2.1. El Niño and the Southern Oscillation in the Western Tropical Pacific
 ENSO is an interannual mode of climate variability associated with a disturbance of the Walker circulation over the tropical Pacific Ocean [Rasmusson and Carpenter, 1982]. Over the late 20th century, observations from the tropical Pacific have shown that at 2–7 year intervals, the easterly trade winds relax, and this allows the pool of warm western Pacific surface waters to migrate toward the central and eastern tropical Pacific. As these warm surface waters shift eastward during an El Niño, so does the zone of strong atmospheric convection. The sea surface temperature anomalies associated with El Niño disrupt atmospheric circulation within the Hadley cell and in doing so, influence weather systems globally. As wind stress decreases over the eastern tropical Pacific during an El Niño, the strength of upwelling also decreases and hence, the thermocline deepens across the central and eastern tropical Pacific. In the western Pacific, the subsurface response is just the opposite: the thermocline shoals during an El Niño. In contrast to El Niño, La Niña conditions are characterized by intensified trade winds, warmer surface conditions than usual, and a deepening of the thermocline in the western tropical Pacific. The atmosphere responds to these oceanic conditions with a seesaw pattern in sea level pressure anomalies between Tahiti and Darwin (Australia), which is referred to as the Southern Oscillation. Thus, the coupled ocean-atmosphere process has been termed the El Niño Southern Oscillation (ENSO).
 Two different flavors of “El Niño” [Trenberth and Stepaniak, 2001] have been identified, based on the location of the maximum positive SST anomalies. Although there are differences among studies that attempt to characterize these different modes of ENSO [i.e., Ashok et al., 2007; Kug et al., 2009; Kumar et al., 2006; Larkin and Harrison, 2005a, 2005b; Trenberth and Stepaniak, 2001], the cold tongue El Niño (thereafter, EP El Niño), similar to the canonical El Niño described by Rasmusson and Carpenter , exhibits a maximum sea surface warming in the eastern Pacific. In contrast, some events show positive SST anomalies mostly confined to the Niño 4 region. These El Niños have been referred to as “dateline El Niño” [Larkin and Harrison, 2005a], “central Pacific (CP) El Niño” [Kao and Yu, 2009], “warm pool El Niño” [Kug et al., 2009], or “El Niño Modoki” [Ashok and Yamagata, 2009; Ashok et al., 2007]. This central Pacific warming may be accompanied by a weak cold tongue warming [Kug et al., 2009] or cooling [Ashok and Yamagata, 2009; Ashok et al., 2007; Yeh et al., 2009]. Both events are characterized by cooler surface waters than normal in the western tropical Pacific. CP El Niño has occurred more frequently since 1976 in association with a weakening of the easterly trade winds and a flattening of the equatorial thermocline gradient [Ashok et al., 2007; Yeh et al., 2009]. Conversely, CP La Niña is characterized by large negative SST anomalies in the central equatorial Pacific flanked on both sides by positive SST anomalies [Ashok and Yamagata, 2009]. Furthermore, the EP type of ENSO tends to produce strong El Niño events but relatively weak La Niña events. A reverse tendency has been shown for CP ENSO events [Kao and Yu, 2009].
 The mechanisms leading to the CP ENSO are not well understood. Recently, Yu et al.  demonstrated that some ENSO events are actually of both types. For instance, the 1982/83 El Niño was a weak CP event followed by a strong EP event. On the other hand, Ashok et al.  suggested that CP El Niño is not part of the traditional El Niño evolution and that these two types of ENSO behavior are fundamentally different phenomena, especially after the 1970s. Indeed, the transition mechanism and dynamical subsurface structure of a CP El Niño is inconsistent with the traditional delayed oscillator/discharge concept [Kao and Yu, 2009; Kug et al., 2009] and so is the pattern, amplitude, and even sign of the extratropical atmospheric teleconnections associated with these two types of ENSO [i.e., Ashok et al., 2007; Larkin and Harrison, 2005a]. For instance, changes in the extent and location of tropical Pacific warming have been shown to alter the teleconnection with rainfall and temperature anomalies over Asia, North America and Australia [Ashok et al., 2007; Hendon et al., 2009; Larkin and Harrison, 2005b; Lim et al., 2009; Wang and Hendon, 2007; Weng et al., 2007, 2009], the Indian monsoon [Kumar et al., 2006], and the frequency of tropical cyclones in the North Atlantic [Kim et al., 2009]. These differences in teleconnection patterns between cold tongue and warm pool El Niño can introduce a large bias in paleo-ENSO reconstructions from extratropical locations, stressing the need for high-resolution records from all regions of the tropical Pacific. In order to capture the pre-historic occurrence of the two types of ENSO, these reconstructions should include records from the western tropical Pacific. This is because temperature anomalies in the eastern tropical Pacific are of opposite signs during CP and EP ENSO events.
2.2. The Oceanography of the Indonesian Seas
 The MD77 sediment core (Figure 1) was collected as part of the IMAGES program using the R.V. Marion Dufresne. The core location is at the northern entrance of the Makassar Strait (1.4°N, 119°E, and 968m depth) in the Sulawesi Sea. The Indonesian Sea, including the Sulawesi Sea is part of the Indo-Pacific Warm Pool, the largest reservoir of warm surface waters in the tropical Pacific. This is the source of moist static energy for the rising limb of the Hadley and Walker circulation cells [Qu et al., 2005]. Therefore, changes in SSTs in the region can directly impact global atmospheric circulation. Surface ocean variability within the Makassar Strait is strongly influenced by the seasonal monsoons and the position of the Inter-Tropical Convergence Zone (ITCZ), and to a lesser extent by ENSO [Gordon, 2005]. The Makassar Strait is part of the so-called Indonesian Seaway, which is the primary conduit for exchange of Pacific waters with the Indian Ocean and therefore, a critical part of the thermohaline circulation system [Bray et al., 1996; Gordon, 1986]. The transport of upper ocean waters from the North Pacific through the Makassar Strait to the Indian Ocean, referred to as the Indonesian Throughflow (ITF), occurs primarily within the thermocline (100–200m, [Sprintall, 2009; Susanto and Gordon, 2005]). Approximately 80% of the ITF flows through the Makassar Strait [Gordon, 2005]. At the MD77 core location, the bulk of the ITF is composed of North Pacific subtropical water that flows southward from the Mindanao Current, east of the Philippines [Sprintall, 2009, Figure 1]. Thermocline temperature and salinity changes are primarily controlled by the seasonal reversal of winds that accompany the East Asian and Australian monsoons and also ENSO on an interannual timescale [Gordon, 2005].
 Seasonal changes in ITF properties and the net transport of water from the Pacific to the Indian Ocean are largely driven by the seasonal changes in monsoonal winds. The ITF transport through the thermocline intensifies during both monsoon seasons [Gordon, 2005; Gordon et al., 2008]. During the Australian (northwestern) monsoon, wind-forcing results in a transport of the lower-salinity and more buoyant Java Seawater into the Makassar Strait, creating a surface “plug,” which is accompanied by an increased transport within the thermocline [Gordon, 2005; Gordon et al., 2003; Susanto and Gordon, 2005]. The southeastern monsoon winds constrain southward surface water flow out of the Makassar Strait, also enhancing thermocline transport from July to September [Gordon, 2005; Gordon et al., 2003]. Transport of waters within the thermocline is highest and shallowest (lowest and deepest) during the southeastern (northwestern) monsoon [Gordon et al., 2008].
 Since the ITF is primarily driven by the difference in sea level height between the western tropical Pacific and the Indian Ocean, the transport and water mass properties of the ITF also vary with ENSO. The weakening of the trade winds and the associated displacement of the western Pacific warm pool that occurs during an El Niño reduces the sea level gradient between the two ocean basins, reducing the ITF transport [Bray et al., 1996; England and Huang, 2005; Fieux et al., 1996; Gordon, 2005; Gordon and Fine, 1996; Gordon and Susanto, 1999; Gordon et al., 1999; Meyers, 1996; Sprintall, 2009]. A 15-year database of XBT (Expendable Bathythermograph) data has shown that thermocline temperatures in the Makassar Strait are highly correlated (r = 0.77) with the SOI [Ffield et al., 2000]. Unfortunately, there is no long, continuous in situ temperature and salinity data at the MD77 site. We therefore rely on reanalysis data to evaluate the impact of ENSO and the seasonal cycle on temperature and salinity at the surface and within the thermocline at the core location. We use the European Center for Medium-Range Weather Forecasts (ECMWF) ocean reanalysis system 3 covering the period 1959–2009 (thereafter, ORA-S3, [Balmaseda et al., 2008]). Our motivation for using the ORA-S3 reanalysis compared to other reanalysis data sets stems from the fact that the artificial vertical velocity fields produced by most ocean models within a few degrees of the equator [Bell et al., 2004] are partially corrected for through the use of an online bias-correction scheme in the pressure field [Balmaseda et al., 2007, 2008]. In the absence of this bias-correction algorithm, the artificial velocity field can produce inaccuracies in the temperature estimates [Bell et al., 2004], making an evaluation of ENSO dynamics at the MD77 location impossible. We acknowledge that the reanalysis data may not be completely accurate (a fact further discussed in section 3) but it provides a basis for discussing the impact of ENSO at the MD77 site.
 Temperature and salinity anomalies were calculated for the ORA-S3 data set by removing the annual mean temperature (salinity) of the entire reanalysis period at each depth. Any long-term trends in the data were not removed beforehand in order to evaluate the relative strength of the interannual signal and background climate variability. Temperature anomalies associated with ENSO dynamics at the MD77 location are greater within the thermocline than in the mixed layer. The maximum anomalies occur at ∼100m depth and average ∼ ± 2°C. In the upper thermocline, the temperature anomalies associated with ENSO are on the order of ∼±1°C (Figure 1b). On the other hand, salinity changes are largely independent of ENSO (Figure 1c). At the surface, the lack of an ENSO-salinity relationship is explained by the long-term decrease (∼1psu) in salinity from 1959 to 1982 apparent in the ORA-S3 reanalysis data set, which masks the interannual variability. Changes in salinity in the sub-surface are small (<0.1psu) and are not systematically associated with ENSO, although there is a general tendency toward slightly lower (higher) salinities during La Niña (El Niño). However, the interannual anomalies in salinity are of the same magnitude as those associated with the seasonal cycle (Figure 1e). On the other hand, the seasonal changes in temperature, which can conceal the surface expression associated with ENSO, are reduced in the thermocline, which enhances the expression of ENSO (Figure 1d). At the core location the thermocline temperature (and, to a lesser extent, salinity) anomalies that accompany CP and EP types of ENSO are of the same magnitude and most importantly, of the same sign. This means that the MD77 record is particularly well-suited for documenting the evolution of both types of ENSO events during the past millennium.