4.1. Control Run
 The control run captures the general δ13C features observed in the water column of the present-day Atlantic and in LH sediment cores. Compared to GLODAP, control run δ13C values are higher in the North Atlantic. This is most likely explained by the Suess effect, which has recently been shown to penetrate down to 2500 meters depth and southwards to 30°N in the North Atlantic [Olsen and Ninnemann, 2010]. Two cores near the Bahamas by Slowey and Curry , OCE205-33GGC and OCE205-100GGC, have higher δ13C values than are seen in the control run. GLODAP δ13CDIC measurements, however, contradict such high δ13Cforam values. The control run correlates very well with LH sediment data in the Northern Atlantic (see Figure 4), but underestimates the observations in some areas of the South Atlantic. Model δ13C values in the upper 500–2000 m in the South Atlantic are lower than foraminiferal δ13C by more than 0.5‰ PDB. This seems to be a general model problem, that is also present in the glacial model scenarios (see below).
 An altered Redfield stoichiometry for Southern Ocean and Antarctic surface and intermediate depth waters (as suggested by Zahn and Keir  or Lynch-Stieglitz et al. ) could be the source for the lower-than-observed Antarctic Intermediate Water (AAIW) δ13C in the model. Zahn and Keir  argued that isotopically depleted CO2(aq) readily enters the atmosphere once the upwelled waters come into contact with the sea surface. At the same time nutrients are not taken up by photosynthesis to such a degree that would ensure a constant Redfield stoichiometry, which in turn would lead to higher-than-expected δ13C values for AAIW.
 Following Broecker and Maier-Reimer  and Lynch-Stieglitz et al.  we investigate the contribution of the air-sea exchange (δ13Cas) to the carbon isotopic composition of surface water by removing the biological component from the total modeled or observed δ13C signal:
where PO4 is the phosphate concentration. Our control run yields positive values of δ13Cas in the AAIW formation area, which indicates that the isotopic signature of our modeled AAIW is substantially influenced by isotopic air-sea fluxes (Figure S4). However, our model values of δ13Cas (about 0.2‰ PDB) are significantly smaller than observations (of up to 1‰ PDB according to Mackensen et al. ) which suggests that our carbon cycle model underestimates the isotopic air-sea exchange in the AAIW formation area. This is probably due to the air-sea exchange formulation in the model which does not explicitly depend on the wind speed but employs a globally averaged gas transfer velocity value. This model deficit is also corroborated by the results from a numerical sensitivity study carried out by Broecker and Maier-Reimer  who (employing an earlier version of our model) found that by doubling the air-sea exchange rate of CO2 the δ13C values in the formation region of AAIW would increase by 0.4‰ PDB.
4.2. LGM Runs
 The model-data difference plots (Figures 5 and 6) indicate that model results are systematically higher than observations in the deep South Atlantic and along the North American coast. In the deep Southern Ocean south of 40°S GS simulates δ13C values that are higher by 0.23 to 0.84‰ PDB than a whole suite of sediment cores (PS1745-3 and PS2082-1 [Mackensen et al., 1994], TTN057-6 [Hodell et al., 2003], RC15-93, RC15-94, TN057-21 and V22-108 [Ninnemann and Charles, 2002], see Figures 5 and 6). The δ13C values in the cores reported by Ninnemann and Charles  are based on both C. wuellerstorfi and C. mundulus. Hodell et al.  showed that C. kullenbergi (which is the same species as C. mundulus [Yu et al., 2008]) records systematically lower δ13C values than C. wuellerstorfi. Extrapolating the δ13Ckullenbergi data scatter to −0.80‰ PDB suggests a δ13Cwuellerstorfi ≈ 0‰ PDB [see Hodell et al., 2001, Figure 1]. Some of the difference seen in the model-data comparison could therefore be taken up. Another possible influence is the phytodetritus effect [Mackensen et al., 1993], which causes foraminifera to record lower-than-expected δ13C values, typically explaining 0.4‰ PDB. Cores that lie close to an oceanic front are potentially affected. The coarse model resolution does not permit to capture steep oceanographic gradients such as oceanic fronts, which may also partly explain the model-data offset in the Southern Ocean. LGM reconstructions of oceanic fronts in the Southern Ocean by Gersonde et al. [2003, 2005] suggest that the Polar Front (PF), the Sub-Antarctic Front (SAF) and the Sub-tropical Front (STF) shifted northward by 3–5°. Careful comparison of the frontal positions with the locations of the relevant cores shows that PS1745-3 and RC15-93 fall exactly on the reconstructed PF, whereas PS2082-1 and TTN57-6 coincide with the reconstructed SAF. Cores RC15-94 and V22-108 fall in between the reconstructed PF and SAF, TN057-21 lies between the reconstructed SAF and STF. Since oceanic fronts meander about their mean position, the latter cores may also be affected by the phytodetritus effect. Hence, both factors, measurements on epibenthic species other than C. wuellerstorfi and changes in frontal positions with the associated phytodetritus effect, may explain model-data differences in the deep Southern Ocean.
 In the western North Atlantic below 4 km and between 20–30°N there are four cores which have δ13C values that are lower than GS by 0.40 to 0.62‰ PDB (KNR140-12JPC, KNR140-22JPC and KNR140-28GGC [Keigwin, 2004], EN120-1GGC [Boyle and Keigwin, 1987], see Figure 6). The horizontal flow fields of model runs CS and GS (not shown) reveal an AABW influx in the deep western North Atlantic. The model δ13C signal, however, is still too high. Keigwin  stresses that measurements of Holocene δ13C values in this location below 3 km do not agree with present-day DIC measurements, which may point toward yet unknown problems with these cores.
 There are also areas where model results indicate lower δ13C values than observational δ13C values. This is particularly true for the intermediate depth South East Atlantic, the Brazil margin cores, the central North Atlantic (see Figure S2), and the central South Atlantic at 3 km water depth.
 In the South East Atlantic all model scenarios (including the control run) yield lower δ13C values by about −1‰ PDB compared to the sediment data (175-1087A [Pierre et al., 2001], ODP1085A [Bickert and Mackensen, 2004], IOW226920-3 [Mollenhauer et al., 2002], and KW-31 [Sarnthein et al., 1994]). Sediment δ13Cforam data are consistently higher than any model simulation δ13C, and model results of temperature and salinity (not shown) do not point toward any anomalous water mass here. The low model δ13C is likely to be an artefact which may partly be caused by underestimation of carbon isotope air-sea exchange in the formation region of AAIW (see control run discussion above). For the LGM scenarios this misrepresentation might be exacerbated due to generally stronger and seasonally more varying glacial winds, which may cause the anomalously low δ13C signal in the South East Atlantic. One might also speculate about a Mediterranean influence: Zahn et al.  find that Mediterranean Outflow Water (MOW) has a δ13C signature that was higher during the LGM (greater than 1.6‰ PDB) when compared to today (1.3‰ PDB). Bickert and Mackensen  show that MOW extends southwards after leaving the Strait of Gibraltar. It cannot be verified, however, that MOW extends further south than 10°N at depths above 2500 meters, as there are no sediment cores at these depths away from the continental slope. Moreover, the difference in MOW δ13C between today and the LGM is small (0.3‰ PDB) compared to the difference seen in the model-data comparison in the middepth South East Atlantic (1‰ PDB). There is no Mediterranean in our LGM setup.
 Most of the Brazil margin cores of Curry and Oppo  between 25–35°S contain δ13C values that are higher by up to 0.60‰ PDB than model values in GS (e.g., CHN115-70PC, CHN115-89PC, or CHN115-91PC, see Figure 6f). The same data-model difference holds for the LH data and our control simulation (Figure 3c). The most likely culprit is again poor carbon isotope air-sea exchange in the model (see control run discussion above), which is likely to be more pronounced in our LGM runs. Additionally, upwelling of NADW-derived waters south of Cape Frio would introduce much higher δ13C values [Acha et al., 2004], but the model cannot resolve such local upwelling features.
 Several cores in the central North Atlantic between 25–40°W and at 2–3.5 km water depth are enriched in 13C with respect to either GS or CS (e.g., CHN824115 [Boyle and Keigwin, 1987], or T86-15P [Sarnthein et al., 1994]) with δ13C values relative to GS that are higher by 0.42 to 0.66‰ PDB (see Figure S2). This points toward a model NADW flowpath that is too shallow in the central North Atlantic.
 Analogously, there are four cores in the central South Atlantic by Bickert and Mackensen  for which scenario GS simulates δ13C values that are lower than observations by 0.58 to 0.88‰ PDB (GeoB3808-6, GeoB5115-2, GeoB5121-2, and GeoB2016-1). Again, one may speculate about a NADW signal in the sediments that neither model scenario captures as model-NADW is shoaling too much. Additionally, the model problems seen in the South East Atlantic may contribute by extending into the central South Atlantic.
4.3. Anomalies or Δδ13C
 The Δδ13C plots in Figures 8 and 9 have the advantage that systematic errors such as constant offsets in δ13C values in the sediments due to, e.g., upwelling, or model artefacts such as the one seen in the South East Atlantic, are reduced. Scenario GB performs poorly when compared to the sediment data (Figures 8a, 8c, 9a, and 9c). Δδ13C is similar in both CS and GS. For the sediments the Δδ13C = 0 line lies close to 2 km water depth. The same holds for GS, but not for CS, where the zero-line is above 1.4 km water depth (Figures 8b and 9b). In addition, the average sediment Δδ13C signal below 2 km water depth is less than 0.75‰ PDB, which is similar to GS. In CS this value is mostly above 0.75‰ PDB. This further strengthens the good agreement of scenario GS with the observations. The sediment cores south of 40°S are only affected by frontal upwelling during the LGM and not during the LH (see above). Therefore, the effect is not systematic, and the high Δδ13C in the sediments comes as no surprise.
 The differences between our three model scenarios are summarized in Figure 7. Scenarios GS and CS both correlate very well with the sediment data. GS, however, correlates better in the North, West and East Atlantic. Additionally, the variance in GS is closer to that of the reconstructions. Scenario GB performs poorly in the model-data comparison.
 The altered fresh water balance in the Southern Ocean which is employed in both, GS and CS, seems to be a crucial feature in our LGM simulations. It is caused by (1) enhanced northward sea ice export and melting away from the sea ice production zone, which in turn causes (2) a relative increase in brine rejection when new sea ice is forming. The overturning cell in the North Atlantic is shoaling and weakening for both scenarios, indicating another important LGM feature (Figure 3). The strength of the positive overturning cell, however, is less well constrained: 12 Sv for GS contrast with only 8 Sv for CS, although both scenarios show a good fit for the North Atlantic. This significant difference in the response of the model to the two different SST reconstructions deserves further explanation. It is important to note that the SST reconstructions have an impact on atmospheric wind patterns and evaporation/precipitation patterns that the atmospheric model generates, which in turn have an impact on freshwater and heat fluxes into the ocean, ocean circulation, and air-sea gas exchange [see also Romanova et al., 2004]. Scenario GS, for instance, has surface waters south of Iceland that are saltier by more than 2 PSU when compared to CS (not shown). This causes stronger downwelling and is very likely the reason for the more rigorous AMOC in scenario GS compared to CS. Since δ13C is not a purely kinematical tracer, it can only be used to reconstruct the geometries of water masses. The strength of the overturning cell cannot be assessed.
4.4. Relation to Previous Studies
 Previous model-data comparisons have either used a much reduced number of observations, or not employed a 3D OGCM. Winguth et al.  used ad hoc circulation fields and a limited amount of mostly East Atlantic observations. Their glacial first guess scenario yields a reduced North Atlantic overturning circulation, which is compensated for by an increased influx of Southern Ocean deep waters. This result is similar to what we find for scenario GS, but there are conceptual differences. Winguth et al.  prescribed estimated salinity fields which were additionally modified in high latitudes to reduce the model-data misfit. In our model setup, salinity is a fully prognostic variable, which is physically more consistent.
 Tagliabue et al.  employ an OGCM and a biogeochemistry model forced by different LGM boundary conditions. Their model scenario that agrees best with observations (CircA) has a reduced ventilation in the North Atlantic and reduced AABW export. Their increased AABW export scenario (CircB) does not agree well with observations. This is the opposite of what our δ13C model results show: our two best fitting scenarios arrive at increased AABW export from the Southern Ocean. Since increased AABW inflow into the North Atlantic is also supported by radiocarbon, grain size, and Pa/Th studies [Robinson et al., 2005; Hall et al., 2011; Negre et al., 2010] we believe that our GS and CS model scenarios are well suited to describe the LGM Atlantic Ocean state.
 So far most modeling studies have focused on changing the freshwater balance in the North Atlantic [Roche et al., 2007; Kageyama et al., 2009; Otto-Bliesner and Brady, 2010], with mixed successes regarding the integrity with observational data. Changes in freshwater production in the Southern Ocean [Adkins et al., 2002] have attracted comparatively less attention, but seem to be important [Stocker et al., 1992; Fichefet et al., 1994; Winguth et al., 1999; Seidov et al., 2001; Shin et al., 2003b; Schmittner, 2003; Butzin et al., 2005]. Schmittner  increased rates of sea ice formation and northward export while keeping the AMOC strength at present-day levels. This resulted in saltier and denser AABW, increased its formation rate, and led to a higher consistency with reconstructions of glacial bottom water properties. Shin et al. [2003a] modeled enhanced northward sea ice export in the Southern Ocean in a fully coupled ocean-atmosphere circulation model and arrived at a shoaled and weakened AMOC. Butzin et al.  found that modeling radiocarbon in the glacial ocean with a changed freshwater balance in the Southern Ocean agrees best with observations. Our study with its widespread collection of δ13C values puts these modeling efforts on a more comprehensive observational base and further highlights the Southern Ocean's role in influencing global glacial climate.