February

1992

T. Yasunari

Roleof theAsian

Monsoonof theon

GlobalBy Tetsuzo

andY. Seki177

theInterannual

ClimateSystemYasunari

and

Yuji

Variability

Seki1

Institute of Geoscience, University of Tsukuba, Tsukuba, Ibaraki 305, Japan(Manuscript received9 September 1991, in revised form 3 December 1991)

Abstract

The role of the Asian summer monsoon on the interannual variability of the global climate system
particularly relevant to the ENSO time scales is discussed, by examining the statistical and dynamical links between the Asian summer monsoon, the atmosphere/ocean system in the tropics and thewesterly flow regimes in the extratropics.
The Asian monsoon, the ocean and the atmosphere in the tropical Pacific are tightly linked together
as one climate system, named here as the MAOS (Monsoon and the Coupled Atmosphere/Ocean
System). The MAOS prominently shows the biennial oscillatory nature which tends to have anomalousstates starting in the northern summer monsoon season and persisting for about one year (Yasunari,
1990a: 1991).
The anomalous state of the MAOS produces the anomalous atmospheric circulation over the subtropics and the extratropics of the north Pacific during summer through the early winter, throughthe modulation of the subtropical high and the stationary Rossby wave propagation mechanism. In
the mid winter, this anomalous circulation over the north Pacific is evolved to the hemispheric winteranomalous circulation with wavenumber-one and/or-two structure.

The anomalous circulation over Eurasia associated with this hemispheric anomalous flow regime
seems to provide a favorable condition for the extensive (or diminished) snow cover area over central
Asia, which in turn is responsible for the reversed anomalous state of the next Asian summer monsoon
and the MAOS. That is, the biennial nature of the climate system in the northern hemisphere may
be due, at least partly, to this two-way interactions between the tropics and the extratropics. In theseprocesses, the Asian monsoon plays a key role as a transmitter of climate signals between the tropics
and the extratropics through the land/atmosphere/ocean interaction in the seasonal cycle.
In addition, it is strongly suggested that the North Atlantic Oscillation (NAO), in reality, plays
a crucial role in the timing of the occurrence of the ENSO event, by stochastically amplifying or
damping the biennial oscillation of this coupled climate system. That is, the more or less irregular
ENSO cycle may result from this interaction between the MAOS and the NAO, where the former
seems to have the nature of an almost-intransitive climate system, while the latter seems to represent
the more chaotic nature of the westerly flow regime.

1. Introduction
The Asian summer and winter monsoon form huge
circulation systems in the general circulation of the
global atmosphere (Krishnamurti, 1971; 1973). Simply speaking, this system is characterized as a zonal
asymmetry as well as a meridional asymmetry inthe tropics. Figure 1 shows the monthly mean sea
level pressure and surface wind vectors (a) and themonthly mean velocity potential and divergent windvector at 200mb (b) in July 1990 (JMA, 1990).
This figure apparently shows a predominant systemwith the zonally-oriented east-west (or Walker) circulation over south Asia through the southern sub1Present affiliation:
The Weather News (Co . Ltd).
(C)1992, Meteorological Society of Japan

tropical Pacific, as well as the meridionally oriented
monsoon circulation over the south Asia throughthe Indian Ocean. This remarkable zonal asymmetry of the circulation has proved to be maintained and modulated through the interaction between the Asian monsoon and the coupled atmosphere/ocean system in the equatorial Pacific (Yasunari, 1990a). This interaction seems to play a
key role on the mechanism of the El Nino/Southern
Oscillation (ENSO), as has been discussed by Barnett (1985, 1988) and Yasunari (1987). That is, aweaker (stronger) than normal Asian summer monsoon is very favorable for triggering the El Nino
(anti-El Nino or La Nina (Philander, 1985))state of
the equatorial Pacific through the weaker (stronger)
east-west circulation in the tropics.
47

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Journal

Fig.

1.
(b)

Monthly
for

July

mean

sea

1990.

Contourlevelof thepressureintervalin

Meteorological

(a)

and

(b)

is 1.06×10m2/s.

velocity

In addition, Yasunari (1990b) and Kodera andChiba (1989) pointed out a possible link of theEurasian snow cover the preceding spring to theanomalous state of the tropical and subtropical Pacific, presumably through the anomalous state ofheating contrast between the continent and theoceans. The observational evidence of the association between the Eurasian snow cover extent andthe succeeding summer monsoon shown firstly byHahn and Shukla (1976) was confirmed by the general circulation model (GCM) studies by Barnett
et al. (1989) and Yasunari et al. (1991). These results strongly suggest that the extra-tropical forcing
through the Asian summer monsoon should be incorporated into the physical process of the ENSO.
On the other hand, recent numerous studies haveshown that the anomalous state of the equatorial
Pacific, either El Nino or La Nina, greatly influences the anomalous state of the circulation in theextra-tropics through stationary Rossby wave propagations (e. g., Hoskins and Karoly, 1981; Webster,
1981; Horel and Wallace, 1981; Palmer and Mansfield, 1984). These results have even brought the48

Societypotential(JMA,of JapananddivergentVol. 70, No.

windvector

at

1

200mb

1990)

idea that the potential predictabilityof the longrange forecasting of the extra-tropicalclimate owes
its major part to the anomalous state of the tropicalclimate system.
How, then, may the ENSO cycle, or more generally, the interannual variability of the global climate
system, occur through the interaction between thetropics and the extra-tropics?
This paper attempts
to solve this problem, by reviewing and re-examining
each process of the interactions in more detail. Particularly, the role of the Asian summer monsoon will
be focused on as a transmitter of climatic signals between the tropics and the extra-tropics
in the seasonal cycle.

2. Asian monsoon and the coupledmosphere system in the tropicalocean/atPacificSince the early part of this century (Walker andBliss, 1932) the relationship between the Asian summer monsoon and the ENSO has been manifested
by numerous studies, showing that the occurrenceof the ENSO is closely associated with the preceding weak Asian summer monsoon as shown in Fig. 2.

February

1992

T. Yasunari

and

179

Y. SekiFig. 2. Interannual variation of all-India monsoon rainfall. The ENSO years are indicated with black bars.
(Mooley and Shukla, 1987)

Fig. 3. Time series of Indian monsoon rainfall anomaly (thick solid line) and sea water temperature anomaly at 20m (thickdashed line) and 100m (thin dashed line)
depth averaged for 137E line (2N-10N) in
the succeeding January. (Yasunari, 1990a)

However,Meehl (1987) and Yasunari (1990a) notedthat the Indian monsoon activity seems to play anactive rather than a passive role in determining theanomalous state of the warm water pool in the western Pacific through the anomalous east-west circulation in the followingautumn and winter seasons,as shown in Fig. 3. These studies further substantiated the tight link between the Asian summer monsoon fluctuation and the anomalous state of the coupled ocean/atmosphere system in the tropical Pacific, and suggested that the Asian summer monsoon and the coupled system in the tropical Pacific should be understood as a climate system combined together in the tropics. Hereafter, we refer tothis climate system by the abbreviation MAOS (i. e.,
Monsoon/Coupled Atmosphere Ocean System).
A prominent feature of the MAOS is its oscillatory nature on the quasi-biennial time scale. In addition, an anomalous state of this MAOS shows astrong phase lock to the seasonal cycle, as is apparently seen in Fig. 4. This strong seasonality of
the anomalous state of the MAOS has prompted us
to introduce a concept of the "monsoon year" (Yasunari, 1991) as a unit climatic year in the tropics.
In this context, an ENSO event may be understoodas a considerably amplified phase of the oscillationin the MAOS.

Fig. 4. Lag-correlations between the Indian
monsoon rainfall anomaly and the sea surface temperature anomaly in the western(0-8N, 130E-150W) and the eastern (08N, 170W-150W) Pacific. The reference
monsoon season is shown with a thickblack bar. Y(0) denotes the year of reference monsoon year and Y(-1) (Y(+1))
denotes the year before (after) Y(0). (Yasunari, 1990a)

3. InfluenceMAOSof theextra-tropics

on

the

The Asian summer monsoonis forced primarily by
the differential heating between the Eurasian continent and the surrounding oceans, though the moist
process in the atmosphere is essential for maintenance of the gigantic monsoon circulation. However,this monsoon circulation is supposed to be very sensitive to the surface heating condition (Charney andShukla, 1981). Blanford (1884) first noted this aspect, by examining the relationship between the Himalayan snow cover in winter and the followingIndian summer monsoon. Hahn and Shukla (1976),
Dickson (1984), Dey and Bhanu Kumar (1984)
and others have re-examined this Eurasian snow
cover-Indian monsoon connection by using the dataof satellite-derived snow cover extent. Morinaga(1992) pointed out that the snow cover over central
Asia near the Caspian Sea in late winter throughspring shows significant negative correlation with
49

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70, No.

1

Fig. 5. Time series of the snow cover extent anomaly over central Asia in April(Kodera and Chiba, 1989) and the sea water temperature anomaly at 125m depth
in the western tropical Pacific in the following Janurary for the period of 1967 to1988.

the Indian monsoon rainfall (IMR) (Parthasarathy,1987).
The anomalous state of the surface conditions(e. g., snow cover, soil moisture) of the continent
in winter and spring, therefore, should affect theanomalous state of the MAOS for each monsoon
year. The apparent time-lag correlation between thesnow cover extent in central Asia in April (Kodera
and Chiba, 1989) and the oceanic mixed layer temperature of the tropical western Pacific in the succeeding January as shown in Fig. 5, strongly supports this idea. Undoubtedly, this time-lag correlation should physically be based upon the anomalousstate of Asian monsoon and the associated coupled
ocean/atmosphere system in the tropical Pacific persisting from summer through winter, which may be
induced by the anomalous snow cover over the continent.
There seem to be two physical processes responsible for the time-lag correlation between the wintersnow cover and the summer monsoon. One is thealbedo effect and another is the snow-hydrologicaleffect, as shown schematically in Fig. 6. The latter
effect involves the melting of an anomalous snow
mass, moistening of the soil and an anomalous
evaporation, which substantially reduces the heating of the atmosphere from the ground surface inthe warmer seasons. The importance of this process was noted by Yeh et al. (1983), and was further substantiated in the GCM experiments by Yamazaki (1989), Barnett et al. (1989) and Yasunari
et al. (1991). These recent GCM studies also notedthat the combined albedo/snow hydrological effect50

Fig. 6. Schematic diagram for the albedo feedback and the hydrological feedback of snow
cover during the seasonal march from winter to summer. (Yasunari et al., 1991)

is actually important for this "time-lagged" effect.
On the other hand, Morinaga and Yasunari (1987)
deduced the atmospheric circulation pattern in winter responsible for the maximum (minimum) snow
cover extent over central Asia. Figure 7 shows thelag-correlation pattern between the snow cover extent over central Asia in February and the 500mb
height anomaly of the northern hemisphere in thepreceding December. This figure indicates that theheavy snowfall (and the large snow cover extent)over central Asia is associated with the anomalous
deep trough over there combined with the anomalous ridges over Europe and east Asia. This pattern
may be identified as a sort of "Eurasian (EU)" pattern (Wallace and Gutzler, 1981). That is to say,the anomalous winter circulation pattern over theEurasian continent shown here seems to be a precursor signal affecting the anomalous state of theMAOS in the following monsoon year.

4. Subtropical and extratropicalthe MAOS

responses toSince Bjerknes (1969), it has been noted that
the El Nino event, or more exactly, the anomalous state of the atmosphere over the equatorial
central/eastern Pacific associated with the sea surface temperature (SST) anomaly there, produces
the anomalous circulation over the north Pacificthrough the north American sector. Horel and Wallace (1981) demonstrated the characteristic circula-February

1992

T. YasunariFig. 7. Lag-correlation

map of theandY. Seki

181

500mb

geopotential height anomaly in December and the snow cover extend anomaly,
over central Asia in the following February. (Morinaga and Yasunari, 1987)

tion patterns (e.g., PNA, NAO, WP, Eu etc.) associated with the ENSO events. Hoskins and Karoly(1981) and Webster (1981) gave a theoretical basis for these observations as a forced Rossby wave
propagation from the tropics. Many GCM studies(e.g., Blackmon et al., 1976 etc.) have verified theextra-tropical responses to the SST anomalies in theequatorial Pacific with the El Nino condition.
However,the extra-tropical responses to the SST
anomalies in the equatorial Pacific should be examined more generally with reference to each anomalous state of the MAOS, since the extra-tropical response is not only limited to the El Nino conditions.
Palmer and Mansfield (1984), for example, showedthat the atmospheric response over the north Pacific is as strong as, and even stronger during the La
Nina conditions (namely, higher SST anomalies in
the western Pacific) than the El Nino conditions.
Figure 8 shows the monthly wind stress anomaly
in the northern Pacific subtropics and mid-latitudes,
composited for weak summer monsoon years minus strong summer monsoon years, to show moreclearly the features for the weak monsoon year. It
is apparent that cyclonic circulation anomalies are
dominant
170°W‑140°W)

in

the

northeastern
from

July

Pacific
through(40°N‑50°N,
October,indi‑

Gating the weaker subtropical high during summer
through the following autumn. This result is wellcorrelated with the westerly anomalies along theequatorial belt, which implies a weaker state of theMAOS with the weaker east-west circulation (i. e.,
weaker trade wind and upper tropospheric wind sys-

Fig. 8. Anomaly surface wind stress vectors
composited for July through December for
weak Indian monsoon years minus strong
Indian monsoon years. (Yasunari, 1990b)tem) in the tropics. The persistent cyclonic circulation anomaly in the same region in November
and December, on the other hand, corresponds to astronger than normal, or more equatorward shift of
the Aleutian low.
To deduce the dominant regional circulation patterns in the northern mid-latitudes associated withthe anomalous state of the MAOS, the time coefficients of some localized dominant circulation (orteleconnection) patterns are correlated to the Indian summer monsoon index. These patterns are obtained by applying the varimax-rotated empirical or51

182

JournalFig.

9. Examples
mb

Fig.

geopotentialof dominantheight

10. Lag-correlations
teleconnection

patternsof theteleconnection
anomalies

betweenshown

Meteorologicalpatterns

for 45 years

the Indianfrom

monsoonin Fig. 9. Correlations

thogonal functions (EOFs) analysis to the monthly
500mb height anomaly data for the period of 1946
to 1990 (Yasunari and Ueno, 1992). Figure 9 shows
the spatial patterns of six dominant circulation patterns. This objective method could successfully deduce most of the well-known teleconnection patterns
such as NAO, PNA, WP, EU (Wallace and Gutzler,
52

Society

deducedVol.

of Japanby varimax

rotatedEOFs

of monthly

70, No.

500

1946

to 1990.

rainfallanomaly

and time

coefficientsof the principal

valuesless than

0.2 areshown.

with

not

1

1981). Figure 10 shows the time sequence of lag correlation between the Indian monsoon rainfall index
and the time coefficients of some dominant teleconnection patterns for each month before through after the reference monsoon year. It is noteworthy tostate that the significant correlations (above the 95% level) related to the Indian summer monsoon (orFebruary

1992

T. Yasunari

andY. Seki183

Fig. 12. Same as Fig. 11 except for Decemberof (a) strong monsoon year and (b) weak
monsoon year.

with the pattern for the snow cover anomaly overEurasia, as shown in Fig. 7.
To find out the hemispheric anomalous circulation patterns associated with the anomalous stateof the MAOS, the monthly 500mb height anomalies are composited for each month during and after the weak and strong summer monsoon years,
respectively, as shown in Fig. 11. As is readily deduced from this figure, in the weak monsoon years(Fig. 11 right), for example, the PNA pattern (ofnegative height anomaly over the north Pacific andthe northeastern part of north America with positive anomaly over the Alaska/Rockies) is significant
in August through October, implying an intensified
Aleutian low and the meandering flow on the downstream side over the north American sector. In thestrong monsoon years, in contrast, the reversed PNA

Fig. 11. Composite
anomaliesof 500mb
geopotential
height from August to
November of strong (left) and weak (right)
Indian monsoon years. Contours are 10
gpm and negative values are shown withdashed lines.

the MAOS) persistently appear during and after thesummer monsoon seasons (i. e., autumn and winter)in the area of the north Pacific and the north American sector, as shown in the series of lag-correlations
for the PNA, WP and PNA-II patterns. This strong
seasonality of the persistent correlations seems to
be consistent with the concept of the corresponding
monsoon year in the topics. Other significant correlation peaks of the EU, PNA and PNA-II patterns
appear in the previous autumn through early winter of Y(-1), which seems to be closely associatedpattern is dominant in the same period, implyingthe
weakened Aleutian low with more zonal flow overthis region. These (reversed) PNA patterns in thisseason with the intensified (or weakened) Aleutian
low may not necessarily be due to the Rossby wave
propagation mechanism, but may be more directly
related to the weakened (or intensified) north Pacific high in the subtropics, as shown in Fig. 8. This
weakened (or intensified) high may facilitate (or obstruct) the southward intrusion of the westerly flow
over the north Pacific.
5. Feedback processes
to the MAOS

from the extratropics

Interestingly, the anomalous circulation with the(reversed) PNA pattern in late autumn seems to develop into the anomalous planetary-scale flow regime
with wavenumber one plus two through the courseof the seasonal evolution of the polar air mass inearly winter (November and December), as shown in
Figs. 11 and 12. That is, in the weak monsoon year,
the PNA pattern over the north American sector
tends to develop the strong seasonal trough over thenortheast north America (and that over east Asia)
while more zonal flow develops over the Eurasiancontinent. In the strong monsoon years, by con53

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JournalFig.

13. Schematic
betweenthe

Rossby-wave

diagramMAOSof thein thepropagationinterannual

tropics,
andof theandland-surface

Meteorologicalvariation
theof theflow

regimeprocesses

overEurasian

owe its mechanismat least partly to the two-way interactions between the tropics and the extra-tropics
in different seasons of the year, in which a negative
feedback exists in the propagation of climatic signals.
The interactions of the climate system between
the MAOSin the tropics and the extra-tropics, andbetween the seasons with a biennial oscillatory nature are summarized schematically in Fig. 13. One
persistent anomaly in the MAOS from summer towinter seems to be responsible for the corresponding winter anomalyin the westerlyflowregime in thenorthern extra-tropics, which provides the anomalous land-surface condition (i. e., snow cover andpossiblysoil moisture) over the Eurasian continent.
This anomalous land-surface condition, in turn, produces the opposite anomaly of the followingnorthern summer monsoon (and the MAOS) to the previous year.

of Japanclimatewesterly

trast, more zonal flow develops over the north Pacificthrough the north American sector, while an anomalous trough develops over the central part of theEurasian continent. These anomalous flow regimes
in winter may be due partly to the linear stationary
Rossby-wave response to the persistent forcing of
the anomalous MAOS, but also to the non-linear interaction between the changing basic-state flow andtransient eddies through the seasonal march.
Now, it should be noted that this anomalous pattern in early winter (Fig. 12 left) is nearly identicalto the pattern favorable for a large snow cover extent
over central Asia, as shown in Fig. 7. Namely, thestrong (or weak) summer monsoon condition tends
to produce the weak (or strong) summer monsoon
condition for the next year, through the anomalous
westerly flow regime and the associated anomalous
snow cover condition in the following winter. In
other words, the biennial nature of the MAOS may

54

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interactions, by meansof theseasonal
cycle .

6. Role of the NAO on the MAOS and the
ENSO
In reality, however, the year-to-year variabilitiesof the MAOS as well as the westerly flow regimes
do not necessarily show the biennial nature exactly,
as is seen, e.g., in Fig. 5. The ENSO event (or the
ENSO cycle), in fact, shows a preferred periodicityof 4 to 6 years, rather than 2 to 3 years (Rasmusson
and Carpenter, 1982 etc.), though it seems to basically contain the nature of the biennial oscillation.
We will further examine this problem.
Here, to focus on the difference of the circulation patterns between the ENSO years and thenon-ENSO years among the weak summer monsoon
years, the time coefficientsof some dominant patterns (Fig. 9) are composited for three years centered by the weak monsoon year (Y(0)) with the
ENSO event (a) and without the ENSO event (b)(Yasunari, 1988), as is reproduced here in Fig. 14.
The SOI is also composited in the same manner andincluded in this diagram. It should be noted that the
SOI apparently shows a biennial cycle in the periodof Y(-1) through the beginning of Y(+1) particularly in the case of the ENSO. Associated with thisfeature, a remarkable difference is also seen in the
SOI tendency from the winter of Y(-1)/Y(0) to thesummer of Y(0) between the two cases.
To examine the teleconnection patterns associatedwith this remarkable feature, we focus on the difference of the dominant patterns appearing in the winter of Y(-1)/Y(0). The negative phase of PNAII pattern, with negative height anomaly over thenortheast north America and positive anomaly overthe northeast Pacific seems to be a common feature in both cases. This may presumably be related to the strong or non-weak summer monsoonof Y(-1), as is suggested in Figs. 11 and 12. A
remarkable difference between the two cases is thephase or polarity of the NAO pattern, as noticed

February

1992

T. Yasunari

andY. Seki185

Fig. 15. Composite anomalies for 500mb
height (north of 20N) for the precedinglate winter (January and February) of the
ENSO event among the weak Indian monsoon years. The contour interval is 5gpm
and negative values are dashed. 5% (orless) significant area is shaded.

Fig. 14. Composite time coefficients of somedominant teleconnection patterns associated with (a) the ENSO years and (b)non-ENSO years among the years of weak
Asian summer monsoon. Time sequence ofcomposite SOI for each case is also plotted.
(Yasunari, 1988)

in the sign of the time coefficients in the same period. In the case of the ENSO years, negative values
are significant, which implies the phase of the NAO
with north/negative and south/positive anomalies.
In the case of the non-ENSO years, in contrast, large
positive values are noticeable.
It is noteworthy to state that since Walker andBliss (1932) the NAO has been noted as an oscillation independent from the Southern Oscillation(SO). A more sophisticated complex EOF analysis of
the near-global sea level pressure date made by Barnett (1985) also shows the apparent orthogonality of
the NAO to the SO. Certainly Fig. 10 shows no significant linear correlation between the Indian monsoon rainfall index and the NAO, though the NAO
is one of the most dominant teleconnection patterns
in the northern hemisphere. However, this does notnecessarily imply that the NAO is not physicallylinked with the SO. Figure 14 suggests, on the contrary, that the polarity of the NAO, coupledwith thenegative phase of the PNA-II may play a crucial rolein the occurrence or non-occurrence of the ENSO
even later on. The seasonal 500mb height anomaly
composited for the late winter (Jan. to Feb.) of Y(0) with the ENSO event (Fig. 15) apparently showsa large positive anomaly over the Aleutian low area
and the middle of the north Atlantic and large negative anomalies over the northern part of north America and central Asia. This pattern is, indeed, very
similar to the typical anomalous pattern for largesnow cover extent over central Asia (Fig. 7), except
for the prominent NAO pattern over the north Atlantic and the more southward extent of negative
anomaly toward the Arabian Sea. This anomalous
pattern strongly suggests that in the winter preceding the ENSO event a more-zonally oriented strong
jet is predominant
over north America through thenorth Atlantic while a deep trough with a southward
intrusion of the jet is located over central Asia.
A question may arise as to how this slight difference of the preceding winter circulation anomaly
over the north Atlantic through the Eurasian continent triggers the ENSO event in later seasonsor no, though both conditions are undoubtedly
responsible for a weaker summer monsoon condition.
One speculative explanation may be a difference of

55

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Journalof theMeteorological

Rossby-wave train path and associated storm tracks
over central and south Asia, which may (or maynot) facilitate the cyclogenesis and westerly windburst over the equatorial Indian Ocean in late winterthrough the pre-monsoon season, possibly throughthe mechanism suggested by Hiskins et al. (1990).
Namely, the NAO seems to play a dynamical roleof modulating the intensity and distribution of thejet stream over the north Atlantic, which in turnchanges the Rossby wave propagation characteristics on the downstream-side, i, e., over Eurasia andthe Indian Ocean.
We could hypothesize, in any event, that if only
one anomalous state of the NAO with the polarityof negative/north and positive/south is additionalin the preceding winter to the anomalous state of
the MAOS with the weak Asian summer monsoon,a very favorable condition for triggering the ENSO
event may be produced.
7. Concluding

remarks

The role of the Asian monsoon on the global-scale
climatic variability with the biennial and ENSO
time scales is discussed, by examining the two-way
interactions between the tropics and the extratropics in the seasonal cycle. In this process, theAsian monsoon functions as part of one climate system referred to as the MAOS (Monsoon and thecoupled Atmosphere/Ocean System) over southeast
Asia through the tropical Pacific.
The MAOS prominently shows a biennial oscillatory nature which tends to have one anomalous state
starting in the northern summer monsoon season

and persisting for about one year (i. e., one "monsoon year") with the mature phase in winter (Yasunari, 1990a; 1991). This anomalous state of theMAOSis combinedwith the anomalous atmosphericcirculation over the subtropics and the extratropics(i. e., PNA or reversed PNA) of the north Pacificfrom late summer through winter, through the modulation of the subtropical high and the stationary
Rossby wave propagation mechanism.
The anomalous circulation over the north Pacific thus produced in the early winter seems to
evolve to the hemispheric anomalous circulation
regime with wavenumber-oneand/or-two structure
in late winter. That is, the PNA pattern (negative height anomalies over the north Pacific andthe northeast north America with positive anomaly
over Alaska/Rockies), followed by the weak summer monsoon, tends to produce on anomalous cold
trough over northeast north America with more
zonal flow over Eurasia. By contrast, the reversed
PNA pattern, followedby the active monsoon, tends
to produce more zonal flow over the north Pacificthrough north Atlantic sector with an anomalous
cold trough over Eurasia.
The anomalous circulation thus produced over56

Societyof JapanVol.

70, No. 1

Fig. 16. Schematic diagram of the interactions between the MAOS and the chaotic
westerly flow regime (the Chaos), similar
to Fig. 13 but focusing more on the spatial structure. The role of the NAO is also
specified.

Eurasia seems to provide a favorable condition for
the extensive (or diminished) snow cover area andmass over central Asia, which in turn is responsible for the reversed anomalous state of the following
Asian summer monsoon and the MAOS, presumably
through the albedo and snow-hydrological effect ofanomalous snow cover. That is, the biennial natureof the MAOS and the climate system in the northern hemisphere may be due, at least partly, to thistwo-way interaction between the MAOS in the tropics and the westerly flow regime in the extratropics
in the seasonal cycle.
In addition, the significant polarity change of the
NAO in the preceding winter depending on the
ENSO or the non-ENSO event suggests that theNAO over the north Atlantic seems to play a rolein the climatic chaos, by stochastically amplifying
or damping the biennial nature of the MAOS. The
timing of the ENSO event or the ENSO cycle may,therefore, be modulated by this interaction between
the MAOS and NAO, both of which are basicallyindependent of each other. The MAOS seems to behave more or less as a quasi-transitive or almostintransitive climate system (Lorenz, 1968), while
the NAO may represent the chaotic behavior of thewesterly regime. The whole view of the MAOS, thechaos in the westerly regime and their interactions
through the Rossby wave propagation and the landsurface processes (snow cover, soil moisture etc.)
over the Eurasian continent is schematically shownin Fig. 16.
The mechanism of the interannual variability of
the climate system discussed here should be compared to the recent model result by Lorenz (1990).
He argued that the interannual variability of the climate system could result from an interaction between the chaotic winter circulation and the intransitive summer circulation but shows a totallychaotic behavior because of a randomly-renewed ini-February

1992

T. Yasunari

tial condition for summer by the chaotic winter condition. This conceptual model for the propagationof climatic signals in the seasonal cycle may bepartly compatible with the observational evidence
presented here. However, in reality, the anomalousstate of the MAOS seems to give persistent anomalous forcing to the westerly regime from the northernsummer to the following winter, to produce a particular planetary flow pattern as a "climatic signal",
though it is more or less disturbed by the chaotic nature of the westerly flow regime. This climatic signal in winter is, in turn, transmitted to the following
summer, through the memory of land-surface (andpossibly ocean-surface) processes during winter andspring. As a consequence, the variability of the realclimate system seems to have one maximum power
spectra in the biennial through the ENSO time scale,rather than the red or white noise spectra.
These results suggest strongly that the Asianmonsoon is manifested as a global-scale land/
atmosphere/ocean interaction between the largest
continent and the largest ocean on the earth, andplays a key role as a transmitter of climate signals between the tropics and the extratropics inthe seasonal cycle. There are, however, still many
problems to be solved. Particularly, the interactive physical processes of the Asian monsoon withthe coupled ocean/atmosphere system in the Pacific/Indian ocean sector and with the land-surface
condition (e. g., snow cover, soil moisture etc.) overthe Eurasian continent may be a central issue tobe intensively studied both observationally and theoretically. These problems should be undoubtedly
one of the main targets for the ongoing TOGA andthe forthcoming GEWEX.
Acknowledgements
The present authors wish to thank Ms. Yuki
Morinaga for computing assistance and helpful discussion. They are also indebted to Dr. M. Koderaof the Meteorological Research Institute (MRI) for
supplying the arranged snow cover data.
The
two anonymous reviewers' comments were extremely
helpful in the revision of the paper. This study has
been supported by the Grant-In-Aid for Scientific
Research from the Ministry of Education, Science
and Culture No. 01460055.

andY. Seki187

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地 球 気候 シス テムの 年 々変動 に おけ るア ジア モ ンスー ンの役 割
安 成 哲 三 ・関 祐 治1
(筑波大 学 地球 科学 系)

ENSOの

時 間 ス ケ ー ル の 地 球 気 候 シ ス テ ム の 年 々 変 動 に 果 た す ア ジ ア モ ンス ー ンの 役 割 を、 ア ジ ア の 夏の モ ン ス ー ン、 熱 帯 太 平 洋 の 大 気 ・海 洋 結 合 系 及 び 中 緯 度 の 偏 西 風 レ ジ ー ム の 間 の 統 計 的 、 力 学 的 な 関 係
を 調 べ る こ と に よ り、 考 察 し た。
ア ジ ア モ ンス ー ン は 、 熱 帯 太 平 洋 域 で の 大 気 ・海 洋 結 合 系 と密 接 に リ ン ク して お り、 モ ン ス ー ン/大 気 ・
海 洋 結 合 系(略して 、MAOSと仮 称)と(Yasunari,1990a)。
こ のMAOSはの モ ン ス ー ン頃 か ら始 ま り、 約1年
こ のMAOSのも呼 べ る 一 つ の シス テ ム を な して い る こ と が 明 らか と な っ て い る、 準2年 周 期 の 振 動 特 性 を持 っ て お り、 あ る 偏 差 状 態 は 、 ア ジ ア の 夏
持 続 す る と い う季 節 性 を示 す(Yasunari,1991)。

偏 差 状 態 は 、 亜 熱 帯 高 気 圧 の 強 弱 や ロ ス ビ ー 波 の 伝 播 と い う機 構 を 通 し て 、 北 太 平 洋 の亜 熱 帯 ・中 緯 度 の 夏 か ら秋 に か け て の 大 気 循 環 に 、 大 き な影 響 を 与 え て い る こ とが わ か っ た 。 即 ち 、 モ ン
ス ー ン の 弱 い(強い)年に は、(逆)PNAパ

タ ー ンが 卓 越 す る。 そ して 、 引 き続 く冬 の 半 球 ス ケ ー ル の 偏

西 風 循 環 は 、 こ の 秋 の 大 気 循 環 の 偏 差 が 初 期 条 件 とな っ た よ う な波 数1ま
即 ち 、PNAパ

た は2のパ ター ンが卓越 す る。

タ ー ン に よ り、 北 米 東 岸 あ るい は 極 東 の トラ フ が 発 達 し、 ユ ー ラ シ ア 大 陸 上 は よ り帯 状 流 的な 流 れ の パ タ ー ン とな る 。 反 対 に 、 逆PNAパ

タ ー ン で は 、 北 太 平 洋 か ら 北 米 域 が よ り帯 状 流 的 と な る一

方 、 ユ ー ラ シ ア 大 陸 上 の トラ フ が 発 達 しや す くな る 。
ユ ー ラ シ ア 大 陸 上 の トラ フ の 発 達 ・未 発 達 は 、 さ ら に 、 そ こ で の 冬 か ら春 の 積 雪 面 積 の 偏 差 の 形 成 と い

1現 在 所 属:(株)ウ58

ェ ザ ー ・ニ ュ ー スFebruary

1992

T. Yasunari

and Y. Seki189

う物 理 過 程 を 通 して 、 次 の 夏 の ア ジア モ ン ス ー ンの 偏 差 に影 響 す る こ とが 示 さ れ た。 即 ち 、MAOSと風 レ ジ ー ム が 結 合 した こ の 気 候 シ ス テ ム で は、 弱 い(強い)夏タ ー ンが 持 続 し、 続 く冬 に はユ ー ラ シ ア 大 陸 上 に 少(多)雪の 夏 の モ ンス ー ン は 、 強 い(弱に 、MAOSとい)状

偏西の モ ンス ー ンの 後 の 秋 に は、(逆)PNAパ
を も た らす 循 環 場 が 卓 越 す る こ と に よ り、 次

態 に な る と い う、2年 振 動 的 傾 向 を持 つ こ とが しめ さ れ た 。 この よ う

中 ・高 緯 度 の 偏 西 風 レ ジ ー ム を含 む 気 候 シ ス テ ム の 準2年

振 動 的 変 動 の機 構 は 、 ア ジ ア モ ン

ス ー ン を媒 介 と した 、 熱 帯 と中 ・高 緯 度 の あ い だ の 、 季 節 を違 え た相 互 作 用 に よ る こ とが 強 く示 唆 さ れ る。
さ ら に 、 現 実 の よ り非 定 常 的 な シ ス テ ム の 振 る 舞 い と、ENSOの状 態 の 物 理 的 な 理 解 に は 、ENSOとよ う に上 記 の 準2年

振 動が 増幅 されたは 全 く独 立 な振 動 系 と して指 摘 さ れ て い る北 大 西 洋 振 動(NAO)の、こ の シス テ ム へ の ス トカ ス テ ィッ ク な 強 制 が 非 常 に 重 要 で あ る こ と を 示 唆 す る観 測 的 事 実 も提 示 さ れ た。

59