With these primary motivations in mind, I will review what we do know about changes in the deep ocean circulation from paleo‐tracer data. The paper is not a review of mechanisms for glacial to interglacial CO 2 change [ Sigman and Boyle , 2000 ; Sigman et al ., 2010 ] or a review of proposed mechanisms for rapid climate change. Instead, I hope to show what data are available from the interior of the ocean and how we might quantify them, with an emphasis on the last glacial maximum (LGM). The paper does not start from the same place as an inverse approach with the question “what was the past circulation, and can I establish within error that it was different than today” [ LeGrand and Wunsch , 1995 ]. Instead, I think from the above discussion, it is clear that the deep sea has a crucial role to play in past climate changes, and, accordingly, it is essential to review what it is we can extract from the tracer information to better understand this role. The question of what drives changes in the deep ocean circulation is a fascinating and open research topic in modern oceanography and one where paleoclimate studies could help provide an answer in the future. However, the motivation for understanding the rate of deep ocean ventilation is not necessarily to reconstruct the full circulation. For example, the pCO 2 changes require an understanding of how ocean carbon and alkalinity are transported and diffused in the fluid, just like any tracer affected by circulation. To this end, I will also review a relatively new method of using tracer balances to constrain important parameters of the past circulation after summarizing what we know about the tracer distributions themselves. At the end, I will propose a new idea for how the LGM circulation pattern might have come about and what role it played in glacial to interglacial climate change.

Most, but not all [ Gildor and Tziperman , 2003 ; Timmerman et al ., 2003 ], of the current theories for why there are ice ages and/or rapid climate changes within glacial stages rely on variations in the overturning rate of the deep ocean, though the exact mechanisms at work are not yet understood and they tend to focus on NADW on or off modes. From the methane synchronization in ice cores [ Blunier and Brook , 2001 ], it is clear that temperature increases in Antarctica associated with the North Atlantic Heinrich Events (so‐called “A events”) preceded the rapid warmings seen in Greenland by 1–3000 years. A time difference this large is not the same mechanism as the classic “see‐saw” described above and must originate from a piece of the climate system with large inertia. Only the cryosphere itself and the deep ocean are large enough reservoirs to cause this very long time constant [ Imbrie et al ., 1992 ].

Broecker introduced the notion of a “salt‐oscillator” in the North Atlantic to explain both the rapid shifts and the glacial‐interglacial differences seen in paleo records. He hypothesized that North Atlantic Deep Water (NADW) was close to an instability threshold with respect to salinity. In this early idea, the opposing forces of fresh water export out of the Atlantic basin over Central America via the atmosphere and salt export via NADW transport out of the basin lead to oscillations in the climate system. While NADW formation is active, its associated cooling of the ocean warms the atmosphere and greatly alters the heat budget of the North Atlantic. Coupling this fresh water balance to the observation of phasing between Greenland and Antarctic ice cores resulted in a modified salt‐oscillator idea, the “bipolar seesaw” [ Broecker , 1998 ; Stocker , 1998 ]. In its simplest form, this idea relies on radiation that has warmed the southern Atlantic to cross the equator and be released to the atmosphere as NADW sinks. In this way, a warm mode in the North Atlantic leads to a cooling of the Southern Hemisphere, an idea first put forward by Crowley [ Crowley , 1992 ].

But changing pCO 2 is the not the only motivation for examining changes in the deep circulation through glacial cycles. Over the last 20 years, studies of past climate have undergone a profound change in emphasis from a glacial‐interglacial approach to the study of rapid shifts in the system. By cross correlating isotopic records of atmospheric temperature in ice cores drilled at Greenland's summit, we now recognize the global prevalence of large amplitude climate shifts on very short time scales [ Dansgaard et al ., 1993 ; GRIP , 1993 ; Grootes et al ., 1993 ; Steffensen et al ., 2008 ]. What were formerly considered uncorrelated wiggles in the Greenland archives are now well established to be real climate signal [ Alley et al ., 1995 ]. While Milankovitch cycles are the fundamental pacemakers of climate at 20, 40, and 100 thousand year periods, research emphasis has switched to the observations of much more rapid fluctuations at the centennial to decadal timescale. Twenty‐two separate interstadial events—brief returns to warmer climate—have been recognized in the Greenland summit ice core records. Large volumes of ice rafted debris are recognized roughly every 7000 years in North Atlantic sediments as armadas of icebergs were shed from continental ice sheets in so‐called “Heinrich Events” [ Heinrich , 1988 ; Hemming , 2004 ; Ruddiman et al ., 1980 ]. And the data are not limited to the North Atlantic, with Greenland‐like variation in the Santa Barbara Basin [ Behl and Kennett , 1996 ], the Southern Ocean [ Charles et al ., 1996 ], the Indian Ocean [ Schulz et al ., 1998 ], and the Chinese Monsoon [ Wang et al ., 2001 ]. Collectively these observations represent a challenge for paleoclimate studies. One of the largest outstanding questions in our field is the issue of what mechanisms cause these large and rapid shifts (Heinrich, deglacial, and interstadial) and, more specifically, what is the role of deep ocean circulation in their realization [ Boyle , 2000 ]?

Sensitivity of pCOto changes in biological production and deep‐ocean overturning rate as calculated in a multibox model [after]. It is not possible to move from preindustrial values of ~280 ppmV to LGM values of 180–200 ppmV without changing both the biology of the Southern Ocean (bold green arrow) and the deep ocean overturning strength (bold blue arrows). While there are several caveats to these model results discussed in the text and examined in detail in. [], results for a single productivity constant and varying the Southern Ocean overturning show a simple dependence of pCOon the integrated amount of preformed phosphate ion in the deep ocean [after.,]. A full description of the model shown here will appear in a future paper.

The quantitative role that the deep ocean circulation plays in the lowering of glacial pCO 2 has been an important area of research since the mid‐1980s. With the publication of the so‐called “Harvardton Bears” box models [ Knox and McElroy , 1984 ; Sarmiento and Toggweiler , 1984 ; Siegenthaler and Wenk , 1984 ], the joint dependence of the system on deep ocean ventilation rates and high latitude biological productivity became clear (Figure 2 ). These simple models pointed out that the Southern Ocean is the atmospheric window to the large volumes of the deep ocean; thus, the processes in this region must mediate the storage of carbon in the deep ocean (and its removal from the atmosphere during the glacial periods). Toggweiler [ Toggweiler , 1999 ] expanded on this work by building a series of box models that moved the deep ocean upwelling from the low latitudes to the high latitudes (four boxes), emphasized the importance of intermediate waters (six boxes), and divided the deep box into two regions ventilated from the south and the north separately (seven boxes) (Figure 2 a). The paleo data were better explained with each of these sequential enhancements to the simplest box models, but the basic implication that both productivity changes and deep ocean circulation rate changes are necessary to produce significant atmospheric CO 2 change is clearly consistent across models, regardless of configuration Figure 2 b.

Ice core records of (a) pCOand (b) δD of the ice along with summer solar insolation at 65°N (c). Marine Isotope Stages are labeled in Figure 1 A. and Antarctic “A‐events” are labeled in Figure 1 B. The “sawtooth” character of the ice core temperature proxy δD has a strong 100 ka signal that is not easily seen in the insolation forcing curve. Carbon dioxide can work as an amplifier of insolation forcing during terminations and might provide the “kick” for rapid deglacial warmings. If COmatters in the glacial‐interglacial climate system, then the deep ocean nutrient and circulation dynamics must play a role in these climate changes.

The role of the deep ocean in climate change has been a central theme of paleoceanography from the early days of the field. As the locus of nearly all the mass, thermal inertia, and carbon in the ocean‐atmosphere system, the deep ocean's behavior is an important parameter in understanding past climates—whether considering either the steady states or the propensity for rapid shifts. There are two observations from ice cores that particularly motivate the investigation of past deep ocean circulation in the context of past climate change. One of these is the carbon cycle changes demanded by the past atmospheric CO 2 data from Antarctica, which show significant variability across glacial to interglacial cycles [ Barnola et al ., 1987 ; Neftel et al ., 1985 ]. In these records, air temperature and CO 2 closely follow each other over the last 850,000 years [ EPICA , 2004 ]. However, at glacial terminations, there is a strong nonlinearity in each of these variables relative to the proposed insolation forcing during boreal summers (Figure 1 ). The classic “sawtooth” structure of glacial cycles [ Broecker and van Donk , 1970 ] is punctuated by rapid increases in sea level, temperature, and CO 2 that require a nonlinear response of the ocean‐atmosphere system to solar forcing. While CO 2 does not seem to initiate deglaciations [ Caillon et al ., 2003 ], its role as a greenhouse gas is one of the most important amplifiers to temperature change during these climate transitions. Because of the acid dissociation constants of the carbonate system and the alkalinity content of the ocean—today there is 60 times more carbon in the deep ocean than in the atmosphere—any past changes in atmospheric CO 2 must involve the carbon dynamics of the deep ocean reservoir.

These sorts of simple model ages have been replaced in modern oceanography with large data sets and sophisticated Ocean General Circulation Models (OGCMs). Both Transit Time Distributions (TTDs) [ DeVries and Primeau , 2011 ; Hall and Haine , 2002 ] and inverse models attempt to use dynamical constraints on the circulation as well as tracer information to extract properties of the deep ocean. Following the pioneering work of Waugh et al. in the stratosphere [ Waugh and Hall , 2002 ], the TTD approach gives rise to the notion of tracer age spectra. A single parcel contains water that arrived at that sampling site from many different pathways and many different transit times. The two end‐member assumption in Figure 4 is a useful simplification, but studies of the full modern ocean suggest at least six [ Johnson , 2008 ] if not many more [ Gebbie and Huybers , 2010 ] end‐members. Inverse models have produced flux estimates, with error bars, for deep water masses [ Ganachaud and Wunsch , 2000 ] and constrained carbon export fluxes [ Schlitzer , 2002 ]. The inverse approach will ultimately provide important information about the paleo deep ocean, but at the moment the tracer data set is limited enough that other approaches may produce more insight (see below).

With a known remineralization rate of organic matter, and its typical phosphorous content, the deviation from conservative mixing seen in Figure 4 b could be used to calculate the mean time for deep water to arrive at this site of 50/50 mixing, and thus derive circulation information from chemical tracers. Unfortunately, the integrated flux of organic matter out of the surface ocean into the deep is a difficult number to constrain in the modern, let alone the past. Radiocarbon, on the other hand, contains this rate information with no other knowledge of the system. Figure 4 c plots the Δ 14 C distribution versus potential temperature along the same isopycnal as in Figures 4 a and 4b. In the Atlantic, there is about a 100‰ range in the Δ 14 C of DIC. High salinity North Atlantic Deep Water (NADW) forms with a value of ~ −65‰ while fresher Antarctic Bottom Water (AABW) begins spreading at ~ −165‰ [ Broecker et al ., 1995 ]. These numbers are somewhat uncertain because of the presence of nuclear bomb produced radiocarbon in all ocean surface waters today and the somewhat arbitrary nature of picking end‐member values. Figure 4 c shows that most of this 100‰ spread is due to mixing of deep water masses. Points below the pure end‐member mixing line represent loss of 14 C due to in situ decay. The largest offset from mixing is about 30‰, corresponding to 240 14 C years of isolation from the atmosphere (for this choice of two end‐members). This is only a model age of the actual mass flux of water, but it is useful in the context of other tracers. Because we know the rule for nonconservative behavior of Δ 14 C a priori—that it follows first‐order radioactive decay—this tracer is much more useful than PO 4 at informing the circulation rate of the deep ocean. Radiocarbon tells us that today the Atlantic is a young ocean.

Tracer versus tracer plots for the deep Atlantic isopycnal σ θ = 37.0. (a) As both potential temperature and salinity are conservative tracers, the linear relationship implies that there are only two end‐members mixing along this isopycnals at these depths. A slight offset between the “true” AABW end‐member (blue circle, picked by the author) and the value at the end of the line is due to the deep mixing choke point at the Vema Passage in the South Atlantic (black circle). The NADW end‐member (red circle) is at the opposite end of each line. (b) Potential temperature versus [PO 4 ]. Extra phosphate above the conservative mixing line comes from remineralization of raining organic matter along this isopycnal. (c) Curvature from the nonconservative Δ 14 C versus conservative potential temperature plot arises from radioactive decay of 14 C along the deep isopycnal trajectory. While there is a large range of Δ 14 C in the modern deep Atlantic, most of this variance is due to water mass mixing and NOT due to older ventilation ages. At a 50/50 mixture of NADW and AABW/Vema Passage end‐members, the Δ 14 C is ~30‰ lower than the conservative mixing line, implying about 250 years of in situ aging. Deep waters in the Western Atlantic are relatively young.

Because the effects of temperature and salinity on seawater density oppose each other in the two dominant deep ocean water masses, there is a class of isopycnals in the Atlantic that outcrop in both the northern and southern source water formation regions (Figure 3 d). At the surface, NADW is denser than AABW due to its higher salinity. But because of the pressure dependence of the thermal expansion coefficient, the warmth of NADW makes it less dense than AABW at depth. This feature of the thermodynamics of seawater results in the fact that some isopycnal surfaces can be characterized by both nearly pure NADW and nearly pure AABW. This is true even when comparing neutral density surfaces instead of isopycnals [ Jackett and McDougall , 1997 ]. Gradients in these conservative tracers along a single isopycnal give rise to a very useful diagnostic tool in modern hydrography. Plots of a conservative tracer versus a chemically active, or nonconservative, tracer like phosphate, along these dual outcropping isopycnals reveal the degree to which [PO 4 ] changes are due to chemical reactions or due to simple mixing of water masses with different initial [PO 4 ] (Figure 4 b). NADW and AABW leave the surface with very different preformed phosphate values due to the large differences in the efficiency of the biological pump in the surface waters that make up each end‐member. As these water masses mix together, they gain PO 4 due to the remineralization of raining organic matter from above. This process leads to deviations from a conservative mixing line in Figure 4 b. At a 50/50 mixture of NADW and AABW along the 1032.1 kg/m 3 isopycnal, the chemical conversion of solid organic matter to dissolved phosphate increases the [PO 4 ] by ~15%.

Sections of conservative tracers in the Western Atlantic from the GLODAP database. (a) Potential temperature, (b) salinity, (c) potential density relative to the surface pressure (σ θ ), and (D.) a calculation how σ θ evolves for end‐member T and S values of NADW and AABW. In the deep ocean, there is clear mixing between warm/salty NADW and cold/fresh AABW. The largest densities relative to the surface pressure are associated with NADW and are NOT found at the bottom of the ocean in the South Atlantic. At the surface, NADW is denser than AABW. At depth, the warmth of NADW causes it to be less dense than AABW such that this southern sourced water mass actually fills most of the abyssal ocean and much of the western basin of the South Atlantic. The change in sign of the relative densities of NADW and AABW leads to along isopycnals mixing of tracer in the modern deep Atlantic for a large volume of water.

There are many good reviews on the modern distribution of deep water masses and their dynamics [ Reid , 1981 ; Rintoul et al ., 2001 ; Talley et al ., 2011 ; Warren , 1981 ], but it is instructive to consider how modern tracer information can be used in a simple way to infer deep ocean processes. Many modern studies diagnose the fluxes of heat and moisture at the ocean's surface to constrain the volumes and rates of deep water mass circulation [ Speer and Tziperman , 1992 ]. As temperature and salinity are conservative tracers in the ocean interior, their rates of change at water mass outcrop regions are useful diagnostics. In paleoceanography, on the other hand, we rarely have access to these surface fluxes, but we do have information about tracers in the ocean interior. To this end, a tracer view of deep circulation begins with the distribution of salinity and potential temperature (theta). As the off axis geothermal heat flux is ~100 mW/m 2 , the assumption of conservative behavior is not strictly true for potential temperature [ Emile‐Geay and Madec , 2009 ; Joyce et al ., 1986 ], but the changes induced by bottom heating in the modern ocean are small. Sections of temperature and salinity in the Atlantic Ocean (Figures 3 a and 3 b) show the classic interleaving of warm/salty North Atlantic Deep Water (NADW) and cold/fresh Antarctic Bottom Water (AABW). NADW appears to “split” the southern source waters into AABW and its slightly warmer and fresher counterpart Antarctic Intermediate Water (AAIW).

3 History

3.1 Biogeochemistry‐Based Tracers One of the most robust results of deep water paleoceanography is that the mixing ratio of northern and southern source waters in the Atlantic varied on glacial‐interglacial timescales. Today this result is based on a relatively large number of chemical tracer data but the first evidence came from benthic foraminifera faunal abundances. Streeter and Shackleton [Streeter and Shackleton, 1979] noticed that the percentage of Uvigerina peregrina relative to other benthic species in core V29‐179 (44.7 N, 24.5 W, 3331 m) varied with the δ18O isotope stages. Core top assemblages show that this species prefers waters originating from the Southern Ocean. Combining the down core and modern observations, Streeter and Shackleton concluded that “Uvigerina does not like NADW” and that during interglacials the retreat of Uvigerina in the Atlantic implies a dominance of waters from the north, but during glacial periods the Uvigerina compatible AABW was more dominant than it is today. Chemical tracers sensitive to the relative mixtures of Antarctic and Northern sourced waters placed this faunal‐based idea in firmer context. In the early 1980s, two papers laid the foundation for much of what we think of as paleo‐tracer oceanography today. Boyle and Keigwin [Boyle and Keigwin, 1982] measured Cd/Ca ratios and δ13C in benthic foraminifera from the deep (~3200 m) North Atlantic and found large differences between glacial and interglacial times. Like phosphate in the modern ocean, the value of these tracers at any one location in the deep Atlantic is strongly influenced by the relative mixing ratios of northern (low Cd/Ca and high δ13C) and southern (high Cd/Ca and low δ13C) waters. Boyle and Keigwin concluded that the intensity of water mass production from the north decreased by about a factor of two as compared to deep water production from the south during severe glaciations. At nearly the same time, Curry and Lohmann [Curry and Lohmann, 1983] analyzed benthic δ13C data from a depth transect of cores in the Eastern Equatorial Atlantic. These authors recognized that lower δ13C values at depth during the LGM corresponded to a lower flux of [O 2 ] rich bottom waters into the Eastern Basin, as well as an increase in the overlying rain of organic matter. What these two studies have in common is a recognition that the sedimentary record could be used to reconstruct past chemical tracer information. In the same way, the modern distribution of chemical species in the deep ocean reflects ocean circulation and biogeochemistry (see above); past ocean tracer distributions can be interpreted in this framework. These studies also share the assumption that Cd/Ca and δ13C in benthic foraminifera are tracers of past nutrient content of the deep waters. This interpretation is driven by the correlation in the modern ocean of low Cd/Ca and high δ13C with low [PO 4 ] and vice versa [Boyle, 1988; Kroopnick, 1985]. For all three tracers, biological production in the surface ocean depletes the nutrient signature, and remineralization in deeper waters enriches the nutrient content. This latter process is what gives rise to the curvature in Figure 3b, but it is also what sets the different northern and southern end‐members themselves. Modern NADW is formed from waters of subtropical origin that have been depleted of their PO 4 , Cd, and 12C relative to 13C. Modern AABW, on the other hand, is formed from waters of higher nutrient content that experience lower overall biological nutrient utilization efficiency, thus leaving newly sinking deep waters with a higher PO 4 , Cd, and 12C to 13C ratio. While the processes that affect these nutrient tracers are fundamentally biological, their different deep water end‐member signatures provide a “quasi‐conservative” tracer of the circulation. A second important assumption of this work is that the Cd/Ca ratios and δ13C values recorded in benthic foraminifera are proportional to the values of the tracers in contemporaneous bottom waters. Core top calibrations of these correlations in the modern ocean are the best confirmation that this assumption is true [Duplessy et al., 1984]. Research into the biomineralization processes that actually set the tracer values in biogenic carbonate attempts to understand how these correlations come about [Duplessy et al., 1970; Erez, 1978; 2003]. The question, “Are foraminifera good physical chemists” is beyond the scope of this review, but the observation that modern forams produce good correlations for many different tracers is the basis for using them to constrain past circulation [Bemis et al., 1998; Boyle, 1988; Marchal and Curry, 2008]. These early works on sedimentary tracers of paleo‐circulation opened the field to many new studies using both Cd/Ca and δ13C. Several papers emphasized that the Pacific and the Atlantic basins were always chemically distinct from one another [Boyle and Keigwin, 1985/6; Shackleton et al., 1983], though work in and around the Southern Ocean established that this glacial “end‐member” looked much more like the Pacific than it does today [Charles and Faibanks, 1992; Oppo and Fairbanks, 1987]. Studies of this kind for the LGM Atlantic were compiled in a landmark paper from Duplessy et al. [Duplessy et al., 1988], which attempted to “map” the distribution of δ13C in a north‐south Western Basin section (Figure 5). This iconic figure has been redrawn several times—first by Broecker [Broecker and Denton, 1990] to eliminate the influence of high latitude North Atlantic planktonics, then by Sarnthein [Sarnthein et al., 1994] to emphasize the Eastern Basin, then by Curry and Oppo [Curry and Oppo, 2005] based on the addition of more South Atlantic depth transects, and finally by Marchitto and Broecker [Marchitto and Broecker, 2006] to depict Cd/Ca and δ13C side by side. Figure 5 is modified from Curry and Oppo and shows the main features of the result. At the LGM, the deep Atlantic is filled with an isotopically depleted (nutrient‐rich) water mass from the south that invades a region that is today filled by isotopically enriched (nutrient poor) northern source water. Glacial North Atlantic Intermediate Water (GNAIW) was volumetrically less important than the Glacial Southern Source Water (GSSW). This result has been interpreted many times to be the result of a reduced flux of NADW at the LGM as compared to AABW, but this interpretation ignores the importance of tracer diffusion and end‐member variability in setting the values of any tracer and emphasizes the role of transport [Marchal and Curry, 2008; Wunsch, 2003]. Strategies for how one can more quantitatively extract circulation information from these tracers will be taken up in a later section. Figure 5 Open in figure viewer PowerPoint 13C section in the Western Atlantic [after Curry and Oppo, 2005 Lund et al. 2011a 18O CaCO3 . LGM δC section in the Western Atlantic [after]. White dots are locations of data compiled and reported by Curry and Oppo. A shoaled and volumetrically more important southern sourced water mass is apparent when compared to the modern (see Figure 3 ). The depth transects at ~30°N and ~30°S were used into estimate the LGM profile of the conservative property δ Intermediate depth records during the last deglaciation from both the North and South Atlantic show a mirror image pattern compared to time series of the abyssal water variability [Came et al., 2003; Marchitto et al., 1998; Oppo and Horowitz, 2000; Rickaby and Elderfield, 2005]. In the competition for space in the deep Atlantic between northern and southern source waters, it appears that each time the abyss is ventilated from the south (high nutrient signal) the intermediate depths see an increase in the GNAIW signature (low nutrients). Before the LGM, there is sparse coverage of deep tracer dynamics at a variety of depths, but the evidence from Marine Isotope Stage (MIS) 3 and 4 changes in deep water properties in the North Atlantic do show a strong response to Heinrich Events [Vidal et al., 1997]. These periods seem to be a near cessation of deep input from northern sourced water masses. There is also some evidence for D/O events in the deep ocean, with Skinner's Mg/Ca data from the North Atlantic maybe the best example [Skinner and Elderfield, 2007] and Charles' benthic δ13C data from the South Atlantic the first to point to these changes in the deep [Charles et al., 1996]. Heinrich events seem to have a more robust deep water signature than the D/O transients. This feature of deep circulation is best demonstrated in cores where planktonic δ18O records from the Atlantic show characteristics of one polar ice core archive, while the benthic foraminifera records seem similar to the opposite polar ice core record Greenland ice core‐like signals while the benthic stable isotopes from the same core look much more like the Antarctic record [Charles et al., 1996; Curry and Oppo, 1997; Shackleton et al., 2000]. Other biogeochemical tracers have been developed to move beyond the mapping of paleo nutrients, mostly with the goal of understanding the ocean's role in setting pCO 2 variability. Most of these are focused on constraining pieces of the past carbon system, and while they are not strictly circulation tracers, they do respond to changes in deep water mass structure. Using Ba/Ca ratios from benthic foraminifera, Lea et al. established this deeply regenerated species as a paleo alkalinity tracer [Lea and Boyle, 1989; Lea and Boyle, 1990]. When combined with nutrient data, the record from 4200 m deep in the South Atlantic (core RC13‐229 in the Cape Basin) shows inferred carbonate ion concentrations that were anticorrelated with pCO 2 over two glacial cycles [Lea, 1995]. Marchitto et al. have developed the combined use of Zn/Ca and Cd/Ca in benthic forams to estimate past [CO 3 ] as well. Here the link to the carbonate system is based on the idea that Zn and Cd have different incorporation ratios into biogenic calcite at different calcite under saturation states [Marchitto et al., 2000]. Two aspects of boron chemistry have also been used to construct similar information. Boron isotope values in planktonic foraminifera seem to follow ambient pH in culture data [Hönisch et al., 2003; Sanyal et al., 1996] and show a glacial to interglacial signal that is consistent with the ice core pCO 2 data [Hönisch et al., 2009; Sanyal et al., 1995]. The extension of this tracer into benthic forams is a promising new advance to study deep ocean carbonate chemistry [Rae et al., 2011]. Recently the B/Ca ratios themselves have been implicated as a deep water saturation state tracer [Yu and Elderfield, 2007] and have been usefully employed to study the atmosphere and ocean carbon reservoir exchanges during the last deglaciation [Yu et al., 2010]. The details of these and other similar studies are easily the subject of a separate review paper of their own.

3.2 Physically Based Tracers of Water Mass Mixing Ratios Analysis of the fluxes that set the modern tracer distributions in the deep ocean is greatly aided by having both a conservative tracer, like T and S, and a nonconservative tracer, like [PO 4 ], [Cd], or δ13C. As described above, in paleoceanography the nonconservative tracers can be interpreted in a “quasi‐conservative” framework, but the tracer information content would be much more clear if we could make tracer‐tracer plots (like those in Figure 3) with a species that only responds to the mixing of deep water masses. The combination of δ13C and Cd/Ca data has been proposed as one such tracer. Recognizing that surface δ13C is set by both nutrient dynamics and air‐sea gas exchange, while Cd/Ca is only set by nutrient fluxes, Charles et al. [Charles et al., 1993] proposed that the “thermodynamic” component of δ13C, called δ13C as , should be a significant fraction of the total δ13C variability in surface and intermediate water DIC. In the intermediate waters south of Australia, Lynch‐Stieglitz et al. combined Cd/Ca and δ13C to estimate the δ13C as and use it as a conservative mixing tracer [Lynch‐Stieglitz et al., 1995]. Unfortunately the data are very scattered and few studies have followed up on this tracer. The isotopes of Nd (ε Nd ) found in sediments [Piotrowski et al., 2005; Rutberg et al., 2000], forams [Elmore et al., 2011], deep‐sea corals [van der Flierdt et al., 2010], or fish teeth [Martin and Scher, 2004] have been proposed to reflect, with various degrees of success, the isotopic composition of Nd in the deep waters that bathe them. In principle this water column ε Nd value is only affected by mixing of deep waters and could be used as an effective x axis for paleo tracer versus tracer plots. Northern source water has an unradiogenic value and is set in the surface North Atlantic like T and S. But the other ε Nd “end‐member” is set by sedimentary exchange in the Pacific with a radiogenic signal. This boundary condition in the deep Pacific is very different than the Southern Ocean surface fluxes of heat and salt, which define the southern deep water end‐member. Unlike ε Nd , other water mass tracers like Δ14C, CFCs, and [O 2 ] depend on surface properties near the boundary conditions of T and S and can aid in constraining mixing ratios. A bottom up tracer has been a goal of the oceanography community since at least the days of GEOSECS where much time and effort were spent to characterize the inputs of Ra to the ocean in the deep Pacific. If ε Nd could provide this view of the deep circulation, it would contain orthogonal information to T and S and be very useful in diagnosing water mass interaction. However, recent work has shown that as deep waters flow along margins they evolve in their ε Nd value due to boundary exchange fluxes over and above their “end‐member” inputs [Lacan and Jeandel, 2005; Lacan et al., 2012]. These processes add a nonconservative element to the water column ε Nd values. Overall the use of ε Nd as a tracer of water mass mixing is still under development. It has the distinct advantage of there being little chance of a biologically induced “vital effect,” but the basics of the tracer in the modern water column are still to be uncovered [Arsouze et al., 2007; Arsouze et al., 2010]. Important work on where in the water column the ε Nd value of foraminifera is set [Elmore et al., 2011] should be emphasized in the near term development of this tracer. Ideally the deep water mixing ratio tracer would come from a measure of salinity, temperature, or a proxy of either one. Indeed, one of the oldest problems in paleoceanography is how to separate the dual effects of temperature and δ18O water (salinity) from the single measurement of biogenic δ18O. For separate reasons, Emiliani [Emiliani, 1966] and Dansgaard [Dansgaard and Tauber, 1969] each assumed that the bulk of the glacial to interglacial signal in planktonic foraminifera is due to temperature. In contrast, Shackleton compared benthics and planktonics from the same core to show that δ18O water changes are actually about two third of the whole LGM signal [Shackleton, 1967]. Labeyrie used high latitude cores, where the freezing point is a defined temperature extreme, to make a similar point [Labeyrie et al., 1987]. Comparisons with the sea level record derived from surface corals have also been a useful way to normalize for the global portion of the δ18O water changes in benthic δ18O data. Chappell and Shackleton [Chappell and Shackleton, 1986] demonstrated that the deep Pacific cooled by 2°C at the MIS 5e to 5d transition and then warmed again by the same amount only at the last termination. Adopting a similar strategy, and a better constrained coral‐based sea level curve, Cutler et al. confirmed the Chappell and Shackleton result and found the deep Equatorial Atlantic warmed by ~4°C at the last termination [Cutler et al., 2003]. The difference between a 2°C cooling at the 5e/5d shift and this 4°C Termination 1 warming is made up by about 2°C of cooling over the glacial period that was largely paced by Milankovitch cycles. These types of records, and some others from high‐resolution benthic δ18O data, are shown in Figure 6. An assumption in this work is that the scaling of δ18O water with sea level at the LGM [Schrag et al., 2002] holds for the whole glacial cycle. Waelbroeck extended this approach by constructing a global δ18O water curve for the last four glacial cycles using relative sea level changes measured during just the last glacial cycle and applying separate δ18O benthic ‐sea level trends for glaciations and deglaciations [Waelbroeck et al., 2002]. These approaches have led to some very robust results (see Figure 6 and the discussion below), but they require that the global δ18O water signal be well mixed in the ocean during the slide into glacial maxima, an assumption that probably does not hold during deglaciation [Gebbie, 2012; Skinner and Shackleton, 2005]. Shackleton deconvolved the δ18O w signal from benthic forams by using the record of δ18O in atmospheric O 2 and a tuning strategy to line up the sediment and ice cores [Shackleton, 2000]. His assumption of a constant Dole effect over the last glacial cycle makes his analysis less robust than others described here, but a principal conclusion that temperature comprises a significant fraction of the benthic foraminifera δ18O variability has held up to subsequent scrutiny. Figure 6 Open in figure viewer PowerPoint 18O records (black diamonds, TR163‐22, 2830 m [Lea et al., 2002 Hodell et al., 2003 Shackleton et al., 2000 Curry and Oppo, 1997 Chappell and Shackleton, 1986 Cutler et al., 2003 18O w /sea level relationship for each core based on the local estimate of δ18O w at the LGM using pore fluids results [Adkins et al., 2002a 18O w = 1.1‰; ODP1089 δ18O w = 1.2‰; MD95‐2042 δ18O w = 0.9‰, EW9209‐1JPC δ18O w = 0.9‰). This procedure eliminates the interlab offset in δ18O measurements that can be seen in the MD95‐2042 data by scaling the measured glacial to interglacial amplitude of d18O rather than the absolute values. There is a clear difference in the history of deep temperature between the Pacific (TR163‐22) and the Atlantic (MD95‐2042 and EW9209‐1JPC) basins. Large Atlantic temperature changes at the MIS 5/4 boundary, and dropping into the LGM, are not found in the Pacific. The Southern Ocean (ODP1089, Atlantic Sector) is always the coldest water, though the deep North Atlantic cools to near freezing temperatures at the LGM, too. Deep ocean temperature changes from four high‐resolution benthic δO records (black diamonds, TR163‐22, 2830 m [.,]; red triangles with black outline, ODP1089, 4621 m [.,]; blue squares, MD95‐2042, 3146 m [.,]; green circles, EW9209‐1JPC, 4056 m []). Temperatures were estimated following the procedure of Chappell and Shackleton [] and Cutler et al. [.,]. I scaled the slope of the δ/sea level relationship for each core based on the local estimate of δat the LGM using pore fluids results [.,] and a total sea level change of 120 m (TR163‐22 δ= 1.1‰; ODP1089 δ= 1.2‰; MD95‐2042 δ= 0.9‰, EW9209‐1JPC δ= 0.9‰). This procedure eliminates the interlab offset in δO measurements that can be seen in the MD95‐2042 data by scaling the measured glacial to interglacial amplitude of d18O rather than the absolute values. There is a clear difference in the history of deep temperature between the Pacific (TR163‐22) and the Atlantic (MD95‐2042 and EW9209‐1JPC) basins. Large Atlantic temperature changes at the MIS 5/4 boundary, and dropping into the LGM, are not found in the Pacific. The Southern Ocean (ODP1089, Atlantic Sector) is always the coldest water, though the deep North Atlantic cools to near freezing temperatures at the LGM, too. Independent estimates of both the LGM deep ocean's δ18O water value and its salinity can be derived from pore fluid measurements of δ18O and [Cl] in suitable cores. This idea of McDuff's [McDuff, 1985] was first successfully employed by Schrag and DePaolo [Schrag and DePaolo, 1993] and Schrag et al. [Schrag et al., 1996] in the deep Atlantic. Treating the sediment column like a sand packed pipe where tracer can diffuse and advect in only the vertical direction allows the measured profile in δ18O and [Cl] at any site to constrain a model of the LGM bottom water values of temperature and salinity. The [Cl] to salinity conversion is straightforward and the δ18O water to temperature conversion only requires some benthic δ18O CaCO3 measurements in the same core. Using a limited, but globally distributed set of cores, Adkins et al. [Adkins et al., 2002a] were able to constrain the T/S plot for the LGM ocean using this technique (Figure 7a). This work confirmed the Schrag et al. inference that much of the deep ocean cooled to the freezing point at the LGM and found the surprising result that salty deep waters were produced in the Southern Ocean, not the North Atlantic, during this same time. The deep Atlantic salt gradient seen in Figures 3 and 4 evidently reversed sometime during the last glacial period. Exploring the implications of this result is part of the motivation for the last section of this paper. Figure 7 Open in figure viewer PowerPoint 18O w [Adkins et al., 2002a 18O CaCO3 (dashed). Lines of constant density curve with changing temperature because of the temperature dependence of the thermal expansion coefficient of seawater (known as “cabbeling”). At warm temperatures, isolines of δ18O and σ θ are parallel and benthic oxygen isotopes are a good proxy for density. At cold temperatures, the isolines cross. However, as the plot demonstrates, δ18O CaCO3 is a linear combination of T and S (using the “Mixed Deep Water” line of Legrand and Schmidt [LeGrande and Schmidt, 2006 18O water ) and can be used as a conservative tracer of water mass properties. The relationship between temperature and salinity in the deep ocean and how these two work in the equation of state of seawater. (a) T/S plot based on the pore fluid measurements of [Cl] and δ.,]. Deep stratification in the modern ocean is largely controlled by temperature, but at the LGM salinity is the chief cause of density differences at these core sites. In addition, the modern meridional gradient of salt in the Atlantic reversed at the LGM to a dominant salty source from the Southern Ocean. (b) A calculated T/S plot with lines of constant sigma‐theta (solid) and constant δ(dashed). Lines of constant density curve with changing temperature because of the temperature dependence of the thermal expansion coefficient of seawater (known as “cabbeling”). At warm temperatures, isolines of δO and σare parallel and benthic oxygen isotopes are a good proxy for density. At cold temperatures, the isolines cross. However, as the plot demonstrates, δis a linear combination of T and S (using the “Mixed Deep Water” line of Legrand and Schmidt [] Table 1 to convert salinity into δ) and can be used as a conservative tracer of water mass properties. Other tracers for deep ocean temperature have also been investigated. The Mg/Ca ratio in planktonic foraminifera has seen widespread use in recent years as a robust (if postdepositional dissolution effects can be minimized or calculated) tracer of SST [Elderfield and Ganssen, 2000; Rosenthal et al., 2000]. However, only a few studies have tried to use this technique in benthic forams. Martin et al. measured the last glacial cycle in the deep Pacific and found temperature changes that are broadly consistent with Chappell's and Cutler's results [Martin et al., 1998]. During MIS 3 and 4 these authors found deep ocean temperature shifts that correlate with pCO 2 changes measured in Antarctic ice cores. An elegant comparison of the CO 2 ‐temperature sensitivity during Termination 1 and the glacial period demonstrates that nearly all of the glacial age pCO 2 changes can be attributed to the temperature dependence of CO 2 solubility in the ocean [Martin et al., 2005]. Terminations, on the other hand, have too steep a temperature sensitivity to be explained by thermodynamics alone and must have CO 2 exchanges between reservoirs as part of their explanation. Skinner et al. have used the Mg/Ca ratio measured in the benthic species Globobulimina affinis to constrain time series of deep ocean temperature off the Iberian margin [Skinner and Elderfield, 2007]. They find large (~1.5°C) shifts associated with Heinrich Events where the water at ~2600 m both warms and cools by about 1°C. Recently, Elderfield et al. have shown, from a long record of coupled benthic δ18O and Mg/Ca from the South Pacific, that each glacial maximum reached very close to the freezing point for the last 1.4 Ma [Elderfield et al., 2012]. There are several promising new tracers of ocean temperature that could play an important role in future reconstructions of deep ocean circulation. Mg/Li ratios in foraminifera [Bryan and Marchitto, 2008] and deep‐sea corals [Case et al., 2010] are correlated with bottom water temperature. As both Mg and Li show large increases in the centers of calcification of biogenic carbonate, taking their ratio helps to normalize out this nontemperature‐dependent aspect of Mg incorporation into skeletons. The tracer has not been used to monitor past ocean conditions yet, but it clearly holds some promise. Noble gas measurements in terrestrial aquifers have been usefully employed to constrain the LGM tropical temperature change [Stute et al., 1995]. In a new twist on this idea, Kr/N 2 gas concentration ratios in ice cores are proposed to constrain the mean deep ocean temperature at any one time [Headly and Severinghaus, 2007; Ritz et al., 2011]. Assuming that these two gases are conserved, the gravitational fractionation effects can be accounted for, and employing a mass balance model, LGM values from GISP2 show a mean ocean temperature change of 2.7 ± 0.6°C. Unlike the aquifers, the inherent stratigraphy of ice cores should allow for this approach to provide time series of average deep ocean change in the future. Finally, a new approach to the stable isotopes of oxygen and carbon in biogenic skeletons has been pioneered by John Eiler and his group at Caltech. The relative propensity for 13C and 18O atoms to form bonds with each other in CaCO 3 , over and above a stochastic distribution based only on their relative abundances, has a temperature‐dependent equilibrium constant [Ghosh et al., 2006]. Carbonates from inorganic precipitation in the lab and a wide variety of biogenic skeletons show the same sensitivity to temperature change for this clumped isotope, or Δ 47 , ratio. Due to the analytical precision required to get 1–2°C error bars, the technique is sample intensive and will probably not be a good high‐resolution paleo‐temperature tool in the near future. However, the thermometer is independent of the δ18O water value and does show great promise in systems where large samples can be collected [Eiler, 2011; Thiagarajan et al., 2011]. This new technique is sure to advance our understanding of past deep ocean behavior via judicious application to problems where a few data points will make a lasting impact.