Chapter 8 Response of the uk Meteorological Office General Circulation model To Sea-Surface Temperature Anomalies in the Tropical Pacific Ocean

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CHAPTER 8

RESPONSE OF THE UK METEOROLOGICAL OFFICE GENERAL CIRCULATION MODEL TO SEA-SURFACE TEMPERATURE ANOMALIES IN THE TROPICAL PACIFIC OCEAN T.N. PALMER

ABSTRACT A number of tropical Pacific SST anomaly experiments, run on the UK Meteorological Office 11level general circulation model in perpetual January mode, are described. It is found that the model's extratropical response can be statistically significant as far downstream as the European continent depending on the anomaly used, but is sensitive to the specification of orography in the model. A realistic response to a composite El NiRo SST anomaly is obtained in the extratropics provided envelope orography replaces the standard orographic specification. A negative East Pacific anomaly run revealed that aspects of the tropical response are not readily explained by linear theory and some possible reasons for this are proposed. The extratropical response appears in some experiments to be qualitatively consistent with a downstream Rossby wavetrain and this has been tested using a barotropic model. In other experiments, and in the Southern Hemisphere, this is less clear. The extent to which the extratropical response is maintained by cyclogenesis in mid-latitudes remains to be firmly established though preliminary results suggest it is important. The extratropical response is sensitive to relatively weak SST anomalies in the tropical West Pacific.

1. INTRODUCTION

Some details are given in this paper of a number of tropical Pacific SST anomaly experiments run on the UK Meteorological Office high-resolution 1I-level (grid point) GCM. The model has been described by Saker (1975), and the horizontal resolution used for the experiments described below is 23" x 3;". The GCM has a penetrative convection scheme (Lyne and Rowntree, 1976) and the version used here has zonally averaged noninteractive cloud amounts. SSTs and sea-ice are fixed throughout the integration. The integrations are summarised in Table 1. They consist of experiments run with TABLE 1 Integrations on the 11-level model Experiment

SST anomaly

Orography

a

control + Fig. l a -Fig. l a Fig. l b control Fig. l b Fig. l c

standard standard standard standard 20 envelope 2a envelope 20 envelope

b C

d e

f

g

30"

15'

0"

15~

30~5

90'

Fig. 1. (a) SST anomaly (K) for experiments b and c (December 1982); (b) SST anomaly (K) for experiments d and f (enhanced Rasmusson and Carpenter, 1980, composite, as used by Blackmon et al., 1983); and (c) SST anomaly (K) for experiment g (warm West Pacific).

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both plus and minus the observed SST anomaly for December 1982 (Fig. la); experiments with an enhanced version of the composite Rasmusson and Carpenter (1982) El Nino SST anomaly (as used by Blackmon et al., 1983; Fig. lb), both with standard and envelope orography; and an experiment with a small (idealised) warm SST anomaly in the tropical West Pacific (Fig. lc). The integrations have all been initialised with data from 28 December 1972, and run for 540 days in a perpetual January mode, except for experiment g which has, at present, been run for 90 days.

2. EXPERIMENTS

2.1. Experiments a, b and c; a strong El Nino anomaly in the East Pacific Addition of the observed SST anomaly for December 1982 (Fig. la) to climatological SST (Fig. 2) produced a strong absolute maximum in SST near 120"W. The associated anomalous rainfall is illustrated in Fig. 3a, with a maximum value of 12mm per day, lying close to this SST maximum. To the west of this maximum, 850 mb wind anomalies are westerly with strengths up to about 10 m s-l (Fig. 3b) cancelling the normal easterly trade winds and converging into the region of maximum precipitation. The observed 850 mb wind anomaly and satellite-sensed outgoing longwave radiation anomalies (OLR) are shown in fig. 4 of Rasmusson and Wallace (1983). The agreement in the general strength and direction of the wind anomalies is good, although, for the period DecemberFebruary, the time-averaged maximum OLR anomaly is positioned about 30" further west than the model's rainfall anomaly. The 540-day mean 200 mb geopotential height anomalies (in dm) for the experiments with positive and negative SST anomalies are shown in Fig. 4a and b, respectively. In Fig. 4a there are positive anomaly centres to the north and south of the rainfall anomaly maximum, at 15"N and 15"s and about 120"W. In the Northern Hemisphere there is some indication of a wavetrain with centres over the western United States, eastern Canada, and the British Isles. The response in mid-latitudes is approximately equivalent barotropic. The response in the Southern Hemisphere is somewhat different, corresponding to a dominant signal in the anomalous zonal mean flow. In Fig. 4b the anomaly field in the tropics is similar to the negative of Fig. 4a except

12O~E

180~

Fig. 2. Climatological winter SSTs (K) in the tropical Pacific.

120.w

60'W

86

W

I 180'

ISO'W

1

. I

120'

1

I

PO'

1 1

I 60'

I I 3( W

Fig. 3. (a) convective rain anomaly (mm/day); and (b) 850 mb wind anomaly for experiments b-a (December 1982).

that the cyclonic centres are nearly 15 degrees to the west of the corresponding anticyclonic centres, as is the maximum negative rainfall anomaly relative to the corresponding maximum for the positive SST anomaly experiment. It is interesting to note that, whereas the magnitude of the anomalous cyclonic centres is similar to that of the corresponding anticyclonic centres, the magnitude of the anomalous precipitation is not, with the negative rainfall anomaly (not illustrated) some two times smaller than the corresponding positive anomaly. In mid-latitudes, the response t o the negative SST anomaly is quite uncorrelated with the pattern in Fig. 4a. In fact, the pattern in Fig. 4b is qualitatively similar to the Pacific/North American (PNA) teleconnection pattern, discussed by Horel and Wallace (19811, with anomaly centres over the Aleutian Islands, Hudson Bay and the southeastern states of America. WhiIst the response in Fig. 4a does not show a strong PNA teleconnection pattern characteristic of the observed DecemberFebruary 200 mb height anomaly composite (Quiroz, 1983), the December 200 mb height anomaly did not itself show a strong PNA pattern. Preliminary results from a recent experiment with the observed December-February 1982/83 composite SST anomaly (see Quiroz, 1983), and envelope orography (see Section 2.2) show a stronger

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PNA pattern which agrees well with observations. It is interesting to note in this respect that the December-February composite SST anomaly was weaker, and positioned further west than the December only SST anomaly (see Quiroz, 1983). In order t o test that the difference fields illustrated in Fig. 4 are statistically significant, t-statistics of the six 90-day mean fields have been produced, treating them as independent samples. Of course this latter assumption cannot be completely justified. Recent tests (D.A. Mansfield, pers. commun., 1984) have shown that a GCM’s evolution depends strongly on its initial conditions up to about day 15 of an integration. Hence, only about 2 / 3 of a given 90-day period can be thought of as essentially independent of its neighbouring periods. A second problem in relating a model-generated t-statistic to the atmosphere, concerns the model’s low-frequency variability. The model’s low-frequency variability is smaller than corresponding atomospheric values so for a given size of anomaly, the model may show stronger significance than the atmosphere. The t-statistics for the fields shown in Fig. 4a and b are illustrated in Fig. 5a and b respectively. With each experiment comprising six 90-day periods, there are ten degrees of freedom in the t-test. A two-sided significance of 1% with this number of degrees of freedom corresponds to a value of about three. (With half this number of degrees of freedom the corresponding value is about four.) In Fig. 5a, in the Northern Hemisphere, the low centres over the United States and the British Isles are strongly significant. The tropical high centres are also significant at 1%, though the high over the Canadian east coast is not. Much of the southern hemispheric response is significant at 1%. In Fig. 5b the tropical low centres and the Northern Hemisphere mid-Pacific high are strongly significant. The high/low dipole downstream of these anomaly centres is barely significant. Parts of the Southern Hemisphere response also appear to be significant at 1%. Bearing in mind the caveats discussed above, it appears that with such large SST anomalies, the atmospheric response can be strongly significant, not only over the Pacific, but further downstream as far as the European continent. In discussing reasons for the asymmetry between the positive and negative SST anomaly experiments, it is important to bear in mind that the anomaly fields do not represent differences from a zonally averaged basic state. In the control integration SSTs are highest in the tropical West Pacific. Associated with this there is large-scale ascent and a precipitation maximum over the West Pacific with general descent over the East Pacific, forming part of the so-called Walker circulation. A decrease in SST east of the dateline may not substantially change the rainfall field in the East Pacific, as it is already small. The largest rainfall differences are likely to occur near the western edge of such a negative SST anomaly where its addition may reverse the sense of large-scale vertical motion. On the other hand, a large increase in SST in the East Pacific may completely reverse the Walker Circulation in the East Pacific and give rise to substantial anomalous convection over the SST anomaly. Hence, the longitudinal position of the latent heating anomaly would not be expected to be identical for positive and negative SST anomaly fields. As we have noted above (see also Sections 2.2 and 2.3), the extratropical PNA pattern appears to be more readily excited the further west, over the Pacific, the latent heating anomaly occurs.

Fig. 5. (a) t-statistics of 200 mb height anomaly for experiments b-a; and (b) t-statistics of 200 mb height anomaly for experiments c-a.

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2.2. Experiments d-f; Rasmusson and Chrpenter composite El Niiio anomaly The convective precipitation associated with this SST anomaly (Fig. Ib) is illustrated in Fig. 6. Maximum values occur near the dateline, again consistent with the position of warmest waters in the anomaly experiment. The tropical 200 mb wind anomalies are illustrated in Fig. 7, for day 1 (Fig. 7a), for days 1-10 averaged (Fig. 7b), and for days 1-540 (Fig. 7c). It can be seen that the westerly anomalies over Africa and the Indian Ocean spread eastward from the SST anomaly. The easterlies that spread westwards, on the other hand, appear to be weak and possibly contained within the West Pacific. The anticyclone pairs are positioned to the east of the precipitation anomaly. The 200mb geopotential height anomaly, illustrated in Fig. 8a, shows in the PNA region a deep low over the Gulf of Alaska and a weak ridge over eastern North America. This pattern does not compare particularly well with the PNA pattern, or the response to a similar SST anomaly, reported by Blackmon et al. (1983), or Shukla and Wallace (1983). In common with many other high-resolution GCMs, the 11-level model develops a winter-time climate drift with polar temperatures several degrees colder than observed. Associated with this, zonal winds are too strong, and baroclinic waves are not sufficiently damped over land. These systematic errors affect the results of our sensitivity experiments. However, Wallace et al. (1983) have shown that a respecification of orography in a GCM can substantially alleviate these errors (although other effects due to sub-grid scale orographically induced gravity wave drag in the upper troposphere and lower stratosphere may also help alleviate these errors; see Palmer and Shutts, 1984). The 200 mb geopotential height anomaly for the experiment with the enhanced Rasmusson and Carpenter SST anomaly, and twice standard deviation (213) envelope orography (minus control with 20 envelope orography) is shown in Fig. 8b. (With 20 envelope orography, the model orographic heights are enhanced by twice the subgridscale standard deviation of orographic height, as determined by a ten-minute resolution specification of orography; Wallace et al., 1983.) Figures 8a and b show substantial differences in the extratropical Northern Hemisphere. With envelope orography there is a low at about 45'N, 160°W, a high over central Canada, and a low to the east of the U.S.A. This pattern conforms more closely with the PNA teleconnection pattern, and Blackmon et al.'s response. The notion that the correct positioning of storm-track activity is important in obtaining a realistic time-mean response to an SST anomaly is suggested by F i g . 9 h i c h shows the high-pass filtered anomalous eddy streamfunction forcing -V-' [V-(v'(')](using a poor man's filter (Lorenz, 1979) with three-day variances) for two 90-day periods of the composite SST anomaly runs. Figure 9a, for the standard orography runs, shows anomalous cyclonic eddy forcing in excess of 15 m2 s-' (shaded) extending across the Pacific into the U.S.A. (an eddy forcing of - 1 0 m 2s-' would spin-up a time-mean cyclonic streamfunction of - 1 x lo6 m2 s-l in about one day). For the envelope orography experiments, there is strong cyclonic forcing in the mid-Pacific near 165"W, though in the East Pacific, the eddy forcing is anticyclonic. The differences between Figs. 9a and b appear to be consistent with the difference in the positions of the negative 200mb height anomalies over the North Pacific, in Figs. 8a and b. The sensitivity of the response to different orographic specifications highlights the

Fig. 6. Convective rain anomaly (mmlday) for experiment d-a (RC composite).

Fig. 7. 200 m b wind anomaly for experiments d-a (RC composite); (a) day 1; (b) days 1-10; and (c) days 1-540.

Fig. 8. 200rnb geopotential height anomaly (drn); (a) experiments d-a (RC composite, standard orography); and (b) experiments f-e (RC composite, envelope orography).

Fig. 9. Anomalous high-pass eddy streamfunction forcing at 250 mb (m2 s-’) over a 90-day period (values less than - 15 mz s-* are shaded); (a) experiments d-a (RC composite, standard orography); and (b) experiments f-e (RC composite, envelope orography).

100

Fig. 10. 500 mb geopotential height anomaly (dm) for experimentsg-e (warm West Pacific).

importance of the model's basic climatology in the determination of the impact in the extratropics of tropical SST forcing.

2.3. Experiments g and e; the response to a warm tropical West Pacqic SST anomaly The precipitation anomaly associated with the warm West Pacific SST anomaly has positive values near the SST anomaly with a maximum of 10mm per day at about 5"N, 155"'. Figure 10 illustrates the 500mb geopotential height anomaly. There is a high centre at 30"N, 155"E, to the north of the SST anomaly, and a downstream response with centres just east of the dateline, and over Alaska. Note also the anomalous easterly flow in high latitudes. This pattern is remarkably similar to that occurring in the severe El Nino winter of 1976/77. Palmer and Mansfield (1984) have argued that the difference between the extratropical response to the 1976/77 El Niiio event, and the response to the 1972/73 or 1982/83 El Nino events may be due to small differences in SST in the

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tropical West Pacific. (An analysis of SST in the tropical West Pacific in January 1977, from Meteorological Office historical SST archives, showed a pool of anomalously warm water in excess of 1 K; see Palmer and Mansfield, 1984.) The response to a warm West Pacific anomaly reported in Palmer and Mansfield is similar to that shown here except that the distance between anomaly centres over the Pacific is less here. This is consistent with the reduction in stationary Rossby wavelength brought about by the influence of envelope orography in weakening the zonal mean flow. Integration g is continuing to 540 days. 3. COMPARISON OF RESULTS WITH LINEAR THEORY

3.1. Tropical response Gill’s (1980) model (see also Heckley and Gill, 1984) describes the linear response of the tropical atmosphere to a diabatic heating anomaly. This type of model has been used to provide the atmospheric component of tropical coupled atmosphere/ocean models, so it is clearly important to understand the extent to which it can describe the atmospheric response to a fixed SST anomaly. The time evolution of the tropical 200mb wind anomaly associated with the Rasmusson and Carpenter SST anomaly was illustrated in Fig. 7. The propagation of a westerly mode with small meridional wind component, to the east of the precipitation anomaly, is consistent with the Kelvin mode component of Gill’s solution. At 850 mb this mode is much weaker, however (see, e.g., Fig. 3b); whereas the low-level linear solution is equal and opposite to the upper-level solution. The time evolution shown at 200 mb also shows a weaker, and less extensive mode to the west of the precipitation anomaly. The easterly wind anomalies do not appear to extend into the Indian Ocean. Furthermore, the anticyclonic anomalies are located to the northeast and southeast of the precipitation anomalies (cf. Figs. 6 and 8). All these facts suggest that the linear tropical Rossby mode response may not be simulated in the GCM. In the linear model the anticyclonic doublet is part of the Rossby mode propagating to the west of the diabatic heating, A related problem of accounting for the tropical GCM response in terms of linear theory is posed by results from the positive and negative 1982 El Nino anomaly experiments. As discussed, the magnitude of the anomalous 200 mb cyclonic doublet associated with the negative SST anomaly experiment is equal to the corresponding anticyclonic doublet associated with the positive SST anomaly experiment, despite the fact that the magnitude of the anomalous precipitation was two to three times smaller for the negative experiment. It is possible, therefore, that nonlinear modifications to Gill’s model may be essential. For example, the forcing in his model is provided (with standard notation) by the linearised form, -f V- v, of the divergence forcing - (f + {)V. v. In the tropics where f is small, {-f, providing a basic asymmetry between positive and negative anomalous diabatic heating. For negative forcing, 5 will reinforce f to enhance the linearised effect; conversely for positive forcing { will cancel f to reduce the linearised effect (see also the discussion in Section 3.2). The importance of nonlinearity in the vorticity budget of the

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real atmosphere during the 1982/83 winter has been stressed recently by Sardeshmukh and Hoskins (1984). A second effect which may be relevant in explaining the apparent asymmetry between Kelvin and Rossby modes is the zonal inhomogeneity of the basic state flow. In particular it is found that the upper level absolute vorticity is small over the tropical West Pacific in the control run with very weak meridional gradient north of the equator. It is possible that propagation of the equatorial Rossby mode over the West Pacific was inhibited by the presence of this weak gradient in the control integration.

3.2. Extratropical response The notion of equivalent barotropic tropically forced Rossby wavetrains propagating into the extratropics has been developed by Hoskins and Karoly (1981) and Webster (1981) and extended by Simmons (1982) and Simmons et al. (1983), to explain the extratropical response to anomalous diabatic heating. Linearised about a zonally symmetric superrotational basic state, the theory predicts that these wavetrains should propagate on great circle paths. It is clear that the GCM experiments by no means provide unequivocal support for the notion that the wavetrain theory accounts for the major extratropical response. Whilst results from the West Pacific anomaly experiment (Section 2.3) and the positive December 1982 El Niiio anomaly experiment (Section 2.1) suggest a wavetrain-like response into the Northern Hemisphere, results from the other experiments (and the Southern Hemisphere responses) are less clear. One component of the general circulation which is likely to be essential in accounting for the extratropical response is the (nonlinear) interaction of the transient eddies (extratropical cyclones for example) with the stationary flow, as discussed by Opsteegh and Vernekar (1981), and Kok and Opsteegh (1984). For example, one important systematic error in the climatology of the model with standard orography is the inability to fill baroclinic lows over land. This effect is largely removed by the introduction of envelope orography. It was commented in Section 2.2 that the introduction of envelope orography into the Rasmusson and Carpenter SST anomaly experiments tended to position the low anomaly centre further west, moving it from the west coast of Canada to the Aleutian Islands. It is quite possible that this resulted from eddy mean-flow interaction with the correct positioning of storm tracks in runs which included envelope orography (see Fig. 9). The remainder of the PNA pattern over the North American continent may correspond to an enhanced ridge and downstream trough forced by the strengthened mean-flow upstream over the Pacific flowing over (the envelope of) the Rockies. On the other hand, results using the barotropic model described by Simmons (1982) suggest that the forced Rossby mode response can reproduce the qualitative extratropical pattern of the GCM response to at least one SST experiment. Figure l l a shows the GCM’s anomalous 200 mb streamfunction response to the positive December 1982 El Nino anomaly (Fig. la). Figure 1 Ib shows the barotropic model’s response to (an idealisation of) the GCM’s anomalous 200mb divergence over the tropical East Pacific. The barotropic model’s basic state was provided by the (540-day mean) 200 mb streamfunction of the control run (a). A five-day damping was applied so that the forced solution reached a steady state, and no unstable modes were allowed to grow. The

I03

m

Fig. 11. (a) 200mb streamfunction anomaly for experiments b-4 (+ December 1982); and (b) streamfunction response from barotropic model forced by the 200 mb divergence anomaly in experiments b-a over the tropical East Pacific.

180"

150"w

120"

90"W

150'W

120"

90"W

30"N

0"

30's

180'

Fig. 12. (a) anomaly of V-v at 200 mb (contour interval 0.5 X anomaly of (ft f ) V . v / 2 n at 200mb (contour interval 0.2 X

s-') for experiments b-a; and (b)

for experiments b-a.

SKI)

105

qualitative correspondence between Figs. 1l a and 1l b is good. Note, for example, the low centres over the United States and the British Isles in both experiments. One should treat this result with caution, however. As we have discussed above, the anomalous nonlinear divergence forcing ( f + c)V. v may be considerably weaker than the linear counterpart used to force the barotropic model. For example, Fig. 12 shows the anomalous values of V - v and ( f + <)V-v/2i2 for the positive December 1982 anomaly experiment at 200 mb. Whereas the anomalous divergence clearly shows the effect of the tropical SST anomaly, contours of (f+ {)V- v / 2 n do not. It is possible that the tropical height anomaly is in a (nonlinear modon-like) balanced state, with diabatic heating acting to offset radiative and other nonconservative effects, but not acting as a direct forcing for the nondivergent flow. As before, anomalous synoptic-scale eddies to the north may provide a substantial forcing to the time-averaged mid-latitude downstream response. In this respect, it is interesting to speculate that growth of midlatitude cylones over the East Pacific may be directly affected by the enhanced moisture field in the tropical Pacific as a result of the SST anomaly. Work is now in progress analysing the present model runs to establish the interplay between these different components of the general circulation. 6 . CONCLUSIONS

In this paper we have discussed results from experiments designed to test the sensitivity of the wintertime climatology of the UK Meteorological Office general circulation model to sea-surface temperature anomalies in the tropical Pacific Ocean. The following summarises the conclusions of this paper. (1) The GCM extratropical response is more sensitive to SST anomalies in the tropical West Pacific than in the tropical East Pacific. Small differences (- 1 K) in SST in the tropical West Pacific may explain the variability of the response from one El Nino event to another (cf. Palmer and Mansfield, 1984). (2) The response in the model t o a large SST anomaly in the East Pacific can be statistically significant, not only over the Pacific Ocean, but also downstream as far as the European continent. (3)The model’s response to a tropical SST anomaly does depend strongly on the model’s basic climatology. With standard orographic specification, the Northern Hemisphere wintertime flow is too strong and baroclinic waves are not damped sufficiently over land. By altering the orographic specification in such a way as to reduce the strength of the flow and position more correctly the extratropical storm tracks, the response to a composite El Nino anomaly corresponds more closely with observed atmospheric anomaly patterns. (4)Certain aspects of the model’s tropical response t o an SST anomaly cannot be readily explained by linear theory on a zonally homogeneous basic state. Possible modifications of such theory to account for the model results may include the relative vorticity contribution, {, to the divergence forcing (f+{)V*V, and the imposition of weak basic state absolute vorticity gradients over the tropical West Pacific. (5) The anomalous tropical divergence forcing from one experiment has been used to generate a downstream wavetrain response in a linear barotropic model. The results

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compare favourably with the GCM response, both in the subtropics and in mid-latitudes. However, the forcing f V * v used in the linear model appears to be substantially larger than the nonlinear forcing (f + {)V. v. It is possible that enhanced cyclogenesis to the north of the SST anomaly may act to maintain the mid-latitude time-averaged response. A more detailed account of these, and other experiments is in preparation. ACKNOWLEDGEMENTS

My thanks to Dr. A.J. Simmons for help in carrying out the integration described in Section 3, to Dr. D.A. Mansfield for considerable assistance (in producing the eddy diagnostics in particular) and to Dr. G.J. Shutts for helpful discussions. This research was partially supported by the EEC contract, Number CLI-034-8 lUK(H). REFERENCES Blackmon, M.L., Geisler, J.E. and Pitcher, E.J., 1983. A general circulation model study of January climate anomaly patterns associated with interannual variations of equatorial Pacific sea surface temperatures. J . Atmos. Sci., 40: 1410-1425. Gill, A.E., 1980. Some simple solutions for heat-induced tropical circulation. Q. J.R. Meteorol. SOC., 106: 447-462. Heckley, W.A. and Gill, A.E., 1984. Some simple analytical solutions to the problem of forced equatorial long waves. Q. J.R. Meteorol. SOC.,110: 203-218. Horel, J.D. and Wallace, J.M., 198 1. Planetary-scale atmospheric phenomena associated with the Southern Oscillation. Mon. Weather Rev., 109: 813-829. Hoskins, B.J. and Karoly, D.J., 1981. The steady linear response of a spherical atmosphere to thermal and orographic forcing. J. Atmos. Sci., 38: 1179-1196. Kok, K. and Opsteegh, J.D., 1985. On the possible causes of anomalies in seasonal mean circulation patterns during the 1982/83 El Nino event. J. Atmos. Sci. (in press). Lorenz, E.N., 1979. Forced and free variations of weather and climate. J . Atmos. Sci., 36: 13671376. Lyne, W.H. and Rowntree, P.R., 1976. Development of a convective parameterization using GATE data. Meteorol. Off. 20 Technical Note 11/70. UK Meteorological Office, Bracknell. Opsteegh, J.D. and Vernekar, A.D., 1982. A simulation of the January standing wave pattern including the effects of transient eddies. J. Atmos. Sci., 39: 734-744. Palmer, T.N. and Mansfield, D.A., 1984. Response of two atmospheric general circulation models to sea-surface temperature anomalies in the tropical East and West Pacific. Nature, 310: 483485. Palmer, T.N. and Shutts, G.J., 1984. Preliminary results of the effect of a parameterization of gravity wave drag in the Meteorological Office 11-layer operational model. Meteorol. Off. 13 Branch Memorandum 147, UK Meteorological Office, Bracknell. Quiroz, R.S., 1983. The climate of the “El Nino” Winter of 1982-83. A season of extraordinary climate anomalies. Mon. Weather Rev., 111: 1685-1706. Rasmusson, E. and Carpenter, T., 1982. Variations in tropical sea surface temperature and surface wind fields associated with the Southern Oscillation/El Nino. Mon. Weather Rev., 110: 354384. Rasmusson, E. and Wallace, J.M., 1983. Meteorological aspects of the El Niiio/Southern Oscillation. Science, 222: 1195-1202. Rowntree, P.R., 1972. The influence of tropical east Pacific Ocean temperatures on the atmosphere. Q. J.R. Meteorol. SOC.,98: 290-321.

107 Saker, N.J., 1975. An 11-layer general circulation model. Meteorol. Off. 20 Technical Note 11/30, UK Meteorological Office, Bracknell. Sardeshmukh, P.D. and Hoskins, B.J., 1985. Vorticity Balances in the Tropics during the 1982-3 El-Nifio Southern Oscillation event. Q. J. R. Meteorol. SOC.(in press). Shukla, J. and Wallace, J.M., 1983. Numerical simulation of the atmospheric response to equatorial Pacific sea surface temperature anomalies. J. Atmos. Sci., 40: 1613-1630. Simmons, A.J., 1982. The forcing of stationary wave motion by tropical diabatic heating. Q. J. R. Meteorol. SOC.,108: 503-534. Simmons, A.J., Wallace, J.M. and Branstator, G . , 1983. Barotropic wave propagation and instability and atmospheric teleconnection patterns. J. Atmos. Sci., 40: 1363-1392. Wallace, J.M., Tibaldi, S. and Simmons, A.J., 1983. Reduction of systematic forecast errors in the ECMWF model through the introduction of an envelope orography. Q. J.R. Meteorol. SOC., 109: 683-718. Webster, P.J., 1981. Mechanisms determining the atmospheric response to sea surface temperature anomalies. J. Atmos. Sci., 38: 554-571.