pacific decadal variability: a review · the pdo pattern is similar to the sst variability...

Minobe et al. PDV review

Pacific Decadal Variability: A Review

Shoshiro Minobe1, Niklas Schneider2,

Clara Deser3, Zhengyu Liu4, Nathan Mantua5,

Hisashi Nakamura6,7, and Masami Nonaka7

1: Division of Earth and Planetary Sciences, Graduate School of Science, Hokkaido University, N-10, W-8,

060-0810, Sapporo, Japan.

2: International Pacific Research Center, University of Hawaii at Manoa, 1680 East West Road, Honolulu,

HI 96822 USA.

3: NCAR / Climate and Global Dynamics Division, P.O. Box 3000, Boulder, CO, 80307 USA.

4: Center for Climatic Research, Gaylord Nelson Institute for Environmental Studies, University of

Wisconsin-Madison, 1225 W. Dayton St., Madison, WI 53706-1695 USA.

5: University of Washington, Climate Impacts Group, Box 354235, Seattle WA 98195-4235 USA.

6: Department of Earth, Planetary Science, Graduate School of Science, University of Tokyo, Tokyo,

113-0033 Japan.

7: Frontier Research Center for Global Change, Japanese Agency for Marine, Earth Science and

Technology, Showa-machi, Kanazawa-ku, Yokohama, Kanagawa 236-0001, Japan.

Submitted to Journal of Climate, 1 Sept., 2004.

Corresponding author address

Shoshiro Minobe

Rigaku 8th Building,

Division of Earth and Planetary Sciences, Graduate School of Science, Hokkaido University,

N-10, W-8, 060-0810, Sapporo, Japan.

e-mail: [email protected]

tel.: +81-11-706-2644

fax.: +81-11-746-2715


Contents

Abstract ........................................................................................................................................................... 1 1. Introduction ............................................................................................................................................. 1 2. Spatial patterns of SST and overlying atmosphere.................................................................................. 2

2.1. PDO................................................................................................................................................. 2 2.2. Other SST patterns .......................................................................................................................... 3

3. Climatic regime shifts ............................................................................................................................. 3 3.1. 1970s climatic regime shift ............................................................................................................. 3 3.2. 1920s and 1940s regime shifts and 1980s change........................................................................... 4 3.3. Possible regime shift in the late 1990s ............................................................................................ 5

4. Timescale dependent signatures .............................................................................................................. 5 4.1. Quasi-decadal (10-yr) oscillations................................................................................................... 6 4.2. Bidecadal (20-yr) oscillation ........................................................................................................... 6 4.3. Pentadecadal (50–70-yr) oscillation ................................................................................................ 7

5. Impacts .................................................................................................................................................... 7 5.1. Land air-temperature ....................................................................................................................... 7 5.2. Precipitation .................................................................................................................................... 7 5.3. Storm track ...................................................................................................................................... 8 5.4. Subsurface temperature and salinity................................................................................................ 8 5.5. Ocean currents................................................................................................................................. 9 5.6. Biochemistry ................................................................................................................................. 10 5.7. Decadal modulation of interannual variability .............................................................................. 10

6. Midlatitude origin.................................................................................................................................. 11 6.1. Atmospheric stochastic forcings on ocean mixed layer ................................................................ 11 6.2. Atmospheric stochastic forcings on ocean propagation (advection or waves) .............................. 12 6.3. Intrinsic variability in the ocean.................................................................................................... 13 6.4. Mid-latitude delayed oscillator...................................................................................................... 14 6.5. Possibility of mid-latitude ocean-to-atmosphere feedback............................................................ 14

7. Tropical origin....................................................................................................................................... 15 7.1. ENSO residual and ENSO bursting............................................................................................... 16 7.2. Tropical delayed oscillator ............................................................................................................ 16 7.3. Basin mode.................................................................................................................................... 16

8. Coupled midlatitude-tropical origin ...................................................................................................... 17 8.1. Advected temperature changes in the ocean.................................................................................. 17 8.2. Transport change in the ocean (Subtropical Cell) ......................................................................... 17 8.3. Influence of mid-latitude atmosphere on the tropics ..................................................................... 18

9. External forcing to the global climate system ....................................................................................... 18 9.1. Solar radiation and lunar-solar tide ............................................................................................... 18 9.2. Possible linkage to global warming............................................................................................... 18


10. Predictability ..................................................................................................................................... 19 11. Discussion ......................................................................................................................................... 19 Acknowledgements ....................................................................................................................................... 20 References ..................................................................................................................................................... 21 Figure and Figure Captions ........................................................................................................................... 34

Minobe et al. 1 PDV review

Abstract

For decadal climate variability over the Pacific Ocean, observed features, mechanisms, and

predictability are reviewed. For the observed variability, we summarize the spatial patterns, and temporal

evolution, including so-called climatic regime shifts, as well as impacts on the atmosphere and ocean.

Various hypotheses of mechanisms of the Pacific decadal variability are described according to their

domain of influence as mid-latitudes, tropics, and pan-Pacific. In addition, possible influences of natural

and anthropogenic external forcings are considered, and current state of predictability is reviewed.

1. Introduction Decadal-to-centennial (decadal hereafter) climate variability in the atmosphere and ocean over

and around the Pacific Ocean has attracted large attention in the past decade and a half. The substantial

impacts of decadal variability on climate, ecosystem, and society clearly indicate the importance of

understanding of Pacific Decadal Variability (PDV), its mechanism and predictability. In this review, we

aim to summarize the present state of knowledge of decadal climate variability over the Pacific Ocean, and

hence also to update recent summaries of this topic by Latif (1998), Miller and Schneider (2000), Mantua

and Hare (2002), and an unpublished summary by Sarachick and Vimont1. In this review, we document

important aspects of observed decadal variability, summarize proposed mechanisms, and discuss

predictability.

Due to space limitations, we will not give an overview of results from paleo-climate studies. A

noteworthy review of tree-ring reconstructions on PDV is given in Mantua and Hare (2002). Coral-based

information on PDV may be found in Linsley et al. (2000), Urban et al. (2000), Cole et al. (2000) and

Evans et al. (2002).

Also, we will not describe the impacts of decadal climate variability on ecosystems. In particular,

dramatic impacts of decadal climate variability on marine ecosystems are widely accepted, with prime

examples from salmon (e.g., Mantua et al. 1997; Hare and Mantua 2000) and sardine (e.g., Yasuda et al.

1999), and fisheries see also the recent review by Miller et al. (2004) and Bakun (2004).

The rest of this review is organized as follows. Sections 2−5 are devoted to descriptions of

observational variability. Spatial patterns of SST and the overlying atmosphere are described in section 2.

Regime shifts are documented in section 3, and section 4 shows timescale dependent signatures. Impacts on

climate parameters are described in section 5. Mechanisms are described in sections 6−8, with mid-latitude

origin in section 6, tropical origin in section 7, and interaction between these two regions in section 8.

Possible external forcings are described in section 9. Predictability is discussed, and a prediction scheme of

SSTs over Kuroshio-Oyashio Extension (KOE) region is highlighted in section 10. Discussion is presented

in section 11.

1 Available from http://www.aos.wisc.edu/%7Edvimont/Papers/pdv.pdf


2. Spatial patterns of SST and overlying atmosphere 2.1. PDO

A widely cited index for North Pacific decadal variability, the “Pacific (inter-)Decadal

Oscillation” (PDO) is based on the leading Empirical Orthogonal Function (EOF) of monthly Sea-Surface

Temperature (SST) anomalies (departures from the long-term monthly means) north of 20°N. Before the

EOF calculation, global mean SSTs were subtracted from the SST field (Mantua et al. 1997). The PDO

pattern forms an “oval” in the western and central North Pacific, surrounded by a “horseshoe” of opposite

signed anomalies off California, in the Gulf of Alaska, and towards the tropics, where SST anomalies form

a “triangular” shape (Figs. 1 and 2). The PDO pattern is similar to the SST variability associated with

ENSO events to some extent, but the PDO has larger amplitudes in mid-latitudes rather than low-latitudes

and a broader width of equatorial anomalies than those of ENSOs (see Fig. 2 in Mantua et al. 1997). Zhang

et al. (1997) obtained essentially the same pattern as the PDO, as the regression map onto the principle

component of global SSTs linearly removed interannual ENSO variability. Zhang et al. (1997) called this

pattern of low-frequency SST variability “ENSO-like interdecadal variability”. Mantua et al. (1997) noted

that the PDO may be viewed as ENSO-like interdecadal variability. By repeating methodology used by

Zhang et al. (1997) for global reanalysis data of National Centers for Environmental Prediction– National

Center for Atmospheric Research (NCEP-NCAR) (Kalnay et al. 1996), Garreaud and Battisti (1999)

showed that the SST and atmospheric patterns are meridionally symmetric with respect to the equator using

the all the calendar months (Fig. 2). In winter hemisphere, however, has stronger atmospheric signals.

The principal component of the PDO, the PDO index, shows pronounced low frequency

variability with transitions from negative to positive values in 1924/25 and in 1976/77, and to negative

values in 1947/48 and possibly in 1998 (Fig. 1). There are also a few minor transitions, for example in the

late 1950s and around 1990.

The atmosphere co-varies with the PDO index, suggesting a coupled phenomenon (Mantua et al.

1997). The regression of sea level pressure and winds on the PDO index (Fig. 1) shows that cool SST

anomalies in the central North Pacific are associated with a deepened Aleutian Low covering much of the

North Pacific, and enhanced westerlies in the central North Pacific. The atmospheric circulation anomalies

associated with the PDO resembles the Pacific/North American (PNA) pattern (Wallace and Gutzler 1981;

Horel and Wallace 1981) more than those related to ENSOs (Zhang et al. 1997).

Decadal variability seems to have little seasonal modulation in SSTs. Separate EOF analysis of

SSTs during winter and summer seasons (Zhang et al. 1998c; Norris et al., 1998; Barlow et al. 2001)

showed that summer SST anomalies are a little stronger than their wintertime counterpart, with similar

pattern in the North Pacific. Norris et al. (1998) suggested that local marine stratiform clouds play a role in

maintaining the summertime SST anomalies. On the other hand, Barlow et al. (2001) suggested the

importance of subsurface temperature anomalies in mid-latitudes. Newman et al. (2003) argued that the

winter to summer persistency in mid-latitudes is due to continuous tropical forcings.

Other names used for the phenomenon called the PDO are the Interdecadal Pacific Oscillation

(IPO) (Power et al., 1997, 1999) and the North Pacific Oscillation (NPO) (Gershunov and Barnett 1998).

Please note this North Pacific Oscillation is different from the meridional dipole of Sea-Level Pressure


(SLP) anomalies called also the North Pacific Oscillation parallel to the North Atlantic Oscillation (NAO)

by Walker and Bliss (1932) and Rogers (1981), and also different from the North Pacific mode obtained by

Barlow et al. (2001) as described below.

2.2. Other SST patterns

The PDO is not the only pattern of Pacific SST variability on decadal timescales. Barlow et al.

(2001), using rotated EOF analysis, identified three dominant patterns in the Pacific Ocean, one associated

with interannual ENSO variability, and two on decadal timescales (PDO and a North Pacific mode) (Fig. 3).

Their PDO pattern primarily occupies the eastern Pacific with comparatively weak anomalies of opposite

polarity in the central North Pacific, though the original PDO pattern has sizable amplitudes also in the

western North Pacific. The North Pacific mode has largest anomalies at 40°N from the coast of Asia to

140°W, with weaker opposing anomalies along American coast, without substantial signatures in the tropics.

Similar modes were obtained by an EOF analysis applied to low pass filtered SST anomalies in a limited

region over the central-to-western North Pacific (150°E−140°W, 20°N−50°N); the first and second modes

have the large amplitudes in the subpolar front and subtropical front, respectively (Nakamura et al. 1997).

Tomita et al. (2001) confirmed the robustness of these patterns, analyzing correlation maps for

area-averaged SST time series. The pattern of the subpolar frontal mode is similar to the North Pacific

mode of Barlow et al. (2001) and the 2nd EOF mode of north of 20°S by Deser and Blackmon (1995). The

subtropical frontal mode of Nakamura et al. (1997) resemble to Barlow’s PDO. These studies suggested

that two spatial modes are robust features for the decadal variability over the Pacific Ocean. Enfield and

Mestas-Nuñez (1999), however, reported that the subpolar variability was liked to the central equatorial

SST anomalies in their EOF analysis of reconstructed SST data of Kaplan (1998) during 1856–1991.

The subpolar frontal mode by Nakamura et al. (1997) is closely related to the deepened Aleutian

low pattern at the sea surface and PNA pattern aloft. The atmospheric circulation anomalies for the

subtropical frontal mode resemble the East Pacific pattern characterized by a meridional dipole between the

mid-latitude eastern North Pacific and Alaska (Nakamura et al. 1997).

3. Climatic regime shifts 3.1. 1970s climatic regime shift

One of the interesting features of the decadal variability over the Pacific Ocean is the so-called

climatic “regime shift”, the 1976/77 regime shift being a prime and well studied example. Although several

studies investigated decadal variability over the Pacific Ocean before the late 1980s (e.g., Namias 1969;

Dickson and Namias 1976; Douglas et al. 1982), studies for decadal variability were sporadic before the

series of studies for the 1976/77 regime shift. Kashiwabara (1987) reported negative geopotential height

anomalies over the North Pacific after 1977 and possible linkage to ENSO. Namias et al. (1988) showed

that the decadal fluctuation of the first EOF mode of SSTs over the North Pacific coherent with the PNA

index. Nitta and Yamada (1989) showed that the dramatic changes of global SST, with essentially the same

pattern as the PDO, and more frequent occurrences of the positive phase of the PNA pattern after 1977, and


suggested that a tropical Pacific origin for the 1976/77 shift, based on SST, rainfall and SLP. Trenberth

(1990) emphasized the step-like nature with a transition in 1976/77 between two regimes especially for

Aleutian low strengths in winter, and showed that Aleutian low was deepened and shifted eastward after

1977. Trenberth and Hurrell (1994) investigated more closely the structure of atmospheric variability

associated with the 1970s regime shift. Graham (1994) concluded that the tropical SST anomalies are the

source of the extratropical atmospheric circulation anomalies, based on data analysis and numerical

modeling.

The 1976/77 cooling in the central North Pacific and warming off California and in the Gulf of

Alaska (Nitta and Yamada 1989) was caused by changes of surface heat flux, mixing and Ekman advection

(Miller et al. 1994, Seager et al. 2001) associated with the increases in wintertime mid-latitude westerlies

(Trenberth and Hurrell 1994). Winter mixed-layer depth were increased by 20-30 m in the central North

Pacific, and reduced in the east (Polovina et al. 1995; Deser et al. 1996; Miller and Schneider 2000). Also,

Sea ice in the eastern Bering Sea was reduced after 1977 (Niebauer 1998).

3.2. 1920s and 1940s regime shifts and 1980s change

The 1970s climatic regime shift is not a unique phenomenon. Minobe (1997) and Mantua et al.

(1997) suggested that the regime shift occurred in the 1920s, 1940s and 1970s, with alternating polarities,

based on SLP, SST, air-temperature, precipitation data over and around the North Pacific. A climatic regime

shift may be defined as a transition from one climatic state to another within a period substantially shorter

than the lengths of the individual epochs of each climate states (Minobe 1997). Historically, Kutzbach

(1970) already documented the large and rapid SLP changes in the 1920s and 1950s, and Yamamoto et al.

(1986) documented significant changes in several atmospheric parameters, including air-temperature and

precipitations in Japan in the late 1940s, and called this phenomenon “climatic jump”, which is essentially

the same concept as the climatic regime shift. Cooper et al. (1989) showed that the substantial differences

of atmosphere and ocean conditions in the tropical Pacific from 1950s to the mid-1970s compared with the

earlier data.

The 1920s and 1940s shifts exhibited similar patterns to the 1970s, with some differences. The

1940s and 1970s SST changes are generally characterized by similar SST patterns with the features of the

PDO (Zhang et al. 1997), and 1940s shift also have maximal amplitudes in the subarctic and subtropical

fronts (Minobe and Maeda 2004) as well as the shift in the 1970s (Nakamura et al. 1997). On the other

hand, the 1940s SST warming was centered near Japan in contrast much weaker signatures in the 1970s

shift there, while the 1970s shift influenced more strongly air-temperatures over North America (Minobe

2002; Deser et al. 2004). SLP patterns for these shifts are characterized by the overall strengthening and

weakening of Aleutian lows in winter, but in spring SLP signatures shifted eastward (Minobe 2002; Deser

et al. 2004). Coherent tropical SST changes were observed in these shifts (Zhang et al. 1997; Minobe 1997;

Mantua et al. 1997). Very recently, Deser et al. (2004) showed the concurrent changes in SST, SLP, land

precipitation and ocean cloud cover, which may represent precipitation over the ocean, were observed in

the tropical Indian Ocean and Pacific Ocean associated with the 1920s, 1940s and 1970s climatic regime

shifts. Figure 4 shows that the North Pacific Index (NPI), which is defined averaged SLP anomalies over a


region 30°–65°N, 160°E–140°W (Trenberth and Hurrell 1994), and tropical index proposed by Deser

(2004) exhibited prominent and abrupt changes corresponding to the three regime shifts. Minobe (1999;

2000) suggested that these three regime shifts can be interpreted as simultaneous phase reversals of the

bidecadal and pentadecadal oscillations explained latter, and these two oscillations arise from different

physical mechanisms, because the two oscillations exhibit different seasonalities (Fig. 5).

Another prominent change is observed in 1988/89 with the weakened Aleutian lows (Trenberth

and Hurrell 1994), which is accompanied by corresponding SST changes and decreased ice cover in the Sea

of Okhotsk (Tachibana et al. 1996), and various parameters of climate and biological indices (Overland et

al. 1999a; Hare and Mantua 2000). This change in the Pacific sector is strongly associated with the Arctic

and North Atlantic atmospheric circulation changes as clearly seen both in NAO index (Hurrell 1995) and

Arctic Oscillation (AO) index (Thompson and Wallace 1998). Watanabe and Nitta (1999) reported that the

sharpness of decadal changes in 1989 arose from synchronous phase shifts of interdecadal variations over

the Pacific Ocean and quasi-decadal variations over the North Atlantic. In contrast to other regime shifts,

tropical SSTs did not exhibit consistent changes for the 1988/89 change (Yasunaka and Hanawa, 2003).

3.3. Possible regime shift in the late 1990s

A series of changes in the physical environment occurred over the North Pacific around 1998/99.

In 1999, California and Alaska air-temperatures were the lowest in decades against prevailing warm

anomalies after the 1970s shift (Schwing and Moore 2000; Minobe 2000). SSTs warmed in the northern

and eastern North Pacific and cooled in the central and western North Pacific, corresponding roughly the

negative PDO pattern against the positive PDO anomalies in the 1970s shift (Minobe 2002; Schwing et al.

2002), consistent with a phase reversal of the PDO index (Fig. 1). Chavez et al. (2003) and Peterson and

Schwing (2003) reported substantial changes of zooplankton, salmon, anchovy, and sardine in the eastern

North Pacific, opposing to what occurred at the 1970s regime shift. However, a closer look revealed

differences between the 1998/99 change and the 1970s shift. Bond et al. (2003) reported that the SST

change in 1998/99 was characterized by a meridional seesaw of SST anomalies in the western and central

North Pacific, different from the PDO. Minobe (2002) showed that wintertime SLP difference between

1999-2002 and 1977-1998 resembles the East Pacific pattern, rather than the overall Aleutian low strength

changes associated with the PDO, consistent with a sea-level rise in the central and western North Pacific

around 30°−40°N. The anomalous climate condition starting from 1999 at least temporarily terminated in

2002 to 2003 (not shown), based on the data to 2004. The nature and mechanism of the anomalous

condition from 1999 to 2002 still need to be explored.

4. Timescale dependent signatures Several studies examined timescale dependent structures, motivated by a working hypothesis that

different mechanisms may have different preferred timescales. Tanimoto et al. (1993) and Zhang (1997)

described differences in spatial structures between the decadal and interannual SST variability. On

decadal-to-centennial timescales, commonly reported oscillations are Quasi-Decadal Oscillation (QDO;


8−15-yr), BiDecadal Oscillation (BDO; 16−25-yr), and pentadecadal oscillation (50−70-yr) (e.g., Mann

and Park 1994, 1996; Chao et al. 2000; Minobe 2000). Here we summarize studies on QDO, BDO and

pentadecadal oscillation.

4.1. Quasi-decadal (10-yr) oscillations

A QDO may have occurred involving the tropics. Brassington (1997) showed that a QDO

(8−15-yr) is detected in equatorial SSTs, Southern Oscillation Index (SOI), and SOI reconstructed from

tree-rings. Tourre et al. (2001) conducted a joint Multi-Taper Method/Singular Value Decomposition

analysis for SST and SLP data of Kaplan et al. (1998; 2000) north of 30°S in the Pacific Ocean, and

obtained SST pattern somewhat similar to PDO but slightly shifted to the east, accompanied by zonal

dipole SLP anomalies in the North Pacific. Examining observed subsurface temperature and surface wind

anomalies, Luo and Yamagata (2001) showed that temperature anomalies for a 14-yr oscillation in the

equatorial Pacific were traced back to the Southern Hemisphere, and might initiate the occurrence of the

1970s regime shift. Giese et al. (2002) obtained similar conclusions. On the other hand, White et al. (2003)

and Hasegawa and Hanawa (2003) reported westward propagating subsurface temperature signatures in the

Northern Hemisphere on timescales of 12–13-yr along 16°−18°N, which appear to lead QDO in equatorial

SSTs. White and Tourre (2003) proposed that the tropical QDO is accompanied by a globally distributed

eastward propagating wavenumber-1 pattern of SLP.

Probably independent from the QDO over the tropics, another QDO was reported in mid- and

high-latitudes in the North Pacific. Mann and Park (1996) showed that the Northern Hemisphere QDO is

related to the NAO (Hurrell 1995) with some signature in the North Pacific. Xie et al. (1999) showed that

the QDO of the NAO strongly influenced SSTs over the northern Japan Sea and northern North Pacific

roughly along the subpolar front, with a similar pattern to AO. Consistently, influence of the quasi-decadal

variability shared by NAO/AO is also detected in the subsurface temperatures in the Japan Sea (Minobe et

al. 2004) and Okhotsk Sea (Minobe and Nakamura 2004). These results showed that the quasi-decadal

variability in the northwestern North Pacific and adjacent seas are related to the quasi-decadal oscillation in

AO/NAO.

4.2. Bidecadal (20-yr) oscillation

Climate signals on a 20-yr timescale over the North Pacific and also in North America has been

analyzed from various aspects. BDO signatures were captured in the basin-scale or global analyses of

surface temperatures (Ghil and Vautard, 1991; Kawamura, 1994; Mann and Park, 1994; Chao et al., 2000),

and analyses of surface temperatures and SLPs (Polovina et al., 1995; Mann and Park, 1996; White et al.,

1997; White and Cayan, 1998; Zhang et al., 1998b; Tourre et al., 1999, 2001; Minobe 1999, 2000; Minobe

et al. 2002). Most of these studies showed similar spatial patterns in SSTs or SLPs over the North Pacific.

BDO distributes globally, but has the largest amplitude in the central North Pacific in association with the

variability of the Aleutian Low (e.g., Mann and Park 1996; Minobe et al. 2002), accompanied SST

anomalies similar to that of PDO in the northern hemisphere with meridional symmetry about the equator

(e.g., White and Cayan 1998). BDO signatures was also observed in Alaska air- and water-temperatures


(Royer 1989), mixed layer depth in the central North Pacific (Polovina et al., 1995), and 400 m

temperatures centered in KOE (Deser et al., 1999). The first complex EOF mode of the upper layer heat

contents exhibited clockwise propagation in the subtropical gyre on the bidecadal timescale (Tourre et al.

1999). The evidence of the interdecadal variability on about a 20-year timescale was also observed in

tree-ring records over North America (Mitchell et al. 1979; Ware, 1995; Cook et al. 1997; Ware and

Thomson, 2000; Biondi et al., 2001).

4.3. Pentadecadal (50–70-yr) oscillation

Multidecadal variability is prominent in PDO index (Fig. 1) and in the NPI (Fig. 5), characterized

by the aforementioned three regime shifts in the 1920s, 1940s and 1970s. Minobe (1997; 2000) suggested

that these shifts are likely to be associated with a 50−70-yr (pentadecadal) oscillation, which is detected

both in instrumental data and a tree-ring reconstruction. Consistently, Chao et al. (2000) detected a 70-yr

oscillation in SST anomalies. Analyzing the spectrum of the NPI and aforementioned tropical index

(section 3.2), Deser et al. (2004) showed that the 50-year spectral peaks are commonly observed in the two

time series, and also in a coral record in the Indian Ocean in the 19th and 20th centuries (Fig. 6). The NPI

spectral peak is consistent with the previous spectral analysis (Minobe, 1997; Percival et al. 2001) and

wavelet analysis (Minobe 1999, 2000). It is noteworthy, however, the 50-year frequency is almost the

lowest frequency that can be estimated from 100-yr data, and hence it is impossible to discriminate whether

an apparent 50-yr peak is accompanied by a spectral trough at lower frequencies (true peak) or not (false

peak). Despite this limitation, a statistically significant spectral trough between decadal and interannual

frequency bands in Fig. 6 strongly suggests that the decadal variability arise from different mechanisms

from those for the interannual variabilities (see also Trenberth and Hurrell, 1994).

5. Impacts 5.1. Land air-temperature

Reflecting the similarity of spatial pattern between PDO and ENSO, air-temperature anomalies

due to PDO are generally similar to those connected to ENSO (Mantua et al. 1997). In association with

PDO, wintertime temperature over Alaska and western Canada and the Pacific Northwest are higher, and

temperature in Mexico and the southeastern US is reduced (Mantua and Hare, 2002). Minobe (2002)

reported that the pentadecadal variability of PDO is observed in mid-latitude western North America in

spring, but not in winter. Consistently, Cayan et al. (2001) showed that spring came earlier since the late

1970s in western United States (US), based on the blooming of lilac and honeysuckle bushes and the timing

of snowmelt-runoff pulses.

5.2. Precipitation

Boreal winter precipitation shows a relation to PDO; a positive phase of PDO is associated with

increases of rain and river runoff in the coastal ranges of Alaska, and in the southwestern US and Mexico,

while precipitation in Canada and Siberia is reduced (Mantua and Hare 2002). PDO are associated with


more rain in subtropics, but drier tropics and midlatitudes in North and South America (Dettinger et al.

2000). On the western side of the Pacific, increases of temperature and reduction of rain fall in Australia

occurs in conjunction with cool central North Pacific temperatures (Power et al. 1997). Although absolute

correlations of precipitations onto the PDO index are generally smaller than those of temperatures,

long-term precipitation changes were still closely related to the PDO. McCabe et al. (2004) showed that the

drought frequency in 20-yr moving window over US is strongly related to the multidecadal variability of

PDO, along with the contributions of the Atlantic Multidecadal Oscillation (Enfield et al. 2001). The PDO

is also likely to influence the Indian summer monsoon rainfall (Krishnan and Sugi 2003). A warm (cold)

phase of the PDO is characterized by a decrease (increase) in the monsoon rainfall and a corresponding

increase (decrease) in the surface air temperature over the Indian subcontinent. Other studies also showed

long-term precipitation changes in North America (e.g., Dettinger and Cayan, 1995; Cayan et al., 1998; Hu

et al., 1998; Higgins and Shi, 2000) and Australia (Latif et al., 1997). Minobe and Nakanowatari (2002),

focusing on the bidecadal timescale, showed that the bidecadal precipitation variability occurred over and

around the Pacific Ocean in boreal winter, accompanied by significant coherency peaks with the NPI, and

suggested that the BDO in precipitation was responsible for a long-lasting wintertime Hawaiian drought in

the 1990s (Fig. 7).

5.3. Storm track

Storm track activity also exhibits prominent decadal variability. Nakamura et al. (2002) showed

that wintertime storm track activity over the northwestern Pacific is enhanced during the period 1987−1994

compared with the period 1980−1986 in association with the weakening of east Asian winter monsoon.

Chang and Fu (2002) found that winter storm track activity over the North Pacific was significantly

correlated to the PDO index from 1948 to 1999. Also, the PNA pattern was largely linked to the eastern

Pacific cyclone frequencies in winter, and controls cyclone activity over the Gulf region and the North

American coast during the last two decades (Gulev et al. 2001). In summer, storm track activity indicated

increasing trend, accompanied by similar trend in SSTs (Norris 2000). Focusing on timescales slightly

longer than the timescale of synoptic scale eddies dominating storm tracks, Nakamura (1996) showed that

variability on timescales from 10 days to season was weakened after the 1976/77 regime shift over the

North Pacific. Storm track activity is important for the mean balance of atmospheric circulations (e.g.,

Holton J. R., 1992), and hence decadal variability of the storm track is expected to have wide range of

impacts on atmospheric parameters (Nakamura et al. 2004).

5.4. Subsurface temperature and salinity

Upper 3000-m temperature anomalies in the Pacific Ocean are dominated decadal oscillatory

signatures superposed on an overall warming trend (Levitus et al. 2000). The 1970s regime shift in upper

150 m temperatures exhibits a pan-Pacific distribution, and 400 m temperature already began warming in

1968 (Stephens et al. 2001).

Anomalous surface conditions are recorded in water masses subducting to thermocline. In the

central North Pacific, associated with the 1970s regime shift, cooler surface waters subducted (Deser et al.


1996; Yasuda and Hanawa 1997) and propagated as anomalies of thermocline depth in the southwestward

direction to the subtropical Pacific (Schneider et al. 1999; Tourre et al. 1999). The water mass carrying the

temperature anomalies from the central North Pacific is the central mode water, which is one of major

water masses in the North Pacific (Nakamura 1996; Suga et al. 1997). The subsurface temperature

anomalies originating from the central North Pacific somewhat continuous to the equator in appearance

(Zhang et al. 1998a), but tropical temperature anomalies are likely to be driven by local winds and not due

to temperature anomalies forced in the mid-latitudes (Schneider et al. 1999). An Ocean General Circulation

Model (OGCM) experiment by Inui et al. (1999) suggested that both wind and heat flux anomalies

contributed to anomalous temperatures of subducted waters. Ladd and Thompson (2002) concluded that the

buoyancy forcing is of primary importance in the variability of mode water formation.

Wind anomalies over the central North Pacific caused the western North Pacific subsurface

oceanic responses with a 5-year lag via Rossby waves excited in the central North Pacific (Miller et al.

1998, Deser et al. 1999). This response of the mid-latitude gyre changed temperature anomalies in the

Kuroshio-Oyashio Extension (KOE) region (Deser et al. 1999; Miller and Schneider 2000, Seager et al.

2001; Schneider et al. 2002). The subsurface temperature variability in KOE region becomes more decadal

with increasing depth (Deser et al. 1999). Subtropical mode water south of Kuroshio extension west of the

dateline also exhibited decadal variability (Yasuda and Hanawa 1997; Hanawa and Kamada 2000). The

superposition of the propagating subducted signal and the delayed response in the Kuroshio extension can

yield to a clockwise propagation of heat content anomalies (Zhang and Levitus, 1997: Tourre et al. 1999).

Salinity changes are not well understood, because much fewer observational data are available

for salinity than for temperature. Using the data of Japanese repeat section along 137°E, Shuto (1996)

reported a decadal salinity changes in North Pacific Intermediate Water and North Pacific Tropical Water.

For the latter water mass, Suga et al. (2000) attributed a step-like increase of salinity in 1978 to the 1976/77

shift, and proposed that the changes of advection speeds rather than surface water fluxes are responsible. In

the Gulf of Alaska, surface freshening of 0.6 Practically Salinity Unit (PSU) was observed in the 1980s

compared with the 1970s due to enhanced surface fresh water flux (Overland et al. 1999b). Observations

from north of Hawaii indicated decadal freshening and cooling in the upper thermocline from 1991 to 1997

(Lukas 2001), possibly in association with the decadal precipitation changes (Minobe and Nakanowatari

2002) and storm track activity changes (Nakamura et al. 2002). Sacrificing temporal resolution but

retaining spatial coverage, Joyce and Dunworth-Baker (2003) produced mean temperature and salinity

maps for two periods of 1945−1975 and 1976−1998 over the North Pacific, and showed that the largest

salinity and temperature changes between the periods occurred around the Kuroshio bifurcation front at

160°E. They further suggested that the meridional shift of the front related to the PDO is responsible for the

epoch difference.

5.5. Ocean currents

Anomalous winds also influenced current transports. Qiu and Joyce (1992) showed that Kuroshio

transport gradually increased from the mid 1970s to the early 1980s by 10.8 Sv (1970–80 to 1982–88) with

mean transport of 51.7 Sv (1969−88). Also, geostrophic transport of Kuroshio extension for 142°–150°E


was increased, consistent with a Sverdrup balance (Deser et al., 1999). On the other hand, abnormal

southward penetration of Oyashio inferred from coastal SST data suggested that the deepened Aleutian

Low also enhances the Oyashio on interdecadal timescales (Sekine 1988; Minobe 1997). McPhaden and

Zhang (2002) showed a slow-down tendency of meridional overturning circulation between the tropics and

subtropics in the 1980s and 1990s, but their updated analysis showed a rebound from 1998 to 2003

(McPhaden and Zhang 2004)

5.6. Biochemistry

Decadal variability in physical climate on decadal timescale substantially influenced biochemical

parameters in the ocean. Ono et al. (2001) showed that prominent bidecadal oscillation in Apparent Oxygen

Utilization (AOU) is observed in addition to trend-like increase component with an out-of-phase relation to

BDO in NPI from 1968−1998 in northwestern North Pacific on density surfaces of North Pacific

Intermediate Water (Fig. 8c). Similar bidecadal variability of oxygen concentration was observed in the

deep water in the Japan Sea, accompanied by out-of-phase oscillation of phosphate and temperature (Fig.

8ab) (Watanabe et al., 2003). Andreev and Watanabe (2002) showed that the AOU in the subarctic North

Pacific exhibited coherent decadal changes consistent with the NPI.

5.7. Decadal modulation of interannual variability

The climate regime shifts in the 1920s, 1940s and 1970s influenced interannual climate variations.

Minobe and Mantua (1999) showed that the interannual variability in the wintertime Aleutian Low is strong

in a regime with a stronger regime-mean Aleutian low, with coherent signatures in the SLP, 500 hPa

geopotential height, SST, and wind fields. Coherent modulation of interannual SLP variability were also

observed over the northern North Atlantic and Arctic Sea (Minobe and Mantua 1999). The modulation of

interannual variability prevailing both over the North Pacific and the North Atlantic may be related to the

prominence of so-called “Aleutian low and Icelandic low seesaw” on interannual timescales (Honda et al.

2001), because the periods of the prominent negative correlations identified by Yamane et al. (2002)

approximately corresponded to the regimes of stronger regime-mean Aleutian lows.

PDO also impacts ENSO’s teleconnected response on precipitation and temperature over North

America, Australia and India. Gershunov and Barnett (1998) showed that the occurrence of El Niño (La

Niña) during a positive (negative) phase of the NPO and La Nina during a negative phase, leads to strong

changes in precipitation frequency over US, while the El Nino (La Nina) and negative (positive) NPO tend

to produce weaker anomalies. The NPO is essentially the same as the PDO as mentioned in section 2.1.

Also, when El Niño (La Niña) events occur during warm (cold) phases of PDO, the dry (wet) Indian

monsoon is more likely to prevail (Krishnan and Sugi, 2003). Power et al. (1999) indicated that the ENSO

influence on Australian rainfall is dependent on the IPO or PDO. Such modulation suggests a possibility of

better prediction skill of interannual variability by knowing the status of decadal variability. However,

using a Coupled General Circulation Model (CGCM), Pierce (2002) concluded that the association between

NPO SSTs and ENSO’s effects on North America found by Gershunov and Barnett (1998) is primarily due

to the fact that both are responding to the same internal atmospheric variability. In such a case,


incorporating accurate predictions of NPO SSTs into ENSO prediction schemes would have little ability to

improve forecasts of ENSO’s effects.

The character of ENSO appears to change on interdecadal to multidecadal timescales. ENSO

amplitudes were large from 1885 to 1915 and after 1960 (e.g., Gu and Philander, 1995). Wang (1995)

showed that the characteristics of the onset of the El Niño changed since the late 1970s using the data from

1950 to 1992. An and Wang (2000) reported that the oscillation period of ENSO increased from 2–4 yr

(high frequency) during 1962–75 to 4–6 yr (low frequency) during 1980–93. They further indicated that the

period increase is related to the eastward shift of the zonal wind stress with respect to the SST anomalies,

using a Zebiak and Cane ocean model coupled with empirical atmospheric models. Wang and An (2002)

suggested that the changes in the mean currents and upwelling reduce the effect of the zonal temperature

advection while enhancing that of the vertical advection; thus, the prevailing westward propagation

(Rasumusson and Carpenter 1982) was replaced by eastward propagation or standing oscillation.

6. Midlatitude origin Mid-latitude amplitudes relative to the tropics for decadal variability are larger than those for

interannual variability related to the ENSO. This suggests that mid-latitudes may play more important roles

for decadal variability than those for interannual variability. In the last decade, studies of mid-latitude

processes for the decadal variability substantially progressed with several important concepts, such as

reemergence mechanism, stochastic resonance, as explained below. However, whether the mid-latitude

ocean influences the overlying atmosphere or not remains to be answered. This ocean-to-atmosphere

feedback is essential for a delayed negative feedback loop in mid-latitudes without involving the tropics

studied by a number of papers.

6.1. Atmospheric stochastic forcings on ocean mixed layer

In the extratropics, the oceanic response to intrinsic atmospheric variability is the “null hypotheses”

for decadal oceanic variability. The prototype of this dynamical system is a constant-depth reservoir of

temperature T driven by an air-sea heat flux (Hasselmann 1976, Frankignoul and Hasselmann 1977), as

given by,

dTT Fdt

λ + = (1)

where T is the temperature anomalies in the mixed layer, F is the stochastic heat flux forcing, λ is surface

heat flux response to anomalies of SST. Note that density, mixed layer thickness, and heat capacity of sea

water are absorbed into F and λ. Temperature anomalies are then given by a one sided, exponential

low-pass filter of the surface heat flux:

( ')( ) ( ') 't t tT t F t e dtλ− −

−∞= ∫ . (2)

Forcing at time scales shorter than the damping e-folding scale, λ-1, leads to a weak, damped response that

lags the forcing by 90°. As the time scale of the forcing increases beyond the damping timescale, the


temperature response mimics the forcing, with a phase shift approaching zero.

The power spectrum of temperature to white noise forcing is proportional to

2 2

1λ ω+

(3)

Thus, a temperature spectrum is inversely proportionally to the square of the frequency (a spectral slope of

‘-2’) at shorter periods, and a flat spectrum for longer periods. This red spectrum is a powerful hypothesis

consistent with observations of SST of the extratropical ocean (Frankignoul and Hasselmann 1977).

An interesting modification of model (1) is inclusion of atmospheric temperature adjustments to

changes of SSTs (Barsugli and Battisti 1998). In this case, the atmosphere and ocean temperatures follow

each other, and the air-sea heat flux vanishes at low frequencies. The vanishing air-sea flux at the sea

surface implies a reduced damping, and low-frequency oceanic and atmospheric temperature variance are

three-times increased from those without thermal coupling between the ocean and the atmosphere. This

reduction of damping and the increased variance are found in atmospheric general circulation models when

coupled to prescribed surface temperatures is replaced by a slab ocean model (Bladé 1997, Lau and Nath

1996, Pierce 2001).

The mixed layer examined in (1) was assumed to be a constant heat reservoir, but in reality is

subject to seasonal changes. Mixed layer temperature anomalies formed in one winter are covered by a

shallow newly formed mixed layer in summer, and exposed to the surface again in the next autumn, when

the mixed layer deepens. This so-called “reemergence mechanism” was investigated by Alexander and

Deser (1995) and Alexander et al. (1999). Recently, Deser et al. (2003) successfully reproduced observed

SST and heat content autocorrelation functions taking account of the reemergence mechanism.

6.2. Atmospheric stochastic forcings on ocean propagation (advection or waves)

Stochastic atmospheric forcings can cause spectral peaks in the ocean, when the dynamical ocean

processes for signal propagations such as advection (Saravanan and McWilliams, 1997) or waves (Jin

1997; Weng and Neelin, 1999) are included. A prototype equation of this model is given by,

( )/ / ( , )t P c P x F x tγ + ∂ ∂ + ∂ ∂ = (4)

where P is the amplitude of propagating signal, such as thermocline depth of Rossby waves or temperature

anomalies advected by mean currents, c is the propagation speed, and x is the coordinate in the direction of

propagation. When coordinates along the propagation characteristics ( , / )x s t x cξ = = − are introduced,

(4) results in

( , )PP c F sγ ξξ∂

+ =∂

. (5)

This is equivalent to (1), and the solution is given by

0

0

( ') /( ) / 0

0'( , ) ', ' ,

x x cx x x c

x

x xx x eP x t F x t dx e P x tc c c

γγ

− −− − −− = − + −

∫ . (6)

If spatial damping scale, c γ/ , is much shorter than the forcing scale in association with, e.g., slow

propagation speed of Rossby waves in high-latitudes, the system is essentially the same as (1). In the


Alaskan gyre, low frequency anomalies of dynamic height and pycnocline depth can indeed be attributed to

anomalies of Ekman pumping with (1) and a decay time of one year for dynamic height (Lagerloef 1995)

and one to two years for pycnocline depth (Cummins and Lagerloef 2002).

If damping scale is comparable or longer than the forcing scale, responses of (6) and associated

spectra are qualitatively dependent on spatial structures of atmospheric forcings. A monopole atmospheric

forcing results in a red-noise spectrum in the ocean with a -2 spectral slope, as in (3) (Frankignoul et al.

1997). On the other hand, stationary atmospheric forcings with two or more poles in the direction of

oceanic propagation can cause a spectral peak in the ocean due to “spatial resonance” (Frankignoul and

Reynolds, 1983; Jin 1997, Saravanan and McWilliams 1997, Neelin and Weng 1999). The resonance with

spectral peaks occur at a timescale of 2 /Wx c , where Wx is the distance between the any pair of poles

with alternating polarities, because along the wave trajectory the forcing never changes its sign. This

implies that even if the ocean responds only passively to atmospheric forcing, spectral peaks can occur in

the ocean -- the presence of spectral peaks in the ocean does not necessarily imply a coupled mode with

significant feedbacks from the ocean to the atmosphere. It should be noted, however, that a spectral peak in

the atmosphere, such as in the NPI described in section 4.3, implies the ocean-to-atmosphere feedback.

Consequently, physical implications of spectral peaks in the atmosphere and those in the ocean are

considerably different.

Recent studies showed two additional mechanisms whereby stochastic atmospheric forcing

results in oceanic spectral peaks due to oceanic propagation. One is the basin mode described in section 7.3.

The other is “preferential amplification” of zonal flows, such as Kuroshio extension, due to meridional

difference of zonal propagation speeds of ocean Rossby waves (Qiu 2003) (Fig. 9). The speed difference

causes the southwest-northeastward tilt of wave phase, and tilted surface elevation anomalies result in

enhanced zonal current anomalies on specific timescales. Qiu (2003) suggested that a 12-yr spectral peak

may have occurred due to this mechanism in the Kuroshio extension west of the dateline.

6.3. Intrinsic variability in the ocean

Intrinsic oceanic variability associated with energetic western boundary currents can also

generate low frequency variability via nonlinearity with eddy activities as reviewed by Miller and

Schneider (2000). Also, readers who are interested in the theoretical developments of this topic including

feedback from the ocean to the atmosphere may refer to Meacham (2000), Dewar (2001) and Kravtsov and

Robertson (2002).

If this type of variability occurs in the Pacific Ocean, a prime candidate region may be

Kuroshio-Oyashio Extension (KOE). Qiu (2000) described a contraction and expansion of the region of

high eddy variance in the KOE on interannual timescales using satellite altimeter data, but it is not clear

whether such variability is also important on decadal timescales. Qiu and Miao (2000) suggested that

interannual variations of Kuroshio path south of Japan since 1975 was a self-sustained oscillation, which

started due to a background condition change of the enhanced wind-driven Sverdrup transport associated

with the 1970s regime shift. This suggests that intrinsic oscillations in the ocean may be modulated by

atmospheric forcings on decadal timescales. Eddy variability may play a role for decadal oceanic variability


in other regions. Roemmich and Gilson (2001) reported that observed meridional heat flux due to eddies

varies on decadal time scales. Schneider et al. (2004) found that observed low frequency variations of

salinity derive from changes of the California current caused by anomalies of the eddy field in this region.

6.4. Mid-latitude delayed oscillator

One early influential hypothesis for PDV was a delayed oscillator in mid-latitudes with

ocean-to-atmosphere feedback in the western North Pacific (Latif and Barnett 1994; 1996). Specifically,

they suggested that North Pacific decadal SST anomalies are maintained by a positive feedback with the

air-sea fluxes of heat and Ekman advection. The associated anomalies of Ekman pumping generate Rossby

waves that, after multi-year delay, alter the heat transport of the western boundary current, counteracting

the initial temperature perturbation. This scenario, reminiscent of Bjerknes (1964), leads to a delayed

oscillator where the decadal time scale is determined by the Rossby wave propagation in the subtropical

gyre. As noted earlier in section 5.4, lagged response of the heat content and SST anomalies were

confirmed by observations (Miller et al. 1998; Deser et al. 1999). A number of idealized models have been

advanced that encompass these dynamics, such as Jin (1997), Weng and Neelin (1999), and Cessi (2000).

Latif (1998) and Miller and Schneider (2000) reviewed the numerical and theoretical studies of this

mechanism. A few more recent studies examined mid-latitude ocean dynamics with chaotic Lorenz

atmospheric model (Masuda 2002; Ferrari and Cessi, 2003).

However, recent studies based on long coupled integrations indicate that feedbacks are not

required to explain the decadal variance in the western North Pacific. For example, Frankignoul et al.

(2000) concluded that coupled feedbacks account for only a minor fraction of the decadal variance, in

contrast to an earlier analysis of a shorter segment of the same run (Robertson 1996). Similarly, a longer

integration of the coupled model used by Latif and Barnett (1994, 1996) failed to yield a peak at decadal

time scales (Schneider et al. 2002). Instead, spectra of ocean pressure and SST in the Kuroshio extension fit

those expected from pure stochastic forcing, indicating no evidence for the delayed negative feedback.

Although the verification failed for earlier models that involve delayed oscillators with ocean to

atmosphere feedback in mid-latitudes, a recent study by Wu et al (2003) showed that this feedback is

important for decadal variability in a CGCM. Using rotated EOF analysis of low-pass filtered SSTs in a

CGCM, Wu et al. (2003) detected a North Pacific mode and a tropical mode (Fig. 10), whose patterns are

reminiscent of observed REOF modes by Barlow et al. (2001). In order to identify the key regions and

processes, Wu et al. (2003) employed partial coupling and partial blocking strategy, and concluded that the

mid-latitude air-sea coupling is important for the North Pacific mode on multidecadal timescales.

6.5. Possibility of mid-latitude ocean-to-atmosphere feedback

An important issue for the mid-latitude origin hypotheses is whether oceanic variability

feedbacks to the atmosphere or not. In general, atmospheric forcing locally controls the extratropical North

Pacific SSTs, as shown by the correlation between SST and sum of latent and sensible heat fluxes (Cayan

1992; Deser and Timlin 1997). However, the western part of the subarctic front, or KOE, is the region

where the surface heat flux anomalies are controlled by SSTs on the decadal timescale (Tanimoto et al.


2003; Fig. 11). SST anomalies in this region are strongly influenced by ocean circulation in an OGCM (Xie

et al. 2000). Consistently, recent decadal variations in SST in the KOE are due to acceleration, rather than a

meridional shift proposed by Seager et al. (2000), of the currents (Qiu 2003), echoing the heat budget

analysis of Vivier et al. (2002) and Tomita et al. (2002). The SST anomalies in this region involve

substantial decadal changes of surface frontal structures (Nakamura and Kazmin 2003). Tanimoto et al.

(2003) hypothesized that the heat flux from the KOE influences the storm track over the North Pacific,

which in turn contributes to cause large scale atmospheric circulation anomalies.

The response of Atmospheric General Circulation Models (AGCMs) to prescribed extratropical

SST anomalies is generally weak (see the recent review by Kushnir et al. 2002), sensitive to season (Peng

and Whitaker, 1997), and involves changes of the mid-latitude storm tracks (Peng and Robinson 2001;

Inatsu et al. 2003). For example, Pierce et al. (2001) prescribed a typical SST anomaly pattern in an AGCM

and concluded that the atmospheric response projects weakly, but significantly, on the wind stress curl

pattern associated with the decadal variability.

However, prescribing anomalies of SST is problematic because only a part of the decadal SST

pattern represents the responses to oceanic variability, and the remainder is forced by the contemporaneous

atmospheric variability. An alternative approach is to examine the response of coupled atmosphere-mixed

layer ocean system to perturbations of the ocean heat flux convergence (Yulaeva et al. 2001; Sutton and

Mathieu 2002). Yulaeva et al. (2001) found that in response to an idealized heating perturbation in the KOE

region, sea level pressure exhibits a down-stream low, as expected from linear theory, supporting generally

a mid-latitude coupled mode.

These conclusions based on AGCMs or CGCMs have to be viewed with some caution, due to

limited resolutions and physics. A recent study of high-resolution satellite observations in the Kuroshio

extension indicates a clear signature of synoptic-scale SST influences on wind speed in the atmospheric

boundary layer (Nonaka and Xie 2003, see also Xie 2004 for other regions). Further, Feliks et al. (2004)

has shown that mid-latitude oceanic fronts significantly influence the free atmosphere, using a numerical

atmospheric model consisting of a boundary layer component and a quasi-geostrophic component. These

effects may not be fully represented in the current AGCMs due to insufficient horizontal and vertical

resolutions and physics in the boundary layer.

7. Tropical origin Tropical origin of North Pacific decadal variability was suggested by a number of papers

including those described in sections 2−4. Recently, Newman et al. (2003) suggested that the most of

annual mean PDO variability can be attributed to tropical variability, stochastic atmospheric forcing in

mid-latitudes and the reemergence mechanism (Fig. 12). Their simple model successfully hindcasted PDO

variability, with the correlation between modeled and observed PDO indices as high as 0.74. It should be

noted, however, the tropical SST changes can involve different mechanisms from those responsible

interannual ENSOs, and can arise from within the tropics as described in this section or influenced by

extratropics as explained in the next section.


7.1. ENSO residual and ENSO bursting

Several studies have suggested that decadal tropical climate variations are generated by tropical

atmosphere–ocean dynamics alone (e.g., Zebiak and Cane 1987; Münnich et al. 1991; Kirtman and Schopf

1998; Timmermann and Jin 2002). In particular, Timmermann et al. (2003) showed that a low-order ENSO

model was able to simulate realistic ENSO “bursting”, i.e., warm events with anomalously large amplitudes,

and suggested such bursting may be responsible for strong El Niños in 1972/73, 1982/83 and 1997/98. In

contrast, Thompson and Battisti (2001) suggested that the inclusion of stochastic forcing in a delayed

damped oscillator model of ENSO results in both interannual and decadal variability in the tropics.

7.2. Tropical delayed oscillator

Coupled numerical models have shown that decadal variability in the tropics involve a delayed

oscillator mechanism similar to that described for ENSO, but with slower Rossby waves at higher latitudes

(Kirtman 1997) or of higher vertical mode number (Liu et al. 2002). Knutson and Manabe (1998) and

Yukimoto et al. (2000) showed that their respective coupled models simulate the pattern of the ENSO-like

decadal mode characterized by triangular SST anomalies in the eastern tropical Pacific and opposing

anomalies in subtropics in the western Pacific. Westward propagation of ocean heat content anomalies in

these models occurred around 12°N in Knutson and Manabe (1998) (Fig. 13), but 20°–30°N in Yukimoto et

al. (2000). Wind stress curl anomalies also propagated westward along 20°N in Yukimoto’s coupled model,

while similar westward propagation was observed along 14°N by Capotondi and Alexander (2001). After

reaching the western boundary, the temperature signal appears to propagate equatorward and then eastward

along the equator, contributing to thermocline depth variability, which affects SSTs. These SST anomalies

induce atmospheric teleconnections in the extratropics with opposing polarity from the original anomalies

(Fig. 13).

7.3. Basin mode

The communication of the equatorial and extraequatorial regions of the ocean via waves allow

for the existence of a basin-scale mode of variability. Upon reflection off a western boundary, anomalous

mass of an extratropical Rossby wave propagates equatorward as a coastal Kelvin wave, and brings

extratropical anomalies to the equator (Lysne et al. 1997). At the eastern boundary of the basin, equatorial

disturbances propagate poleward as coastal Kelvin waves, shedding Rossby waves along the way (e.g.,

Jacobs et al., 1994). Taken together, these waves form basin modes with eigenfrequencies set by the longest

crossing time of the extratropical Rossby wave (Cessi and Louazel, 2001). Liu (2002, 2003) showed that

the tropical mode (Jin 2001) and basin mode (Cessi and Louazel, 2001) are essentially the same.

The basin mode, if forced by stochastic forcing at mid-latitudes, can cause a decadal oscillation

with period as long as 20 years for the first vertical mode, again unrelated to the ocean-to-atmosphere

feedback (Cessi and Louazel 2001; Liu 2003). On the other hand, Wang et al. (2003) showed that the

tropical mode if coupled with a prescribed atmospheric response can result in a 10-15 year oscillation in the

tropics; they interpreted this variability as a decadal version of the recharge/discharge paradigm for ENSO


proposed by Jin (1997bc).

8. Coupled midlatitude-tropical origin 8.1. Advected temperature changes in the ocean

Gu and Philander (1997) proposed a delayed negative feedback loop between the mid-latitudes

and tropics (see review by Miller and Schneider 2000) as follows; A stronger than normal Aleutian Low

causes SST cooling in the central North Pacific. The cooler waters subduct in the thermocline and

propagate toward the equator following the mean ocean circulation, expressed as the subtropical cells

(McCreary and Lu 1994; Liu 1994). At the equator, these waters upwell to the surface and cool SSTs in the

eastern equatorial Pacific, which then weaken the Aleutian Low via the atmospheric bridge (Lau and Nath

1996; Lau 1997; Alexander et al. 2002), reversing the phase and allowing another half cycle with opposing

polarity. Although Gu and Philander (1997) assumed that the advected temperature anomalies were

imposed at the ocean surface, Schneider et al. (2000) showed that temperature anomalies without density

perturbations due to canceling salinity anomalies can arise from subsurface anomalous advections, and play

an important role in a quasi-decadal oscillation in a CGCM. He called this mode “spiciness mode”.

As described in section 5.4, subducted waters are advected from the North Pacific to the

southwestward (Deser et al. 1996), but the temperature anomalies can be traced only to 18°N (Schneider et

al. 1999). Consistently, Nonaka and Xie (2000) showed that temperature anomalies starting from the central

North Pacific do not reach to the equator in their ocean GCM (OGCM). Also, a recent adjoint assessment

of OGCM (Fukumori et al. 2004) showed that the most of the waters in the eastern equatorial Pacific Niño

3 water do not originate in the central North Pacific, but from the southeastern portion of the North Pacific.

Therefore, advection of the temperature anomalies from the North Pacific is not likely to form a oscillatory

feedback loop.

As noted earlier (section 4.1), Luo and Yamagata (2001) reported that a QDO signature of

subsurface temperature anomalies propagate from the Southern Hemisphere to the equator. They suggested

this is a part of a delayed negative feedback loop responsible for the QDO, combined with possible

ocean-to-atmosphere feedback in the equatorial eastern Pacific, which yields subsurface temperature

anomalies in the southeastern South Pacific. Luo et al. (2003) detected a similar feedback loop in a CGCM.

However, other observational studies (section 4.1, Hasegawa and Hanawa 2003; White et al. 2003) and

numerical studies (section 7.2, Knutson and Manabe 1998; Yukimoto et al. 2000) suggested that

propagating signatures from the North Hemisphere are important in the QDO in the equatorial eastern

Pacific.

8.2. Transport change in the ocean (Subtropical Cell)

In contrast to the changes of advected temperature anomalies, advection speed changes may play

a role in controlling equatorial SSTs. Kleeman et al. (1999) showed that a 3.5 layer ocean model coupled to

a statistical atmosphere model exhibited quasi-decadal (12-yr) variability. Their mechanism is similar to the

mechanism of Gu and Philander (1997), since in both models off-equatorial winds influenced the equatorial


SST via the subtropical cells. However, the major factor in Kleeman et al. (1999) is change in the amount

of cold-water transport into the tropics by the subtropical cell, but the major factor in Gu and Philander

(1997) is the temperature changes that are advected equatorward by the mean subtropical cells. Nonaka et

al. (2002) showed that about half of the simulated SST variability in the eastern equatorial Pacific is forced

by off-equatorial winds (20°−8°S, 8°−25°N) on decadal timescales, via changing the subtropical cells (Fig.

14). This mechanism is supported by following numerical studies (Klinger et al. 2002; Solomon et al. 2003)

and also by the observed spin-down of the subtropical cells from the 1970s to 1990s and the associated

increasing equatorial SST (McPhaden and Zhang 2002).

8.3. Influence of mid-latitude atmosphere on the tropics

Variability in the extratropical atmosphere impacts the tropical Pacific zonal wind stress, and

thereby alters the east-west slope of the equatorial thermocline to preconditions the ocean for altered ENSO

(Barnett et al. 1999, Pierce et al. 2000). This physics involve a coupled ocean-atmosphere process and the

seasonal cycle, called the “seasonal footprinting mechanism” (Vimont et al. 2001). During winter,

atmospheric anomalies in the North Pacific, reminiscent of an intensification of the Aleutian Low, impart

an SST anomaly pattern in the subtropics. During subsequent summer, this SST pattern forces an

atmospheric circulation that influences the equatorial zonal wind stress. Thus, this process communicates

extratropical intrinsic variability to the equator, and initiates coupled processes akin to ENSO. In one

coupled model, the footprinting mechanism accounts for a majority of the decadal variability in the tropical

Pacific (Vimont et al. 2002), and provides stochastic perturbations to ENSO (Vimont et al. 2003).

9. External forcing to the global climate system 9.1. Solar radiation and lunar-solar tide

Several studies suggest that external forcing may induce sizable decadal-to-centennial

oscillations in earth’s climate system, i.e., 11-yr cycle in solar radiation for the QDO (White et al. 1997),

19-yr lunar-solar tide (Royer 1989) or 22-yr double solar cycle or Hare solar cycle (Mitchell et al. 1979;

Cook et al. 1997) for BDOs. The precise mechanisms whereby the solar cycles affect climate are not well

understood. However, although the 11-yr cycle solar radiation change is small (∼0.1%), these radiation

changes influence ozone concentrations and associated solar heating in the stratosphere, which influence

the tropospheric circulation indirectly, at least in one AGCM study (Shindell et al. 1999). The relative

importance of these external forcings and internal variability in the earth’s climate system upon PDV

remains to be examined. There is as yet no proposed external forcing mechanism for the pentadecadal

variability.

9.2. Possible linkage to global warming

Another external forcing that can interplay with natural climate modes may be radiative forcing

due to increasing concentrations of greenhouse gases, sulfate aerosols, and volcano eruptions. Most

CGCMs show that global warming induces an SST anomaly pattern similar to that during the warm phase


of ENSO (see Boer et al. 2004 and references therein). If this is the case, the PDO may have positive trend

in the future, since it is associated with ENSO pattern. Also in the past, ENSO-like decadal variability can

be related to global warming. A CGCM with increasing greenhouse gases showed that weaker warming in

the equator accompanied by stronger warming in midlatitudes, i.e., overall warming plus La Niña-like SST

pattern from 1881 to 1950, but replaced by a overall warming plus El Niño-like pattern from 1951 to 1997

(Cai and Whetton, 2001). Wang and Schimel (2003) suggested that recent warming signatures from 1980 to

2000 can be related to positive trends in the PDO and NAO, and pointed out a possibility that the global

warming-related climate changes may manifest themselves as the primary dynamical climate modes.

10. Predictability

The various mechanisms described in the previous section imply some degree of predictability for

Pacific decadal variability. Stochastic atmospheric forcing of an ocean mixed layer (sections 6.1) implies

SST predictability associated with the thermal inertia of the mixed layer, which can be several years

associated with the reemergence mechanism (Deser et al. 2003). This type of predictability is based on

persistence of a given polarity rather than oscillatory behavior in time. Note that this system implies no

predictability of the atmosphere, as in the conceptual model of the PDO by Newman et al. (2003), unless

there are local feedbacks from the ocean to the atmosphere.

When atmospheric forcing combined with on oceanic propagation (wave or currents, section 6.2),

phase-reversals of oceanic variables can be predicted in the region of downstream of propagation.

Schneider and Miller (2001) proposed a prediction scheme of this type for the January−March KOE SST

based on westward propagating oceanic Rossby waves excited in the central North Pacific, and obtained a

forecast skill (variance reduction) of 0.50 at a lead-time of one-year (Fig. 15). They assumed that the

vertical displacements of thermocline influenced the SSTs, but by taking account of zonal advection of

SSTs due to Rossby waves (Qiu 2003), the prediction skill may be further improved. It should be noted that,

however, Schneider and Miller’s scheme again provide the prediction of oceanic parameters, and prediction

of the atmospheric variability remains untouched.

11. Discussion Previous observational studies revealed decadal variability over and around the Pacific Ocean

and impacts on various physical and environmental parameters. A dominant spatial structure is that

described by the PDO or ENSO-like decadal variability, but two or more modes can be at work. One

candidate is subarctic frontal mode (Nakamura et al. 1997) or north Pacific mode (Barlow et al. 2001, Wu

et al. 2003), and another candidate consists of tropical triangular pattern accompanied by opposing

anomalies in the subtropical front, and is called the subtropical frontal mode (Nakamura et al. 1997), PDO

(Barlow et al. 2001) or tropical Pacific mode (Wu et al. 2003). For timescale dependent patterns or specific

timescales of variability, three major timescales were commonly reported, i.e., quasi-decadal, bidecadal,

and pentadecadal timescales. Future studies are needed to determine whether these variabilities arise from

air-sea coupling, passive ocean responses to stochastic atmospheric forcing, or artifacts of analyses.


As described in sections 6−9, a wide range of hypotheses of mechanisms for the decadal

variability over the Pacific Ocean have been proposed. Most of mechanisms may be summarized, however,

in a single schematic diagram (Fig. 16), which expresses how information is exchanged between different

regions, different vertical layers, and different feedbacks between the ocean and atmosphere. Other

combinations of information exchange than those described above are also possible. This schematic

illustrates the complexity of possible mechanisms for Pacific decadal variability. Thus, the future task for

identifying mechanisms of observed Pacific decadal variability is to evaluate the relative importance of

different combinations of proposed processes, which can then provide a solid foundation for predictability

of the system.

Acknowledgements

1. The authors thank CLIVAR for encouraging us writing this paper, Drs. Bruce Cornuelle, Arthur J.

Miller, David W. Pierce and Jim Potemra for helpful discussions and comments on the manuscript.

This study was supported by grant-in-aid for scientific research (kaken-hi #15540417, to SM) and by

21st century center of excellence program on “Neo-Science of Natural History” lead by H. Okada (to

SM) both from the Ministry of Education, Culture, Sports, Science and Technology, Japan, by the

National Science Foundation (OCE00-82543) and the Department of Energy (DE-FG03-01ER63255)

(to NS). This is IPRC contribution XXXX and SOEST contribution XXXX.


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Figure and Figure Captions

Fig. 1. (top) Anomalous climate conditions associated with the warm phases of the Pacific Decadal Oscillation (PDO), and (bottom) November-March average values of the PDO index. Values shown are °C for sea surface temperature (SST), millibars for sea level pressure (SLP) and direction and intensity of surface wind stress. The longest wind vectors represent a pseudostress of 10 m²/s². Actual anomaly values for a given year at a given location are obtained by multiply the climate anomaly by the associated index value. (after Mantua and Hare 2002).


Fig. 2. Global reanalyzed fields regressed upon a global residual (GR) time series. GR is defined by the

leading principle component of global monthly anomalies of SST, with a 6-yr high-pass filtered cold

tongue index was linearly removed. (top) Sea surface temperature (SST). Contour interval is 0.1 K per

standard deviation of GR (std dev)-1. The zero contour is omitted and negative contours are dashed.

(middle) Surface winds. The reference vector is 1.5 m s-1 (std dev)-1. (bottom) Sea level pressure (SLP).

Contour interval is 0.25 hPa (std dev)-1. The zero contour is omitted and negative contours are dashed.

Regressions are based on all calendar months, from Jan 1958 to Dec 1993. (after Garreaud and Battisti

1999)


Fig. 3. The spatial patterns for the three leading modes of Pacific SST variability during 1945–93 obtained

from rotated principal component analysis: (a) ENSO, (b) Pacific decadal, and (c) North Pacific. (after

Barlow et al. 2001)


Fig. 4. The (inverted) North Pacific SLP Index (NPI; updated from Trenberth and Hurrell, Climate

Dynamics, 1994), a measure of the strength of the wintertime atmospheric circulation over the North

Pacific, and a “Tropical Index”, a synthesis of 6 tropical records from the Indo-Pacific basin, including SST,

SLP, rainfall and cloudiness defined by Deser et al. (2004), during 1900-1997. Both records have been

normalized and smoothed with a 3-point binomial filter. Each tick mark on the ordinate represents one

standard deviation. The correlation between the two time series is 0.72, indicating that approximately half

of the variance in the NPI is accounted for by the Tropical Index. Note the similar “regime” transitions in

both records that occurred in 1925, 1947 and 1977 (dashed vertical lines). (after Deser et al. 2004)


1900 1920 1940 1960 1980 2000-2-1012

a Winter-Spring

hPa

Filtered NPI

1900 1920 1940 1960 1980 2000

-2

0

2

b Winter

hPa

1900 1920 1940 1960 1980 2000-2-1012

c Spring

Year

hPa

Fig. 5. (a) The NPI averaged over winter-spring (Dec.–May) seasons (solid curve), and the respective

averages for 1899–1924, 1925–1947, 1948–1976 and 1977–1997 (dashed lines). Filtered NPI (b) in the

winter–spring, (c) in the winter and (d) in the spring season. The green curves indicate the 10–80-year

band-pass filtered NPI data, the red curves indicate the 10–30-year band-pass filtered (bidecadal filtered)

NPI data, and the blue curves indicate the 30–80-year band-pass filtered (pentadecadal filtered) NPI data.

The regime shifts in the 1920s, 1940s and 1970s are prominent in the winter-spring and winter NPI, but not

in the spring NPI. The shift is marked by the simultaneous phase reversals of the pentadecadal and

bidecadal filtered time series in winter-spring and spring seasons. (after Minobe 1999)

1900 1920 1940 1960 1980 2000

-4-202

hPa

Winter-Spring NPI

Year

d

c

b

a


Fig. 6. (Top) Power spectra of the NPI and

(middle) Tropical Index based upon detrended

data for the period 1900-1997. The thin curves

represent the power spectrum and its 95%

confidence limits for a “red noise” null

hypothesis based upon a first-order

autoregressive process with the same

autocorrelation as the observed time series. Both

spectra exhibit enhanced variance at the lowest

frequencies (periods > ~ 20 years) and a deficit of

power around 8–14 years. The Tropical Index

also exhibits an enhancement of power at

inter-annual frequencies (~ 4–7 years). The

increased power at inter-decadal periods in both

records is significant at the 95% confidence level

based upon a first-order autoregressive model

null hypothesis. (bottom) Squared-coherence

between the NPI and the Tropical Index. The

95% confidence level is 0.6 (thin horizontal line).

The cross-spectrum exhibits high

squared-coherence (~ 0.8) at periods > 20 years

and between 4 and 7 years [values exceeding 0.6

(0.75) are statistically significant at the 95%

(99%) confidence level]. (after Deser et al. 2004)


0.00 0.02 0.04 0.06 0.08 0.10 0.120.0

0.2

0.4

0.6

0.8

1.0

cycle yr-1

Squ

ared

Coh

eren

cy

E-China/S-Japan, Hawaii, Florida, mid-lat. E-N-Am Samoa, SE-Austraila SW-Austraila

90%95%97%99%

1900 1920 1940 1960 1980 2000

-500

50

-2-1012

mm

Mon

th-1

hPa

A. Band-PassedHawaii

Fig. 7. (a) Squared coherency between the NPI and area-averaged gauge precipitation time series from

1900–1995 over eastern-China/southern-Japan (solid black line), Hawaii (solid red line), Florida (solid blue

line), and mid-latitude eastern North America (solid green line), Samoa Islands (dashed black line),

southeastern Australia (dashed red line), and southwestern Australia (dashed blue line) in boreal winter. (b)

Precipitation time series of Hawaii averaged for gridded gauge data (solid red line), which is extended from

1995−2001 by the average of three gauge-measurement stations (Honolulu, Lihue and Hilo) (dashed red

line), along with the NPI (green line, right-hand axis). (after Minobe and Nakanowatari 2002).


Fig. 8. Dissolved oxygen concentration, phosphate concentration, and temperature observed in (a) eastern

Japan sea basin, (b) Yamato basin both in the Japan Sea, and (c) subpolar western North Pacific near Japan,

along with NPI. Smooth curve is subject to bidecadal filter plus trend components. (after Watanabe et al.

2003)


Fig. 9. Schematics showing the sea-surface height phase maps under a monochromatic wind forcing in the

east. The forcing period is short in (a), intermediate in (b), and long in (c). In the schematic, L denotes the

length of the boundary jet, C is the distance from the forcing to the center longitude of the jet, and A and B

denote the southern and northern latitudes bounding the jet, respectively. Each phase contour is separated

by π the initial phase line (θ=0) is assumed to be meridional. (after Qiu 2003)


Fig. 10. The Tropical Pacific Mode (TPM) and North Pacific Mode (NPM) in a 400-year control simulation

in the Fast Ocean Atmosphere Model. The pattern of (a) TPM and (b) NPM are obtained as the 1st and 3rd

REOF of the low-passed (> 7 years) SST in the Pacific basin. The power spectra are shown for the

projections of the Pacific SST on (c) TPM and (d) NPM. Levels of 50% and 95% statistical confidence are

indicated. As in the observation, the TPM has an ENSO-like pattern, while the NPM has the dominant

variability in the western North Pacific; at lower frequency regime, the NPM is dominated by multi-decadal

variability while the TPM by decadal variability. (after Wu et al., 2003).

(c) TPM Power Spectrum

(b) NPM: CTRL REOF3 (9%) (d) NPM Power Spectrum

(a) TPM: CTRL REOF1 (21%)


Fig. 11. (a) Composite difference maps of upward latent heat flux (W m-2) at the surface for Dec.–Jan.

between the two 4-year periods of 1968/69–1971/72 and 1982/83–1985/86, corresponding to the typical

flux anomalies (doubled) for the warm period of the subarctic frontal zone (indicated with a rectangle)

around 1970 associated with the North Pacific decadal variability. Regions of enhanced heat loss from the

ocean are indicated with dark colors. Coloring convention is indicated below. Note that enhanced

(suppressed) heat loss over cold (warm) SST anomalies over the Gulf of Alaska (central North Pacific)

indicates atmospheric forcing on SST anomalies in these regions, whereas enhanced heat loss in the warm

subarctic frontal zone suggests oceanic influence on the atmosphere. (b) As in (a), but for a direct

contribution to (a) from local SST anomalies. (after Tanimoto et al. 2003)


Fig. 12. Time series of ‘‘forecast’’ and observed PDO. A forecast PDO in a year is estimated from a ENSO

index in the same year and PDO in the previous year. (after Newman et al. 2003)


Fig. 13. Time-lagged composite ocean heat content and surface wind anomalies for the decadal timescales

from the R30 coupled model control experiment. The sequences of maps depict approximately one-half of

a life cycle of the composite anomalies. The ocean heat content anomalies are defined as the vertically

integrated temperature anomaly over model layers 3–5 (69–237 m), with a contour interval of 30°C m.

Values less than -30 and greater than +30 are depicted by light and dark shading, respectively. For the

surface wind anomalies, a reference vector of 3 m s-1 is shown at the bottom of the diagram. (after Manabe

and Knutson 1998)


Fig. 14. (a) Observed annual-mean (thin lines with dots) and decadal (thick lines) SST anomalies from the

Comprehensive Ocean Atmosphere Data Set (black curves) and temperature anomalies from the control run

at 15 m (red) in the eastern equatorial region (140°–90°W, 1°S–1°N). The decadal curves are obtained by

applying a 9-year, weighted, running-mean filter. (b) As in (a), except showing decadal anomalies for the

control run (red) and the NoEQ (forced by off-equatorial winds, green) and EQ (forced by equatorial winds,

purple) solutions. (c) As in (b), except showing interannual anomalies for the control (red) and NoEQ

(green) solutions. Trends are removed from all curves. (after Nonaka et al. 2002).


Fig. 15. Time series of FMA SST anomalies in the KOE along 40°N, 140°–170°E. Connected red dots are

observations from the reanalysis; thick red line is the 3-yr average. Solid black line is the hindcast of SST

anomalies from Eq. (6). Purple line is observed (White 1995) 100–400-m temperature, a proxy for

thermocline depth. Blue diamonds are the forecast for FMA of 2002, 2003, 2004, and 2005 obtained from

reanalysis winds up to May 2001. (after Schneider and Miller 2001)


Fig. 16. Schematic showing how information carried by different physical processes in different layers

(thermocline, ocean mixed layer, atmosphere) between different regions. Combinations of items shown in

this figure correspond to mechanisms proposed by previous studies. For example, mechanism of Latif and

Barnett (1994; 1996) can be expressed by the Rossby wave from the central-eastern Pacific to

Kuroshio-Oyashio Extension, which feedback to the atmosphere, and atmosphere influences the

central-eastern Pacific. Gu and Philander (1997) mechanism can be expressed by advection arrow from the

central-eastern Pacific to equatorial upwelling region, from where teleconnection feedbacks the

central-eastern Pacific.

pacific decadal variability: a review · the pdo pattern is similar to the sst variability...

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