Late Miocene through early Pliocene deep water circulation and climate change viewed from the sub-antarctic South Atlantic

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1 Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 Late Miocene through early Pliocene deep water circulation and climate change viewed from the sub-antarctic South Atlantic Katharina Billups College of Marine Studies, University of Delaware, 700 Pilottown Road, Lewes, DE 19958, USA Received 21June 2001; accepted 12 April 2002 Abstract Benthic foraminiferal stable isotope records for the past 11 Myr from a recently drilled site in the sub-antarctic South Atlantic (site 1088, Ocean Drilling Program Leg 177, 41 S, 15 E, 2082 m water depth) provide, for the first time, a continuous long-term perspective on deep water distribution patterns and Southern Ocean climate change from the late Miocene through the early Pliocene. I have compiled published late Miocene through Pliocene stable isotope records to place the new South Atlantic record in a global framework. Carbon isotope gradients between the North Atlantic, South Atlantic, and Pacific indicate that a nutrient-depleted watermass, probably of North Atlantic origin, reached the sub-antarctic South Atlantic after 6.6 Ma. By 6.0 Ma the relative proportion of the northernprovenance watermass was similar to today and by the early Pliocene it had increased to greater than the modern proportion suggesting that thermohaline overturn in the Atlantic was relatively strong prior to the early Pliocene interval of inferred climatic warmth. Site 1088 oxygen isotope values display a two-step increase between V7.4 Ma and 6.9 Ma, a trend that parallels a published N 18 O record of a site on the Atlantic coast of Morocco. This is perhaps best explained by a gradual cooling of watermasses that were sinking in the Southern Ocean. I speculate that relatively strong thermohaline overturn at rates comparable to the present day interglacial interval during the latest Miocene may have provided the initial conditions for early Pliocene climatic warmth. The impact of an emerging Central American Seaway on Atlantic^Pacific Ocean upper water exchange may have been felt in the North Atlantic beginning in the latest Miocene between 6.6 and 6.0 Ma, which would be V1.5 Myr earlier than previously thought. ß 2002 Elsevier Science B.V. All rights reserved. Keywords: Miocene; Ocean Drilling Program site 1088; paleoceanography; Pliocene; stable isotopes 1. Introduction Numerous studies have focused on early Pliocene climate because of a widespread interest in * Tel.: , ext. 4349; Fax: address: kbillups@udel.edu (K. Billups). assessing climate sensitivity to changes in greenhouse gas levels and oceanic heat transport during an interval of time of relative global warmth (e.g., Raymo et al., 1996; Dowsett et al., 1992; Billups et al., 1998a; Ravelo and Andreasen, 2000). By now much evidence exists that enhanced oceanic heat transport played a signi cant role in sustaining high latitude warmth during that period of / 02 / $ ^ see front matter ß 2002 Elsevier Science B.V. All rights reserved. PII: S (02)

2 288 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 time (e.g., Dowsett et al., 1992, Billups et al., 1998a, 1998b, 1999; Ravelo et al., 1997; Kwiek and Ravelo, 1999; Ravelo and Andreasen, 2000). Observational and modeling studies suggest that northward heat transport and thermohaline overturn was not as e ective as today until the closure of the Central American Seaway (CAS) during the early Pliocene (e.g., Keigwin, 1982; Mikolajewicz and Crowley, 1997; Haug and Tiedemann, 1998). However, other studies have shown that modern-type deep water formation in the North Atlantic commenced earlier, during the late Miocene (Wright et al., 1991), perhaps associated with subsidence of the Greenland^Scotland Ridge (Wright and Miller, 1996). In contrast to early Pliocene climatic warmth, the late Miocene is thought to be a time of global cooling (e.g., Mercer, 1976; Shackleton and Kennett, 1975; Ciesielski et al., 1982). Ice sheets expanded on East Antarctica and the West Antarctic ice sheet became grounded below sea level (Ciesielski et al., 1982). Increases in biogenic silica accumulation rates and increases in subpolar planktonic foraminiferal species indicate cooling of the circum-antarctic region (Brewster, 1980; Loutit, 1981; Kennett and Vella, 1975). By some accounts, these climatic changes eventually lead to the isolation and desiccation of the Mediterranean (the Mediterranean salinity crisis) via glacioeustatic sea level falls of up to 40^60 m (Shackleton and Kennett, 1975; McKenzie and Oberhaensli, 1985; Hodell et al., 1986; Hodell and Kennett, 1986; Hodell et al., 1994). However, a late Miocene increase in global ice volume is still being debated because it is not ubiquitously recorded in late Miocene benthic foraminiferal N 18 O records (Hodell et al., 2001). In this study, I revisit late Miocene through early Pliocene deep water circulation and climate change. Previous studies have focused either on early Pliocene or late Miocene paleoceanography, but thus far a long-term study that includes both intervals has been lacking. I present a new benthic foraminiferal stable isotope record from the Atlantic sector of the sub-antarctic Southern Ocean that spans the past 11 Myr. I take advantage of recently published high resolution benthic foraminiferal carbon isotope records from the North Atlantic and Paci c oceans to illustrate that the relative rate of thermohaline overturn reached modern proportions by V6 Ma and exceeded modern conditions during the early Pliocene. The oxygen isotope values increase during the late Miocene with a maximum at 6.9 Ma, which is initially most easily explained by an increase in global ice volume, but after a closer look the trend may be more consistent with a cooling of watermasses sinking in the Southern Ocean. Results from site 1088 provide evidence that thermohaline overturn was relatively strong prior to the interval of early Pliocene climatic warmth, which leads me to speculate that the closing history of the CAS in uenced North Atlantic deep water (NADW) formation beginning between V6.6 and 6.0 Ma Deep water circulation Site 1088 (V2000 m water depth), which is located on the Agulhas Ridge, is situated within the modern mixing zone of upper circum-polar deep water (CDW), a nutrient-enriched watermass representative of the high latitude Southern Ocean, and NADW, a nutrient-depleted watermass created in the high latitude North Atlantic (Fig. 1, Table 1). Oceanic nutrient gradients, as re ected in the N 13 Cof4CO 2 of deep water and preserved by benthic foraminiferal N 13 C values, are commonly used to reconstruct deep water circulation patterns for the past (Duplessy et al., 1984; Mix and Fairbanks, 1985; Oppo and Fairbanks, 1987; Raymo et al., 1990; Charles and Fairbanks, 1990; Wright et al., 1991). Benthic foraminiferal N 13 C values record the relative amount of respired carbon in the deep ocean and therefore provide a proxy for watermass age as well as mean oceanic nutrient levels (Broecker and Peng, 1982). Depending on relative rates of deep water formation and the amount of nutrients in the deep ocean, carbon isotope gradients between the Atlantic, Southern Ocean, and Paci c can be large, as in the modern ocean (Table 1), or small, as during the early Miocene (e.g., Woodru and Savin, 1989; Delaney and Boyle, 1987; Wright et al., 1992). The challenge is to distinguish N 13 C variations caused by regional deep watermass mixing

3 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ Fig. 1. Schematic cross-section through the Atlantic sector of the Southern Ocean along the Leg 177 transect showing locations of sites with respect to the vertical distribution of water masses (Shipboard Scienti c Party, 1999, Leg 177). Abbreviations are as follows: AABW = Antarctic Bottom Water; AAIW = Antarctic Intermediate Water; SASW = sub-antarctic Surface Water; SAF = sub-antarctic Front; PF = Polar Front. from mean ocean N 13 C uctuations caused by changes in the mean ocean nutrient inventory. In theory, this distinction can be made if the N 13 C value of the NADW source water and the N 13 C value of mean ocean deep water (e.g., the deep Paci c Ocean) can be separately characterized (e.g., Oppo et al., 1990; Wright et al., 1991). In practice, however, determining the N 13 C value of an endmember watermass is di cult. For example, in the modern eastern equatorial Paci c regional changes in surface water productivity Table 1 Site summary Site Location Water depth N 13 a Cof4CO 2 Time Reference (m) Interval (Ma) PN, 15 W ^10 Hodell et al. (2001) N, 33 W ^3 Raymo et al. (1989) PN, 23 05PW ^ 1Wright et al. (1991) Sale 34 N, 06 47PW N/A N/A 5^9 Hodell et al. (1994) PN, 43 29PW ^5 Billups et al. (1998b) PN, 42 54PW ^7 Shackleton and Hall (1997) PS, 18 05PE ^11 Wright et al. (1991) S, 15 E ^11 This study S53P, 07 25PE ^5 Hodell and Venz (1992) 5^10 Mueller et al. (1991) 10^7 Wright et al., (1991) PS, PE ^ 1Wright et al., (1991) PS, PE 1533 N/A 4.3^14 Wright et al. (1992); Zachos et al. (2001) PS, 90 49PW ^6.5 Shackleton et al. (1995) PN, PW ^5 Mix et al. (1995) a Kroopnick (1985); Ostlund et al. (1987).

4 290 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 and recirculation of deep water in restricted basins can lower benthic foraminiferal N 13 C values with respect to the global mean (Mix et al., 1995). In the North Atlantic, deep waters form from a number of individual components that sink in different source regions (e.g., Schmitz and McCartney, 1993). Any changes in the relative proportion of the individual components or the 13 C/ 12 C ratio of the source water through time would a ect the N 13 C value of the watermass that exits the deep North Atlantic and enters the circum-polar South Atlantic. Unfortunately, in reconstructions of past deep water circulation patterns one is limited by the availability of records to represent the source regions. However, since the last synthesis of late Miocene deep water circulation (Wright et al., 1991) a number of stable isotope records have become available from the North Atlantic and Paci c (Fig. 2, Table 1). I will take advantage of these recently published records to estimate the varying e ects of choosing a particular endmember site to reconstruct deep water circulation changes in the sub-antarctic Southern Ocean Details on the site selection strategy I am comparing the site 1088 N 13 C record to a number of published records from the North Atlantic, Southern Ocean, and Paci c (Fig. 2, Table 1). To represent the late Pliocene/Pleistocene North Atlantic, I use the V3.5-Myr-long record from Ocean Drilling Program (ODP) site 607 located on the western ank of the Mid-Atlantic Ridge because it lies close to the modern core of NADW (Raymo et al., 1992). For the early Pliocene, I use the ODP site 925 N 13 C record from the western tropical Atlantic because it has already been established that this site remained within the core of NADW during that time (Billups et al., 1998b). Two sites are available that span the later portion of the late Miocene: ODP site 926, also located in the deep western tropical Atlantic, and site 982 at intermediate water depth on the Rockall Plateau in the subpolar North Atlantic (Fig. 2). Because of the large spatial as well as vertical separation between these sites, it will be possible to directly assess the e ect of the choice of North Atlantic endmember N 13 C values on the deep water reconstruction, at least for the period of overlap between 5 and 7 Ma. The earlier late Miocene is represented at Deep Sea Drilling Project (DSDP) site 608, the same site used by Wright et al. in their 1991 (Wright et al., 1991) synthesis of late Miocene evolution of deep water circulation. I also compare the site 1088 record with previ- Fig. 2. Location of sites discussed in this study (see also Table 1).

5 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ ously published records from the Atlantic sector of the Southern Ocean. Together with DSDP site 360 to the north, site 1088 forms a vertical transect through deep watermasses (Table 1, Fig. 2). To create a spatial transect, I compare the records of ODP site 704 to the south on the Meteor Rise with that of site 1088 (Table 1, Fig. 1). Published benthic foraminiferal N 13 C records from sites 360 and 704 illustrate that these locations were sensitive to deep water circulation changes during the late Miocene (Wright et al., 1991) and the Plio/ Pleistocene (site 704 only, Hodell and Venz, 1992; Raymo et al., 1992; Hodell, 1993). To constrain global ocean N 13 C variability during the Plio/Pleistocene, I use ODP site 849. Site 849 is located in the eastern equatorial Paci c to the west of the Paci c Rise and is the most suitable record to constrain global ocean N 13 C variability (Mix et al., 1995). But it only spans the past 5 Myr. For the late Miocene, I can use the western equatorial Paci c N 13 C record from DSDP site 289 (Wright et al., 1991), but this record does not extend into the early Pliocene. This leaves an V1Myr gap. I could bridge this gap using eastern equatorial Paci c site 846 located in the Peru Basin (Fig. 2). However, in this region modern ocean deep water N 13 C values may be lower than the global mean because of high surface water productivity and restricted deep water circulation (Mix et al., 1995). As I shall show, the most conservative approach to link the latest Miocene with the earliest Pliocene is to interpolate between site 289 and site 849 benthic foraminiferal N 13 C values. 2. Methods 2.1. Sampling and stratigraphy At site 1088 three holes were drilled that together comprise a V230-m-thick sedimentary sequence. Site 1088B, the focus of this study, spans the Holocene through middle upper Miocene. To complete the upper Miocene sequence, I included sediments from site 1088C. Recovery of the entire sedimentary sequence was better than 95%, but gaps exist at core breaks. Upper Miocene sediments consist primarily of nannofossil ooze with an increasing component of foraminifera toward the Holocene (Shipboard Scienti c Party, 1999). Sedimentation rates estimated from shipboard biostratigraphy remain relatively constant at V24 m/myr during the upper Miocene and decrease to V9 m/myr just below the Miocene/Pliocene boundary (Fig. 3). I initially sampled the cores at 150 cm and 75 cm, respectively, to achieve an average time step of V70 kyr. I later doubled the temporal resolution of the upper-most Miocene through Pliocene portion of the record, and I added a few measurements adjacent to intervals where rapid and large changes in the stable isotope stratigraphy became apparent. All data are presented in Appendix I and archived electronically at the World Data Center-A. Overall the sampling strategy su ces to fully characterize long terms trends in the stable isotope records. The additional measurements allow me to more precisely describe the timing and character of deep water circulation changes. The age model of site 1088 is based on shipboard calcareous nannofossil biostratigraphic datums reported on the Berggren et al. (1995) time scale (e.g., Fig. 3). I linearly interpolated between these datums to convert from meters composite depth to geologic time. As demonstrated below (3.2. Carbon isotope stratigraphy), the carbon isotope stratigraphy at site 1088 can be used to re ne a portion of the late Miocene age model Analytical Bulk sediments were oven-dried, disaggregated in a sodium metaphosphate solution, and washed through a 63-Wm sieve to remove the clay fraction. One to ve tests of benthic foraminifera Cibicidoides wuellerstor and Cibicidoides kullenbergi were picked from the s 150-Wm-size fraction. Prior to stable isotope analyses, foraminiferal tests were ultrasonically cleaned in deionized water and oven-dried. Stable isotopic analyses were carried out using an VG Optima mass spectrometer equipped with a common acid bath at 90 C at the University of California at Santa Cruz. All N 13 C values ( þ 0.05x) and N 18 O values ( þ 0.08x) were calibrated to Vienna Pee Dee be-

6 292 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 Fig. 3. ODP Leg 177 shipboard biostratigraphic datums versus depth (solid circles). Sedimentation rates (labeled) change at 53 and at V142 m composite depth. I interpolate between these datums to construct the time series shown in Fig. 4. The two open squares indicate additional control points based on a comparison of the site 1088 N 13 C record to the high resolution N 13 C record from North Atlantic site 982 (Hodell et al., 2001). These two control points can be used to re ne a portion of the late Miocene age model (3.2. Carbon isotope stratigraphy). lemnite via National Bureau Standard-19 and an in-house standard. 3. Results 3.1. Site 1088 stable isotope records Overall, here is no trend in the site 1088 benthic foraminiferal N 18 O record until the late Pliocene when N 18 O values display the expected increase due to the growth of large-scale Northern Hemisphere glaciation (Fig. 4A). On the ner scale, during the late Miocene there is a distinct trend toward increasing N 18 O values across the Tortonian/Messinian boundary, which indicates some growth of continental ice volume and/or a cooling of the sub-antarctic circum-polar watermass. During the Messinian, N 18 O values decrease abruptly but continue to uctuate, with the exception of a brief interval during the early Pliocene, with increasing amplitude toward the present. This observation is consistent with other deep sea records and indicates increasingly larger glacial to interglacial climate variability as Northern Hemisphere ice sheets began to grow larger during the late Pliocene (e.g., Raymo et al., 1989; Hodell and Venz, 1992). It is well known that late Miocene foraminiferal N 13 C records display a long term decrease, the late Miocene N 13 C shift (e.g., Hodell and Kennett, 1986; Wright et al., 1991; Hodell et al., 2001), and site 1088 benthic foraminiferal N 13 C values clearly follow the global pattern (Fig. 4B). Just after the late Miocene N 13 C minimum N 13 C values increase rapidly, uctuate about a stationary mean until the Miocene/Pliocene boundary, and then display a further, more gradual increase during the early Pliocene. Thereafter, N 13 C values decrease and begin to uctuate with a relatively

7 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ large amplitude during the latest Pliocene and Pleistocene Carbon isotope stratigraphy Assuming that the late Miocene N 13 C shift (e.g., Fig. 4B) was globally synchronous it can be used as a stratigraphic marker (e.g., Keigwin, 1987; Hodell and Kennett, 1986; Wright et al., 1991). The recently published high resolution N 13 C record from subpolar North Atlantic site 982 (Hodell et al., 2001) illustrates that the timing of the late Miocene N 13 C shift can be constrained to between 7.62 Ma (N 13 C maximum) and 6.65 Ma (N 13 C minimum) (e.g., Fig. 5A). Site 982 s astronomically tuned chronology, which uses the most recently developed insolation curves as the tuning target (Hodell et al., 2001), therefore provides the most up-to-date age estimate for this important late Miocene event. The age of the N 13 C maximum (7.62 Ma) and N 13 C minimum (6.65 Ma) can be used to re ne the relevant portion of the site 1088 age model if the site 982 astrochronology is consistent with the Berggren et al. (1995) time scale. To check whether the two time scales are consistent, I compare the ages of the Tortonian/Messinian boundary and the Messinian/Zanclean (early Pliocene) boundary (Table 2). The tuned age of the Tortonian/Messinian boundary is V100 kyr older than that of the Berggren et al. (1995) time scale, but the age of the Messinian/Zanclean boundary is the same. Comparing the N 13 C control points with the initial shipboard biostratigraphy illustrates that the timing of the late Miocene N 13 C minimum is well suited to re ne the latest Miocene change in sedimentation rates at site 1088 (Fig. 3). The age of the site 982 N 13 C maximum, on the other hand, does not agree with the age predicted from adjacent biostratigraphic datums; Fig. 4. Site 1088 benthic foraminiferal (A) oxygen isotope record and (B) carbon isotope record plotted versus age obtained from shipboard biostratigraphic datums reported on the Berggren et al. (1995) time scale (solid circles in Fig. 3).

8 294 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 Fig. 5. Comparison of the site 1088 benthic foraminiferal N 13 C record to records from the (A) North Atlantic, (B) Southern Ocean as a spatial transect, (C) Southern Ocean as a depth transect, (D) Paci c Ocean. Age scale re ects a re ned alignment of the late Miocene N 13 C records from sites 1088, 704, 360, and 289 to the high resolution N 13 C record of site 982 (Tables 3 and 4).

9 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ Table 2 Age model comparison Boundary Berggren et al. (1995) Astrochronology site 982 a Tortonian/Messinian 7.14 Ma Ma Messinian/Zanclean 5.32 Ma 5.32^5.33 Ma a a The age of the Messinian/Zanclean is estimated from Hodell et al. (2001). it is too young (Fig. 3). This discrepancy cannot be explained by the nature of the two di erent age models because the site 982 astrochronology would predict an age of the N 13 C maximum V100 kyr older than on the Berggren et al. (1995) time scale. More likely the discrepancy is caused by the relatively coarse resolution of the biostratigraphic datums at site I use the N 13 C maximum and the N 13 C minimum as additional age control points (Table 3) in order to avoid o sets between the site 1088 and site 982 N 13 C records. I also use the same two N 13 C control points to adjust the age models of other (low resolution) late Miocene N 13 C records (Table 4). Although the revised age models only apply to the interval de ned by the N 13 C shift (7.65 Ma and 6.62 Ma), the temporal resolution of published late Miocene records is too coarse to justify any other adjustments. For simplicity, I make a constant adjustment to the portion of the records outside the two control points (Table 4). No adjustments are necessary to reconcile the ages of site 608; its age model has been updated to the Berggren et al. (1995) time scale by Zachos et al. (2001). It is not necessary to make changes to any of the Plio/Pleistocene records used to reconstruct N 13 C gradients. These records have astronomically tuned age control, although based on di erent insolation curves. The Berggren et al. (1995) time scale incorporates astronomically tuned age control points derived from Plio/Pleistocene oxygen isotope stratigraphy of Shackleton et al. (1990) and Hilgren (1991). Discrepancies among the Plio/Pleistocene records and between the Plio/ Pleistocene records and site 1088 may arise from the use of di erent insolation curves as tuning targets. These di erences are small during the Pleistocene and increase to V20^40 kyr by the earliest Pliocene (e.g., Tiedemann and Franz, 1997). Because the temporal resolution of the site 1088 record is V50 kyr during the Plio/Pleistocene, age model discrepancies of V40 kyr are not important Comparison of benthic foraminiferal N 13 C records Applying the N 13 C-based age model achieves excellent agreement among N 13 C records across the Tortonian/Messinian boundary (Fig. 5). Site 1088 N 13 C values parallel those of the North Atlantic although the absolute magnitude of the N 13 C shift is larger at site 1088 (Fig. 5A). Southern Ocean and Paci c N 13 C records overlap during the late Miocene N 13 C shift, and N 13 C gradients develop among the various sites shortly after the N 13 C minimum at 6.6 Ma (Fig. 5B^D). A rst-order comparison of the records also illustrates that within the North Atlantic source water di erentiation occurred at V6.6 Ma. A Table 3 Late Miocene portion of the site 1088 age model Event a Depth Age Time scale (m composite depth) (Ma) LO Discoaster quinqueramus a Berggren et al. (1995) N 13 C minimum b Site 982 (astrochronology) FO Amaurolithus spp. a Berggren et al. (1995) N 13 C maximum b Site 982 (astrochronology) FO Discoaster quinqueramus a Berggren et al. (1995) a Late Miocene biostratigraphic control points with corresponding ages based on Berggren et al. (1995), the entire age model can be found in Proceedings of the ODP Initial Reports 177; LO stands for last occurrence; FO stands for rst occurrence. b Additional control points from carbon isotope stratigraphy, ages re ect astrochronology at site 982.

10 296 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 Table 4 Summary of age adjustments (Myr) based on late Miocene N 13 C records Site N 13 C minimum N 13 C maximum Above N 13 C min Below N 13 C max 360 a a N/A b a a Wright et al. (1991). b Mueller et al. (1991). N 13 C gradient developed at V6.6 Ma between the two North Atlantic sites, with the deeper and more southern site 926 recording the higher N 13 C values (Fig. 5A). The N 13 C gradient suggests either a chemical di erentiation among individual source waters, or an increase in the relative contribution of a 13 C-enriched source watermass. Because it is not possible to tell what blend of watermasses contributed to the one at 1088, the e ect of source water di erentiation will be taken into consideration in the reconstruction of the relative ux of NADW to the Atlantic sector of the Southern Ocean during the late Miocene. The uniformity of N 13 C values among all Southern Ocean sites suggests that they were bathed by the same watermass during the late Miocene, and this watermass also extended into the deep western equatorial Paci c. Site 1088 N 13 C values overlap with those at nearby sub-antarctic site 704 (from Wright et al., 1991) (Fig. 5B), the deeper and more northern South Atlantic site 360 (Fig. 5C), and with values at site 289 (Fig. 5D). Site 1088 N 13 C values tend to be lower than those late Miocene site 704 values published by Mueller et al. (1991) (Fig. 5B), but in light of the good agreement between site 1088 and site 704 of Wright et al. (1991), this discrepancy may not have environmental signi cance. During the Pliocene and Pleistocene it appears that small changes in the geometry of the mixing front between NADW and CDW altered the N 13 C gradient between sites 1088 and 704 (Fig. 5B). During the late Miocene, the uniformity of N 13 C values among all Southern Ocean sites suggests that they were bathed by the same watermass, and that this watermass also extended into the Paci c (Fig. 5D). Unfortunately, there is no single continuous isotope record that illustrates how mean ocean N 13 C values decrease after this point until the beginning of the site 849 record at the Miocene/Pliocene boundary. Nearby eastern equatorial Paci c Ocean site 846 N 13 C values are clearly lower that those at site 849 during the entire Pliocene. 4. Discussion 4.1. Oceanic N 13 C gradients and deep water circulation To discuss long-term changes in the carbon isotope gradients and to calculate the relative proportion of NADW at site 1088, I have rst smoothed the records using a 10% Gaussian distribution and then interpolated values at 0.5-Myr intervals (Fig. 6). Smoothing highlights the presence of a N 13 C gradient between the two North Atlantic sites 982 and 926 beginning at V6.6 Ma (Fig. 6A). Smoothing results in a loss of detail at site 608 during the late Miocene N 13 C shift. But for the interval prior to this event the smoothed record can be used as one of the North Atlantic endmembers. The smoothed records also highlight the N 13 C di erence between the two eastern equatorial Paci c sites 849 and 846 (Fig. 6A). Site 846 N 13 C values are lower on average by 0.2x, which casts doubt on the value of this site for constraining mean Paci c N 13 C values during the latest Miocene to earliest Pliocene climate transition. The most conservative approach to estimate mean ocean N 13 C values is to interpolate between

11 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ site 849 and site 289 (Fig. 6A, dotted line). In this way, mean global N 13 C values may be overestimated, which would underestimate the calculated % NADW (see Eq. 1 below). This approach implies that mean ocean N 13 C values decreased between 6.5 and 5.0 Ma, a trend that is largely supported by the N 13 C records from the other sites (Fig. 6A). To obtain an estimate of North Atlantic source water N 13 C variability (at least for the late Miocene), I splice together the North Atlantic records in two ways. One, I combine the records of sites 608 and 926 with a short portion of the site 982 record between 7.5 and 7 Ma to obtain a record of maximum source water N 13 C values. Two, I combine the records of site 608 (until 8.5 Ma) and all of site 982 to obtain a record of minimum source water N 13 C values. Both curves are used to calculate the Atlantic to Paci c N 13 C gradient (Fig. 6B) and the estimated proportion of NADW to site 1088 (Fig. 6D). Accordingly, the modern N 13 C gradient between the North Atlantic and Paci c Ocean developed during the latest Miocene between 8 and 7 Ma and remained similar to, or perhaps slightly larger than, the modern gradient through the Pliocene. For comparison, the modern sub-antarctic Southern Ocean (site 1088) to Paci c N 13 C gradient developed 1Myr later at V6 Ma, and it remained higher than the modern gradient throughout the Pliocene (Fig. 6C). Following the methodology of Oppo and Fairbanks (1987), Hodell (1993), and Wright et al. (1991), I estimate the % NADW that bathed site 1088 using the following equation: %NADW ¼ðN 13 C N 13 C Pacific = N 13 C North Atlantic 3N 13 C Pacific ÞU100 ð1þ Based on modern N 13 Cof4CO 2 in the sub-antarctic South Atlantic, the North Atlantic and Paci c (Table 1), the watermass at site 1088 consists of V33% NADW (Fig. 6D). This formula assumes that the sub-antarctic Southern Ocean re- ects an intermediary deep water location downstream of the North Atlantic source and upstream of the Paci c Ocean mean. As such, it does not apply to times when the Southern Ocean N 13 C values were higher than those in the North Atlantic (e.g., between 11 and 10 Ma, Fig. 6A striped region), nor to times without a signi cant di erence between North Atlantic source water N 13 C values and the global mean represented by the Paci c Ocean (e.g., prior to V7 Ma in this compilation). The % NADW calculation does, however, con- rm the absence of a nutrient-depleted watermass at site 1088 prior to 6.6 Ma (Fig. 6D). And it highlights the rather sudden change in deep water distribution patterns brought about by an increase in the relative ux of NADW to site 1088 between 6.6 and 6.0 Ma. By 6.0 Ma, the relative ux was comparable to today, and it was even higher during the early Pliocene, regardless of the choice of the North Atlantic endmember. The % NADW increase during the Pleistocene is meaningless, it is an artefact of a low North Atlantic to Paci c N 13 C gradient (Fig. 6B). The new comparison of carbon isotope records presented in this study agrees well with the rst synthesis of the evolution of late Miocene N 13 C gradients of Wright et al. (1991), who propose a late Miocene onset of modern-type deep-watermass production in the North Atlantic. Results from site 1088 stress that prior to the late Miocene N 13 C minimum (6.6 Ma), NADW did not extend into the Atlantic sector of the Southern Ocean. Wright et al. also illustrate that the relative proportion of NADW to the sub-antarctic Southern Ocean increased signi cantly during the late Miocene. Site 1088 N 13 C records show that once the relative ux of NADW reached the sub-antarctic Southern Ocean, it quickly increased to proportions similar to today by 6 Ma. The relative ux increased further 1Myr later and remained enhanced with respect to the modern value through the early Pliocene. These observations lead me to suspect that if enhanced thermohaline overturn is to explain early Pliocene climatic warmth, the initial trigger must be sought in the late Miocene Late Miocene ice volume changes? Perhaps one of the most prominent paleoceano-

12 298 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 Fig. 6. (A) Smoothed (using a 10% Gaussian distribution) benthic foraminiferal N 13 C records. The dotted line is to bridge the gap between site 849 and 289. (B) The North Atlantic to Paci c N 13 C gradient is calculated with sites 608 (11^7.5 Ma), 982 (7.5^7 Ma) and 926 (7^5 Ma) to obtain a record of maximum North Atlantic source water N 13 C values (open circles). The North Atlantic to Paci c N 13 C gradient is also calculated from sites 608 (11^8.5 Ma) and site 982 (8.5^5 Ma) to obtain a record of minimum source water N 13 C values (solid circles). (C) The site 1088 to Paci c gradient, and (D) % NADW using Eq. 1 (see text). All gray horizontal bars re ect the modern conditions using the regional N 13 Cof4CO 2 (Ostlund et al., 1987, Table 1).

13 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ graphic events of the late Miocene was the repeated isolation and desiccation of the Mediterranean, the salinity crisis. Restriction of surface water in ow from the Atlantic may have been due to tectonic events and/or a glacioeustatic sea level lowering (for a review see Hodell et al., 1994). Hodell et al. (1994) in particular, discuss a late Miocene (Tortonian/Messinian, 7.3 to 6.9 Ma) increase in ice-volume apparent as a two-step increase in benthic foraminiferal N 18 O values at a site located on the Atlantic coast of Morocco (the Sale Briqueterie drill core, Fig. 2). Hodell et al. (2001) now question the existence of a long-term increase in ice-volume during this time interval because benthic foraminiferal N 18 O values at North Atlantic site 982 remain relatively constant and do not accommodate an increase in ice volume (e.g., Fig. 7). Rather, Hodell et al. (2001) argue, the Sale N 18 O record re ects cooling and/ or increased salinity caused by a change in intermediate water circulation in response to the tectonic isolation of the Mediterranean. Site 1088 N 18 O values parallel the N 18 O record from Sale across the Tortonian/Messinian boundary, although the absolute N 18 O values are higher at the Southern Ocean site (Fig. 7). Thus the late Miocene N 18 O increase observed at Sale must be of extra-regional importance. The simplest explanation is that the N 18 O increase re ects an increase in continental ice volume. However, this interpretation implies that intermediate water temperature changes in the subpolar North Atlantic have perfectly masked the global signal, which seems serendipitous. It is perhaps more likely that the N 18 O increase at site 1088 re ects a cooling of upper deep waters sinking in the Southern Ocean. If cooling also occurred in Antarctic intermediate waters, which ow north toward the subtropical Atlantic in the modern ocean (e.g., Pickard and Emery, 1990), then the N 18 O increase at Sale can be explained by the dominance of Antarctic intermediate water. This interpretation would imply that the watermass at this site was not in uenced by mixing with water owing out of the Mediterranean, which would be consistent with a constricted gateway through the Ri an Straits at about this time (e.g., Hodell et al., 2001 and references therein). However, this interpretation leaves the conundrum of the Sale N 13 C record being identical to that of site 982 (Hodell et al., 2001), while Southern Ocean deep water records have distinctly lower N 13 C values than the North Atlan- Fig. 7. Comparison of the site 1088 N 18 O record with the N 18 O records from site 982 (North Atlantic) and the Sale Briqueterie section from the Atlantic coast of Morocco (Hodell et al., 2001) between 5 Ma and 9 Ma. Note the di erence in scale between the site 1088 and the Sale records.

14 300 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 tic at this time (e.g., Figs. 5 and 6). But Southern Ocean intermediate watermasses can have relatively high N 13 C values because of vigorous air^ sea interactions (Charles and Fairbanks, 1990), and unexpectedly high benthic foraminiferal N 13 C values at southwestern subtropical Paci c site 588 (Fig. 2, Table 1) during the Miocene have been explained in this manner by Wright et al. (1992). A quick comparison of the N 13 C records of sites 1088, 982, and 588 to the Sale record con rms that the Sale N 13 C values, although very similar to those recorded in the North Atlantic, also overlap with those recorded at site 588 (Fig. 8). The comparison is consistent with the notion that the Sale record represents an intermediate water mass signal. Therefore, I concur with Hodell et al. (2001) that the late Miocene N 18 O increase re ects a temperature signal rather than ice volume growth, which supports those studies that have called for a tectonic control on a late Miocene eustatic sea level drop The role of tectonics A number of other geologic events occurred during the Neogene that a ected boundary conditions and likely changed deep water circulation and climate patterns. The sill depth of the CAS shoaled from 1000 m during the middle Miocene to near breaching of the sea surface during the early Pliocene redirecting the ow of warm and salty water to the north within the Atlantic (e.g., Keigwin, 1982; Keller et al., 1989; Duque-Caro, 1990; Haug and Tiedemann, 1998). A decrease in mantle plume activity on the Greenland^Scotland Ridge may have lowered the sill depth between the North Atlantic and Norwegian^Greenland Seas thereby allowing over ow of Arctic waters into the North Atlantic stimulating NADW formation (Wright and Miller, 1996). And prior to the northward drift of New Guinea to its present location between 5 and 3 Ma, the Indonesian Through ow was deeper, which would imply weaker zonal sea surface temperature gradients, a weaker Walker circulation, but enhanced heat transport out of the tropics relative to today (Cane and Molnar, 2001). Of these tectonic events the closure of the CAS and the sinking of the Greenland^Scotland Ridge provide the more direct control on Atlantic deep water distribution patterns. For example, the timing of the late Miocene onset of NADW presence in the sub-antarctic Southern Ocean (6.0 Ma) as inferred from the diminished N 13 C gradient be- Fig. 8. Comparison of sites 1088, 982, and Sale N 13 C records to the N 13 C record from the intermediate water depth, southwest Paci c site 588 (Wright et al., 1992 with ages updated to Berggren et al. (1995) by Zachos et al. (2001)). The site 588 N 13 C values are high, they overlap with those in the North Atlantic. Relatively high N 13 C values in intermediate water masses sinking in the Southern Ocean can be attributed to vigorous air^sea exchange (Charles and Fairbanks, 1990), which can explain high N 13 C values at site 588 during the Miocene (Wright et al., 1992).

15 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ tween the North Atlantic and Southern Ocean has been attributed to the sinking of the Greenland^ Scotland Ridge (Wright and Miller, 1996). Based on the results from site 1088, the contribution of northern source waters to the sub-antarctic South Atlantic during this pulse was equivalent to that observed during the present-day interglacial interval. Sinking of deep waters in the North Atlantic may have brought about an enhanced delivery of warmer subtropical surface waters, which, after V1Myr, may have lead to early Pliocene climatic warmth observed at high northern latitudes. Although the depth of the Greenland^Scotland Ridge may have controlled the amount of water spilling into the North Atlantic, the CAS may have had a relatively larger impact on North Atlantic hydrography and climate by strengthening the western boundary current and associated poleward transport of warm and saline waters. Between 6.6 Ma and 6.0 Ma, an emerging CAS may have su ciently limited thermocline water exchange between the Atlantic and Paci c oceans to stimulate thermohaline overturn in the Atlantic. Near breaching of the sea surface by the emerging seaway need not be invoked, and this model is not necessarily in con ict with the initiation of upper NADW formation in the Labrador Sea at 4.6 Ma, or the evolution of surface water salinity gradients between the Atlantic and Paci c after V4 Ma(Haug and Tiedemann, 1998; Keigwin, 1982; Haug et al., 2001). I speculate that the development of the N 13 C gradient between the subpolar and tropical Atlantic at 6.6 Ma is linked to deep water formation in di erent source regions analogous to Labrador Sea Water and Nordic Sea Water in the modern ocean (e.g., Schmitz and McCartney, 1993). Regardless, relatively strong thermohaline overturn during the latest Miocene suggests that an emerging CAS may have in uenced NADW formation V1.5 Ma earlier than previously thought. 5. Conclusion Site 1088 provides a new opportunity to monitor the evolution of late Miocene through early Pliocene deep water circulation and thermohaline overturn. To this end I have compiled published N 13 C records from the North Atlantic, Southern Ocean and Paci c. Placed in this context, benthic foraminiferal N 13 C records from site 1088 provide evidence for the presence of a nutrient-depleted watermass proportional to modern values at 6.0 Ma and in excess of the modern proportion during the early Pliocene. Site 1088 benthic foraminiferal N 18 O values display a two-step increase between V7.4 Ma and 6.9 Ma, a trend that parallels the N 18 O record of at a site on the Atlantic coast of Morocco. Cooling of Southern Ocean upper circum-polar and intermediate watermasses best explains this observation. I conclude that relatively strong thermohaline overturn at rates comparable to the present day interglacial interval during the latest Miocene may have provided the initial conditions for early Pliocene climatic warmth. The impact of an emerging CAS on Atlantic^Paci c Ocean upper water exchange may have been felt in the North Atlantic beginning in the latest Miocene between 6.6 and 6.0 Ma, which would be V1.5 Myr earlier than previously thought. Acknowledgements I thank Christina Ravelo for stable isotope analyses and Bill Chaisson for helpful comments on an earlier version of the manuscript. Peggy Delaney and Tom Crowley provided thoughtful and constructive reviews that substantially improved this paper. This research used samples provided by the ODP. ODP is sponsored by the U. S. National Science Foundation (NSF) and participating countries under the management of Joint Oceanographic Institutions (JOI), Inc. This research was funded by NSF Grant OCE to A.C. Ravelo and by an NSF postdoctoral fellowship to K.B. ODP Leg 177 science was supported by a JOI/USSAC Grant 177-F (to J. Zachos) and is currently being supported by NSF Grant OCE (to K.B.).

16 302 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 AppendixI Summary of site 1088 stable isotope data Site Core Sec Top MCD Age biostratigraphy Age N 13 C stratigraphy N 13 C N 18 O (Ma) (Ma) (x) (x) 1088B 1H not applicable B 1H not applicable B 1H not applicable B 1H not applicable B 1H not applicable B 1H not applicable B 1H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 2H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 3H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 4H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable

17 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^ Appendix I (Continued). Site Core Sec Top MCD Age biostratigraphy Age N 13 C stratigraphy N 13 C N 18 O (Ma) (Ma) (x) (x) 1088B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 5H not applicable B 6H not applicable B 6H not applicable B 6H not applicable B 6H not applicable B 6H not applicable B 6H not applicable B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 6H B 7H

18 304 K. Billups / Palaeogeography, Palaeoclimatology, Palaeoecology 185 (2002) 287^307 Appendix I (Continued). Site Core Sec Top MCD Age biostratigraphy Age N 13 C stratigraphy N 13 C N 18 O (Ma) (Ma) (x) (x) 1088B 7H B 7H B 7H B 7H B 7H B 7H B 7H B 7H B 7H B 7H B 7H B 8H B 8H B 8H B 8H B 8H B 9H B 9H B 9H B 9H B 9H B 10H B 10H B 10H B 10H B 11H B 11H B 11H B 11H B 11H not applicable B 12H not applicable B 12H not applicable B 12H not applicable B 12H not applicable B 13H not applicable B 13H not applicable B 13H not applicable B 13H not applicable B 14H not applicable B 14H not applicable C 2H not applicable C 2H not applicable C 3H not applicable C 3H not applicable C 3H not applicable C 3H not applicable C 3H not applicable C 3H not applicable C 3H not applicable C 4H not applicable C 4H not applicable C 4H not applicable C 4H not applicable