changes in the brewer-dobson circulation for 1980–2009 revealed in merra reanalysis data
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Asia-Pac. J. Atmos. Sci., 50(S), 73-92, 2014 pISSN 1976-7633 / eISSN 1976-7951
DOI:10.1007/s13143-014-0051-4
Changes in the Brewer-Dobson Circulation for 1980-2009 Revealed in MERRA
Reanalysis Data
Jong-Yoon Kim*, Hye-Yeong Chun, and Min-Jee Kang
Department of Atmospheric Sciences, Yonsei University, Seoul, Korea
(Manuscript received 17 July 2014; accepted 9 October 2014)© The Korean Meteorological Society and Springer 2014
Abstract: Changes in the Brewer-Dobson circulation (BDC) during
the 30 years 1980-2009 are investigated using Modern Era Retrospective-
analysis for Research and Applications (MERRA) reanalysis data.
The mass streamfunction that is induced by wave forcings in the
transformed Eulerian-mean (TEM) equation through the downward-
control principle is used as a proxy for the BDC. The changes in the
BDC are investigated using two aspects: the wave propagation
conditions in the stratosphere and the wave activity in the upper
troposphere. They are compared in the first (P1) and second (P2) 15-
year periods. The resolved wave forcing, expressed by the Eliassen-
Palm (EP) flux divergence (EPD), is significantly enhanced during
the December-January-February (DJF) season in P2 in both the
Northern Hemisphere (NH) high latitudes and the Southern
Hemisphere (SH) mid- and high latitudes. The increased zonal mean
zonal wind at high latitudes in the SH, caused by ozone depletion,
leads to an upward shift of the Rossby-wave critical layer and this
allows more transient planetary waves to propagate into the
stratosphere. In the NH, the enhanced EPD in DJF leads to an
increase in the frequency of Sudden Stratospheric Warming (SSW)
events. The gravity wave drag (GWD) is smaller than the EPD and
the change in it between the two time periods is insignificant. The
residual term in the TEM equation is similar to the GWD in the two
periods, but its change between the two periods is as large as the
change in the EPD. Among the four components of the EP flux at
250 hPa, the meridional heat flux played a dominant role in the
enhancement of the BDC in P2.
Key words: Brewer-Dobson circulation, climate change, wave
forcing, wave propagation condition
1. Introduction
The Brewer-Dobson circulation (BDC) is the chemical trans-
port circulation of the stratosphere (Brewer, 1949; Dobson,
1956), which includes mean meridional circulation and the
quasi-horizontal two-way mixing (Plumb, 2002; Shepherd,
2002; Shepherd, 2007; Birner and Bonisch, 2011). The
stratospheric meridional circulation, consisting of upwelling in
the tropics and downwelling in the extratropics (Andrews et
al., 1987; Plumb, 2002; Shepherd, 2007), is induced by wave
forcing propagating from the troposphere (Holton et al., 1995;
Plump and Eluszkiewicz, 1999; Semeniuk and Shepherd, 2001)
and it controls the mass exchange between the troposphere and
the stratosphere (Holton, 1990). The mean meridional cir-
culation also has a significant influence on the adiabatic
warming/cooling in the stratosphere. In particular, tropical
upwelling affects the water vapor input into the stratosphere by
determining the tropopause temperature (Kerr-Munslow and
Norton, 2006). Thompson and Solomon (2009) found con-
trasting latitudinal structures of recent stratospheric tempera-
ture and ozone trends by using satellite observation data. These
results are consistent with the enhancement of the BDC,
causing adiabatic warming/cooling in the stratosphere. Ac-
cording to Solomon et al. (2010), the stratospheric water vapor
concentration has decreased by 10% since 2000, because of
the tropopause cooling caused by the enhancement of tropical
upwelling. This decrease slowed down the rate of increase in
the global surface temperature between 2000 and 2009 by
approximately 25% in comparison to that which would have
occurred due to only carbon dioxide and other greenhouse
gases (GHGs). Based on the research results noted above,
changes in the BDC are very important subjects of study
because they have an influence on stratospheric and tropo-
spheric climate change.
Several studies have analyzed the mean meridional circul-
ation in the changed climate based on chemistry-climate model
(CCM) results, and suggested that the increase in the wave
propagation caused by the changes in the zonal mean temper-
ature and zonal wind led to the strengthening of the BDC
(Butchart and Scaife, 2001; Butchart et al., 2006, 2010;
Fomichev et al., 2007; Garcia and Randel, 2008; Li et al.,
2008; McLandress and Shepherd, 2009; Shepherd and
McLandress, 2011). According to Butchart et al. (2006, 2010),
by comparing various CCM results using the Intergovernmental
Panel on Climate Change (IPCC) Special Report on Emissions
(SRES) A1B scenario (Nakicenovic and Swart, 2000), the
annual mean tropical upwelling in the lower stratosphere
increases by almost 2% per decade on average, with 59% of
this trend being forced by the parameterized orographic gravity
wave drag (GWD) in the models, although there are differences
in the reason and intensity among the models. This is a
consequence of the eastward acceleration of the subtropical
jets, which increases the upward flux of the gravity-wave
*Current affiliation: International Cooperation Department, KoreaMeteorological Industry Promotion Agency, Seoul, Korea.Corresponding Author: Hye-Yeong Chun, Department of Atmos-pheric Sciences, Yonsei University, 134 Shinchon-dong, Seodaemun-ku, Seoul 120-749, Korea.E-mail: chunhy@yonsei.ac.kr
74 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
momentum that reaches the lower stratosphere in these
latitudes. Cagnazzo et al. (2006), using the Met Office United
Model, suggested that the cooling due to Arctic ozone
depletion in spring leads to a strengthening of the polar vortex
and this weakens the BDC by reducing wave propagation into
the polar region. They also suggested that the weakened BDC
causes a reduction in the ozone transport into the polar region
and adiabatic warming, which results in further cooling in the
polar region. Li et al. (2008) found that BDC strengthening is
most significant in the SH summer by using the Geophysical
Fluid Dynamics Laboratory coupled CCM. The reason for this
significant strengthening is that the Antarctic ozone depletion
increased the meridional temperature gradient through the
cooling of the Antarctic lower stratosphere during the past 4
decades, and the resulting westerly shift of the zonal wind
increased the planetary wave activity in the stratosphere. This
led to an increased downward mass flux. Shepherd and
McLandress (2011) suggested that the GHG-induced tropo-
spheric warming pushes the Rossby-wave critical layers
upward and this allows more planetary waves to penetrate into
the subtropical lower stratosphere.
Previous studies analyzed changes in the mean meridional
circulation, representing the BDC, using CCM simulations, but
the results varied among the different studies. In particular, the
contribution of planetary and gravity waves to the BDC
strengthening was widely different. It is known that the
difference comes from the basic structure of the models pro-
ducing the planetary waves and the reality of the gravity wave
parameterization scheme (Kim et al., 2003; Alexander et al.,
2010). Therefore, we investigate the changes and causes of the
BDC for the time period between 1980 and 2009 that are
revealed in the Modern Era Retrospective-analysis for Research
and Applications (MERRA) reanalysis data (Rienecker et al.,
2011). MERRA reanalysis is likely one of the best reanalysis
data sets for analyzing the influence of planetary waves and
gravity waves on BDC changes, because it provides not only
3-hourly data, but also GWD data that the other reanalysis data
sets do not provide. Hu and Tung (2002) and Iwasaki et al.
(2009) analyzed the BDC changes using various reanalysis
data. However, Hu and Tung (2002) analyzed only specific
latitudes (50oN and 60°N) in the NH and there is no detailed
analysis of the various wave forcings that drive the BDC. The
data by Iwasaki et al. (2009) was also restricted, because they
used a short time period of less than 20 years. In the present
study, the BDC is represented by the mean meridional
circulation and resultant mass fluxes in the tropics and ex-
tratropics.
This paper is organized as follows. Section 2 provides a
general description of the MERRA reanalysis data and section
3 describes the methodology. Section 4 describes the results.
We analyze the propagation conditions of the stationary and
transient planetary waves in the stratosphere and the wave
activity in the upper troposphere. The propagation conditions
are analyzed by the changes in the Rossby-wave critical levels
for the transient waves and the refractive index of the
stationary planetary waves. The changes in the wave activity in
the upper troposphere are investigated by analyzing the vertical
components of the Eliassen-Palm (EP) flux at 250 hPa and 70
hPa. Finally, section 5 summarizes our findings and provides
the conclusions.
2. Data
The MERRA reanalysis data has 42 levels with the top level
at 0.1 hPa. The horizontal resolution is 1.25o latitude by 1.25o
longitude. We use the 3-hourly data from January 1980 to
December 2009. The MERRA reanalysis data has realistic
zonal wind variability in the lower stratosphere (Rienecker et
al., 2011). The variables used to analyze the BDC are tem-
perature, the wind in the horizontal and vertical directions, and
the GWD. The Goddard Earth Observing System Model
Version 5 (GEOS-5) used for MERRA incorporates both an
orographic GWD scheme based on McFarlane (1987) and a
scheme for non-orographic waves based on Garcia and Boville
(1994). The GWD term provided by the MERRA reanalysis
data is the sum of the two terms. Figure 1 illustrates the annual
and zonal mean temperature anomaly between 1980 and 2009
from their 30-year mean. There is a weak warming trend in the
troposphere. In the stratosphere, positive deviation is shown
before 1995, while negative deviation is recorded after 1995.
The tropospheric warming is more significant after 1995 when
the stratospheric cooling began. Therefore, there may be
significant dynamical changes in the troposphere and the
stratosphere beginning in 1995. We investigate the BDC and
the change in the BDC for two separate 15-year periods before
and after 1995, which we denote as P1 and P2, respectively,
hereafter.
3. Methodology
a. Mass streamfunction
In this study, the BDC is expressed by the mass streamfunc-
tion as follows:
, (1)
, (2)
, (3)
where f is the Coriolis parameter (f = 2ωsinφ, ω is the earth
rotation rate, and φ is the latitude) χdc is the downward control
mass streamfunction derived by wave forcing (Haynes et al.,
1991), and χdirect is the mass streamfunction computed by the
residual mean circulation. The residual meridional and vertical
velocities are as follows (Andrews et al., 1987):
, (4)
χdc φ z,( ) cosφ
f-----------– EPD φ z',( ) GWD φ z',( ) X φ z',( )+ +
z
∞
∫ ρ0dz'=
χdirect φ z,( ) cosφ v*
z
∞∫ ρ0dz'–=
f f 1 acosϕ( )⁄ ∂ ∂φ⁄( ) ucosφ( )–=
v*
v1
ρ0
-----∂∂z-----
ρ0v'θ '
dθ dz⁄--------------⎝ ⎠⎛ ⎞–=
30 November 2014 Jong-Yoon Kim et al. 75
, (5)
where θ is the potential temperature, z is the log-pressure
height [z ≡ −Hln(p/ps), H is the scale height, p pressure, and ps
= 1000 hPa], ρ0 is the background density [ρ0 = ρSexp(-(z/H)),
ρs is the density at pS], a is the Earth’s radius, and ϕ is the
latitude. Assuming a steady-state, χdc and χdirect are equal, and
this assumption might hold for most of the condition in a
w*
w1
acosϕ--------------
∂∂ϕ------
v'θ 'cosϕ
dθ dz⁄--------------------⎝ ⎠⎛ ⎞+=
Fig. 1. Latitude-height cross sections of the annual mean, zonal mean temperature anomalies between 1980 and 2009 from their 30-yearmean.
76 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
realistic atmosphere (Haynes et al., 1991). However, the steady-
state assumption might cause some problems in the inter-
pretation of trends in the residual circulation of a realistic
stratosphere. Chun et al. (2011) found that dU/dt contributes to
the annual cycle of the residual mean vertical velocity and
have a clear semi-annual cycle, with the maxima at equinoxes
and minima at solstices. In the present study, we calculate χdc
based on the steady-state assumption, following the original
downward-control principle. The overbar and prime in Eq. (1)
through (5) denote the zonal mean and the departure of each
variable from its zonal-mean, respectively. In (1), EPD and X
denote the EP flux divergence and the residual in the
Transformed Eulerian mean (TEM) equation as follows:
.
(6)
The residual term incorporates not only the small scale
turbulences and GWD drag, which cannot be produced by the
GWD parameterization scheme, but also the imbalance caused
by the incremental analysis. We will discuss this issue further
in section 5. We calculated every fields including mass stream-
function up to 0.1 hPa, the top of MERRA reanalysis data,
although the results are shown up to 10 hPa in the present
study to focus on the stratosphere. It is noteworthy that the
streamfunction at certain level is determined by the wave
forcing that is integrated from that level to the top of atmo-
sphere (or model top), based on the downward-control
principle, and the streamfunction in the present study is
negligible above 10 hPa (not shown).
The strength of the BDC can be quantified using the net
upward mass flux in the tropical region determined by the
mass streamfunction at the turnaround latitude (Holton, 1990).
Based on Holton (1990), the area-averaged extratropical vertical
mass flux across a pressure surface in the NH ( ) and SH
( ) can be expressed as:
, (7)
. (8)
In this study, and are the turnaround latitudes in
each hemisphere, which are located at the minimum and
maximum of the mass streamfunction, respectively. Here,
either χdc or χdirect can be used to calculate FNH and FSH. Using
(7) and (8), the net upward mass flux in the tropical region
( ) can be estimated using the following equation:
. (9)
This calculation assumes that there is only upward mass flux
between the turnaround latitudes and only downward mass flux
in the other region. This implies that small changes in the mass
streamfunction at the turnaround latitudes are likely to cause
significant differences in the mass flux (Chun et al., 2011).
b. Eliassen-Palm flux Divergence (EPD)
The resolved planetary wave forcing represented by the EPD
is the primary forcing for the BDC. Following the work of
Andrews et al. (1987), the EPD is calculated using the
following equation:
. (10)
In this study, F(φ)and F(z) are the horizontal and vertical com-
ponents of the EP flux, respectively, and each term consists of the
sum of heat and momentum flux as follows (Chun et al., 2011):
, (11)
, (12)
, (13)
, (14)
where fa is the absolute vorticity given by the following
equation:
. (15)
In this study, the 3-hourly MERRA reanalysis data are first
daily averaged and only the zonal wavenumbers from 1 to 16
are calculated using a Fast Fourier Transform (FFT). The
wavenumber decomposition makes it possible to divide the
EPD into the planetary-scale component (sum of the zonal
wavenumbers 1-3) and the synoptic-scale component (sum of
the wavenumbers 4-16). We separate the EPD into its stationary
(monthly mean) and transient component (deviations from the
monthly mean) as well. In the EPD analysis, the results may be
sensitive to the choice of the period used for defining the
stationary component. In this study, we define the stationary
part as a monthly mean to insure consistency with previous
studies (McLandress and Shepherd, 2009; Shepherd and
McLandress, 2011).
c. Spectral analysis of transient Rossby waves
Following Randel and Held (1991) and Shepherd and
McLandress (2011), we calculate the cospectral density of the
momentum and heat flux in (11)-(14), say CS(ω, k), using the
FFT, where ω and k are the frequency and zonal wavenumber
of waves, respectively. The frequency is smoothed using a
Gaussian function with an e-folding width of 0.1 day−1 and
transformed into the phase velocity c and k space using CS(ω,
k)dω = CS(c, k)dc, where c = ωacosφ/k is the phase velocity.
∂u∂t------ f
1
acosφ--------------
∂∂φ------ ucosφ( )–⎝ ⎠
⎛ ⎞v* w*∂u∂z------–
1
ρ0acosφ-------------------∇+ F⋅ GWD X+ +=
FNH
↓
FSH
↓
FSH↓ 2πa
2ρ0 w
*
π 2⁄–
φTL
SH
∫ cosφ dφ 2πaχ φTLSH( )= =
FNH↓ 2πa
2ρ0 w
*
φTL
NH
π 2⁄∫ cosφdφ 2– πaχ φ
TLNH( )= =
φTLSH φ
TLNH
FTR
↑
FTR↑ FNH
↓ FSH↓+( )–=
EPD1
ρ0acosφ-------------------∇ F⋅=
1
ρ0acosφ-------------------
1
acosφ--------------
∂∂φ------ F
φ( )cosφ( ) ∂F
z( )
∂z----------+=
Fφ_1
ρ0acosφ u'v'–( )=
Fφ_2
ρ0acosφdu
dz------
v'θ '
dθ dz⁄--------------⎝ ⎠
⎛ ⎞=
Fz_1
ρ0acosφ fav'θ '
dθ dz⁄--------------⎝ ⎠
⎛ ⎞=
Fz_2
ρ0acosφ u'w'–( )=
fa f1
acosφ--------------
∂∂φ------ ucosφ( )–=
30 November 2014 Jong-Yoon Kim et al. 77
Fig. 2. Time series of the mass flux calculated by χdirect
at 70 hPa for: (a) tropical upward mass flux, (b) NH downwardmass flux, and (c) SH downward mass flux. The colored lines in (a)-(c) represent each season (blue: DJF (December-January-February), green: MAM (March-April-May), red: JJA (June-July-August), orange: SON (September-October-November), and black: annual average), (d)-(o) are the same as in (a)-(c) except for the time series in each season afterthe seasonal 30-year mean is subtracted. The dashed lines in (d)-(o) indicate the linear trends of the annual variation.The slope of the trend is represented as α, and * mark denotes statistically significant trend at a 95% confidence levelusing t-test.
78 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
d. Refractive index
The square of the refractive index (ηk
2), which characterizes
the propensity for stationary planetary wave propagation, is
given in spherical quasi-geostrophic form by the following
equation (Andrews et al., 1987):
, (16)
where the meridional derivation of the potential vorticity ( )
is as follows:
. (17)
N denotes the buoyancy frequency and uz is the vertical
shear of the zonal mean zonal wind. In general, the refractive
index increases with the vertical wind shear and decreases with
the zonal mean zonal wind (Tung and Lindzen, 1979; Hu and
Tung, 2002).
4. Results
a. Changes in the Brewer-Dobson circulation
Figure 2 shows the time series of the mass flux at 70 hPa
calculated using the mass streamfunction. As shown in (9),
(Fig. 2a) is equal and opposite to the sum of (Fig.
2b) and (Fig. 2c). In this figure, the curves with different
colors represent different seasons. Figures 2d-o are the same as
Figs. 2a-c, except that the 30-year mean is subtracted, and the
linear trends are denoted by dashed lines. increases with
time primarily in DJF and JJA and these increases are due to
the increased and the increased , respectively. In
particular, the rapid increase in the after 1995 in JJA is
due to a rapid increase in the . The increasing trend is also
evident in SON and this is due to the increasing trend in the
downward mass fluxes in both hemispheres. In Fig. 2,
statistical significance of the trend is examined by performing
the t-test that takes into account autocorrelation of the time
series using the methods by Wilks (2006) and Krishnamurthy
and Kirtman (2009). Based on this test, increasing trends in the
tropical upward mass flux in DJF, NH downward mass flux in
DJF, and NH downward mass flux in SON are statistically
significant at 95% confidence level. In this study, changes in
the BDC are investigated based on two aspects: wave
propagation conditions in the stratosphere and wave activity in
the upper troposphere. These results are compared in the first
(1980-1994, P1) and the second (1995-2009, P2) 15-year time
period.
ηk2 y z,( )
qϕ
u----- k
acosϕ--------------⎝ ⎠⎛ ⎞
2
–f
2NH-----------⎝ ⎠⎛ ⎞
2
–=
qϕ
qϕ2Ωa
-------cosϕ1
a2
----ucosϕ( )ϕcosϕ
---------------------ϕ
–f2
ρ0
----- ρ0uz
N2
-----⎝ ⎠⎛ ⎞
Z
–=
FTR↑ FNH
↓
FSH↓
FTR↑
FNH↓ FSH
↓
FTR↑
FSH↓
Fig. 3. Latitude-height cross sections of the zonal mean zonal wind averaged over (a) MAM, (b) JJA, (c) SON, and (d) DJF.The contours and colored shading denote the mean in P1 (contour intervals of 5 m s−1) and the statistically significant differencebetween P2 and P1 above a 95% confidence level, respectively.
30 November 2014 Jong-Yoon Kim et al. 79
b. Wave propagation conditions in the stratosphere
The changes in the zonal mean zonal wind between P1 and
P2 revealed in the MERRA reanalysis data for 1980-2009 in
this study are somewhat smaller than those of previous studies
using CCM simulations over 100 years due to a short analysis
period (Li et al., 2008; McLandress and Shepherd, 2009).
Figure 3 shows the zonal mean zonal wind in each season
averaged in the P1 period (contours) and the difference bet-
ween the P2 and P1 periods (colored shading indicates
statistically significant differences on the 95% confidence
level). Significance levels for the differences of the means are
computed using the Student’s t test. Except in the tropical
regions, statistically significant changes, more than 3.5 m s−1,
primarily appear in DJF, while there are small changes, less
than 1 m s−1, during the other seasons. Therefore, in this
section we focus on the change in the zonal mean zonal wind
in DJF. In the SH, a decrease in the temperature in the
Antarctic caused by ozone depletion strengthens the zonal
mean zonal wind at mid-latitudes during P2 through the
thermal wind relationship. It is noteworthy that Antarctic
ozone in P2 is still less than in P1 (http://ozonewatch.
gsfc.nasa.gov), although there is a significant positive trend
after 1996 when dynamically-induced changes of ozone are
removed (Salby et al., 2011), and consequently temperature is
the SH polar region in P2 is colder than in P1. In the NH,
however, there is a weakened zonal mean zonal wind in the
mid- to high latitudes. The decrease in the zonal mean zonal
wind in the NH mid- to high latitudes in DJF is related to the
more frequent occurrence of the SSW events in P2. According
to the World Meteorological Organization (WMO), the SSW is
a phenomenon in which the latitudinal mean temperature
increases poleward from 60oN and an associated circulation
reversal is observed at 10 hPa or below (McInturff, 1978).
When we select a SSW of which the zonal mean zonal wind at
60oN and 10 hPa is easterly during the NH winter (DJF), the
number of SSWs selected is 7 in P1 and 12 in P2, with strong
and long-lasting SSWs occurring more frequently after 2000.
Our analyses reveal (not shown) that the frequent occurrence
of SSWs in P2 stems from the enhanced meridional heat flux
in December, a month before most SSWs occur, primarily due
to planetary waves with a zonal wavenumber one. Although
SSWs are not the primary subject of the current study, changes
in the occurrence of SSWs between P1 and P2 is an interesting
research topic, given that it is strongly related to the planetary
waves propagated from the troposphere into the stratosphere.
This topic will be investigated in a future study.
Figure 4 shows the mass streamfunction induced by the
wave forcing averaged over DJF for P1 (Fig. 4a) and the
difference between P2 and P1 by the total forcing (Fig. 4b),
EPD, GWD, and residual terms (Figs. 4c-h). The red lines
denote the turnaround latitudes of the direct mass streamfunc-
tion in P1. Student’s t-test was used for statistical comparison,
and the shadings in Figs. 4b, d, f and h denote areas of
statistically significant differences between P2 and P1 at a 95%
confidence level. The BDC in P1 induced by the total wave
forcing (Fig. 4a) indicates that the NH circulation is stronger
than the SH circulation, as expected, because the wave forcing
in the winter hemisphere is stronger than that in the summer
hemisphere. However, the difference between P2 and P1 of the
BDC (Fig. 4b) from 100 hPa to about 30 hPa indicates that the
magnitude of the change signal is similar in both of the
hemispheres. Therefore, the BDC changes in the SH are also
important. It is clear that EPD is the primary forcing of the
BDC in both hemispheres as represented. The influence of the
EPD changes on the BDC changes (Fig. 4d) indicates that
there is statistically significant strengthening in the SH and this
contributes to the increase in the mass flux, because the
changes occur at the turnaround latitude. In the NH there is a
statistically meaningful increase in the BDC by the EPD at
high latitudes and this result is in agreement with the more
frequent occurrence of the SSW events in P2. In the NH mid-
Fig. 4. Latitude-height cross sections of the mass streamfunctionaveraged over DJF for (left) P1 and (right) the difference between P2and P1 (first row) by the total forcing, (second row) by the EPD, (thirdrow) by the GWD, and by the residual terms (fourth row). The contourintervals are 5 kg m−2 s−1 and 1 kg m−2 s−1 for the left and right panels,respectively. The shaded regions in (b), (d), (f), and (h) represent thestatistically significant anomalies at a 95% confidence level. The redlines denote the turnaround latitudes in P1.
80 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
latitudes where the turnaround latitude is included, however,
the BDC is weakened by the EPD, although it is not
statistically significant. The contribution of the GWD to the
BDC in P1 (Fig. 4e) is much smaller than that by the EPD
(Fig. 4c) and the change in P2 is also negligible. This result is
not consistent with the results of previous studies based on the
CCM simulations, which suggest that the GWD contributes to
the changes in BDC by more than 50 percent at 70 hPa (e.g.,
Butchart et al., 2010; Okamoto et al., 2011). This result is
caused by the difference of the GWD parameterization scheme
among the models, and we will discuss this issue further in
section 5.
The contribution of the residual term [X in Eq. (6)] in P1 is
generally larger than that by the GWD, especially in the NH
and SH subtropics, but the contribution is still much smaller
than that by the EPD. The sign of the mass streamfunction by
the residual term in the SH is generally the opposite to the sign
of the mass streamfunction by the total forcing term. However,
the changes in the BDC by the residual term are significant
and they are similar to the changes in the BDC by the EPD. In
the NH subtropics to mid-latitudes, a positive change by the
residual term make the total change nearly zero at 35oN after
compensating for the negative change by the EPD from 70 hPa
to 10 hPa. In the SH, the contribution of the residual term
equatorward of 45oS to the total mass streamfunction change is
predominant among the three forcing terms. In addition, the
residual change contributes to the increase in the mass flux,
because statistically significant changes occur at the turnaround
latitude. In previous studies using reanalysis data that did not
provide the GWD forcing term, the effects of the GWD were
estimated from a residual term without explicitly taking the
GWD, which is equivalent to GWD + X in the present study,
into consideration. Therefore, the contribution of the residual
term that was thought to expose the GWD to the changes in
the BDC was significant (e.g., Okamoto et al., 2011). This is
consistent with the results of previous studies using CCMs.
However, the present case is quite complicated using the
MERRA reanalysis data that provides the GWD output
explicitly. Originally, it is expected that the GWD forcing term
to be significant, at least more than half of the residual term.
This is because the dominant process of the residual term was
thought to be small-scale gravity waves that could not be
represented explicitly in model or reanalysis data grids. When
we found the GWD to be much smaller than the residual term
in the present MERRA data, we developed three hypotheses:
(i) The small-scale gravity waves represented by the GWD
parameterization used in the GEOS-5, the base model of
MERRA, are underestimated. Considering that no global
observational data of the GWD are available for use in the
assimilation process, the GWD output provided by MERRA is
purely model output, which relies significantly on the reality of
the GWD parameterization [Although observational estimation
of global GW momentum flux is available from satellite data
(e.g., Ern et al., 2004), no feasible GWD information is
available primarily due to lack of direction information of the
satellite-observed GW momentum flux.]; (ii) The residual term
also includes an imbalance caused by the incremental analysis
during the assimilation process, which may depend on the
numerical methodology; and (iii) The numerical diffusion
from the model can be another source of the residual term and,
in general, it can be relatively larger for higher resolution
models due to the use of relatively larger diffusion coefficients.
At this point, there is no straightforward method to estimate,
even roughly, the contributions of physical and numerical
processes to the residual term and caution needs to be taken
when interpreting this term in BDC studies.
Figure 5 shows the mass streamfunction averaged over JJA.
Figure 5a, which shows the BDC in P1 induced by the total
wave forcing, indicates that the SH (winter hemisphere) circul-
ation is stronger than the NH (summer hemisphere) circulation,
as expected from Fig. 4a. The difference between P2 and P1 of
the BDC (Fig. 5b) shows that the magnitude of change is
similar in both of the hemispheres in JJA. However, statistically
significant changes appear only by the residual term (Fig. 5h),
not by the EPD (Fig. 5d) and the GWD forcing (Fig. 5f) terms
from 70 hPa to 10 hPa. This is likely due to the fact that the
zonal mean zonal wind in JJA does not change much (less than
± 1 m s−1) in the P2 period compared to P1 period, as re-
Fig. 5. The same as in Fig. 4 except for JJA.
30 November 2014 Jong-Yoon Kim et al. 81
presented in Fig. 3b, and, consequently, the propagation
conditions of the planetary waves do not change much in the
P2 period.
Figure 6 shows , and calculated from the mass
streamfunction based on the downward-control principle
shown in Figs. 4 and 5. There are total forcing (blue) and
individual forcing of EPD (red), GWD (orange), and residual
term (green) at 70 hPa during DJF (Fig. 6a) and JJA (Fig. 6b),
along with those calculated from the direct mass stream-
function (black) as shown in Fig. 2. Several interesting features
can be found from Fig. 6. First, the tropical upward mass flux
at 70 hPa is larger in DJF than in JJA, due to larger values of
the downward mass fluxes in the respective winter and
summer hemisphere. Second, contribution by the EPD to the
tropical upward mass flux is predominant, due to its primary
contribution to the winter hemisphere downward mass flux.
Third, contribution by the GWD and the residual term to the
tropical upward mass flux is less than 10%, but contribution by
the GWD to the respective summer hemisphere downward
mass flux is about one third (half) of that by the EPD in DJF
(JJA). Fourth, tropical upward mass flux and extratropical
downward mass fluxes increase in P2 in both DJF and JJA,
due to the increase in mass fluxes by the EPD forcing and the
residual term in the NH in DJF and by the residual term in JJA,
as also evident from Figs. 4 and 5.
It is noteworthy that the turnaround latitudes in the NH and
SH used for the mass flux calculations in Figs. 2 and 6 are
those obtained from the direct mass streamfunction, which are
shown in the red curves in Figs. 4 and 5. The turnaround
latitudes are 27oN (26oN) and 47oS (48oS) in P1 (P2) in DJF
and 48oN (47oN) and 26oS (24oS) in P1 (P2) in JJA at 70 hPa,
however, there is no statistically significant shift of the turn-
around latitude between P2 and P1. These turnaround latitudes
are similar to those of the downward-control mass stream-
function by total forcings, but different from those by
individual forcing terms, as shown in Figs. 4 and 5. The
turnaround latitudes by direct mass streamfunction and down-
ward control mass streamfunctions by total and individual
FTR↑ F
NH↓ F
SH↓
Fig. 6. Tropical upward mass flux (left) and NH (middle) and SH (right) downward mass fluxes calculated from the direct massstreamfunction (black) and mass streamfunction based on the downward-control principle by total forcing terms (blue), EPD (red),GWD (orange), and residual (green) terms at 70 hPa during (a) DJF and (b) JJA. The sign of NH and SH downward mass fluxes isreversed for better comparison. Solid and stippled bars denote results in P1 and P2, and * mark denotes statistically significantchange at a 95% confidence level using t-test, respectively.
82 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
forcing terms and the mass fluxes calculated using turnaround
latitudes of direct and each forcing-induced mass stream-
fcuntions are shown in Table 1. As mentioned in Chun et al.
(2011), it is clear that the magnitude of the mass flux is
sensitive to the turnaround latitudes, especially for those
obtained by minor forcing terms such as the GWD and the
residual term. Among the differences in the turnaround
latitudes, those by the residual term in JJA during the P1
period are most significantly different from the turnaround
latitudes of the direct streamfunction, 79oN and 41oS vs. 48oN
and 26oS, and this causes the negative tropical upward mass
flux by the residual term shown in Fig. 6b. With the
turnaround latitudes of mass streamfunction by the residual
term, 79oN and 41oS in the NH and SH, downward mass fluxes
in the NH and SH are about 0.00 and 0.24 (× 109 kg s−1) in P1,
respectively. Using the turnaround latitudes by the residual
term in P2, the tropical upward mass flux is 0.48 × 109 kg s−1
in JJA, and consequently the change in the mass flux by the
residual term between P2 and P1 is about 0.24 × 109 kg s−1,
which is similar to that shown Fig. 6b. The mass fluxes by the
GWD are also different when they are calculated using the
turnaround latitudes of the downward control mass stream-
function, especially in the winter hemisphere. However, the
changes in the mass fluxes by individual forcing terms
between P2 and P1 are similar to that shown in Fig. 6.
Figure 7 shows the mass streamfunction averaged over DJF
for P1 and the difference between P2 and P1 (a) due to the
total EPD forcing, (c) due to the stationary and planetary-scale
EPD component, (e) due to the stationary and synoptic-scale
EPD component, (g) due to the transient and planetary-scale
EPD component, and (i) due to the transient and synoptic-scale
EPD components. The planetary-scale component is calculated
by the sum of the zonal wavenumbers 1-3 and the synoptic-
scale component is calculated by the sum of the wavenumbers
4-16. In the NH, the BDC in P1 induced by the EPD (Fig. 7a)
is explained primarily by the stationary and planetary-scale
EPD component (Fig. 7c) and the transient and planetary-scale
EPD component (Fig. 7g). In the SH, the BDC in P1 is
induced primarily by the transient and synoptic-scale EPD
components (Fig. 7i). The difference between P2 and P1 (right
panels), however, is explained primarily by the stationary and
planetary-scale EPD components at NH high latitudes (Fig.
7d) and the transient and planetary-scale EPD components at
SH high latitudes (Fig. 7h). By taking into consideration the
fact that the NH and the SH high latitudes are the two primary
regions of the zonal mean zonal wind change between P2 and
P1 (Fig. 3d), the changes in the mass streamfunction by each
component of the EPD may be related to the propagation
conditions of each of the wave components. In order to
examine this possibility, we calculate two parameters: the
critical level of the transient Rossby waves and the refractive
index of the stationary waves, which are described in sections
3c and 3d, respectively.
Figure 8 shows the EPD cospectra, indicating the time-space
structure of the wave forcing, with respect to the phase
velocity and latitude for the transient and planetary waves at
70 hPa for DJF in P1 (Fig. 8a), P2 (Fig. 8b), and the difference
between P2 and P1 (Fig. 8c). Superimposed on Figs. 8a, b are
the zonal mean zonal wind averaged for DJF (blue line) and
the ± 1 standard deviation of the daily zonal mean zonal wind
for the DJF mean (blue shading). The zonal wind profiles in
P1 (blue) and P2 (red) are shown in Fig. 8c, together with the
Table 1. Tropical upward mass flux and NH and SH downward mass fluxes in DJF and JJA calculated from the direct mass streamfunction(Direct), total downward-control streamfunction (DC-total), and downward-control streamfunctions by EPD (DC-EPD), GWD (DC-GWD), andresidual (DC-residual) terms at turnaround latitude (TL) of each term. MF* denotes mass flux calculated at turnaround latitude from direct massstreamfunction. Here, unit of TL and MF (and MF*) are degree and 10
9 kg s
−1, respectively.
DJF JJA
NH SH TR NH SH TR
P1 P2 P1 P2 P1 P2 P1 P2 P1 P2 P1 P2
DirectTL 26.9 25.6 46.9 48.1 48.1 46.9 25.6 24.4
MF* 4.9 5.2 1.3 1.4 6.2 6.6 0.5 0.6 3.4 3.6 3.9 4.2
DC-total
TL 26.9 25.6 46.9 48.1 48.1 46.9 25.6 24.4
MF 4.8 5.1 1.2 1.3 6.0 6.4 0.4 0.6 3.5 3.7 3.9 4.3
MF* 4.8 5.1 1.2 1.3 6.0 6.4 0.4 0.6 3.5 3.7 3.9 4.3
DC-EPD
TL 26.6 23.1 41.9 49.4 50.6 35.6 23.1 23.1
MF 4.5 4.6 0.9 1.0 5.4 5.6 0.4 0.4 3.5 3.4 3.9 3.8
MF* 4.4 4.5 0.9 1.0 5.3 5.5 0.4 0.4 3.4 3.4 3.8 3.8
DC-GWD
TL 36.9 36.9 46.9 46.9 46.9 46.9 46.9 46.9
MF 0.7 0.8 0.3 0.3 1.0 1.1 0.2 0.2 0.2 0.2 0.4 0.4
MF* 0.3 0.2 0.3 0.3 0.6 0.5 0.2 0.2 0.1 0.1 0.3 0.3
DC-residual
TL 10.6 11.9 78.1 43.1 79.4 75.6 40.6 11.9
MF 0.9 0.4 0.0 0.1 0.9 0.5 0.0 0.0 0.2 0.5 0.2 0.5
MF* 0.1 0.3 0.0 0.0 0.1 0.3 -0.2 0.0 0.0 0.3 -0.2 0.3
30 November 2014 Jong-Yoon Kim et al. 83
regions where the differences in the EPD are statistically
significant at a 95% confidence level (heavy gray shading) and
a 90% confidence level (light gray shading). Most of the
transient and planetary components of the EPD are located
along or on the negative side of the zonal wind profile [(c-U) <
0, where U is the zonal mean zonal wind], which could be seen
more clearly in the NH than in the SH, and this is indicative of
the existence of critical-level absorption (Randel and Held,
1991). There is a significant increase in the zonal mean zonal
wind, more than 3.5 m s−1, in the SH at mid- to high latitudes,
and the statistically significant strengthening of the transient
and planetary components of the EPD are found in the same
region (Fig. 8c). In the NH, the zonal mean zonal wind
increases slightly at 30-47oN and decreases at 47-80oN.
Although the transient and planetary components of the EPD
increase slightly at 30-45oN at a significance level of 95%, the
phase velocity changes are too small to be related to the
increase in the zonal mean zonal wind. The reduction in the
cospectra in P2 poleward of 60oN for a wide phase velocity
range on the negative side of the zonal wind profile is related
to the reduced zonal mean zonal wind in P2, although its
change is not statistically significant. Figures 8d-f show the
phase velocity-height cross sections at 70oS in P1 (Fig. 8d), P2
(Fig. 8e), and the difference between P2 and P1 (Fig. 8f).
Figure 8f shows that the increase in the zonal mean zonal wind
in P2 from 100 to 10 hPa shifts the critical levels of the
planetary waves upward and causes more planetary waves to
propagate in the stratosphere.
Figure 9 shows the same as in Fig. 8 except for the transient
and synoptic components of the EPD. Between 30oS and 60oS,
large values of the EPD cospectra by the transient and synoptic
components of the EPD exist on the positive and negative side
of the zonal wind profile and similar features can be seen in
the NH subtropics. Therefore, it is unclear whether or not the
critical level restricts the propagation of the synoptic-scale
waves. In addition, Fig. 9c, which shows the difference
between the two periods, indicates that there are no changes in
the transient and synoptic components of the EPD in the SH
mid- to high latitudes where the zonal mean zonal wind
increases. The statistically significant decrease near 20-40oS
and inverse near 60-70oN with phase velocity of 0-10 m s−1
during P2 are not likely related to the shift in the Rossby-wave
critical levels. Although the zonal mean zonal wind increases
in P2 at 30oN throughout the height from 100 to 10 hPa (Fig.
9f), the increase in the transient synoptic-scale waves in P2 due
to the increase in the zonal mean zonal wind is not evident,
except near 100 hPa with phase velocity ranging from 10 to 20
m s−1. In summary of Figs. 8 and 9, the strengthening of the
zonal mean zonal wind in the SH mid- to high latitudes leads
to an upward shift in the Rossby wave critical layers for the
transient and planetary waves. This change allows for more
transient and planetary waves to propagate into the stratosphere,
resulting in the enhancement of the BDC in the P2 period.
Figure 10 shows the refractive index of the stationary and
planetary waves of zonal wavenumbers 1, 2, and 3 in DJF for
P1 (contour) and P2 (shading). In the winter hemisphere (NH),
the refractive index is positive in wavenumbers 1 and 2, except
near the pole, largely due to the westerly zonal wind, and
according to Eq. (16), negative area is widen as the wave-
number increases. In the summer hemisphere (SH), the
refractive index is negative above 30 hPa where the zonal
mean zonal wind changes from westerly to easterly, and
negative area is widen as the wauationvenumber increases, as
in the NH. That is the reason why more stationary planetary
waves propagate into the stratosphere in the winter hemisphere
Fig. 7. Latitude-height cross sections of the mass streamfunctionaveraged over DJF for (left) P1 and (right) the difference between P2and P1 by the total EPD forcing (first row), by the stationary andplanetary-scale EPD component (second row), by the stationary andsynoptic-scale EPD component (third row), by the transient andplanetary-scale EPD component (fourth row), and by the transient andsynoptic-scale EPD component (fifth row). The contour intervals are 5kg m−2
s−1
and 1 kg m−2
s−1
for the left and right panels, respectively.The shaded regions in (b), (d), (f), and (h) represent statisticallysignificant anomalies at a 95% confidence level. The red lines denotethe turnaround latitudes in each hemisphere in P1.
84 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
and the EPD by stationary waves is stronger in the winter
hemisphere, as shown in Fig. 7c. The areas of the positive
refractive index for wavenumber 1 extend upward in the SH
stratosphere almost to all latitudes poleward of 15oS in P2, and
this allows for more stationary waves of wavenumber 1 to
propagate into the SH stratosphere, especially in the mid- to
high-latitudes. However, refractivity indices for wavenumbers
2 and 3 are nearly equal or somewhat reduced in P2 in the SH
mid- to high-latitudes in the upper troposphere and strato-
sphere. The changes in the refractivity indices of stationary
waves with different zonal wavenumbers between P2 and P1
are related to the changes in the vertical propagation of EP flux
by each wave components.
Figure 11 shows the EP flux vectors by stationary and
planetary waves of zonal wavenumbers 1 (first row), 2 (second
row), and 3 (third row) in the SH in DJF for (left) P1 and
(right) the difference between P2 and P1. The horizontal
component of EP flux vector is divided by 1000 to emphasize
the vertical propagation. Figures 11a, b show that vertical
component of EP flux for wavenumber 1, which has a primary
maximum near 60oS and a secondary maximum near 10oS and
100 hPa in P1, generally increases in P2 poleward of about
60oS at most altitudes, but decreases in P2 between 30oS-40oS.
The EP flux for wavenumber 2 (Figs. 11c, d) is maximum near
15oS with a secondary maximum at 60oS and 100 hPa in P1,
and the vertical component of EP flux increases in P2
poleward of about 60oS below 40 hPa. The EP flux for
wavenumber 3 (Figs. 11e, f) is maximum near 20oS with a
Fig. 8. The EPD cospectra with respect to the phase velocity and the latitude for the transient and planetary waves with zonalwavenumbers 1-3 at 70 hPa for DJF in (a) P1, (b) P2, (c) the difference between P2 and P1, and the (d), (e), (f) phase velocity-height cross section at 70oS from (a), (b), and (c). The contour interval is 0.01 m s−1 day−1 in (a), (b), (d), and (e) and 0.005 m s−1day
−1 in (c) and (f). The zero line and ± 0.01 line are omitted. Superimposed on (a), (b), (d), and (e) are the zonal mean zonal wind
(blue line) and ± 1 standard deviation of the daily zonal mean zonal wind from the mean (blue shading). The zonal wind profiles inP1 (blue) and P2 (red) are shown in (c) and (f), together with the regions where the differences in the EPD are statisticallysignificant at a 95% confidence level (heavy gray shading) and a 90% confidence level (light gray shading).
30 November 2014 Jong-Yoon Kim et al. 85
secondary maximum near 45oS in P1 at 100 hPa, and it de-
creases equatorward of about 60oS in P2, especially below 50
hPa. The results shown in Figs. 10, 11 demonstrate that the
changes in the propagation condition of individual stationary
waves between P2 and P1 are indeed related to the changes in
the vertical propagation of EP flux during the two periods.
Given that the stationary and planetary waves are defined in
the present study by the sum of the three stationary waves of
wavenumers 1-3, changes in the EPD forcing by the stationary
and planetary waves in the P2 period are various with altitude
and latitude, which does not lead to strengthen the mass
streamfunction in the SH mid- to high-latitudes in DJF, as
shown in Figs. 7c, d.
In the present study, we investigate the change in the pro-
pagation conditions by analyzing the Rossby wave critical
layer and refractive index. In summary, the strengthening of
the zonal mean zonal wind in the SH mid- to high latitudes by
ozone depletion in the SH polar region causes (i) an upward
shift of the Rossby-wave critical layers, which allows for more
transient planetary waves to propagate into the stratosphere,
and (ii) an increase in the refractive index in the SH strato-
sphere, which leads to more stationary planetary wave pro-
pagation into the stratosphere in P2. The results noted above
correspond well with the outcomes using the NCEP/DOE
reanalysis data such as trends of the upward shift of the
Rossby-wave critical layers and increase in the refractive index
in SH stratosphere (not shown).
c. Wave activity in the upper troposphere
The BDC can be strengthened by increasing the wave
activity in the upper troposphere without a change in strato-
spheric propagation conditions. In order to investigate the
wave activity in the troposphere, we examine each component
of the EP flux, given by the Eqs. (11)-(14). Figure 12 shows
the EPD in P1, P2, and the difference between the two periods
in DJF by the total EPD, by the F φ_1 component, by the FZ_1
component, and by the FZ_2 component. The Fφ_2 component in
Fig. 9. The same as in Fig. 8 except for the transient and synoptic-scale waves with zonal wavenumbers 4-16.
86 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
(12) is not represented, because it is of negligible magnitude.
Figure 12 shows that the total EPD in P1 (Fig. 12a) and P2
(Fig. 12b) can be explained by the F Z_1 component (Figs. 12g,
h) and F φ_1 component (Figs. 12d, e), and that the F Z_2 com-
ponent has a similar magnitude only in the lower stratosphere.
The difference between the two periods (Fig. 12c) indicates an
increase in the wave forcing in the stratosphere, which is in
accord with the difference in the streamfunction induced by
the EPD as represented in Fig. 4d. The Fφ_1 component is
intensified in the NH high latitudes and weakened in the NH
mid-latitudes. This is likely due to the fact that the SSWs occur
more frequently in the P2 and the weakened polar vortex
associated with the SSWs allows more waves to propagate into
the high latitudes. It is not straightforward, however, to
confirm this causality, given that the SSWs depend strongly on
the planetary waves propagating into the stratosphere and that
there might be a two-way feedback between the SSWs and
planetary waves. The difference in the F Z_1 component (Fig.
12i) indicates an overall increase in the wave forcing in the
stratosphere. We note that the F Z_1 component is the primary
forcing for the total EPD change from 250 hPa to 10 hPa (Fig.
Fig. 10. Latitude-height cross sections of the refractive index ofstationary and planetary-scale waves of zonal wavenumber (a) 1, (b) 2,and (c) 3 in DJF in P1 (contour) and P2 (shading).
Fig. 11. Latitude-height cross sections of the EP flux vectors ofstationary and planetary-scale waves of zonal wavenumber (first row)1, (second row) 2, and (third row) 3 in the SH in DJF in (left) P1 and(right) the difference between P2 and P1. The horizontal component ofthe EP flux is divided by 1000 to emphasize the vertical propagation.
30 November 2014 Jong-Yoon Kim et al. 87
12c), especially in the SH mid- to high latitudes. In order to
examine whether the increase in the F Z_1 component of the EP
flux in P2 is related to the increase in the wave activity in the
upper troposphere, further analysis is required.
Figure 13 shows the annual cycle of F Z_1 in P1, P2, and the
difference between the two periods at 70 hPa and 250 hPa. In
Fig. 12. Latitude-height cross sections of the EPD in (left) P1, (middle) P2, and (right) the difference between P2 and P1 by the(first row) total EPD forcing, (second row) F φ_1
component, (third row) Fz_1
component, and (fourth row) Fz_2
component. Contourintervals are ± 0.2, ± 0.5, ± 1, ± 2, ± 5, and ± 10m s
−1day
−1 in the first and second columns and ± 0.04, ± 0.1, ± 0.25, ± 0.5, ± 1, and
± 2 m s−1 day−1 in the third column. Red and blue denote positive and negative, respectively. The regions where the differences arestatistically significant at a 95% confidence level are shaded.
88 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
this study, we focus on the changes in the mid-latitudes,
because the EP flux difference between the two periods at
turnaround latitude is important. Although significant change
is shown in some months in the tropical stratosphere, the
changes are largely influenced by quasi-biennial oscillation,
which is beyond scope of the current study. During P1, the
maximum value occurs primarily in the wintertime in the NH
mid-latitudes at both levels. In the SH mid-latitudes, however,
the maximum and minimum value at 70 hPa are found in SON
and DJF, respectively, while the maximum and minimum
value at 250 hPa are found in March-April and SON, re-
spectively. This implies that, in the SH, F Z_1 propagated into
the stratosphere is influenced more by the propagation
conditions than the wave activity in the upper troposphere.
Differences between the two periods at 70 hPa (Fig. 13c)
indicate that there is a statistically significant increase in
January-February-March (JFM) in the SH between 50-70oS,
September in the SH centered at 50oS, and in most months
except in JFM in the NH subtropics. Among those regions
with a statistically significant increase in F Z_1 at 70 hPa, a
statistically significant increase at 250 hPa is shown in
February-March in the SH between 50oS and 70oS and
September near 50oS, but there is no evident increase in
January, relatively. This implies that the increase in the EPD
near 50oS in September is related to the increase in the wave
activity in the upper troposphere (250 hPa). In order to examine
this hypothesis, we investigate the heat flux composing F Z_1 at
50°S.
According to the Eq. (13), F Z_1 is determined by the heat
flux. Figure 14 shows the longitudinal distribution of the heat
flux at 50oS in P1, P2, and the difference between the two
periods in September and January. In September, the lon-
gitudinal distributions of the heat flux at 70 hPa (solid line)
and 250 hPa (dashed line) are generally similar to each other.
The amplitude at 70 hPa is much larger in P2, especially
eastwards of 60oW, except for a few longitudes (near 90oW-
10oW in the P1 period and 60oE-150oE in the P2 period). The
increase in the heat flux during P2 is also observed at 250 hPa,
with a longitudinal structure similar to at 70 hPa, except near
180-120oW. These enhancements in the heat flux in P2 at both
the 70 hPa and 250 hPa levels are shown clearly in Fig. 14c. In
January, when the change of F Z_1 is different between the 70
hPa and 250 hPa, the longitudinal distributions of the heat flux
at 70 hPa and 250 hPa are significantly different from each
other and from those in September. The amplitudes are much
smaller at both levels, as expected in the summer hemisphere,
Fig. 13. Annual cycle in the EP flux Z_1 component in (first column) P1, (second column) P2, and (third column) the differencebetween the two periods: (upper) 70 hPa; (lower) 250 hPa, respectively. Contour intervals are 2.5 × 10
4m
2s−2
in (a) and (b),0.5 × 104 m2 s−2 in (c), 12.5 × 104 m2 s−2 in (d) and (e), and 2.5 × 104 m2 s−2 in (f). The regions where the differences are statisticallysignificant at a 95% and 90% confidence level are shaded with heavy and light gray shadings, respectively.
30 November 2014 Jong-Yoon Kim et al. 89
and the longitudinal distribution of the heat flux at 70 hPa and
250 hPa is primarily out-of-phase. In addition, there is no clear
enhancement in the heat flux during P2, except near 180oW-
130oW at both levels. Figure 14 demonstrates that the en-
hancement of F Z_1 in P2 at 50oS in September (Fig. 13) is due
to the increase in the heat flux at 250 hPa at the same latitude
without experiencing a significant filtering process between
250-70 hPa, while that in January is likely related to the
changes in the wave propagation conditions.
In summary, Figs. 8-11 demonstrate that the change in the
Fig. 14. Longitudinal distribution of the heat flux (K m s−1) at 50oS in (first column) P1, (second column) P2, and (third column) thedifference between the two periods in (upper) September and (lower) January. The solid and dashed lines denote the values at 70hPa and 250 hPa, respectively.
Fig. 15. Schematic diagram of the dynamical processes of the BDC changes in DJF.
90 ASIA-PACIFIC JOURNAL OF ATMOSPHERIC SCIENCES
propagation conditions, which is represented by the critical
layer of transient waves and the refractive index of stationary
waves, leads to the EPD enhancement in SH mid-to high
latitudes in DJF, while the EPD enhancement in the SH mid-
latitudes in September is likely due to the enhanced wave
activity in the upper troposphere. However, it is not straight-
forward to distinguish between wave activity and propagation
conditions precisely, because there is a close interaction bet-
ween waves and the mean flow. In addition, in the present
study wave propagation is considered based on the mono-
chromatic wave theory of which the phase speed remains
during the propagation. However, when we consider Rossby
wave packets, the phase speed can be changed following rays
if the background wind and stability change horizontally and
timely (Andrewes et al., 1987). Such a complicated situation
could not be considered in the present study.
5. Summary and conclusions
In this study, changes in the BDC during the 30 years from
January 1980 to December 2009 are investigated using the
MERRA reanalysis data. In the present study, the BDC is
represented by the mean meridional circulation, although it
includes quasi-horizontal two-way mixing processes (Shepherd,
2002). The changes in the BDC are analyzed using two
aspects: the wave propagation conditions in the stratosphere
and the wave activity in the upper troposphere. They are then
compared in the first (P1) and the second (P2) 15 years in
which the stratospheric temperature shows positive and
negative deviation, respectively.
The wave forcing determining the strengthening of the BDC
is stronger in the NH than in the SH during DJF, because the
wave forcing is generally stronger in the winter hemisphere
(Andrew et al., 1987). However, the magnitude of the BDC
change in DJF between the two periods (P2-P1) is similar in
both of the hemispheres from 100hPa to 10hPa but the reasons
for the changes are different. Figure 15 shows a schematic
diagram of the dynamical processes of the BDC changes in
DJF. In the SH, the increased zonal mean zonal wind in the
mid- to high latitudes, caused by ozone depletion in the polar
region, leads to an upward shift of the Rossby-wave critical
layer and this allows more transient planetary waves to
propagate into the stratosphere. This causes the increase in
mean meridional circulation and mass flux in the stratosphere,
because the changes in the wave forcing occur at the turn-
around latitudes. In the NH, there is a statistically significant
increase in the EPD at high latitudes, and this results in the
more frequent occurrence of the SSWs in P2. Although there is
a decrease in the EPD in the NH mid-latitudes, this is not
statistically significant. The GWD, which is provided ex-
clusively by the MERRA reanalysis data, is much smaller than
the EPD and its change between the two periods is
insignificant. On the other hand, the residual term in the TEM
equation is similar to the GWD in the NH and is much larger
than the GWD in the SH. In addition, the change in the
residual term between the two periods is as large as that in the
EPD in both hemispheres. The changes in the residual term
strengthen mean meridional circulation and increase the mass
flux, because the change occurs at turnaround latitudes.
It is noteworthy that the present results related to planetary
waves and their divergence with high-order derivatives may
not be the same as those using different reanalysis data sets,
although the zonal-mean structure of the wind and temperature
of each reanalysis data set may be similar. When we calculate
the mass stream function and the wave forcing terms using the
NCEP/DOE reanalysis data (not shown), the results are
generally consistent with those in the present study, except that
the mass streamfunction calculated by the synoptic and
transient components of the EPD are slightly larger in the SH
at high latitudes. There are several previous studies, using
other reanalysis or observation data set, that are somewhat
different from the present result. For instance, Seviour et al.
(2012), using the ERA-Interim reanalysis data from 1989 to
2009, found a negative trend in upwelling mass flux at 70 hPa,
although with somewhat inconsistent negative temperature
trend at 70 hPa, which is induced mainly by adiabatic process
of the positive trend in upwelling. Diallo et al. (2012), using
the ERA-Interim reanalysis data from 1989 to 2010, found a
significant negative trend of the age of air in the lower strato-
sphere and insignificant positive trend in the mid stratosphere,
implying the strengthening of the BDC only in the lower
stratosphere. Futhermore, Engel et al. (2008), using balloon-
borne measurements of stratospheric trace gases from 1975 to
2005, suggested that there is no trend in the age of air in the
stratosphere. Thus, examining the robustness of the current
results using various reanalysis data sets remains to future study.
The changes in the wave activity in the upper troposphere
are investigated by analyzing the first vertical component of
the EP flux FZ_1 that plays an important role in the BDC
change. In September, the increase in F Z_1 at 70 hPa, 50oS is
related to the increase in F Z_1 at 250 hPa. In January, however,
the increase in F Z_1 at 70 hPa and 50oS is not directly related to
the increase in the wave activity at 250 hPa, but instead it is
related to the wave propagation conditions.
Some of the aforementioned results related to the GWD are
different from those in the previous studies using climate
models or reanalysis data. According to Butchart et al. (2010),
using the 11 CCMs, the influence of the GWD on the BDC
trend is larger than that of the EPD and on average it
accounted for 59% of the trend in the annual mean upwelling.
However, the level of contribution is significantly different
between each model. In the GEOS-5, the base model of the
MERRA reanalysis data, the GWD accounts for less than 20%
of the trend (Butchart et al., 2010). Recent climatological
simulations using the Whole Atmosphere Community Climate
Model (WACCM) including three gravity wave drag (GWD)
parameterizations (orographic, non-stationary background, and
convective GWD parameterizations) show that the GWD’s
contribution to the tropical upwelling is 19% at 70 hPa (Chun
et al., 2011). Therefore, the uncertainty of the GWD’s
30 November 2014 Jong-Yoon Kim et al. 91
contribution to the BDC is very high depending on how well
the GWD parameterization scheme represents the reality.
In most previous studies on the BDC using reanalysis data,
the GWD forcing has been estimated from the residual term in
the TEM equation and this may contain considerable un-
certainties in the contribution of the GWD forcing on the BDC
and its trend. The residual term includes small scale
turbulences, the GWD that cannot be presented by the
parameterization scheme, and the imbalance caused by the
incremental analysis. By investigating the contribution of the
GWD to the BDC using the current MERRA reanalysis data
that provides the GWD forcing output (sum of the orographic
GWD and the non-orographic GWD) explicitly, we first are
able to calculate the GWD forcing in the TEM equation and its
contribution to the BDC, separately from those by the residual
term. The contribution of the GWD forcing to the BDC is
much smaller than that by the residual term, especially during
the P2 period. However, this does not directly implicate the
small contribution of the GWD on the BDC, given that the
residual term may have included the GWD forcing that cannot
be represented well by the GWD parameterization schemes in
the global model. In order to reduce such uncertainties, the
development of more realistic GWD parameterization schemes
for climate models with integrated efforts on the observations,
theories, and numerical simulations is required (Alexander et
al., 2010), especially taking the GWD sources into con-
sideration explicitly (Chun et al., 2011).
Acknowledgments. We would like to thank Young-Ha Kim,
and Wook Jang for their help. HYC was funded by the Korean
Meteorological Administration Research and Development
Program under Grant CATER_2012-3054.
Edited by: Hong, Kim and Yeh
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