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The very strong coastal El Niño in 1925 in the far-eastern Pacific

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Abstract

The 1925 El Niño (EN) event was the third strongest in the twentieth century according to its impacts in the far-eastern Pacific (FEP) associated with severe rainfall and flooding in coastal northern Peru and Ecuador in February–April 1925. In this study we gathered and synthesised a large diversity of in situ observations to provide a new assessment of this event from a modern perspective. In contrast to the extreme 1982–1983 and 1997–1998 events, this very strong “coastal El Niño” in early 1925 was characterised by warm conditions in the FEP, but cool conditions elsewhere in the central Pacific. Hydrographic and tide-gauge data indicate that downwelling equatorial Kelvin waves had little role in its initiation. Instead, ship data indicate an abrupt onset of strong northerly winds across the equator and the strengthening/weakening of the intertropical convergence zones (ITCZ) south/north of the equator. Observations indicate lack of external atmospheric forcing by the Panama gap jet and the south Pacific anticyclone and suggest that the coupled ocean–atmosphere feedback dynamics associated with the ITCZs, northerly winds, and the north–south SST asymmetry in the FEP lead to the enhancement of the seasonal cycle that produced this EN event. We propose that the cold conditions in the western-central equatorial Pacific, through its teleconnection effects on the FEP, helped destabilize the ITCZ and enhanced the meridional ocean–atmosphere feedback, as well as helping produce the very strong coastal rainfall. This is indicated by the nonlinear relation between the Piura river record at 5°S and the SST difference between the FEP and the western-central equatorial Pacific, a stability proxy. In summary, there are two types of EN events with very strong impacts in the FEP, both apparently associated with nonlinear convective feedbacks but with very different dynamics: the very strong warm ENSO events like 1982–1983 and 1997–1998, and the very strong “coastal” EN events like 1925.

Keywords

Coastal El Niño ENSO Eastern Pacific Wind-evaporation-SST feedback Peru Ecuador 

1 Introduction

“El Niño” was first introduced to the scientific community in reference to the anomalous climatic event that took place in 1891 along the coast of Peru, described as an abnormal intrusion of warm oceanic water from the north, replacing the normally cold coastal-upwelled water and favoring the occurrence of strong rainfall and flooding in the otherwise arid northern coast of Peru (Carranza 1891). The warm southward ocean flow was named “Corriente del Niño” (Child’s current) in reference to the weaker climatological version of this current that is normally present after Christmas time (Carrillo 1893).

Nowadays, the term “El Niño” (EN) is used as shorthand for referring to the warm phase of the large-scale El Niño–Southern Oscillation (ENSO) coupled ocean–atmosphere phenemenon, characterized by anomalously high SST in the central–eastern equatorial Pacific (El Niño) and the reduction of the zonal gradient in sea level pressure across the basin (Southern Oscillation; SO). However, the relation between EN and the SO is not always strong (Deser and Wallace 1987), while the SO has been shown to exist even without the ocean dynamics generally associated with EN (Clement et al. 2011). In practice, the definition of “El Niño” is therefore more a matter of convenience of its users than a strict scientific result (Trenberth 1997).

Although the essential physics of ENSO have been largely identified (e.g. Neelin et al. 1998), recent research is focused on understanding the diversity among the individual events, for which a popular procedure is to classify these based on whether the maximum SST anomalies are predominantly found in the central or eastern Pacific (see review by Capotondi et al. 2015), although this classification is somewhat arbitrary (Takahashi et al. 2011). One possibly “true” distinct type of EN could consist of the extreme EN of 1982–83 and 1997–98, as they appear to correspond to a different dynamical regime from the rest of EN due to the nonlinear activation of deep convection in the cold eastern Pacific (Takahashi et al. 2011; Takahashi and Dewitte 2016). These extreme events have been associated with intense warming in the FEP and disproportionally large rainfall anomalies in the arid western coast of South America (e.g. Woodman 1985, 1999; CAF 2000).

From the perspective of the FEP impacts, the only other “very strong” event in the last century was the 1925–26 EN (Quinn et al. 1987; Fig. 1a) and by several other measures in this region, this EN can be considered among the three strongest, with 1982–83 and 1997–98 (Fig. 1b–f). It was due to the detailed report of this event by Murphy (1926) that “El Niño burst onto the international scientific scene” as a legitimate research topic (Cushman 2004), leading to the discovery by Berlage in 1929 (Cushman 2004) of the statistical relation between EN variability, using an index of rainfall in northern Peru (Eguiguren 1894) as a proxy, and the large-scale atmospheric Southern Oscillation (Walker 1924), culminating with the concept of ENSO, an essentially coupled ocean–atmosphere phemonenon (Bjerknes 1969).
Fig. 1

El Niño-related indices for the twentieth century: a El Niño magnitudes estimated according to its coastal manifestations (Quinn et al. 1987, updated by Quinn 1992; the 1997–98 event was added), b annual rainfall (cm) in Guayaquil, coastal Ecuador [Sep(−1)–Aug(0)], annual mean discharge (m\(^3\)/s) of the Zaña (c) and d Viru rivers, e Prosopis pallida annual growth ring width (rainfall proxy) from Casma, coastal Peru (Rodriguez et al. 2005), f SST anomaly in Puerto Chicama (°C), g monthly detrended sea level anomaly (cm) at Balboa, h the equatorial SST “cold tongue index” (°C), and i the Southern Oscillation Index (reversed axis). The vertical grey lines correspond to January 1925, 1983, and 1998

Although the 1925–26 EN was relatively well documented at the time, some of the ideas that appeared well justified at the time, particularly the role of northerly winds (Schott 1931), have been discarded in the subsequent years with the establishment of the ENSO paradigm in the 1970s and 1980s (Wyrtki 1975; Wallace et al. 1998; Neelin et al. 1998), but without taking a close look at the 1925 EN. Thus, it is timely to revisit the 1925–26 EN in an integrated way, under the light of modern theory and expanded datasets, to recover potentially valuable information and insights on the nature of EN and its diversity.

2 Data sources and processing

Monthly series for Puerto Chicama SST (7.7°S, 79.4°W, 1925–2002), the Piura river discharge (1925–1998), Piura rainfall (1932–2008), and estimated EN magnitudes of Quinn (1992) were obtained from the JISAO data archive (http://jisao.washington.edu/data). The monthly Piura discharge data for March and April 1925 were absent, but these values were estimated as discussed in "Appendix B".

Monthly precipitation for Milagro (1921–1981) and Guayaquil (1915–2000) and mean air temperature for Iquique (1900–1988) from the NOAA NCDC GHCN v2 database were obtained from the IRI Data Library (http://iridl.ldeo.columbia.edu/). Annual precipitation values were calculated based on the hydrological year from September of the previous year to August. We also used the annual tree-ring width series for an individual of Prosopis pallida (locally known as “algarrobo”) in San Rafael, Casma (near the coast at 9.5°S), for 1908–2002, a proxy for annual precipitation (Rodriguez et al. 2005).

Daily “research quality” mean sea surface height (SSH) for Balboa (8.97°N, 79.57°W, 1907–2012), on the Pacific side of the Panama Canal, was obtained from the University of Hawaii Sea Level Center (http://uhslc.soest.hawaii.edu/). Monthly means were calculated from this data. The climatology for daily anomalies was calculated using six harmonics of the annual period fitted to the daily data. Anomalies were linearly detrended over the full period.

We used ship-based ocean surface data from the ICOADS database (Worley et al. 2005; Woodruff et al. 2011) in two formats. Firstly, we used the gridded ICOADS 2-degree Enhanced v. 2.5.2 monthly summaries, specifically the mean SST, wind, and cloudiness, which we obtained from the NOAA ESRL website (http://www.esrl.noaa.gov/psd/data/gridded). This dataset consists averages of individual observations over the corresponding 2° grid boxes, with no spatial interpolation. For broad-scale mapping, we constructed 3 month-mean SST and surface wind anomalies during the 1925–26 EN from the gridded monthly ICOADS dataset, with no smoothing or interpolation, limiting the results to those grid cells that contained at least 3 and 10 observations for SST and wind, respectively (wind has stronger high frequency variability). The climatology was constructed from the same data by first averaging temporally and then filling the spatial gaps with a first guess and 9-point smoothing ten times to merge the filled values with the averaged ones. Additionally, we constructed monthly time-series for equatorial segments along four shiptracks that had good data coverage (Fig. 3a).

We also used the individual ship observations from ICOADS Release 2.5 (dataset ds540.0 at NCAR CISL RDA) to produce monthly averages every degree latitude from 9° to 30°S along a well-transited shipping route from Panama to the coasts of Ecuador, Peru and Chile, that started after the opening of the Panama Canal in 1914. Because of this and the onset of World War II in the Pacific, there is good data availability along this track between the years 1920 and 1942. Therefore, unless explicitly indicated, our base period for climatologies for all variables is 1920–1939.

We used a database of approximately 2500 news articles from the Peruvian newspaper “El Comercio” for the period from January 1925 to December 1926, focusing particularly on mentions of meteorological or hydrological phenomena in the coast of Peru (Chang 2014), as well as a database of articles from the newspaper “El Tiempo” from the Piura region (Rojas-Rosas 2014). Numerical meteorological and river discharge data was also retrieved from the newspapers, particularly from the Hipólito Unanue meteorological station in Lima and the daily raingauge data from the Harvard Observatory in Arequipa. Since El Comercio is based in Lima and communications with the northern coast were not in real-time, the news articles often lag the actual events and date and time were seldom reported precisely (Chang 2014), so we indicate the publication date and page number for the relevant newspaper articles in footnotes.

We obtained 18 temperature profiles from the Arcturus expedition (Beebe 1926), which were made in a broad region between Panama and the Galapagos islands between March 30 and June 9, 1925 (Table S1). Four oceanographic profiles of temperature and salinity made in the upper 200 m by a British cruise (NODC code GB012817) along the coast between Ecuador and Panama between April 17–19, 1925 (Table S1) were retrieved from the World Ocean Database 2013 (WOD13; http://www.nodc.noaa.gov/OC5/SELECT/dbsearch/dbsearch.html). We also digitized selected meteorological, hydrological, and oceanographic data from tables and graphs in Murphy (1926), Zegarra (1926), Berry (1927), Zorell (1929), Bailey (1930), Schott (1931), Sheppard (1930, 1933), Petersen (1935), Schaeffer et al. (1958), Woodman (1985), and Reparaz (2013).

Two long-term indices used to represent the basin-scale ENSO variability are the cold tongue index (CTI; SST anomaly in 6°S–6°N, 180°–90°W, 1845–2011) from the JISAO data archive (http://jisao.washington.edu/enso/), and the Southern Oscillation Index (SOI; normalized pressure difference between Darwin and Tahiti for 1866–2013) from the CRU website (http://www.cru.uea.ac.uk/cru/data/soi/, based on Ropelewski and Jones 1987).

We complemented the analysis with the following reconstructed observational products, which were trusted only to the extent that they were consistent with actual in situ data: The Hadley Centre Global Sea Ice and Sea Surface Temperature (HadISST) v1.1 (Rayner et al. 2003) and the NOAA Extended Reconstructed SST (ERSST) v3b, Smith et al. 2008) SST products, the SODA 2.2.4 ocean reanalysis (Giese and Ray 2011; http://apdrc.soest.hawaii.edu/), and the NCEP twentieth century Reanalysis v2 (hereafter 20CRv2; Compo et al. (2006, 2011); http://www.esrl.noaa.gov/psd/data/gridded).

3 Ocean–atmosphere evolution and processes

The large-scale evolution of the 1925–26 EN followed approximately the development phases of the “canonical El Niño” (Rasmusson and Carpenter 1982, hereafter RC82; Harrison and Larkin 1998). The “onset phase” in November 1924–January 1925 was characterized by anomalously cool conditions and easterly equatorial wind anomalies in the central and eastern equatorial Pacific (Figs. 2a, 3b–e), followed by strong anomalous warming near the coast of South America peaking in March 1925 (Figs. 2b, 3b).
Fig. 2

Seasonal mean SST (°C) from ICOADS (colors, shown only for at least three observations per grid cell) and surface wind vectors anomalies (shown only for at least ten observations per grid cell and a minimum magnitude of 1 m/s). Also shown is the SST anomaly reconstruction HadISST 1.1 (contours; interval: 0.5 °C, slight smoothing). The averaging periods are indicated in each panel

Important departures from the RC82 and HL98 composites are that (1) the warming off Peru took place a couple of months earlier than the corresponding “peak phase” of RC82, so that it coincided with the warmest months (Takahashi 2005) that is most favorable for deep convection (e.g. Takahashi 2004; Huaman and Takahashi 2016), while (2) cool conditions remained in the central-eastern Pacific, expanding to the west, until June 1925 (Figs. 2b, c, 3e), which we later argue is also important for the convective dynamics in the FEP (Sects. 3.2, 3.4).
Fig. 3

a Segments of the four ship tracks with the best data coverage in the central and eastern equatorial Pacific during the period 1920–39 (numbered 14 westward from the coast; standard Niño 1+2 and 3.4 regions indicated) and raw (thin) and 1-2-1-smoothed (colors, thick) monthly be SST (°C) and fi zonal wind (m/s) anomalies along each of the ship tracks based on the gridded ICOADS data

The warming was confined to the coast in February–March 1925 but then spread westwards, reaching the central Pacific (~170°W) around August 1925 and peaking by the end of the calendar year (Fig. 3b–e), corresponding to the “mature phase” of the canonical EN (RC82). Westerly wind anomalies started developing in the central Pacific in March 1925 (Fig. 3h–i), indicating the onset of the Bjerknes feedback associated with the FEP warming (Dewitte and Takahashi 2016) and leading to the establishment of the warm ENSO phase.

In the next subsections we address specific mechanisms relevant to the development of this EN event.

3.1 Lack of initial Kelvin wave forcing

In the ENSO paradigm, the coastal warming is associated with downwelling equatorial Kelvin waves (e.g. Chiodi et al. 2014) that depress the thermocline and raise the sea level. However, in contrast to other EN events, the sea level data at Balboa, Panama, does not show positive anomalies in early 1925, and in fact they are negative in February (Fig. 4). Furthermore, we find that large sea level anomalies (e.g. >10 cm) in other EN events are associated with substantially weaker monthly SST anomalies than the one observed at Puerto Chicama in March 1925 (Fig. 4), with the exception of the peaks of the extreme 1982–83 and 1997–98 EN (Fig. 4d, f). Thus, a sea level anomaly associated with a downwelling Kelvin wave sufficiently strong to account for the observed warming in early 1925 could be expected to have been greater than 20 cm.
Fig. 4

Daily detrended sea surface height anomaly at Balboa (cm; colors) and monthly sea surface temperature anomaly (°C; grey lines) for Puerto Chicama for selected El Niño events

On the other hand, the second warming peak in 1925, starting in November, coincides with a 20 cm sea level anomaly, indicating that a downwelling equatorial Kelvin wave pulse was responsible (Fig. 4a). Downwelling Kelvin waves at this time are consistent with the westerly wind anomalies in the central Pacific (Fig. 3h–i), although this data cannot resolve specific westerly wind events.

Predominantly positive sea level anomalies were observed starting in May 1925 (Fig. 4a), after which the positive SST anomalies extended southward to 30°S (Fig. 5b). Particularly, the warming pulses in August–September 1925 and December 1925–February 1926 coincide with the positive sea level pulses around those times (Fig. 4a), providing further support to their interpretation as downwelling Kelvin wave pulses that were able to propagate the warming signal polewards into Chile, whereas the warming peak in March 1925 was restricted to north of 20°S (Fig. 5b).
Fig. 5

a Monthly sea surface temperature (shading, °C) and surface wind (m/s), and b the anomalies (with respect to 1920–39) from ICOADS along the near-coastal track. The zero meridional wind contour is solid white in a

Additional evidence of the absence of downwelling equatorial Kelvin wave forcing is provided by the hydrographic measurements from the Arcturus expedition east of the Galapagos (March 30–June 9, 1925) and the ship GB012817 along the coast of Colombia and Ecuador (April 17–19, 1925). The data from both indicate SST anomalously above 27 °C near the equator, sharply decreasing to 19–20 °C at the 50 m depth (Fig. 6a, c). The equatorial anomalies calculated from the SODA climatology transition from generally positive at the surface to negative at 100 m (Fig. 6b, d, e). Conversely, strong downwelling Kelvin waves would result in a deep (>100 m) warm layer (Cucalon 1987; Garcés-Vargas et al. 2005; see profiles for 1983 and 1998 in Fig 7b, d, f, h).
Fig. 6

ac Temperature and df its anomalies at a, d 0 m, b, e 50 m, and c, f 100 m based on hydrographic measurements from the Arcturus (circles; March 30–June 9, 1925, starting on the east) and ship GB012817 (times symbols; April 17–19, 1925, starting on the south). The anomalies are calculated with respect to the SODA 1920–1939 climatology, except around the Galapagos were SODA is less reliable

Fig. 7

a Location of four oceanographic stations from the GB012817 ship (labeled 14) in April 17–19, 1925 (starting with station 4). The surface temperature (°C) and salinity (pps) are indicated in red and blue, respectively. Observed profiles of b, d, f, h) temperature and c, e, g, i) salinity (solid) are shown for stations b, c 1, d, e 2, f, g 3, and h, i 4. Also included are the climatology for April (1920–39, thick short dashed), and the data for April 1983 (long dashed) and 1998 (dot-dashed) for the nearest grid cell from SODA 2.2.4

The zonal wind showed weak positive anomalies (<1 m/s) in the FEP prior to the warming (Fig. 3f, g), which could have helped with the warming.

3.2 Northerly winds and the ITCZ

The most outstanding aspect of the atmospheric circulation in early 1925 was the extreme southward extension of the Panama wind jet, which climatologically reaches the equator in February and March (Fig. 8b, c) but in 1925 extended to 8–9°S and fanned out to the Galapagos (Fig. 8f, g). Similarly, the Papagayo jet further west extended almost to the equator near 95°W (Fig. 8f, g). The northerly anomalies were on the order of 2–3 m/s (Figs. 3j, k, 5b, 9), substantially stronger than in the EN composite of (Harrison and Larkin 1998) for February (~0.6 m/s around 3°N in their Fig. 6) and the RC82 composite for March–May (Fig. 19b in RC82). The onset of these northerly winds from January to February was abrupt, as was their retreat from March to April (Figs. 3j, k, 5a, b, 8a, b).
Fig. 8

ICOADS sea surface temperature (shading, °C) and surface winds (m/s) for January through April for the 1920–39 climatology (top row), and year 1925 (bottom row). The 26 °C isotherm (solid and light purple) and the zero meridional wind contour (dashed) are included

The enhancement of the northerly winds could have been the result of external atmospheric forcing. In the case of the gap jets through Central America, the forcing could be associated with the SLP difference between the Atlantic and the Pacific (Karnauskas et al. 2008). The ICOADS data shows positive SLP anomalies around Central America in both February and March 1925, with negative anomalies off Ecuador only clearly in March (Fig. 9a, b). However, although these (absolute) northerly winds were connected to the northeasterlies in the Caribbean, the northerly anomalies themselves were limited to the Pacific (Figs. 5b, 9). Furthermore, in situ observations of the integrated January–April northerly wind at Balboa and Cristobal, at both ends of the Panama Canal, indicate that 1925 did not have unusually high northerly winds (Schaeffer et al. 1958; Fig. S1). On the other hand, in February–March 1925, the Southern Oscillation was in a positive state (Fig. 1i), suggesting an enhanced South Pacific anticyclone, while the actual wind data does not show substantial subtropical wind anomalies (Fig. 5b), indicating that the northerly wind anomalies were not forced from the south Pacific. Therefore, atmospheric forcing from the Caribbean or the south Pacific does not appear to have had a key role in driving the northerly wind anomalies.
Fig. 9

Monthly anomalies of surface wind (m/s; vectors) overlayed on a, b cloud cover (octas; shading) and c, d) sea level pressure (hPa; shading anomalies from ICOADS for a, c February and b, d March 1925. Only grid cells with at least 15 observations are shown (slight spatial smoothing). Anomaly vectors with magnitude larger than 1 m/s are darker

Another possibility is that local air–sea interaction amplified and maintained the northerly wind anomalies that could have been initiated by weak external atmospheric or oceanic forcing. Particularly, the meridional gradient in absolute SST near the coast was reversed, with the warmest/coolest waters found south/north of the equator (Figs. 5a, 8f, g). The SST anomalies presented a meridional dipole pattern with maxima at °S and 5°N and the northerly wind anomalies in-between (Fig. 5b). This anomalous SST gradient could directly reinforce the northerly winds via thermally-induced pressure gradients (Lindzen and Nigam 1987; Battisti et al. 1999) and, perhaps more importantly, the anomalously high SST south of the equator could strengthen the southern hemisphere ITCZ that climatologically is present in February–April around 5°S (Huaman and Takahashi 2016). The latter is suggested by the positive cloudiness anomaly in the ICOADS data around 2°S in February and, most notably, around 7°S in March (Fig. 9c, d). This enhancement of the SH ITCZ is consistent with the surface wind convergence off the coast of northern Peru (Fig. S2), as well as with the heavy rainfall observed in the otherwise arid northern-central coast of Peru (Sect. 3.4, "Appendix A"). Additionally, the barometric pressure measured at Chicama (7.7°S) at 7 a.m. (to reduce the diurnal land heating effect; the 4 pm data has similar variability but with lower values) indicates a large 9–10 hPa drop from around 1017 hPa in the beginning of January to an average of 1008 hPa in the second half of March (Fig. 10b), consistent with the establishment of the equatorial trough off the coast of Peru, with lower pressure than the March mean value of 1010 hPa according to the TAO buoy at 5°S, 95°W for the years 2001–2003, probably an indication of the intensity of the anomalous SH ITCZ. On the other hand, the ICOADS data shows that the fanning of the Panama jet is associated with net surface wind divergence in the eastern Pacific north of the equator in February and March (but not in the Caribbean) instead of the climatological convergence (Fig. S2), consistent with the reduced cloudiness in the NH in March (Fig. 9d). The above suggests an anomalous local meridional overturning cell in the FEP in February–March 1925 with ascent/descent in the southern/northern hemisphere.
Fig. 10

Daily series for January–April 1925: a sea surface temperature in Puerto Chicama and Callao (°C; open circles and dots, respectively), b barometric pressure in Chicama (hPa; 1-2-1 smoothed), c precipitation in Zorritos (mm; Petersen 1935), d Piura river discharge (m\(^3\)/s; thick circles are discharge reconstructed from water height; small and large black triangles indicate days with moderate and heavy rainfall, respectively, in the city of Piura, Table S2), e Virú (black) and Chicama (grey) river discharge (m\(^3\)/s; Zegarra, 1926), f precipitation in Trujillo (mm; *the accumulated value for March 7–9 was reported on March 9), and g precipitation in Lima (mm; El Comercio, 1925). Data was digitized from Murphy (1926) unless explicitly indicated. Days with missing data are shaded (except for the gaps in the Chicama discharge), and, in the case of precipitation, they are assigned a value of zero (following Petersen 1935). In bf, thick black lines indicate when the near-coastal SST analysis of Schott (1931) exceeded 26 °C at the corresponding locations (for Zorritos, dashed indicates extrapolation)

The existence of an approximate threshold SST for the activation of deep convection (Graham and Barnett 1987; Johnson and Xie 2010; Takahashi and Dewitte 2016; Jáuregui and Takahashi 2017) introduces a nonlinearity that can explain the abruptness of the onset of the SH ITCZ and northerly wind (Xie and Philander 1994, hereafter XP94; Wang and Wang 1999), which would take place more easily in the warm seasonal peak driven by insolation (Takahashi 2005). In this sense, the coastal EN could be described as an amplification of the seasonal cycle.

The wind speed anomalies associated with the northerly wind anomalies present a dipole pattern (not shown), with reduced/enhanced speed in the southern/northern hemisphere, contributing to reduced/enhanced surface evaporation and, therefore, enhanced/reduced SST. This positive wind speed-evaporation-SST (WES) feedback (XP94) probably was key in establishing this coastal EN event, which would imply that it would correspond to shallow solar warming. Another possible mechanism is meridional warm advection, which is discussed in the next subsection.

3.3 The “Corriente del Niño”

Meridional warm advection associated with an anomalously strong southward “Corriente del Niño” (El Niño Current; Carrillo 1893) was the first EN mechanism identified (Carranza 1891; Schott 1931). In 1891, in addition to warm water along the northern coast of Peru, carcasses of crocodiles and tree debris from north of 3°S were found at 8°S (Carranza 1891). In 1925, the report at 4.6°S of a lizard not previously found in Peru but abundant off Ecuador at 3.2°S (Murphy 1926) and, in 1926, of seeds at 4.6°S of mangroves that are found to the north of 3.6°S (Berry 1927), are also indications of southward advection.

From the end of January through February 1925, ship drift data along the coastal track indicate southward flow north of the equator, with a speed on the order of 30–50 cm/s (Zorell 1929). Murphy (1926) also reported coastal measurements of southward flow of around 50 cm/s further south (5 and 2°S) within the same period. Schott (1931) discussed the southward progression of warm SST fronts in early 1925 using cruise data and coastal stations, with the warmest front reaching Puerto Chicama in February 27, Callao in March 12 (Fig. 10a) and as far south as Pisco in March 16 (14°S), from which he inferred a southward propagation speed of 40–50 cm/s, although onshore advection, as observed in the 1982–83 and 1997–98 events (Morón 2011), is another possibility.

Having discarded the possibility of strong downwelling Kelvin waves (Sect. 3.1), the strong northerly winds are the most likely forcing of the warm countercurrent [e.g. Philander and Pacanowski (1981)]. We can produce a rough estimate of the wind-driven surface current u based on the wind observations by neglecting the Coriolis force near the equator and considering the frictional balance \(r_s \mathbf{{u}}={\tau }/\rho H\), where \(r_s\approx\) (2 days)\(^{-1}\) is a frictional dissipation rate (Zebiak and Cane 1987; Dewitte 2000), \(\rho \approx 10^3\) kg/m3 is the water density, \({\tau }=\rho _a C_D|\mathbf{u}_a|\mathbf{u}_a\) is the surface wind stress with \(|{u}_a|{u}_a\) indicating the surface wind pseudo-stress, \(\rho _a=1.2\) kg m\(^{-3}\) the air density and \(C_D=2\times 10^{-3}\) the drag coefficient (Perigaud et al. 2000). Based on the individual ICOADS observations, the near-coastal pseudostress had a mean northerly component of \(15.5\pm 4.7\) m\(^2\)/s\(^2\) around the equator (1°S–1°N) in February–March 1925. Assuming an Ekman layer depth of \(H=30\) m (Oerder et al. 2015), we calculate the equatorial southward current speeds as 18 ± 5 cm/s. This is a lower bound, since the high SST implies reduced atmospheric stability, i.e. larger \(C_D\) and wind stress, while the ocean stratification associated with the shallow fresh warm surface (Fig. 7h, i) suggests a shallower H. Additionally, changing observational practices (i.e. Beaufort scale vs anemometers) and anemometer heights due to increasing ship sizes (Cardone et al. 1990) introduces further underestimation of the wind stress. Thus, the estimated wind-driven current speed provides a lower bound that is consistent with the other observational estimates for the “Corriente del Niño” speed.

Additionally, the northerly wind anomaly enhanced the upwelling in the Panama bight (Alory et al. 2012), with cold water extending almost to the surface, except for a very shallow warm layer (Fig. 7b), and salinities typical of the 100 m depth (Fig. 7c). This is similar to what was observed in 1891 (Schott 1931) but contrasts sharply with the extreme EN conditions in April 1983 and 1998, which featured a deep fresh warm layer of around 80 m depth (Fig. 7a–g).

3.4 Local and remote sea surface temperature control on eastern Pacific precipitation

The dependence of coastal precipitation of SST approximately follows the threshold model of XP94 (see Sect. 3.2), but with a threshold SST of around 26 °C (Woodman 1999; Takahashi 2004; Ramos 2014). This model provides an adequate qualitative explanation for the temporal evolution of the periods of strong rainfall and river discharge events at different latitudes along the Peruvian coast (Fig. 10c–g), as they approximately coincide with the periods in which the local near-coastal SST [based on the ship analysis of Schott (1931)] was greater than 26 °C (indicated with lines in Fig. 10c–g). These conditions were progressively established from north to south and ended in the reverse order. The details of the severe impacts at various sites along western Peru and Ecuador associated with heavy rainfall and flooding in 1925, which can provide important guidance for paleoclimatic and historical EN reconstructions as well as for risk management, are presented in Appendix A.

On the other hand, the cool central Pacific SST is also known to enhance the precipitation in the mid and upper basins along the western Andes of Peru (Lavado-Casimiro and Espinoza 2014). This is explained by the connection of SST in the western Pacific warm pool connects to the whole tropical free troposphere via deep convection, cooling it during basin-scale La Niña conditions (Yulaeva and Wallace 1994; Chiang and Sobel 2002), which reduces the tropospheric stability and facilitates convection (Vecchi and Soden 2007; Jáuregui and Takahashi 2017), while also producing easterly near-equatorial upper-air wind anomalies over South America that are also favorable for convection (Kousky and Kayano 1994; Vuille et al. 2000).

The role of local and remote forcing is verified for the Piura river by the relatively high positive correlation between its annual mean discharge (reconstructed as described in Appendix B) with February–March mean SST in the FEP (\(r=0.69\) in the Niño 1+2 region; Fig. 11a, b), and negative values in the western–central equatorial Pacific (\(r=-0.17\) in the region we denote as \(T_w\) (155°E–175°W, 5°S–5°N; Fig. 11a, c). On the other hand, we see in the scatter plots that the two relationships are not linear and that the correlation with Niño 1+2 is strongly dominated by the 1982–83 and 1997–98 events, which are outliers in the correlation with \(T_w\) (Fig. 11b, c). However, if we simply subtract \(T_w\) from Niño 1+2, providing a rough index of tropospheric stability in the eastern Pacific, we not only find an enhanced correlation (\(r=0.72\)), but generally a more monotonic nonlinear relationship (Fig. 11d). The high SST in \(T_w\) during the warm ENSO phases in 1926 and 2016 explain why the river discharge was not as high in those years despite the high Niño 1+2 SST, whereas 1925 and 2008 had high discharges due to the low SST in \(T_w\) and high Niño 1+2 SST. The 2008 case is interesting because it was primarily regarded as a basin-scale La Niña (cool ENSO) event (Bendix et al 2011). The nonlinearity of the relation between Niño 1+2 \(-T_w\) and the discharge, with a sharp increase in slope above −1.5 °C, can also explain why the strong EN (warm ENSO) failed to produce rainfall in northern Peru as strong as in 1983 or 1998 despite having similar Niño 3.4 SST (L’Heureux et al. 2016).
Fig. 11

a Linear correlation between the annual discharge averaged for the Piura river with the February–March SST from HadISST 1.1 (1925–2016). Scatter plots between the same discharge and the SST averaged over b the Niño 1+2 region, c the Tw region (155°E–175°W, 5°S–5°N), as well as with d the difference between the two (Niño 1 + 2 minus Tw)

We hypothesize that a similar process reduced the stability of the seasonal southern hemisphere ITCZ in 1925, which in the context of the XP94 model could have been through a combination of lowering of the threshold SST and/or enhancing the rate of the precipitation increase with SST due to basin-scale La Niña (cool ENSO) conditions, which cools the free troposphere and produces easterly upper-air anomalies (both of which are indicated by the 20CRv2 in 1925, not shown). In this view, this type of coastal EN would be the result of the interannual destabilization of the seasonal cycle in the FEP by the cool ENSO phase.

4 Discussion

The ENSO paradigm is based on the interaction between equatorial ocean dynamics and zonal winds through SST. Historically, however, the association of northerly winds in the FEP with EN had been noted by Eguiguren (1894) and Murphy (1926), while Schott (1931) went further to propose that these winds were the forcing of the coastal EN. This hypothesis subsequently was countered by the finding that the coastal winds tend to strengthen during EN (Wyrtki 1975; Enfield 1981; Rasmusson and Carpenter 1982) and Wooster (1980) argued that Schott failed by “underestimating the magnitude of the time and space scales involved”. But none of this later studies explicitly analyzed the 1925 and their failure was in implicitly assuming that the same mechanisms act in the same way in every event, despite Wyrtki’s (1975) conclusion that “El Niño certainly does not have only a single cause”.

Nevertheless, the RC82 and the Harrison and Larkin (1998) EN composites do show northerly wind anomalies during the coastal warming phase, but weak compared to 1925. The seasonality of these anomalies appears to be critical for the strong feedback between SST, the ITCZ and the northerly wind, as the SST and the ITCZ off Peru peak climatologically around March (Takahashi 2005; Huaman and Takahashi 2016). We argue that this feedback was made more effective in 1925 by the dominant cold conditions in the rest of the equatorial Pacific, which promoted convection in the FEP by the destabilition of the troposphere and with moist easterly advection from the Amazon. However, strong northerly anomalies were also observed in early 1926, around the peak of the warm ENSO phase (Fig. 12a). This suggests that perhaps longer-term changes, like a lower convective threshold for convection (Johnson and Xie 2010), could have made this mechanism more effective in the past. In fact, the latitude of the trade-wind confluence in these two years has been the lowest in the 1920–2012 period, including the extreme 1982–83 and 1997–98 events and there is a slight (but not significant) northward trend in this latitude, perhaps a response to the stabilization associated with the long-term tropical tropospheric warming (Johnson and Xie 2010; Jáuregui and Takahashi 2017). Consistent with this, the low tropospheric stability estimated as the difference between the 700 hPa potential temperature from the 20CRv2 and SST in the Niño 1+2 region (not shown) also has a small albeit not significant positive trend. On the other hand, the SST difference between the Niño 1+2 and the \(T_w\) region, a stability proxy for the FEP (see Sect. 3.4) does not show a clear trend (Fig. 12b), although uncertainty in SST reconstructions is an issue for trends in the zonal SST gradient in the tropical Pacific (Deser et al. 2010). On the other hand, many climate future climate change projections with global climate models indicate a joint trend in increasing rainfall and northerly wind anomalies off northern Peru similar to the proposed for the 1925 EN (Belmadani et al. 2014), but the models in general continue to have strong biases in this region, particularly the double ITCZ syndrome (Zhang et al. 2015).
Fig. 12

a Latitude of zero meridional wind along the near-coastal track based on the February–March mean wind from ICOADS, the b February–March mean Niño 1+2 minus Tw SST (°C) from HadISST 1.1 (black, circles) and ERSST v3b (grey, crosses) and c the annual (Jan–Dec) mean Piura river discharge (m3/s). Linear fits are shown dashed. Vertical lines indicate the years 1925, 1983, and 1998

If we focus on the extreme EN impacts in western tropical South America (e.g. Figs. 1a–f, 12c), both the large-scale version (1982–83, 1997–98) or this coastal version have similar signatures. So not only is the record too sparse to identify trends that differentiate between the two types, but the two trends would be responding to different processes. This is important for long-term reconstructions of ENSO diversity. The interpretation of paleoclimatic and historical records of extreme EN impacts in the FEP (high temperatures, flooding, impacts on marine ecosystems) would need to carefully consider these two types, ideally contrasting with proxies from other regions that help with the discrimination of the two.

5 Conclusions

In this study we revisited the very strong 1925 El Niño (EN), the third strongest, after 1982–83 and 1997–98, in terms of its impacts on the far-eastern Pacific (FEP) associated with heavy rainfall on the coasts of southern Ecuador and Peru in February–April 1925. In situ instrumental records and extensive newspaper reports allowed us to reconstruct the ocean–atmosphere evolution in the tropical Pacific, particularly in the FEP, as well as the large extent of the impacts associated with heavy rainfall and flooding in coastal Peru and Ecuador, which extended as far south as 12°S.

The 1925 EN event was the one that introduced the concept of El Niño to the scientific community. However, this very strong and archetypical EN event in terms of its FEP signature was restricted to this region and coincided with anomalously cold conditions in the rest of the equatorial Pacific. This “coastal” EN took place in February–April 1925, in the transition from a cool to a warm ENSO state. However, the coastal warming was not associated with the equatorial zonal dynamics that are the essence of ENSO. Instead, in situ hydrographic data in the FEP indicate that the warming was shallow and the tide-gauge data at Balboa indicates that sea level was lower than normal, so downwelling Kelvin waves were unlikely to have played a significant role in producing the coastal warming. On the other hand, in situ ship-based wind data indicate strong northerly winds that reached 7°S on average in February−March 1925, which climatologically only extend to the equator near the coast. The meridional asymmetry in SST was also reversed, i.e. with a warmer southern than northern hemisphere. The ship data also indicates reduced cloudiness, increased sea level pressure, and net wind divergence, north of the equator, and the opposite off northern Peru, indicative of the weakening/strengthening of the ITCZ to the north/south of the equator. The above then indicates a reversal of the north–south asymmetry relative to the equator in the coupled ocean–atmosphere system that includes the ITCZ, meridional wind, and SST.

The abrupt onset of this coastal EN, with the southern ITCZ and northerly anomalies, suggests strong external forcing and/or strong nonlinear coupled feedbacks. Regarding the former, the in situ data indicates that the Panama gap jet was not anomalously stronger and that the south Pacific anticyclone was not weaker in this period, indicating that the associated external forcing was not strong. Thus, it seems more likely that the nonlinearity associated with a threshold in SST for the activation of the ITCZ south of the equator was responsible for the abrupt transition, as in the model of Xie and Philander (1994, XP94). The mechanism for the ocean warming probably involved a combination of the wind-evaporation-SST mechanism (XP94) and oceanic southward warm advection, i.e. the enhanced “Corriente del Niño”. The existence of this warm countercurrent was shown with ship drift data and is consistent with a calculation of a frictional current driven by the observed northerly winds. To the extent that a nonlinear coupled feedback underlies the onset and retreat of the southern ITCZ, the very strong coastal EN can be described as an enhancement of the insolation-driven seasonal cycle, particularly through the destabilization of the seasonal SST-ITCZ-meridional wind dynamics.

We propose that an important mechanism for destabilizing the ITCZ involves the zonal equatorial SST gradient in the Pacific, particularly the contrast between the SST in the FEP and in the western–central Pacific, as the latter affects the temperature in the free tropical troposphere and, therefore its static stability, as well as the zonal upper air moisture flow in the FEP that affects convection. This is shown to be the case using a river discharge record on the coast of Peru at 5°S, the climatological latitude of the southern ITCZ band, which shows a strong nonlinear relation with the difference in SST between the FEP and western-central Pacific, with more strongly enhanced discharge when the former is less than ~1.5 °C colder than the latter. In early 1925, the cold conditions in the central-western Pacific and the warm eastern Pacific both contributed to high river discharges, whereas during the warm ENSO phase the effects of the two regions oppose each other as in 1926 and notably during the strong 2015–16 EN (L’Heureux et al. 2016).

In the context of ENSO diversity, there have been several studies that have classified EN events according to their spatial pattern. But if we focus on those EN events that have very strong impacts on coastal Peru and Ecuador with very high SST and coastal rainfall, we can identify two major types:
  • Very strong warm ENSO events, e.g. 1982–83 and 1997–98, that are associated with the zonal dynamics in the equatorial Pacific and a nonlinear Bjerknes feedback that enhances their growth (Takahashi and Dewitte 2016), which are potentially predictable several months in advance,

  • Very strong “coastal” EN, e.g. 1891 and 1925, with cold to neutral conditions in the rest of the equatorial Pacific and associated with meridional dynamics in the FEP involving the abrupt enhancement of the ITCZ and warming south of the equator and strong northerly winds. Based on our current knowledge, this type of event is not as predictable as the warm ENSO events.

Further studies will be needed to verify the proposed mechanisms, but this will prove challenging as climate models continue to have strong biases in this region, particularly associated with a warm coastal bias and an excessively strong southern ITCZ, perhaps similar to this very strong coastal EN. In situ data in this region, particularly observations of the atmospheric circulation associated with the southern ITCZ, would be very valuable for the validation of both the mechanisms and the models themselves.

Footnotes

  1. 1.

    Due to missing data in June 1925, this is probably slightly underestimated (see Fig. S3b).

  2. 2.

    El Comercio 1925-04-03, p. 1–2.

  3. 3.

    El Comercio 1925-04-18 p. 11.

  4. 4.

    El Comercio 1925-04-17, p. 8.

  5. 5.

    El Comercio 1925-02-12, p. 3; 1925-02-13, p. 2.

  6. 6.

    El Comercio 1925-04-17, p. 8.

  7. 7.

    El Comercio 1925-02-12, p. 3; 1925-02-14, p. 10; 1925-02-21, p. 10.

  8. 8.

    El Comercio 1925-02-17, p. 1; El Comercio 1925-02-27, p. 2; 1925-04-17, p. 8.

  9. 9.

    El Comercio 1925-04-17, p. 8.

  10. 10.

    El Comercio 1925-02-17, p. 1.

  11. 11.

    El Comercio 1925-04-17, p. 8.

  12. 12.

    El Comercio 1925-04-17, p. 8.

  13. 13.

    El Comercio 1925-02-28, p. 6; 1925-04-17, p. 8.

  14. 14.

    El Comercio 1925-03-02, p. 1; 1925-03-06, p. 6; 1925-03-07, p. 3.

  15. 15.

    El Comercio 1925-03-07, p. 3; 1925-04-12 p. 6.

  16. 16.

    El Comercio 1925-04-17, p. 8.

  17. 17.

    El Comercio 1925-04-12 p. 6.

  18. 18.

    El Comercio 1925-04-17, p. 8.

  19. 19.

    El Comercio 1925-04-12 p. 6.

  20. 20.

    El Comercio 1925-04-12 p. 6.

  21. 21.

    El Comercio 1925-04-17, p. 8.

  22. 22.

    El Comercio 1925-04-17, p. 8.

  23. 23.

    El Comercio 1925-03-15 p. 6; 1925-03-24, p. 3.

  24. 24.

    El Comercio, 1925-03-24, p. 6; 1925-04-17, p. 8.

  25. 25.

    El Comercio, 1925-03-24, p. 6.

  26. 26.

    El Comercio 1925-03-24, p. 6.

  27. 27.

    El Comercio 1925-03-30 p. 1.

  28. 28.

    El Comercio 1925-03-30, p. 1, 4; El Comercio 1925-04-14, p. 7; 1925-04-17, p. 8.

  29. 29.

    El Comercio 1925-03-30 p. 1; 1925-04-17, p. 8.

  30. 30.

    El Comercio 1925-03-30 p. 1.

  31. 31.

    El Comercio 1925-04-12 p. 6; 1925-04-17, p. 8.

  32. 32.

    El Comercio 1925-04-12 p. 6.

  33. 33.

    El Comercio 1925-04-15 p. 1,4; 1925-04-17, p. 8.

  34. 34.

    El Comercio 1925-05-11, p. 4.

  35. 35.

    El Comercio 1925-05-11, p. 4.

  36. 36.

    El Comercio 1925-02-18, p. 4.

  37. 37.

    El Comercio 1925-02-28, p. 4

  38. 38.

    El Comercio 1925-02-25, p. 1.

  39. 39.

    El Comercio 1925-03-14, p. 2.

  40. 40.

    El Comercio 1925-03-25 p. 4; 1925-03-27, p. 5; 1925-03-14, p. 2; 1925-04-13, p. 4.

  41. 41.

    El Comercio 1925-03-27, p. 4; 1925-04-03, p. 8; El Comercio 1925-04-13, p. 1,4; 1925-05-11, p. 4.

  42. 42.

    El Comercio 1925-05-11, p. 4.

  43. 43.

    El Comercio 1925-03-27, p. 4,5; 1925-03-28, p. 2; 1925-03-30, p. 2; 1925-04-30, p. 4; 1925-05-11, p. 4.

  44. 44.

    El Comercio 1925-04-03, p. 8; 1925-04-13, p. 1; 1925-04-13, p. 1; 1925-05-11, p. 4.

  45. 45.

    El Comercio 1925-04-21, p. 7

  46. 46.

    El Comercio 1925-05-11, p. 4.

  47. 47.

    El Comercio 1925-04-08, p. 2; 1925-04-13, p. 4.

  48. 48.

    El Comercio 1925-02-25, p. 1; 1925-02-27, p. 3; 1925-03-01, p. 5.

  49. 49.

    El Comercio 1925-03-12, p. 12; 1925-03-15, p. 12; 1925-03-17, p. 1; 1925-03-18 p. 3

  50. 50.

    El Comercio 1925-03-25, p. 5.

  51. 51.

    El Comercio 1925-03-17 Mp.3 2do; 1925-03-18 p. 1; 1925-03-23, p. 2; 1925-03-24 p. 1.

  52. 52.

    El Comercio 1925-03-18 p. 1.

  53. 53.

    El Comercio 1925-03-27, p. 4; 1925-03-29 p. 2; 1925-04-09 p. 11.

  54. 54.

    El Comercio 1925-05-02, p. 8; 1925-06-02, p. 1.

  55. 55.

    El Comercio 1925-03-25, p. 5.

  56. 56.

    El Comercio 1925-03-25, p. 5; 1925-03-25, p. 2.

  57. 57.

    El Comercio 1925-04-02, p. 2; 1925-05-02, p. 8; 1925-06-02, p. 1.

  58. 58.

    El Comercio 1925-04-02, p. 2.

  59. 59.

    El Comercio 1925-04-02, p. 2.

  60. 60.

    El Comercio 1925-02-26, p. 1, 5.

  61. 61.

    El Comercio, 1925-03-21, p. 7; 1925-03-21, p. 3; 1925-03-29, p. 8; 1925-04-02, p.4; 1925-04-29, p. 4.

  62. 62.

    El Comercio 1925-04-02, p. 4.

  63. 63.

    El Comercio 1925-03-27, p. 4.

  64. 64.

    El Comercio 1925-04-04, p. 8.

  65. 65.

    El Comercio 1925-03-30, p. 2.

  66. 66.

    El Comercio 1925-03-30, p. 2; 1925-04-03, p. 1.

  67. 67.

    El Comercio 1925-04-16, p. 6.

  68. 68.

    El Comercio, 1925-03-12.

  69. 69.

    Based on a vapor pressure of 21–23 hPa from his Fig. 4 (corrected by a factor of 10).

  70. 70.

    El Comercio 1925-02-09, p. 7.

  71. 71.

    El Comercio 1925-02-18, p. 2; 1925-02-22, p. 2; 1925-03-02, p.3; 1925-03-20, p. 1.

  72. 72.

    El Comercio 1925-03-21, p. 6.

  73. 73.

    El Comercio 1925-02-18, p. 2; 1925-02-22, p. 2; 1925-03-02, p. 3; 1925-03-20, p. 1.

  74. 74.

    El Comercio 1925-03-09, p. 1.

  75. 75.

    El Comercio 1925-03-18, p. 3.

  76. 76.

    El Comercio 1925-03-27, p. 2; El Comercio 1925-03-24, p. 5.

  77. 77.

    El Comercio 1925-02-10 p. 6; 1925-02-12 p. 4; 1925-02-18 p. 4 ; 1925-02-27 p. 2.

  78. 78.

    El Comercio 1925-02-27 p. 2.

  79. 79.

    El Comercio 1926-02-27.

  80. 80.

    El Comercio 1925-02-03, p. 3; El Comercio 1925-02-04, p. 4; 1925-02-05, p. 8; 1925-02-09, p. 1; 1925-02-11, p. 2; 1925-02-17, p. 2; 1925-02-26, p. 9; 1925-02-27, p. 3.

  81. 81.

    El Comercio 1925-02-14, p. 1; 1925-02-25, p. 2; 1925-02-26, p. 8; 1925-02-27, p. 3.

  82. 82.

    1925-03-12, p. 4.

  83. 83.

    El Comercio 1925-06-09, p. 2.

  84. 84.

    El Comercio 1925-02-09, p. 1.

  85. 85.
  86. 86.

    El Comercio 1925-04-17, p. 8.

Notes

Acknowledgements

This study was supported by the Manglares-IGP Project (IDRC 106714) and the Peruvian PP068 program. We thank A. Chang and C. Rojas-Rosas for the “El Comercio” and “El Tiempo” newspaper clippings, respectively, as well as I. Montes and A. Levy for their help acquiring the Arcturus (Beebe 1926) and Chicama data, respectively. We thank K. Mosquera, R. Woodman, J. Cole, B. Dewitte, A. Watkins, R. Rodríguez, A. Timmermann, J. Apaéstegui, J. C. Espinoza, and S. Morera for useful discussions. Data analysis and plots were done with GNU Octave and GrADS.

Supplementary material

382_2017_3702_MOESM1_ESM.pdf (650 kb)
Supplementary material 1 (PDF 650 kb)

References

  1. Aceituno P, Prieto MR, Solari M, Martínez AG, Poveda G, Falvey M (2009) The 1877–1878 El Niño episode: associated impacts in South America. Clim Change. doi: 10.1007/s10584-008-9470-5
  2. Alory G, Maes C, Delcroix T, Reul N, Illig S (2012) Seasonal dynamics of sea surface salinity off Panama: the far Eastern Pacific fresh pool. J Geophys Res: Oceans 117(C4):C04028CrossRefGoogle Scholar
  3. Autoridad Nacional del Agua (ANA) (2010) Estudio Hidrológico y Ubicación de la Red de Estaciones Hidrométricas en la Cuenca del Río Rímac, 2, 134 ppGoogle Scholar
  4. Bailey SI (1930) Observations made at the Arequipa station 1896–1925. Ann Astron Obs Harvard Coll 87(2A):179–217Google Scholar
  5. Battisti DS, Sarachik ES, Hirst AC (1999) A consistent model for the large-scale steady surface atmospheric circulation in the tropics. J Clim 12(10):2956–2964CrossRefGoogle Scholar
  6. Beebe W (1926) The Arcturus oceanographic expedition. Zoologica 8:1–45Google Scholar
  7. Belmadani A, Echevin V, Codron F, Takahashi K, Junquas C (2014) What dynamics drive future wind scenarios for coastal upwelling off Peru and Chile? Clim Dyn 43(7):1893–1914. doi: 10.1007/s00382-013-2015-2 CrossRefGoogle Scholar
  8. Bendix J, Trachte K, Palacios E, Rollenbeck R, Göttlicher D, Nauss T, Bendix A (2011) El Niño meets La Niña—anomalous rainfall patterns in the “traditional” El Niño region of southern Ecuador. Erdkunde 65(2):151–167CrossRefGoogle Scholar
  9. Berry E (1927) Meteorological observations at Negritos, Peru, December 1924, to May, 1925. Mon Weather Rev 55:75–78CrossRefGoogle Scholar
  10. Bjerknes J (1969) Atmospheric teleconnections from the equatorial Pacific. Mon Weather Rev 97(3):163–172CrossRefGoogle Scholar
  11. CAF (Corporación Andina de Fomento) (2000) El Fenómeno El Niño 1997–1998. Memoria, Retos y Soluciones, vol 5, p 293Google Scholar
  12. Capotondi A, Wittenberg AT, Newman M, Di Lorenzo E, Yu JY, Braconnot P, Cole P, Dewitte B, Giese B, Guilyardi E, Jin FF, Karnauskas K, Kirtman B, Lee T, Schneider N, Xue Y, Yeh SW (2015) Understanding ENSO diversity. Bull Amer Met Soc 96(6):921–938. doi: 10.1175/BAMS-D-13-00117.1 CrossRefGoogle Scholar
  13. Cardone V, Greenwood JG, Cane MA (1990) On trends in historical marine wind data. J Clim 3:113–127CrossRefGoogle Scholar
  14. Carranza L (1891) Contra-corriente maritima observada en Paita y Pacasmayo. Bol Soc Geogr Lima 1(9):344–345Google Scholar
  15. Carrillo CN (1893) Hidrografía oceánica. Bol Soc Geogr Lima, pp 72–110Google Scholar
  16. Chang A (2014) La cobertura periodística del Fenómeno El Niño de 1925–1926 en el diario El Comercio de Lima. Masters Thesis in History, Pontificia Universidad Católica del PerúGoogle Scholar
  17. Chiang JCH, Sobel AH (2002) Tropical tropospheric temperature variations caused by ENSO and their influence on the remote tropical climate. J Clim 15:261–2630CrossRefGoogle Scholar
  18. Chiodi A, Harrison D, Vecchi GA (2014) Subseasonal atmospheric variability and El Niño waveguide warming: observed effects of the Madden-Julian oscillation and westerly wind events. J Clim 27(10):3619–3642CrossRefGoogle Scholar
  19. Clement AC, DiNezio PN, Deser C (2011) Rethinking the oceans role in the Southern Oscillation. J Clim. doi: 10.1175/2011JCLI3973.1
  20. Compo G, Whitaker J, Sardeshmukh P (2006) Feasibility of a 100-year reanalysis using only surface pressure data. Bull Am Metereol Soc 87(2):175–190. doi: 10.1175/BAMS-87-2-175 CrossRefGoogle Scholar
  21. Compo GP, Whitaker JS, Sardeshmukh PD, Matsui N, Allan RJ, Yin X, Gleason BE, Vose RS, Rutledge G, Bessemoulin P, Bronnimann S, Brunet M, Crouthamel RI, Grant AN, Groisman PY, Jones PD, Kruk MC, Kruger AC, Marshall GJ, Maugeri M, Mok HY, Nordli X, Ross TF, Trigo RM, Wang XL, Woodruff SD, Worley SJ (2011) The twentieth century reanalysis project. QJRMS 137:1–28. doi: 10.1002/qj.776 CrossRefGoogle Scholar
  22. Cucalon E (1987) Oceanographic variability off Ecuador associated with an El Niño event in 1982–1983. J Geophys Res 92(C13):14309–14322CrossRefGoogle Scholar
  23. Cushman GT (2004) Enclave vision: foreign networks in Peru and the internationalizaiton of El Niño research during the 1920s. Hist Meteorol 1:65–74Google Scholar
  24. Deser C, Wallace JM (1987) El Niño events and their relation to the Southern Oscillation: 1925–1986. J Geophys Res 92:14189–14196CrossRefGoogle Scholar
  25. Deser C, Phillips AS, Alexander MA (2010) Twentieth century tropical sea surface temperature trends revisited. Geophys Res Lett 37:L10701. doi: 10.1029/2010GL043321 CrossRefGoogle Scholar
  26. Dewitte B (2000) Sensitivity of an intermediate ocean–atmosphere coupled model of the tropical Pacific to its oceanic vertical structure. J Clim 13:2363–2388CrossRefGoogle Scholar
  27. Dewitte B, Takahashi K (2016) Diversity of moderate El Niño events evolution: role of air–sea interactions in the eastern tropical Pacific. Clim Dyn (in review)Google Scholar
  28. Douglas M, Mejia J, Ordinola N, Boustead J (2009) Synoptic variability of rainfall and cloudiness along the coasts of northern Peru and Ecuador during the 1997–8 El Niño event. Mon Weather Rev 137:116–136. doi: 10.1175/2008MWR2191.1 CrossRefGoogle Scholar
  29. Eguiguren V (1894) Las lluvias en Piura. Bol Soc Geogr Lima 4:241–258Google Scholar
  30. Enfield DB (1981) Thermally driven wind variability in the planetary boundary layer above Lima, Peru. J Geophys Res 86(C3):2005–2016CrossRefGoogle Scholar
  31. Falvey M, Garreaud RD (2005) Moisture variability over the South American Altiplano during the South American low level jet experiment (SALLJEX) observing season. J Geophys Res 110(D22):D22105. doi: 10.1029/2005JD006152 CrossRefGoogle Scholar
  32. Garcés-Vargas J, Schneider W, Abarca del Río R, Martínez R, Zambrano E (2005) Inter-annual variability in the thermal structure of an oceanic time series station off Ecuador (19902003) associated with El Niño events. Deep-Sea Res I. doi: 10.1016/j.dsr.2005.05.008
  33. Garreaud RD, Aceituno P (2001) Interannual rainfall variability over the South American altiplano. J Clim 14:2779–2789CrossRefGoogle Scholar
  34. Giese B, Ray S (2011) El Niño variability in simple ocean data assimilation (SODA), 1871–2008. J Geophys Res 116(C2):C02024. doi: 10.1029/2010JC006695 CrossRefGoogle Scholar
  35. Goldberg RA, Tisnado G, Scofield RA (1987) Characteristics of extreme rainfall events in northwestern Peru during the 1982–1983 El Niño period. J Geophys Res 92(C13):14225–14241CrossRefGoogle Scholar
  36. GRA La Libertad (2010) La Libertad: Clima y Ríos en Cifras. Estadísticas de Seis Décadas, p 67Google Scholar
  37. Graham NE, Barnett TP (1987) Sea surface temperature, surface wind divergence, and convection over tropical oceans. Science 238(4827):657–659CrossRefGoogle Scholar
  38. Harrison DE, Larkin NK (1998) El Niño-Southern Oscillation sea surface temperature and wind anomalies, 1946–1993. Rev Geophys 36(3):353–399CrossRefGoogle Scholar
  39. Holstein O (1927) Chan-Chan: capital of the Great Chimu. Geogr Rev 17(1):36–61CrossRefGoogle Scholar
  40. Horel JD (1986) Cornejo-Garrido AG (1986) Convection along the coast of northern Peru during 1983: spatial and temporal variation of clouds and rainfall. Mon Weather Rev 114:2091–2105CrossRefGoogle Scholar
  41. Huaman L, Takahashi K (2016) The vertical structure of the Eastern Pacific ITCZs and associated circulation using the TRMM precipitation Radar and in situ data. Geophys Res Lett 43:8230–8239. doi: 10.1002/2016GL068835 CrossRefGoogle Scholar
  42. Huertas L (2001) Diluvios andinos: a través de las fuentes documentales. Pontificia Universidad Católica del PerúGoogle Scholar
  43. Jáuregui YR, Takahashi K (2017) Simple physical-empirical model of the precipitation distribution in the tropical oceans and the effects of climate change. Clim Dyn (in review)Google Scholar
  44. Johnson N, Xie SP (2010) Changes in the sea surface temperature threshold for tropical convection. Nat Geosci 3:842. doi: 10.1038/ngeo1008 CrossRefGoogle Scholar
  45. Karnauskas K, Busalacchi A, Murtugudde R (2008) Low-frequency variability and remote forcing of gap winds over the east Pacific warm pool. J Clim 21(19):4901–4918. doi: 10.1175/2008JCLI1771.1 CrossRefGoogle Scholar
  46. Kousky VE, Kayano MT (1994) Principal modes of outgoing longwave radiation and 250-mb circulation for the South American sector. J Clim 7(7):1131–1143CrossRefGoogle Scholar
  47. Lavado-Casimiro W, Espinoza JC (2014) Impactos de El Niño y La Niña en las lluvias del Perú (1965–2007). Rev Bras Meteor 29(2):171–182CrossRefGoogle Scholar
  48. León KB (2014) Análisis espacio-temporal de las precipitationes y caudales durante los eventos El Niño (1982–83 y 1997–98) en la costa norte peruana. Thesis, Agric. Eng, Universidad Nacional Agraria La Molina, Peru, p 146Google Scholar
  49. L’Heureux ML, Takahashi K, Watkins AB, Barnston AG, Becker EJ, Di Liberto TE, Gamble F, Gottschalck J, Halpert MS, Huang B, Mosquera-Vásquez K, Wittenberg AT (2016) Observing and Predicting the 2015–16 El Niño. Bull Am Meteorol Soc. doi: 10.1175/BAMS-D-16-0009.1 Google Scholar
  50. Lindzen RS, Nigam S (1987) On the role of sea surface temperature gradients in forcing low-level winds and convergence in the tropics. J Atmos Sci 44(17):2418–2436CrossRefGoogle Scholar
  51. Morón O (2011) Climatología de la salinidad superficial del mar frente a la costa peruana. 1960–2008. Inf IMARPE 38(1):7–39Google Scholar
  52. Murphy RC (1926) Oceanic and climatic phenomena along the west coast of South America during 1925. Geogr Rev 16:26–54CrossRefGoogle Scholar
  53. Nials FL, Deeds EE, Moseley ME, Pozorski SG, Pozorksi TG, Feldman R (1979) El Niño: the catastrophic flooding of coastal Peru. Field Museum Nat Hist Bull 50(7):4–14, (8):4–10Google Scholar
  54. Neelin J, Battisti DS, Hirst AC, Jin FF, Wakata Y, Yamagata T, Zebiak S (1998) ENSO theory. J Geophys Res 103(C7):14261–14290CrossRefGoogle Scholar
  55. Oerder V, Colas F, Echevin V, Codron F, Tam J, Belmadani A (2015) Peru–Chile upwelling dynamics under climate change. J Geophys Res. doi: 10.1002/2014JC010299
  56. O’Connor H (1988) Investigación del Huayco de Chosica 1987, sus efectos y medidas de mitigación. Thesis in Civil Engineering, Universidad Nacional de Ingeniería, Lima, 99 ppGoogle Scholar
  57. Perigaud C, Melin F, Cassou C (2000) ENSO simulated by intermediate coupled models and evaluated with observations over 1970–98. Part I: Role of the off-equatorial variability. J Clim 13:16051634CrossRefGoogle Scholar
  58. Petersen G (1935) Estudios climatológicos del noroeste peruano. Bol Soc Geol Peru VII:1–141Google Scholar
  59. Philander SGH, Pacanowski RC (1981) The oceanic response to cross-equatorial winds (with application to coastal upwelling in low latitudes). Tellus 33:201–210CrossRefGoogle Scholar
  60. Quinn WH, Neal VT, Antunez de Mayolo SE (1987) El Niño occurrences over the past four and a half centuries. J Geophys Res 92(C13):14449–14461CrossRefGoogle Scholar
  61. Quinn WH (1992) A study of Southern Oscillation-related climatic activity for AD 622–1900 incorporating Nile River flood data. In: Diaz HF, Markgraf V (eds) El Niño historical and paleoclimatic aspects of the Southern Oscillation. Cambridge Univ. Press, Cambridge, pp 119–149Google Scholar
  62. Ramos Y (2014) Estimación del efecto del cambio climático en la precipitación en la costa norte del Peruú usando simulaciones de modelos climáticos globales. Thesis, Meteor. Eng, Universidad Nacional Agraria La Molina, Peru, p 168Google Scholar
  63. Rasmusson EM, Carpenter TH (1982) Variations in tropical sea surface temperature and surface wind fields associated with the Southern Oscillation/El Niño. Mon Weather Rev 110:354–384CrossRefGoogle Scholar
  64. Rayner NA, Parker DE, Horton EB, Folland CK, Alexander LV, Rowell DP, Kent EC, Kaplan A (2003) Global analyses of sea surface temperature, sea ice, and night marine air temperature since the late nineteenth century. J Geophys Res 108(D14):4407. doi: 10.1029/2002JD002670 CrossRefGoogle Scholar
  65. Reparaz G (2013) Los Ríos de la Zona Árida Peruana. Universidad de Piura, Peru/Institut Cartogràfic de Catalunya, Spain, 352 ppGoogle Scholar
  66. Rocca L (2000) Impactos de “El Niño” en el sector rural: Lambayeque (siglo XX). SEPIA VIII—Mesas Regionales, p 57Google Scholar
  67. Rodriguez R, Mabres A, Luckman B, Evans M, Masiokas M, Ektvedt TM (2005) “El Niño” events recorded in dry-forest species of the lowlands of northwest Peru. Dendrochronologia 22:181–186. doi: 10.1016/j.dendro.2005.05.002 CrossRefGoogle Scholar
  68. Rojas-Rosas C (2014) La fotografía como elemento informativo del diario “El Tiempo” de Piura: Evolución de uso en la cobertra del fenómeno “El Niño‘’ de 1925 y 1983. Thesis, Communications, Universidad de PiuraGoogle Scholar
  69. Ropelewski CF, Jones PD (1987) An extension of the Tahiti–Darwin southern oscillation index. Mon Weather Rev 115:2161–2165CrossRefGoogle Scholar
  70. Schaeffer MB, Bishop YMM, Howard GV (1958) Some aspects of upwelling in the Gulf of Panama. Inter Am Trop Tuna Comm Bull 3(2):79–132Google Scholar
  71. Schott G (1931) Der Peru-Strom und seine nördlichen Nachbargebiete in normaler und anormaler Ausbildung. Ann Hydrogr Mar Meteor 59, 161–169, 200–213, 240–257. Translated to Spanish in Bol Cia Admin Guano, IX, 3–4, 65–110 (1933)Google Scholar
  72. Sheppard G (1930) Notes on the climate and physiography of southwestern Ecuador. Geogr Rev 20(3):445–453CrossRefGoogle Scholar
  73. Sheppard G (1933) The rainy season of 1932 in southwestern Ecuador. Geogr Rev 23(2):210–216CrossRefGoogle Scholar
  74. Smith T, Reynolds R, Peterson T, Lawrimore J (2008) Improvements to NOAAs historical merged land–ocean surface temperature analysis (1880–2006). J Clim 21(10):2283–2296CrossRefGoogle Scholar
  75. Takahashi K (2004) The atmospheric circulation associated with extreme rainfall events in Piura, Peru, during the 1997–1998 and 2002 El Niño events. Ann Geophys 22:3917–3926CrossRefGoogle Scholar
  76. Takahashi K (2005) The annual cycle of heat content in the Peru Current region. J Clim 18:4937–4954CrossRefGoogle Scholar
  77. Takahashi K, Dewitte B (2016) Strong and moderate nonlinear El Niño regimes. Clim Dyn. doi: 10.1007/s00382-015-2665-3
  78. Takahashi K, Montecinos A, Goubanova K, Dewitte B (2011) ENSO regimes: reinterpreting the canonical and Modoki El Niño. Geophys Res Lett 38:L10704. doi: 10.1029/2011GL047364 CrossRefGoogle Scholar
  79. Terneus A, Gioda A (2006) In search of colonial El Niño events and a brief history of meteorology in Ecuador. Adv Geosci 6:181–187. doi: 10.5194/adgeo-6-181-2006 CrossRefGoogle Scholar
  80. Trenberth KE (1997) The definition of El Niño. Bull Am Meteorol Soc 78(12):2771–2777CrossRefGoogle Scholar
  81. Vecchi GA, Soden BJ (2007) Effect of remote sea surface temperature change on tropical cyclone potential intensity. Nature 450(7172):1066–1070CrossRefGoogle Scholar
  82. Vuille M, Bradley RS, Keimig F (2000) Interannual climate variability in the Central Andes and its relation to tropical Pacific and Atlantic forcing. J Geophys Res: Atmos 105(D10):12447–12460CrossRefGoogle Scholar
  83. Walker GT (1924) Correlation in seasonal variations of weather. IX. A further study of world weather. Mem India Meteorol Dep 24(9):275–333Google Scholar
  84. Wallace JM, Rasmusson EM, Mitchell TP, Kousky VE, Sarachik ES, von Storch H (1998) On the structure and evolution of ENSO-related climate variability in the tropical Pacific: lessons from TOGA. J Geophys Res 103(C7):14241–14259CrossRefGoogle Scholar
  85. Wang B, Wang Y (1999) Dynamics of the ITCZ–equatorial cold tongue complex and causes of the latitudinal climate asymmetry. J Clim 12:1830–1847CrossRefGoogle Scholar
  86. Woodman RF (1985) Recurrencia del fenómeno del Niño con intensidad comparable a la del Niño 1982–1983. El Fenómeno El Niño, CONCYTEC, Ciencia, Tecnología y Agresión Ambiental, pp 301–332Google Scholar
  87. Woodman RF (1999) Modelo estadístico de pronóstico de las precipitaciones en la costa norte del Peru. El Fenómeno El Niño. Investigación para una prognosis, 1er encuentro de Universidades del Pacíifico Sur: Memoria, pp 93–108Google Scholar
  88. Woodruff SD, Worley SJ, Lubker SJ, Ji Z, Freeman JE, Berry DI, Brohan P, Kent EC, Reynolds RW, Smith SR, Wilkinson C (2011) ICOADS Release 2.5: extensions and enhancements to the surface marine meteorological archive. Int J Climatol 31(7):951–967. doi: 10.1002/joc.2103 CrossRefGoogle Scholar
  89. Wooster WS (1980) Early observations and investigations of El Niño: the event of 1925. Oceanography: the past. Springer, New York, pp 629–641Google Scholar
  90. Worley SJ, Woodruff SD, Reynolds RW, Lubker SJ, Lott N (2005) ICOADS release 2.1 data and products. Int J Climatol 25:823–842. doi: 10.1002/joc.1166 CrossRefGoogle Scholar
  91. Wyrtki K (1975) El Niño—the dynamic response of the equatorial Pacific Ocean to atmospheric forcing. J Phys Oceanogr 5:572–584CrossRefGoogle Scholar
  92. Xie SP, Philander SGH (1994) A coupled ocean–atmosphere model of relevance to the ITCZ in the eastern Pacific. Tellus 46A:340–350CrossRefGoogle Scholar
  93. Yulaeva E, Wallace JM (1994) The signature of ENSO in global temperature and precipitation fields derived from the microwave sounding unit. J Clim 7:1719–1736CrossRefGoogle Scholar
  94. Zebiak SE, Cane MA (1987) A model El Niño–Southern Oscillation. Mon Weather Rev 115:2262–2278CrossRefGoogle Scholar
  95. Zegarra JM (1926) Las lluvias y avenidas extraordinarias de 1925 y su influencia sobre la agricultura del departamento de La LibertadGoogle Scholar
  96. Zhang X, Liu H, Zhang M (2015) Double ITCZ in coupled ocean-atmosphere models: from CMIP3 to CMIP5. Res Lett Geophys. doi: 10.1002/2015GL065973
  97. Zorell F (1929) La corriente del Niño en 1925. Bol Soc Geogr Lima XLVI:1–18Google Scholar

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Authors and Affiliations

  1. 1.Instituto Geofísico del PerúLimaPeru

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