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Abstract
Circulations in the vertical-lateral and vertical-longitudinal planes in an unstable planetary boundary layer are compared through the use of tetroon trajectories. On the average, the circulation in the transverse plane is 40% greater than in the longitudinal, but in the afternoon the transverse circulation is twice as great, providing evidence for the existence of longitudinal roll vortices at this time. The absolute magnitude of the transverse circulation increases uniformly with increase in wind speed and increase in the depth of the well-mixed layer, but the longitudinal circulation does not. The tetroon-derived stress increases with increase in wind speed and increase in absolute transverse circulation, suggesting that longitudinal roll vortices represent an efficient mechanism for the earthward transport of momentum. Comparisons are made between these tetroon results and results obtained by Deardorff from a three-dimensional numerical model of the unstable planetary boundary layer.
Abstract
Circulations in the vertical-lateral and vertical-longitudinal planes in an unstable planetary boundary layer are compared through the use of tetroon trajectories. On the average, the circulation in the transverse plane is 40% greater than in the longitudinal, but in the afternoon the transverse circulation is twice as great, providing evidence for the existence of longitudinal roll vortices at this time. The absolute magnitude of the transverse circulation increases uniformly with increase in wind speed and increase in the depth of the well-mixed layer, but the longitudinal circulation does not. The tetroon-derived stress increases with increase in wind speed and increase in absolute transverse circulation, suggesting that longitudinal roll vortices represent an efficient mechanism for the earthward transport of momentum. Comparisons are made between these tetroon results and results obtained by Deardorff from a three-dimensional numerical model of the unstable planetary boundary layer.
Abstract
Between June and November of 1970, 26 GHOST-type constant-level balloons were released from Ascension Island (8S) for flight at 30 and 50 mb. The balloons were positioned by the Interrogation, Recording and Location System (IRLS) aboard the Nimbus D satellite. Eight of the flights at 50 mb and three of the flights at 30 mb were tracked for more than one month, and one 50-mb flight was tracked continuously for more than five months while making seven circumnavigations of the earth. During the period June 1970 to March 1971, the 50-mb flights drifted northward at a mean speed of ∼0.1 m sec−1. The northward drift was a maximum in the Northern Hemisphere winter, suggesting a weak upward extension of the Hadley cell to at least 50 mb (the 30-mb data were insufficient for such an analysis). Superimposed an this drift were oscillations in meridional velocity having an approximate two-month period, with these oscillations also being most pronounced in the Northern Hemisphere winter. Small (1–3 m set−1) short-period fluctuations in meridional velocity were evident directly above the equator at 50 rob. Thew waves appear to move wesiwaid at speeds of 30–40 m sec−1 and to have a wavelength of about 90’ longitude. They were responsible for transporting small amounts of westerly momentum into the winter hemisphere. Kelvin waves were also delineated by flights near the equator. These waves appear to move eastward at speeds of 30–40 m sec−1 and to have a wavelength of 360° longitude. Some comparisons are made between these IRLS data and those obtained from GHOST balloon flights at the same heights in early 1969.
Abstract
Between June and November of 1970, 26 GHOST-type constant-level balloons were released from Ascension Island (8S) for flight at 30 and 50 mb. The balloons were positioned by the Interrogation, Recording and Location System (IRLS) aboard the Nimbus D satellite. Eight of the flights at 50 mb and three of the flights at 30 mb were tracked for more than one month, and one 50-mb flight was tracked continuously for more than five months while making seven circumnavigations of the earth. During the period June 1970 to March 1971, the 50-mb flights drifted northward at a mean speed of ∼0.1 m sec−1. The northward drift was a maximum in the Northern Hemisphere winter, suggesting a weak upward extension of the Hadley cell to at least 50 mb (the 30-mb data were insufficient for such an analysis). Superimposed an this drift were oscillations in meridional velocity having an approximate two-month period, with these oscillations also being most pronounced in the Northern Hemisphere winter. Small (1–3 m set−1) short-period fluctuations in meridional velocity were evident directly above the equator at 50 rob. Thew waves appear to move wesiwaid at speeds of 30–40 m sec−1 and to have a wavelength of about 90’ longitude. They were responsible for transporting small amounts of westerly momentum into the winter hemisphere. Kelvin waves were also delineated by flights near the equator. These waves appear to move eastward at speeds of 30–40 m sec−1 and to have a wavelength of 360° longitude. Some comparisons are made between these IRLS data and those obtained from GHOST balloon flights at the same heights in early 1969.
Abstract
The mean monthly polar stereographic map analyses of the Free University of Berlin terminated at the end of 2001. This paper summarizes the changes in size of the 300-mb north circumpolar vortex, and quadrants, for the full period of record, 1963–2001, where the size has been defined by planimetering the area poleward of contours in the jet stream core. A contracted vortex has tended to be a deep vortex in winter but a shallow vortex in summer. During 1963–2001 there was a statistically significant decrease in vortex size of 1.5% per decade, the decrease in size of Western Hemisphere quadrants being twice that of Eastern Hemisphere quadrants. A significant increase in Arctic Oscillation (AO) index accompanies the significant decrease in vortex size, but since the vortex contracts appreciably in all four seasons, whereas the positive trend in the AO index is mainly in winter, the vortex cannot serve as a proxy for the AO index. The evidence for vortex contraction at the time of the 1976–77 regime shift is not conclusive, but there is good evidence for a 6% increase in vortex size due to the 1991 Pinatubo eruption. There is little change in vortex size following the 1982 El Chichon eruption, however. Because on average there is a significant 4% contraction of the vortex following an El Niño, it is proposed that the vortex expansion to be expected following the 1982 El Chichon eruption has been contravened by the contraction following the strong 1982–83 El Niño. There is little relation between vortex size and phase of the quasi-biennial oscillation (QBO), and the evidence for a contracted vortex near 11-yr sunspot maxima is tenuous because the vortex record extends through only three full sunspot cycles. There is a highly significant tendency for opposite vortex quadrants 0°–90°E and 90°W–180° to vary in size together, indicating either a pulsating polar vortex or the propagation of planetary wavenumber 2.
Abstract
The mean monthly polar stereographic map analyses of the Free University of Berlin terminated at the end of 2001. This paper summarizes the changes in size of the 300-mb north circumpolar vortex, and quadrants, for the full period of record, 1963–2001, where the size has been defined by planimetering the area poleward of contours in the jet stream core. A contracted vortex has tended to be a deep vortex in winter but a shallow vortex in summer. During 1963–2001 there was a statistically significant decrease in vortex size of 1.5% per decade, the decrease in size of Western Hemisphere quadrants being twice that of Eastern Hemisphere quadrants. A significant increase in Arctic Oscillation (AO) index accompanies the significant decrease in vortex size, but since the vortex contracts appreciably in all four seasons, whereas the positive trend in the AO index is mainly in winter, the vortex cannot serve as a proxy for the AO index. The evidence for vortex contraction at the time of the 1976–77 regime shift is not conclusive, but there is good evidence for a 6% increase in vortex size due to the 1991 Pinatubo eruption. There is little change in vortex size following the 1982 El Chichon eruption, however. Because on average there is a significant 4% contraction of the vortex following an El Niño, it is proposed that the vortex expansion to be expected following the 1982 El Chichon eruption has been contravened by the contraction following the strong 1982–83 El Niño. There is little relation between vortex size and phase of the quasi-biennial oscillation (QBO), and the evidence for a contracted vortex near 11-yr sunspot maxima is tenuous because the vortex record extends through only three full sunspot cycles. There is a highly significant tendency for opposite vortex quadrants 0°–90°E and 90°W–180° to vary in size together, indicating either a pulsating polar vortex or the propagation of planetary wavenumber 2.
Abstract
Based on the 120-station North American radiosonde network, temperature trends for 100–50-mb (low stratosphere), 300–100-mb (tropopause), and 850–300-mb (troposphere) layers, and Earth’s surface, are evaluated for six 10° latitude bands extending from 20°–30°N to 70°–80°N for the 20-yr interval 1975–94. Confidence estimates are indicated by two standard errors of the least squares regression. In the average for the six latitude bands, the 100–50-mb annual temperature trend is −0.5°C decade−1 and the 850–300-mb trend is 0.2°C decade−1. In spring at 70°–80°N, the 100–50-mb and 300–100-mb layers cool by almost 2°C decade−1. The 300–100-mb layer cools by 0.7°C decade−1 relative to the 850–300-mb layer at 70°–80°N, but the two layers have the same warming trend at 20°–30°N, indicating the transition from the 300–100-mb layer being mostly in the stratosphere in polar regions to mostly in the troposphere in the northern subtropics. The surface warms much more than the troposphere at 70°–80°N (showing that surface temperature trends are not representative of tropospheric trends in polar regions) and slightly more at 20°–30°N, but surface warming is less than tropospheric warming in the 40°–70°N belt. At the surface at the radiosonde sites the 1200 UTC (morning) temperature cools relative to the 0000 UTC (evening) temperature by 0.05°C per decade on average, but the 850–300-mb temperature trends at 0000 UTC and 1200 UTC are essentially the same. The 0.7°C decade−1 cooling of low stratosphere relative to troposphere increases to 0.9°C decade−1 when adjustment is made for the stratospheric warming and tropospheric cooling following El Chichon and Pinatubo eruptions. The temperature trends obtained from 11 North American radiosonde stations in a 63-station global network agree well with the trends based on the entire 120-station network, and the latter are fairly representative of zonally averaged trends based on the 63-station network and microwave sounding unit data. Comparison with Canadian ozonesonde data shows that, in the low stratosphere and high troposphere during 1975–94, a decrease in temperature of 1°C decade−1 was associated with a decrease in ozone of about 10% decade−1.
Abstract
Based on the 120-station North American radiosonde network, temperature trends for 100–50-mb (low stratosphere), 300–100-mb (tropopause), and 850–300-mb (troposphere) layers, and Earth’s surface, are evaluated for six 10° latitude bands extending from 20°–30°N to 70°–80°N for the 20-yr interval 1975–94. Confidence estimates are indicated by two standard errors of the least squares regression. In the average for the six latitude bands, the 100–50-mb annual temperature trend is −0.5°C decade−1 and the 850–300-mb trend is 0.2°C decade−1. In spring at 70°–80°N, the 100–50-mb and 300–100-mb layers cool by almost 2°C decade−1. The 300–100-mb layer cools by 0.7°C decade−1 relative to the 850–300-mb layer at 70°–80°N, but the two layers have the same warming trend at 20°–30°N, indicating the transition from the 300–100-mb layer being mostly in the stratosphere in polar regions to mostly in the troposphere in the northern subtropics. The surface warms much more than the troposphere at 70°–80°N (showing that surface temperature trends are not representative of tropospheric trends in polar regions) and slightly more at 20°–30°N, but surface warming is less than tropospheric warming in the 40°–70°N belt. At the surface at the radiosonde sites the 1200 UTC (morning) temperature cools relative to the 0000 UTC (evening) temperature by 0.05°C per decade on average, but the 850–300-mb temperature trends at 0000 UTC and 1200 UTC are essentially the same. The 0.7°C decade−1 cooling of low stratosphere relative to troposphere increases to 0.9°C decade−1 when adjustment is made for the stratospheric warming and tropospheric cooling following El Chichon and Pinatubo eruptions. The temperature trends obtained from 11 North American radiosonde stations in a 63-station global network agree well with the trends based on the entire 120-station network, and the latter are fairly representative of zonally averaged trends based on the 63-station network and microwave sounding unit data. Comparison with Canadian ozonesonde data shows that, in the low stratosphere and high troposphere during 1975–94, a decrease in temperature of 1°C decade−1 was associated with a decrease in ozone of about 10% decade−1.
Abstract
A 63-station radiosonde network has been used for many years to estimate temperature variations and trends at the surface and in the 850–300-, 300–100-, and 100–50-mb layers of climate zones, both hemispheres, and the globe, but with little regard for the quality of individual station data. In this paper, nine tropical radiosonde stations in this network are identified as anomalous based on unrepresentatively large standard-error-of-regression values for 300–100-mb trends for the period 1958–2000. In the Tropics the exclusion of the 9 anomalous stations from the 63-station network for 1958–2000 results in a warming of the 300–100-mb layer rather than a cooling, a doubling of the warming of the 850–300-mb layer to a value of 0.13 K decade−1, and a greater warming at 850–300-mb than at the surface. The global changes in trend are smaller, but include a change to the same warming of the surface and the 850–300-mb layer during 1958–2000. The effect of the station exclusions is much less for 1979–2000, suggesting that most of the data problems are before this time. Temperature trends based on the 63-station network are compared with the Microwave Sounding Unit (MSU) and other radiosonde trends, and agreement is better after the exclusion of the anomalous stations. There is consensus that in the Tropics the troposphere has warmed slightly more than the surface during 1958–2000, but that there has been a warming of the surface relative to the troposphere during 1979–2000. Globally, the warming of the surface and the troposphere are essentially the same during 1958–2000, but during 1979–2000 the surface warms more than the troposphere. During the latter period the radiosondes indicate considerably more low-stratospheric cooling in the Tropics than does the MSU.
Abstract
A 63-station radiosonde network has been used for many years to estimate temperature variations and trends at the surface and in the 850–300-, 300–100-, and 100–50-mb layers of climate zones, both hemispheres, and the globe, but with little regard for the quality of individual station data. In this paper, nine tropical radiosonde stations in this network are identified as anomalous based on unrepresentatively large standard-error-of-regression values for 300–100-mb trends for the period 1958–2000. In the Tropics the exclusion of the 9 anomalous stations from the 63-station network for 1958–2000 results in a warming of the 300–100-mb layer rather than a cooling, a doubling of the warming of the 850–300-mb layer to a value of 0.13 K decade−1, and a greater warming at 850–300-mb than at the surface. The global changes in trend are smaller, but include a change to the same warming of the surface and the 850–300-mb layer during 1958–2000. The effect of the station exclusions is much less for 1979–2000, suggesting that most of the data problems are before this time. Temperature trends based on the 63-station network are compared with the Microwave Sounding Unit (MSU) and other radiosonde trends, and agreement is better after the exclusion of the anomalous stations. There is consensus that in the Tropics the troposphere has warmed slightly more than the surface during 1958–2000, but that there has been a warming of the surface relative to the troposphere during 1979–2000. Globally, the warming of the surface and the troposphere are essentially the same during 1958–2000, but during 1979–2000 the surface warms more than the troposphere. During the latter period the radiosondes indicate considerably more low-stratospheric cooling in the Tropics than does the MSU.
Abstract
Discussed is the use of tetroons to obtain estimates of the Reynolds stresses, the rate of production of eddy kinetic energy through these stresses, the coefficient of eddy viscosity, and the viscous dissipation within the planetary boundary layer. At an average height of about 3000 ft. (admittedly in the upper portion of the boundary layer) the tetroons yield a mean value for the zonal Reynolds stress of 1.5 dynes/cm2. This value increases by a factor of three as the lapse rate increases by a factor of one-half. The tetroons yield an average value for the rate of production of eddy kinetic energy through Reynolds stresses of 5.4 cm.2/sec.2, but this value is probably an underestimate inasmuch as the tetroons appear systematically to underestimate the wind shear in the vertical. The rate of production of eddy kinetic energy through Reynolds stresses decreases radically with increase in lapse rate. On the average, both of the above parameters increase with increase in the 3000-ft wind speed.
The mean tetroon-derived value for the coefficient of eddy (dynamic) viscosity is 0.57×103 gm. cm.−1 sec.−1 This mean value appears somewhat large (in accord with the above-mentioned underestimate of the vertical shear) and the scatter of inividual flight values suggests that this may be too sensitive a parameter to estimate from the tetroon data.
Equating the rate of production of eddy kinetic energy through Reynolds stresses to the dissipation appears to yield too large a value for the dissipation. Consequently, buoyancy and flux divergence terms probably can not be neglected and it may be necessary to place temperature instruments on the tetroons before acceptable dissipation estimates can be obtained.
Abstract
Discussed is the use of tetroons to obtain estimates of the Reynolds stresses, the rate of production of eddy kinetic energy through these stresses, the coefficient of eddy viscosity, and the viscous dissipation within the planetary boundary layer. At an average height of about 3000 ft. (admittedly in the upper portion of the boundary layer) the tetroons yield a mean value for the zonal Reynolds stress of 1.5 dynes/cm2. This value increases by a factor of three as the lapse rate increases by a factor of one-half. The tetroons yield an average value for the rate of production of eddy kinetic energy through Reynolds stresses of 5.4 cm.2/sec.2, but this value is probably an underestimate inasmuch as the tetroons appear systematically to underestimate the wind shear in the vertical. The rate of production of eddy kinetic energy through Reynolds stresses decreases radically with increase in lapse rate. On the average, both of the above parameters increase with increase in the 3000-ft wind speed.
The mean tetroon-derived value for the coefficient of eddy (dynamic) viscosity is 0.57×103 gm. cm.−1 sec.−1 This mean value appears somewhat large (in accord with the above-mentioned underestimate of the vertical shear) and the scatter of inividual flight values suggests that this may be too sensitive a parameter to estimate from the tetroon data.
Equating the rate of production of eddy kinetic energy through Reynolds stresses to the dissipation appears to yield too large a value for the dissipation. Consequently, buoyancy and flux divergence terms probably can not be neglected and it may be necessary to place temperature instruments on the tetroons before acceptable dissipation estimates can be obtained.
Abstract
The areal coverage to be expected from 300-millibar constant-pressure balloons (CPB) is examined for various numbers of releases per day from various stations. The meteorological variables which can be measured are discussed. The accuracy of balloon fixes required for useful determinations of velocity and acceleration is determined. It is shown, both theoretically and by computations from nine flights conducted by the Naval Research Laboratory in 1953, that the CPB data are adequate for the representation of the wind and pressure fields over the area covered by the trajectories. Together with radiosonde and rawin data they enable determination of the vertical velocity, but measurement of all the necessary parameters by CPB alone has not been achieved yet.
There is presented an analysis of the geostrophic deviations computed from the trajectory data, including their variation with latitude, wind speed, and pressure-gradient force.
Abstract
The areal coverage to be expected from 300-millibar constant-pressure balloons (CPB) is examined for various numbers of releases per day from various stations. The meteorological variables which can be measured are discussed. The accuracy of balloon fixes required for useful determinations of velocity and acceleration is determined. It is shown, both theoretically and by computations from nine flights conducted by the Naval Research Laboratory in 1953, that the CPB data are adequate for the representation of the wind and pressure fields over the area covered by the trajectories. Together with radiosonde and rawin data they enable determination of the vertical velocity, but measurement of all the necessary parameters by CPB alone has not been achieved yet.
There is presented an analysis of the geostrophic deviations computed from the trajectory data, including their variation with latitude, wind speed, and pressure-gradient force.
Abstract
Dynamical links of the Northern Hemisphere stratosphere and troposphere are studied, with an emphasis on whether stratospheric changes have a direct effect on tropospheric weather and climate. In particular, downward propagation of stratospheric anomalies of polar temperature in the winter–spring season is examined based upon 22 years of NCEP–NCAR reanalysis data. It is found that the polar stratosphere is sometimes preconditioned, which allows a warm anomaly to propagate from the upper stratosphere to the troposphere, and sometimes it prohibits downward propagation. The Arctic Oscillation (AO) is more clearly seen in the former case. To understand what dynamical conditions dictate the stratospheric property of downward propagation, the upper-stratospheric warming episodes with very large anomalies (such as stratospheric sudden warming) are selected and divided into two categories according to their downward-propagating features. Eliassen–Palm (E–P) diagnostics and wave propagation theories are used to examine the characteristics of wave–mean flow interactions in the two different categories. It is found that in the propagating case the initial wave forcing is very large and the polar westerly wind is reversed. As a result, dynamically induced anomalies propagate down as the critical line descends. A positive feedback is that the dramatic change in zonal wind alters the refractive index in a way favorable for continuous poleward transport of wave energy. The second pulse of wave flux conducts polar warm anomalies farther down. Consequently, the upper-tropospheric circulations are changed, in particular, the subtropical North Atlantic jet stream shifts to the south by ∼5 degrees of latitude, and the alignment of the jet stream becomes more zonal, which is similar to the negative phase of the North Atlantic Oscillation (NAO).
Abstract
Dynamical links of the Northern Hemisphere stratosphere and troposphere are studied, with an emphasis on whether stratospheric changes have a direct effect on tropospheric weather and climate. In particular, downward propagation of stratospheric anomalies of polar temperature in the winter–spring season is examined based upon 22 years of NCEP–NCAR reanalysis data. It is found that the polar stratosphere is sometimes preconditioned, which allows a warm anomaly to propagate from the upper stratosphere to the troposphere, and sometimes it prohibits downward propagation. The Arctic Oscillation (AO) is more clearly seen in the former case. To understand what dynamical conditions dictate the stratospheric property of downward propagation, the upper-stratospheric warming episodes with very large anomalies (such as stratospheric sudden warming) are selected and divided into two categories according to their downward-propagating features. Eliassen–Palm (E–P) diagnostics and wave propagation theories are used to examine the characteristics of wave–mean flow interactions in the two different categories. It is found that in the propagating case the initial wave forcing is very large and the polar westerly wind is reversed. As a result, dynamically induced anomalies propagate down as the critical line descends. A positive feedback is that the dramatic change in zonal wind alters the refractive index in a way favorable for continuous poleward transport of wave energy. The second pulse of wave flux conducts polar warm anomalies farther down. Consequently, the upper-tropospheric circulations are changed, in particular, the subtropical North Atlantic jet stream shifts to the south by ∼5 degrees of latitude, and the alignment of the jet stream becomes more zonal, which is similar to the negative phase of the North Atlantic Oscillation (NAO).
Abstract
The utility of a “first difference” method for producing temporally homogeneous large-scale mean time series is assessed. Starting with monthly averages, the method involves dropping data around the time of suspected discontinuities and then calculating differences in temperature from one year to the next, resulting in a time series of year-to-year differences for each month at each station. These first difference time series are then combined to form large-scale means, and mean temperature time series are constructed from the first difference series. When applied to radiosonde temperature data, the method introduces random errors that decrease with the number of station time series used to create the large-scale time series and increase with the number of temporal gaps in the station time series. Root-mean-square errors for annual means of datasets produced with this method using over 500 stations are estimated at no more than 0.03 K, with errors in trends less than 0.02 K decade−1 for 1960–97 at 500 mb. For a 50-station dataset, errors in trends in annual global means introduced by the first differencing procedure may be as large as 0.06 K decade−1 (for six breaks per series), which is greater than the standard error of the trend. Although the first difference method offers significant resource and labor advantages over methods that attempt to adjust the data, it introduces an error in large-scale mean time series that may be unacceptable in some cases.
Abstract
The utility of a “first difference” method for producing temporally homogeneous large-scale mean time series is assessed. Starting with monthly averages, the method involves dropping data around the time of suspected discontinuities and then calculating differences in temperature from one year to the next, resulting in a time series of year-to-year differences for each month at each station. These first difference time series are then combined to form large-scale means, and mean temperature time series are constructed from the first difference series. When applied to radiosonde temperature data, the method introduces random errors that decrease with the number of station time series used to create the large-scale time series and increase with the number of temporal gaps in the station time series. Root-mean-square errors for annual means of datasets produced with this method using over 500 stations are estimated at no more than 0.03 K, with errors in trends less than 0.02 K decade−1 for 1960–97 at 500 mb. For a 50-station dataset, errors in trends in annual global means introduced by the first differencing procedure may be as large as 0.06 K decade−1 (for six breaks per series), which is greater than the standard error of the trend. Although the first difference method offers significant resource and labor advantages over methods that attempt to adjust the data, it introduces an error in large-scale mean time series that may be unacceptable in some cases.
