JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 111, D10106, doi:10.1029/2005JD006497, 2006
Variability in the altitude of fast upper tropospheric winds over the Northern Hemisphere during winter Courtenay Strong1,2 and Robert E. Davis1 Received 14 July 2005; revised 3 November 2005; accepted 13 February 2006; published 31 May 2006.
[1] The surface of maximum wind (SMW) is used as a frame for examining spatial and temporal variability in the vertical position of fast upper tropospheric winds over the Northern Hemisphere during winter. At a given observation time in a gridded data set, the SMW is defined as the surface passing through the fastest analyzed wind above each grid node, with a vertical search domain restricted to the upper troposphere and any tropospheric jet streams extending into the lower stratosphere. Documenting how SMW pressure varies spatially and temporally guides the operational and climatological analysis of tropospheric jet streams and enables more informed interpretation of isobaric wind speed signals (i.e., isobaric speed variability can be caused by both jet core pressure changes and jet core speed changes). Using six-hourly Northern Hemisphere (NH) NCEP- NCAR reanalysis data from 1958 to 2004, the mean three-dimensional structure of the NH SMW is mapped for winter, and vertical cross sections are used to show the position of the SMW relative to the mean isotachs, the mean lapse-rate tropopause, and the mean dynamic tropopause. The Arctic Oscillation (AO) and El Nin˜o/Southern Oscillation (ENSO) are identified as the leading contributors to SMW pressure variability over the NH using principal components analysis. During the AO positive phase, mean SMW pressures are decreased at high latitudes and increased at middle and subtropical latitudes. The SMW response to the AO is elliptical with the centers of pressure change displaced more equatorward over Asia and the Atlantic. During the warm phase of ENSO, eastward expansion and strengthening of the Pacific jet stream is associated with SMW pressures up to 27 hPa lower near 30 N. The relatively low SMW pressures of the east Pacific SMW are flanked, during the warm phase of ENSO, by slower upper tropospheric winds and higher SMW pressures near the equator and the Gulf of Alaska. Citation: Strong, C., and R. E. Davis (2006), Variability in the altitude of fast upper tropospheric winds over the Northern Hemisphere during winter, J. Geophys. Res., 111, D10106, doi:10.1029/2005JD006497.
1. Introduction and the development of tornados [Fawbush et al., 1951; 1.1. Perspective Skaggs, 1967]. Jet streams also produce gravity waves associated with clear air turbulence [Knox, 1997] and polar [2] Much effort has been aimed at documenting and understanding variability in the horizontal position of con- stratospheric clouds [Hitchman et al., 2003; Buss et al., centrated bands of fast upper tropospheric winds known as 2004]. jet streams. Knowing the jet stream’s location has proven [3] In addition to having immediate practical relevance, valuable to aviation [Reiter, 1958] and short- and long-term jet stream position variability can be an important indicator forecasting [Riehl et al., 1954]. Jet stream variability also of climate change. Extratropical jet streams are closely tied influences stratosphere-troposphere exchange [Hoerling et to the polar front [Palmen, 1948; Palmen and Newton, al., 1993; Langford, 1999; Wimmers et al., 2003], the 1948] and strengthen and expand equatorward during win- intensity and track of synoptic-scale storms [Nakamura, ter [Namias and Clapp, 1949]. On interannual or longer 1992; Christoph et al., 1997], the development of severe time scales, changes in the thermal structure of the atmo- convective storms [Uccellini and Johnson, 1979; Kloth and sphere can exert an influence on the shape and intensity of Davies-Jones, 1980], the occurrence of heavy rain [Smith fast tropospheric flows. The linkage between temperature and Younkin, 1972] and snow [Uccellini and Kocin, 1987], and the geometry and location of the fast mid- to upper- tropospheric circulation has been demonstrated through study of the circumpolar vortex [Angell and Korshover, 1Department of Environmental Sciences, University of Virginia, 1977; Davis and Benkovic, 1992, 1994; Burnett, 1993]. Charlottesville, Virginia, USA. 2 Vortex variations at 700, 500 and 300 hPa collectively Now at Department of Soil, Environmental, and Atmospheric account for almost two-thirds of midlatitude tropospheric Sciences, University of Missouri-Columbia, Columbia, Missouri, USA. temperature variability [Frauenfeld and Davis, 2003]. Copyright 2006 by the American Geophysical Union. [4] Because of the jet stream’s close relationship to 0148-0227/06/2005JD006497$09.00 tropospheric and stratospheric thermal fields, it has been
D10106 1of15 D10106 STRONG AND DAVIS: FAST UPPER TROPOSPHERIC WIND ALTITUDE D10106 observed to vary in conjunction with climate phenomena signals. Climate studies considering upper tropospheric that are linked to temperature variability. The Arctic Oscil- wind speed variability often utilize an isobaric surface lation (AO), for example, is the meridional seesaw of presumed to lie close to the jet stream level. Some inaccu- atmospheric mass between the polar region and midlatitudes racy may result when data obtained from a single isobaric [Lorenz, 1951; Kutzbach, 1970] manifested as the leading surface are used to evaluate climate trends in jet streams or principal component of wintertime sea level pressure fast upper tropospheric winds in general [Strong and Davis, [Thompson and Wallace, 1998]. For zonally averaged data 2005]. At any one time, no isobaric surface can capture a jet over the Northern Hemisphere (NH), the positive phase of stream whose altitude varies spatially in the hemisphere, the AO is associated with a column of tropospheric warm- and, at any one location, no isobaric surface can capture a ing near 45 N, cooling in the upper troposphere and jet stream whose vertical position varies temporally. stratosphere over the polar region, and slower (faster) zonal [8] In prior research, we identified statistically significant wind speeds near 30 N (60 N) from the surface into the temporal trends in wind maxima pressure (as large as stratosphere [Thompson and Wallace, 2000]. 30 hPa decade 1) over the Northern Hemisphere tropics [5] Planetary circulation is also influenced by the El for summers 1958–2004, indicating decadal jet stream Nin˜o/Southern Oscillation (ENSO) – the coupled variation descent consistent with observed changes in the temperature of sea surface temperature and sea level pressure over the structure of the troposphere and lower stratosphere [Strong tropical Pacific [Bjerknes, 1969]. The ENSO warm phase and Davis, 2006]. Here, we document spatial and temporal corresponds to anomalously low pressure and high surface variability in the vertical location of wind maxima close to water temperatures over the eastern Pacific. Associated the tropopause during winter, including those associated tropical phenomena include weakening or reversal of the with the tropopause jet streams. In the Introduction easterly trade winds [Ichiye and Petersen, 1963], increases section 1.2, we give theoretical consideration to the rela- in tropical tropospheric temperature [Horel and Wallace, tionship between horizontal temperature gradients and the 1981; Yulaeva and Wallace, 1994; Sobel et al., 2002], water vertical location of wind speed maxima. Section 2 describes vapor mass, and precipitation intensity [Soden, 2000]. An the data and the surface of maximum wind (SMW) [Strong ENSO-related intensification and equatorward displacement and Davis, 2005] used to track the pressure of upper tropo- of the mid- to upper-tropospheric circulation over the NH spheric wind maxima. Spatial variability of the SMW is eastern Pacific has been detected in upper tropospheric wind considered in section 3.1 and principal contributors to joint speed [Arkin, 1982; Rasmusson and Mo, 1993; Yang et al., space time variability of the extratropical and tropical SMW 2002], middle and upper tropospheric geopotential height are considered in section 3.2. topography [Horel and Wallace, 1981; Hastenrath, 2003], storm tracks [Chen and Van den Dool, 1999], second order 1.2. Theory statistics indicative of storm track regimes [Straus and [9] The geostrophic wind speed’s rate of change with Shukla, 1997], and the circumpolar vortex [Angell and respect to pressure can be expressed as (details in Korshover, 1985; Frauenfeld and Davis, 2000; Angell, Appendix A) 2001]. The influence of ENSO on extratropical circulation @V R R @T is also manifested as variations in the equatorial zonal g ¼ d r~T cos b ¼ d v ð1Þ Walker cell, the tropical meridional Hadley cell, the extra- @ ln p f v f @n tropical meridional Ferrel cell, and the midlatitude zonal cell [Wang, 2002]. where Rd is the dry air gas constant, f is the Coriolis [6] The vertical position of fast upper tropospheric wind parameter, Tv is virtual temperature, the del operator is features has also received some attention. Initial interest in applied on an isobaric surface, b is the angle between the jet stream altitude produced cross sections of zonally geopotential gradient ( r~F) and the virtual temperature ~ averaged summer and winter wind speed data from all gradient ( rTv), and n is oriented normal to V~g and defined latitudes [Mintz, 1954], physical explanation of the jet positive to the right of the flow direction. Equation (1) stream’s location [Staff Members, 1947], and detailed map- indicates that wind speed will increase with height (@Vg/@ ln ping of the layer of maximum wind’s altitude relative to jet p will be negative) when higher temperatures on an isobaric stream axes [Reiter, 1958]. The U.S. Navy began opera- surface are collocated with the higher values of geopotential tionally mapping the altitude and vertically averaged speed (i.e., b < p/2). The temperature gradient in the b < p/2 case of the three-dimensional layer of maximum wind (LMW) is associated with stronger gradients of geopotential in the beginning at the end of 1958 for aviation purposes [Reiter, overlying layer in accordance with the hypsometric 1961]. relationship, supporting an increase in the overlying [7] Little is known, however, about climatological vari- geostrophic wind speed. In the b p/2 case, lower ability in the pressure of jet streams and other upper temperatures are collocated with higher values of geopo- tropospheric wind maxima. Significant interannual variabil- tential, indicating weaker overlying geopotential gradients ity in the pressure of tropospheric jet streams is an intrin- and geostrophic wind speeds. ~ sically interesting and important phenomenon in the [10] Since krTvkcos b and @Tv/@n are equivalent, the atmosphere. From a practical perspective, mapping how more compact notation @Tv/@n will be used in subsequent the pressures of upper tropospheric wind maxima vary sections of this manuscript. In this section, however, the spatially guides operational and climatological analyses of behavior of equation (1) in the atmosphere will be consid- ~ tropospheric jet streams, and documenting how the pres- ered using the krTvkcos b form to stress that equation (1) sures of upper tropospheric wind maxima vary temporally may be forced to zero (marking the local wind maxima) by enables more informed interpretation of isobaric wind speed either the angle between the geopotential and temperature
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Figure 1. (a) Meridional cross section showing virtual temperature isopleths (labels in K) and isotachs (grayscale shading every 7 m/s up to 56 m/s) along 47.5 E on 1 January 2000. Bold lines show the position of the polar front, the subtropical front, and the Reanalysis tropopause. The dashed vertical line shows the position of the profiles in Figures 1b–1d (47.5 N, 47.5 E). (b) The magnitude of the isobaric ~ ~ ~ virtual temperature ascendent (krTvk), and the angle b between rTv and rF. (c) From equation (1), the change of the geostrophic wind speed with respect to ln p so that a positive value indicates speed increasing with height. (d) Reanalysis wind speeds. The shaded box across Figures 1b–1d shows the layer where the wind maximum likely occurred. gradients or the magnitude of the temperature gradient in the upper, poleward branch of the Hadley circulation itself. As an example, Figure 1a shows a cross section of [Krishnamurti, 1961a, 1961b]. virtual temperature isopleths and isotachs along 47.5 Eon [11] Figures 1b–1d show profiles of the terms from 1 January 2000 from the NCEP-NCAR Reanalysis [Kalnay equation (1) evaluated on the dashed vertical line passing et al., 1996]. The virtual temperature gradients associated through the PFJ core in Figure 1a. The magnitude of the ~ with the polar front integrate vertically to produce the isobaric virtual temperature ascendent (krTvk) was nonzero isobaric geopotential gradients supporting the polar front and fluctuated around 6 Â 10 6 Km 1 below the PFJ core jet stream (PFJ: 250 hPa, 45.7 N). Likewise, the isobaric (Figure 1b; Figure 1a shows the virtual temperature iso- ~ virtual temperature gradients associated with the subtropical pleths associated with the @Tv/@y component of krTvk). front integrate vertically to produce the isobaric geopoten- The geopotential gradients were generally aligned with the tial gradients contributing to the subtropical jet stream (STJ: virtual temperature gradients up to 250 hPa (b < p/2, 200 hPa, 22.5 N). The STJ also receives energy from Figure 1b), indicating an increase in geopotential gradient ~ tropical heating and the conservation of angular momentum magnitude with height (positive krTvkcos b, Figure 1c),
3of15 D10106 STRONG AND DAVIS: FAST UPPER TROPOSPHERIC WIND ALTITUDE D10106 and wind speed increasing with height (Figure 1d). At some use of satellite data after 1978 [Kistler et al., 2001]. pressure between 250 and 200 hPa, the virtual temperature Kanamitsu et al. [1997] found that satellite data impacted gradient was nonzero but orthogonal to the geopotential the Northern Hemisphere troposphere Reanalysis less than gradient (b = p/2, Figure 1b), the first derivative of wind it impacted data sparse areas such as the stratosphere and speed with respect to pressure went to zero (Figure 1c), and eastern oceanic areas of the Southern Hemisphere. The the wind speed reached a local maximum (Figure 1d). potential for perceiving an artifact of changes to the Re- Above 200 hPa, the virtual temperature gradients opposed analysis observing system as climate variability is reduced the geopotential gradients (b > p/2, Figure 1b), and wind here because conclusions are not based solely on temporal speed decreased with height (Figure 1d). variability in the Reanalysis. Rather, results obtained from [12] The ageostrophy prevalent in the vicinity of jet the Reanalysis are contextualized through correlation with streams [Newton, 1959; Shapiro and Kennedy, 1981] vio- existing, well-studied climate phenomena and indices. lates the assumption of geostrophic balance, so equation (1) [15] Climate indices or time series employed here include will not always provide a reliable estimate of the true the Arctic Oscillation Index (AOI), the Multivariate El change of wind speed with pressure. Nonetheless, the Nin˜o/Southern Oscillation Index (MEI), and the mean NH geostrophic thermal wind relationship indicates how tem- surface temperature (T NH). AOI values were obtained from perature influences the vertical distribution of geopotential the Climate Prediction Center, and were developed by gradients and is thus a useful theoretical model for inves- projecting the monthly mean 1000 hPa geopotential height tigating jet core pressure variability. The pressure variability anomalies onto the first principal component from the of wind maxima is determined in this research without principal component analysis (PCA) of the monthly mean assuming geostrophic balance (i.e., the wind field will be 1000 hPa geopotential height field north of 20 N. MEI used rather than the fields of geopotential and temperature). values were obtained from the Climate Diagnostic Center, Equation (1) will then be used to link the observed wind and capture variability in sea-level pressure, zonal and maxima pressure variability to atmospheric temperature meridional components of the surface wind, sea surface variability. A change in jet core pressure represents a change temperature, surface air temperature, and total cloudiness in the position of the local wind speed maximum on the fraction of the sky over the tropical Pacific [Wolter and pressure axis and is therefore associated with nearby Timlin, 1993, 1998]. Monthly mean T NH data were obtained changes in the first derivative @Vg/@ ln p. from the NASA Goddard Institute for Space Studies, and were calculated by combining meteorological station mea- 2. Data and Methods surements over land with sea surface temperatures obtained primarily from satellite measurements [Reynolds and Smith, 2.1. Data 1994; Smith et al., 1996]. We averaged the December– [13] We use wind speed on seven isobaric surfaces (500, February monthly values for each predictor time series for 400, 300, 250, 200, 150, and 100 hPa) resolved every six winters 1958–2004 to match the temporal structure of the hours on a 2.5 grid from the NCEP-NCAR Reanalysis, and surface of maximum wind winter seasonal means. Reanalysis lapse-rate tropopause pressure data at the same temporal and horizontal spatial resolution. In addition to 2.2. Surface of Maximum Wind using the Reanalysis lapse-rate tropopause, we also identify [16] We illustrate the algorithm for locating the upper a ‘‘dynamic tropopause’’ based on the rapid increase in tropospheric SMW using data from 1200 UTC on 9 January potential vorticity between the troposphere and stratosphere 2004 along 175 E (Figure 2). Data in the upper troposphere [Danielson, 1968; Danielson and Hipskind, 1980; Hoerling and lower stratosphere along the meridian are first evaluated et al.,1991;Morgan and Nielsen-Gammon, 1998]. To for SMW candidacy, and then the fastest wind speed from locate the dynamic tropopause, we use Reanalysis data the candidates above each node is selected as a point on the and an expression for Ertel’s potential vorticity in isobaric SMW. From the column of data above a node (above a tick coordinates [Bluestein, 1992] mark on the abscissa in Figure 2), each wind speed from 500 to 100 hPa is a SMW candidate unless the datum is Rd @v @T @u @T @q above the measurement level closest to the lapse-rate Pq ¼ g z þ f þ ; ð2Þ sp @p @x @p @y p @p tropopause and outside of a 25.7 m/s (50-knot) contour that captures at least one tropospheric datum. In other words, the where relative vorticity z = @v/@x @u/@y, q is potential search domain for the SMW is restricted to the upper temperature, and s is a static-stability parameter troposphere and any tropospheric jet streams that extend into the lower stratosphere. RT @ ln q [17] For the example in Figure 2, points that are not SMW s ¼ : ð3Þ p @p candidates are left blank, points that are SMW candidates are shown with open circles, and the SMW is shown with Various values of Pq have been suggested for locating the solid circles. No SMW candidates appear outside 500 to dynamic tropopause ranging from 1 to 3.5 PVU [Hoerling et 100 hPa, and all points from 500 hPa to the lapse-rate al., 1991; Spaete et al., 1994], where 1 PVU is defined as tropopause are SMW candidates. The point at 40 N and 10 6 m2 s 1 Kkg 1. In section 3, we calculate and show the 200 hPa is a SMW candidate because it is bounded by a winter mean pressure of the 1, 2, 3, and 3.5 PVU isopleths. 25.7 m/s wind speed contour that captures at least one [14] Artificial trends or discontinuities may exist in the tropospheric datum. The point at 87.5 N and 100 hPa is not a Reanalysis because of variations in the observing system SMW candidate because its 25.7 m/s contour does not including changes in the density of rawinsonde data and the capture at least one tropospheric datum. Properties associated
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the extratropics (poleward of 20 N). For each PCA, the variables are Reanalysis grid nodes and the cases are the seasonal mean SMW pressures at each node over 47 years. Since jet stream level thicknesses may be more than 10% larger in the tropics than the polar region, the correlation matrix is used as input into the PCA (rather than the variance-covariance matrix). To ensure proper representa- tion in the correlation matrix, data are standardized and weighted by the square root of the cosine of latitude. Singular value decomposition is used to obtain the eigen- values and unit-length eigenvectors of the area-weighted correlation matrix. The time series produced by projecting the weighted, standardized data onto the mth eigenvector will be referred to as ‘‘principal component m’’ (PC m), and the contoured map of the mth eigenvector’s elements will be referred to as the PC m ‘‘loading pattern.’’ For display and comparison with existing climate indices, each principal component time series is standardized by subtracting its Figure 2. Wind speed (contours), lapse-rate tropopause mean and dividing by its standard deviation. The variance pressure, the maximum wind level (MWL), and surface of accounted for by PC m is eigenvalue m divided by the maximum wind (SMW) along 175 E at 1200 UT on summation of all the correlation matrix eigenvalues. 9 January 2004. Open circles show data that are candidates [21] Correlations are measured with least squares linear for being SMW pressures, and solid circles show the SMW. regression and reported p statistics are based on degrees of The gray areas show wind speeds greater than or equal to freedom (df) adjusted to account for Pearson lag-1 autocor- 25.7 m/s (50 knots). relation in the regression residuals. Where the results of correlation tests are mapped (Figures 5a and 8a), the fraction of the map area with significant results is verified with the SMW will be denoted by a tilde so that SMW to be field significant at a = 0.05 using Monte Carlo pressure is Pe and SMW wind speed is Ve. simulations [Livezey and Chen, 1983] confirming there is [18] A dynamic tropopause based on isentropic potential less than a five percent probability that the amount of area vorticity [Hoerling et al., 1991, and references therein] with significant local results occurred by chance given the could alternatively be used in applying the SMW method. field’s spatial degrees of freedom. The Reanalysis lapse-rate tropopause is used here because it is also viable in the tropics. The algorithm presented in 3. Results this section could likewise be adapted to locate the surface of maximum geostrophic wind using the fields of temper- [22] Section 3.1 is focused on the spatial variability of the ature and geopotential to interpolate the pressure where winter SMW. In section 3.2, PCA is used to uncover the AO equation (1) reaches zero near the analyzed geostrophic and ENSO as principal contributors to SMW joint space- wind maximum. time variability. The relationship between the SMW and AO [19] It is the SMW’s flexible upper bound that distin- is explored in section 3.3, and the relationship between the guishes it from the operationally defined layer of maximum SMW and ENSO is explored in section 3.4. wind (LMW) [Reiter, 1958] and the maximum wind layer 3.1. SMW Spatial Variability (MWL) available in the Reanalysis. The MWL has a fixed e [23] Figures 3a and 3b show the 47-year mean winter P vertical domain between 500 and 70 hPa, and is shown for e comparison in Figure 2. In this example, differences be- and V for all locations over the NH. SMW pressures tween the SMW and MWL in the tropics and through the jet generally decrease from the pole toward the position of core near 30 N result from the MWL’s use of cubic splines the polar front jet (PFJ) in the midlatitudes. The winter for interpolation. Differences in higher latitudes are more SMW maintains a broad shelf of high pressure in the Arctic substantial and result from the SMW algorithm’s exclusion because of the expanded polar high, and is steeply sloped of stratospheric wind data outside a tropospheric jet stream near the fast jet stream cores in the midlatitude western (25.7 m/s contour). In particular, the winter polar night jet Pacific, off the east coast of the United States and east of the Mediterranean Sea (Figure 3a). Just equatorward of the stream is excluded by the SMW algorithm. In the summer, e the MWL tends to capture stratospheric jets excluded by the steep SMW pressure decreases, high values of V associated SMW over the tropical lapse-rate tropopause and the SMW with the presence of jet cores are found in the west Pacific, and MWL have better agreement in mid- to high-latitudes eastern United States, and to the south of the Mediterranean (not shown). Sea (Figure 3b). The depression in the SMW topography (area of relatively high pressure) in the equatorial western 2.3. Statistical Methods Pacific (Figure 3a) is associated with a corridor of relatively e [20] Principal components analysis (PCA) is used to low V (Figure 3b) and will be discussed in the context of identify uncorrelated patterns of simultaneous anomalies ENSO in section 3.4. that account for variability in winter mean Pe. A separate [24] The cross section in Figure 3c shows how the PCA is performed in the tropics (equatorward of 20 N) and 47-year mean SMW relates to the 47-year mean lapse-rate
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Figure 3. (a) Mean surface of maximum wind pressure (Pe) and (b) mean surface maximum wind speed (Ve) for winters 1958–2004. The lines in Figures 3a and 3b mark the meridian used for the cross sections in Figures 3c and 3d. (c) Cross section showing Pe (bold dashed line) with mean winter tropopause pressure (bold solid line) and isotachs (dotted lines at 10 m/s interval) along 135 E for winters 1958– 2004. Dashed lines bound the 10th to 90th percentile of Pe and the gray area bounds the 10th to 90th percentile of seasonal mean tropopause pressure. (d) Same as Figure 3c, except the tropopause is replaced with the dynamic tropopause as shown by isopleths of 1, 2, 3, and 3.5 potential vorticity units (PVU) with shading between 1 and 3.5 PVU. tropopause and isotachs using data from 135 Easan near 35 N, and resides generally below it over high example. The seasonal mean SMW at 135 E is below the latitudes. tropopause at the equator, undulates in the tropical latitudes, [26] The tendency for a nearly vertical tropopause near jet and extends laterally poleward through the lapse-rate tro- streams is well known [e.g., Platzman, 1949; Defant and popause in the vicinity of the western Pacific jet’s 70 m/s Taba, 1957; Morgan and Nielsen-Gammon, 1998], and the mean isotach (200 hPa, 35–45 N, Figure 3c). The range of dynamic tropopause and lapse-rate tropopause are both winter Pe at 135 E is minimized close to the jet core at steeply sloped through the jet core near 35 N. The SMW, 32.5 N, reflecting the stability of this feature’s seasonal by contrast, is most steeply sloped poleward of the fastest mean pressure. wind speeds (Figure 3c, also seen in the contour packing in [25] The SMW also has a structural relationship to the Figure 3a just inside the spiral-shaped corridor of fast winds dynamic tropopause. Figure 3d shows various isosurfaces in Figure 3b). used to represent the dynamic tropopause ranging from 1 to 3.5 PVU with gray shading over intervening pressures. 3.2. Contributors to SMW Temporal Variability [27] In the extratropics, the first principal component of The SMW’s relationship to the Reanalysis lapse rate e tropopause is similar to the SMW’s relationship to the winter mean SMW pressure (extratropical P PC 1, Figure 4a) accounts for 28% of the extratropical variance. 3.5-PVU dynamic tropopause recommended for the extra- e tropics by Hoerling et al. [1991] (compare Figures 3c and The P PC 1 loading pattern has oppositely-signed rings in 3d). Specifically, the SMW resides below the 3.5-PVU the polar and midlatitudes, suggesting an AO-related phe- dynamic tropopause in the tropics, extends through it into nomenon, and produces scores correlated with the AOI the lower stratosphere at the 70-m/s wind speed maximum (Figure 4b). The AO has its strongest positive and negative
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Figure 4. (a) Loading pattern of the first principal component (PC 1) of winter mean surface of maximum wind pressure (Pe) in the extratropics (unshaded area poleward of 20 N). Contour interval is 0.01 with negative values dashed. The +AO and –AO labels relate to the loading pattern of the Arctic Oscillation and are referenced in the text. (b) Time series of the extratropical Pe PC 1 score and the Arctic Oscillation Index. (c) Loading pattern of tropical (unshaded area from 0–20 N) Pe PC 1 contoured at 0.01 with negative values dashed. (d) Time series of the tropical Pe PC 1 score and the multivariate El Nin˜o/ Southern Oscillation Index (MEI). The effective degrees of freedom are denoted ‘‘df’’ and the mean square error is denoted ‘‘MSE.’’ loadings (Figure 4a: labeled +AO and AO, respectively) df = 43, mean square error [MSE] = 0.36) followed by an where the extratropical Pe PC 1 loadings are relatively weak. ENSO-like phenomenon (NH Pe PC 2 [9% of the hemi- The AO and Pe loading centers of action have a latitudinal spheric variance], score versus MEI: p < 0.01, r2 = 0.29, df = offset because changes in the AO reflect temperature 29, MSE = 0.71). The leading principal component’s variability (related to pressure variability) whereas changes correlation with the AOI does not markedly depend on in Pe reflect temperature gradient variability. In the tropics, the inclusion or exclusion of the tropics during the PCA. the first principal component of winter mean SMW pressure However, the leading principal component of the tropical- (tropical Pe PC 1, Figure 4c) accounts for 14% of the tropical only PCA is better correlated with the MEI than the second variance. The loading pattern suggests an ENSO-like phe- principal component of the hemispheric analysis, perhaps nomenon with a pair of oppositely signed loading centers because manifestations of ENSO outside the tropics can where convection shifts from Indonesia toward the central vary among events [Rasmusson and Wallace, 1983]. The Pacific during El Nin˜o, and the loading pattern produces third principal component of the hemispheric PCA is scores correlated with the MEI (Figure 4d). significantly correlated with mean NH surface temperature e [28] A PCA was also performed using 47 years of winter (NH P PC 3 [8% of the hemispheric variance], score versus 2 mean SMW pressure at each node over the entire NH as T NH: p < 0.01, r = 0.36, df = 34, MSE = 0.64). Attention input (not shown). An AO-like phenomenon was again the here will be restricted to the AO (section 3.3) and ENSO leading contributor to SMW variability (NH Pe PC 1 [20% (section 3.4) as the two leading contributors to SMW space of the NH variance], score versus AOI: p < 0.01, r2 = 0.64, time variability.
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Figure 5. (a) Pearson correlation coefficient (r) for winter (December–February) mean surface of maximum wind pressure (Pe) versus the Arctic Oscillation Index (AOI) for 1958–2004. Contours are at 0.3 with negative values dashed. Results locally significant at a = 0.05 account for 53% of the map area (including, but not limited to, all of the contoured area), and the map of results is field significant at a = 0.05. The hemisphere is divided into sectors (e.g. East Pacific) discussed in the text. (b) For each of the sectors in Figure 5a, the mean winter surface of maximum wind pressure (Pe) for positive Arctic Oscillation Index (AOI) minus Pe for negative AOI. The arrows in Figure 5a indicate where the sample meridians are located, and all results outside of the gray region are individually significant at a = 0.05.
3.3. SMW and the AO organization of Figures 6 and 7 will facilitate their discus- [29] SMW pressure tends to seesaw with the AO, and the sion. Each column shows data from one of the four sectors map of the Pe versus AOI correlation is field significant at in a meridional cross section collocated with an arrow at the a = 0.05 (Figure 5a). During the positive phase of the AO, equator in Figure 5a. The three panels in each column show, SMW pressures are generally decreased over the high from the bottom to the top, the AO-related change in latitudes (negative Pe versus AOI correlation) and increased temperature (DT, bottom panel), the AO-related change in e 1 over the midlatitudes or subtropics (positive P versus AOI the Rdf @Tv=@n term from equation (1) (middle panel), correlation). Figure 5a has been divided into five sectors the AO-related change in wind speed (DV, upper panel), based on the pattern of the SMW versus AO relationship and the mean SMW during the positive phase of the AO and and the nature of the underlying tropospheric temperature during the negative phase of the AO (also upper panel). changes associated with the AO (shown later). The pattern [31] In the East Pacific at 120 W (Figure 6, left column), of Pe versus AOI correlation is elliptic with a longer axis the temperature changes associated with the AO are largely oriented through the Atlantic and Asian sectors (Figure 5a). stratospheric, consisting of a high latitude cooling and The ellipticity of the SMW response to the AO is further midlatitude warming (Figure 6c). The intervening increase illustrated in Figure 5b where we show the SMW pressure in @Tv/@n near the tropopause at 70 N (Figure 6b) raises the altitude where @T /@n goes to zero and is associated under positive AO minus the SMW pressure under negative v e AO for a sample meridian from each of the five sectors. The decreased values of P (Figure 6a). Likewise, decreased Atlantic and Asian zones of AO-related SMW pressure @Tv/@n near the tropopause at 30 N (Figure 6b) is associ- changes (55 Nand20 N, respectively) are displaced ated with increased SMW pressures (Figure 6a). Although equatorward from the corresponding belts of SMW pressure the pressure of the SMW varies with the AO in the East change in the East Pacific and European sectors (Figure 5b). Pacific sector, there is little change in wind speed at the The SMW relationship to the AO in the Central Pacific SMW (Figure 6a) because weak changes in @Tv/@n through sector is relatively weak, with only isolated significant the depth of the underlying troposphere (Figure 6b) verti- differences poleward of 20 N. cally integrate to cause only weak changes in upper tropo- [30] The ellipticity of the SMW response and the presence spheric geopotential gradients. AO-related wind speed or absence of accompanying upper tropospheric wind speed changes do occur over the East Pacific, but are predomi- changes can be understood by examining AO-related nantly stratospheric with faster speeds above the column of changes in the thermal structure of the atmosphere. This positive D@Tv=@n and slower speeds above the column of is illustrated in Figures 6 and 7, beginning with the East negative DTv=@n (Figures 6a and 6b). Pacific sector and moving east through the Asian sector (the [32] The Atlantic sector (Figure 6, right column) also temperature changes in the Central Pacific sector are exhibits high latitude cooling north of warming, but here the weaker, but spatially similar to, those in the Atlantic sector temperature changes appear in the stratosphere and tropo- and will not be shown separately). Some comments on the sphere (Figure 6f). The deep temperature changes over the
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Figure 6. The symbol D indicates mean value for the positive phase of the Arctic Oscillation Index (AOI) minus mean value for the negative phase of the AOI for 1958–2004, and each panel shows the 47-year mean tropopause with a thin gray line. Figures 6a–6c are for 120 W, and Figures 6d–6f are for 60 W. (a, d) DV contoured gray at 1.5 m/s with negative values dashed. The AOI positive phase surface of maximum wind pressure (Pe, solid black line) and AOI negative phase Pe (dashed black line) with 1 circles showing differences significant in a t test at a = 0.05. (b, e) D[ Rdf @Tv=@n] contoured at 1.5 m/s with negative values dashed. (c, f) DT contoured at 0.5 K with negative values dashed.
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Figure 7. Same as Figure 6, but for (a–c) 0 E and (d–f) 110 E.
Atlantic sector support a vertically tilted column of in- @Tv/@n (Figure 6e). Warming of the column at 40 N during creased @Tv/@n from the surface at 45 N into the strato- the positive phase of the AO (Figure 6f) weakens @Tv/@n sphere (Figure 6e) and decreased SMW pressures near in the equatorward column at 30 N (Figure 6e) and 55 N. There are also speed increases on the SMW near decreases the wind speeds at the overlying SMW 55 N (Figure 6d) because of the deep underlying changes in (Figure 6d). Near-zero values of D@Tv=@n close to the
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Figure 8. (a) Pearson correlation coefficient (r) for winter (December–February) mean surface of maximum wind pressure (Pe) versus the Multivariate El Nin˜o/Southern Oscillation Index (MEI) for 1958–2004. Contours are at 0.2 with negative values dashed. Results locally significant at a = 0.05 account for 38% of the map area (including, but not limited to, all of the contoured area), and the map of results is field significant at a = 0.05. (b) Mean winter Pe for MEI 1 (El Nin˜o warm phase) minus mean winter Pe for MEI 1 at 7.5 N and 30 N. Differences significant in a t test at a = 0.05 are shown with circles.
SMW (Figure 6e), however, are associated with little or no contributing to the ellipticity of the SMW versus AO change in the mean SMW pressure (Figure 6d). Side by side relationship. comparison of the @Tv/@n changes in the East Pacific and [35] Farther south in the Asian sector, the positive phase of Atlantic sectors (Figure 6b versus Figure 6e) reveals that the the AO is also associated with higher temperatures near the centers of maximum positive and negative @Tv/@n change steep slope in the tropopause (Figure 7f, 25–45 N). Al- are approximately 15 closer to the equator in the Atlantic though the associated, equatorward changes in @Tv/@n are sector, consistent with the relative positions of the SMW large (10–35 N, Figure 7e), the SMW pressure differences pressure changes. The equatorward displacement in the are small (10–35 N, Figure 7d) because the pressure of the Atlantic sector contributes to the ellipticity of the SMW maximum change in @Tv/@n is very close to the SMW versus AOI relationship shown in Figure 5. pressure at this latitude (indicating generalized slowing of [33] In the European sector (Figure 7, left column), the wind speed without a relocation of the speed maximum). For positive phase of the AO is associated with tropospheric comparison, consider the pressure offset between the SMW warming poleward of 40 N and tropospheric cooling equa- and the center of @Tv/@n change near 65 N in Figures 7d and torward of 40 N (Figure 7c). The warming poleward of 7e (or any other SMW pressure change shown in the 40 N is associated with a strengthening of @Tv/@n in a manuscript). The comparison illustrates two points: (1) sig- column from 60 N at the surface up into the polar strato- nificant SMW pressure differences occur when there is a sphere (Figure 7b), SMW pressures decrease and wind pressure offset between the maximum @Tv/@n change and the speeds are generally faster (Figure 7a). Between the col- SMW, and (2) SMW pressure changes are always associated umns of tropospheric warming and cooling, @Tv/@n is with nearby changes in @Tv/@n, but changes in @Tv/@n near weakened (40 N, Figure 7b) during the positive phase of the SMW are not always associated with SMW pressure the AO, and the overlying SMW is slower and lower changes. Both points follow from the general relationship (Figure 7a). between the position of a function’s maximum (in this [34] The Asian sector features high latitude stratospheric application, the pressure of the fastest wind speed) and the cooling and a shallow tropospheric warming near 65 N function’s first derivative (in this application, equation (1)). during the positive phase of the AO (Figure 7f). Tropo- spheric virtual temperature gradients are large in this region 3.4. SMW and ENSO during winter and the shallow tropospheric warming [36] SMW pressure is significantly correlated with the MEI amounts to little change in @Tv/@n and little change in wind at nodes amounting to more than one-third of the NH surface area and more than one-half the tropics equatorward of 20 N speed between the positive and negative phases of the AO. e The stratospheric cooling, however, increases @Tv/@n near (Figure 8a), and the map of the P versus MEI correlation is the SMW and raises the altitude at which @Tv/@n goes to field significant at a = 0.05. At 30 N, the warm phase of zero, supporting a broad zone of SMW pressure decrease ENSO is associated with SMW pressure decreases from between 50 and 70 N (Figure 7d). Side by side comparison China across the Pacific toward the United States of the Asian and European sector @Tv/@n changes (Figure 8a) with changes in mean pressure as large as (Figures 7b and 7e) shows that the location of the largest 27 hPa (Figure 8b). At 7.5 N, the warm phase of ENSO is near-tropopause increase in @Tv/@n is closer to the pole over associated with up to 60 hPa of SMW pressure increase near Europe and closer to the equator over the Asian sector, 120 E and 160 W with relatively low SMW pressures over
11 of 15 D10106 STRONG AND DAVIS: FAST UPPER TROPOSPHERIC WIND ALTITUDE D10106 intervening longitudes (Figure 8b). The low-Ve, high-Pe de- during the AO positive phase, mean SMW pressures are pression in the west-Pacific SMW topography introduced in decreased at high latitudes and increased at middle and section 3.1 (Figure 3a) is a stable feature with lower variance subtropical latitudes. The magnitude of the SMW response than the surrounding SMW (not shown), and exhibits a weak to the AO is largest in the European and East Pacific sectors pressure correlation with the MEI (Figure 8a). and smallest in the Central Pacific. The SMW response to [37] The relationship between SMW pressure variability the AO is elliptical with the centers of SMW pressure and ENSO shown in Figure 8 can be understood in the change displaced most equatorward over Asia and the context of atmospheric temperature changes. Beginning Atlantic. The position and strength of the tropospheric and with the east Pacific at 140 W (Figures 9a–9c), the warm stratospheric temperature changes associated with the AO phase of ENSO is associated with tropospheric warming are used to explain why the belts of SMW pressure change centered at 15 N and 250 hPa and tropospheric cooling are arranged elliptically about the pole, and why each is not centered at 40 N and 400 hPa (Figure 9c). The east Pacific consistently associated with speed changes. Over the At- warming center at 15 N is consistent and collocated with lantic and Europe, AO-related temperature changes cause ENSO-related downward vertical velocity anomalies in the deep tropospheric changes in @Tv/@n and variation of the Hadley circulation cell [Wang, 2002]. Likewise, the cooling overlying upper tropospheric winds is correspondingly at 40 N in Figure 9c is consistent with ENSO-related large. Over the east Pacific and Asia, AO-related tempera- upward vertical velocity anomalies identified by Wang ture changes cause largely stratospheric changes in @Tv/@n, [2002]. @Tv/@n is strengthened in the intervening column resulting in SMW pressure changes with comparably small at 25 N (Figure 9b), supporting the relatively low SMW changes in upper tropospheric wind speeds close to the pressures and faster wind speeds in the overlying jet stream SMW. (Figure 9a). Equatorward of the tropospheric warming [41] The dominant pattern of SMW space time variability center at 15 N, @Tv/@n is weakened at 7.5 N and 300 hPa in the tropics is related to ENSO. Changes in the tropo- (Figure 9b), resulting in SMW pressure increases and slower spheric temperature structure alter @Tv/@n such that the overlying wind speeds (Figure 9a). Poleward of the cooling tropical SMW topography is somewhat flatter (more iso- center at 40 N, @Tv/@n is weakened (Figure 9b) and the baric) during the warm phase of ENSO. Changes in tropo- overlying SMW is slower and lower (Figure 9a). spheric @Tv/@n align well with changes in upper [38] While ENSO-related changes in the east Pacific tropospheric wind speed, and the well-known eastward SMW are significant as far north as 65 N (Figure 9a), expansion and strengthening of the Pacific jet stream is ENSO-related changes in the west Pacific are more tropical, associated with SMW pressures up to 27 hPa lower near extending instead to 30 N (Figures 9d and 8a). The upper 30 N. The east Pacific SMW pressure decrease is flanked troposphere in the west Pacific exhibits ENSO-related cool- by slower upper tropospheric winds and SMW pressure ing at 30 N and 300 hPa (Figure 9f), collocated and increases near the equator and the Gulf of Alaska. consistent with an ENSO-related upward vertical velocity [42] Some of the interannual differences in SMW pres- anomaly in the Hadley circulation cell [Wang, 2002]. @Tv/@n sure declared significant in section 3 are small (20 to 50 hPa) is weakened (strengthened) in the column immediately relative to the vertical spacing of the Reanalysis data within poleward (equatorward) of the cooling. Upper tropospheric the SMW domain (50 or 100 hPa). The Reanalysis vertical winds are faster and SMW pressures are lower over the resolution is sufficient to support the statistical significance column of increased @Tv/@n at 20 N. Warming near the of these conclusions for two reasons. First, a large sample lapse-rate tropopause at 25 N associated with ENSO weak- size informs each seasonal mean: 360 (or 364 for leap years) ens @Tv/@n in the equatorial troposphere above which individual SMW pressure measurements are used for each near tropopause winds are slower and SMW pressures are seasonal mean calculation at each node. Second, and more increased. Considering Figures 9a and 9d together, the important, the within-season pressure variability of the warm phase of ENSO is associated with a flattening of the SMW is sufficiently large relative to the vertical spacing SMW topography equatorward of 30 N when viewed along of the Reanalysis data. Using winter 2004 as an example, meridians. Warm phase related flattening of the SMW calculation of the seasonal mean SMW pressure involved topography is also apparent in Figure 8b which views the seven isobaric surfaces at nodes accounting for more than ENSO-related changes in SMW pressure along parallels. half the NH surface area, and no single node incorporated fewer than four isobaric surfaces in its seasonal mean SMW 4. Summary and Discussion pressure calculation. For winter 2004, the 95% confidence intervals on the estimate of the seasonal mean are typically [39] Mapping the mean three-dimensional structure of the Pe ± 8.1 hPa and rarely exceed Pe ±16.0hPa.The winter SMW reveals a surface that generally slopes upward seasonal mean SMW pressure results also exhibit a from the pole to the equator with a steep midlatitude logical relationship to seasonal mean wind speed on transition close to the polar front jets. The SMW is generally nearby isobaric surfaces [Strong and Davis, 2005], and 50 to 100 hPa below the lapse-rate tropopause except within the seasonal mean SMW pressure results correlate logi- fast jet cores that extend laterally into the stratosphere cally with independently-developed climate indices such through tropopause folds or breaks (e.g. the winter jet stream as the AO (section 3.3). over the western Pacific). The SMW’s structural relationship [43] The geostrophic thermal wind example discussed in with the dynamic tropopause is similar to the SMW’s the Introduction (Figure 1 and section 1.2) makes clear the structural relationship with the lapse-rate tropopause. dependence of SMW pressure on the temperature structure [40] The dominant pattern of SMW space-time variability of the underlying atmosphere. PCA was used to identify the in the winter extratropics is related to the AO such that, AO and ENSO as the two leading contributors to SMW
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Figure 9. The symbol D indicates mean value for winters when multivariate El Nin˜o/Southern Oscillation Index (MEI) 1 minus mean value for years when MEI 1 for 1958–2004. Each panel shows the 47-year mean tropopause with a thin gray line. Figures 9a–9c are for 140 W, and Figures 9d– 9f are for 120 E. (a, d) DV contoured gray at 4 m/s with negative values dashed. Mean winter surface of maximum wind pressure (Pe) for winters with MEI 1 (solid black line) and MEI 1 (dashed black 1 line) with circles showing differences significant in a t test at a = 0.05. (b, e) D[ Rdf @Tv=@n] contoured at 3 K with negative values dashed. (c, f) DT contoured at 0.5 K with negative values dashed.
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~ ~ ~ pressure variability over the NH, and their influences on the where b is the angle between rTv and rF. Since krTvkcos SMW were studied in the context of associated changes in b is the portion or component of the virtual temperature tropospheric and stratospheric temperature. The upward gradient vector that operates in the direction of the trend in mean NH surface temperature could also contribute geopotential height gradient, (A2) may be equivalently to SMW and jet core pressure variability. Surface temper- expressed in natural coordinates ature increases have been largely confined to the dry, cold- @V R @T core anticyclones of North America and Siberia [Balling et g ¼ d v ðA3Þ al., 1998; Michaels et al., 2000], likely altering the thick- @ ln p f @n ness structure and distribution of isobaric temperature ~ gradients in the overlying atmosphere. Indeed, for the where n is oriented normal to V g and defined positive to the winter mean SMW pressure PCA performed using the entire right of the flow direction. Equation (A3) is also obtained if the isobaric geostrophic wind magnitude expressed in NH as input, the third principal component is significantly 1 correlated with NH mean surface temperature, identifying a natural coordinates (Vg = f @F/@n) is differentiated with relationship needing further study. The SMW sensitivity to respect to pressure. upper tropospheric and lower stratospheric temperature [46] Acknowledgments. Comments from three anonymous reviewers gradients motivates its investigation in the context of trends helped to improve this manuscript. Discussions with J. W. Nielsen- and variability in the atmospheric thermal structure [Angell, Gammon regarding the dynamic tropopause were very helpful and are 1998; Pielke et al., 1998; Christy et al., 2000] and their gratefully acknowledged. C. Strong was supported under a National relationship to surface temperature variability [Hurrell and Science Foundation Graduate Research Fellowship. Trenberth, 1998; Gaffen et al., 2000; Michaels et al., 2000]. References [44] Although the mean wind speed on the SMW is Angell, J. K. (1998), Contraction of the 300 mbar north circumpolar vortex shown here, its variability is not explicitly considered. during 1963–1997 and its movement into the eastern hemisphere, When the geostrophic thermal wind equation in natural J. Geophys. Res., 103(D20), 25,887–25,894. coordinates (equation (A3)) is integrated from the surface Angell, J. K. (2001), Relation of size and displacement of the 300 mbar north circumpolar vortex to QBO, El Nin˜o, and sunspot number, 1963– to the SMW, we may expect a faster (slower) wind maxi- 2000, J. Geophys. Res., 106, 31,787–31,794. mum when the integration is performed over a deeper Angell, J. K., and J. Korshover (1977), Variation in size and location of the (shallower) layer unless there is a compensating change in 300 mb North circumpolar vortex between 1963 and 1975, Mon. Weather Rev., 105, 19–25. the horizontal gradient of temperature. Indeed, when upper Angell, J. K., and J. Korshover (1985), Displacements of the north circum- tropospheric wind speed changes accompanied the exam- polar vortex during El Nin˜o, 1963 –1983, Mon. Weather Rev., 113, ples of SMW pressure change examined here, the underly- 1627–1630. Arkin, P. A. (1982), The relationship between interannual variability in the ing changes in @Tv/@n were of sufficient magnitude and 200 mb tropical wind field and the Southern Oscillation, Mon. Weather vertical depth that deceleration most often occurred with Rev., 110, 1393–1404. SMW pressure increases and acceleration most often oc- Balling, R. C., P. J. Michaels, and P. C. Knappenberger (1998), Analysis of curred with SMW pressure decreases. 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