INTERNATIONAL JOURNAL OF CLIMATOLOGY

Int. J. Climatol. 23: 143–155 (2003)

Published online in Wiley InterScience (www.interscience.wiley.com). DOI: 10.1002/joc.868

SEASONAL MODULATION OF THE EL NINO–SOUTHERN OSCILLATIONRELATIONSHIP WITH SEA LEVEL PRESSURE ANOMALIES OVER THE

NORTH ATLANTIC IN OCTOBER–MARCH 1873–1996

V. MORON,a,b,* and I. GOUIRANDa,b

a UFR des Sciences geographiques et de l’Amenagement, Universite de Provence (Aix-Marseille I), Franceb CEREGE, UMR–CNRS 6635, France

Received 10 January 2002Revised 13 September 2002

Accepted 15 September 2002

ABSTRACT

The seasonal modulation of the relationship between the sea-level pressure anomalies (SLPAS) over the North Atlanticregion (100 °W–50 °E; 20–70 °N) and the sea surface temperature anomalies (SSTAS) in the tropical Pacific (120–290 °E;20 °N–20 °S) is investigated in the northern winter (October to March) from 1873 to 1996, using singular-valuedecomposition and composite analyses. Both methods show that the pattern of the North Atlantic SLPA associated withthe tropical Pacific SSTA in November–December is different from that found in January–March. The surface covered bya significant SLPA is larger in November–December and February–March than in January. In November–December, thewarm El Nino–southern oscillation (ENSO) events are associated with negative SLPAS extending from the Hudson Bay toScandinavia and positive SLPAS over the Azores high. The cold ENSO events are associated with a positive SLPA betweenGreenland and western Europe. In January, and mainly in February–March, the warm ENSO events are associated witha positive SLPA north of 50 °N and a negative SLPA extending from the southeastern USA toward western and centralEurope. The cold ENSO events exhibit almost reversed SLPA patterns. The change between November–December andJanuary–March is also observed at the hemispheric scale. In November–December, the SLPAS associated with the warmminus cold ENSO composite form a hemispheric north–south dipole pattern with positive (negative) anomalies south(north) of 40–45 °N. In January–March, the SLPA pattern associated with the warm minus cold ENSO composite isclose to the tropical Northern-Hemisphere pattern. Copyright 2003 Royal Meteorological Society.

KEY WORDS: El Nino–southern oscillation; Tropical Pacific sea surface temperature anomalies; sea-level pressure anomalies; NorthAtlantic; Europe; monthly time scale

1. INTRODUCTION

El Nino–southern oscillation (ENSO) is the major coupled air–sea phenomenon on the 2–8 year time scale(Philander, 1990). This coupled mode is characterized by a seesaw in sea-level pressure (SLP) between theeastern and western tropical Pacific associated with extremes in sea surface temperature (SST) anomalies(SSTAs) in the central and eastern equatorial Pacific (Rasmusson and Carpenter, 1982). Warm ENSO eventscorrespond to one phase of the phenomenon during which the pressure difference across the tropical Pacific isreduced and the SSTAs are positive in the central to eastern tropical Pacific (Philander, 1990). The contrastingcold ENSO events have an enhanced SLP gradient across the tropical Pacific and negative SSTAs in the centralto eastern tropical Pacific. These changes are associated with global-scale climatic anomalies through the wholetropical zone, and often over much of the extratropics, mainly around the North Pacific (e.g. Ropelewski andHalpert, 1987, 1996; Hamilton, 1988; Kiladis and Diaz, 1989; Halpert and Ropelewski, 1992; Yulaeva andWallace, 1994; Diaz et al., 2001).

* Correspondence to: V. Moron, CEREGE, Europole mediterranean de l’Arbois, BP 80, 13545 Aix en Provence, France;e-mail: moron@cerege.fr

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144 V. MORON AND I. GOUIRAND

The relationships between the ENSO phenomenon and climatic anomalies across the North Atlantic andEurope are weak (Rogers, 1984; Fraedrich, 1990, 1993; Fraedrich and Muller, 1992; Fraedrich et al., 1992;Moron and Ward, 1998; Pozo-Vazquez et al., 2001) and appear to have been unstable throughout the 20thcentury (Rodo et al., 1997; van Oldenborg et al., 2000). Previous studies have claimed to have identified aweak, albeit statistically significant, SLP anomaly (SLPA) pattern in December–February. During the coldENSO events, this SLPA pattern is close to the positive phase of the North Atlantic oscillation (NAO) (Pozo-Vazquez et al., 2001) and is associated with a poleward shift of the Atlantic storm track (May and Bengtsson,1998) and more anticyclonic days over western and central Europe (Fraedrich, 1990, 1993). During the warmENSO events, the SLPAs are less significant (Pozo-Vazquez et al., 2001) with negative SLPAs stretchingfrom the southeastern USA to Europe, an equatorward shift of the Atlantic storm track (May and Bengtsson,1998) and more frequent cyclonic days over western and central Europe (Fraedrich, 1990, 1993). Most ofthe studies mentioned above consider the relationship for fixed 3 month (December–February for the borealwinter) seasons. Exceptions are Fraedrich (1990) and Laıta and Grimalt (1997), who used the two months ofNovember–December and January–February for the northern winter. The arbitrary pooling of months togetherinto a priori seasons can miss genuine ENSO influences (Gouirand and Moron, 2000). This possibility israised by Huang et al. (1998), who demonstrated that the sign of the NAO − ENSO connection changesbetween December and January.

This paper focuses on the seasonal modulation of the ENSO relationships with SLPAs across the NorthAtlantic area during the boreal winter. Our strategy is based on the study of the SLPAs associated withwarm and cold ENSO events at the monthly time scale. We analyse carefully the consistency between theconsecutive months to avoid possible mixing of different patterns. Section 2 presents the data and methodsused. The results are presented in Section 3. Singular value decomposition analysis (SVDA) is used to studythe linear relationship between the tropical Pacific SSTAs and the North Atlantic SLPAs. Composites of NorthAtlantic SLPAs for warm and cold ENSO events are analysed to reveal a possible non-linear response of theNorth Atlantic atmospheric circulation. A summary and discussion of the results (Section 4) close the paper.

2. DATA AND METHODS

2.1. Data

The data sets used in this study span the 1873–1996 winter period (October–March; the year refers tothe months of January–March). For SLP, the study employs the Northern Hemisphere regular 5° × 5° griddata set of Basnett and Parker (1997), used recently by Pozo-Vazquez et al. (2001). Monthly SLPA fields arecomputed for each grid point as departures from the mean of the whole 1873–1996 period. Missing SLPAvalues are then replaced by zero. The NAO index used in this study is generated from the Azores minusIceland standardized SLPA. SST data are taken from the KAPLAN data set (Kaplan et al., 1998), availableon a 5° × 5° grid since January 1856.

2.2. Methods

SVDA is used to extract the first mode that maximized the covariance between two climatic fields (Brether-ton et al., 1992). The ‘left’ field (matrix L) is the monthly tropical Pacific (120 °E–90 °W; 20 °N–20 °S) SSTA,and the ‘right’ field (matrix R) is the monthly North Atlantic (100 °W–50 °E; 20–70 °N) SLPA. In the fol-lowing, the vectors and matrices are indicated in boldface and the scalars are in italic. The matrix L oforder Nl × T (with Nl columns and T rows) and the matrix R of order Nr × T are firstly normalized to zeromean and unit variance and weighted by the square root of the cosine of latitude. Let 〈f(t)〉 denote the timeaverage of a time series f(t) over the T observation times, the covariance matrix Clr is defined as (Brethertonet al., 1992).

Clr = 〈LT × R〉Copyright 2003 Royal Meteorological Society Int. J. Climatol. 23: 143–155 (2003)

ENSO NORTH ATLANTIC SEA LEVEL PRESSURE RELATIONSHIP 145

The exponent T denotes the transpose of the matrix and L × R is the matrix product of L and R. Thenthe covariance matrix C of order Nl × Nr is then decomposed using the singular value decomposition(Bretherton et al., 1992):

Clr = U × � × VT

where the Nl × Nl matrix U (Nr × Nr matrix V) is an orthonormal set of vectors called the left (right) singularvectors, and � is a diagonal matrix containing the singular values. There are S nonnegative (S ≤ min(Nl, Nr))singular values ordered in decreasing magnitude. For each pair of patterns, the expansion coefficients A (orderNl × T ) and B (order Nr × T ) are calculated as weighted linear combinations of the grid-point data:

A = L × U

B = R × V

The spatial structure of the first SVDA mode is illustrated here by the homogeneous correlation patterns (i.e.the correlation between the expansion coefficients and the original data). The strength of coupling betweenL and R is indicated by the total squared covariance fraction (SCF, Bretherton et al., 1992) and the temporalcorrelations (TCs) between their expansion coefficients. The significance of TC is assessed with the ‘random-phase’ test (Ebizusaki, 1997). Two Fourier series, with random phases and the same power spectrum as theoriginal series, are created 1000 times and the significance of the observed TC is then determined from thedistribution of these 1000 correlations. In the following, one, two and three asterisks indicate a significantTC at the two-sided 80%, 90% and 95% levels respectively.

Composite analyses are used to check the robustness of the SVDA results and also to reveal a possible non-linear response of the North Atlantic SLPA to ENSO events (Montroy et al., 1998; von Storch and Zwiers,1998). Warm and cold ENSO events are defined with the first empirical orthogonal function (EOF) and theassociated principal component (PC) of the seasonal (October–March) SSTA in the tropical Pacific, whichexplains 67% of the total variance (not shown). This EOF reflects traditional ENSO events (i.e. Rasmussonand Carpenter, 1982; Philander, 1990). Warm (cold) ENSO events are identified as the 10, 20, 30 highest(lowest) scores of the first PC (Table I). We then construct the composite 6 month temporal evolution ofNorth Atlantic SLPA for the warm (cold) ENSO events during 1873–1996. The difference between theSLPAs associated with the warm ENSO events and those associated with the cold ones is computed andcalled hereafter the warm minus cold ENSO composite.

To assess the significance of the composite fields, a local Student’s t-test is applied to the warm andcold ENSO composites for each SLPA grid-point, to ascertain which values are significantly different fromzero at the two-sided 90% level. The 124 monthly SLPA fields are then randomly reshuffled 1000 times,the composites recomputed, and local Student’s t-tests reapplied to the resulting SLPA composite means.The field significance is then assessed as the percentage of reshufflings that have larger significant area thanthe observed composite. The same procedure is used to assess the field significance of the warm minus cold

Table I. Years that are included in the 10, 20 and 30 year warm and coldcomposites (the year refers to the months of January–March). The boldunderlined dates are the ten warmest and ten coldest years. The bold dates

are the 11–20 warmest and coldest years

Warm Cold

1878 1889 1897 1900 1903 19051906 1912 1914 1915 1919 19201926 1931 1940 1941 1942 19581966 1969 1970 1973 1977 19801983 1987 1988 1992 1993 1995

1874 1875 1876 1880 1887 18901893 1894 1904 1909 1910 19111917 1918 1921 1925 1934 19391943 1950 1951 1955 1956 19651968 1971 1974 1975 1976 1989

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146 V. MORON AND I. GOUIRAND

ENSO composite. In that case, the local Student’s t-tests are applied to ascertain which values are significantlydifferent at the two-sided 90% level.

The consistency between composites maps is measured by pattern correlation (PC hereafter). Thesignificance of PC is computed in the following way. For example, to assess the significance of the PCbetween the 30-year warm ENSO composite of November and December: the 64 fields of November andDecember SLPA, which belong neither to the warm nor to the cold ENSO composites, are reshuffled exactlyin the same way 1000 times. Two different 30 year samples are extracted at each step and their means arecorrelated. In each case, the selected months of November and December belong to the same 30 winterseasons. The significance of the observed PC between the 30 year warm ENSO composite of November andthat of December is then determined from the distribution of the 1000 random PCs. In the following, one,two and three asterisks indicate a significant PC at the two-sided 80%, 90% and 95% levels.

3. SEASONAL MODULATION OF ENSO RELATIONSHIP WITH NORTH ATLANTIC SLPA

3.1. SVDA between the tropical Pacific SSTA and the North Atlantic SLPA

The SCF of the first SVDA mode (Figure 1) equals 44%, 54%, 61%, 52%, 70% and 62% and the monthlyTC between the pair of expansion coefficients reached values of 0.41***, 0.40***, 0.49***, 0.35***, 0.47***,0.51*** from October to March.

The first SVDA mode of tropical Pacific SSTA is broadly similar across the whole winter (Figure 1(a), (c),(e), (g), (i) and (k)), with strong positive values over the central and eastern tropical Pacific and weak negativeones over the western tropical Pacific, resembling the typical ‘mature’ phase of ENSO events (Rasmussonand Carpenter, 1982). This pattern is almost identical to the leading EOF of the seasonal tropical PacificSSTA used to define the extreme cold and warm ENSO events (not shown).

The first SVDA mode of North Atlantic SLPA (Figure 1(b), (d), (f), (h), (j) and (l)) changes during thewinter. In November–December (Figure 1(d) and (f)), the SSTAs in the central-eastern tropical Pacific arecorrelated negatively with SLPA over the Icelandic low and positively with those over the Azores high. Inother words, warm (cold) ENSO events are synchronous with stronger (weaker) than usual westerlies acrossthe central Atlantic. The TCs between the expansion coefficients of the leading SVDA mode of the NorthAtlantic SLPA (tropical Pacific SSTA) and the NAO index equal 0.69*** and 0.76*** (0.18** and 0.28***)in November and December respectively.

The SLPA pattern is almost reversed from January (Figure 1(h)) with a westward shift of the highest absolutehomogeneous correlations. In February–March (Figure 1(j) and (l)), the lowest homogeneous correlations arelocated off the southeast US coast and the highest ones extend themselves from Baffin Island to Greenland.The TCs between the expansion coefficients of the first SVDA mode of the North Atlantic SLPA (tropicalPacific SSTA) and the NAO index equal −0.85***, −0.70***, −0.44*** (−0.22***, −0.14*, 0.02) fromJanuary to March respectively. So, Figure 1 suggests that the North Atlantic region experiences weaker(stronger) than usual westerlies in January–March during warm (cold) ENSO events.

3.2. Amplitude of SLPA and dependence on polarity of tropical Pacific SSTA

The SVDA of Section 3.1 could only reveal linear relationships between the atmospheric circulation overthe North Atlantic area and the tropical Pacific SSTA variability. In this section, the amplitude and degreeof linearity of the North Atlantic SLPA associated with warm, cold and warm minus cold ENSO compositesare studied, using a composite analysis.

Table II presents the percentage of the area locally significant at the two-sided 90% level and the fieldsignificance of the composites for each month. The field significance is never reached in October (exceptfor the 30 year warm minus cold ENSO composite). The results are almost independent of the size of thecomposites (Table II). The PCs between the three samples (i.e. the 10, 20, and 30 years) for a given compositeare usually above 0.5* and more frequently above 0.7**. The PCs between the three samples are weakestin January and particularly in October (not shown). The area, locally significant at the two-sided 90% level,

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Figure 1. Homogeneous correlation of the first SVDA mode of the tropical Pacific SSTA field (a,c,e,g,i,k) and the North Atlantic SLPAfield (b,d,f,h,j,l) in October (a,b), November (c,d), December (e,f), January (g,h), February (i,j), and March (k,l). The contour interval

is 0.2 and negative values are dashed

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148 V. MORON AND I. GOUIRAND

Table II. Percentage of area (100 °W–50 °E; 20–70 °N) that is locally significant at the two-sided 90% levelfor (the first six rows) and between (the last three rows) the 10, 20 and 30 warm and cold ENSO events.One, two, three and four asterisks denote field significance at the one-sided 80%, 90%, 95% and 99% levels(null hypothesis is that the surface which is locally significant is zero) according to a Monte Carlo simulation

(see text)

October November December January February March

Warm (10) 8.1 28.0∗∗ 24.2∗ 14.2∗ 21.0∗ 19.8∗Warm (20) 2.2 18.9∗ 27.7∗∗∗ 17.8∗∗ 19.6∗∗∗ 30.9∗∗∗∗Warm (30) 10.7 21.4∗∗ 18.2∗∗ 15.4∗ 17.0∗∗ 17.8∗∗Cold (10) 6.7 13.6 24.2∗∗ 1.6 21.3∗ 23.6∗Cold (20) 5.8 15.2 13.8 3.8 28.7∗∗ 18.6∗Cold (30) 8.0 9.1 15.3∗ 3.5 20.4∗∗ 26.2∗∗Warm — Cold (10) 5.3 28.5∗∗ 35.7∗∗∗ 9.4 30.8∗∗∗ 43.2∗∗∗∗Warm — Cold (20) 7.1 30.3∗∗∗ 32.9∗∗∗ 22.9∗ 31.6∗∗∗ 43.1∗∗∗∗Warm — Cold (30) 21.1∗ 26.9∗ 30.5∗∗∗ 18.9∗ 27.4∗∗∗ 36.2∗∗∗∗

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Figure 2. PC between the warm (full line) and cold (dashed line) ENSO composites of consecutive months. The dotted lines indicate the80% confidence interval based on random reshufflings of 1000 30-year samples belonging neither to the warm nor to the cold ENSO

composite (see text)

is usually greater for the warm composite than for the cold one in November–January, and the reverse istrue for February–March (Table II). Figure 2 presents the PC between the 30 year cold and warm ENSOcomposites of consecutive months. The warm and cold composites of the months of September and April areincluded here to check the consistency with the months outside the winter semester. The consistency betweensuccessive months is highly similar for the warm and cold ENSO composites (Figure 2). The warm and cold

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ENSO NORTH ATLANTIC SEA LEVEL PRESSURE RELATIONSHIP 149

ENSO composites of October are positively (negatively) correlated with those of September (November)(Figure 2). The warm and cold ENSO composites of November (February) are positively correlated withthose of December (March) (Figure 2). In that case, the PCs are usually significant at the two-sided 80%level and independent of the size of the composites (not shown). The PCs between the composites of Januaryand those of February (December) are usually positive (negative), but not significant at the two-sided 80%level (Figure 2). In the following, we discard the analysis of the month of October and concentrate onNovember–December and January–March.

In the warm ENSO composites of November–December (Figure 3(b) and (c)), negative SLPAs arelocated over the northern North Atlantic, from Greenland to Scandinavia, and positive ones south of40–50 °N. This pattern suggests the presence of stronger-than-usual westerlies (i.e. positive phase ofthe NAO) across the North Atlantic. A negative SLPA, which appeared 1 month ahead, grows strongerover the southeastern USA in January (Figure 3(d)). This pattern is stronger in February (Figure 3(e)),then decays in March (Figure 3(f)). Positive SLPAs are located from the Hudson Bay to Iceland andalso over North Africa and the Iberian Peninsula in February–March (Figure 3(e) and (f)), and neg-ative SLPAs stretch from the southeastern US coast to Scandinavia across the central North Atlantic(Figure 3(e) and (f)). The warm ENSO composite in February–March suggests that the westerlies

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Figure 3. Mean SLPA (hPa) in October (a), November (b), December (c), January (d), February (e), and March (f) observed in the 30warm ENSO seasons. Negative values are dashed and shading indicates significant values at the two-sided 90% level according to a

Student’s t-test (null hypothesis is that the mean of the sample is zero)

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150 V. MORON AND I. GOUIRAND

are weaker than usual and/or there is a more southern path of the westerlies in the western NorthAtlantic sector.

In November–December (Figure 4(b) and (c)), the cold ENSO composite is characterized by a positiveSLPA from Greenland to western Europe and a negative SLPA mainly over the eastern part of the Azoreshigh. Even if this pattern is consistent with a weakening of the westerlies (i.e. negative phase of the NAO),it is not strictly symmetric with the warm ENSO composite. This pattern changes in January (Figure 4(d)),with a negative SLPA from Newfoundland to Scandinavia and a positive SLPA south of 35–50 °N overthe Atlantic and over the middle of Europe. In February–March (Figure 4(e) and (f)), this pattern growsstronger (Table II), with a positive SLPA from the southeastern USA extending towards western Europe (andScandinavia in March) and a negative SLPA north of 60°N, suggesting a strengthening and/or a more northernpath of the westerlies in western and central parts of the North Atlantic (Figure 4(e) and (f)).

3.3. The seasonal modulation of the hemispheric association between SLPA and Tropical Pacific SST

Sections 3.1 and 3.2 revealed a seasonal modulation between November–December and January–Febru-ary–March. This section examines only the warm minus cold ENSO composite (for the 30 year sample) at thehemispheric scale to make clearer the exact nature of the relationship found over the North Atlantic (Figure 5).The patterns of November (Figure 5(b)) and December (Figure 5(c)) are close to each other (PC = 0.69∗∗).

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Figure 4. Same as Figure 3, except for the 30 cold ENSO seasons

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The pattern is thus characterized by a north–south dipole on each side of 40–45 °N latitude over the Pacificand the Atlantic Oceans with positive (negative) SLPA south (north) of these latitudes. A negative SLPAappears over the northwestern North Pacific in October (Figure 5(a)). The SLPAs are of the same amplitudeover both oceans in November–December (Figure 5(b) and (c)). The anomalies of the Aleutian and Icelandlows are in phase, with negative (positive) SLPAs in both areas during the warm (cold) ENSO episodes.

From January (Figure 5(d)), the pattern changes (PC = 0.39 between the composites of December andJanuary), with a wave-like pattern around North America and a southwest–northeast elongated pattern overthe North Atlantic and Europe described in Sections 3.1 and 3.2. The wave-like pattern across North Americaresembles the tropical North Hemisphere (TNH) pattern (Barnston and Livezey, 1987). The pattern over theeastern Atlantic and Europe is quite different to the typical TNH pattern (see figure 5 of Barnston and Livezey(1987)). A positive SLPA moves westward from Scandinavia in January (Figure 5(d)) to Greenland in February(Figure 5(e)), then to the northeast of Canada in March (Figure 5(f)) while the negative SLPA remains almoststationary over the northeast North Pacific and the southwest North Atlantic (Figure 5(d)–(f)). The positiveSLPAs seem even to appear over northern Scandinavia in December, even though not significant at the 90%level (Figure 5(c)). The pattern of March (Figure 5(f)) appears as a residual of the previous 2 months. Itshould be noted that the PC between the warm minus cold ENSO composite of January and that of February(those of February and March) equals 0.67*** (0.81***). In January–March (Figure 5(d)–(f)), the SLPAs ofthe Aleutian and Iceland lows are out of phase, with negative (positive) SLPAs over the first (second) areaduring the warm ENSO events. So, the seasonal modulation found over the North Atlantic is a regional partof a hemispheric pattern, which also changes between November–December and January–March.

4. DISCUSSION AND CONCLUDING REMARKS

This paper has addressed the seasonal modulation of the relationship between ENSO and the North AtlanticSLPA during the northern winter.

The relationship between tropical Pacific SSTA and SLPA in the North Atlantic region explains only asmall amount of the total variance when the whole 1873–1996 period is studied. In a linear sense, thealternation between warm and cold ENSO events in the central and eastern tropical Pacific explains, at best,16–25% of the whole interannual variance at a monthly time scale (e.g. in the southwestern North Atlanticin February–March). This conclusion is consistent with previous results of Rogers (1984), Fraedrich (1993)and Pozo-Vazquez et al. (2001), amongst others.

We show here that the relationship between ENSO episodes and SLPA over the North Atlantic changesstrongly between November–December and January–March. The relationship in October is not significant,and is quite different from that found in November–December. In November–December, the warm ENSOepisodes are characterized by a significant drop of SLP between Greenland and the British Isles and asignificant increase on the southern margins of the Azores high. The cold ENSO events lead to weaker-than-usual westerlies, with a positive SLPA stretching from Greenland to western Europe. In January, and mainlyFebruary–March, the strongest SLPA response is shifted to the western North Atlantic region. The warmENSO events are then associated with negative SLPAs extending from southeastern USA toward the easternNorth Atlantic and western Europe, and positive SLPAs from northeast Canada toward Greenland and thenortheast North Atlantic. The cold ENSO episodes lead to almost the reverse in SLPA. The conclusions ofPozo-Vazquez et al. (2001), that no statistically significant SLPAs are found in December–February duringthe warm ENSO events, seem to be related to their seasonal pooling. Thus, the usual seasonal pooling of themonths of December, January and February into boreal winter (Fraedrich, 1990; Moron and Ward, 1998; Pozo-Vazquez et al., 2001) is misleading when the relationships with ENSO are studied in this area, because it mixesthe different patterns of December and February. The best seasonal pooling seems to be November–Decemberon the one hand and January–March (or February–March) on the other hand. We agree with Fraedrich (1990)and Pozo-Vazquez et al. (2001), who reported (for January–February and December–February respectively)that the North Atlantic climate was influenced more strongly by cold than by warm ENSO events, but thisstudy shows that this is mainly the case in February–March and that the opposite situation is observed inNovember–January.

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Figure 5. Difference between the warm and cold ENSO composite (30 years in each sample) in the Northern Hemisphere (in hPa) inOctober (a), November (b), December (c), January (d), February (e) and March (f). Negative values are dashed and shadings indicatethe significant values at the two-sided 90% Student’s t-test (null hypothesis is that the mean of warm and cold composites is equal)

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Warm (cold) ENSO events are associated with the positive (negative) phase of the NAO in Novem-ber–December and this association is inverted in January and mainly February–March. But, the classicalNAO index (i.e. Azores minus Iceland SLPA) is not an optimal index when the relationships between theSLPA of the North Atlantic region and the ENSO are studied (e.g. Rogers, 1984). In fact, the highest absolutehomogeneous correlations (in SVDA) and absolute SLPA (in warm and cold composites) are usually shiftedrelative to Reykjavik and Ponta Delgada, or even relative to the centres of the Icelandic low and of theAzores high.

We show here that the major change between November–December and January–March is also seen atthe hemispheric scale. In November–December, the warm-cold ENSO composite is dominated by a globalnorth–south dipole with negative (positive) SLPA north (south) of 40–45 °N peaking in the western andcentral North Pacific and the central and eastern North Atlantic. In January–March, the North Atlantic’srelationship is a regional part of a hemispheric pattern dominated by a wave across North America. Even ifthe January–March composites (Figure 5(d)–(f)) just represent the mean state of the Northern HemisphereSLPA conditional on the warm minus cold ENSO difference, and not necessarily a mode of these SLPAs, itis interesting to note that they resemble the TNH pattern (Barnston and Livezey, 1987). A notable differencewith the TNH pattern is the positive centre migrating from Scandinavia to northeastern Canada.

The explanation of the seasonal modulation observed at the hemispheric scale between Novem-ber–December and January–March deserves further study, but several hypotheses could be put forward.It is unlikely that this seasonal modulation can be explained by the annual cycle of the ENSO alone, becausethe SSTAs in the central and eastern Tropical Pacific are highly persistent between October and March, beforethe ‘spring predictability barrier’ in April–May (Balmaseda et al., 1995). Other hypotheses are related to therole of the North Atlantic’s atmospheric basic state and the relationships between the North Pacific and theNorth Atlantic sectors. The speed of the low-level westerlies could modify the length of the Rossby wavetrain. The development of a wave, such as the one found here in January–March, is perhaps only possiblewhen the mean westerlies reach a given threshold. The time evolution of the warm minus cold ENSO com-posite (Figure 5) is quite consistent with the formation of the seesaw between the Aleutian and Icelandic lowsrecently described by Honda et al. (2001). In particular, they show that the out-of-phase relationship betweenthe intensity of both lows during the ‘peak period’ (from 31 January to 16 March) is preceded by a weakin-phase association during the beginning of the winter, i.e. in November–December (see their Figure 6(a)and (b)). They also demonstrate that the seesaw in February–mid-March is robust even after the influence ofENSO has been removed. This seesaw is initiated by the amplification of the Aleutian low anomalies, thenthe propagation of a Rossby wave train across North America and formation of stationary anomalies over theNorth Atlantic from January (Honda et al., 2001).

It is possible that ENSO triggers the above oscillation, which is essentially initiated by mid-latitudeprocesses. This hypothesis is consistent with the results of Straus and Shukla (2000) and Shukla et al. (2000),who analysed 9 runs from a 39-year atmospheric general circulation model forced by annually varying SST.They demonstrated that, over the period beginning 20 December, the dominant SST-forced mid-latitudecirculation pattern around North America is not the Pacific–North America pattern but a pattern that isquite similar to the January–March warm minus cold composite shown in Figure 5 (e.g. see figure 11 ofStraus and Shukla (2000)). This atmospheric response is linked to ENSO-related diabatic heating (May andBengtsson, 1998; Shukla et al., 2000). The intensification (reduction) of the Aleutian low during warm (cold)ENSO events is associated with a larger (smaller) advection of warm maritime air over North America,then a reduction (intensification) of the continent–ocean thermal contrast downstream over the northwesternAtlantic. In consequence, cyclones develop less (more) frequently over this area, leading to a possiblereduction (intensification) of the Icelandic low during the warm (cold) ENSO events (Fraedrich, 1993; Mayand Bengtsson, 1998). This mechanism could explain the SLPA pattern observed from January and particularlyin February–March, but it is less relevant in November–December. The covariability of the Aleutian andIcelandic lows is weak during that period (Honda et al., 2001) and other mechanisms could be more efficient.For example, the anomalous warmth (coldness) of the tropical troposphere during the warm (cold) ENSOepisodes (Wu and Newell, 1998; Diaz et al., 2001) could increase (reduce) the meridional north–south pressuregradient between tropical and extratropical latitudes, and then strengthen (weaken) the westerlies in the North

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154 V. MORON AND I. GOUIRAND

Pacific and North Atlantic sectors. The above hypotheses do not exclude other mechanisms, such as theseasonal modulation of the forcing of the SSTA in the central and eastern tropical Pacific on the tropicalAtlantic SSTA, then on the Azores high (Giannini et al., 2000), or the seasonal modulation of the localatmospheric response to North Atlantic SSTA (Peng et al., 1995).

It is important to recognize that the ENSO relationships highlighted in Section 3 represent statisticalaverages over the last 125 years. The non-stationarity and decadal variations in observed ENSO relationshipswith the North Atlantic SLPA will be explored in future studies.

ACKNOWLEDGEMENTS

Thanks to Tracy Basnett (UKMO, Bracknell, UK), who kindly provided us with the SLP monthly fields. Wewould also like to thank Luc Beaufort (CEREGE, Aix en Provence, France) and Gilles Reverdin (LEGOS,Toulouse, France), who kindly commented on a first draft of this paper. We greatly appreciated the carefulreading of Alessandra Giannini (NCAR, Boulder, USA) and M. Neil Ward (IRI, New York, USA) on anintermediate version of this paper and the efforts of the reviewers to improve the paper.

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