Download - Changes in the Brewer-Dobson circulation for 1980–2009 revealed in MERRA reanalysis data

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: [email protected]

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.


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( )–=


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.


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.


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.


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.


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.


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


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.


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).


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.


(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.


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.


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.


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.


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


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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