August

1991

Local

T. Yasunari,andRemoteAppearing

A. Kitoh

Responses
in theandT. Tokioka

to Excessive

Snow

Northern

A Study

Springwiththe

By Tetsuzo

and

MRI *

473

Mass

Summerover

Eurasia

ClimateGCM

Yasunari

Institute of Geoscience, University of Tsukuba, Tsukuba, Ibaraki 305, Japan
and
Akio

Kitoh

and

TatsushiMeteorological Research Institute,

Tokioka

Tsukuba, Ibaraki 305, Japan(Manuscript received 7 September 1990, in revised form 11 June 1991

Abstract

The effect of excessive snow mass over the Eurasian continent on the spring and summer climate
is investigated by using the MRI *GCM. The ensemble mean of the four runs (SNOW runs) with theexcessive snow mass of 5 cm (water equivalent) at the beginning of March over the snow cover area ofthe continent is compared with that of the control runs, to deduce the effect of the snow mass on the
climatic parameters in the later seasons.

The main results are summarized as follows:(1) In spring, the albedo effect is dominated in the lower latitudes particularly over the Tibetan
Plateau. The reduced net radiation by the anomalous snow cover balances the reduced surface sensible
and latent heat fluxes, which account for the significant decrease of surface temperature, cloudinessand total diabatic heating over there in the SNOW runs.
(2) In summer, in contrast, the snow-hydrologicaleffectis significant, particularly in the mid-latitudes.
The increase of ground wetness in the SNOW runs causes anomalous cooling and higher pressure near
the surface. A moderate signal of the weakened Asian summer monsoon is also obtained. However,
the increase of evaporation activates cumulus convection, which partly compensates for the decreaseof total diabatic heating by the cooling near the surface. This evaporation/convection feedback seems
to work, on the other hand, to sustain the increased ground wetness throughout the summer.
(3) The atmospheric teleconnection patterns induced by the anomalous snow cover over the Tibetan
Plateau and east Asia significantly appear over the north Pacific and the North American continent inspring through late summer. These anomalous circulations cause the considerable decrease of surface
temperature over the northeastern part of North America.
(4) The implication of these results for the Ice Age issue is also briefly discussed.
1. Introduction
At the end of the last century, Blanford (1884)
first pointed out the possible relationship between
the Himalayan spring snow cover and the succeeding summer monsoon rainfall over India, though his
result attracted little attention from meteorologists
for a long time afterwards. By using satellite snow
cover data over Eurasia, Hahn and Shukla (1976)
re-demonstrated the link, showing the apparent reverse relationship between the Eurasian snow coverextent in winter and the India monsoon rainfall inc1991,

MeteorologicalSocietyof Japanthe following summer. Several succeeding studies(Dey and Bhanu Kumar, 1983: Dickson, 1984, etc.)
confirmed their results.
In addition, recent observational studies havestrongly suggested that the snow cover over Eurasia
plays an important role in producing the anomalous
condition of the coupled ocean/atmosphere systemin the equatorial Pacific (e.g., E1 Nino) by means of
the Asian monsoon (Barnett, 1985; Yasunari, 1987,
1990, 1991).
These observational relationships between theEurasian snow cover in winter and spring and thetropical climate system in later seasons are sug-

474

Journalof theMeteorologicalgested to be based principally on two physical processes. One is the albedo effect, as suggested by
Hahn and Shukla (1976). The high albedo of thesnow reduces the total incoming solar radiation,
which may in turn reduce the heating of the groundsurface of the continent from spring to summer season. Some general circulation model (GCM) studies with the Ice Age surface boundary conditions(Williams,1975; Manabe and Hahn, 1977; Prell andKutzbach, 1988, etc.) have supported this idea.
The other process is the snow-hydrological effect,
which was first examined by Yeh et al. (1983) by using a simplified GCM. Miyakoda and Chao (1982)
pointed out the importance of this effect on longrange forecasting. Recently, Barnett et al. (1989)
attempted to discuss this effect as well as the albedo
effect more comprehensively by the aid of GCM experiments with full physics and geography. That is,
the snow mass not only consumes the solar energy
during the snow melting, but also reduces the heating of the ground after the snowmelt by moistening
the surface soil layer with the melted snow. This effect may provide a plausible explanation of the timelag relation between the winter snow cover and thesummer monsoon. Yamazaki (1989) has suggested
that the combination of these two effects is actuallyimportant for producing the anomalous atmosphericcondition over the continent in summer. These two
effects of snow cover are schematically shown in Fig.
1.
On the other hand, recent observational studies
have suggested some significant time-lagged teleconnection patterns in the mid-latitudes induced by theanomalous snow cover over Eurasia. By using lagcorrelation synoptics Morinaga (1987) and Morinaga
and Yasunari (1987) deduced some significant lagcorrelation patterns between the anomalous snow
cover over central Asia and the atmospheric anomalies, particularly on the downstream-side of the continent. Kodera and Chiba (1989) showed that the
Okhotsk high intensity at 500 mb in June, which isan important factor in controlling the Baiu (frontalrain) activity over Japan, is highly correlated to thesnow cover anomaly over central Asia in April, presumably by means of stationary wave trains overeastern Eurasia. Iwasaki (1991) pointed out the apparent 1-year lag correlation between the snow cover
over eastern Eurasia and that over the north American continent.
Furthermore, Kuhle (1987) argued the possibilityof the "Tibetan Ice Sheet" during the Quaternary
Ice Age and its role on triggering the global-scaleglaciation itself.
The results of these observational and paleoclimatic studies urge us to assess more quantitatively the effects of the Eurasian snow cover anomaly
on the local as well as remote or global-scale responses under different basic states of atmosphere.

I)

2)

Society

ALBEDOSNOW

Fig.

of JapanVol. 69, No. 4

FEEDBACKHYDROLOGICALFEEDBACK

1. Schematic

diagramfor the albedo

back and the hydrologicalfeedback

cover

march

duringthe seasonal

feedof snow

from

win-

ter to summer.

For example, it has not been clear how and to what
extent the snow mass anomalies of each season andeach area produce atmospheric circulation anomalies via the anomalous energy balance on the groundsurface and in the atmosphere. Seasonal as well as
spatial dependencies seem to be crucial for the twophysical processes (Fig. 1). Particularly, the relative importance of snow cover anomalies over the
Tibetan Plateau should be investigated in more detail from these points of view. How the snow coverprocess interacts with other physical processes, e.g.,ground-hydrology, evaporation and cumulus convection, may also be a key factor for producing atmospheric anomalies.
This study focuses on these problems through theanalysis of the GCM experiments with anomalous
snow cover in spring over Eurasia.
2. Design of GCM experiments
2.1 The Model
The GCM utilized for this study is a troposphericversion of the MRI •GCM-I (Tokioka et al., 1984)
produced at the Meteorological Research Institute(MRI). The model atmosphere consists of 5-layers in
the vertical with a top at 100 mb, and has a horizontal grid resolution of 5* in longitude and 4* in latitude. The formulation of diabatic processes includesAugust1991

T. Yasunari,

A. KitohandT. Tokioka

475

the cumulus parameterization of Arakawa and Schubert (1974), interacting with the planetary boundary layer (PBL) and the radiative process. The diurnal and seasonal variations of solar insolation aretaken into account.
The land surface processes in the model are as
follows: The soil moisture is predicted by using the
one-layer bucket model with a field capacity of 15
g cm-2. The heat and water budgets at the modelsurface are governed by*

Tg is the ground temperature, C the bulk heat capacity of the ground, RS the solar flux absorbed at
the surface, RL the net upward longwave flux atthe surface, FS the sensible heat flux, Fw the latent
heat flux (i.e., Fw = L x E, L is the latent heat ofwater vapor), Hi the snow melt and/or heat conduction through sea ice, W the available soil moisture,
Pr the precipitation rate, E the evapotranspiration
rate, R the surface runoff, Sm the snowmelt, theevapotranspiration efficiency, Ep the potential evapotranspiration rate, GW the ground wetness, Wmax
the maximum available soil moisture, and Dw thesoil moisture deficiency. Here, Wmax=15gcm-2 isadopted for all land grid points. Surface runoff R isa function of Pr and Dw (see Eq. (6)), so that it canoccur when rainfall occurs even if the ground is not
saturated.
The seasonally varying land surface albedo isadopted from Matthews (1983) depending upon vegetation conditions. Surface albedo over the ocean istaken from Posey and Clapp (1963). Surface albedo(a) over the snow is a function of snow mass andmasking depth of snow which depends on the terrain height as follows:*

where *b is the bare land albedo by Mathews (1983),
s is the deep snow albedo (=0.70), SN is the* snow
mass. The masking depth of snow dsw is defined asfollows:

Fig. 2. (Upper) simulated precipitation andwind vector at 850 mb in the control run
for the northern summer. The shaded(dotted) area denotes precipitation of morethan 5 mm (less than 1 mm) per day.
(Lower) observations based on the 197987 NMC wind and the Jaeger (1976) precipitation.
where h is the surface terrain
height in km.
The sea surface temperature
is prescribed
the monthly
ocean.
Thewithclimatologicaldataover the wholesnowfall rate is predictedby using the

precipitation
rate and the air temperature
at thesurface.
The snow mass over land is predicted
as aresult of heat balance of the land surface.

2.2 Experiments
The model is integrated for three and a half years
starting from March 1 as a control run (hereafter referred to as the CNTL run). The model performanceof the CNTL run, particularly for the Asian monsoon is essential as a premise of the evaluation of
the anomaly run discussed later. This CNTL run
is fundamentally the same as Kitoh and Tokioka(1986) and Kitoh et al. (1988), which successfully
reproduced the Asian and African monsoons. Fig.
2 shows the mean wind vector field at 850 mb andtotal precipitation for summer (June, July, August)from (a) the CNTL run and (b) the observations.
The large-scale features of the Asian monsoon, suchas the southwest monsoon flow over the Arabian sea,

476

Journalof theMeteorologicalSocietyof JapanVol. 69, No. 4

Fig. 3. Snow mass distribution in the CNTL run (left) and observed snow cover extent from
NOAA/NESDIS data (right) for March through June. In the left diagrams, values of more than 1.0 cm(water equivalent depth) are shown with solid lines, and values of more than 5cm are shaded. Values
between 1.0cm and 0.1cm are heavily dotted, and those between 0.1cm and 0cm are lightly dotted.
In the right diagrams, snow cover frequency of more than 50* (during the 21 years) are indicated.
rainfall distribution over the monsoon Asia and central Asia, subtropical high over the Pacific etc., aresimulated well in the CNTL run, though the rainfall and westerly wind over the equatorial western
Pacific is considerably less than the observations.
In the anomalous snow cover run (hereafter referred to as the SNOW run), 5cm (water equivalent) snow mass anomaly was added only over thesnow-covered area between 30*N and 60*N of theEurasian continent on March 1 of every CNTL run
year as an initial condition to minimize the shockof abrupt change in the integration. The model is
integrated for 184 days up to September 1. Maximum snow mass in the CNTL run appeared at theend of February, which is nearly coincident withthat of the observation. A remarkable feature in the
CNTL run is a very reasonably simulated seasonalchange of snow cover area if we refer to the observation from NOAA/NESDIS data, as shown in Fig.
3. Although the best climatology of (water equivalent) snow depth over Eurasia is not yet available,
the maximum snow depth distribution in the CNTL
run seems to be highly compatible with the observed
distribution from Schutz and Bregman (1987) overthe major part of the continent. The snow depth

in the model tends
the lower latitudes
continent.

to show a larger value locally inand in the easternpart of theFour half-year (March to August) sets of the
SNOW run are thus obtained as shown in Fig. 4,
and the ensemble means of the outputs from thefour SNOW runs are compared with those from thefour corresponding CNTL runs, to deduce the effects of the anomalous snow cover in spring (March
1). It may be a tough question as to how reasonably
we can give the snow mass anomaly in the model,
since we have so far very little information about
the interannual as well as seasonal variability of realsnow mass (depth) on the continental scale. However, recent satellite-derived snow depth informationfrom Nimbus-7 SMMR data for 7 years (Chao et al.,1987) suggested the areal-mean year to year maximum snow depth over the whole snow-covered area
ranges 5cm to 8cm (water equivalent), which implies that a 5cm (water equivalent) anomaly overthe major part of Eurasia seems to be a very reasonable value for the interannual variability in thepresent climate. The mean snow mass distribution
in the CNTL run and the anomaly added in the
SNOW run on March 1 are shown in Fig. 5.

August1991

T. Yasunari,

A. KitohandT. Tokioka

477

Fig. 4. Design of the GCM experiment forthe control (CNTL) runs and the anomaly(SNOW) runs.

Fig. 5. Distribution of anomalous snow cover(5cm water equiv alent depth) added in
the SNOW run on March 1.
3. Changes
sia

in thesufacecondition

over

Eura-

Since the two physical processes, i.e., the albedo
effectand the snow-hydrologicaleffect,are supposed
to be essential for producing the anomalies in theatmosphere as shown in Fig. 1, the differences insurface parameters over the Eurasian continent between the CNTL run and the SNOW run are examined in this section. Figure 6 shows the seasonal change (fromMarch to August) of the monthly
mean values of the two runs and the anomaly(SNOW run-CNTL run) of snow mass and albedo
averaged over the Eurasian continent east of 75*E.
The increases of albedo in the SNOW run are noticeable along the southern periphery of snow coverarea of each month in spring (March to May), mostly
correspondingto the delayed snowmelt. This albedo
change is, however,,hardly recognizable after June,
when the snow-covered area over Eurasia almost
completelydisappeared in the SNOW as well as the
CNTL runs.
Figure 7 shows the same diagram as Fig. 6 but
for surface air temperature (Ts) and soil moisture(GW). The TS (and GW) anomaly shows totally
negative (and positive) values particularly in the
mid-latitudes (30*N-50*N),undoubtedly associatedwith the excesses in snowmass and albedo (Fig. 6).
It should be noted that there seem to be two stages
in the seasonal march of the negative Ts anomaly,
one in the lower latitude (around 30*N) in spring(March and April), and the other in the higher lat-Fig. 6. Seasonal change (from March to August) of monthly mean snow mass (upper)
and albedo (lower) of each latitude for the
CNTL run (left), SNOW run (middle) and
its difference (SNOW-CNTL) (right). The
unit of snow mass is cm (water equivalent). The heavily (lightly) dotted area insnow mass shows 0.1-1.0 (0-0.1)cm depth.
The unit of albedo is %. Areas of 5cm(40*) or larger are shaded in snow mass
(albdo) diagrams of CNTL and SNOW
runs. Negative values are shaded in the
SNOW-CNTL diagrams.

Fig. 7. Same as Fig. 6 but for surface air temperature (TS) and ground wetness (GW).
The unit of TS is C. The unit of GW is% of the full water storage (15cm waterequivalent).

itudes (40*N-50*N)in summer (June to August). A
similar feature in the positive GW anomaly is also
noted, but in this case the anomaly in the second
stage (in summer) is far larger than that in the firststage (in spring).

478

Journalof theMeteorologicalSocietyof JapanVol. 69, No. 4

anomaly in mid-latitudes (40*N-50*N) is mostly related to the anomaly in east Asia to the northeastof the Tibetan Plateau.
The distribution of the GW anomaly is complicated, but the area of the large negative Ts anomaly
between 80*E and 120*E totally corresponds to thepositive GW anomaly there. However, there is noalbedo anomaly there. It is worthwhile to statethat the large anomalies in the second stage (overeast Asia in summer) appear completely after thesnowmelt over the whole of Eurasia while those in
the first stage (over the Tibetan Plateau in spring)seem to be related to the in-situ snowmass andalbedo anomaly. These features suggest that the effect of snow cover appears in a different way depending upon seasonal timing and geographical location,
as is also pointed out by Barnett et al. (1989). This
aspect will be discussed further in the next section.

4. Changes in heat and water balance overEurasiaFig. 8. Distributions of albedo (%), GW (%)
and Ts (*) anomalies over the northern hemisphere in (a) April and (B) August. Areas of positive (negative) anomaly
with the significant level of 5% or less areshaded (dotted).
The distributions of albedo, GW and Ts anomalies over the northern hemisphere in April and August are shown in Fig. 8, to see more precisely thedifference in the spatial distribution of these anomalies between spring and summer. It is apparent
that the negative Ts anomaly as well as the positive
albedo and GW anomaly in the subtropics (30*N)in spring (April) are mostly due to the significant
anomalies over the Tibetan Plateau. In summer(August), on the other hand, the large negative Ts

The large negative Ts anomalies at the two stages( i. e., in spring and in summer) are supposed to result from the different physical processes at the surface, since the anomalies of other related parameters (GW, albedo, etc.) have shown considerably
different features between the two stages. To examine this aspect more in detail, the seasonal changesof heat and water balance at the surface and in
the troposphere, and their anomalies (SNOW runCNTL run) are computed for the typical areas ateach stage, that is, over the Tibetan Plateau (32°N40*N, 75*N-105*E),and over east Asia (40*N-48*N,
90*E-120*E),as shown in Fig. 9 and Table 1, respectively.
Over the Tibetan Plateau in April (the first stage)
the albedo effect is dominant as shown in Table 1.
This stage is obviouslycharacterized by the fact thatsnow exists where normally it would not. In the
SNOW run the net short-waveradiation (RS) is significantly reduced due to the large positive snowmass (SN) or albedo (*) anomaly. Because of therelatively cooled surface, both latent heat flux (FW)and sensible heat flux (FS) are considerably suppressed. Less cloudiness (N) and precipitation (Pr)associated with more stable atmosphere and relatively downward motion may be responsible for thesignificant reduction of total heating (Q) in the atmospheric column.
Over east Asia in August (the second stage), in
contrast, although there is no SN and * anomaly,
large negative Ts and Tg anomalies are still apparent. This is due to the relative cooling of the surface
basically by the increase of FW (evaporation) andthe decrease of FS, which are directly related to theincrease of GW after the snowmelt of the excessive
snowmass. In addition, less RS due to the increase ofcloud cover (N) through the increased evaporation

August1991

T. Yasunari,

A. Kitohand T. Tokioka

479(1974), where cumulus convection is very sensitive toevaporation, or more directly, the moisture amountin the planetary boundary layer (PBL).
5. AsianFig. 9. Surface heat and water balance andcloudiness over (a) Tibetan Plateau in
April and (b) east Asia in August. The left
(right) column of each pair represents thevalue in the CNTL (SNOW) run. Units
and notations are indicated in Table 1.

may also be partly responsible for the large negative Ts (or Tg) anomaly. This snow-hydrologicaleffect on Ts (and Tg) becomes more significant inlate summer (July and August) than in early summer (June), presumably because evaporation nonlinearly increases as Ts (or Tg) increases following
Clausius-Clapeyron's law.
Although the decrease in Ts and Tg is extremely
large in the SNOW run, there is no significant
change in Q, which implies that the increased diabatic heating in the upper and middle part of thetroposphere by cumulus convection, caused by theincreased evaporation from the surface compensates
the cooling effect near the surface. The increased
Pr, on the other hand, has a role in the maintenance of wet ground throughout the summer. Thatis, the snow-hydrologicaleffect in this model seems
to be very effective for cooling in the lower troposphere, but not for the reduction of total heating in
the whole troposphere. This characteristic natureof the model response to the excess snowmass andwet ground seems to be closely associated with the
cumulus parameterization of Arakawa and Schubert

summermonsoonanomalies

The anomalies in some Asian summer monsoon
indices, induced by the excess snow mass in the
SNOW run are described here. As is pointed out
in the previous section, the snow-hydrological effect in this model to the total heating of the atmosphere is more or less diluted by the evaporation/cumulus convection feedback. Nevertheless,some signals of a weakened Asian summer monsoon
are found, particularly in late summer (July andAugust). Sea level pressure (SLP) over the continent is a measure of land surface heating. Figure 10
shows the monthly mean SLP in the CNTL run andthe anomaly (SNOW run-CNTL run) for August.
There is an indication of a weaker monsoon troughover northern India (positive SLP anomaly) in the
SNOW run, although it is not significant. It is interesting to note that a large negative SLP anomaly
area is seen along the northern and eastern periphery of the positive SLP anomaly, and another broad
area of positive anomaly exists over Alaska and thenorth polar region. These significant anomalies are
discussed further below.
The other parameters particularly related to themonsoon activity (i. e., wind, precipitation, diabatic
heating, divergence)show the weakening of the monsoon in the SNOW run, though most of the anomalies are marginally significant. Figure 11 shows thesame diagram as Fig. 10 but for the wind field at
850mb. The easterly anomaly over the Arabian sea
is noticeable, indicating the weaker southwest monsoon flow in the SNOW run. This weaker monsoon
wind is consistent with the weaker monsoon troughover northern India.
Another important measure of the monsoon isprecipitation. Figure 12 shows the same diagram as
Fig. 10 but for the total precipitation. The CNTL
run has simulated well the broad feature of the observed distribution of Asian summer monsoon rainfall (e.g., refer to Legates (1987)). The anomaly
distribution shows a general decrease of rainfall of
about 20% in the area of maximum monsoon rainfall over southeast Asia. Over northeast Asia andcentral Asia, in contrast, a positive anomaly of precipitation is noticeable, associated with the snowhydrological feeadback as discussed in Section 4.
The total heating of the atmosphere over India
through southeast Asia is also considerably reduced.
Associated with the reduced atmospheric heating
and convective rainfall, the velocity potential fieldin the upper troposphere (200mb) shows a weakened divergent monsoon circulation over south andsoutheast Asia, as shown in Fig. 13. Although thisweakened monsoon circulation in the SNOW run is480

Journalof theMeteorologicalSocietyof JapanVol. 69, No. 4

Table 1. Anomaly (SNOW run-CNTL run) of heat and water budget over (a)the Tibetan Plateau and (b)Mongolia (east Asia). Gothic values are those with the significant level of 5% or less.
Heat and Water Budgeta) Tibetan Plateau (32-40N, 75-105E)b) Mongolia (40-48N, 90-120E)

SN
1- *
RS
RL
FW
FS
HG
N

snow depth (cm water)
1- surface albedo (%)
net shortwave radiation (W/m2)
net longwave radiation (W/m2)
latent heat flux (W/m2) = L*E
sensible heat flux (W/m2)
net heat budget at the surface
cloudiness (%)

Fig. 10. Monthly mean SLP in the CNTL run(upper) and the anomaly (SNOW runCNTL run) for August. In the upper diagram, pressure values of 1020mb or more
(1000 mb or less) are shaded (dotted). In
the lower diagram, positive (negative) values with the significant level of 5% or lessare shaded (dotted).

Ts
Tg
Q
E
R
Pr
GW

surface air temperature (*)
ground temperature (*)
total diabatic heating (W/m2)
evaporation (mm/d)
runoff (mm/d)
precipitation (mm/d)
soil moisture (cm water)

Fig.

11. Same

as Fig.

10 butfor windvectors

at 850 mb.

typically shown in August, similar features associated with the weakened convective activity over thesoutheast Asia through the western Pacific are also
apparent during late spring (May) through summer.
Thus, the weakening of the Asian summer monsoon

August1991

T. Yasunari,

A. KitohFig. 12. Same as Fig. 10 but for precipitation. In the upper diagram, values of 7
mmday-1 or more (1 mm day-lor less) areshaded (dotted). In the lower diagram,
positive (negative) values with the significant level of 5% or less are shaded (dotted).
principally by the snow-hydrological effect of theanomalous snowmass has been confirmed. These
results for the Asian summer monsoon, are fundamentally the same as obtained by Barnett et al.
(1989), though the details of the experimental procedures and the results are quite different from eachother. Further comparisons between the present results and theirs will be made later.

6. Teleconnections
tudes

in middle and high lati-

It has been established both from observations
and theories (Hoskins and Karoly, 1981; Webster,1982 etc.) that the anomalous diabatic heating induced by the anomalous surface heating condition
produces anomalous local and remote responses in
the atmospheric circulation field. Some observational studies have suggested that the anomalous
snow cover over Eurasia induces the in-situ and remote anomalous circulation referred to in the introduction. Here the anomalous circulation field(SNOW runs-CNTL runs) is examined from thispoint of view.
Figure 14 shows the monthly mean geopotential
height anomaly at 500 mb from March to August.
In spring the anomalous patterns are very similar toeach other, particularly in March and April. A negative anomaly is significant over east Asia. It should
be noted that to the northeast of this negative
anomaly there exist a distinct positive anomaly area
over Kamchatka Peninsula/far east Siberia and also
another negative anomaly over the Labrador/Baffin
Bay area, suggesting the propagation of a Rossby
wave train forced by the anomalous vorticity overeast Asia. The negative height anomaly over east
Asia can be explained as a local response to theandT. Tokioka

Fig.

13.
tential

481

Same

as

Fig. 10 but

for velocitypo-at 200 mb.

increased diabatic cooling over and around the Tibetan Plateau (see Table 1 and Fig. 8) and the increase of diabatic heating over the north Pacific near
the east coast of Asia, as theoretically pointed out
by Smagorinsky (1953). Interestingly, the same typeof anomaly pattern was obtained in the observed
anomaly pattern for April, relevant to the anomalous snow cover over central Asia (Morinaga, 1987).
In June an area of large negative anomaly appears
over the Kamchatka Peninsula where the Okhotsk
high is located in this season. An area of positive anomaly is also found over the Canadian arctic
region, to the northeast of the negative anomaly,
and another area of negative anomaly is marked
over the northeast of North America, which again
shows a Rossby wave-train pattern originated fromthe Okhotsk high area.
The change of the anomaly pattern over east
Asia from the previous spring months may be related to the change of the seasonal basic state there.
The monthly mean total heating field shows that inthis month the heating/cooling contrast between thecontinent and the ocean is completely reversed fromthe winter and spring condition to the early summer condition, i, e., heating over the continent andcooling over the Okhotsk sea. This seasonal change
in horizontal contrast of diabatic heating seems
to be essential for the formation of the Okhotsk
high or ridge in the middle troposphere (Asakura,1968). The anomalous cooling over the continent
by the snow-hydrological effect weakens the heating contrast, and may result in the weakening of
the Okhotsk high or ridge. Surprisingly enough, theanomaly pattern shown here is quite similar to that
deduced from the observed lag-correlation pattern
between the anomalous snow cover over central Asia
in April and 500 mb height anomaly in June (Kodera
and Chiba, 1989), as reproduced here in Fig. 15.
In July there is no significant change in the circulation patterns as a whole. This seems to be related
to the basic state of the zonal wind and the spatial482

Journalof theMeteorologicalSocietyof JapanVol. 69, No. 4

Fig. 15. Correlation coefficients (100) between April snow cover over central Asia
(50*-80*E, 50*-60*N) and 500 mb heightin June, based on the data for 1968-86.
Positive values (of 0.45 or more) are dotted (shaded). (Kodera and Chiba, 1989).

Fig. 16. Moisture flux anomaly (SNOW runCNTL run) at 850mb in August. The unit
is gms-1.
Asiathrough

remarkable

Fig. 14. Monthly mean geopotential heightanomaly at 500mb in March through August. Positive (negative) values with thesignificant level of 5% or less are shaded(dotted).
distribution of total heating in this month. The midlatitude westerly wind in this month over Eurasia is
the weakest in the year, and the area of maximum
heating associated with the Asian monsoon is confined to the easterly zone in the lower latitudes bothin the CNTL run and the SNOW run. Under thiscondition the anomalous cooling near the surface by
the anomalous soil moisture in east Asia seems to
be too small to produce effectively the anomalous
circulation in the mid/high latitudes.
In August the anomalous circulation with the
Rossby-wave-like pattern is again evolved over eastthefeaturearcticnorthin thisAmericanmonthis theregion.

A

anomalousheat source and upper-level divergence (Fig. 13) overthe coastal area of east Asia, i. e., over Japan islands. This anomalous heat source seems to be related to the activated baroclinic disturbances associated with the intensified warm and moist advection
from the southeast ahead of the intensified troughthere, as shown in Fig. 16. It is noteworthy that thisintensified trough seems to have been non-linearly
amplified by the interaction of cold air advection
along the north-eastern periphery of the anomalous
cold high in the lower troposphere over the Mongolian region and the warm air along the east coast of
the Eurasian continent, as is demonstrated in Fig.
10 and Fig. 11. The mean westerly wind in the troposphere over this region has become stronger thanthat in July, which enables a stronger teleconnection
on the downstream side.
Barnett et al. (1989) also noted the teleconnection
patterns over east Asia through the north American
continent in their model results for the anomalous
snow cover over Eurasia, though the spatial scales of
their anomaly patterns are generally far larger thanthose of ours. This difference between the two model
results may be due to the difference of the spatial
scales as well as the thickness of snow cover anomaly
given in each model. The snow mass anomaly intheir model seems to have a whole-continent scale,
while that in our model is more regionally confinedAugust1991

T. Yasunari,

A. Kitoh

to the interior of the continent.
This aspectof thesnow cover issue, i, e., the sensitivityof the teleconnection pattern
to the size and volume of snow coveranomaly,
may be one of the important
problems
tobe investigatedfurther.
7.

Discussion

7.1 Seasonality in the snow feedbacks
The aforementioned results have elucidated thatthe feedback of anomalous snowcover over theEurasian continent produces atmospheric anomalies
through the anomalies in the surface heat and waterbalance. The albedo effect is proved to be significant
in spring in the lower latitude, particularly over the
Tibetan Plateau, where the incoming solar radiation in this season is extemely strong. Although thedirect response of the atmosphere (e.g., lower Ts,
higher SLP etc.) to this effect is limited over andaround the Tibetan Plateau, the indirect responseby way of the teleconnection appeared nearly hemispherically as described in Section 6. This result at
least partly contrasts with the argument by Barnett
et al. (1989).
During summer months, the snow-hydrologicalfeedback has induced a weakening of the Asian summer monsoon to some extent, particularly in the
SLP, precipitation and divergence field. The anomalies of these parameters over the monsoon region are,however, generally less significant and more localized in size than those appearing in the experiments
by Barnett et al. (1989). This may be partly due
to the enormous snow mass anomaly in their experiments, where the snow mass difference between
their SNOW (D) run and the CNTL (C) run in Juneaveraged over Asia was 10 cm. In our experiments
this difference in June was only traceable over thelimited area of northern Asia, as shown in Fig. 6,
which seems to be more realistic if we compare theseasonal change of our model snow mass distribution
to that of the observed snow mass distribution estimated from Nimbus-7 SMMR data (Morinaga andIgarashi, 1991).
7.2 Evaporation/convection feedback
Another reason for a less significant monsoon in
our SNOW runs may be due, as commented in Section 4, to stronger evaporation/cumulus convection
feedback in our model. That is, the areas of positive
soil moisture anomaly in the SNOW run correspondwell with the areas of positive anomaly in evaporation, water vapor content in the PBL, precipitation,and total diabatic heating in the air column particularly in summer months (not shown). This effectmay at least partly dilute the cooling effect of evaporation near the surface to the total heating of theatmosphere, by adding the more cumulus heating in
the middle and the upper atmosphere.
This result contrasts with that by Barnett et al.

andT. Tokioka

483(1989), where surface cooled by surplus moisture
flux suppresses convection there. This notable difference between the two model results may be partly
due to the difference of cumulus parameterizationin the models. In our model with the ArakawaSchubert scheme, convection seems to be more sensitive to the moisture amount in the PBL thantheir scheme ( i. e., Kuo's parameterization), where
it needs large-scale moisture convergence.

On the other hand, this evaporation/convection
feedback enhances and prolongs the snowhydrological feedback near the surface by increasing cloudiness and precipitation, which tend to sustain reduced Rs and increased GW and, as a result,lower Ts and higher SLP throughout the summer.
Fundamentally the same results as ours are notedin the GCM experiment by Yeh et al. (1984), who
examined the effect of the excessivesoil moisture on
the climate anomaly. They pointed out that the excessivesoil moisture over the mid-latitude rain belt
tends to prolong the moisture anomaly and increase
precipitation by recycling water between land surface and atmosphere. By contrast, this effect is relatively weak over the dry regions in the subtropics.
We have so far very little information about thereal feature of cumulus convection and precipitationin the interior of the Eurasian continent in summer. We should emphasize, however, that this evaporation/convection feedback is supposed to be feasible particularly over the seasonally-heated continent in summer, where large-scale convergence inthe lower troposphere is generally dominant. In
fact, the GCM experiment by Delworth and Manabe (1989) with the realistic geography also suggeststhat this feedback by soil moisture anomaly is significantly large in the middle latitudes over the continent in summer. Once this feedback works overeast Asia through the interior of the continent in the
mid-latitudes, localized circulation anomalies may
be produced there by the anomalous diabatic heating in the middle and upper troposphere.
It seems to be crucial for the snow-hydrologicalfeedback, therefore, how the increased evaporation by the excessive snow affects convection/precipitation processes. In other words, the characteristic nature of the combined snow-hydrology/evaporation/convection feedback seems to provide
a clue for the late effect of snowmass to the summer
climate. This may really be a new but important
aspect of the snow cover issue that should be examined further. It includes the assessment of, e.g., thelarge-scalewater cycle, and of a more realistic cumulus parametarization over the interior of the summer
continent, associated the snow-hydrologicalprocessduring spring through summer.

484

JournalFig.

17. Surface
areair temperature

anomalyof theMeteorologicalin March

through

Society

August.

of JapanContours

Vol. 69, No. 4

are 1*and negativevaluesdashed.

7.3 Teleconnection over North America and the Ice
Age issue
The PNA-like teleconnection pattern during
spring through summer, induced by the anomalous
forcing over east Asia associated with the anomalous
snow mass over Eurasia, may be one of the important features of the general circulation in this experiment. This model result is in good agreementwith the observed circulation pattern deduced by
Kodera and Chiba (1989) (Fig. 15). Yasunari (1988)
has also offered observational evidence for the dominance of this pattern particularly in late summer ofa weak Indian monsoon year. Large month to monthchanges in the teleconnection patterns may be mostprobably due to the seasonal changes of the zonal
mean state, as has been noted by Kang (1990).
Associated with this pattern, an intense anoma-

lous trough at 500 mb over the northeast part of theNorth American continent (Fig. 14) has appeared
during spring through summer in the SNOW run.
This persistent trough may be very important for
the issue of the Ice Age initiation, since this troughseems to provide a favorable condition (i.e., persistent negative Ts anomaly) for the snow accumulation there, as shown in Fig. 17. We will note here
the important implication of the present result for
the Ice Age problem.
Lamb (1977), for example, suggested that the formation of an ice sheet over the North American
continent could lead the earth to the stage of full
glaciation by triggering another ice sheet over Scandinavia by way of the downstream adjustment ofstationary planetary waves. He commented further
that the persistent cold trough over this region evenAugust1991

T. Yasunari,

A. Kitohthrough the warmer seasons may be a key factor for
the formation of the Laurentide Ice Sheet, though
he did not mention how the cold trough could havebeen maintained. Kuhle (1987) suggested, on theother hand, the key role of the cooled Eurasiancontinent, particularly the completely-glaciated Tibetan Plateau, in triggering hemispheric cooling andglaciation, though he did not mention the possible
dynamical process of hemispheric cooling.
The present result may suggest a common answer
to their questions. That is to say, the persistentanomalous snow cover over Eurasia including the
Tibetan Plateau, which might be the first signal of
reduced solar radiation by some external or internal
forcing, could be responsible for the anomalous circulation over North America noted by Lamb (1977),
by means of the downstream teleconnection deduced
here. In fact, a very recent model study by Short etal. (1991) has suggested that the maximum responseof surface temperature change associated with the
Milankovitch forcing occurs over the interior of theEurasian continent, not over the polar regions. A recent GCM study by Oglesby (1990) also showed thatin addition to Alaska, western Canada and Siberia,
the Tibetan Plateau is one of the "glaciation sensitive" regions to the initial imposition of heavy snow
cover. These problems should, of course, be further
examined by the long-term integration of the models
with interactive oceans.
8. Conclusions
The effect of excessive snow mass over theEurasian continent on the spring and summer climate has been investigated by performing the MRI
GCM model experiments. The ensemble mean offour SNOW runs with the excessive snow mass of
5cm (water equivalent) on March 1 over the snow
cover area of the continent has been compared withthat of the CNTL runs, to deduce the effect of thesnow mass on the climatic parameters. The main
results are summarized as follows:(1) In spring, the albedo effect is dominated in
the lower latitudes particularly over the Tibetan
Plateau. The reduced net radiation by the anomalous snow cover balances the reduced surface sensible and latent heat fluxes. Cloudiness is also reduced
and the total diabatic heating in the atmosphere isconsiderably diminished. Although the direct response of the atmosphere to this effect is more orless localized, the teleconnection pattern induced bythis localized anomaly seems to be an important feature of the total influence of the snow cover.
(2) In summer, the snow-hydrological effect is significant particularly in the mid-latitudes. The increase of ground wetness in the SNOW run causes
anomalous cooling and higher pressure near the surface. However, the change in the total diabatic
heating is not significant mainly due to the evap-

andT. Tokioka

485

oration/cumulus convection feedback, which partly
counteracts the cooling near the surface. This evaporation/convection feedback seems to work, on theother hand, to sustain the increased ground wetness
throughout the summer.
(3) Some parameters of the Asian summer monsoon
activity indicate a weaker monsoon in the SNOW
run, although the snow-hydrological effect seems tohave been diluted to some degree by the evaporation/convection feedback as mentioned above. Associated with this weak monsoon signal, the sea levelpressures have increased over the Asian monsoon region through the equatorial Indian Ocean as a wholein early northern summer. This feature is consistent
with the observed initial stage of an ENSO event.
(4) A remarkable feature in this experiment is theatmospheric teleconnection patterns over the northPacific through the North American continent, induced by the anomalous snow cover over the Tibetan
Plateau and east Asia during spring through latesummer. These circulation patterns are responsible for the considerable decrease of surface temperature particularly over the northeastern part of North
America. These results are partly supported by theobserved circulation anomalies associated with the
weak Asian monsoon condition (Yasunari, 1988).
The signatures in this experiment also suggest thepossible role of the anomalous snow cover over Eurasia on the initiation of the Ice Age.
The results obtained here are, however, still tentative since the parameterization of ground hydrology,
which may be a key factor of the snow-hydrologicalfeedback, is too simple in this model. The interaction of snow-hydrology feedback with the evaporation/cumulus convection feedback may also be
an important problem to be further investigated forthis snow cover/climate issue. The remote responses
related to the ENSO and the Ice Age issues may
be . very interesting aspects of the Eurasian snow
cover, but these problems should be re-examined by
the more sophisticated coupled atmosphere/ocean
GCMs.
Acknowledgements
This work was done as a co-opearative
study
betweenthe Instituteof Geoscience,
UniversityofTsukubaand the MeteorologicalResearch
Institute.
It was also partly supportedby the Grant-in-Aid
for
Scientific
Researchby the Ministryof Education,
Scienceand Culture
No. 01460055.
The authors
would like to thankDr. T. Barnett
for his criticalcomments,
which substantially
improved
the paper.
References
Arakawa A. and W.H. Schubert,
1974: Interactionof acumulus cloud ensemble with the large-scale environment. Part I. J. Atmos. Sci., 31, 674-701.

486

Journalof theMeteorological

Asakura, T., 1968: Dynamic climatology of atmosphericcirculation over east Asia centered in Japan. Papers
Meteor. Geophys., 19, 1-68.
Barnett, T.P., 1985: Variations in near-global sea levelpres sure. J. Atmos. Sci., 42, 478-501.
Barnett, T.P., L. Dumenil, U. Schlese, E. Roeckner and
M. Latif, 1989: The effect of Eurasian snow coveron regional and global climate variations. J. Atmos.
Sci., 46, 661-685.
Blanford, H.F., 1884: On the connexion of Himalayansnowfall and seasons of drought in India. Proc. Royal
Soc. London, 37, 3-22.
Chao, B.F., W.P. O'Conner, A.T.C. Chang, D.K. Halland J.L. Foster, 1987: Snow load effect on the earth's
rotation and gravitational field, 1979-1985. J. Geophys. Res., 92, B9, 9415-9422.
Delworth, T. and S. Manabe, 1989: The influence of soil
wetness on near-surface atmospheric variability. J.
Climate, 2, 1447-1462.
Dey, B. and O.S.R.U. Bhanu Kumar, 1983: Himalayan
winter snow cover area and summer monsoon rainfall
over India. J. Geophys. Res., 88, 5471-5474.
Dickson, R.R., 1984: Eurasian snow cover versus Indian
monsoon rainfall-An extension of the Hahn-Shukla
results. J. Clim. Appl. Meteor., 23, 171-173.
Hahn, D.J. and J. Shukla, 1976: An apparent relationship between Eurasian snow cover and Indian monsoon rainfall. J. Atmos. Sci., 33, 2461-2462.
Hoskins, B.J. and D.J. Karoly, 1981: The steady linear
response on a spherical atmosphere to thermal andorographic forcing. J. Atmos. Sci., 38, 1179-1196.
Iwasaki, T., 1991: Year-to-year variation of snow coverarea in the Northern Hemisphere. J. Meteor. Soc.
Japan, 69, 209-217.
Jaeger, L., 1976: Monatskarten des Niederschlags fur
die ganzeErde. Berichte Deutsches Wetterdienstes,
Nr. 139, Offenbach, 38pp.
Kang, I.-S., 1990: Influence of zonal mean flow change
on stationary wave fluctuations. J. Atmos. Sci., 47,
141-147.
Kitoh, A. and T. Tokioka, 1986: A simulation of thetropospheric general circulation with the MRI atmospheric general circulation model. Part II: The July
performance. Pap. Meteor. Geophys., 37, 145-168.
Kitoh, A., K. Yamazaki and T. Tokioka, 1988: Influenceof soil moisture and surface albedo changes over the
African tropical rain forest on summer climate investigated with the MRI . GCM-I. J. Meteor. Soc. Japan,
66, 65-86.
Kodera, K. and M. Chiba, 1989: Western Siberian springsnow cover and east Asian June 500mb hieght. Papers Meteor. Geophys., 40, 51-54.
Kuhle, M., 1987: Subtropical mountain- and highlandglaciation as Ice Age triggers and the waning of theglacial period in the Pleistocene. Geolournal, 14.4,
393-421.
Lamb, H.H., 1977: Climate. Present, past and future.
Vol. 2 Climatic history and future. 835 pp. William
Clowers & Sons, Ltd., Beccles.
Legetes, DR., 1987: A climatology of global precipitation. Publications in Climatology, 40-1, 85 pp.
Manabe, S. and D.G. Hahn, 1977: Simulation of tropicalSocietyof JapanVol. 69, No. 4

climate of an Ice Age. J. Geophys. Res., 82, 38893911.
Matthews, E., 1983: Global vegetation and land use:
new high resolution data bases for climate studies.
J. Clim. Appl. Meteor., 22, 474-487.
Miyakoda, K. and J.-P. Chao, 1982: Essay on dynamical
long-range forecasting of atmospheric circulation. J.
Meteor. Soc. Japan, 60, 292-308.
Morinaga, Y., 1987: Interaction between snow cover andatmospheric circulation in the Northern Hemisphere.
Master thesis, Institute of Geoscience, University ofTsukuba, 55pp.
Morinaga, Y. and T. Yasunari, 1987: Interactions between the snow cover and the atmospheric circulations in the northern hemisphere. IAHS Publications
No. 166, 73-78.
Morinaga, Y. and H. Igarashi, 1991: Relationships between snow cover extent and snow mass derived from
Nimbus-7 SMMR data over the Eurasian continent.
Climatological Notes, Univ. of Tsukuba, No. 40, 101110.
Oglesby, R.J., 1990: Sensitivity of glaciatin to initialsnow cover, CO2, snow albedo, and oceanic roughness in the NCAR CCM. Climate Dynamics, 4, 219235.
Posey, J.W. and P.F. Clapp, 1964: Global distribution
f normal surface albedo. Geofis. Int., 4, 33-48.
oPrell, W.L. and J.E. Kutzbach, 1988: Monsoon variability over the past 150,000 years. J. Geophys. Res.,92, D7, 8411-8425.
Schutz, C. and LB. Bregman, 1987: Global annual snow
accumulation by months. RAND Note N-2551-AS.
Short, D.A., J.G. Mengel, T.J. Crowly, W.T. Hide and
G.R. North, 1991: Filtering of Milankovitch cycles
by earth's geography. Quaternary. Res., 35, 157173.
Smagorinsky, J., 1953: The dynamical influence of largescale heat sources and sinks on the quasi-stationary
mean motions of the atmosphere.
Quart. J. Roy.
Meteor. Soc., 79, 342-366.
Tokioka, T., K. Yamazaki, I. Yagai and A. Kitoh, 1984:
A description of the Meteorological Research Institute atmospheric general circulation model (The
MRI. GCM-I) . Tech. Rep, of the MRI. No. 13.
249pp.
Webster, P., 1982: Seasonality in the local and remote atmospheric response to sea surface temperature anomalies. J. Atmos. Sci., 39, 41-52.
Williams, J., 1975: The influence of snowcover on the atmospheric circulation and its role in climatic change:
A analysis based on results from the NCAR global
circulation model. J. Appl. Meteor., 14, 137-152.
Yamazaki, K., 1989: A study of the impact of soil moisture and surface albedo changes on global climate
using the MRI . GCM-I. J. Meteor. Soc. Japan, 67,
123-146.
Yasunari, T., 1987: Global structure of the E1
Nino/Southern Oscillation. Part II. Time evolution.
J. Meteor. Soc. Japan, 65, 81-102.
Yasunari, T., 1988: Global teleconnections associatedwith Indian monsoon and the ENSO. Meteor. Res.
Report, Univ. of Tokyo, 88-1, 30-38.

August1991

T. Yasunari,

A. Kitoh

Yasunari, T., 1990: Impact of Indian monsoon on the
coupled atmosphere/ocean system in the tropical Pacific. Meteor. * Atmos. Phys., 44, 29-41.
Yasunari, T., 1991: Monsoon and ENSO: A coupled
ocean/land/atmosphere system. TOGA Notes, No.
2, 9-13.
Yeh, T.C., R.T. Wetherald and S. Manabe, 1983: A

and

Yeh,

T. Tokioka

487

model

study

of thelogic

effectsof sudden

Wea.

Rev.,

T.C.,
effect111,

R.T.

short-term

cover

and

hydro-

removal.

Mon.

1013-1024.

Wetheraldof soil moisturehydrologychange-A

Wea.

112,

Rev.,

climaticsnow
andS. Manabe,on the short-term
numerical

1984:
climateexperiment.

The
and
Mon.

474-490.

ユーラシア大陸の積雪が北半球暖候季の大気 に与 え る影響 について
‑MRI・GCMによる数値実験‑安成哲三(筑波大学地球科学系)
鬼頭昭雄・ 時岡達志
（ 気象研究所気候研究部)

ユ ー ラ シ ア 大 陸 の 冬 の 終 わ りの積 雪 が 、 北 半 球 の 引 き続 く暖 候 季(春・夏)の大 気 状 態 に ど の よ うな 影

響 を及 ぼ す か を 、 気 象 研 究 所 の 大 気 大 循 環 モ デ ル を用 い た 数 値 実 験 に よ り調 べ た 。
三 月初 め に ユ ー ラ シ ア 大 陸 の 積 雪 域 に5cm(水わ りま で積 分 して 、 コ ン トロー ル ・ラ ン(CNTL

当 量)の 積 雪 を上 乗 せ した ラ ン(SNOWrun)を 夏 の 終run)と 比 較 す る こ とに よ り、 正 の積 雪 偏 差 が 春 か ち 夏 にか け て の 大 気 状 態 を ど の よ う に 変 え るか を評 価 し た結 果 、 以 下 の こ とが 明 か と な っ た 。
(1)春 に は 、 ア ル ベ ー ド効 果 が 、 比 較 的 低 緯 度 で 日射 の 強 い チ ベ ッ ト高 原 付 近 を中 心 に 顕 著 に現 れ 、顕・ 潜 熱 フ ラ ッ ク ス の 減 少 、 地 上 気 温 の 低 下 、 大 気 全 加 熱 量 の 減 少 を も た らす 。
(2)一 方 夏 に は 、 よ り高 緯 度 側 の 中 央 ア ジ ア・ 東 ア ジ ア の 平 原 を 中 心 に 、 融 雪 水 文 学 的 過 程 に よ る
土 壌 水 分 量 の 増 加 と地 上 気 温 の 低 下 が 卓 越 す る。 しか しな が ら 、 蒸 発 量 の 増 大 は 、 積 雲 対 流 を
活 発 化 させ る こ とに よ り、 地 表 付 近 で の 顕 熱 フ ラ ッ ク ス の 減 少 に よ る大 気 全 加 熱 量 の 減 少 を 、部 分 的 に補 償 す る役 割 を 果 し て い る 。 こ の 蒸 発 ・対 流 フ ィー ドバ ッ クは 、 しか し、 地 表 面 の よ
り湿 っ た(低温 の)状

態 を 夏 の 間 を通 して 持 続 させ る働 き も して い る 。 ア ジア モ ン ス ー ン は 、

全 体 と し て 弱 ま っ て い る。
(3)こ れ ら の 地 域 で の 大 気 の(相対 的 の)冷

源 は 、 局 地 的 な大 気 循 環 の 偏 差 を作 り出 す だ け で は な

く、 定 常 ロ ス ビー 波 の 伝 播 に よ り、 春 ・夏 を通 じて 北 太 平 洋 か ら北 米 大 陸 上 に 、 顕 著 な テ レ コ
ネ ク シ ョ ン ・パ ター ン を形 成 す る。 こ の 循 環 パ ター ン に よ り、 北 米 大 陸 部 は 低 気 圧 性 循 環 が 卓
越 し、 暖候 季 を通 じて 低 温 傾 向 が 維 持 さ れ る。
こ れ らの 結 果 が 、 氷 期 論 に 与 え る示 唆 に つ い て も、 簡 単 な議 論 を加 え た 。