (approximately 1000–850 hPa), (b) the next 5 sigma levels (approximately 850–550 hPa), (c) the next 6 sigma levels (approximately 550–200 hPa), and (d) the full vertical integral (sigma levels 1–28). For comparison with Rasmusson (1967), the standard vector length in (d) is approximately 15 × 102 g (cm s)−1. Values of meridional flux greater than 1 m s−1 g kg−1 (less than −1 m s−1 g kg−1) are shaded dark (light).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 12.
Fig. 12.

Map sequence of observed precipitation represented as the composite mean (1963–94) for 7-day periods centered at (a) day −10, (b) day −3, (c) day +4, and (d) day +11 relative to monsoon onset. The contour interval is 0.5 mm day−1. Values greater than 1.5 mm day−1 are shaded.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 13.
Fig. 13.

(a) Longitude–time diagram of the composite mean (1963–94) observed precipitation anomalies (departures from the JJA 1963–94 time mean) averaged between 34° and 38°N. Results are shown for a 3-day running mean. (b) Map of observed precipitation represented as the composite mean (1963–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1). In (a) and (b), the contour interval is 0.25 mm day−1, the zero contour is omitted for clarity, and values greater than 0.25 mm day−1 (less than −0.25 mm day−1) are shaded dark (light).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 14.
Fig. 14.

Maps of NCEP (a) 200-hPa wind (units: m s−1) and divergence (units: 10−6 s−1), and (b) 500-hPa vertical velocity (units: μb s−1) represented as the composite mean (1979–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1). In (a), the contour interval is 0.5 × 10−6 s−1, and in (b), the contour interval is 0.05 μb s−1.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 15.
Fig. 15.

Map sequence of the NCEP 500-hPa height (units: m) and 500-hPa specific humidity (units: g kg−1) represented as the composite mean (1979–94) for 3-day periods centered at (a) day −12, (b) day −9, (c) day −6, (d) day −3, (e) day 0, (f) day +3, (g) day +6, (h) day +9, and (i) day +12 relative to monsoon onset. The contour interval is 20 m, except for heights above 5880 m where it is 10 m. The light (dark) shading denotes values of specific humidity greater than 2 g kg−1 (greater than 2.5 g kg−1).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 16.
Fig. 16.

Maps of (a) surface specific humidity and (b) 500-hPa specific humidity represented as the composite mean (1979–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1) and normalized by the summer time mean (JJA 1979–95) value (units: percent). The contour interval is 5%.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 17.
Fig. 17.

Map sequence of the NCEP precipitable water (units: mm) represented as the composite mean (1979–94) for 7-day periods centered at (a) day −10, (b) day −3, (c) day +4, and (d) day +11 relative to monsoon onset. The contour interval is 3 mm, and values greater than 24 mm are shaded.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 18.
Fig. 18.

Analysis of the NCEP July minus June (a) percent frequency of criterion 1 jets (Bonner 1968) during the nighttime (0600 UTC and 1200 UTC combined) and (b) vertically integrated, nighttime moisture flux for criterion 1 jets. In each case, the analysis is for 1985–89. The units are (a) in percent and (b) m s−1 g kg−1. The standard vector length in (b) is 2 m s−1 g kg−1.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 19.
Fig. 19.

Maps of NCEP (a) precipitation, (b) evaporation, (c) evaporation minus precipitation, and (d) moisture flux divergence represented as the composite mean (1979–94) difference between the 45-day period after onset (day 0 to day +44) and the 45-day period before onset (day −45 to day −1). In each case, the units are mm day−1 and the contour intervals are . . . , −4, −3, −2, −1, −0.5, −0.25, 0.25, 0.5, 1, 2, 3, 4, . . . . In (a) and (b), values greater than 0.25 mm day−1 (less than −0.25 mm day−1) are shaded dark (light); in (c) and (d), the shading is reversed.

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 20.
Fig. 20.

Map sequence of the NCEP meridional moisture flux (shaded) and vector moisture flux (units: m s−1 g kg−1) for sigma levels 22–28 (approximately 1000–850 hPa) represented as the composite mean (1979–94) for 7-day periods centered at (a) day −10, (b) day −3 (c) day +4, and (d) day +11 relative to monsoon onset. The standard vector length is 15 m s−1 g kg−1. Values of meridional flux greater than 1 m s−1 g kg−1 (less than −1 m s−1 g kg−1) are shaded dark (light).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 21.
Fig. 21.

Same as Fig. 20 except for sigma levels 17–21 (approximately 850–550 hPa). The standard vector length is 10 m s−1 g kg−1. Values of meridional flux greater than 1 m s−1 g kg−1 (less than −1 m s−1 g kg−1) are shaded dark (light).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Fig. 22.
Fig. 22.

Same as Fig. 20 except for sigma levels 11–16 (approximately 550–200 hPa). The standard vector length is 3 m s−1 g kg−1. Values of meridional flux greater than 0.5 m s−1 g kg−1 (less than −0.5 m s−1 g kg−1) are shaded dark (light).

Citation: Journal of Climate 10, 10; 10.1175/1520-0442(1997)010<2600:IOTNAM>2.0.CO;2

Table 1.

Date of onset of monsoon precipitation for each year, based on the daily monsoon precipitation index for the southwestern United States defined in section 3.

Table 1.
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