Constraints on glaciation in the middle Eocene (46–37 Ma) from Ocean Drilling Program (ODP) Site 1209 in the tropical Pacific Ocean


  • Caroline F. Dawber,

    1. Department of Earth Sciences, University of Cambridge, Cambridge, UK
    Search for more papers by this author
  • Aradhna K. Tripati

    1. Department of Earth Sciences, University of Cambridge, Cambridge, UK
    2. Department of Earth and Space Sciences, Department of Atmospheric and Oceanic Sciences, and Institute of Geophysics and Planetary Physics, University of California, Los Angeles, California, USA
    Search for more papers by this author


[1] The presence of glacial ice in the late middle Eocene has been vigorously debated. Recently published sedimentary data from the high latitudes is suggestive of episodic cooling events and near-freezing sea surface temperatures during parts of the late middle Eocene. Constraints on ice volumes and the significance of excursions in open ocean foraminifera and seawater δ18O reconstructions are less clear, and there are few high-resolution δ18O records. We present a new detailed record of benthic foraminiferal δ18O from Site 1209 that exhibits variations (Δδ18Obenthic) of 0.6‰–1.3‰. Different approaches have previously been used to interpret Δδ18Obenthic, including (1) an a priori assumption of a 50% contribution of temperature, similar to what is reconstructed for the Last Glacial Maximum–recent change; (2) applying Oligocene calibrations between apparent sea level (ASL) and Δδ18Obenthic; or (3) assuming temperature and seawater δ18O contributions can be partitioned through comparison with benthic Mg/Ca. Using assumption 1, the record from Site 1209 indicates changes in seawater δ18O of 0.3‰–0.7‰, equivalent to ∼33–72 m (m) of ASL (assuming mean ice δ18O of ∼−45‰). Using assumption 2 and two different end-member calibrations, the δ18Obenthic record implies changes in ASL of 23–50 m or 50–108 m. The third approach yields changes in seawater δ18O of up to 0.6‰ to 1.4‰. We explore the compatibility of the results of each of these approaches with other studies that discuss evidence for ephemeral glaciations during the middle Eocene with variable ice storage at one or both poles.

1. Introduction

[2] The history of glaciation in both hemispheres has been a topic of debate for decades [e.g., Kennett, 1977; Shackleton and Opdyke, 1977; Bartek et al., 1992; Raymo, 1994; Tripati et al., 2001, 2005, 2008; Zachos et al., 2001; Miller et al., 2005; Coxall et al., 2005; Shevenell and Kennett, 2007; Edgar et al., 2007; Spielhagen and Tripati, 2009]. The nature of climate variability during the Eocene (∼55–34 Ma) has been particularly controversial, as most proxy records are low resolution and/or punctuated. The uncertainty over the history of ice volume during the middle Eocene in particular has resulted in the interval being termed the “doubthouse” [Browning et al., 1996].

[3] Our understanding of Eocene climate and of the middle Eocene doubthouse has been significantly advanced by the recovery of new sedimentary sequences, as well as the publication of several records of benthic foraminiferal δ18O [Wade and Kroon, 2002; Tripati et al., 2005; Sexton et al., 2006a; Edgar et al., 2007; Burgess et al., 2008; Zachos et al., 2008; Bohaty et al., 2009], sequence boundaries [Miller et al., 2005; Pekar et al., 2005; Röhl et al., 2004; Gale et al., 2006; Peters et al., 2010; Dawber et al., 2011], temperature reconstructions [Tripati et al., 2003, 2005; Billups and Schrag, 2003; Dutton et al., 2005; Hollis et al., 2009; Liu et al., 2009; Spielhagen and Tripati, 2009] and ice-rafted debris [Moran et al., 2006; Eldrett et al., 2007; St. John, 2008; Tripati et al., 2008; Stickley et al., 2009]. High pCO2 “greenhouse” conditions are thought to have characterized the early Eocene [Pearson and Palmer, 2000; Pagani et al., 2005; Tripati et al., 2005], and polar climates during the early Eocene are generally thought to have been warm [Dingle et al., 1998; Francis, 1999; Francis and Poole, 2002]. However, a new record of seawater δ18O from the midlatitude southwest Pacific (derived from paired carbonate δ18O and Mg/Ca measurements on benthic foraminifera) exhibits large changes (of ∼0.5‰–1‰) between 50 and 48 Ma, which could support the presence of transient ephemeral ice sheets on Antarctica during the early Eocene [Creech et al., 2010].

[4] Proxy records for the middle and late Eocene have been used to argue that mean pCO2 values were high, with support for large-amplitude variations in pCO2 and seawater carbonate chemistry during this time [Pearson and Palmer, 2000; Pagani et al., 2005; Tripati et al., 2005]. On the Antarctic Peninsula, there is evidence from records of vegetation and the oxygen isotope composition of marine cements and mollusks that cooling in this region had commenced by the early middle Eocene [Pirrie et al., 1998; Dingle and Lavelle, 1998; Francis, 1999; Francis and Poole, 2002; Ivany et al., 2008]. There are also indications that there may have been a change to cooler and more arid weathering regimes on the Antarctic continent during the middle Eocene based on clay mineral assemblages in Southern Ocean cores [Ehrmann and Mackensen, 1992]. Data from the Arctic implies the temperature history of the region was spatially variable during the Eocene. Floral evidence suggests more temperate conditions in Arctic Canada [e.g., Jahren, 2007]. In a sequence from Svalbard, glendonites (pseudomorphs of ikaite, a mineral which requires near-freezing water temperatures to form) are found in early Cenozoic sediments [Spielhagen and Tripati, 2009]. Recent studies have also reported sea ice diatoms and ice-rafted debris (IRD) in the central Arctic Ocean, in samples as old as 47 Ma [Moran et al., 2006; St. John, 2008; Stickley et al., 2009]. Additionally IRD has been reported in middle and late Eocene sediments from the Greenland Sea back to 44 Ma [Tripati et al., 2008]. However, it is debated whether the lithic fragments found in Eocene sediments were in fact ice rafted, and if so, whether ice-rafted debris was associated with glacial ice or seasonal sea ice.

[5] Based on records of carbonate δ18O, it has traditionally been argued that the middle and late Eocene were characterized by little or no ice storage, with the abrupt 1‰ increase in benthic and seawater δ18O values close to the Eocene-Oligocene boundary (Oi-1) thought to reflect first appearance of substantial continental ice sheets on Antarctica [Lear et al., 2000; Zachos et al., 2001; Coxall et al., 2005; Lear et al., 2008; Liu et al., 2009]. In contrast, the stratigraphic architecture of Eocene sequences from the New Jersey Coastal Plain (NJCP) has been used to argue for high-frequency sea level variations on the order of tens of meters (10–40 m) driven by small ice sheets on Antarctica as early as ∼49 Ma [Browning et al., 1996; Miller et al., 2005] and also during some intervals of the Cretaceous [e.g., Miller et al., 2005]. Other studies of sequence boundaries have also called for significant variations in sea level during the Eocene [Pekar et al., 2005; Röhl et al., 2004; Gale et al., 2006; Peters et al., 2010; Dawber et al., 2011].

[6] However, the significance of ∼0.6‰–1.3‰ shifts in δ18Obenthic reported in middle and upper Eocene sediments from several basins [Bohaty and Zachos, 2003; Tripati et al., 2005; Sexton et al., 2006a; Edgar et al., 2007] has not been resolved. Based on the comparison of a Pacific deep-water δ18Obenthic record with low-latitude surface δ18Ocarbonate records, Tripati et al. [2005] argue that the majority of the Δδ18Obenthic reflects seawater δ18O, supported by Mg/Ca-based estimates of seawater δ18O. In contrast, some have assumed that Eocene δ18Obenthic reflects approximately equal contributions from temperature and seawater δ18O and have questioned whether the geochemistry of benthic foraminifera reported by Tripati et al. [2005] reflects contamination, a site-specific artifact, or “noise” [Edgar et al., 2007]. Thus there is still considerable debate over (1) whether middle Eocene climate was relatively stable, characterized by gradual cooling, or more variable/dynamic; (2) whether the late middle Eocene was ice-free or not (i.e., whether the greenhouse-icehouse transition began then or began at the Eocene-Oligocene transition); and (3) if there was ice present ephemerally, whether there was some ice stored at just one or at both poles.

[7] Below we present high-resolution (average resolution of 1 sample/5 cm, generally equivalent to ∼15 ka) records of benthic foraminiferal δ18O and Mg/Ca for the middle Eocene from Ocean Drilling Program (ODP) Site 1209 in the Pacific Ocean. We show that the benthic foraminiferal δ18O record at Site 1209 exhibits variations of 0.5‰–1.3‰. Several different approaches are used to interpret these records and constrain temperature and seawater δ18O. To interpret changes in δ18Obenthic, we (1) use an a priori assumption of a 50% contribution from temperature (similar to what is reconstructed for the Last Glacial Maximum–recent change, and also the approach adopted by Edgar et al. [2007]), (2) apply Oligocene calibrations between apparent sea level (ASL) and Δδ18Obenthic (equivalent to ∼24%–65% temperature contributions to Δδ18Obenthic [Pekar et al., 2006]), or (3) assume temperature and seawater δ18O effects can be deconvoluted through comparison with benthic foraminiferal Mg/Ca. To interpret changes in benthic foraminiferal Mg/Ca, we consider the effects of changing (1) temperature, (2) seawater Mg/Ca, and (3) carbonate saturation. Irrespective of which of the three approaches are used to interpret benthic δ18O, we find there is evidence for changes in seawater δ18O of 0.3‰ to 0.6‰ or more. We discuss the implications of these records and explore how compatible these reconstructions are with sedimentary and geochemical proxy records from several published studies [e.g., Bohaty and Zachos, 2003; Tripati et al., 2005, 2008; Miller et al., 2005; Edgar et al., 2007]. We conclude there is evidence for major glaciations and deglaciations at ∼44–43 Ma, ∼42–40 Ma, and ∼39–38 Ma. We also suggest that excursions in the δ18O of benthic foraminifera and seawater at ∼41 and ∼38.6 Ma are apparently too large to be accommodated by changes in ice storage on Antarctica. Reported variations in δ18O could potentially be explained by the occurrence of short-lived glacial episodes during the Eocene that were associated with some ice storage in the Northern Hemisphere, and would be consistent with previous reports of IRD in Arctic sediments [St. John, 2008; Tripati et al., 2008].

2. Paleoceanographic Setting and Stratigraphy

[8] ODP Site 1209 (32°39.108′N, 158°30.3564′E), situated on the southern high of Shatsky Rise in the Pacific Ocean, was located in the northern subtropics throughout the middle and late Eocene (refer to auxiliary material). The Pacific Ocean basin was the largest oceanic basin during the Eocene, and consequently Pacific reconstructions of seawater δ18O may be a better approximation for mean seawater δ18O (and ice volume) than records from the continental shelf or slope environments, or from open ocean settings in smaller basins such as the Atlantic Ocean or Indian Ocean.

[9] Site 1209 was located above the carbonate compensation depth (CCD) throughout the greenhouse-icehouse transition (refer to auxiliary material [Rea and Lyle, 2005; Tripati et al., 2005]), with a paleodepth of ∼1.9–2.5 km during the Eocene [Dutton et al., 2005; Bohaty et al., 2009]. The shallower depth of Site 1209 may reduce potential artifacts related to the impact of changing carbonate ion concentration and saturation on foraminiferal δ18O and Mg/Ca [Spero et al., 1997; Elderfield et al., 2006], although proxy records for carbonate dissolution support a relatively shallow Pacific lysocline during middle Eocene [Hancock and Dickens, 2005].

[10] Chronostratigraphic datums for Site 1209 are based on biostratigraphic markers and indicate that the studied interval (140–164 revised meters composite depth (rmcd)) is middle Eocene in age (refer to auxiliary material [Bralower, 2005; Petrizzo et al., 2005]). Owing to high planktonic fragmentation in parts of the studied interval, many of the widely used foraminiferal marker species are absent and/or their distribution could not be precisely defined [Petrizzo et al., 2005]. In contrast, many of the calcareous nannofossil marker species are present, although their preservation has been described as moderate [Bralower, 2005]. We report data as a function of depth (Figure 1) and age on the Berggren and Pearson [2005] timescale (Figures 2, 3, 4 and 5), using published calcareous nannofossil datums (summarized in the auxiliary material [Bralower, 2005]), and note that the age model for Site 1209 may be open to reinterpretation if the calibrations of calcareous nannofossil datums at the low latitudes are revised. The use of planktonic foraminifera datums [Petrizzo et al., 2005] results in the age model being adjusted by several million years, although the temporal range of several of the secondary zonal markers have since been recalibrated [Pearson et al., 2006].

Figure 1.

ODP Site 1209 data reported versus revised meters composite depth. Shown (from bottom to top) is a photographic splice (dashed lines) of core material from 1209A and 1209B used in this study; ship board color reflectance (L*) data [Bralower et al., 2003]; records of foraminiferal dissolution indices based on the percentage of benthic foraminifera (pink) and planktonic fragmentation (gray). Dashed lines denote data from Hancock and Dickens [2005] (based on material from the >38 μm fraction), and solid lines are data from this study (using material from the >63 μm fraction). The benthic foraminiferal δ18O record is based on measurements of Oridorsalis umbonatus (blue circles). The taxa Nuttallides truempyi (blue squares) and Cibicidoides spp. (blue triangles) were substituted for ∼20% of the δ18O record, where O. umbonatus was not present in sufficient quantities. The δ18O values have been normalized to Cibicidoides spp. using the species corrections of Katz et al. [2003] and corrected to equilibrium [Graham et al., 1981]. Also shown is the raw Mg/Ca data (black) and the bottom water temperature reconstruction (green) that uses the respective species calibrations of Lear et al. [2002] and a constant value of 2.25 mol/mol for Eocene seawater Mg/Ca. For reference, the stratigraphic range at Site 1209 of some zonal calcareous nannofossil marker species are shown [Bralower, 2005].

Figure 2.

Comparison of the middle Eocene benthic foraminiferal δ18O record from Site 1209 (blue) to published records from deep Pacific Site 1218 (black [Tripati et al., 2005; Lear et al., 2004]) and deep Atlantic Site 1260 (purple [Edgar et al., 2007]). The data for sites 1209 and 1260 are reported on the Berggren and Pearson [2005] GPTS, and data for 1218 are on the Pälike et al. [2006] timescale. We note that there may be small (few tens to hundreds of ka) between these age scales. Shown also is the long-term trend in “global” benthic δ18O that is based on a compilation of data from multiple sites and using multiple species [Zachos et al., 2008].

Figure 3.

Middle Eocene benthic foraminiferal Mg/Ca record for Site 1209 is compared to published data from the Southern, Indian, and Pacific oceans [Lear et al., 2000; Billups and Schrag, 2003; Dutton et al., 2005]. All data is reported on the Berggren and Pearson [2005] GPTS.

Figure 4.

Illustration of the sensitivity of the long-term trends and absolute paleotemperature estimates based on foraminiferal Mg/Ca. (a) A range of Eocene seawater Mg/Ca histories that may be consistent with existing independent constraints (diamonds denote constraints from midocean flank carbonate veins [Coggon et al., 2010], circles are constraints from a cation flux model [Wilkinson and Algeo, 1989], squares denote constraints from fluid inclusions in halite crystals, closed squared are from Horita et al. [2002], open squares are from Lowenstein et al. [2001]). (b) The bottom water temperature reconstructions for ODP Site 1209 based on the (color-coded) seawater Mg/Ca histories shown in Figure 4a. (c) The best guess bottom water temperature reconstruction for Site 1209 based on the green seawater Mg/Ca model that is compatible with the cation flux constraints.

Figure 5.

Comparison of the Site 1209 benthic foraminiferal δ18O and Mg/Ca-based bottom water temperature reconstruction (using the green seawater Mg/Ca model in Figure 4), with carbonate accumulation data and a reconstruction of the carbonate compensation depth from the deep equatorial Pacific [Tripati et al., 2005]. Carbonate accumulation events (CAE) are from Tripati et al. [2005]. Shown also are proxy estimates of atmospheric pCO2 and their estimated uncertainties (diamonds, alkenones [Pagani et al., 2005; Freeman and Hayes, 1992]; circles, boron isotopes [Demicco et al., 2003]; inverted triangle, stomatal indices [McElwain, 1998; Kürschner et al., 2001]; open circle with dot, paleosol carbon isotopes [Ekart et al., 1999]). The seawater δ18O reconstruction from Site 1209 (blue) based on combined estimates of benthic foraminiferal δ18O and Mg/Ca bottom water temperatures. The amplitude of seawater δ18O excursions discussed in the text and our interpretations of ice budgets are based on the smoothed seawater δ18O reconstruction.

3. Methods and Materials

3.1. Analytical Procedure

[11] Benthic foraminiferal oxygen isotope ratios were determined on two gas source mass spectrometers (a Micromass Prism or Thermo Scientific 253) in the Department of Earth Sciences, University of Cambridge. Long-term analytical precision based on replicate analyses of an in-house standard is 0.08‰. Calcite δ18O ratios were determined primarily on samples of Oridorsalis umbonatus. Nuttallides truempyi and Cibicidoides spp. were substituted where O. umbonatus was absent, and δ18O values have been adjusted using published species offsets [Katz et al., 2003].

[12] Mg/Ca ratios were measured on samples of O. umbonatus and Cibicidoides spp. picked from the >150 μm size fraction. Samples were cleaned to remove clays, coccoliths, and organic matter using the standard Cambridge cleaning protocol that is based on the work of Barker et al. [2003], and measured on a Varian Vista inductively coupled atomic emission spectrophotometer located in the Department of Earth Sciences. Trace metal ratios were calculated from intensity ratios following the method outlined by de Villiers et al. [2002]. A suite of samples were cleaned with an additional reductive cleaning step, and trace metal ratios determined on an inductively coupled plasma mass spectrometer at Cambridge University. Mg/Ca values were adjusted by 0.1 mmol/mol to correct for the lowering of Mg/Ca ratios during this corrosive step. Samples were screened for the presence of potential Mg-contaminating ferromanganese overgrowths, clay minerals and silicate particles based on the abundance of Al, Fe, Mn, and Si [Barker et al., 2003]. Analytical precision based on measurements of splits of foraminiferal samples is 3%.

3.2. Middle Eocene Temperature and Seawater δ18O Reconstruction

[13] Bottom water temperatures are calculated using published Mg/Ca temperature calibrations for O. umbonatus and Cibicidoides spp. [Lear et al., 2002], combined with model-derived estimates of Cenozoic seawater Mg/Ca ratios [Wilkinson and Algeo, 1989] (seawater Mg/Ca ratios range from ∼3.57 mol/mol at 45 Ma to ∼3.84 mol/mol at 35 Ma, see discussion The Wilkinson and Algeo [1989] seawater Mg/Ca reconstruction has been used by several other studies [Lear et al., 2000; Tripati et al., 2003, 2005; Billups and Schrag, 2003; Okafor et al., 2009], allowing us to directly compare our temperature estimates to those in other publications.

[14] Seawater δ18O is estimated from combining benthic carbonate δ18O and Mg/Ca-based paleotemperatures using a standard δ18O-paleotemperature equation [Shackleton, 1974]. We calculate the uncertainty of seawater δ18O estimates to be ∼0.4‰, based on the propagated error of individual uncertainties of Mg/Ca analyses (external precision of ±3%) and δ18O analyses (external precision of ±0.08‰), the reported uncertainty in δ18O species corrections (±0.2‰ [Katz et al., 2003]), and the uncertainty in Mg/Ca-temperature calibrations (equivalent to ±0.3‰ [see Rohling, 2007, and references therein]). We also include an estimate for the error in temperature that might result from a large change in bottom water carbonate saturation (∼20 μmol/kg [Tripati and Elderfield, 2005]).

4. Results and Discussion

4.1. The Site 1209 Benthic Foraminiferal δ18O Record

[15] The middle Eocene δ18Obenthic record from Site 1209 (Figures 3 and 4) exhibits a long-term increase of ∼1.3‰, consistent with the global composite [Zachos et al., 2001]. High-frequency variations of greater than 0.6‰ are also observed, and occur at 158.5 rmcd (∼44 Ma), 155 rmcd (∼42 Ma), 153.3 rmcd (∼41.2 Ma), 152.7 rmcd (∼41 Ma), 147 rmcd (∼38.5 Ma) and 144 rmcd (∼38 Ma). These long- and short-term variations are interpreted to predominantly reflect changes in temperature and/or ice volume. Given the potential uncertainties associated with the age model, it is not clear whether these excursions correspond to any of the orbital frequencies that have been inferred to pace δ18O cycles during the Oligocene [Pälike et al., 2006].

4.2. Assessing Ice Budgets From the Site 1209 δ18Obenthic Record

[16] To assess the significance of the Site 1209 δ18Obenthic record and evaluate middle Eocene ice budgets, we consider several approaches (Table 1).

Table 1. Estimates of Apparent Sea Level (ASL) Variations as Constrained by Changes in Benthic Foraminiferal δ18Oa
Δδ18ObenthicScenario 1Scenario 2
Δδ18OseawaterASL (m) 0.12‰ δ18Osw per 10 m1ASL (m) End-Member 1 (0.12‰ per 10 m)2ASL (m) End-Member 2 (0.26‰ per 10 m)2
  • a

    Scenario 1 assumes a 50% temperature contribution to Δδ18Obenthic and estimates ASL using a calibration between seawater δ18O-ASL based on earliest Oligocene reconstructions [Katz et al., 2008]. Scenario 2 ASL estimates use end-member δ18Obenthic-ASL calibrations based on compilations of Oligocene and Miocene records [Pekar et al., 2006]. ASL end-member calibrations are equivalent to ∼24% and 65% temperature contribution to Δδ18Obenthic, respectively (based on a subtracted seawater δ18O estimate calculated from 0.091‰ per 10 m, as described by Pekar et al. [2006]).


4.2.1. Approach 1: Assume Equal Partitioning of δ18Obenthic Into Temperature and δ18Oseawater Components

[17] Assuming that δ18Obenthic records comprise a 50% contribution from temperature (Table 1, scenario 1), the Site 1209 record would support variations of 0.30‰–0.65‰ in δ18Oseawater, equivalent to 36–78 m ΔASL. The lower range (0.30‰) of these estimates are similar to those reported by Edgar et al. [2007] based on a δ18Obenthic record from the deep equatorial Atlantic (ODP Site 1260) and the same assumption of a 50% temperature contribution (Figure 2). Edgar et al. [2007] argue that Δδ18Oseawater of this magnitude indicates variations in continental ice budgets that could be accommodated by ice storage solely on Antarctica. Although this calculation does not preclude the occurrence of ephemeral glaciations in the Northern Hemisphere as previously proposed by Tripati et al. [2005].

[18] The higher range of the calculated Δδ18Oseawater at Site 1209 greatly exceeds that reported for Site 1260 (Figure 2 [Edgar et al., 2007]) and would support significantly larger ice budgets. Variations in δ18Obenthic ranging from 0.6‰ to 1.2‰ are also observed in middle Eocene records from ODP Site 1218 (deep Pacific) [Tripati et al., 2005] and at multiple sites in the Southern Ocean [Bohaty and Zachos, 2003]. Using this same framework to interpret δ18Obenthic, this amplitude of variability would also indicate changes in δ18Oseawater of 0.3‰–0.6‰ during the late middle Eocene, and also indicate substantially greater ice budgets than proposed by Edgar et al. [2007].

4.2.2. Approach 2: Applying a δ18Obenthic–Sea Level Calibration

[19] A second approach to estimating ice budgets from δ18Obenthic is to use a published calibration based on contemporaneous measurements of apparent sea level and δ18Obenthic [e.g., Pekar et al., 2006] (scenario 2; Table 1). We use calibrations derived from Oligocene and Miocene records [Pekar et al., 2006], as these likely provide more appropriate climatic analogies than Pleistocene records. However, we note that the propagated error for the calibration is likely to be large, as it will reflect uncertainties associated with the δ18Obenthic records, the backstripped eustatic estimates and conversion of eustasy to apparent sea level. Based on end-member calibrations, the δ18Obenthic record for Site 1209 supports variations of 23–108 m in apparent sea level (Table 1). These wide-ranging estimates of changes in apparent sea level support the occurrence of glaciation during the late middle Eocene, but could be accommodated by different ice budget scenarios.

4.2.3. Approach 3: Using Coupled δ18Obenthic and Benthic Foraminiferal Mg/Ca-Paleotemperature Reconstructions to Independently Constrain δ18Oseawater

[20] Benthic foraminiferal Mg/Ca ratios and Mg/Ca derived estimates of bottom water temperatures from Site 1209 (Figures 3 and 5) are similar to values previously reported in a low-resolution study [Dutton et al., 2005]. Mean temperatures are estimated to be ∼9°C during the middle Eocene (based on a constant seawater Mg/Ca ratio of 2.25 mol/mol, see discussion below), with variations of up to 2°C. The Mg/Ca-based temperature reconstruction for Site 1209 likely reflects variations in Southern Ocean sea surface temperatures, as the neodymium isotopic composition of fish teeth at this site has been cited as evidence of a Southern Ocean source for intermediate bottom waters in the Pacific Ocean during the middle Eocene [Thomas, 2004]. These estimates are consistent with Mg/Ca-based temperatures reconstructed for several high-latitude sites, including in the Weddell Sea (ODP Site 689 [Lear et al., 2000; Billups and Schrag, 2003]) and Indian Ocean (ODP Site 757 [Billups and Schrag, 2003]), and to Arctic Ocean sea surface temperatures reconstructed for an overlapping interval from alkenones (44.5 Ma: ∼8°C–11°C) [Weller and Stein, 2008].

[21] At Site 1209, the long-term and short-term trends in benthic foraminiferal Mg/Ca are notably different to the benthic δ18O record (Figure 1). Taken at face value, this observation suggests that there may have been significant increases in δ18Obenthic that were accompanied by either no temperature change or an increase in temperature at Site 1209. As a result, the majority of the δ18Obenthic record for Site 1209 could be attributed to changes in δ18Oseawater (ice volume). Previous studies have also documented little or no temperature decrease at low-latitude open ocean sites associated with the large increase in δ18Obenthic observed across the Eocene-Oligocene boundary (Mg/Ca reconstruction of Lear et al. [2000]; TEX86 and U37k′ reconstructions of Liu et al. [2009]) and during the middle Miocene (Mg/Ca reconstruction of Holbourn et al. [2005]). In contrast, organic proxy-based temperature reconstructions for the high latitudes support significant cooling across the Eocene-Oligocene boundary, which may indicate that the temperature change at this time was very heterogeneous, or may reflect inaccuracies in one or both temperature reconstructions [e.g., Sachs et al., 2000; Elderfield et al., 2006; Turich et al., 2007; Lipp et al., 2008].

[22] The inferred relationship between ice growth and Pacific Ocean warming at ∼2 km water depth (and Southern Ocean warming) is enigmatic, and contrasts with Quaternary records, which show ice volume, and high-latitude sea surface temperatures are positively correlated [e.g., Martin et al., 2002; Elderfield et al., 2010]. For the purpose of this discussion, we adopt a “straw man” approach to evaluating the fidelity of the Site 1209 Mg/Ca data set as a record of water temperature. Testing End-Member Interpretations for the Benthic Mg/Ca Record at Site 1209

[23] In an exercise, we assume that benthic foraminiferal δ18O, Mg/Ca, and Mg/Ca-based temperatures should positively covary, as observed in Pleistocene records [e.g., Martin et al., 2002; Elderfield et al., 2010]. We therefore consider the additional parameters (e.g., seawater Mg/Ca, diagenesis, carbonate ion effect) that may influence foraminiferal Mg/Ca [Lear et al., 2000, 2004; Billups and Schrag, 2003; Elderfield et al., 2006] and explore the extent to which these parameters must have changed in order to reconcile the benthic δ18O and Mg/Ca records. We then evaluate whether these “competing” parameters can account for some or all of the observed variability observed in the Mg/Ca record from Site 1209 (Table 2).

Table 2. Estimates of the Potential Impact of Parameters Other Than Temperature on Mg/Ca Record at Site 1209
ModelEvidence ForEvidence Against
Changing seawater Mg/Ca overprints short-term (<106 years) variations in benthic Mg/Ca record Long residence time for Mg and Ca in seawater (>106 years)
Changing seawater Mg/Ca overprints long-term (>106 years) variations in benthic Mg/Ca recordAll reconstructions and modeled histories of seawater Mg/Ca show increase of up to 0.5 mol/mol; an increase of 0.2 mol/mol could overprint to 3°C of cooling 
Bottom water Δ[CO32−] overprints short-term (<106 years) variations in benthic Mg/Ca record No significant variations observed in carbonate accumulation; potential Δ[CO32−] bias on Mg/Ca temperatures during CAE events acts to suppress rather than amplify the apparent negative covariation of benthic δ18O and temperature.
Bottom water Δ[CO32−] overprints long-term (>106 years) variations in benthic Mg/Ca recordIncreasing planktonic foraminiferal fragmentation in late middle EocenePlanktonic fragmentation not necessarily representative of bottom water Δ[CO32−]
Diagenetic alternation overprints variations in benthic Mg/Ca recordBenthic foraminifera appear “frosty” and increasingly fragmented in the late middle Eocene. The uncertainty is difficult to accurately quantify.Intersample foraminifera preservation variable; the uncertainty is difficult to accurately quantify

[24] We show below that changing seawater Mg/Ca probably cannot account for the observations at Site 1209, and that the discrepancies between the δ18Obenthic and foraminiferal Mg/Ca records cannot be reconciled with existing proxy estimates of carbonate dissolution at Site 1209. The effect of preservation on the Site 1209 bottom water temperature reconstruction is difficult to estimate, and although benthic foraminifera at Site 1209 do not exhibit the exquisite preservation found in continental shelf and slope environments, the record is less likely to contain dissolution-related artifacts than records from deeper sites. Consideration of Nontemperature Effects on Foraminiferal Mg/Ca: Seawater Mg/Ca Ratios

[25] Absolute temperature estimates based on benthic foraminifera Mg/Ca are sensitive to the precise value of seawater Mg/Ca used in calculations [Tripati et al., 2003; Billups and Schrag, 2003; Sexton et al., 2006b], and there are several different reconstructions for the history of Cenozoic seawater Mg/Ca (Figure 4). It is unlikely that short-term variations (<1 Myr) in foraminiferal Mg/Ca records from the middle Eocene reflect changes in seawater Mg/Ca because of the long residence times of Mg and Ca in the oceans (107 and 106 years, respectively [Broecker and Peng, 1982]). However, as the total duration of the benthic foraminiferal Mg/Ca record for Site 1209 exceeds the residence time of these ions in seawater, some component of the long-term (>1 million years) trends may reflect variations in seawater Mg/Ca.

[26] We consider a range of seawater Mg/Ca histories and evaluate the effect on bottom water temperature estimates (Figure 4). In order to reconcile the benthic foraminiferal δ18O and Mg/Ca records (i.e., to have them positively covary), seawater Mg/Ca would have had to increase during the middle Eocene (Figure 4). If deep ocean temperatures at Site 1209 cooled by ∼3°C over the middle Eocene (45–36 Ma), to attain this magnitude of cooling from the Site 1209 Mg/Ca record, seawater Mg/Ca would have had to increase by ∼0.2 mol/mol over 9 million years (green model in Figure 4).

[27] Existing reconstructions suggest that seawater Mg/Ca may have varied by as much as 60% over the past 65 million years [Sandberg, 1983; Wilkinson and Algeo, 1989; Lowenstein et al., 2001; Dickson, 2002; Horita et al., 2002; Creech et al., 2010]. As discussed previously by Tripati et al. [2003], Billups and Schrag [2003], and Sexton et al. [2006b], over tens of millions of years there are notable differences between published seawater Mg/Ca reconstructions that are based on fluid inclusion concentrations in halite crystals [e.g., Horita et al., 2002] (closed squares in Figure 4a) and models based on midocean ridge spreading rates [e.g., Stanley and Hardie, 1998] and cation fluxes [Wilkinson and Algeo, 1989] (circles in Figure 4a). During the middle Eocene, interpolations from all proxy- and model-based reconstructions suggest a relatively small change (<0.5 mol/mol) in seawater Mg/Ca may have occurred, consistent with what is needed to reconcile the Site 1209 benthic Mg/Ca record with a long-term 3°C cooling.

[28] Within the considered range of seawater Mg/Ca histories, there are a number of different of scenarios that are compatible with both the estimated long-term cooling of 3°C and proxy/model-based data (green and light blue models in Figure 4). However, the different scenarios result in different absolute temperature estimates. The cation flux-based seawater Mg/Ca reconstruction of Wilkinson and Algeo [1989] predicts average bottom water temperatures at Site 1209 of ∼9°C. We assume that the Wilkinson and Algeo [1989] seawater Mg/Ca reconstruction is most appropriate for the Eocene (Figure 4c) as the absolute temperature estimates are most compatible with other independent temperature estimates from alkenones [Weller and Stein, 2008].

[29] Although it is possible to reconcile the long-term change in foraminiferal Mg/Ca–bottom water temperatures at Site 1209 with the decline in benthic foraminiferal δ18O by changing seawater Mg/Ca, there are still notable discrepancies on shorter timescales (<1 Myr). Given the residence time of these ions in seawater, it is unlikely that the negative covariation of Mg/Ca and benthic δ18O on relatively short timescales (<106 years) results from changing seawater Mg/Ca. Bottom Water Carbonate Saturation

[30] The influence of bottom water carbonate saturation (Δ[CO32−]) on foraminiferal Mg/Ca ratios may introduce some bias in paleotemperature reconstructions, although the exact cause of this effect is not yet well understood and its magnitude is of debate [Elderfield et al., 2006; Rosenthal et al., 2006; Lear, 2007; Yu and Elderfield, 2008; Elderfield et al., 2010]. Recent work suggests that the sensitivity of foraminiferal Mg/Ca ratios to Δ[CO32−] may differ significantly between species [Elderfield et al., 2006; Yu and Elderfield, 2008], with infaunal species exhibiting much weaker sensitivities in comparison to epifaunal species [Elderfield et al., 2010]. The reason for the apparent differences in the species specific sensitivity of Mg/Ca ratios to Δ[CO32−] is unclear and likely complex, but Elderfield et al. [2010] suggest that the weaker sensitivity of infaunal species may in part reflect the fact that they calcify in pore waters rather than bottom waters. Although pore water temperatures in the upper few centimeters of the sediment (where O. umbonatus resides [Corliss, 1985]) will be the same as bottom waters, pore water Δ[CO32−] tends away from bottom water values toward zero [Martin and Sayles, 1996]. As a result, the Mg/Ca ratio of infaunal species may be less sensitive to changes in Δ[CO32−]. In fact, there is some indication from Mg/Ca values of modern Oridorsalis umbonatus that this species may not be sensitive to changes in saturation state [Rathmann and Kuhnert, 2008]. This study was based on comparing estimates of pore water Δ[CO32−] to test Mg/Ca ratios in a limited number of samples (n = 6), and therefore additional study is necessary to test whether this conclusion is robust.

[31] Using the sensitivity of benthic foraminiferal Mg/Ca to bottom water Δ[CO32−] established for Cibicidoides wuellerstorfi [Elderfield et al., 2006] we include in our error propagation a conservative estimate of the error in bottom water temperature that would result from a large change in carbonate ion saturation of ∼20 μmol/kg [Tripati and Elderfield, 2005]. Based on a new core top calibration for Oridorsalis umbonatus (C. F. Dawber and A. K. Tripati, Relationships between bottom water carbonate saturation and element/Ca ratios in core top samples of the benthic foraminifera Oridorsalis umbonatus, submitted to Paleoceanography, 2011), we note that this sensitivity may overestimate the contribution of changes in carbonate saturation to the Mg/Ca record of Oridorsalis umbonatus.

[32] Previous studies have demonstrated that there were large oscillations in Pacific deep-water carbonate preservation during the middle Eocene, likely linked to variations in global ice storage as evidenced by contemporaneous increases in benthic foraminiferal and seawater δ18O [Tripati et al., 2005]. These carbonate accumulation events (CAE) support substantial changes in the carbonate compensation depth (CCD) of up to 1 km [Tripati et al., 2005], which may have introduced some bias into deep water Mg/Ca–bottom water temperature reconstructions [Tripati et al., 2005]. A criticism of the records from Site 1218 is that they are from sites that are near the CCD [Edgar et al., 2007]. Site 1209, however, had a paleodepth of ∼1.9–2.5 km during the middle Eocene [Bralower et al., 2003], which is ∼2 km above the estimated average depth of the CCD and ∼1 km above the estimated CCD at its shallowest point. It is unclear if such large changes in deep-water carbonate saturation would have propagated to intermediate-depth bottom waters. Carbonate accumulation data for Site 1209 does not support large changes in intermediate-depth bottom water carbonate saturation [Hancock and Dickens, 2005; Bohaty et al., 2009], with one possible exception at ∼40.1 Ma [Bohaty et al., 2009]. Other carbonate dissolution proxy data have been cited as evidence for a relatively shallow Pacific lysocline during the middle Eocene (Hancock and Dickens, see discussion below, section We therefore consider whether the Mg/Ca bottom water temperature estimates at Site 1209 may be biased by changing carbonate saturation (Figure 5).

[33] At the peak of CAE-3 (∼41 Ma, the largest carbonate accumulation event), a transient, but substantial (0.7‰) decrease is observed in the Site 1209 benthic foraminiferal δ18O record (Figure 5), accompanied by a small (∼1°C) decrease in Mg/Ca-based bottom water temperatures. An increase in carbonate saturation at this time may have a positive bias on foraminiferal Mg/Ca values, resulting in an overestimation of the change in bottom water temperature. Thus, the observed temperature decrease should be considered as a minimum estimate. If changing carbonate saturation is causing a bias in foraminiferal Mg/Ca at Site 1209 during CAE-3, it is acting to suppress, and not amplify, the enigmatic relationship between benthic foraminiferal δ18O and Mg/Ca-based intermediate bottom water temperatures.

[34] During CAE-4, an increase of 0.6‰ is observed in the Site 1209 benthic δ18O record (Figure 5). The change in Mg/Ca-based temperature estimates across this event can be described by two trends. Between the onset of CAE-4 and peak accumulation, temperatures at Site 1209 are invariant, while between peak accumulation and the termination of CAE-4, temperatures decline. It may be possible that a decline in bottom water temperature between the onset of CAE-4 and peak accumulation is obscured by a positive bias resulting from an increase in carbonate saturation. If we assume the temperature decline should contribute approximately half of the benthic δ18O increase, the change in bottom water carbonate saturation needed to reconcile a ∼0.15 mmol/mol change in foraminiferal Mg/Ca is on the order of ∼+20 μmol/kg (based on the sensitivity established for C. wuellerstorfi [Elderfield et al., 2006]). This value is similar to the magnitude (although opposite in direction) to that calculated by Yu and Elderfield [2007] for the change in Atlantic deepwater carbonate saturation between the Last Glacial Maximum and the Holocene, which was accompanied by an 800 m shoaling of the carbonate saturation depth. The amplitude of the CCD deepening across CAE-4 (∼700 m) may be consistent with a potential carbonate saturation bias on foraminiferal Mg/Ca only if the change in intermediate-depth bottom water carbonate saturation is similar in amplitude (and direction) to deep water. This hypothesis is difficult to test at present given the difficulty in interpreting changes in carbonate accumulation in carbonate-dominated sediment. Interestingly, two proxies commonly used to infer carbonate dissolution, planktonic foraminiferal fragmentation, and the relative abundance of benthic foraminifera, suggest increased dissolution at Site 1209 between 39.5 Ma and 38.5 Ma [Hancock and Dickens, 2005] (Figure 1, 151.5 – 148 rmcd). If these proxy reconstructions are accurate, they suggest that the lysocline and carbonate saturation depth may have been decoupled during CAE-4. As a result, decreased carbonate saturation at intermediate sites may result in a negative carbonate ion bias on foraminiferal Mg/Ca–bottom water temperatures, which again is acting to suppress the apparent negative covariation of benthic δ18O and bottom water temperatures.

[35] There likely were changes in Pacific intermediate and deep-water carbonate saturation during the middle Eocene. However, if our present understanding of the sensitivity of foraminiferal Mg/Ca to changing carbonate saturation is appropriate for the Eocene, and carbonate dissolution reconstructions at Site 1209 are accurate, then changes in intermediate water carbonate saturation cannot be solely responsible for the apparent negative correlation between records of benthic δ18O and Mg/Ca at Site 1209 on timescales of <106 years. Dissolution and Diagenetic Alteration

[36] Diagenetic processes including dissolution and recrystallization may also influence the geochemistry of foraminiferal calcite, although the impacts are difficult to robustly determine. The influence of dissolution should be minimal given the relatively shallow water depth of Site 1209 relative to the carbonate compensation depth. Two proxies for dissolution have been cited as evidence for a shallower lysocline and enhanced dissolution at Site 1209 during parts of the middle and late Eocene [Hancock and Dickens, 2005]. We developed high-resolution records of these indices for this study (Figure 1); these are broadly consistent with the results from Hancock and Dickens [2005] and exhibit cumulative increases during the late middle Eocene.

[37] The tests of planktonic foraminifera may be more susceptible (than benthics) to postmortem dissolution as they calcify in surface waters that are likely to be more saturated (with respect to Δ[CO32−]) than their depositional environment and they have highly porous tests. In addition, dissolution in planktonic foraminifera begins as they descend through the water column [e.g., Schiebel et al., 2007]. The test structure of Eocene planktonic foraminifera also appears to influence the susceptibility to dissolution [e.g., Petrizzo et al., 2008]. Unless planktonic fragments are identified to a genera level, records of this index will be sensitive to changes in the faunal assemblage [Petrizzo et al., 2008]. We note that planktonic fragmentation in early Eocene sediments from Site 1209 could not unequivocally be attributed to carbonate dissolution [Petrizzo et al., 2008]. As the benthic foraminiferal abundance index is quoted with respect to the number of whole planktonic foraminifera, it may also be biased by these same factors. Nevertheless, increased planktonic fragmentation and benthic abundance at Site 1209 correlates with darker core material (Figure 1), which may indicate a reduction in the carbonate:organic carbon ratio. Although a few studies have documented notable heterogeneity of Mg/Ca in deep-water benthic foraminifera [e.g., Rathmann et al., 2004], the sensitivity of the Mg/Ca thermometer to dissolution is still poorly constrained.

[38] Scanning electron microscope (SEM) images of benthic foraminifera indicate an increasing effect of dissolution in the upper part of the studied interval. However, within a single sample, the preservation of individual tests varies significantly (refer to auxiliary material). Care was taken to select the best preserved (i.e., intact and nonchalky) foraminiferal tests for analysis and specimens that were fragmented or had obvious holes were not selected. None of the benthic foraminiferal δ18O or Mg/Ca excursions discussed below corresponds to a major change in foraminiferal preservation (Figure 1).

[39] It is likely that any secondary diagenetic calcite would have formed in pore waters similar in temperature and seawater δ18O to bottom waters. An inability to quantify the amount of “neomorphosed” spar present in tests, and the lack of consensus over appropriate partition coefficients for Mg in diagenetic calcite [Tripati et al., 2003; Sexton et al., 2006b], means that recrystallization represents a source of uncertainty that is difficult to accurately quantify.

4.3. Variations in Seawater δ18O During the Middle Eocene

[40] The fidelity of early Cenozoic seawater δ18O reconstructions depends largely on the accuracy of foraminiferal Mg/Ca–bottom water temperature reconstructions. With this caveat and those discussed above, and with the calculated minimum propagated error of 0.4‰ on each individually reconstructed value of seawater δ18O, we consider the evolution of glaciation during the middle Eocene. The seawater δ18O reconstruction from Site 1209 (Figure 5) supports large variations in ice volume back to the early middle Eocene (∼45.4 Ma). The long-term increase in seawater δ18O of ∼1.1‰ between ∼44 Ma (159.5 rmcd) and ∼37.5 Ma (141 rmcd) is consistent with the gradual build up of continental ice. Superimposed on this trend are a number of large amplitude (>0.6‰) excursions. Three major positive excursions are observed at approximately 44 Ma (∼158 rmcd), 41 Ma (∼153 rmcd) and 38 Ma (∼145 rmcd), some of which are also observed in records of benthic δ18O at this site and at other sites (1218 and 1260), and which could reflect the relatively rapid (<0.5 Ma) growth of large volumes of ice. The amplitude of the seawater δ18O shift at Site 1209 between ∼41 Ma and 39 Ma (∼1.3 ± 0.3‰, error based on the standard error of preexcursion and postexcursion means) is similar to that reported from Site 1218 in the deep Pacific at ∼41.6 Ma (∼1.5‰ [Tripati et al., 2005]).

4.3.1. Why Might Temperature and Seawater δ18O at Site 1209 Vary Inversely?

[41] The observed trends in δ18Obenthic and foraminiferal Mg/Ca suggest that temperature and seawater δ18O may have been negatively correlated, with warming at Site 1209 associated with increases in global ice storage. It is possible that a climatically associated change in the water mass bathing the site could potentially result the observed correlation. For example, in the modern ocean, Circumpolar Deep Water and North Atlantic Deep Water have similar densities, but differ in water δ18O by ∼0.4‰ and in temperature by ∼2°C. Data for Atlantic Site 1260 [Edgar et al., 2007] may be consistent with large water mass gradients in the late middle Eocene, as benthic δ18O values at the site are on average ∼0.5‰ lower than those observed at Site 1209. However, challenges in interpreting these apparent basinal differences as simply a function of water mass properties stems from the different species that were used, and the shallow infaunal nature of O. umbonatus (the primary taxon examined at Site 1209). It may eventually be possible to confidently assess possible contributions to Δδ18Obenthic from water mass reorganization using records of δ18O from more epifaunal benthic taxa, and/or from other tracers such as basinal gradients in δ13Cbenthic or ɛNd. However, variations in the δ13Cbenthic for O. umbonatus is likely to reflect changes in pore water conditions as well as water masses. Although there have been studies of Eocene deep water circulation published using epifaunal δ13Cbenth and other tracers [Wright and Miller, 1993; Thomas, 2004; Via and Thomas, 2006; Scher and Martin, 2006; Cramer et al., 2009], there is not yet a consensus on the number of water masses, their source regions, their geometry, their physical properties, or on the timing of major circulation changes (e.g., the opening of the Drake Passage and Tasman Gateway).

[42] Alternatively, if taken at face value, the trends in δ18Obenthic and foraminiferal Mg/Ca at Site 1209 may support the hypothesis that early Cenozoic glaciation may have been influenced by parameters other than temperature, such as high-latitude moisture supply [Prentice and Matthews, 1991]. To date, supporting evidence for the so-called “snow gun hypothesis” is equivocal. Sedimentary and geochemical data from the La Meseta Formation, exposed on the northern Antarctica Peninsula, support a strongly seasonal wet climate during the middle and late Eocene replaced by a frost-prone drier climate in the latest Eocene [Dingle et al., 1998], although it is unclear how widespread these conditions may have been across the continent given the maritime influence on the climate of peninsula and the lower latitude of the peninsula relative to the Antarctic continent. Unfortunately, Antarctica fossil vegetation records for the middle and late Eocene (45–38 Ma) are highly punctuated and may be limited by uncertainties associated with the “nearest modern analog” approach to estimating temperature and precipitation. However, additional support for the presence of a (relatively) wet climate on the continent comes from model simulations of glacial conditions under 2x preindustrial pCO2, which yield higher mean annual precipitation on the peninsula compared to the whole continent average [Thorn and DeConto, 2006]. The recovery of new middle Eocene sequences should allow for more rigorous determinations of Antarctic glacial history and the testing of the snow gun hypothesis.

4.3.2. Estimating Ice Budgets From Seawater δ18O Reconstructions

[43] The estimation of ice volume from seawater δ18O assumes reconstructed seawater δ18O values are accurate and reflect changes in whole ocean composition. A calibration of 0.09‰/10 m sea level equivalent has been proposed for the greenhouse-icehouse transition [DeConto and Pollard, 2003], assuming a mean ice δ18O of ∼ −35‰ (V-SMOW). This calibration at least partly accounts for warmer low- and high-latitude oceanic and atmospheric temperatures in the Eocene [Lear et al., 2000; Tripati et al., 2003, 2005; Tripati and Zachos, 2002; Pearson et al., 2007] and is consistent with a seawater δ18O–sea level calibration derived from proxy reconstructions of earliest Oligocene glaciation [Katz et al., 2008]. Climate simulations suggest that during the early Cenozoic, the mean isotopic composition of ice would be much heavier than modern. Mean values of approximately −30‰ were derived from modeling experiments [DeConto et al., 2008]. For comparison, the isotopic composition of ice on Antarctica today ranges from −20‰ to −50‰ [Morgan, 1982]. However, as the isotopic composition of an ice sheet can evolve over time, the approach of using a mean δ18O value for ice may be over simplistic [DeConto et al., 2008], although it probably is a reasonable first-order assumption.

[44] Based on this calibration, the increase in whole-ocean δ18O arising from fully glaciating Antarctica is estimated to have been 0.5‰ based on an assumed paleotopography [DeConto and Pollard, 2003]. However, new constraints on the paleotopography of West Antarctica indicate that there may have been significantly more land area available for ice storage during the Eocene [Wilson and Luyendyk, 2009], and therefore this threshold may be somewhat higher. In addition, given the large uncertainties in reconstructing seawater δ18O from paired measurements of foraminiferal Mg/Ca and δ18O, the method of using absolute seawater δ18O values to determine ice volume will only be able to resolve the presence of large amounts of ice. We use a “best guess” threshold of 0.9‰ to assess when glacial ice may have occurred off of the Antarctic continent, which is based on the published value of 0.5‰ [DeConto and Pollard, 2003] and the uncertainty in reconstructed seawater δ18O that we estimate as 0.4‰. The seawater δ18O reconstruction from Site 1209 exhibits multiple, large-amplitude shifts that are in excess of the published threshold at ∼44 Ma, ∼41.4 Ma, 41 Ma, ∼38.6 Ma and ∼38 Ma (Figure 5). Of these shifts, those at 41 Ma and 38.6 Ma are in excess of the threshold and estimated uncertainty in seawater δ18O values.

[45] The magnitude of seawater δ18O excursions at Sites 1209 could therefore be consistent with some storage of ice in both hemispheres during parts of the middle Eocene, as previously inferred from other proxy records [Tripati et al., 2005; Moran et al., 2006; Tripati et al., 2008; St. John, 2008; Stickley et al., 2009]. Ice-rafted debris provides direct evidence of glaciers at sea level, and has been documented in the Southern Ocean back to 45 Ma [Ehrmann and Mackensen, 1992], and in middle Eocene sediments from the Arctic and Greenland basins back to ∼46 Ma [Moran et al., 2006; Tripati et al., 2008; St. John, 2008; Stickley et al., 2009]. These records of IRD exhibit large changes in mass accumulation rate, providing evidence for ephemeral glacial episodes. Sea ice–associated diatoms have also been found in Arctic Ocean sediments dating back to 47.5 Ma [Stickley et al., 2009].

[46] It may not be immediately apparent how the stable isotopic data and physical evidence of near-freezing temperatures at sea level, such as IRD which is an indicator of seasonal sea ice and/or glacial ice in the work by Moran et al. [2006], St. John [2008], and Tripati et al. [2008] and sea ice diatoms [Stickley et al., 2009], can be reconciled with reconstructions of warm high-latitude surface ocean temperatures (8°C–15°C) based on organic proxies and planktic foraminiferal Mg/Ca [Hollis et al., 2009; Bijl et al., 2009; Liu et al., 2009]. Polar SSTs of this magnitude are also implied by benthic foraminiferal Mg/Ca data for the late middle Eocene (e.g., this study) [see also Lear et al., 2000; Tripati et al., 2005]. MAT estimates for Arctic Canada at 45 Ma (early middle Eocene) based on the nearest living analog of fossil vegetation and pollen [Jahren, 2007, and references therein] range from 12°C to 15°C.

[47] One possibility is the data simply reflect regional differences in climate (e.g., the occurrence of sea ice diatoms and pollen in early middle Eocene Arctic deposits that are from geographically distant locales). It is also possible that the reconstructions from these various studies are recording conditions during different times (e.g., seawater δ18O and IRD peaks are from transient “glacial” maxima and other data may represent “interglacial” extremes). There is some indirect support for this from sedimentary evidence for alternations between near freezing and cool temperate conditions reported from Paleogene deposits from Svalbard based on the occurrence of glendonites, mollusks, and flora [Spielhagen and Tripati, 2009]. In addition, it is known that at high-latitude sites, proxy reconstructions of SST that are derived from biologically precipitated compounds reflect a bias to the season of growth/maximum abundance. It may be that SST reconstructions derived from organic compounds such as TEX86 and U37k′ [e.g., Liu et al., 2009] may be preferentially recording warm month conditions (not mean annual temperatures) during warm intervals of the Eocene. Finally, it is clear that there are competing (nontemperature) influences on foraminiferal Mg/Ca ratios (i.e., carbonate ion effects, changing seawater Mg/Ca ratios, preservational biases) that can introduce uncertainties in SST reconstructions [e.g., Tripati et al., 2003]. It has also been shown that there are similar biases in organic SST proxies. For example, Turich et al. [2007] and Lipp et al. [2008] discuss the possibility of nontemperature influences on TEX86.

4.3.3. Comparisons With Sedimentary Records

[48] The reconstructed seawater δ18O excursions observed at Site 1209 appear to coincide with known sequence boundaries [e.g., Browning et al., 1996; Miller et al., 2005; Pekar et al., 2005]. Assuming that the age model for Site 1209 is robust, the oldest seawater δ18O excursion correlates with a sequence boundary reported on the South Tasman Rise (Lu3 at ∼44.0–44.2 Ma [Pekar et al., 2005]) and in European Basins [Hardenbol et al., 1998] and a major fall in sea level (∼70 m between 44.2 and 42 Ma) on the New Jersey Coastal Plain [Browning et al., 1996; Miller et al., 2005], which is also recognized in the global sea level curve [Haq et al., 1987]. This event has previously been interpreted to reflect ice growth [Browning et al., 1996; Miller et al., 2005; Pekar et al., 2005]. Large negative excursions of 0.7 ± 0.3‰ and 1.2 ± 0.3‰ are observed in the Site 1209 seawater δ18O reconstruction at ∼41.3 Ma (153.5 rmcd) and 38.6 Ma (147 rmcd), respectively, and correlate with maximum flooding surfaces on the global sea level curve (at ∼41.2 and 38.8 Ma [Haq et al., 1987]).

[49] Although the timing of middle and late Eocene sea level variations estimated from the sequence stratigraphic record [e.g., Browning et al., 1996; Miller et al., 2005; Pekar et al., 2005] and the 1209 foraminiferal-based proxy reconstructions appear to be consistent, the amplitude of variations are dissimilar, as is observed during earliest Oligocene glaciation [e.g., Coxall et al., 2005; Katz et al., 2008]. Some studies [Pekar et al., 2002; Katz et al., 2008] have argued that this discrepancy, in part, reflects the difference between estimates of “apparent” sea level (volume of ocean water, estimated from benthic and seawater δ18O reconstructions) and eustasy (i.e., hydroisostatically compensated sea level relative to a fixed datum, calculated from sequence stratigraphy and backstripping). The discrepancy between existing sea level reconstructions can be reduced by making a published correction of +32% on eustatic estimates [Pekar et al., 2002]. However, others dispute the concept of apparent sea level when applied to whole ocean (see review by Mitrovica [2003]). Even if a correction is applied, middle Eocene sequence stratigraphic estimates of sea level variations are still smaller than those estimated from the Site 1209 reconstructions. This observation may indicate (1) that the large magnitude variations observed in the Site 1209 reconstruction are biased and/or (2) that sequence stratigraphic estimates are minimum estimates.

4.4. Paleoceanographic Records Across the “Middle Eocene Climatic Optimum”

[50] A transient and large amplitude (∼1.0‰–1.5‰) negative excursion has been reported in records of bulk and fine fraction carbonate and planktonic and benthic foraminifera δ18O at ∼40.0 Ma from multiple sites in the Southern, Atlantic, Indian and Tethys oceans [Bohaty and Zachos, 2003; Bohaty et al., 2009; Spofforth et al., 2010]. This excursion, known as the middle Eocene Climatic Optimum (MECO) [Bohaty and Zachos, 2003], has previously been interpreted to reflect a transient warming of up to ∼6°C based on the assumption of ice-free conditions during this interval [Bohaty et al., 2009]. Although there is some evidence to support warming at the high southern latitudes at this time [Houben et al., 2009; Burgess et al., 2009], the respective roles of temperature and seawater δ18O have not been determined at most localities.

[51] Based on a compilation of data from many open ocean localities, Bohaty et al. [2009] noted that the MECO δ18O excursion approximately coincided with the first occurrence of D. scrippsae at the low latitudes and with planktonic foraminiferal zone P13/E12 (O. beckmanni zone [Berggren et al., 1995; Berggren and Pearson, 2005]). At Site 1209 there are significant discrepancies (∼3 m) between the position of these calcareous nannofossil and planktonic foraminifera datums (refer to auxiliary material). Recent work suggests that the temporal range of either/both of these species may be latitudinally diachronous [Bohaty et al., 2009; Edgar, 2008], which could account for the discrepancies between these datums at Site 1209. We also note that O. beckmanni is extremely rare at Site 1209 [Petrizzo et al., 2005] and the nominal range of this species documented at Site 1209 was based on 3 specimens [Petrizzo et al., 2005; M. R. Petrizzo, personal communication, 2009]. We found no additional evidence of this species.

[52] Assuming the calcareous nannofossil age model is robust, the 40.4–40.0 Ma interval at Site 1209 is condensed and the temporal resolution afforded by our sampling is too low to resolve tens-to-hundred thousand year variations in benthic foraminiferal δ18O. In contrast, if the reported range of O. beckmanni at Site 1209 [Petrizzo et al., 2005] is taken at face value and/or the calcareous nannofossil calibrations are inaccurate, the 0.8‰ negative δ18O excursion at ∼153.5 rmcd may correspond to the same MECO isotope excursion identified at other localities [Bohaty and Zachos, 2003; Bohaty et al., 2009; Spofforth et al., 2010]. Interestingly, benthic foraminiferal Mg/Ca values increase slightly at the onset of this negative δ18O excursion, although minimum δ18O values are associated with a cooling of intermediate waters at Site 1209 (Figures 1 and 5). If Mg/Ca paleotemperatures are accurate, these data support attributing the majority of the δ18O excursion observed at ∼153.5 rmcd to a change in seawater δ18O (Figure 5).

[53] At present, it is not possible to precisely correlate the Site 1209 records with the MECO δ18O excursion identified at other localities [Bohaty and Zachos, 2003; Bohaty et al., 2009]. However, the seawater δ18O reconstruction for Site 1209 exhibits multiple, large amplitude (0.6‰–1.3‰) shifts during the Eocene that we interpret to reflect substantial variations in global ice volume with possible ice storage in both hemispheres. This interpretation is supported by the independent evidence of sea ice and glacier IRD in middle Eocene sediments at the high northern latitudes [Moran et al., 2006; Eldrett et al., 2007; Tripati et al., 2008; St. John, 2008; Stickley et al., 2009], episodic near freezing conditions in the Arctic during the Eocene [Spielhagen and Tripati, 2009] and the glacioeustatic interpretation of the seawater δ18O reconstruction for Site 1218 [Tripati et al., 2005]. Together, these lines of evidence suggest that the assumption of ice-free conditions or small/negligible ice volume contributions to benthic foraminiferal δ18O records during late middle Eocene may not appropriate. Additional independent temperature records are required to precisely determine what the MECO δ18O isotope excursion represents at each locality.

5. Summary

[54] A detailed stable isotope record from ODP Site 1209 on Shatsky Rise indicates there were significant variations in benthic foraminiferal δ18O during the late middle Eocene. We use multiple approaches to interpret the benthic δ18O record, including partitioning changes in benthic δ18O to reflect an equal contribution of temperature and seawater δ18O; use of a benthic δ18O-apparent sea level calibration, and the estimation of temperature and seawater δ18O changes using benthic foraminiferal Mg/Ca. Our record does not resolve which approach is more accurate; however, it does support that previously reported excursions in Pacific records of benthic and seawater δ18O, and benthic Mg/Ca, were not site-, contaminant-, or preservation-related anomalies. The excursions at Site 1209 have well-defined structure and are consistent with the hypothesis that there with major glaciations and deglaciations at ∼44–43 Ma, ∼42–40 Ma, and ∼39–38 Ma associated with seawater δ18O changes of 0.3‰ to 0.6‰ or more.


[55] The authors gratefully acknowledge L. Booth, S. Crowhurst, M. Greaves, M. Hall, and C. Sindrey for providing technical assistance and S. Bohaty, T. Bralower, R. Eagle, H. Elderfield, D. Hodell, K. Miller, M. R Petrizzo, D. Spofforth, S. Turchyn, and members of the Department of Earth Sciences at Cambridge for discussions of this work. We would also like to thank Editor C. Charles, reviewer S. Pekar, and several anonymous reviewers for their comments on earlier drafts of the manuscript. C.F.D. was funded by a NERC studentship. A.K.T. was supported by a NERC postdoctoral fellowship, a junior research fellowship from Magdalene College, and the UCLA Division of Physical Sciences. This research used samples and data provided by the Ocean Drilling Program. All δ18O and Mg/Ca data will be archived on the PANGAEA database upon publication.